Géodynamique andine

Andean geodynamics

Geodinámica andina

7th International Symposium on Andean Geodynamics

Université de Nice Sophia Antipolis

2-4 septembre 2008

résumés étendus

extended abstracts

resúmenes ampliados

organisateurs / organisers / organizadores

Institut de recherche pour le développement Université de Nice Sophia Antipolis

____________________________________

IRD Éditions INSTITUT DE RECHERCHE POUR LE DÉVELOPPEMENT

Paris, 2008

ISAG

comité d’organisation organising committee comité organizador

J.-Y. Collot (IRD - Géosciences Azur), T. Monfret (IRD - Géosciences Azur),

T. Sempere (IRD - LMTG - U. Toulouse) B. Delouis (U. Nice - Géosciences Azur), E. Tric (U. Nice - Géosciences Azur),

G. Hérail (IRD, Lima)

comité scientifique et représentants nationaux scientific advisory board and national representatives

comité científico y representantes nacionales

V. Acocella (U. Roma), R. W. Allmendinger (Cornell U., Ithaca), P. Alvarado (U. San Juan), F. Audemard (FUNVISIS, Caracas),

J.-P. Avouac (Caltech, Pasadena), S. Beck (U. Arizona at Tucson), P. Charvis (IRD - Géosciences Azur), E. Flueh (GEOMAR, Kiel), Ll. Fontboté (U. Geneva), R. García (U. San Andrés, La Paz), Y. Gaudemer (IPGP, Paris), G. González (UCN, Antofagasta), A. J. Hartley (U. Aberdeen), E. Jaillard (IRD - U. Grenoble),

A. Kammer (U. Bogotá), J.-L. Le Pennec (IRD, Clermont-Ferrand), J. Martinod (U. Toulouse), O. Oncken (GFZ, Potsdam), L. Ortlieb (IRD, Bondy),

M. Pardo (U. de Chile), V. Ramos (U. Buenos Aires), C. Ranero (ICREA, Barcelona), R. Rodríguez (IGME, Madrid), S. Rosas (PUCP & SGP, Lima), F. Sàbat (U. Barcelona),

U. Schaltegger (U. Geneva), P. Soler (IRD, Marseille), H. J. Tavera (IGP, Lima), C. Vigny (ENS, Paris), W. Winkler (ETH, Zürich), H. Yepes (EPN, Quito)

aides financières / fundings / ayudas económicas

Institut de recherche pour le développement Université de Nice Sophia Antipolis

Institut national des sciences de l’univers (INSU) Région Provence-Alpes-Côte d’Azur

Laboratoire Géosciences Azur Conseil général des Alpes-Maritimes

© IRD, 2008 ISBN : 978-2-7009-1643-1

ISAG

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Sommaire / Contents / Contenido Seismic risk associated with the Magallanes-Fagnano continental transform fault, Tierra del

Fuego, Southern Argentina 13-16 L. ABASCAL & G. GONZÁLEZ-BONORINO Cambrian paleogeography at the western Gondwana margin: U-Pb ages and provenance areas

of detrital zircons of the Mesón Group (Upper Cambrian), Northwest Argentina 17-20 C. J. ADAMS, H. MILLER, G. F. ACEÑOLAZA, & A. J. TOSELLI Assessment of erosion rate modifications during the Neogene incision in the Semiarid Andes

(Northern Chile) using the Black Top Hat function applied to a Digital Elevation Model 21-24 G. AGUILAR, J. DARROZES, E. MAIRE, & R. RIQUELME Preliminary results of a geochemical survey at Lastarria volcano (Northern Chile): Magmatic

vs. hydrothermal contributions 25-28 F. AGUILERA, F. TASSI, O. VASELLI, E. MEDINA, & T. DARRAH Towards a geodynamical model for the “middle” Cretaceous very low-grade metamorphism

in Central Chile: The geochronological approach 29-32 L. AGUIRRE, V. OLIVEROS, D. MORATA, M. VERGARA, M. BELMAR, & S. CALDERÓN The Pichaihue Limestones (Late Cretaceous) in the Agrio fold and thrust belt, Neuquén Basin,

Argentina 33-36 B. AGUIRRE-URRETA, P. J. PAZOS, V. A. RAMOS, E. G. OTTONE, C. LAPRIDA, & D. G. LAZO Paleoseismic investigation on the Boconó fault between Las González and Estanques, Mérida

Andes, Venezuela 37-40 M. J. ALVARADO, F. A. AUDEMARD, J. LAFFAILLE, R. J. OLLARVES, & L. M. RODRÍGUEZ Seismic source study and tectonic implications of the historic 1958 Las Melosas, Central Chile,

crustal earthquake 41-43 P. ALVARADO, S. BARRIENTOS, M. SAEZ, M. ASTROZA, & S. BECK Dendrochronology of the Central Andes of Bolivia 44-47 J. ARGOLLO, C. SOLÍS, & R. VILLALBA An Andean mega-thrust synthetic to subduction?: The San Ramón Fault and associated seismic

hazard for Santiago (Chile) 48-51 R. ARMIJO, R. RAULD, R. THIELE, G. VARGAS, J. CAMPOS, R. LACASSIN, & E. KAUSEL Architecture and style of compressive Neogene deformation in the eastern-southeastern

border of the Salar de Atacama Basin (22°30’-24°15’S): A structural setting for the active volcanic arc of the Central Andes 52-55

F. ARON, G. GONZÁLEZ, E. VELOSO, & J. CEMBRANO Block rotations in the Puna plateau 56-59 C. ARRIAGADA, P. ROPERCH, & C. MPODOZIS Continental growth through protracted subduction and accretionary processes along Western

Gondwana: The case of the Ocloyic Orogeny in southern South America 60-63 R. A. ASTINI, G. COLLO, & F. MARTINA The 2007 Pisco earthquake (Mw=8.0), Central Peru: Preliminary field investigations and

seismotectonic context 64-66 L. AUDIN, H. PERFETTINI, D. FARBER, H. TAVERA, F. BONDOUX, & J.-P. AVOUAC The 2006 eruptions of the Tungurahua volcano (Ecuador) and the importance of volcano

hazard maps and their diffusion 67-70 D. BARBA, P. SAMANIEGO, J.-L. LE PENNEC, M. HALL, C. ROBIN, P. MOTHES, H. YEPES, P. RAMÓN,

S. ARELLANO, & G. RUIZ

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Control of Mesozoic extensional structures on the Andean deformation in the northern Malargüe fold and thrust belt, Mendoza, Argentina 71-74

F. BECHIS, L. GIAMBIAGI, D. YAGUPSKY, E. CRISTALLINI, V. GARCÍA, & J. MESCUA Flat-slab subduction beneath the Sierras Pampeanas in Argentina 75-76 S. BECK, P. ALVARADO, L. WAGNER, M. ANDERSON, H. GILBERT, & G. ZANDT The November 14, 2007, Mw=7.7 Tocopilla (Chile) earthquake: Preliminary results from InSAR and GPS 77-80 M. BÉJAR-PIZARRO, D. CARRIZO, A. SOCQUET, R. ARMIJO, J.-C. RUEGG, J.-B. de CHABALIER,

A. NERCESSIAN, O. CHARADE, & S. BONVALOT Spatial and temporal patterns of exhumation across the Venezuelan Andes from apatite

fission-track analysis: Implications for Cenozoic Caribbean geodynamics 81-83 M. BERMÚDEZ-CELLA, P. VAN DER BEEK, & M. BERNET Seismicity study of the Ecuadorian margin, using combined inshore-offshore seismological

network 84-87 N. BÉTHOUX, B. PONTOISE, V. ALVAREZ, Y. FONT, M. SEGOVIA, J.-Y. COLLOT, P. CHARVIS,

Y. HELLO, K. MANCHUEL, M. RÉGNIER, Y. FONT, J. DÍAZ, A. VILLASEÑOR, & A. GAILLER Geology of the Río Seco region, Deseado massif (48°35´S), Santa Cruz province, Argentina 88-89 P. P. BISCAYART, D. J. PÉREZ, L. E. ECHAVARRÍA, & M. J. CORREA Electrical conductivity beneath the Payún Matrú volcanic field in the Andean back-arc of

Argentina near 36.5°S: Insights into the magma source 90-93 A. BURD, J. R. BOOKER, M. C. POMPOSIELLO, A. FAVETTO, J. LARSEN, G. GIORDANENGO,

& L. OROZCO-BERNAL Crustal structure and tectonic deformation of the northern Chilean margin, 21-23.5ºS 94-96 A. CALAHORRANO, C. R. RANERO, U. BARCKHAUSEN, C. REICHERT, & I. GREVEMEYER Preliminary stratigraphic study of the San Francisco River volcanic sequence, northwestern

Purace volcano, Cauca, Colombia 97-100 E. CAÑOLA, S. M. LÓPEZ, G. E. TORO, & B. PULGARÍN Geochemical characterization of Volatile Organic Compounds (VOCs) in fluid discharges at

Copahue volcano (Argentina) 101-104 F. CAPECCHIACCI, F. TASSI, O. VASELLI, A. CASELLI, & M. AGUSTO The lithosphere of Southern Peru: A result of the accretion of allochthonous blocks during

the Mesoproterozoic 105-108 V. CARLOTTO, J. CÁRDENAS, & G. CARLIER Igneous rocks with adakitic-like signature in South America 109-112 S. I. CARRASQUERO Long-lived constrictional strain field of the inner part of the Andean Orocline: An example

of buttressing effect in oblique subduction curved margin 113-115 D. CARRIZO, G. GONZÁLEZ, & T. DUNAI The interplay between crustal tectonics and volcanism in the Central and Southern volcanic

zones of the Chilean Andes 116-119 J. CEMBRANO, G. GONZÁLEZ, L. LARA, E. VELOSO, E. MEDINA, F. ARON, M. BASSO, V. ORTEGA,

P. PÉREZ, & G. SIELFELD U-Pb geochronologic evidence for the Neoproterozoic-Palaeozoic evolution of the Gondwanan

margin of the North-Central Andes 120-123 D. CHEW, U. SCHALTEGGER, J. KO LER, T. MAGNA, M. J. WHITEHOUSE, C. L. KIRKLAND,

A. MI KOVI , A. CARDONA, & R. SPIKINGS Adakitic rocks and their geodynamic significance: Examples from the Andes of Ecuador and Peru 124-127 M. CHIARADIA, D. MERINO, & B. BEATE Seismotectonic analysis of the Bucaramanga Seismic Nest, Colombia 128-131 G. CHICANGANA & C. A. VARGAS

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Seismotectonic behavior of the Eastern Frontal Fault System: Seismic hazard for the Villavicencio region, Central Colombia 132-135

G. CHICANGANA, C. A. VARGAS, & A. KAMMER The Mw 7.7 Tocopilla earthquake of November 2007: Characteristics of a subduction

earthquake that occurred in the brittle-ductile transition zone of the northern Chile seismic gap 136-139

M. CHLIEH, D. RÉMY, B. DELOUIS, S. BONVALOT, G. GABALDA, T. MONFRET, & M. PARDO Progressive avulsion of the Río Pastaza as a response to topographic uplift and backtilt

of the Ecuadorian Subandean Zone 140-143 F. CHRISTOPHOUL, C. BERNAL, J. DARROZES, J.-C. SOULA, & J. D. BURGOS Superimposed deformational episodes along the migmatitic belt, central portion of the

Sierras Pampeanas Septentrionales, Central Andes, Argentina: An example from the Las Cañas Complex 144-147

C. E. CISTERNA, R. MON, & R. MENA Where is the evidence for Oligocene rifting in the Andes? Is it in the Loncopué Basin of

Argentina? 148-151 P. R. COBBOLD, E. A. ROSSELLO, & F. O. MARQUES Burial history and estimation of ancient thermal gradients in deep synorogenic foreland

sequences: The Neogene Vinchina Basin, south-Central Andes 152-155 G. COLLO & F. M. DÁVILA Coeval subduction erosion and underplating associated with a crustal splay fault at the

Ecuador-Colombia convergent margin 156-159 J.-Y. COLLOT, A. RIBODETTI, B. MARCAILLOU, & W. AGUDELO Active tectonics in the Central Chilean Andes: 3D tomography based on the aftershock

sequence of the 28 August 2004 shallow crustal earthquake 160-163 D. COMTE, M. FARÍAS, R. CHARRIER, & A. GONZÁLEZ Seismic structure of the continental margin offshore the southern Arauco Peninsula,

Chile, at ~38°S 164-167 E. CONTRERAS-REYES, I. GREVEMEYER, E. R. FLUEH, C. REICHERT, & M. SCHERWATH Fractures in the Mejillones Peninsula triggered by the Tocopilla Mw=7.7 earthquake 168-171 J. CORTÉS, D. RÉMY, G. GONZÁLEZ, J. MARTINOD, & G. GABALDA Analyse of the Tarapaca paleolandslide (North Chile) using generalized Newmark approach

and implications on paleosismicity, and on paleoclimate changes 172-175 J. DARROZES, J.-C. SOULA, J. INGLES, R. RIQUELME, & G. HÉRAIL Dynamic topography during flat-slab subduction: A first approach in the south-Central Andes 176-179 F. M. DÁVILA & C. LITHGOW-BERTELLONI Dynamics of the November 3, 2002 eruption of El Reventador volcano, Ecuador: Insights

from the morphology of ash particles 180-183 S. DELPIT, J.-L. LE PENNEC, P. SAMANIEGO, S. HIDALGO, & C. ROBIN Three-dimensional P- and S-wave seismic attenuation models in central Chile - western

Argentina (30°-34°S) from local recorded earthquakes 184-187 P. DESHAYES, T. MONFRET, M. PARDO, & E. VERA Magnetotelluric study of the Parinacota and Lascar volcanoes 188-190 D. DÍAZ, D. BRÄNDLEIN, & H. BRASSE Volcán Llaima (38.7ºS, Chilean Southern Volcanic Zone): Insights into a dominantly

mafic and ‘hyperactive’ subduction-related magmatic system 191-194 M. A. DUNGAN, C. BOUVET de MAISONNEUVE, D. SELLÉS, J. A. NARANJO, H. MORENO,

C. LANGMUIR, O. REUBI, S. GOLDSTEIN, J. JWEDA, S. ESCRIG, O. BACHMANN, & B. BOURDON

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Sedimentologic, paleontologic, and ichnologic evidence for deep-marine Miocene deposition in the present Intermediate Depression of south-central Chile (38°30’-41°30’S) 195-198

A. ENCINAS, K. L. FINGER, & L. BUATOIS Kinematics evolution of the Camisea Subandean thrust belt from apatite fission-track

thermochronology, Peru 199-202 N. ESPURT, J. BARBARAND, S. BRUSSET, P. BABY, M. RODDAZ, & W. HERMOZA Tectonic and glacial forcing of motion along an active detachment fault 203-205 D. L. FARBER & G. S. HANCOCK No subsidence in the development of the Central Depression along the Chilean margin 206-209 M. FARÍAS, S. CARRETIER, R. CHARRIER, J. MARTINOD, A. TASSARA, A. ENCINAS, & D. COMTE Southern Andean (34º-46ºS) tectonic evolution through the inception of Cretaceous to

Neogene shallow subduction zones: A south to north trend? 210-213 A. FOLGUERA & V. A. RAMOS Hypocentral determinations of earthquakes in a 3D heterogeneous velocity model, Ecuador

and Northern Peru: Preliminary results 214-215 Y. FONT, M. SEGOVIA, & H. TAVERA Determination of effective elastic thickness of the Colombian Andes using satellite-derived

gravity data with admittance technique 216-218 R. A. GALÁN & I. F. CASALLAS Numerical modeling of interplay between growth folds and fluvial-alluvial erosion-sedimentation

processes: Application to the Mendoza Precordillera orogenic front (32º30’S) 219-222 V. H. GARCÍA & E. O. CRISTALLINI 3D structure of the subduction zone at the Colombia–Ecuador border 223-226 L. C. GARCÍA-CANO, A. GALVE, P. CHARVIS, A. GAILLER, J.-X. DESSA, B. PONTOISE, Y. HELLO,

A. ANGLADE, & B. A. YATES Block uplift and intermontane basin development in the northern Patagonian Andes (38º-40ºS) 227-230 E. GARCÍA-MORABITO & V. A. RAMOS Pre-Andean deformation in the southern Central Andes (32°-33°S) 231-234 L. GIAMBIAGI, J. MESCUA, A. FOLGUERA, & A. MARTÍNEZ Origin of flat subduction zones: Numerical application to central Chile – western Argentina

between 29°S and 34°S 235-237 G. GIBERT, R. HASSANI, E. TRIC, & T. MONFRET The active upper plate deformation of the Central Andes forearc, northern Chile 238-241 G. GONZÁLEZ, R. ALLMENDINGER, T. DUNAI, J. CEMBRANO, J. MARTINOD, D. RÉMY, D. CARRIZO,

J. LOVELESS, E. VELOSO, F. ARON, & J. CORTÉS Modern geodata management — A tool for interdisciplinary interpretation and visualization 242-244 H.-J. GÖTZE, T. DAMM, & S. SCHMIDT Reflection seismic imaging of the Chilean subduction zone around the 1960 Valdivia

earthquake hypocenter 245-248 K. GROSS, S. BUSKE, S. A. SHAPIRO, P. WIGGER, & the TIPTEQ group Chile Triple Junction migration, mantle dynamics and Neogene uplift of Patagonia 249-252 B. GUILLAUME, J. MARTINOD, & L. HUSSON The dynamic forearc of southern Peru 253-256 S. R. HALL, D. L. FARBER, L. AUDIN, & R. C. FINKEL Oxygen isotopes evidence for crustal contamination and mantle metasomatism in the

genesis of the Atacazo-Ninahuilca magmatic suites, Ecuador 257-260 S. HIDALGO, M.-C. GERBE, H. MARTIN, G. CHAZOT, & J. COTTEN

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Incipient tectonic inversion in a segmented foreland basin: From extensional to piggyback settings 261-264

D. N. IAFFA, F. SÁBAT, O. FERRER, R. MON, & A. A. GUTIÉRREZ Mesozoic backarc basins in western Peru: A brief summary 265-268 J. JACAY, V. ALEJANDRO, A. PINO, & T. SEMPERE Geometric reconstruction of fault-propagation folding: A case study in the western

Cordillera Principal at 34º15’S-34º30’S 269-272 P. JARA, R. CHARRIER, M. FARÍAS, & C. ARRIAGADA Organization and evolution of a segmented deformation front: Llanos foothills, Eastern

Cordillera of Colombia 273-276 A. KAMMER, A. VELÁSQUEZ, A. BELTRAN, A. PIRAQUIVE, & W. A. ROBLES The PUNA passive seismic array in the southern Puna: Tests of lithospheric delamination

in the region of the Cerro Galán ignimbrite 277-280 S. MAHLBURG KAY, B. S. HEIT, B. L. COIRA, E. SANDVOL, X. YUAN, N. A. MCGLASHAN, D. COMTE,

L. D. BROWN, R. KIND, & D. ROBINSON Incision and erosion of the deepest Andean canyons in southern Peru, based on ignimbrites,

remote sensing, and DEM 281-284 A. de LA RUPELLE, J.-C. THOURET, F. ALBINO, T. SOURIOT, T. SEMPERE, & Y. GUNNELL Holocene submarine volcanoes in the Aysén fjord, Patagonian Andes (44ºS): Relations with

the Liquiñe-Ofqui Fault Zone 285-288 L. E. LARA Determination of an arc-related signature in Late Miocene volcanics over the San Rafael block,

Southern Central Andes (34º30´-37ºS), Argentina: The Payenia shallow subduction zone 289-291 V. D. LITVAK, A. FOLGUERA, & V. A. RAMOS Sedimentary constraints on the tectonic evolution of the paired Tumaco–Borbón and Manglares

forearc basins (southern Colombia - northern Ecuador) during the Late Cenozoic 292-294 E. LÓPEZ, J.-Y. COLLOT, & M. SOSSON Compressive active fault systems along the Central Andean piedmont 295-297 J. MACHARÉ, L. AUDIN, C. BENAVENTE, M. SAILLARD, V. REGARD, & S. CARRETIER Tracing a major crustal domain boundary based on the geochemistry of minor volcanic centres

in southern Peru 298-301 M. MAMANI, G. WÖRNER, & J.-C. THOURET Seismicity and structural implications in northern Ecuador from the Esmeraldas experiment 302-305 K. MANCHUEL, B. PONTOISE, N. BÉTHOUX, M. RÉGNIER, & the ESMERALDAS team Influence of trench sedimentation rate on heat flow and location of the thermally-defined

seismogenic zone in the North Ecuador – South Colombia margin 306-309 B. MARCAILLOU, G. SPENCE, K. WANG, J.-Y. COLLOT, & A. RIBODETTI Andesite magma generation at the Quaternary volcanic arc of southwest Colombia 310-314 M. I. MARÍN-CERÓN, T. MORIGUTI, & E. NAKAMURA Estimating building and infrastructure vulnerability in the city of Arequipa, Peru, from

volcanic mass flows: A challenge 315-318 K. MARTELLI, J.-C. THOURET, C. VAN WESTEN, D. FABRE, M. SHERIDAN, & R. VARGAS Metamorphic P-T constraints for the low-temperature assemblages overimposed on

metamorphic and igneous rocks nearby Ñorquinco Lake, Aluminé, North-Patagonian Andes 319-321 C. I. MARTÍNEZ-DOPICO Dynamic topography into the Amazonian basin: Insights from 3-D analogue modelling 322-325 J. MARTINOD, N. ESPURT, S. BRUSSET, F. FUNICIELLO, C. FACCENNA, & P. BABY

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Tectonic control on the 1960 Chile earthquake rupture segment 326-329 D. MELNICK, M. MORENO, D. LANGE, M. R. STRECKER, & H. P. ECHTLER Late Jurassic extensional tectonics in the southwestern Mendoza province, Argentina 330-333 J. F. MESCUA, L. GIAMBIAGI, & F. BECHIS

Moment-tensor inversion of explosion events recorded on the Ubinas volcano, Peru 334-336 J.-P. MÉTAXIAN, V. MONTEILLER, O. MACEDO, G. S. O’BRIEN, E. TAIPE, & C. J. BEAN

Tectono-magmatic evolution and crustal growth along west-central Amazonia since the late

Mesoproterozoic: Evidence from the Eastern Cordillera of Peru 337-338 A. MI KOVI & U. SCHALTEGGER Seismic tomography of the Cotopaxi volcano, Ecuador 339-340 V. MONTEILLER, J.-P. MÉTAXIAN, B. VALETTE, & S. ARAUJO

Analysis of the January 23, 2007 Aysén swarm using joint hypocenter determination 341-343 C. MORA, D. COMTE, R. RUSSO, & A. GALLEGO

Further evidences of Quaternary activity of the Maradona faulting, Precordillera Central,

Argentina 344-347 S. M. MOREIRAS & A. L. BANCHIG Contemporary forearc deformation in south-central Chile from GPS observations (36-39°S) 348-350 M. MORENO, J. KLOTZ, D. MELNICK, H. P. ECHTLER, & K. BATAILLE Regional tephro-stratigraphic correlation in the Ecuadorian coastal region 351-353 P. A. MOTHES, S. VALLEJO, & M. L. HALL Interactions between block rotations and basement tectonics in the Copiapó-Vallenar region,

northern Chile: Preliminary results 354-356 C. MPODOZIS, C. ARRIAGADA, P. ROPERCH, & E. SALAZAR Tracing petrogenetic crustal and mantle processes in zircon crystals from rocks associated with

the El Teniente porphyry Cu-Mo deposit in the high Andes of central Chile: Preliminary results 357-360 Marcia MUÑOZ, R. CHARRIER, V. MAKSAEV, & M. FANNING The brittle/ductile transition in the lithosphere of the Andes region and its relationship with

seismogenesis 361-364 Miguel MUÑOZ Nature of a topographic height in the Tarapacá pediplain, Northern Chile 365-368 V. MUÑOZ, G. HÉRAIL, & M. FARÍAS Stratigraphy of the synorogenic Cenozoic volcanic rocks of Cajamarca and Santiago de Chuco,

northern Peru 369-372 P. NAVARRO, C. CERECEDA, & M. RIVERA Characterization of the Sierras de Córdoba eastern boundary from gravimetry, magnetotelluric

and DEM (Argentina) 373-376 L. A. OROZCO, E. A. ROSSELLO, C. POMPOSIELLO, A. FAVETTO, & C. P. BORDARAMPÉ Crustal seismicity and 3D seismic wave velocity models in the Andes cordillera of Central Chile

(33°-34.5°S) from local earthquakes 377-380 M. PARDO, E. VERA, T. MONFRET, & G. YAÑEZ Why is the passive margin of Argentinean Patagonia uplifting?: An insight by marine terrace

and tidal notches sequences 381-383 K. PEDOJA, V. REGARD, L. HUSSON, J. MARTINOD, & M. IGLESIAS Neotectonics and mass wasting phenomena in the eastern slope of the southern Central Andes

(37º-37º30’S) 384-386 I. M. PENNA, R. L. HERMANNS, & A. FOLGUERA

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Current erosion rates in the Northern and Central Andes: Evaluation of tectonic and climatic controls 387-390

E. PÉPIN, S. CARRETIER, J.-L. GUYOT, E. ARMIJOS, H. BAZAN, P. FRAIZY, L. NORIEGA, J. ORDÓÑEZ, R. POMBOSA, & P. VAUCHEL

The volcanic rocks of the Mondaca river, Cordillera Principal (31°45'S), San Juan province,

Argentina 391-392 D. J. PÉREZ & J. M. SÁNCHEZ-MAGARIÑOS Geophysical modeling of intrusive bodies: A case study in the Fuegian Batholith. Argentina 393-396 J. I. PERONI, A. TASSONE, M. CERREDO, H. LIPPAI, M. MENICHETTI, E. LODOLO, & J. F. VILAS Influence of tectonic and magmatic parameters in the deformation of the Andean subduction

margin in Central Chile based on analogue models 397-400 L. PINTO, F. ALBERT, & R. CHARRIER Structural styles in the Eastern Cordillera, Subandean Ranges - Santa Barbara System transition,

and Lomas de Olmedo Trough (northern Argentine Andes) 401-404 J. POBLET, M. BULNES, R. E. SEGGIARO, N. G. AGUILERA, L. R. RODRÍGUEZ-FERNÁNDEZ,

N. HEREDIA, & J. L. ALONSO Paleomagnetic results from the Antarctic Peninsula and its relation with the Patagonian Andes 405-408 F. POBLETE & C. ARRIAGADA Altiplano-Puna elevation budget and thermal isostasy 409-412 C. PREZZI, H.-J. GÖTZE, & S. SCHMIDT Subduction partitioning evidenced by crustal earthquakes along the Chilean Andes 413-416 J. QUEZADA & K. BATAILLE Constraints on delamination from numerical models 417-420 J. QUINTEROS, V. A. RAMOS, & P. M. JACOVKIS Magmatic history of the Fitz Roy Plutonic Complex, Southern Patagonia (Argentina) 421-422 C. RAMÍREZ de ARELLANO, B. PUTLITZ, & O. MÜNTENER Late Cretaceous synorogenic deposits of the Neuquén Basin (36-39°S): Age constraints from

U-Pb dating in detrital zircons 423-426 V. A. RAMOS, M. PIMENTEL, & M. TUNIK Revisiting accretionary history and magma sources in the Southern Andes: Time variation of

“typical Andean granites” 427-430 C.W. RAPELA, R.J. PANKHURST, J.A. DAHLQUIST, E.G. BALDO, C. CASQUET, & C. GALINDO Recent debris-flows and megaturbidite in a confined basin of the North Ecuador subduction

trench 431-434 G. RATZOV, J.-Y. COLLOT, M. SOSSON, & S. MIGEON Geomorphology of the Fitzcarrald Arch, Peru, and its relationships with the Nazca plate

subduction 435-438 V. REGARD, R. LAGNOUS, N. ESPURT, J. DARROZES, P. BABY, M. RODDAZ, Y. CALDERON,

& W. HERMOZA Orientation of current crustal stresses in the South America plate between 30° and 55°S 439-441 C.-D. REUTHER & E. MOSER New field studies in the Gonzanamá, Catamayo and Malacatos-Vilcabamba basins, Ecuador:

Preliminary results 442-445 P. REYES, F. MICHAUD, P. CARBONEL, & M. FORNARI Petrology of the 2006-2007 tephras from Ubinas volcano, southern Peru 446-449 M. RIVERA, M.-C. GERBE, A. GOURGAUD, J.-C. THOURET, H. MARTIN, J.-L. LE PENNEC, & J. MARIÑO Comparative methodological considerations for estimating fracture parameters 450-453 W. ROBLES, A. KAMMER, M. MARENTES, & W. ESPITIA

10

Subduction control on chemical composition of Oligocene to Quaternary sediments of the Ecuadorian Amazonian foreland basin from major and trace elements and Nd-Sr isotopes 454-457

M. RODDAZ, F. CHRISTOPHOUL, J.-C. SOULA, J. D. BURGOS-ZAMBRANO, P. BABY, & J. DÉRAMOND Neogene erosion and relief evolution in the Central Chile forearc (33°-34ºS) as determined

by detrital heavy mineral analysis 458-460 M. P. RODRÍGUEZ, L. PINTO, & G. HÉRAIL The Loncopué Trough: A major orogenic collapse in the western Agrio fold-and-thrust belt

(Andes of Neuquén, 36º40´-38º40´S) 461-464 E. ROJAS-VERA, A. FOLGUERA, G. ZAMORA-VALCARCE, & V. A. RAMOS The Cordillera Blanca fault system as structural control of the Jurassic-Cretaceous basin

in central-northern Peru 465-468 D. ROMERO Block rotations within the northern Peruvian Altiplano 469-472 P. ROPERCH, V. CARLOTTO, & A. CHAUVIN From steady state to climate-driven denudation across the Central Andes in SE Peru 473 G. M. H. RUIZ & V. CARLOTTO Pleistocene uplift rates variability along the Andean active margin inferred from marine

terraces 474-476 M. SAILLARD, L. AUDIN, G. HÉRAIL, S. HALL, D. FARBER, J. MARTINOD, & V. REGARD 3D structure of the Tres Cruces synclinorium from seismic data and serial balanced

cross-sections, Eastern Cordillera, Argentina 477-480 L. SALAZAR, J. KLEY, E. ROSSELLO, R. MONALDI, & M. WIEGAND Analysis of microseismicity in the Precordilleran Fault System at 21°S in Northern Chile 481-484 P. SALAZAR, J. KUMMEROW, G. ASCH, D. MOSER, & P. WIGGER Relations between plutonism in the back-arc region in southern Patagonia and Chile Rise

subduction: A geochronological review 485-488 A. SÁNCHEZ, F. HERVÉ, & M. de SAINT-BLANQUAT Gravity field analysis and preliminary 3D density modeling of the lithosphere at the

Caribbean-South American plate boundary 489-492 J. SÁNCHEZ, H.-J. GÖTZE, M. SCHMITZ, & C. IZARRA Upper lithospheric structure of the subduction zone in south-central Chile: Comparison for

differently aged incoming plate 493-495 M. SCHERWATH, E. CONTRERAS-REYES, E. R. FLUEH, & I. GREVEMEYER Are the Falkland Plateau and the Deseado Massif part of the same Mesoproterozoic

lithospheric block? 496-499 M. SCHILLING & A. TASSARA Principal results of the Caracas, Venezuela, Seismic Microzoning Project 500-503 M. SCHMITZ, J. J. HERNÁNDEZ, C. MORALES, D. MOLINA, M. VALLEÉ, J. DOMÍNGUEZ,

E. DELAVAUD, A. SINGER, M. GONZÁLEZ, V. LEAL, & the Caracas Seismic Microzoning Project working group

Anatomy of the Central Andes: Distinguishing between western, magmatic Andes and

eastern, tectonic Andes 504-507 T. SEMPERE & J. JACAY Direct versus indirect thermochronology: What do we truly trace? An example from SE Peru

and its implication for the geodynamic development of the Andes 508 D. SEWARD, G. M. H. RUIZ, & J. BABAULT Major mid-Cretaceous plate reorganization as the trigger of the Andean orogeny 509-512 R. SOMOZA

11

Linkage between Neogene arc expansion and contractional reactivation of a Cretaceous fold-and-thrust belt (southern Central Andes, 36º-37ºS) 513-516

M. G. SPAGNUOLO, A. FOLGUERA, & V. A. RAMOS Tectonic response of the central Chilean margin (35°-38°S) to the collision and subduction

of heterogeneous oceanic crust: A thermochronological study 517-520 R. SPIKINGS, M. DUNGAN, J. FOEKEN, A. CARTER, L. PAGE Fluvial responses to regional tectonic and local tectonic evolution of Oxaya anticline in

hyper-arid area, Arica (North Chile) 521-523 M. STRUB, J. DARROZES, L. AUDIN, E. MAIRE, G. HÉRAIL, J.-C. SOULA, & J. DÉRAMOND Tithonian to Aptian volcanism in the central Patagonian Cordillera, Aysén, Chile (45°-46°30’S):

U-Pb shrimp new data 524-525 M. SUÁREZ, R. DE LA CRUZ, & M. FANNING Anatomy of the Andean forearc controlling short-term interplate seismogenesis and

long-term Cordilleran orogenesis 526-529 A. TASSARA, R. HACKNEY, & D. LEGRAND A geochemical survey of geothermal resources in the Tarapacá and Antofagasta regions

(northern Chile) 530-533 F. TASSI, F. AGUILERA, O. VASELLI, & E. MEDINA The June 23, 2001, southern Peru earthquake 534-537 H. TAVERA & I. BERNAL Preliminary petrological, geochemical and stratigraphical characterization of the Sotará

volcano, SW Colombia 538-541 L. TÉLLEZ, M. I. MARÍN-CERÓN, G. TORO, & B. PULGARÍN Quaternary soft-linked fault systems highlighted through drainage anomalies in the

northwestern Precordillera Sur (32ºS), Central Andes of Argentina 542-544 C. M. TERRIZZANO & J. M. CORTÉS Neogene ignimbrites and volcanic edifices in southern Peru: Stratigraphy and

time-volume-composition relationships 545-548 J.-C. THOURET, M. MAMANI, G. WÖRNER, P. PAQUEREAU-LEBTI, M.-C. GERBE, A. DELACOUR,

M. RIVERA, L. CACYA, J. MARIÑO, & B. SINGER The Proterozoic basement of the Arequipa massif, southern Peru: Lithologic domains

and tectonics 549-552 P. TORRES, A. ALVÁN, & H. ACOSTA Trachydacitic domes in the caldera of Pino Hachado, province of Neuquén, Argentina 553-554 C. TUNSTALL, J. E. CLAVERO, & V. A. RAMOS Controls on erosion and clastic sediment flux in the Central Andes during the Late Cenozoic 555-557 C. E. UBA, G. ZEILINGER, & M. STRECKER Diente Verde and Mario, Cañada Honda, San Luis, Argentina: Porphyry-type deposits in the

South Pampean flat-slab region of the Central Andes 558-561 N. E. URBINA & P. SRUOGA Relationship between topography and seismicity in the Peruvian Andes: Influence of

topography on stress field 562-565 V. M. URIBE, L. AUDIN, H. PERFETTINI, & H. TAVERA The Peruvian Pataz, Parcoy and Huachón districts: Evidence for a coherent, 400 km-long,

Carboniferous orogenic gold belt along the Eastern Andean Cordillera? 566-568 E. VÁGÓ & R. MORITZ Chemical and mineralogical characterization of the River Huasco (Norte Chico, Chile) 569-570 A. VALDÉS, M. POLVÉ, & D. MORATA

12

Climatic impact on the erosive dynamics of the Pacific Central Andes revealed by cosmogenic and hydrological records of river sediments 571-572

R. VASSALLO, E. PÉPIN, V. REGARD, J.-L. GUYOT, S. CARRETIER, E. GAYER, L. AUDIN, F. CHRISTOPHOUL, R. RIQUELME, J. J. ORDÓÑEZ, & F. ESCÓBAR-CÁCERES

Thermotectonic history of the Northern Andes 573-576 D. VILLAGÓMEZ, R. SPIKINGS, D. SEWARD, T. MAGNA, W. WINKLER, & A. KAMMER Cenozoic high-strontium andesites in the Eastern Cordillera of Northwestern Argentina,

Central Andes 577-579 J. M. VIRAMONTE, N. SUZAÑO, C. PRESCOTT, R. BECCHIO, J. G. VIRAMONTE, M. ARNOSIO,

& M. M. PIMENTEL Heterogeneous thermal overprint of a Late Palaeozoic fore-arc system in north-central Chile

(32°–31°S) discernible by small scale equilibration and age domains (Ar-Ar; fission track) 580-582 A. P. WILLNER, H.-J. MASSONNE, M. SUDO, & S. THOMSON Upper Pleistocene deglaciation as a conditioning factor for catastrophic mass redistribution

in Las Cuevas basin, Mendoza, Argentina 583-586 C. G. J. WILSON, R. HERMANNS, L. FAUQUÉ, M. ROSAS, V. BAUMANN, & K. HEWITT Timing and causes of the growth of the Ecuadorian cordilleras, as inferred from their

detrital record 587-591 W. WINKLER, C. VALLEJO, L. LUZIEUX, R. SPIKINGS, & N. MARTIN-GOMBOJAV Damage zone and the occurrence of world-class porphyry copper deposits in the active

margin of Chile: Geophysical signatures and tectonomagmatic inferences 592-593 G. YÁÑEZ, O. RIVERA, D. COMTE, M. PARDO, L. BAEZA, & E. VERA AUTHOR INDEX 594-597

The Organizing Committee makes clear that the authors are responsible for the quality of the text of their communication, the relevance and exactness of the references they have cited, and the accuracy of their affiliation and address.

Only abstracts submitted in English have been accepted.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 13-16

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Seismic risk associated with the Magallanes-Fagnano continental transform fault, Tierra del Fuego, Southern Argentina

L. del V. Abascal1 & G. González-Bonorino

2

1

UTN-FRRG, Islas Malvinas 1650, 9420 Río Grande, Tierra del Fuego, Argentina (liliana_abascal@hotmail.com) 2

CADIC-CONICET, B. Houssay 200, 9410 Ushuaia, Tierra del Fuego, Argentina (bonorino@cadic.gov.ar)

KEYWORDS : seismic hazard, Tierra del Fuego, Argentina, Magallanes-Fagnano Fault, Andes

Introduction

The Magallanes-Fagnano (MF) and the North Scotia Ridge fault systems constitute the remnant of a transform

fault extending from the Sandwich Islands mid-ocean ridge to the Chilean subduction trench, along the boundary

between the South American and the Scotia lithospheric plates (British Antarctic Survey, 1985; Pelayo and

Wiens, 1989). The MF fault represents the continental segment of this transform fault in the island of Tierra del

Fuego, in southernmost South America, where it delimits two structural domains: a thin-skinned domain north of

the fault, and a thick-skinned domain south of the fault (Winslow, 1982). The MF fault trace measures about

200 km in length; it strikes EW in eastern Tierra del Fuego, and curves smoothly to the NW in western parts of

the island, in Chilean territory. This fault system comprises distinct tectonic lineaments arranged in an “en

échelon” geometry. The master segments are near-vertical faults (Lodolo et al., 2003). A left-lateral dominant

direction of movement along this fault at a rate of 6.6 mm/yr was documented by Smalley et al. (2003).

From the IRIS database (http:// www.iris.washington.edu/ SeismiQuery/ events.htm) were taken 3993 seismic

events recorded between 1/I/1970 and 25/VIII/2007, within the area between 48ºS and 70ºS, and 20ºW and

76ºW, encompassing the Scotia Arc and southernmost Patagonia; magnitudes ranged from 3.1 to 7.8; aftershocks

were filtered out. Several supplementary seismic events from other sourced were also included. On the basis of

fault geometry and mechanics, and clustering of epicenters, seven seismogenic zones were defined: North Scotia

Ridge zone, South Scotia Ridge zone, Sandwich Islands subduction zone, Schackleton Fault Zone, Chilean

subduction zone, and the MF zone. Only the latter two seismogenic zones lie sufficiently close to urban centers

in Tierra del Fuego to represent a hazard. We wish to state that the possibility of tsunami generation from

seismic activity in the Scotia Arc fault systems was not evaluated in this work.

On December 1949 two earthquakes with similar Richter magnitudes of 7.8 shook the island of Tierra del

Fuego with a 9-hour interval. Secondary effects were large wave setup in Lake Fagnano and downdrop of a

southeastern sector of Lake Fagnano and the adjacent Turbio River deltaic plain, giving rise to a coastal lagoon.

On the basis of personal accounts of damage distribution, it has been assumed that the first event had epicenter to

the east of the second event (Costa et al., 2006). Recent relocation of the 1949 epicenters, as well as that for the

June 1970 earthquake (M=7.0), shows all three clearly aligned with the trace of the MF fault (P. Alvarado, pers.

comm., 2007).

This paper presents the first quantified assessment of seismic hazard for the province of Tierra del Fuego. This

assessment is largely based on local data and takes into account the amplifying effect of the Quaternary deposits.

Two previous studies are of a regional scope. In 1985, the Argentine Institute for Seismic Prevention (INPRES)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 13-16

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divided Argentina into 5 seismic zones. Data for Tierra del Fuego relied on information from accelerographs

located outside the island. Major flaws in the zonation were that the MF fault lay at the boundary between two

seismic zones and that highest hazard was located well away from the MF fault. A more recent assessment of

seismic hazard in Tierra del Fuego is due to the Global Seismic Hazard Assessment Program (GSHAP), which

assigns the island Peak Ground Accelerations (PGA) between 0.8 and 1.6 m/sec2, with an exceedance probability

of 10% in 50 years.

In addition to the regional-scale seismic hazard evaluation, the seismic risk for Tolhuin associated with the MF

fault is considered. Tolhuin, with a population of about 3,000, sits on the eastern end of Lake Fagnano, less than

one kilometer from the trace of the MF fault. Evaluation of the seismic risk for Tolhuin followed the procedures

outlined by the United Nations in Risk Assessment tools for Diagnosis of Urban areas against Seismic disasters

(RADIUS). The RADIUS methodology comprises 5 major steps: 1 – Definition of the seismic scenario, setting

likely epicenter locations and earthquake magnitudes; 2 – Designing, or selecting from preexisting formulas, the

seismic attenuation law to be applied; 3 – Calculating the amplifying effect of substrate layers on the basis of

their geotechnical properties; 4 – Converting Peak Ground Acceleration values to Modified Mercalli Intensity

scale values; and 5 –Applying vulnerability functions according to construction type.

Results

The seismic scenario assumed in the evaluation of the provincial seismic hazard was as follows: a M=8.5

Maximum Credible Earthquake, with epicenter where the MF fault intersects the Chile-Argentina border; such

location corresponds well with recorded earthquakes. Tierra del Fuego is largely covered by thick glacial drift

and by thinner fluvio-glacial deposits and peat bogs. Preliminary geotechnical studies suggest that the drift

behaves as a stiff soil and may represent an amplification

factor in the order of 1.2, mainly due to thicknesses in

excess of 10 meters, whereas the fluvio-glacial and the peat

bog deposits may represent an amplification factor in the

order of 2.0 (estimates were obtaine with EERA software;

Bardet et al., 19 ). Due to insufficient local data, the

attenuation law of Campbell (1997) was used. PGA values

obtained from Campbell`s (1997) attenuation law were

converted to MMI values through the equation MMI =

3,333*(log10(PGA*980) – 0,014) (Trifunac y Brady,

1975). Overall, MMI decreases radially from the epicenter

but noticeable anomalies can be observed in areas of soft

soil (Fig. 1).

Constructions in Tolhuin were classified into two

categories, depending on whether they are seismic-resistant

or not. Tolhuin is a recently developed urban center and

buildings are mostly younger than 10 years old. Public buildings, and household dwellings built by the

provincial and national government agencies, generally fall in the seismic resistant category; that is, they comply

Figure 1. Seismic hazard distribution for Tierra del Fuego Province.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 13-16

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with the CIRSOC-103 regulations issued by the INPRES. The majority of the households do not, however,

having been built on the rush by small local constructors. Tolhuin is built on terraced ground generally sloping

toward Lake Fagnano, the SW. The town center, and the majority of the public buildings are in high ground,

approximately 100 m above lake level. The thicker Quaternary section (>150 m in thickness) underlies the town

center. In this area, however, stiff glacial drift and gravelly glacifluvial deposits dominate the upper levels,

resulting in an amplification factor of about 1.2. Lower parts of the town rest on thinner but more clay-rich

substrate, and resulted in an amplification factor of about 2. Two seismic scenarios were evaluated for Tolhuin,

both for M=8.5 earthquakes located on the MF fault, one with epicenter on the Chilean border, approximately

80 km away, the other distant only 20 km from Tolhuin`s urban center. The results are shown in Figure 3.

Seismic resistant buildings would resist and MMI=8.5 with little damage but precarious construction would

suffer considerable damage, especially those located on the lower slopes.

Acknowledgments

We wish to acknowledge the financial support obtained from Consejo Federal de Ciencia y Tecnología

(COFECYT) through the PFIP 2005-Convenio Nº 063. We also wish to thank the Municipal authorities of

Tolhuin and the private and public organisms that provided useful base information for this study.

A B

Figure 2. Seismic scenarios for Tolhuin. A) Epicenter 20 km from city center. B) Epicenter 80 km from Tolhuin.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 13-16

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References Bardet, J.P., Ichii, K., & Lin, C.H., 2000. EERA, A computer program for equivalent-linear earthquake site response analysis

of layered soil deposits. University of Southern California, Los Angeles. British Antarctic Survey 1985. Tectonic map of the Scotia Arc. British Antarctic Survey, Miscellaneous 3, Cambridge. Campbell, K.W. 1997. Empirical near-source attenuation relationships for horizontal and vertical components of peak ground

acceleration, peak ground velocity, and pseudo-absolute acceleration response spectra. Seismological Research Letters 68(1):154-179.

Costa, C.H., Smalley, R., Schwartz, D., Stenner, H., Ellis, M., Ahumada, E., Velasco, M-S. 2006. Paleoseismic observations of an onshore transform boudary: The Magallanes-Fagnano fault, Tierra del Fuego, Argentina. Revista de la Asociación Geológica Argentina 61 (4):647-657.

Klepeis, K. 1994. The Magallanes and Deseado fault zones: Major segments of the South American-Scotia transform plate boundary in southernmost South America, Tierra del Fuego. Journal Geophysical Research 99:22,001-22,014.

Kraemer, P. 2003. Orogenic shortening and the origin of the Patagonian Orocline (56ºSLat.). Journal of South American Earth Sciences 15: 731-748.

Lodolo, E., Menichetti, M., Bartole, R., Ben-Avraham, Z., Tassone, A., Lippai, H. 2003. Magallanes-Fagnano continental transform fault (Tierra del Fuego, southernmost South America). Tectonics, 22(6), doi 10.1029/2003TC001500

Pelayo, A., Wiens, D., 1989. Seismotectonics and relative plate motions in the Scotia Sea region. Journal of Geophysical Research 94: 7293-7320.

Smalley, R., Jr., Kendrick, E., Bevis, M., Dalziel, I., Taylor, F., Lauría, E., Barriga, R., Casassa, G., Olivero, E., Piana,. E. 2003. Geodetic determination of relative plate motion and crustal deformation across the Scotia-South America plate boudary in eastern Tierra del Fuego. Geochemistry Geophysics Geosystems 4(9) 1070, doi:10.1029/2002GC000446

Trifunac, M.D., Brady, A.G. 1975. On the correlation of seismic intensity scales with the peaks of the recorded ground motion. Bulletin, Seismological Society of America 65:103-145.

Winslow, M. 1982. The structural evolution of the Magallanes Basin and neotectonics of the southernmost Andes. In Craddock, C. (ed.) Antarctic Geoscience, University of Wisconsin: 143-154.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 17-20

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Cambrian paleogeography at the western Gondwana margin: U-Pb ages and provenance areas of detrital zircons of the Mesón Group (Upper Cambrian), Northwest Argentina

Christopher J. Adams1, Hubert Miller

2, Guillermo F. Aceñolaza

3, & Alejandro J. Toselli

3

1 GNS Science, PO Box, Lower Hutt , New Zealand (argon@gns.cri.nz)

2 LMU, Department für Geo- und Umweltwissenschaften, Luisenstr. 37, 80333 München, Germany

(h.miller@lmu.de) 3 Universidad de Tucumán, Facultad de Ciencias, Miguel Lillo 205, 4000 S. M. de Tucumán, Argentina

(insugeo@csnat.edu.ar)

KEYWORDS : Argentina, Cambrian, Mesón Group, U-Pb, zircons

Introduction

In northwest Argentina, sedimentation of the Puncoviscana Formation (uppermost Neoproterozoic-Early

Cambrian) finished with folding, metamorphism, and granitoid magmatism of the Pampean Orogeny in mid-

Cambrian times. Above a pronounced angular unconformity, the turbidites of the Puncoviscana Fm. are overlain

by siliciclastic sedimentary rocks of mostly sandstone, partly conglomerate, siltstone and mudstone grain size,

the Mesón Group that is divided into 3 formations: Lizoite, Campanario, and Chalhualmayoc. The Mesón Group

is the basal unit for the sedimentation of the Famatinian (Ordovician-Devonian) orogenic cycle in northwest

Argentina. The siliciclastic rocks of the Mesón Group were deposited in shallow, coast-near tide-dominated

environments in the form of sand bars (Sánchez & Salfity, 1999, Aceñolaza, 2003, 2005).

Generally, the age of the Mesón Group has been considered Cambrian. On paleontological evidence, Sánchez

& Salfity (1999) and Aceñolaza (2003, 2005) restricted the age to “Middle to Upper Cambrian”. The presence of

late Early Cambrian zircons in part of the underlying Puncoviscana Formation (Adams et al., 2008), and the

early to mid Cambrian zircon ages of the Santa Rosa de Tastil and Cañaní granitoids intruding the Puncoviscana

Formation (513 Ma: Adams, oral com.; 514 - 519 - 536 Ma: Bachmann et al., 1991), indicate that sedimentation

of the Mesón Group did not begin before the Middle Cambrian. The Mesón Group as a lithological unit termi-

nates before the Ordovician, whereas quite similar siliciclastic facies continues through the

Cambrian/Ordovician boundary into the lowermost part of the Tremadocian Santa Rosita Fm. (Aceñolaza,

2005).

In order to define the geotectonic position of the Mesón Group within the Gondwana Pacific margin, its

relation to the underlying Puncoviscana Formation and the Pampean and Famatinian orogeny (Pankhurst &

Rapela, 1998), and to recognize the relation of the provenance area of its sediments to those of the Puncoviscana

Formation, samples have been taken from outcrops close to the Puncoviscana Formation. The aim of this work

was to know,

• if the provenance areas of sediments of the Gondwana margin have changed since deposition of the

Puncoviscana Formation and the Pampean orogeny,

• if erosion of the Puncoviscana Formation and its metamorphic equivalents has much contributed to the

sediments of the Mesón Group, or

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 17-20

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• if there was an important new input from the Brazilian shield, similar to that of the time of deposition of

the Puncoviscana Formation, or from anywhere else.

Figure 1. Left: Occurrence of the Mesón Group (Upper Cambrian) in northwest Argentina. A = Sample N° JJ2A, B = Sample N° SAL1, C = Sample N° PMXX3. Right: Frequency diagrams of U-Pb ages of detrital zircons. For discussion see text. Note the numerous grains from 2200 to 2000 Ma, nearly absent in the underlying Puncoviscana Formation.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 17-20

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Results

JJ2A is an angular clast from local rock debris slopes of the Mesón Group, close to sample N° JJ-2 of the

Puncoviscana Formation on the old road from Jujuy to Salta (Adams et al., 2008) (Fig. 1, A). It shows a

youngest peak at 538 to 519, close to the youngest peak of the close lying sample of the Puncoviscana sample

JJ-2 (555 to 530 Ma). Other minor peaks are around 700 to 600 Ma and from 1000 to 900 Ma. The provenance

of 6 zircon grains between 2200 and 1980 Ma will be particularly considered later.

PMXX3 is a sample from outcrop in the upper part of the Lizoite Fm., in the Quebrada de Humahuaca, north

of Tumbaya, Jujuy province (Fig. 1, C). A pronounced peak from 700 to 560 Ma is noticeable. From 2200 to

1980 Ma eight grains are present.

SAL 1 (La Pedrera) is from the road from Salta to La Quesera, south of the city of Salta (Fig. 1, B) immedi-

ately above the angular unconformity of the Pampean Orogeny. It shows a sharply pronounced peak at 502 Ma,

a broad peak at 640 to 580 Ma, and 6 grains from 2200 to 2000 Ma.

Discussion

The maximum age of the Mesón Group is defined by the age of the youngest parts of the underlying Punco-

viscana Formation (523 Ma: Adams et al., 2008), and the age of the youngest granitoids beneath (Bachmann et

al., 1991: 514 Ma, and Adams, oral comm.: 513 Ma). The upper limit is defined by fossil evidence of uppermost

Cambrian age within the lowermost sector of the overlying Santa Rosita Fm. The youngest peak of detrital zir-

con grains in the Mesón Group at 502 Ma (Fig. 1) corresponds to the Middle/Late Cambrian boundary. We think

that these young zircons are the product of volcanic activity at the beginning of the Famatinian magmatic

activity in the neighborhood (Loewy et al., 2004, Sims et al., 1998, Saavedra et al., 1998). An older peak of 538

to 519 Ma is recognized in sample JJ2A (Fig. 1, A). The ages resemble very much the youngest ages from the

Puncoviscana Fm. In all three samples a peak around c. 600 Ma is prominent, similar to most of the

Puncoviscana Fm. samples (Adams et al., 2008). In two of the samples, a peak between 1000 and 900 Ma is also

present. It is nearly identical with a common peak of the Puncoviscana Fm. that appears mostly between about

1100 and 1000 Ma (Adams et al., 2008). Other, Mesoproterozic, ages occurring sometimes in the Puncoviscana

Fm., are not present in the Mesón Group samples. On the contrary, within all Mesón Group samples, distinctive

ages of 2200 to 2000 Ma occur.

Now the question is: Are the sediments of the Mesón Group mostly composed of recycled material of the

underlying Puncoviscana Formation, or, are both lithological units composed of material proceeding from the

same areas? For the Early Cambrian grains of sample JJ2A the youngest parts of the Puncoviscana Fm. are the

most suitable provenance areas. For the late mid to late Cambrian zircons of sample SAL 1, a provenance from

an early Famatinian volcanism is probable. For the few late and early Neoproterozoic grains, recycling of rocks

of the Puncoviscana Fm. is just as possible as an original provenance from the Neoproterozoic Brasiliano

orogen and from the neighboring Sunsás orogen, respectively.

The explicit occurrence of early Paleoproterozoic zircons sharply limited to the time span of 2200 to 2000 Ma

is unexpected, but significant for all samples. Rapela el al. (2007) recently found such ages in boreholes east of

the Sierra de Córdoba within part of the Río de la Plata craton that also presents such ages in Uruguay and

southeast Brazil. Rapela et al. (2007) suggest a former more northern position of the Río de la Plata craton, with

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 17-20

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translation to its present site by large-scale dextral strike-slip movement. There are no younger Paleoproterozoic

and early Mesoproterozoic rocks in between the Río de la Plata craton and the “Pampean Cycle” orogen of

central and northwest Argentina. This explains the absence of such zircons, except Early Neoproterozoic ones.

These may have been derived from the Puncoviscana Formation by recycling, or from a prolongation of the

West Brazilian Sunsás Orogen beneath the Chaco plain in north Argentina intervening between the Río de la

Plata orogen and the Mesón Group deposition site. Although Sánchez & Salfity (1999) record a sediment

transport to the Mesón Group from mostly western sources, this is not documented by the zircon grain ages.

Loewy et al., 2004) have shown that early Proterozoic magmatic and metamorphic rocks of the Arequipa-

Antofalla terrane in the west of the Mesón Group deposition centres are defined by ages from 2.0 to 1.8 Ga,

whilst the characteristic Paleoproterozoic ages of the Mesón Group detrital zircons are older: 2.2 to 2.0 Ga).

Conclusions

At the Pacific Gondwana margin, after the prominent Pampean orogeny in Middle Cambrian, a shallow water

through developed in NW Argentina. Material came partly from the underlying Puncoviscana Formation and/or

its metamorphic equivalents, but continuous provenance from the Brazilian Shield cannot be excluded. In any

case, frequent grains restricted to 2200 to 2000 Ma are considered to have derived from the Río de la Plata

craton of southeast Brazil and Uruguay, and its extension to the area east of the Sierra de Córdoba (Rapela et al.,

2007).

References Aceñolaza, G.F., 2003 — The Cambrian System in Northwestern Argentina: stratigraphical and palaeontological framework.

Geologica Acta, 1: 23-39. Aceñolaza, G.F., 2005 — The Cambrian System in Northwestern Argentina: stratigraphical and palaeontological framework.

Reply. Geologica Acta, 3 (1): 73-77. Adams, Ch., Miller, H., Toselli, A.J., 2008 — The Puncoviscana Formation of northwest Argentina: U-Pb geochronology of

detrital zircons and Rb-Sr metamorphic ages and their bearing on its stratigraphic age, sediment provenance and tectonic setting. Neues Jahrbuch für Geologie und Paläontologie, in press.

Bachmann, G., Grauert, B., Kramm, U., Lork, A., Miller, H., 1991 — El magmatismo del Cámbrico Medio/Cámbrico Superior en el basamento del Noroeste Argentino: Intrusivos de Santa Rosa de Tastil y Cañaní. Actas X. Congreso Geológico Argentino, Tucumán, 4: 125-127.

Loewy, S.L., Connelly, J.N., Dalziel, I.W.D., 2004 — An orphaned basement block: The Arequipa-Antofalla Basement of the central Andean margin of South America. Geological Society of America Bulletin, 116: 171-187; doi: 10.1130/B25226.1.

Pankhurst, R.J., Rapela, C.W., 1998 — The proto-Andean margin of Gondwana: an introduction. In: Pankhurst, R.J., Rapela, C.W. (eds): The Proto-Andean Margin of Gondwana. Geological Society of London, Special Publications 142: 1-9.

Rapela, C.W., Pankhurst, R.J., Casquet, C., Fanning, C.M., Baldo, E.G., González-Casado, J.M., Galindo, C., Dahlquist, J., 2007 — The Río de la Plata craton and the assembly of SW Gondwana. Earth-Science Reviews, 83: 49-82.

Saavedra, J., Toselli, A., Rossi, J., Pellitero, E., Durand, F., 1998 — The early Paleozoic magmatic record of the Famatina System: a review. In: Pankhurst, R.J., Rapela, C.W. (eds): The Proto-Andean Margin of Gondwana. Geological Society of London, Special Publications, 142: 283-295.

Sánchez, M.C., Salfity, J.A., 1999 — La cuenca cámbrica del Grupo Mesón en el Noroeste Argentino: desarrollo estratigráfico y paleogeográfico. Acta Geológica Hispánica, 34: 123-139.

Sims, J.P., Ireland, T.R., Camacho, A., Lyons, P., Pieters, P.E., Skirrow, R.G., Stuart-Smith, P.G., Miró, R., 1998 — U-Pb, Th-Pb, and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin. In: Pankhurst, R.J., Rapela, C.W. (eds): The Proto-Andean Margin of Gondwana. Geological Society of London, Special Publications, 142: 259-281.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 21-24

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Assessment of erosion rate modifications during the Neogene incision in the Semiarid Andes (Northern Chile) using the Black Top Hat function applied to a Digital Elevation Model

Germán Aguilar1, José Darrozes

2, Eric Maire

2, & Rodrigo Riquelme

1

1 Departamento de Ciencias Geológicas, Fac. de Ingenieria y Ciencias Geológicas, Universidad Católica del

Norte, Avenida Angamos 0610, Antofagasta, Chile (gaguilar@ucn.cl, rriquelme@ucn.cl) 2 Laboratoire desMécanismes de Transfert en Géologie, Universite Paul Sabatier, 14 avenue Eduard Belin,

31400 Toulouse, France (darrozes@lmtg.obs-mip.fr, maire@lmtg.obs-mip.fr)

KEYWORDS : semiarid Andes, erosion rates, glacial valley, Black Top Hat (BTH) function

Introduction and geomorphologic settings

We studied a tributary valley of the Huasco Basin (>2700 m a.s.l.; 29°S), called Potrerillo River (Fig. 1). The

valley extends 30 km in direction east-west, including an area of ~2000 Km2. The geomorphic features of the

studied area point out a glacial landscape: U shape, circus, arêtes, truncated spurs and moraines. The Valley is

flanked by hung pedimentation surface product of the successive uplift and erosional events during the Neogene.

Bissig et al. (2002), by Ar-Ar age and correlation between magmatic rocks and alteration and mineralization

events, defined 3 pedimentation surfaces: Frontera-Deidad (17-15 Myr, 4700-5200 m a.s.l.), Azufrera-Torta (14-

12.5 Myr; 4300-4600 m a.s.l.) y Los Ríos (6-10 Myr, 3800-4250 m a.s.l.). Then, the Semiarid Andes present the

opportunity to read geomorphologic registries of a previous stage of pedimentation and studied the posterior

glacio-fluvial incision of the last 6 million years.

Figure 1: Localization of the Potrerillo Watershed in the Semiarid Andes. Hill shade image with the geomorphologic features, the main faults, and the location of 40Ar-39Ar age. Also, it exhibits the location of the cross-sections of the Figure 3.

We postulate that the incision before 6 Myr is small in comparison with the later incision related to the

beginning of the glacial erosion. Also, that the glacial landscape would have been worked from the moment in

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 21-24

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that the mountain chain raise up to the necessary height/elevation to prevent the flow of the humid fronts from

the pacific ocean (Westerly). This elevation would have been reached ~6 Myr B.P., which initiated the beginning

of the ice accumulation linked to the glacial process in the Semiarid Central Andes. The orographical control of

the rainfalls is an important factor to understand the fluvial dynamics in the Semiarid Andes and the landscape

evolution during the last 6 Myr. Also, the semiarid condition and the differential erosional processes between

valleys and surface have favored the preservation of pediment during the last 6 Myr.

This work synthesizes some of the preliminary results of a study, which identified and quantified the erosion

rates of different temporal scales in the Semiarid Central Andes. The aim tries to understand the impact of the

glacial activity in the long-term and short-term denudation of the mountain. To check our hypothesis, we

quantifie the incised volume of the valleys before and after the 6 Myr. In this goal, we applie the Black Top Hat

(BTH) function to a digital elevation model (DEM) to charactize change in the erosion rates during different

neogene incision stages in different valleys. Finally, we discuss the impact of the orographical control of the

rainfalls and glacial erosion in the landscape evolution and the denudation of the Semiarid Central Andes.

Methodology and results

The Black Top Hat Transform function is a grey level mathematical morphological function which allows

valley extraction in a 1D signal and 2D image. Efficiency of the Top Hat Transform was demonstrated in the

first time by Meyer (1979) for Cytology applications. In geomorphology, this function was applied to a high-

precision DEM as a relevant tool for estimating incision and the amount of material removed by recent fluvial

erosion like in a Pyrenean watershed (Rodriguez et al., 2002) and long-term neogene denudation rates from the

Central Andes (Riquelme et al., 2008). The mathematical formulation is based on a set of mathematical

morphology concepts (for details see Rodriguez et al., 2002 or Riquelme et al., 2008). This formulation is

provided by specialized image processing software.

The application of the BTH function is based on the selection of the width of valley incised (L, Table 1),

which corresponds to the greatest separation between pedimentation surfaces. We map the pedimentation surface

and the glacial morphology over a Landsat TM+ image. The surfaces were correlated by three surfaces defined

and dated by Bissig et al. (2002). We considered an error of ±100 m due to the loss of precision in the limit of

pedimentation surface (Fig. 2). These errors were affecting the exactitude of the incision volume and erosion

rate, representing an interval between 0.7 to 1 Km3 and 0.1 to 0.3 m/Myr respectively. For the calculation of the

erosion rates we considerate the age of pedimentation surfaces calculated by Bissig et al. (2002),

The first results (Table 1) of the BTH measurement have been obtained in the Potrerillo Valley. During the

pedimentation period, between 17 and 6 Myr BP, third pediment surfaces have been identified. The first one

which corresponds to the period 17 to 15 My have an incision volume close to 12 Km3 for an incision period of 1

My. This incision period began at the end of the pedimentation phase i.e. 15 my ago and the associated erosion

rate is 5.9 ± 0.3 m/Myr. The second pediment which correspond to the period 14 to 12.5 My have a very weak

incision volume of 2.9 ± 0.7 km3. The incision period began 12.5 My ago and lasted 2.5 My. The associated

incision rate was calculated to 0.6 ± 0.2 m/Myr. In contrast, for the last pediment (10-6 My) and the last 6 My of

incision, the incision volume increase drastically and correspond to approximately 130 Km3 and the associated

erosion rates is 10.6 ± 0.1 m/Myr. This rate represent the double that the maximum rate during the previous

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 21-24

23

pedimentation period. This contrast is point out in the incised valley cross-section (Fig. 2), where during the last

6 Myr the incision exceeds the 1000 m, while the incision among during 17 to 6 Myr BP is minor than 200 m.

These first results will enable us to discuss the variations of the rate of incision and to seek its significance as

well in morphoclimatic and tectonic terms.

Incision stage L: Incision width (m) Incised volume (km3) Erosion rate (m/Myr) Incision 3 (6 Myr - pte.) 8152.66 ± 100 127.3 ± 1 10.6 ± 0.1

Pedimentation stage (10-6 Myr) Incision 2 (12,5-10 Myr) 9167.71 ± 100 2.9 ± 0.7 0.6 ± 0.2

Pedimentation stage (14-12.5 Myr) Incision 1 (15-14 Myr) 10228.99 ± 100 11.7 ± 0.7 5.9 ± 0.3

Pedimentation stage (17-15 Myr) Table 1: Quantitative results of the application of BTH function in the Potrerillo Valley. The BTH function considering the width of valley incised (L). The resulting is the volume incised during and the erosion rates during three incision stage. The erosion rates estimation consider the age of pedimentation surface (Bissig et al., 2002).

Figure 2: Cross-section of the Potrerillo Valley generated of combination of DEM and incised volume digital models (BTH function). Topographic cross-sections show the principal geomorphologic structuring element of the Potrerillo Valley. Also, showing the width of valley incised. The incised volume cross-section are showing the contrast between incision stage 3 and the incision stage 1 and 2. For the located of the cross section see figure 1.

Discussion and conclusion

The erosion rates calculated (<11 m/Myr) are near to the calculated ones for other publications in the Arid

Central Andes (eg. Scholl et al., 1970; Alpers and Brimhall, 1988; Riquelme et al., 2008). We calculated erosion

rates of ~6 m/Myr and ~0.6 m/Myr between pedimentation periods (17-6 Myr). The first erosion rate calculated

(15 Myr to 14 Myr BP) is around of the erosion rates published for valleys in the Arid Central Andes (Riquelme

et al., 2008) and can be correlated to the beginning of the first period of rapid exhumation defined in the zone by

Cembrano et al. (2003) with apatite fission track (300 m/Myr; 15-10 Myr BP). The second erosion rate between

14 to 12.5 Myr BP implicate a minimal valley incision, more consistent with the erosion rates calculated into

Miocene alluvial fans (Gravas de Atacama) and bedrock surface of the Atacama Desert (<0.1 m/Myr by

cosmogenic nuclides 10Be, 26Al and 21Ne; Nishiizumi et al., 2005) and explain the great preservation of

Azufrera-Torta surface (Bissig et al., 2002).

For the last 6 Myr we calculated an erosion rate of ~11m/Myr. This erosion rate is the double of the more great

calculated between anterior incision periods. Also, the volume of valley incised of 130 Km3 during the last 6

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 21-24

24

Myr contrast with the scantly 15 Km3 between 6 to 10 Myr BP. The beginning of this main incision stage is

correlated to the last rapid exhumation determined by Cembrano et al. (2003) to the 5 Myr BP (200 m/Myr). We

postulate that uplift episode around 6 Myr BP permitted a drastic increase of orographical control of the rainfalls

and the beginning of glacial activity. More precisely, for the numerous mountains which exceed the critic

elevation between 3900-4250 m a.s.l. (Amman et al., 2001) the phenomena of condensation and retention of

pacific wet flows (Westerly) are activated and implicated a major change in the erosive mode.

But which is the impact of the glacial erosion in the incision of the valleys and which is the speed of this

phenomenon? In Sierra Nevada, California, the glacial erosion during the Quaternary glaciations, in basins with

glacial coverage of close to a quarter of his area, is between 1250 and 200 m/Myr (Brocklehurst and Whipple,

2002). Considering these erosion rates, with only in a couple of glaciations of 5000 years it might have

excavated the volume of incision of the last 6 Ma of the Potrerillo Valley. Furthermore, only two glacial cycles

would be enough to shape the morphology of the valleys in the high mountain chain of the semiarid Andean. So,

is possible that during the beginning of glacial activity (6-5 Myr BP) the incised valleys is similar to the present

day incision and only the glacial erosion can explain the rapid exhumation calculated by Cembrano et al. (2003).

We presume that the incision of the valleys in The Semiarid Central Andes is product of the intensification of

the orographical control of the rainfalls and the beginning of the glacial and paraglacial erosion to ~6 Myr ago.

Also, the beginning of glacial and paraglacial erosion is an important factors in the exhumation of The Semiarid

Central Andes. To validate these hypotheses it will be necessary to quantify the erosion rates during the last 6

Myr in others glacial and non glacial valleys of the Central Andes and to confront the erosion rates calculated

from the volume measurement of the glacial and paraglacial deposits.

Acknowledgments. We thank the ECOS-CONICYT and the INNOVA-CORFO scientific programs, and Dr. J. Martinod, Dr. G. Gonzalez, Dr. M. Mardonez, Dr. J. Cembrano and Dr. T. Bissig for the many valuable discussions. References Alpers, C.N., Brimhall, G.H., 1988. Middle Miocene climatic change in the Atacama Desert, northern Chile: Evidence from

supergene mineralization at La Escondida. Geological Society of America Bulletin, 100, 1640-1656. Amman, C.; Jenny, B.; Kammer, K.; Messerli, B. 2001. Late Quaternary response to humidity changes in the arid Andes of

Chile (18-29ºS). Palae. Palae. Palae. 172: 313-326. Bissig, T.; Clark, A.H.; Lee, J.K.W.; Hodgson, C.J. 2002. Miocene landscape evolution in the Chilean flat-slab transect:

uplift history and geomorphologic controls on epithermal processes in the El Indio-Pascua Au (–Ag, Cu) belt. Econ Geol 97:971–996.

Brocklehurst and Whipple, 2002. Glacial erosion and relief production in the Eastern Sierra Nevada, California. Geomorphology 42, 1-24.

Cembrano, J., Zentilli, M., Grist, A., Yañez, G. 2003. Nuevas edades de trazas de fisión para Chile Central (30°-34°S): Implicancias en el alzamiento y exhumación de los Andes desde el Cretácico. 10° Congreso Geológico Chileno, Universidad de Concepción-Chile.

Meyer, F. (1979). Cytologie quantitative et morphologie mathématique, Thèse de docteur ingénieur thesis, Ecole des Mines, Paris, (unpublished).

Nishiizumi, K., M.W., Caffe, R.C., Finkell, G., Brimhall, T., Mote, 2005. Remnants of a fossil alluvial fan landscape of Miocene age in the Atacama Desert of northern Chile using cosmogenic nuclide exposure age dating. Earth and Planetary Science Letters, 237, 3-4, 499-507.

Riquelme, R., Darrozes, J., Maire, E., Hérail, G., Soula, J.C. 2008. Long-term denudation rates from the Central Andes (Chile) estimated from a Digital Elevation Model using the Black Top Hat function and Inverse Distance Weighting: implications for the Neogene climate of the Atacama Desert. Rev. geol. Chilena.

Rodríguez, F., Maire, E., Courjault-Radé, P., Darrozes, J. 2002. The Black Top Hat Function applied to a DEM: A tool estimate recent incision in a mountain watershed (Estiber Watershed, Central Pyrenees). Geophysical Research Letters, vol. 29, No. 0.

Scholl, D. W., Christensen M. N., Von Huene R., Marlow M. S., 1970. Peru-Chile trench sediments and sea floor spreading. Geological Society of America Bulletin, 81, 1339-1360.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 25-28

25

Preliminary results of a geochemical survey at Lastarria volcano (Northern Chile): Magmatic vs. hydrothermal contributions

F. Aguilera1, F. Tassi

2, O. Vaselli

2,3, E. Medina

4, & T. Darrah

5

1 Programa de Doctorado en Ciencias mención Geología, Universidad Católica del Norte, Avenida Angamos

0610, Antofagasta, Chile (faguilera@ucn.cl) 2 Department of Earth Sciences, University of Florence, Via La Pira 4, 50121, Florence, Italy (franco.tassi@unifi.it)

3 CNR-IGG Institute of Geosciences and Earth Resources, Via La Pira 4, 50121, Florence, Italy

(orlando.vaselli@unifi.it) 4 Departamento de Ciencias Geológicas, Universidad Católica del Norte, Avenida Angamos 0610, Antofagasta,

Chile (emedina@ucn.cl) 5 Environmental and Earth Sciences Department, Rochester University, Rochester, USA (tdarrah@gmail.com)

KEYWORDS : Lastarria volcano, fumarolic gases, fluid geochemistry, isotope geochemistry, crustal process

Introduction

InSAR images time series (1992-2006) and GPS measurements (Pritchard and Simons, 2002; 2004; Froger et

al. 2007) have shown that the Lastarria-Cordón del Azufre volcanic complex (northern Chile) was interested by

severe ground deformation, probably initiated in early 1998. It has to be noted that this zone was not previously

recognized as active, with the exception of Lastarria volcano that has been characterized by a permanent

fumarolic activity since the beginning of the twentieth century (Casertano, 1963, González-Ferrán, 1995). To

explain the origin of this phenomenon several processes were suggested: i) injection of magma from depth,

possibly causing melting of crustal rocks, ii) uprising of hydrothermal fluid, and iii) rock volume variations

caused by phase changes related to the evolution of the pre-caldera silicic system (Pritchard and Simons, 2002;

2004; Froger et al. 2007; Ruch et al., 2008). In this study, we present the chemical and isotopic features of fluids

discharged from the fumaroles of Lastarria volcano collected during a geochemical survey carried out in May

2006. The main aim is to investigate the origin of the fumarolic fluids and their relation with the tectonic setting

of the system.

Geological and tectonic setting

Lastarria volcano, located in the southern part of the Central Andean Volcanic Zone (CAVZ), is an andesitic-

to-dacitic stratovolcano that forms part of a complex polygenetic structure. According to Naranjo (1986; 1992),

the volcanic complex is formed by: 1) the Negriales lava field (or Big Joe), situated SW of the main volcanic

structure, composed by andesitic-to-dacitic lava flow successions that represent the oldest structure of the

complex (K-Ar dating between 0.6±0.3 and <0.3 Ma; Naranjo, 1988; Naranjo and Cornejo, 1992); 2) the

southern Spur volcanic edifice, located NE of Negriales; 3) the presently active Lastarria volcano sensu stricto,

constituting the main and youngest structure of the system and formed by 5 NW-SE aligned nested craters

(Fig. 1). The permanent fumarolic activity at Lastarria is mainly from i) the north-westernmost craters (from

crater rim and crater bottom) and ii) the NW-SE trending fracture system along the NW external flank of the

Lastarria edifice (Fig. 1). The most prominent regional structure of the CAVZ correspond to NW-SE alignments

(e.g. Calama-Olacapato-El Toro), where several Miocene magmatic centers are aligned (Matteini et al., 2002a,b;

Acocella et al., 2007). The S and SE zones of the complex, hosting large caldera structures (e.g. Cerro Galán,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 25-28

26

Wheelwright) (Ruch et al., 2008) are instead dominated by NNE-SSW tectonic lineaments.

Results

Thermal fluid discharges of Lastarria volcano can be distinguished in two groups: i) low-temperature (LT)

fumaroles (outlet temperature between 80.1 and 96.1 °C) and ii) high-temperature (HT) fumaroles (outlet

temperature between 217 and 278 °C). The LT group is characterized by higher vapor/gas ratios (between 9 and

12.7) with respect to those of HT (<5.5). The chemical composition of the dry gas fraction of the two groups,

dominated by the presence of CO2 (928,800-990,000 μmol/mol), is also distinct: the LT fumaroles show

relatively low concentrations of N2, HCl, H2 and HF (up to 8,140, 665, 440 and 85 μmol/mol, respectively) and

particularly high H2S concentrations (up to 31,450 μmol/mol). Differently, the HT fumaroles have H2S

concentrations <370 μmol/mol and N2, H2, HCl and HF concentrations up to 30,280, 9,140, 4,200 and

580 μmol/mol. A strong difference is also shown by the concentration of the temperature-dependent CO that is

two orders of magnitude higher in the HT fumaroles (up to 29 μmol/mol). Sulfur dioxide is present in variable

amounts (from 1,400 to 26,000 μmol/mol), while CH4 is comprised between 39 and 69 μmol/mol. The

composition of light hydrocarbons, dominated by compounds pertaining to the alkanes group and characterized

by significant amounts of alkenes, is marked by a very low speciation. The isotopic composition of carbon in the

CO2 (13C) varies between -0.42 and -4.13 ‰ V-PDB, while helium isotopes, expressed as R/Rair values, range

between 4.55 and 5.15.

Figure 1. Map of the Lastarria volcanic complex (northern Chile) and location of the sampled fumaroles.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 25-28

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Discussion and conclusions

The presence of highly acidic compounds clearly indicates that Lastarria fumarolic fluids are affected by

conspicuous contribution from a high-temperature source. Accordingly, as shown in the SO2-10*Ar-H2S ternary

diagram (Fig. 2a), the fumarolic discharges can be considered as the result of a mixing process between two

possible end-members related to 1) a magmatic source and 2) a hydrothermal component. On this basis, a clear

distinction seems to exist between the HT and the LT fumaroles, being the former characterized by significantly

higher contribution from the magmatic end-member. Similar indications can be obtained from the N2/50-CH4*5-

Ne*1000 ternary diagram (Fig. 2b), considering that N2-enrichments are to be related to the addition of fluids

from a magmatic (andesitic) source. In fact, the relative abundances of the non-reactive gas species, i.e. N2, Ar

and He (Fig. 2c), are typical of gas discharges associated with subduction-related andesitic magmatism

(“andesite” field; Giggenbach 1992). It is worthy of noting that the LT fumaroles, especially those in the crater

rim and in the highest portion of the flank fracture (Fig. 1) and characterized by the highest contribution from

hydrothermal fluids, show a clear He enrichment. This could be related to strong fluid-rock interactions

involving granite-type rocks that typically constitute the upper crust of the CAVZ (Lucassen et al., 2001). These

evidences are in agreement with i) the R/Rair values, significantly lower than those measured in other volcanoes

of this sector of the Andean Volcanic Chain (e.g. Lascar volcano) (Tassi et al., 2008), and ii) the 13C-CO2

values, likely produced by the mixing of magmatic and limestone sources, the latter likely constituting at least

part of the sedimentary material involved in the subduction process.

Figure 2.a) SO2-10*Ar-H2S ternary diagram; b) N2/50-5*CH4-1000*Ne ternary diagram, air and ASW (Air Saturated Water) are reported; c) Ar-N2/100-10*He ternary diagram, air and ASW compositions and convergent plate boundaries (“andesite” field) (Giggenbach, 1996) are also reported. Symbols: HT fumaroles (open squares), LT fumaroles in the crater rim (filled circle), LT fumaroles in the NE side of flank fracture (filled diamond), LT fumaroles in the highest portion of flank fracture (open triangles).

On the basis of these results, fluid contribution from a crustal source due to melting processes related to the

evolution of the pre-caldera silicic system, one of the possible mechanisms invoked to explain the observed

ground deformation of this system, may be considered not negligible.

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References Acocella, V. Vezzoli, L., Omarini, R., Matteini, M., & Mazzuoli, R., 2007. Kinematic variations across Eastern Cordillera at

24°S (Central Andes): Tectonic and magmatic implications. Tectonophysics, 434: 81-92 Casertano, L., 1963. General characteristics of active Andean volcanoes and a summary of their activities during recent

centurias. Bull. Seismol. Soc. Am. 53: 1415-1433. Froger, J.L., Remy, D., Bonvalot, S., & Legrand, D., 2006. Dynamic of long term multi-scale inflations at Lastarria- Cordon

del Azufre volcanic complex, central Andes, revealed from ASAR-ENVISAT interferometric data. Earth Planet. Sci. Lett. 255: 148-163.

Giggenbach, W., 1992. Isotopic shifts in waters from geothermal and volcanic systems along margins, and their origin. Earth Planet. Sci. Lett. 113: 495-510.

Giggenbach, W., 1996. Chemical composition of volcanic gases. In Scarpa, R., Tilling, R., (eds): Monitoring and mitigation of Volcano Hazards, Springer – Verlag, Berlin: 222-256.

Gonzalez–Ferrán, O., 1995. Volcanes de Chile. Instituto Geográfico Militar, santiago, 639 p. Lucassen, F., Becchio, R., Harmon, R., Kasemann, S., Franz, G., Trumbull, R., Wilke, H., Romer, R., & Dulski, P., 2001.

Composition and density model of the continental crust at an active continental margin the Central Andes between 21º and 27ºS. Tectonophysics 341: 195-223.

Matteini, M., Mazzuoli, R., Omarini, R., Cas, R., & Maas, R., 2002a. The geochemical variations of the upper Cenozoic volcanism along the Calama-Olocapato-El Toro transversal fault system in central Andes (24°S): petrogenetic and geodynamic implications. Tectonophysics 345: 211-227.

Matteini, M., Mazzuoli, R., Omarini, R., Cas, R., & Maas, R., 2002b. Geodynamical evolution of the central Andes at 24°S as inferred by magma composition along the Calama-Olocapato-El Toro transversal volcanic belt. J. Volcanol. Geotherm. Res. 118: 225-228.

Naranjo, J., 1986. Geology and evolution of the Lastarria volcanic complex, north Chilean Andes. MPh Thesis (Unpublished), The Open University, Milton Keynes, 162 p.

Naranjo, J., 1988. Coladas de azufre en los volcanes Lastarria y Bayo en el norte de Chile: Reología, génesis e importancia en geología planetaria. Rev. Geol. Chile 15: 3-12.

Naranjo, J., 1992. Chemistry and petrological evolution of the Lastarria volcanic complex in north Chilean Andes. Geol. Mag. 129, 723-740.

Naranjo, J., & Cornejo, P., 1992. Hoja Salar de la Isla, escala 1:250.000. Servicio Nacional de Geología y Minería, Nº 72 Pritchard, M., & Simons, M., 2002. A satellite geodetic survey of large-scale deformation of volcanic centres in the central

Andes. Nature 418(6894): 167-171. Pritchard, M., & Simons, M., 2004. An InSAR-based survey of volcanic deformation in the southern Andes. Geophys.

Geochem. Geosys. 5(2): 1-42. Ruch, J., Anderssohn, J., Walter, T., & Motagh, M., 2008. Caldera-scale inflation of the Lazufre volcanic area, South

America: evidence from InSAR. J. Volcanol. Geotherm. Res. Submitted Tassi, F., Aguilera, F., Vaselli, O., Medina, E., Tedesco, D., Delgado Huertas, A., Poreda, R., & Kojima, S., 2008. The

magmatic- and hydrothermal-dominated fumarolic system at the Active Crater of Lascar volcano, northern Chile. Bull. Volcanol. in press.

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Towards a geodynamical model for the “middle” Cretaceous very low-grade metamorphism in Central Chile: The geochronological approach

L. Aguirre1, V. Oliveros

2, D. Morata

1, M. Vergara

1, M. Belmar

1, & S. Calderón

1

1 Departamento de Geología, Universidad de Chile, Casilla 13518, Correo 21, Santiago, Chile

2 Departamento Ciencias de la Tierra, Universidad de Concepción, Casilla 160-C, Concepción, Chile

(voliveros@udec.cl)

KEYWORDS : isotopes, very low-grade metamorphism, Mesozoic, Andes, Chile

Introduction

During the Late Jurassic - Early Cretaceous large volumes of volcanic and volcaniclastic rocks were deposited

in central Chile between 25° and 36°S in a 1200 km long ensialic basin characterized by alternating marine and

terrestrial conditions along the pre-Andean continental margin of South America. The Upper Jurassic - Lower

Cretaceous sequences display as two parallel belts in the western and eastern flanks of a Mesozoic synclinorium

(Fig. 1); the western belt conforming the present day Coastal Cordillera and the eastern one along the Andean

Cordillera. The model of a coeval existence of an intra-arc/back-arc pair has been proposed in order to explain

the existence of these two different belts in central Chile (Vergara et al. 1995).

The use of multiple geochronological methods applied to different minerals has proved to be a useful tool in

trying to unravel the evolution and origin of the very low-grade metamorphism affecting the Upper Jurassic -

Lower Cretaceous volcano-sedimentary successions in this region, and consequently to reconstruct the

geodynamical setting of the entire arc/intra-arc/back-arc evolution.

Geological setting

The flanks of the Mesozoic synclinorium consist of two belts of homoclinal sequences dipping towards each

other (Fig. 1). Major differences between the two belts are: cumulative thickness (7-14 km in the west vs.

3-7 km in the east), proportion of volcanic rocks relative to sedimentary rocks (around 80% in the west vs. 15%

to 50% in the east), and abundance of silicic magmatic rocks, i.e. ignimbrites and epizonal granitoids (common

in the west but virtually absent in the east). In both flanks the volcano-sedimentary sequences have been affected

by very low-grade metamorphic events akin to prehnite-pumpeyllite facies and responsible for the local

formation of actinolite-epidote metabasites. Typical mineral assemblages include epidote, chlorite, prehnite,

pumpellyite, celadonite, titanite, quartz, K-feldspar and calcite, with minor actinolite. Lower Cretaceous plutonic

rocks intrude the volcano-sedimentary sequences to the west, whereas in the east only Miocene plutons are

present (Fig. 1).

In the Coastal Cordillera the Veta Negra and Lo Prado formations contain the majority of Lower Cretaceous

volcanic rocks outcropping in this area –partly metamorphosed andesitic to basaltic-andesite lava flows, tuffs

and breccias- whereas the Las Chilcas Formation is mainly sedimentary. In the Andean Cordillera,

metamorphosed basic flows are present at the lower middle part of the marine Upper Jurassic-Lower Cretaceous

Lo Valdés Formation which concordantly overlies the Río Damas Formation (Kimmeridgian) (Hallam et al

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 29-32

30

1986), this last being predominantly volcanic with numerous lava flows of andesitic composition. The Colimapu

Formation (upper Lower Cretaceous) is in turn mainly sedimentary with minor volcaniclastic rocks intercalated

in a series of coarse to fine grained red beds (Fig. 1). The Oligocene-Miocene volcanic activity that took place

in the present Andean Cordillera is represented by the Abanico, Coya-Machali and Farellones formations

(Fig. 1). These units have been also affected by very low-grade metamorphism at the zeolite and prehnite-

pumpeyllite facies.

Figure 1. Idealized, schematic geological profile of central Chile between 33ºS and 35º (modified after Levi et al. 1989) versus ages of igneous and metamorphic minerals from the Coastal and Andean cordilleras. Ages are from: Boric & Munizaga 1994, Aguirre et al. 1999, Fuentes et al. 2005, Morata et al. 2006a,b, Oliveros et al. 2008a,b; Belmar et al. submitted, and this work. VLJ: late Jurassic volcanism; VEK: early Cretaceous volcanism; VOM: Oligocene-Miocene volcanism. MLK: late Cretaceous metamorphic event; M?P: Paleocene metamorphic event? MEM: early Miocene metamorphic event; MLM: late Miocene local metamorphic event (after Oliveros et al. 2008a and Belmar et al. submitted).

Age of volcanism and metamorphism

Coastal Cordillera

Fresh plagioclase crystals have been dated from volcanic rocks of the Veta Negra Formation in the Coastal

Cordillera, yielding Ar-Ar ages between 120 to 114.7 Ma and 104 ± 2 Ma (Fig. 1). K-feldspar from amygdales

and sericite replacing plagioclase phenocrysts in those same rocks were also dated by the Ar-Ar method yielding

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 29-32

31

ages between 103.9 ± 2.1 and 93.1 ± 0.3 Ma (Fig. 1; Aguirre et al. 1999, Fuentes et al. 2005, Morata et al.

2006b). K-Ar on celadonite filling amygdules from a lava flow in Las Chilcas Formation gave an age of

98 ± 3 Ma (Morata et al., 2006a). Wilson et al. (2003) reported two generations of K-feldspar, hosted in

amygdales and veins from metamorphosed and hydrothermally altered lava flows of the Lo Prado Fm, with

average Ar-Ar ages of 110.3 ±1.7 Ma and 103.3 ± 1.2 Ma (ages not shown in Fig. 1).

Andean Cordillera

So far, no isotopic geochronological data for the Upper Jurassic- Lower Cretaceous volcanic rocks in the main

Cordillera of central Chile are available; this is mainly due to the pervasive alteration in those rocks that

precludes the posibility of obtaining reliable Ar-Ar ages. The paleontological record indicates a reliable time

interval between the Kimmeridigian and Tithonian for the volcanism (Hallam et al. 1986). In contrast, several K-

Ar, Ar-Ar and U-Pb ages have been obtained for metamorphic minerals present in amygdales, veins,

groundmass and replacing phenocrysts of lava flows, tuff and breccias (Fig. 1). Three groups of radiometric data

have been interpreted as follows: 1) a late Cretaceous metamorphic event recorded by ages ranging between

108 ± 4 and 82 ± 3 Ma (Belmar et al. submitted, Oliveros et al. 2008a); 2) an early Miocene metamorphic event

at c. 22-15 Ma (Belmar et al. submitted), and 3) a late Miocene local metamorphic event at c. 8 Ma (Oliveros et

al. 2008a). A fourth group of ages ranging between 61.3 ± 8.5 and 47.3 ± 7.9 has been obtained by the U-Pb

method applied to metamorphic titanite; it could be interpreted as a Paleocene metamorphic event (Fig, 1) but it

remains to determine the validity of the radiometric data (Oliveros et al. 2008b).

Geodynamical model of the metamorphism

In spite of the apparent diachronism between the volcanic activity represented by the de Veta Negra (Aptian-

Albian?) and the Lo Valdés-Río Damas (Kimmeridgian-Tithonian) formations, in both flanks of the Mesozoic

synclinorium these units underwent regional non-deformative very low-grade metamorphism during the early

late Cretaceous, from ca. 110 to 83 Ma. This metamorphic process cannot be undoubtedly linked to the plutonic

activity in the arc/back-arc pair, since intrusions of this age have been reported only in the western side (Fig. 1).

A more likely process that could account for the increasing P-T conditions is the burial of the volcanic pile.

Moreover, it is possible that the basin subsidence, astenospheric upwelling and crustal attenuation that

characterized the tectonic setting during that time (Aguirre et al. 1999, Morata et al. 2005) reached their peak

during the late Cretaceous, at least in central Chile, leading to the isotopic closure of the dated metamorphic

minerals. The scatter of the radiometric data between c. 110 and c. 83 Ma implies two possibilities for the

development of the metamorphic process. One is the continuous but not homogeneous burial processes that

originated appropriate P-T conditions for the formation of the metamorphic minerals; and the second is the

occurrence of two different metamorphic events, the first during the late Albian-Cenomanian and the second

(probably only in the eastern part) during the Coniacian. The fact that the range of ages includes U-Pb, Ar-Ar

and K-Ar data suggest they represent the closure of the isotopic systems after the crystallization of the

metamorphic minerals rather than resetting of these during the subsequent Cenozoic metamorphic events.

While in the Coastal Cordillera the volcano-sedimentary sequences do not seem to be affected by any other

significant metamorphic event after the Cretaceous, in the east at least two more processes triggered the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 29-32

32

transformation of the volcanic rocks, not only Mesozoic but Oligocene-Miocene as well. Thus, very low-grade

metamorphism also linked to burial occurred at c. 15 to 22 Ma (Belmar et al. submitted).

Finally, the Miocene plutonic intrusions seen in the Andean Cordillera could be responsible for local thermal

peaks that affected the hosting volcanic rocks.

References Aguirre L., Féraud G., Morata D., Vergara M. & Robinson, D. 1999. Time interval between volcanism and burial

metamorphism and rate of basin subsidence in a Cretaceous Andean extensional setting. Tectonophysics 313: 433-447. Belmar, M., Morata, D., Schmidt, S.Th., Mullis, J., Aguirre, L., Vergara, M., Oliveros, V. & Waite, K. Submitted. Mineral

Chemistry and P-T-t conditions of Very Low-Grade Metamorphism in Meso-Cenozoic volcanic and volcaniclastic successions in the Andean Cordillera of Central Chile. Mineralogical Magazine.

Boris, R. & Munizaga, F. 1994. Geocronología Ar-Ar y b-Sr del depósito estratoligado de cobre El Soldado (Chile central). Comunicaciones 45: 135-148. Levi, B., Aguirre, L., Nyström, J.O., Padilla, H., Vergara, M. 1989. Low-grade regional metamorphism in the Mesozoic–

Cenozoic volcanic sequences of the Central Andes. Journal of Metamorphic Geology 7: 487–495. Fuentes F., Féraud G., Aguirre, L. & Morata D. 2005 40Ar/39Ar dating of volcanism and subsequent very low-grade

metamorphism in a subsiding basin: example of the Cretaceous lava series from central Chile. Chemical Geology 214: 157-177.

Hallam, A., Biró-Bagóczky, L. & Perez, E. 1986. Facies analysis of the Lo Valdés Formation (Titonian-Hauterivian) of the high Cordillera of central Chile, and the palaeogeographic evolution of the Andean Basin. Geological Magazine 123: 425-435.

Morata, D.; Aguirre, L.; Féraud, G. and Belmar, M. 2005. Geodynamic implications of the regional very low-grade metamorphism in the Lower Cretaceous of the Coastal Range in central Chile. 6th International Symposium on Andean Geodynamics. ISAG 2005. Barcelona (España), Septiembre 200, 531-534.

Morata, D., Féraud, G., Schärer, U., Aguirre, L., Belmar, M & Cosca, M. 2006a. A new geochronological framework for lower Cretaceous magmatism in the Coastal Range of central Chile. Actes XIth Chilean Geological Congress, Antofagasta, Chile: 509-512.

Morata, D., Belmar, M., Pérez de Arce, C., Arancibia, G., Morales, S., Carrillo-Rosúa, F.J. 2006b. Dating K-rich fine-grained phyllosilicates from mafic lithologies. An approach to the constraining of low-temperature processes in central Andes. 5th South American Symposium on Isotopic Geochemistry. V SSAGI 2006. Punta del Este, Uruguay, Abril 24-27, 25-27.

Oliveros. V., Aguirre, L., Morata, D., Simonetti, A., Vergara, M., Belmar, M. & Calderón, S. 2008a. Geochronology of very low-grade mesozoic andean metabasites. An approach through the K-Ar, 40Ar/39Ar and U-Pb LA-MC-ICP-MS methods. Journal of the Geological Society 165: 579-584.

Oliveros V., Simonetti, A., Morata, D., Aguirre, L., Vergara, M., Belmar, M. & Calderón, S. 2008b. In-situ U-Pb dating of very low-grade metamorphic titanite in Upper Jurassic-Lower Cretaceous volcanic rocks of central Chile, using Laser Ablation-MC-ICP-MS. In: Acts VI South American Isotope Geology Symposium, Bariloche, Argentina.

Vergara, M., Levi, B., Nyström, J. & Cancino, A. 1995. Jurassic and Early Cretaceous island arc volcanism, extension, and subsidence in the Coast Range of central Chile. Geological Society of America Bulletin 107: 1427–1440.

Wilson, N.F.S., Zentilli, M., Reynolds, P.H. & Boric, R. 2003 Age of mineralization by basinal fluids at the El Soldado manto-type copper deposit, Chile: 40Ar/39Ar geochronology of K-feldspar. Chemical Geology 197: 161-173.

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The Pichaihue Limestones (Late Cretaceous) in the Agrio fold and thrust belt, Neuquén Basin, Argentina

Beatriz Aguirre-Urreta, Pablo J. Pazos, Victor A. Ramos, Eduardo G. Ottone, Cecilia Laprida,

& Dario G. Lazo

Departamento de Ciencias Geológicas, Universidad de Buenos Aires, Ciudad Universitaria, 1428 Buenos Aires,

Argentina (aguirre@gl.fcen.uba.ar, pazos@gl.fcen.uba.ar, andes@gl.fcen.uba.ar, ottone@gl.fcen.uba.ar,

chechu@gl.fcen.uba.ar, dlazo@gl.fcen.uba.ar)

KEYWORDS : Patagonia, Andes, Maastrichtian, Atlantic transgression, foreland basin

Introduction

The Andean system along the western margin of Gondwana records the development of a complex series of

forearc, intraarc, and retroarc basins of distinctive evolution (Ramos 1999). One of these basins, the Neuquén

Basin, is located at the foothills of the Andes (32°-40°SL). The outcrops form a narrow belt along the Andes in

the north, covering part of the Chilean and Argentine Principal Cordillera, while south of 36°SL, the basin

expands towards the eastern foreland forming a large embayment. It is a retroarc basin with a complex history

mainly controlled by the changing tectonic setting of western Gondwana. It encompasses a Late Triassic-Early

Cenozoic succession of several thousand meters of sediments accumulated in quite a variety of conditions

(Legarreta & Gulisano 1989, Legarreta & Uliana 1991). It is bounded to the NE by the Sierra Pintada Massif and

to the SE by the Somuncurá Massif while its western margin was the volcanic arc.

Towards the end of the Early Cretaceous, the Neuquén Basin became a foreland basin due to the incipient

uplift of the Andes associated with the formation of the Agrio fold and thrust belt. This process produced the

final withdraw of the Pacific Ocean from the basin, and allowed the first marine Atlantic transgression during

Campanian-Maastrichtian times (Ramos 1999, Ramos & Folguera 2005). The foreland basin was filled with the

synorogenic deposits of the Neuquén Group. A second phase of deformation is related to the Malargüe Group,

which had a depositional system controlled by the flexural subsidence as a result of tectonic loading of the

Principal Cordillera (Tunik 2001, 2003).

We report here the first evidence of the Late Cretaceous transgression, represented by the Pichaihue

Limestones, located to the west of the Andean thrust front which is the boundary of the presently known

outcrops of the Malargüe Group. The observations were made near Pichaihue, a locality situated some 55 km

southwest of Chos Malal, near the village of Colipilli (figure 1).

The first Atlantic transgression

The Upper Cretaceous rocks in the Neuquén Basin are mostly represented by the synorogenic deposits of the

widespread Neuquén Group (Cenomanian-Campanian) which are separated by a stratigraphic discontinuity from

the overlying continental to shallow marine Malargüe Group (Campanian-Paleocene) (Legarreta et al. 1989).

The Malargüe Group corresponds to the first Atlantic transgression into the basin (Weaver 1927, Uliana &

Dellapé 1981). It represents a regional change in the foreland slope associated with an eustatic sea level rise

(Barrio 1990). The sediments of this group are extensively located between the mountain front and the foreland

(fig. 1) and can reach 450 m in thickness, while are very rarely preserved within the fold and thrust belt

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 33-36

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(Legarreta et al. 1989). In the study locality, the informally named Pichaihue Limestones, which are in part

equivalent to the Malargüe Group, overlie and interfinger with the volcanic rocks of the Colipilli Group, dated in

the area as Late Campanian by Zamora Valcarce et al. (2006).

Figure 1. The Neuquén Basin of western Argentina, showing the extension of the Malargüe Group, the present exposures, and the studied Pichaihue Limestones, near Colipilli.

The Pichaihue section

In spite of numerous studies performed in the volcanic and pyroclastic rocks of the area (see Llambías &

Rapela 1989 and Leanza et al. 2006, and references therein) the occurrence of intercalated limestones remained

unknown. The succession exposed in the area is shown in figure 2, and corresponds to a section surveyed

northwest of Cerro León, a Paleogene subvolcanic body emplaced in the volcanic sequence. The limestones

form an extended cap in the eastern flank of the Colipilli syncline.

The mixed sedimentary pyroclastic section indicates a deposition close to a volcanic centre and freshwater to

brackish and shallow marine water incursions, probably partially interconnected, that were finally desiccated and

consequently brecciated. The fossil plant assemblage, including palms but also pycnoxilic wood and cycads, is

similar to the rich Campanian-Maastrichtian fossil assemblage from the Allen Formation of Bajo de Santa Rosa,

Río Negro province, which includes podocarpaceous conifers, cycads and palms together with vertebrate

remains (Ottone 2007). There are aggregates of serpulid tubes and the bivalves are represented by small

burrowing heterodonts. Although serpulids tolerate salinity fluctuations, they are indicative of marine deposits.

The ostracods are environmentally similar (non marine) but taxonomically different to those described by

Bertels (1972) from the Early Maastrichtian of Huantraico.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 33-36

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The age constraints of the Pichaihue Limestones are based on the radiometric age of the underlying pyroclastic

flows (72.83 ± 0.83 Ma, Ar-Ar in plagioclase, Zamora Valcarce et al. 2006), and in the correlation of part of the

fossil contents with the lower units of the Malargüe Group exposed in the Huantraico region (figure 1). Both

facts point out to a Maastrichtian age for the continental, brackish to shallow marine Pichaihue Limestones that

could be related to the first Atlantic incursion of the Neuquén Basin, and correlated with the Saldeño Formation

exposed further north (Tunik 2003).

Figure 2. Stratigraphic section of the Colipilli Group (pars) and the Pichaihue Limestones at Pichaihue.

Tectonic implications

The Colipilli volcanics with their typical calcalkaline rocks represent the volcanic front at these latitudes. The

foreland shifting of the magmatic arc between Early and Late Cretaceous was associated with an important phase

of shortening and uplift (Ramos & Folguera 2005). The Agrio fold-and-thrust belt was developed between

Cenomanian and Campanian times linked to the migration of the arc. As a consequence of that, a foredeep with

an axial trough was formed east of the thrust front. That trough was filled with the volcanic products derived

from the arc and had a rapid subsidence that culminated with the deposition of the Pichaihue Limestones. The

main outcrops of the Malargüe Group are now preserved east of the Miocene Andean thrust front, which is

associated with the uplift of the Chihuidos High during the Miocene (Ramos & Kay 2006) that led to the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 33-36

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probable erosion of the marine deposits in the western central part of the basin as depicted in figure 1. The rapid

subsidence of the axial trough linked to a high sea-level in Maastrichtian times allowed the sedimentation of the

limestones in the upper part of the trough. Subsequent deformation shifted the thrust front to the present location.

Concluding remarks

The finding of continental and brackish to marine Pichaihue Limestones in the Agrio fold-and-thrust belt

indicates that the first marine transgression derived from the Atlantic ocean had a much wider distribution in the

central part of Neuquén, and probably covered a large part of the basin. This change from Pacific transgression

to Atlantic ones that has been traditionally related to a regional tilting of the basin is clearly a consequence of the

thrust loading of the Principal Cordillera during Campanian to Maastrichtian times.

References Barrio, C.A., 1990 Paleogeographic control of the Upper Cretaceous tidal deposits, Neuquén Basin, Argentina. Journal of

South American Earth Sciences 3: 31-49. Bertels, A., 1972. Ostrácodos de agua dulce del miembro inferior de la Formación Huantrai-Co (Maastrichtiano Inferior),

Provincia del Neuquén, República Argentina. Ameghiniana 9: 173-182. Legarreta, L. & Gulisano, C., 1989. Análisis estratigráfico secuencial de la Cuenca Neuquina (Triásico superior-Terciario

inferior). In G. Chebli & L. Spalletti (eds.) Cuencas Sedimentarias Argentinas. Facultad de Ciencias Naturales, Universidad Nacional de Tucumán, Correlación Geológica Serie 6: 221-243.

Legarreta, L. & Uliana, M.A., 1991. Jurassic-Cretaceous marine oscillations and geometry of back-arc basin fill, central Argentine Andes. International Association of Sedimentology, Special Publication 12: 429-450.

Legarreta, L., Kokogian, D.A. & Boggetti, D.A., 1989. Depositional sequences of the Malargüe Group (Upper Cretaceous-lower Tertiary), Neuquén Basin, Argentina. Cretaceous Research 10: 337-356.

Llambias, E.J. & Rapela, C.W., 1989. Las vulcanitas de Colipilli, Neuquén (37ºS) y su relación con otras unidades paleóge-nas de la Cordillera. Revista de la Asociación Geológica Argentina 44: 224-236. Leanza, H.A., Repol, D. Hugo, C.A. & Sruoga, P., 2006. Hoja Geológica 3769-31 Chorriaca, provincia del Neuquén.

Servicio Nacional Geológico Minero, Boletín 354: 1-93. Ottone, E.G., 2007. A new palm trunk from the Upper Cretaceous of Argentina. Ameghiniana 44: 719-725. Ramos, V.A., 1999. Plate tectonic setting of the Andean Cordillera Episodes 22(3): 183-190. Ramos, V.A. & Folguera, A., 2005. Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc

and foreland deformation In G. Veiga, L. Spalletti, D. Howell & E. Schwarz (eds.) The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics. The Geological Society, Special Publication 252: 15-35.

Ramos, V.A. & Kay, S.M., 2006. Ramos, V.A. y S.M. Kay, 2006. Overview of the Tectonic Evolution of the Southern Central Andes of Mendoza and Neuquén (35°- 39°S Latitude). In S.M. Kay & V.A. Ramos (eds.) Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S latitude). Geological Society of America, Special Paper 407: 1-18.

Tunik, M., 2001. La primera ingresión atlántica de la Alta Cordillera de Mendoza. Ph. D. Thesis, Universidad de Buenos Aires (unpublished) 258 pp.

Tunik, M., 2003. Interpretación paleoambiental de los depósitos de la Formación Saldeño (cretácico superior), en la alta Cordillera de Mendoza, Argentina. Revista de la Asociación Geológica Argentina 58: 417-433.

Uliana, M.A. & Dellapé, D.A., 1981. Estratigrafía y evolución paleoambiental de la sucesión Maastrichtiana-eoterciaria del engolfamiento Neuquino (Patagonia Septentrional). 8° Congreso Geológico Argentino (San Luis), Actas 3: 673-711.

Weaver, C., 1927.The Roca Formation in Argentina. American Journal of Science 13(5): 417-434. Zamora Balcarce, G., Zapata, T., Del Pino, D. & Ansa, A. 2006. Strucutral evolution and magmatic characteristics of the

Agrio fold-and-thrust belt. In S.M. Kay & V.A. Ramos (eds.) Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S latitude). Geological Society of America, Special Paper 407: 125-145.

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Paleoseismic investigation on the Boconó fault between Las González and Estanques, Mérida Andes, Venezuela

Miguel J. Alvarado1, Franck A. Audemard

2,*, Jaime Laffaille

3, Reinaldo J. Ollarves

2, & Luz M.

Rodríguez2

1 Universidad de Los Andes, Grupo de Investigación Ciencia de la Tierra TERRA, Mérida, Venezuela

(alvaradom@ula.ve) 2 Fundación Venezolana de Investigaciones sismológicas FUNVISIS, Caracas, Venezuela

(faudemard@funvisis.gob.ve) 3 Universidad de Los Andes, Fundación para la Prevención del Riesgo Sísmico FUNDAPRIS, Mérida, Venezuela

* presenting author

Introduction

The Boconó fault is the largest active structural feature in the Venezuelan Andes, to which most of the main

historical earthquakes in the region have been assigned. This fault runs in a NE-SW direction, roughly along the

Andean chain backbone for about 500 km. Several crustal depressions related to the right-lateral fault activity

have been identified and described as “pull-apart basins”. Las González pull-apart basin (LGPAB) is the major

of these features (Schubert 1982). It is located in the southwest of the Mérida Andes range, between Las

Gonzalez and Estanques towns in Mérida State, Venezuela. A detailed morpho-structural mapping of this zone

was made by Alvarado et al. (2006). They concluded that only the north trace of this pull-apart basin is active.

Furthermore, they identified a small pull-apart basin along this trace that named the “Lagunillas Pull-part basin”

(LGPB). The objective of this work is to understand the seismogenic behavior of the Boconó fault in this area

through the analysis of paleo-earthquakes, based on two trenches named Pantaleta and Quinanoque, excavated

across the north and south trace of LGPB respectively (figure 1).

Figure 1. Sketch that shows the location of the Pantaleta and Quinanoque trenches across the active trace of Boconó fault between La Gonzalez and Estanques.

Paleoseismic Investigation

Trench sites were chosen taking into account the factors and conditions presented by Audemard (2003).

Technical issues on trench development used in this work are widely discussed in McCalpin (1996).

GRAFIC SCALE

Quinanoque trench

Pantaleta trench

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 37-40

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Pantaleta trench

Located across the north trace of LGPB (figure 1), north of Lagunillas town, where a little sagpond generated

by a shutter ridge marks the active trace of the Boconó fault at this specific location. The sedimentary deposits

exposed in this trench comprise from bottom to top: a poor sorted basal conglomerate with clast diameter

between 4cm and 33cm contained within a red-sandy matrix, overlain by a black 30-cm-thick organic-rich sandy

silt that corresponds to the sag-pond.

Sedimentary features found in this stratigraphic sequence denote the presence of an active trace and the

occurrence of several earthquakes. Figure 2 displays a detailed log of the trench walls, on which sampling points

for 14C dating are also reported. Each point is accompanied with its respective 14C radiometric age, that was

calibrated to calendar year (dendrochronologic method) with a 95% accuracy.

Figure 2. Detailed logs of both walls of the Pantaleta Trench, across the northern strand of the Boconó fault, at the Lagunillas pull-apart basin.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 37-40

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Analysis made in this trench shows the presence of several pre-historical seismic events, and particularly of 3

earthquakes that were reported in historical seismic catalogs (Centeno-Graü, 1940, Grases et al., 1999). The

1610 and 1894 earthquakes were reported by Soulas et al. (1987, in Audemard et al., 1999), for which they built

an isoseismal map that shows its macroseismic hypocenter over the Boconó fault, close to the trench. Audemard

(1997) corroborated through trenching investigations carried out farther southwest in the Andes that both events

happened on the Boconó fault, specifically on the strand that extends from the north of the Lagunillas pull-apart

basin to the southwest. In addition, the 1674 earthquake studied by Palme y Altez (2002), which had never been

previously associated with the Boconó fault, also appears to be present on this trench.

Two seismic recurrence patterns have been identified: the first one is roughly 850-650 years (for pre-historical

earthquakes); and the second pattern seems to be about 280 years (for historical earthquakes).

Quinanoque trench

It was excavated on the southern splay of the Boconó fault at the LGPB, slightly to the west of the Lagunillas

town (figure 1). The fault trace in this place is marked by pressure ridges, spring (tree) lines and a big sag pond

(like the actual lake placed in the Lagunillas town). The Urao lake has also been dammed by a shutter ridge that

has been progressively displaced by dextral slip along the southern active fault splay.

The stratigraphic sequence as described from the trench’s wall, from bottom to top comprises a poorly sorted

basal conglomerate with clast diameter between 1 and 25 cm contained into a sandy matrix. There is a moderate

sorted conglomerate above, which represents a colluvial bed, in turn overlain by an organic-rich sandy silt bed

(figure 3).

Figure 3. Logging of the east and west walls of the Quinanoque Trench, cut across the southern strand of the Boconó fault, at the Lagunillas pull-apart basin..

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 37-40

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More evidences of seismic activity are exposed in this trench than in the La Pantaleta one. Several paleo-

earthquakes including the 1674 historical earthquake were interpreted (figure 3). Some sample points have been

labeled as “recent”, meaning that some levels in the sedimentary sequence have been rejuvenated by present-day

14C. These levels are interpreted as either free-sliding surface on which the lagoon sediments moved during

earthquake shaking or open crack or piping that allowed ventilation of buried paleo-soils. The estimated

earthquake recurrence interval for the south strand of the Boconó fault at the LGPB is 400-450 years

approximately.

Acknowledgements

We would like to thanks all those that in one way or another contributed with this investigation. Special thanks go to Prof. Carlos Ferrer, Geol. Reina Aranguren and the students from the School of Geophysics of Universidad de Los Andes (ULA) who helped to prepare trench walls. We are very grateful to the land owners who gave permits to work in their property. We are very thankful to Prof. Raúl Estevez and Christl Palme who honoured us by visiting the trench sites. This research is a contribution to project FONACIT 2001002492, FONACIT-ECOS Nord 2003000090 and FONACIT-2002000478 (Geodinos). Funding was provided by FONACIT 2001002492, FUNVISIS and logistics by FUNVISIS and ULA.

References Alvarado M., Audemard F. A., Laffaille J., Ferrer C. 2006. Cartografía neotectónica de la falla de Boconó entre las

poblaciones de La González y Estanques, Estado Mérida, para fines de identificación de sitios propicios para excavaciones paleosísmicas. Informe Interno de FUNVISIS, 36 pp.

Audemard, F. A., 1997. Holocene and Historical Earthquakes on the Boconó Fault System, Southern Venezuelan Andes: Trench Confirmation. In: Hancock, P. & Michetti, A. (eds.), Paleoseimology: understanding the past earthquakes using Quaternary Geology. Symposium on Paleoseimicity at the XIV INQUA Congress, Berlin, August 1995. Journal of Geodynamics. 24 (1-4): 155-167.

Audemard F. 1999. Trench investigation along Mérida section of the Boconó Fault (central Venezuela Andes),

Venezuela. Tectonophysics 308, 1-21

Audemard, F. 2003. Estudios paleosísmicos por trincheras en Venezuela: métodos, alcances, aplicaciones, limitaciones y perspectivas. Revista Geográfica Venezolana 44(1), 11-46

Centeno Graü, M. (1940) Estudios Sismológicos. Litografía del Comercio, Caracas Grases, J., Altez, R., Lugo, M. 1999. Catálogo de sismos sentidos o destructores. Venezuela. 1530–1998, Academia de

Ciencias Físicas, Matemáticas y Naturales/Facultad de Ingeniería Universidad Central de Venezuela, Editorial Innovación Tecnológica

McCalpin, J.P. (Ed.), 1996. Paleoseismology. Academic Press, London (583 pp.). Palme C., Altez R. 2002. Los terremotos de 1673 y 1674 en los Andes venezolanos. Interciencias 27, 5. Schubert, C., 1982. Cuencas de tracción en los Andes merideños y en las montañas del Caribe, Venezuela. Acata Científica

Venezolana 33, 389-395.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 41-43

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Seismic source study and tectonic implications of the historic 1958 Las Melosas, Central Chile, crustal earthquake

Patricia Alvarado1,2

, Sergio Barrientos3, Mauro Saez

2, Maximiliano Astroza

4, & Susan Beck

5

1 CONICET (e-mail: alvarado@unsj.edu.ar)

2 Departamento de Geofísica y Astronomía, Facultad de Ciencias Exactas, Fisicas y Naturales, Universidad

Nacional de San Juan, Meglioli 1160 (S) Rivadavia, San Juan, Argentina 3 Servicio Sismológico Nacional, Universidad de Chile, Blanco Encalada 2002, Santiago, Chile

4 Departamento de Ingeniería Civil, Universidad de Chile, Blanco Encalada 2002, Santiago, Chile

5 Department of Geosciences, Univ. of Arizona, Gould Simpson Bldg., 1040 E 4th st. (85721) Tucson, AZ, USA

KEYWORDS : crustal seismicity, Andean Cordillera, neotectonics, seismic hazard

Although Chile is recognized as the site of the largest megathrust earthquakes related to the coupling between

the subducting Nazca plate and the overriding South American plate like the 1960 Valdivia earthquake,

infrequent crustal earthquakes within the continental plate can be very damaging. The earthquake on

4 September 1958 that occurred in Las Melosas, Central Chile (-33.826°S and -70.140°W; Engdalh et al., 1998)

represents one of the large damaging intraplate events located in the Andean cordillera crust at about 60 km

away from Santiago. In this study, new estimates of fault orientation, depth and size using teleseismic body-

wave modeling of the 1958 Las Melosas earthquake are presented (Fig. 1). Although global seismic catalogues

(BSSA, 1959) include only one earthquake on 4 September 1958, Lomnitz (1960), Piderit (1961) and Pardo and

Acevedo (1984) have reported the occurrence of more than one seismic event (at least three) separated by a few

minutes and of similar sized-magnitudes. In fact, these authors assigned a 6.9 magnitude for the three sub-

events, but Flores et al. (1960) reported 6.9, 6.7 and 6.8, respectively. These estimations of the seismic

magnitudes are mainly based on historical intensity reports (Lomnitz, 1970).

Our results for the first event in the sequence of earthquakes on 4 September 1958 that occurred in Las

Melosas indicate a focal mechanism solution with fault planes of right-lateral displacement on an east-west fault

and left-lateral displacement on a north-south fault and a focal depth of 8 km produce the best fit to teleseismic

long period P-waveforms (Fig. 1). A seismic moment M0 of 0.227 x 1019 N-m associated with a moment-

magnitude Mw of 6.3 has been estimated, which is 0.4 to 0.7 units larger than the surface-wave magnitude Ms

earlier reported. Although no surface rupture was reported, the displacement along east-west structures like that

one suggested for one of the fault plane in our focal mechanism solution of the 1958 event seems to be an

efficient mechanism to accommodate differences in shortening from north to south along the high Andean

Cordillera (Alvarado et al., 2008). New findings on the 1958 Las Melosas earthquake intensities by Sepúlveda et

al. (2008) allow us to compare them with our study about the seismic source of this crustal event in order to put

more constraints on the seismic hazard to which this zone, and others along the western foothills of the Andes, is

exposed.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 41-43

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Figure 1. Best results for the seismic source study of the 1958 Las Melosas earthquake modified from Alvarado et al. (2008). (a) Example of a vertical component seismogram recorded at 76° epicentral distance by the station TUC (Tucson, U.S.A.) for the sequence of earthquakes on 4 September 1958. Three different seismic events (E1, E2 and E3) are observed in a time window of 4 to 5 minutes. Event E1 with a clear arrival is the only one modeled in this study. (b) Distribution of seismic stations used in this study with respect to the 1958 earthquake epicentral location.(c) Our preferred focal mechanism solution (fault plane_1: strike N20°E, dip 70°to the southeast and rake 30°; fault plane_2: strike N80°W, dip 62°to the north and rake 157°) plotted as a lower-hemisphere projection with dark compressional quadrants and P-wave seismic records. (d) Synthetic and observed amplitude misfit errors as a function of focal depth for the best combination of strike, dip and rake and variable focal depth (from 0 to 20 km). (e) Relocated seismicity between 1995 and 2005 from Barrientos et al. (2004). (f) Map of the MSK seismic intensities recalculated by Sepúlveda et al. (2008).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 41-43

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References Alvarado, P., Barrientos, S., Saez, M., Astroza, M., and Beck, S., 2008, Source study and tectonic implications of the historic

1958 Las Melosas crustal earthquake, Chile, compared to earthquake damage, Physics of the Earth and Planetary Interiors, Special Issue “Earthquakes in subduction zones: a multidisciplinary approach” (in press).

Barrientos, S., Vera, E., Alvarado, P. and Monfret, T., 2004, Crustal seismicity in Central Chile, J. South Am. Earth Sci., 16, 759-768.

BSSA, 1959, Seismological notes, Bulletin of the Seismological Society of America, 49(1): 115-118. Engdahl, E.R., van der Hilst, R.D. and Buland, R., 1998, Global teleseismic earthquake relocation with improved travel times

and procedures. Bull. Seism. Soc. Amer., 88, 722-743. Flores, R., Arias, S., Jenshke, V. and Rosenberg, L.A., 1960, Engineering aspect of the earthquakes in the Maipo Valley,

Chile, in 1958, Proceedings of 2nd World Conference in earthquake Engineering, Japan, Vol. 1, pp 409-431. Lomnitz, C., 1960, A study of the Maipo Valley earthquakes of September 4, 1958, Proc. 2nd World Conf. Earthq. Eng.,

Tokyo-Kyoto, Japan, 1, 501-520. Lomnitz, C., 1970, Casualties and behavior of populations during earthquakes, Bull. Seism. Soc. Amer., 60(4): 1309-1313. Pardo, M. and Acevedo, P., 1984, Mecanismos de foco en la zona de Chile Central, Tralka 2 (3), 279-293. Piderit, E., 1961, Estudios de los sismos del Cajón del Maipo en el año 1958, Memoria para optar al Título de Ingeniero

Civil, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile, Santiago, Chile. Sepúlveda, S., Astroza, M., Kausel, E., Campos, J., Casas, E., Rebolledo, S. and Verdugo, R., 2008, New findings on the

1958 Las Melosas earthquake sequence, Central Chile: implications for seismic hazard related to shallow crustal earthquake in subduction zones, Journal of Earthquake Engineering, 12: 432 – 455. DOI: 10.1080/13632460701512951

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 44-47

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Dendrochronology of the Central Andes of Bolivia

Jaime Argollo1, Claudia Solís

1, & Ricardo Villalba

2

1 Laboratorio de Dendrocronologia, Instituto de Investigaciones Geológicas y del Medio Ambiente, Universidad

Mayor de San Andrés, La Paz, Bolivia 2 Departamento de Dendrocronología e Historia Ambiental, IANIGLA, Mendoza, Argentina

Resumen En este trabajo se analiza la historia dendroclimatológica de Polylepis tarapacana (queñoa), pequeño arbolito que habita el

Altiplano Boliviano y zonas adyacentes de Perú, Chile y Argentina (16-22°S) entre los 4000 y 5200 m de altura. Muestras dendrocronológicas fueron colectadas sobre las laderas de los volcanes: Sajama y Caquella. Hasta el presente, las cronologías varían entre 98 y 705 años de extensión y constituyen los registros dendrocronológicos más altos del mundo.

Con el objeto de establecer los parámetros climáticos que controlan el crecimiento de P. tarapacana, las variaciones interanuales en el crecimiento de los árboles fueron comparadas con registros regionales de precipitación y temperatura. Las funciones de correlación indican que el crecimiento radial de P. tarapacana está regulado por la precipitación durante el verano previo al ciclo de formación del anillo de crecimiento. En los sitios muestreados la precipitación explica aproximadamente el 50% de las variaciones interanuales en el crecimiento. Las temperaturas más elevadas del verano, que aumentan la evapotranspiración y reducen el agua en el suelo, están negativamente correlacionas con el crecimiento. La longevidad que alcanzan estos registros y su fuerte relación con el clima permitirán reconstruir las variaciones de la precipitación en el Altiplano durante los últimos 5-7 siglos.

Abstract

In this paper we analized the dendroclimatological history of Polylepis tarapacana (queñoa), an small tree growing in the Bolivian Altiplano and adjacent areas of Peru, Chile and Argentina (16-22°S) between 4000 and 5200 m elevation. Dendrocronological samples were collected on the slopes of the volcanoes: Sajama and Caquella. Presently, the chronologies range between 98 and 705 years in length, and represent the highest tree-ring records worldwide. In order to determine the climatic variables controlling P. tarapacana growth, interannual variations in tree growth were compared with regional records of precipitation and temperature. Correlation functions indicate that the radial growth of P. tarapacana is influenced by precipitation during the summer previous to the ring formation. In the sampling sites, precipitation explains around 50% of the total variance in growth. Summer temperatures, which increase evapo-transpiration and reduce soil water supply, are negatively correlated with tree growth. These records offer the unique opportunity for reconstructing precipitation variations across the Altiplano during the past 5-7 centuries.

Introduction

At the biogeographical Andean region in Bolivia, Polylepis is distributed in the sub region yungueña and sub

region puneña above 3000 m and up to 5000 m. Nine species and 8 subspecies are distributed in the Bolivian

Andes (Kessler, 1995).

In this study, we evaluate the dendrochronological potentiality of Polylepis tarapacana Philippi, that is widely

distributed in the Cordillera Occidental de Bolivia. This study set down the presence of growth rings in P.

tarapacana and its annual feature. The similitude between annual variations in the width of the rings from trees

from the same site has allowed us the development of two chronologies at the Bolivian Altiplano. Finally, to

establish the climatic variations that control the radial growing, we determinated the relations between monthly

climate variations and the growing of Polylepis tarapacana at the Bolivian Puna using correlation functions.

Polylepis tarapacana forests in Bolivia

Polylepis tarapacana appears as a little 1 to 3 m high tree or bush that grows between 4000 and 5200 m of

elevation on the biogeographical floor of the Puna in Peru, Bolivia, Chile and Argentina (Kessler 1995, Fjeldsa y

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 44-47

45

Kessler, 1996, Braun 1997). This bioclimatic floor extends trough the Altiplano and is delimited for the

Cordillera Oriental and Cordillera Occidental. The superior bound of distribution that reaches this specie

represents the maximum altitude in the world to an arborescent shape. This species grows in arid environments

with a range of annual precipitation between 250 and 500 mm, where frost is common the whole year (Fjeldså y

Kessler, 1996)

P. tarapacana forests are located at the hillsides of extinguished volcanoes. They develop next to the common

vegetation of the dry Puna. This forests contribute to the increasing of the soil water retention, decrease erosion

by regulating the water runoffs and help to the storage of sediments and nutrients. Besides, they are refuge and

food source to many animal species, and facilitate the growing of several plants (Fjeldsa y Kessler 1996). These

forests have been an important resource to the Altiplano population since they provide wood for house

constructions and firewood for peasants’ works tames.

At present time, Polylepis forests are fragmented in small surfaces as a result of the degradation and alteration

processes as a result of product of many centuries of human intervention. In this way, these forests´s animal and

vegetation biodiversity are seriously threatened by the negative effects of human activity.

Methods and materials

Sites of study.

The studied Polylepis tarapacana forests are located at the Western and South region of the Bolivian

Altiplano. The northest site is at 4750 m near to the Sajama Volcano (18°09’S, 69°00’W) and Caquella

(21°30’S, 67°52’W) at 4560 m of altitude. These two sites are inactive volcanic complex of Superior Mioceno to

Pelistoceno age, consisting of lava flows rocks and pyroclastic. These volcanic mass have been affected by the

Superior Pleistocene Age glacier activity (Clapperton et al. 1997). Small P.tarapacana vegetation units develop

on poor soils product of rocks physical weathering, mainly cryogenic, which form talus debris where vegetation

grows. In the same way these trees develop on volcanic rock fracture, as much as on glacier deposits. These

forests portions are established on slopes that vary from 20 to 40 grades sloping

Establishment of chronologies

After the co-dating phase, the width of the rings had been measured with 0.001mm precision, generating

temporal series for each tree. Measurement and co-date quality had been controlled with the support of the

program COFECHA (Holmes 1983). This program used correlation analysis to compare each series with a

master series compound by the rest of the dated samples from a site.

In this way, it is possible to detect missing or false rings in a particular sample. After the dating control, the

chronologies were constructed for each site using the program ARSTAN that eliminates the biologic growing

trends, and minimizes the uncommon growing variations (Cook y Holmes 1986). The biologic trends in the

width of ring series were modelated using lineal regressions or negative exponential graphs. These standardized

series were finally averaged to obtain the media chronology for each site. Therefore, the chronologies constitute

a temporal series representing the radial growing annual variations of Polylepis tarapacana on each sampled

sites.

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Results

Chronologies

The two chronologies developed from the Polylepis tarapacana extend between 98 and 705 years, being the

Sajama and Caquella chronologies the shorter and more extended respectively. The number of trees in each

chronology varies between 19 and 25. Values of medium sensibility (a statistical used to evaluate the interannual

variability in the rings width) are similar to those described to other subtropical species in South America

(Villalba et al.,1987, 1992) and relatively greater than those of tempered and cold zones species (Bonisegna

1992).

There is a clear common sign between both chronologies. A correlation’s matrix between the two records for

the common period 1902-1999, shows that all of them are correlated significantly.

Climate – radial growth relation

Exist great similitude in correlation functions between the radial growth of Polylepis tarapacana and the

interannual variations of precipitation and temperature in different studied sites of the Bolivian Altiplano

(Argollo et, al. 2004).

Discussion and conclusions

Dendrochronological records of South America are principally from template and cold regions of Argentina

and Chile (Boninsegna and Villalba 1996, Villalba 2000). Sub tropical records are fewer. In northeast of

Argentina, chronologies were developed from forest species between 22 and 28ºS (Villalba et al. 1992, 1998). In

contrast with the high latitudes records, sub tropical chronologies are shorter, rarely overcoming 300 years of

extension. Polylepis tarapacana, a characteristic species from the Bolivian Altiplano, give us new regional

perspectives in the tropical dendrochronology field work. Our researches show that some trees can reach more

than 500 years and that it’s possible to co-date death wood, it’s been allowed until now to elaborate chronologies

of more than 7 centuries of extension. P. tarapacana grows in altitudes higher than 4000 m, in some areas it can

reach 5200 m. For this reason, chronologies developed from this specie represent the more elevated

dendrochronological records of the world. The statistics used traditionally to measure the quality of the

dendrochronologic series show that P. tarapacana chronologies are adequate to reconstruct the past climatic and

environmental variations. Because of the similitude in the growth standard between chronologies along the

Altiplano, these records can also be used as reference chronologies to date archeological material.

P. tarapacana presents marked growth rings. However, it’s very important to consider attributes like quality of

polishing, illumination and perpendicularity of the woody plan in relation with the examined area to achieve a

better growth rings definition. The coincidence in the chronologic sequence of wide and tight rings between

woods of the same place and between places along the Altiplano shows us that the P. tarapacana rings are

bounded to an annual growing seasonal cycle. Another confirmation of the annual nature of the rings in P.

tarapacana is the relation of themselves with the annual climate variations. The yearly precipitations in

Altiplano show an extensive wintry period where there is little rain or there isn’t any. This dry period coincide to

the period of lower temperatures in the year and the period of more daily thermical amplitude

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 44-47

47

In general, in the resulting functions the number of months significantly correlated with rings width is bigger

for temperature than for precipitation, even when we should wait more influence from precipitation than from

temperature in the P. tarapacana growth considering the arid conditions.

It’s possible to see that P. tarapacana that grows in the Altiplano is strongly controlled by variability of

summer rains. Centennial dendrochronologic records, with strong climatic indication offer the opportunity to use

themselves to rebuild the past variations of precipitation in the Bolivian Altiplano during the last 500-700 years.

Acknowledgements To the economic support provided by the Interamerican Institute for the Study of Global Change Project (IAI), CRN03

“The Assessment of Present, Past and Future Climate Variability in the Americas from Treeline Environments” and to the IRD for their logistic support.

References Argollo, J., Soliz, C., Villalba, R. 2004. Potencialidad dendrocronologica de Polylepis tarapacana en los Andes Centrales de

Bolivia. Ecologia de Bolivia pg. 4-22 Boninsegna, J.A., y R. Villalba. 1996. Dendroclimatology in the Southern Hemisphere: Review and Prospect. Tree Rings,

Environment and Humanity, editado por J.S. Dean, D.M. Meko, and T.W. Swetnam. Radiocarbon, pp. 127-141 Boninsegna JA (1992) South American dendroclimatological records. In Bradley RS, Jones PD (eds) Climate since A.D.

1500. Routledge, London pp 446-462. Braun, G. 1997. Métodos digitales para monitorear patrones boscosos en un ambiente andino: El ejemplo Olylepis. In

desarrollo sostenible de Ecosistemas de Montaña: manejo de areas frágiles en los Andes (eds M. Liberman-Cruz & Baied), pp 285-294. The United Nations University press, La Paz

Clapperton M. C., Clayton D. J., Benn I. D., Marden J. C., Argollo, J. 1997. Late Quaternary glacier advances and Palaeolake highstands in the Bolivian Altiplano. Quaternary International, 38/39: 49-59.

Cook,E.R.& R.L Holmes 1986.Users manual for program ARSTAN. Chronology Series VI, University of Arizona. Fjedsa, J. y M. Kessler. 1996. Conserving the biological diversity of Polylepis woodlands of the Highland of Peru and

Bolivia. NORDECO. Copenhagen, Dinamarca. Holmes, R:L. 1983. Computer-assisted quality control in tree-ring dating and measurement. Tree-Ring Bulletin 43, 69-78 Kessler, M. 1995. The genus Polylepis (Rosaceae) in Bolivia. Candollea 50. Conservatore et Jardin Botaniques de Geneve.

172 pp. Villalba, R. 2000. Dendroclimatology: a Southern Hemisphere Perspective. En: Paleo- and Neoclimates of the Southern

Hemisphere: the state of the arts. P. Smolka y W. Volkheimer (editores). Springer. Pag. 28-57. Villalba, R., Holmes, R.L., y Boninsegna, J.A. 1992. Spatial patterns of climate and tree-growth anomalies in subtropical

Northwestern Argentina. Journal of Biogeography, 19: 631-649. Villalba, R., Grau, H.R., Bonisegna, J.A., Jacoby, G.C. y Ripalta, A. 1998. Climatic variations in subtropical South America

inferred from upper-elevation tree-ring records. International Journal of Climatology, 18: 1463-1478. Villalba, R.; Boninsegna, J.A. and Ripalta, A. 1987. Climate, site conditions and tree-growth in subtropical northwestern

Argentina. Canadian Journal of Forest Research, 17 (12): 1527-1544.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 48-51

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An Andean mega-thrust synthetic to subduction?: The San Ramón Fault and associated seismic hazard for Santiago (Chile)

R. Armijo1, R. Rauld

2, R. Thiele

3, G. Vargas

3, J. Campos

2, R. Lacassin

1, & E. Kausel

2

1 IPGP- 4, Place Jussieu, 75252 Paris Cedex 05, France (armijo@ipgp.jussieu.fr, lacassin@ipgp.jussieu.fr)

2 Departamento de Geofísica, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile, Blanco

Encalada 2002, Santiago, Chile (rrauld@dgf.uchile.cl, jaime@dgf.uchile.cl, ekausel@dgf.uchile.cl) 3 Departamento de Geología, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile, Plaza Ercilla

803, Santiago, Chile (rthiele@cec.uchile.cl, gvargas@ing.uchile.cl)

KEYWORDS : subduction margin, Andean orogeny, synthetic thrust, seismic hazard

The Andean orogeny is considered the paradigm for mountain belts associated with subduction plate

boundaries (e.g., James, 1971). Yet, no mechanical model can explain satisfactorily the Andean mountain

building process as a result of forces applied at its nearby Subduction Margin, along the western flank of the

South America continent (e.g., Lamb, 2006). Part of the problem arises from a geometric ambiguity that is

readily defined by the large-scale topography (Fig. 1): the Andes mountain belt is a doubly vergent orogen that

has developed a large Back-Thrust Margin at its eastern flank, with opposite (antithetic) vergence to the

Subduction Margin. The tectonic concept of Subduction Margin used here is equivalent to the pro-flank (or pro-

wedge) concept used for collisional belts (e.g., Willett et al., 1993) and is preferred to the magmatic concept of

fore-arc, which has nearly coincident horizontal extent (Fig. 1). Similarly, the notion of Back-Thrust Margin is

used as an equivalent to that of retro-flank (or retro-wedge) in collisional belts.

The doubly vergent structure of the Andes mountain belt is defined by distinct orogenic thrust boundaries at

the East and West Andean Fronts (Fig. 1). While the East Andean Front coincides with the basal thrust of the

Back-Thrust Margin, the orogenic West Andean Front is located at significant distance from the basal mega-

thrust of the Subduction Margin. The western foreland (~200 km wide horizontally) separating the orogenic

West Andean Front from the subduction zone is designated here as the Marginal (or Coastal) Block.

Consequently a fundamental mechanical partitioning occurs across the Subduction Margin and the Marginal

Block, between the subduction interface, a mega-thrust that is responsible of significant short-term strains and

the occurrence of repeated large earthquakes, and the West Andean Front thrust that appears important in regard

to the long-term cumulative deformation and other processes associated with the Andean orogeny. However,

very few specific observations are available at present to describe and to model this fundamental partitioning.

It is generally admitted that the high elevation of the Andes and of the Altiplano Plateau result from crustal

thickening (up to ~70 km thickness), which is associated with significant tectonic shortening (up to ~150-300 km

shortening) and large-scale thrusting of the Andes over the South-American craton, at the Back-Thrust Margin

(e.g., Kley et al., 1999). On the other hand, the role of the Subduction Margin and of the West Andean Front in

the thickening processes is often considered negligible (Isacks, 1988). Yet the Andean Subduction Margin stands

as one of the largest topographic contrasts on Earth (up to ~12 km), substantially larger than its Back-Thrust

counterpart (Fig. 1). The present study is aimed at revising our knowledge of the large-scale tectonics of the

Andes and its interaction with subduction processes. So we specifically deal with the overlooked West Andean

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 48-51

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Front associated with the Subduction Margin and we attempt to reassess its relative importance during the

Andean orogeny.

Figure 1. Topography and very rough geology of the Central Andes. Red box locates Fig. 2. Square marked with S locates Santiago. The main tectonic features are identified on two selected profiles (A and B, traces marked in red at 20°S and 33.5°S). Vertical black arrows indicate the present-day Volcanic Arc. The Subduction Margin (synthetic to subduction) coincides with fore-arc extent. The Sub Andean Belt in profile B is part of the Back-Thrust Margin, (antithetic to subduction). The Principal Cordillera (PC) includes the Aconcagua Fold Thrust Belt (AFTB), both made of volcanic/sedimentary rocks of the Andean Basin (AB) overlying basement of the Frontal Cordillera (FC). The relatively shallow Cuyo Basin (CB) overlies the basement of the Proto-Precordillera (PP). The Marginal Block is formed of Central Depression (CD), Coastal Cordillera (CC) and Continental Margin (CM). Profile B depicts in light colours major crustal features deduced from the geology: Triassic and pre-Triassic continental basement (brown), post-Triassic basins (yellow) and oceanic crust (blue). The deep basin represented in the two profiles (A and B) is the Andean Basin (AB) that is crossed by the trace of the West Andean Front. VP is Valparaíso Basin. Vertical exaggeration in profiles is 10. Map and profiles based on

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 48-51

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topographic data from the NASA Shuttle Radar Topography mission (SRTM) and a global grid bathymetry available at http://topex.ucsd.edu/WWW_html/srtm30_plus.html.

Figure 2. Morphology and structure of the West Andean Front at Santiago. 3D view of DEM with Landsat 7 imagery overlaid. Oblique view to NE. The most frontal San Ramón Fault reaches the surface at the foot of Cerro San Ramón, across the eastern districts of Santiago. Rectangle shows approximate area mapped in Fig. 3, which gives details of the fault trace. To the East of Cerro San Ramón, Farellones Plateau is incised ~2 km by Ríos Molina-Mapocho and Colorado-Maipo, which grade to the Central Depression (Santiago basin). Red line marks trace of our E-W section (at ~33°30’S).

Figure 3. Map, satellite SPOT image and sections describing the San Ramón Fault and its piedmont scarp in the eastern districts of Santiago. Map and SPOT image cover same area, as shown in Fig. 2. Sections tentatively interpreted across the fault (labelled A and B) are located in the map. The San Ramón Fault trace is at the foot of a continuous scarp east of which the piedmont is uplifted and incised by streams. The more incised Cerros Calán, Apoquindo and Los Rulos (to the N) expose an anticline made of Early Quaternary sediments, cored by bedrock of the Abanico formation (section A). The gently sloping piedmont that is uplifted in the central part of the segment (section B) is covered with Midlle-Late Pleistocene alluvium containing lenses of volcanic ash correlated with Pudahuel ignimbrites. The map has been compiled and georeferenced at 1:25.000 scale, from original mapping on a DEM at 1:5.000 scale.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 48-51

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We analysed and revised critically the Geomorphology and the Geology of the Andes covering the region near

Santiago between ~32.5°S and ~34.5°S (Fig. 1) and focusing on morphologically active tectonic features to

assess the seismic hazard associated with the West Andean Front. Santiago nestles in the Central Depression,

which for long has been described as an extensional graben, bounded to the east and west by normal faults

(Brüggen, 1950). In our work, we show that the San Ramón Fault, crossing the eastern outskirts of Santiago, is a

major active fault with many kilometres of thrust slip (Rauld, 2002; Armijo et al., 2006). The West Andean Front

as defined by the San Ramón Fault is precisely where the Quaternary and older sediments of the Central

Depression are overthrusted by the deformed rocks of the Principal Andes Cordillera.

Our study of the San Ramón Fault aims at describing fault scarps at a range of scales (metres to kilometres)

along with uplift of datable morphological surfaces to determine slip rates over a range of ages (103 yrs to

107 yrs). We combine high-resolution air photographs and digital topographic data with a detailed field survey to

describe the morphology of the piedmont and fault scarps across it (Figs. 2 and 3). Large cumulative scarps and

single event scarps can be identified and mapped with good accuracy. Fault parameters (length of segments, fault

dip, and possible fault slip rate) can be discussed with a view to assess seismic hazard. The multi-kilometric-

scale folding of the San Ramón structure during the past tens of Myr can be used to constrain the thrust geometry

to depths down to ~10 km and more.

At the large scale, key tectonic observations were gathered and analysed critically throughout the study region,

to incorporate our observations of the San Ramón Fault into a complete tectonic section of the Andes, from the

Chile Trench to the South-American craton (see location in Figs. 1 and 2). This unifying approach allows us to

set together, strictly to scale, the most prominent Andean tectonic features, specifically the West Andean Front,

crossing the Andean Basin between the Marginal Block and the Principal Andes Cordillera. We discuss the main

results emerging from this study, particularly the true geometry and possible tectonic evolution of this segment

of the Andes, which allow us to reassess the role of the Subduction Margin and to suggest a broad range of

implications that challenge the Andean orogeny paradigm.

References Armijo, R., R. Rauld, R. Thiele, G. Vargas, J. Campos, R. Lacassin, and E. Kausel 2006. Tectonics of the western front of the

Andes and its relation with subduction processes: The San Ramón Fault and associated seismic hazard for Santiago (Chile), in International Conference Montessus de Ballore, 1906 Valparaíso Earthquake Centennial, Santiago, Chile.

Brüggen, J. 1950. Fundamentos de la Geología de Chile, Instituto Geográfico Militar, Santiago. Isacks, B.L. 1988. Uplift of the central Andean plateau and bending of the Bolivian orocline, J. Geophys. Res., 93 (B4),

3211–3231. James, D.E. 1971. Plate tectonic model for the evolution of the Central Andes, Geol. Soc. Am. Bull., 82, 3325-3346. Kley, J., C. Monaldi, and J. Salfity 1999. Along-strike segmentation of the Andean foreland: causes and consequences,

Tectonophysics, 301 (1-2), 75-94. Lamb, S. 2006. Shear stresses on megathrusts: Implications for mountain building behind subduction zones, J. Geophys.

Res., 111, B07401, doi:10.1029/2005JB003916. Rauld, R.A. 2002. Análisis morfoestructural del frente cordillerano: Santiago oriente entre el río Mapocho y Quebrada de

Macul, Memoria para optar al título de Geólogo. Departamento de Geología thesis, Universidad de Chile, Santiago. Willett, S., C. Beaumont, and P. Fullsack 1993. Mechanical model for the tectonics of doubly vergent compressional

orogens, Geology, 21 (4), 371–374.

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Architecture and style of compressive Neogene deformation in the eastern-southeastern border of the Salar de Atacama Basin (22°30’-24°15’S): A structural setting for the active volcanic arc of the Central Andes

Felipe Aron1, Gabriel González

1, Eugenio Veloso

1, & José Cembrano

1,2

1 Universidad Católica del Norte, Av. Angamos #0610, Antofagasta, Chile (felipearon@ucn.cl, ggonzale@ucn.cl,

eveloso@ucn.cl) 2 Central Andes Resources, Callao #3785, Santiago, Chile (jcembrano@ucn.cl)

KEYWORDS : Neogene tectonics, fold and thrust belt, volcanic arc, Salar de Atacama Basin, Central Andes

Introduction

Detailed field and structural mapping integrated with digital elevation models (DEM), analyses of satellite

images and previous works carried in and around the Salar de Atacama basin are used here to assess the nature

of intra-arc and inner forearc deformation of the Central Andes (between 23° and 24°S, Northern Chile), during

the development of the Andean orogen. Special emphasis was given to the Neogene tectonic evolution between

the Precordillera and the Western Cordillera (Figure 1) at the time of arc formation.

The deformation of the inner-forearc and arc of the Central Andes

The main structural style of the study area is given by first-order kilometric scale ~NS and east-vergent thrust

faults. These faults have a listric section, with detachments levels located approximately 8 km below the surface

(Muñoz et al., 2002; Arriagada et al., 2006 and own work). Subsidiary to these main faults, there is a second-

order thin-skinned system (Kuhn, 2002) with similar orientation to the first-order structures. This system has

detachments levels located approximately 2-3 km below the surface (Figure 2).

Field observations and previous published works (Ramírez & Gardeweg, 1982; Charrier & Reutter, 1994;

Wilkes & Görler, 1994) indicate that the structures deform the Oligocene-Miocene San Pedro Formation,

Tambores Formation and Quepe beds, and the 3.2 My Tucúcaro-Patao Ignimbrite

The topographic expression of both first- and second-order faults corresponds to a set of subparallel fault-

propagation-folds and fault-bend-folds, which can be seen in the field as prominent NS trending ridges with

heights between 50 and 400 m. The fold and thrust belt architecture controls the landscape of the Precordillera

and the Salar de Atacama Basin. Furthermore, we found evidence of an 80 km long structure along the active

magmatic arc (Figure 1, 2), so-called Miscanti Fault. This fault represents the easternmost expression of the fold-

and-thrust belt. The ca. 400 meters high structural relief of the Miscanti Fault controls the development of intra-

arc lakes (Miscanti and Miñiques lakes) and the local and spatial extension of andesitic-basaltic lavas erupted

from nearby volcanic centers. The geometry and evolution of the folding due to this structure, was modeled with

the TRISHEAR 4.5TM software which is based on algorithms presented in Allmendinger (1998). Such modeling

indicates that the Miscanti Fault belongs to the first-order system, having a detachment level buried ca. 8 km

below the surface.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 52-55

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The pattern of deformation exhibits an eastward migration during the last 28 My, but a nearly steady EW

orientation of the maximum compressional axis for the same time window (Jolley et al., 1990; Charrier &

Reutter, 1994; Wilkes & Görler, 1994; Jordan et al., 2002; Reutter et al., 2006 and our own work).

Evidence of active tectonics (Niemeyer et al., 1984, Jordan et al., 2002; Reutter et al., 2006, and González et

al., this symposium) indicates a similar deformation regime at least from the Pleistocene to the Holocene.

Figure 1: Simplified structural map of the study area compiled after Niemeyer (1984), Jolley et al. (1990), Charrier & Reutter (1994), Wilkes & Görler (1994), Jordan et al. ( 2002), Muñoz et al. (2002), Arriagada et al. (2006), Reutter et al. (2006), and our own work. Yellow lines are boundaries of main morpho-structural units. Topographic base: SRTM GTOPO90.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 52-55

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Discussion

Preliminary data of vertical offsets obtained from the observed main structures, indicates a decreasing rates of

uplift and shortening since the Early-Neogene up to Present. Using available data published by Niemeyer (1984)

and Jordan et al. (2002) we estimate an uplift rate of ca. 0.4 mm/yr for the time period between the Late-

Miocene and the Early-Pliocene (Quechua Tectonic Phase). In contrast and by using our own field observations,

we estimate an uplift rate of ~0.05 mm/yr for the time window between the Early-Pliocene up to the Present;

hence, decreasing one order of magnitude between the two identified phases of deformation.

Figure 2: Schematic 3D block model showing the proposed architecture and style of deformation for the study area. Geology units compiled after Ramírez & Gardeweg (1982), Niemeyer (1984), Charrier & Reutter (1994), Wilkes & Görler (1994), Breitkreuz (1995), Mpodozis et al. (2005) and field observations. Vertical exaggeration: 3X.

The nature of the link between the kinematics and timing of deformation in this portion of the volcanic arc of

the Central Andes is currently under study, with the aim of assessing a better understanding of the precise

feedbacks between deformation and volcanism in convergent tectonic settings.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 52-55

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Acknowledgments The authors thank the funding for researching to Fondecyt: National Research Funding Competition-1060187, 2006. We

also thank to Richard Allmendinger for making TRISHEAR 4.5TM software available, Erik Jensen for valuable commenting of previous drafts, Pia Avalos for exhausting help in the field and Alejandro Álvarez for technical support in the field. Finally, we acknowledge the capabilities for working to Universidad Católica del Norte (UCN).

References Allmendinger, R. W. 1998. Inverse and forward numerical modeling of fault-propagation folds. Tectonics 17 (4): 640-656. Arriagada, C., Cobbold, P. & Roperch, P. 2006. Salar de Atacama basin: A record of compressional tectonics in the central

Andes since the mid-Cretaceous. Tectonics 25 (TC1008), doi:10.1029/2004TC001770. Breitkreuz, C. 1995. The Late Permian Peine and Cas Formations at the eastern margin of the Salar de Atacama, Northern

Chile: stratigraphy, volcanic facies, and tectonics. Revista Geológica de Chile 22 (1): 3-23. Charrier, R., & Reutter, K-J. 1994. The Purilactis group of Northern Chile: boundary between arc and backarc from Late

Cretaceous to Eocene. In Reutter, K-J., Scheuber, E., & Wigger P. (eds): Tectonics of the Southern Central Andes: Structure and Evolution of an Active Continental Margin. New York, Springer-Verlag: 189-201

Jolley, E. J., Turner, P., Williams, G. D., Hartley, A. J., & Flint, S. 1990. Sedimentological response of an alluvial system to Neogene thrust tectonics, Atacama Desert, northern Chile. Journal of the Geological Society, London 147: 769-784.

Jordan, T., Muñoz, N., Hein, M., Lowestein, T., Godfrey, L. & Yu, J. 2002. Active faulting and folding without topographic expression in an evaporite basin, Chile. Geological Society of America Bulletin 114 (11): 1406-1421.

Kuhn, D. 2002. Fold and thrust belt structures and strike-slip faulting at the SE margin of the Salar de Atacama basin, Chilean Andes. Tectonics 21 (4), 10.1029/2001TC901042.

Mpodozis, C., Arriagada, C., Basso, M., Roperch, P., Cobbold, P., & Reich, M. 2005. Late Mesozoic to Paleogene stratigraphy of the Salar de Atacama Basin, Antofagasta, Northern Chile: Implications for the tectonic evolution of the Central Andes. Tectonophysics 399: 125-154.

Muñoz, M., Charrier, R., Jordan, T. 2002. Interactions between basement and cover during the evolution of the Salar de Atacama basin, northern Chile. Revista Geológica de Chile. V. 29, n° 1: 55-80.

Niemeyer, H. 1984. La megafalla Tucúcaro en el extremo Sur del Salar de Atacama: una Antigua zona de cizalle reactivada en el Cenozoico. Departamento de Geología, Universidad de Chile, Santiago. Comunicaciones 34: 37-45.

Ramírez, C. F., & Gardeweg, M. 1984. Hoja Toconao, Región de Antofagasta, 1:250.000. Carta Geológica de Chile, Servicio Nacional de Geología y Minería. 54: 122 pp.

Reutter, K-J., Charrier, R., Götze, H-J., Schurr, B., Wigger, P., Scheuber, E., Giese, P., Reuther, C-D., Schmidt, S., Rietbrock, A., Chong, G., & Belmonte-Pool, A. 2006. The Salar de Atacama Basin: a Subsiding Block within the Western Edge of the Altiplano-Puna Plateau. In Oncken, O., Chong, G., Franz, G., Giese, P., Götze, H-J., Ramos, V., Strecker, M., Wigger, P. (eds): The Andes Active Subduction Orogeny. Berlin-Heidelberg, Springer-Verlag. 14: 303-325.

Wilkes, E. & Görler, K. 1994. Sedimentary and structural evolution of the Salar de Atacama depression. In Reutter, K-J., Scheuber, E., & Wigger P. (eds): Tectonics of the Southern Central Andes: Structure and Evolution of an Active Continental Margin. New York, Springer-Verlag: 171-187.

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Block rotations in the Puna plateau

César Arriagada1, Pierrick Roperch

2, & Constantino Mpodozis

3

1 Departamento de Geologia, Universidad de Chile, Casilla 13518, Correo 21, Santiago, Chile

(cearriag@cec.uchile.cl) 2 IRD, LMTG & Géosciences Rennes, campus de Beaulieu, 35042 Rennes, France (pierrick.roperch@ird.fr)

3 Antofagasta Minerals, Ahumada 11, Oficina 602, Santiago, Chile (cmpodozis@aminerals.cl)

KEYWORDS : Central Andes, Puna, block rotations, paleomagnetism

Introduction

Paleomagnetic studies along the Chilean forearc show systematic clockwise rotation along the forearc south of

22°S (Taylor et al., 2005; Arriagada et al., 2006). Arriagada et al. (2006) suggested that the rotations within the

forearc are partly driven by ongoing deformation to the east in the Puna plateau. However, rotations within the

forearc appear to occur mainly during the

Paleogene in Chile while the deformation

within the Sierras Pampeanas and the Puna is

mainly Neogene. Here we report

paleomagnetic results at 31 sites from 10

localities (Figure 1) in Tertiary sediments and

in Permian to Triassic red beds of the

Paganzo Group. Many previous

paleomagnetic studies have been carried out

in these red beds to define the Apparent polar

wander path of South America but few

studies indicate that these paleomagnetic

results may record a component of local

tectonic rotations (Geuna and Escosteguy,

2004 ).

Previous works in the Sierras Pampeanas

(Aubry et al., 1996), in the Puna (Coutand et

al., 1999) and in the Altiplano (Roperch et

al., 2000), demonstrated that the sediments

record a low anisotropy of magnetic

susceptibility (AMS) associated with Andean

compression. The orientation of AMS

lineations are always N-NE oriented in NW

Argentina while they are NW oriented in northern Bolivia. Here we report new AMS results that confirm the

previous observations (Figure 1).

Figure 1. Paleomagnetic sampling.in the Puna and Sierras Pampeanas. A to J are sampling localities (green circles). The arrows correspond to the orientation of the AMS lineations. Colour coding corresponds to results in Tertiary rocks (orange); Cretaceous (violet) and Permian to Triassic (magenta). Red circles are results reported by Coutand et al., (1999).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 56-59

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Paleomagnetic results

One specimen per core was progressively thermally demagnetized. Anisotropy of magnetic susceptibility

(AMS) measurements were made on one or two specimens per core.

Oligocene sediments were sampled in the Upper Eocene to Oligocene Quiñoas Formation to the west of the

Salar de Antofalla (Locality A). The upper part of the sequence is dated at 28.9±0.8 Ma (Kraemer et al., 1999).

A characteristic direction of reverse polarity was observed. Although the sampling is not sufficient enough for

magnetostratigraphic purpose, the dominant reverse polarity is in agreement with deposition during the late

Eocene – Early Oligocene period during which the earth magnetic field is mainly of reverse polarity.

Figure 2. Paleomagnetic results near the Salar de Antofalla. a) Equal-area projection of characteristic directions with angle of confidence in in situ and after tilt correction. b) Directions of the principal axes of Anisotropy of magnetic susceptibility. Ellipsoids are mainly oblate (c) and the magnetic foliation (kmin pole of magnetic foliation) is mainly horizontal and controlled by bedding. Magnetic lineations correspond to kmax directions.

The mean direction after tilt correction is (D=187.4; I=41.0 95=6.1). Assuming an early Oligocene age, the

rotation for this site is only 13.3±8.1°. Although these sites are only slightly tilted without evidence of internal

deformation a magnetic fabric is recorded (Figure 2). The magnetic lineation is NS oriented in agreement with

bedding strike for the area and the orientation of the tectonic structures. Thus, the magnetic lineation is likely

associated with compression.

Figure 3. Characteristic directions (green circles) and AMS measurements for (a) locality J and (b) locality F. Equal-area projection of mean-site characteristic directions with angle of confidence after tilt correction and directions of the principal axes of Anisotropy of magnetic susceptibility.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 56-59

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Near Valle Ancho (Locality J), 4 sites were drilled in the Astaburuaga Formation. All sites have a characteristic

magnetization with reverse polarity carried by magnetite (Figure 3a). The mean direction is (D=207.3; I=45.5;

95=5.8 indicating a clockwise rotation of nearly 30°. AMS measurements indicate prolate ellipsoids with well-

defined magnetic lineations (Figure 3a).

To the west of Vinchina (Locality F), 18 samples were drilled in a Tertiary sedimentary section. The thick

sedimentary section is folded and bedding strikes are NE-E oriented. AMS lineations are also NE-E oriented and

the characteristic mean direction for the section is also clockwise rotated (Figure 3b). These results confirm

previous observations (Aubry et al., 1993, Coutand et al., 1999, 2001) of significant clockwise rotations of

structures at various scales.

To the south-west of Antofagasta de la Sierra, 3 sites were drilled in Tertiary red beds (Locality B, Figure 1). It

was not possible to determine a characteristic direction. These sediments record a well-defined magnetic fabric

with a NE magnetic lineation (Figure 1).

Several sites were drilled in Permian to Triassic red beds sediments. Reverse polarity is found in sites drilled in

La Cuesta Formation (Locality H) probably deposited during the long reverse Kiaman interval while normal and

reverse polarities are found in the Talampaya Formation and these sediments are younger than the upper

boundary of the Kiaman (~260Ma)(Localities D & E).

Sites with normal polarity have N-NE magnetic declinations while reverse polarity sites have SE-S magnetic

declinations. Part of this variation is due both to the Permian to Triassic drift of the South American plate and to

the uncertainties in the precise age of the red beds. Uncertainties in different configurations of Pangea with the

possible existence of a large strike slip displacement between Gondwana and Laurasia (Pangea B to Pangea A)

during the Permian impede also the determination of a global apparent polar wander path for Gondwana using

the more numerous paleomagnetic data from Laurasia.

Geuna and Ecosteguy (2004) provide well-defined paleomagnetic results (Figure 4) from Upper Carboniferous

– lower Permian rocks sampled near Locality G where we also sampled one site in Triassic red beds. When

Figure 4. Paleomagnetic results from sites in Permian La Cuesta – Patquia formations and Triassic Talampaya red beds. Blue circles correspond to results from Geuna and Ecosteguy (2004) while the red circle is the expected direction calculated from the Permian pole from Tomezoli et al. (2006)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 56-59

59

compared with the result reported by Tomezzoli et al. (2006) from a locality in La Pampa region, a region not

affected by Andean tectonics, a clockwise rotation of nearly 30° is expected at locality G. During the Permian-

Triassic, the South American plate exhibit a shift in paleolatitude toward a position similar to the present-day

position and this explain mainly the dispersion in inclination. Local tectonic rotations contribute significantly to

the dispersion in declination.

The Permian to Triassic red beds of the Paganzo Group record also an AMS lineation which is NE oriented at

localities D and E (Figure 1). Although AMS ellipsoids are coherent with the pattern observed in Tertiary rocks,

the AMS fabric in the Permian rocks may have been acquired prior to the Andean Tertiary deformation.

Our results from NW Argentina confirm clockwise rotations within the Puna plateau but further work is needed

to better define the timing and the spatial distribution of the rotations.

References Arriagada, C., Roperch, P., Mpodozis C., & Fernández, R. 2006. Paleomagnetism and tectonics of the southern Atacama

Desert region (25-28ºS) Northern Chile. Tectonics, 25: TC4001, doi:10.1029/2005TC001923. Aubry, L., P. Roperch, M. Urreiztieta, E. Rossello, and A. Chauvin, 1996. Paleomagnetic study along the southeastern edge

of the Altiplano-Puna Plateau: Neogene tectonic rotations., Journal of Geophysical Research, 101, 17.883-17.899. Coutand, I., P. Roperch, A. Chauvin, P. Cobbold, And P. Gautier, 1999. Vertical-Axis Rotations Across The Puna Plateau

(Northwestern Argentina) From Paleomagnetic Analysis Of Cretaceous And Cenozoic Rocks, J. Geophys. Res., 104, B10, 22965-22984.

Coutand, I., P.R. Cobbold, M. De Urreiztieta, P. Gautier, A. Chauvin, D. Gapais, E.A. Rossello, and O. Lopez-Gamundi, 2001. Style and history of Andean deformation, Puna plateau, northwestern Argentina, Tectonics, 20, 210-234.

Geuna, S., and L. Ecosteguy, 2004. Paleomagnetism of the Upper Carboniferous Lower Permian transition from Paganzo basin, Argentina, Geophys. J. Int., 157, 1071-1089.

Kraemer, B., D. Adelmann, M. Alten, W. Schnurr, K. Erpenstein, E. Kiefer, P. Van den Bogaard, and K. Gorler, 1999. Incorporation of the Paleogene foreland into the Neogene Puna plateau: The Salar de Antofalla area, NW Argentina, Journal of South American Earth Sciences, 12, 157-182.

Muttoni, G., D. V. Kent, E. Garzanti, P. Brack, N. Abrahamsen and M. Gaetan , 2003, Early Permian Pangea ‘B’ to Late Permian Pangea ‘A’ Earth and Planetary Science Letters, 215, 379-394

Roperch, P., M. Fornari, G. Hérail, and G. Parraguez, 2000. Tectonic rotations within the Bolivian Altiplano: Implications for the geodynamic evolution of the central Andes during the late Tertiary, Journal of Geophysical Research, 105, 795-820.

Roperch, P., Sempere, T., Macedo, O., Arriagada, C., Fornari, M., Tapia, C., García, M. & C. Laj, 2006. Counterclockwise rotation of late Eocene – Oligocene fore-arc deposits in southern Peru and its significance for oroclinal bending in the central Andes, Tectonics, 25: TC3010, doi:10.1029/2005TC001882.

Taylor, G.K., Dashwood, B., and Grocott, J., 2005, Central Andean rotation pattern: Evidence from paleomagnetic rotations of an anomalous domain in the forearc of northern Chile: Geology, v. 33, p. 777-780.

Tomezzoli, R.N., R.N. Melchor and W.D. MacDonald, 2006. Tectonic implications of post-folding magnetizations in the Carapacha basin, La Pampa province, Argentina, Earth, Planets Space, 58, 1235-1246.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 60-63

60

Continental growth through protracted subduction and accretionary processes along Western Gondwana: The case of the Ocloyic Orogeny in southern South America

Ricardo A. Astini, Gilda Collo, & Federico Martina

Laboratorio de Análisis de Cuencas, CICTERRA-Universidad Nacional de Córdoba, Av. Vélez Sársfield 1611, 2º

piso, Of. 7, X5016GCA Córdoba, Argentina (raastini@efn.uncor.edu, gcollo@efn.uncor.edu,

fmartina@efn.uncor.edu)

KEYWORDS : accretionary orogen, Ocloyic orogen, composite orogeny, South America, Ordovician

Introduction

Since the Middle Cambrian and after the consolidation of Gondwana through mayor collisional orogens, the

Proto-Andean region of South America faced an open ocean, resulting in quasi-permanent subduction and

development of a protracted exterior-orogen during the Paleozoic (the Terra Australis orogen, Cawood, 2005).

Within such expanded time interval (~300 m.y.) discrete orogenic features imply development of recurrent

processes related to a variety of transient coupling mechanisms. However, Cenozoic structural complexities and

spatial superposition of different age mountain-building processes along the Andean margin has prevented

finding a simple and universal model to explain the architectural and continuity relationships between temporally

constrained orogenies within the Central Andes. Recent stratigraphically and regionally constrained

geochronological work has allowed great improvement in our understanding and discrimination of distinct

orogenic episodes. Nevertheless, understanding of the across and along-strike variations of any particular

orogeny is still unclear. Focussing within the more well-known southern segment of the Central Andes in

Argentina during the Ordovician may help understanding the complexities within an orogenic cycle.

Probably the most compelling and distinct tectonothermal event including metamorphism, deformation,

magmatism and basins development within the Terra Australis orogen occurred during the Ordovician and is

known as the Ocloyic Orogeny (see Ramos, 1986). This is a composite orogenic episode that in South America

can now be traced from Patagonia into Perú (>4000 km) and that has been defined, half a way, in northwestern

Argentina, on the basis of stratigraphic relationships. In the Central Andean basin (Northewest Agentina, Bolivia

and Perú), the erosive effects of the Late Ordovician glaciation amalgamate with the tectonic effects in the

Ocloyic unconformity, implying ~50 m.y. of Proto-Andean history (approximately ranging from 490 to 440 Ma).

Various names have been used to embrace the crustal processes occurring within this interval, probably the more

commonly used being that of Famatinian Orogeny (Aceñolaza & Toselli, 1973). However, in the light of more

recent geochronological research this last term has become broad.

The complex and composite Ocloyic Orogeny

Within the long-lasting feature of the terra Australis Orogen the Ocloyic Orogeny is defined as complex and

composite. Complex because it embraces different crustal processes across different segments of the western

Gondwana margin as a result of varying geometry and plate tectonic situations (e.g., ocean-continent versus

continent-continent convergence) and composite because it has become clearer that separate building stages

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associated to crustal addition or recycling have happened. What is not clear is how different configurations along

a same active margin influence for the development of metamorphism, magmatism and sedimentary basins, and

how these processes operate when arrival of an exotic terrane, or when other alternative features (e.g., aseismic

ridges or any other buoyant fragment or process) change the coupling relationships along a convergent setting.

Why is there a regional signature with a strong Ordovician imprint along the entire South American plate if

different mechanisms operated and how do they really correlate between each other, even across adjacent

segments? Along-strike segmentation within the Ocloyic Orogen seems to relate to various accretionary

mechanisms. The accretion of Precordillera must have locally increased coupling for at least some time, helping

to modify existing stratigraphies and generating its own signatures. North and south of the Jagüé slip zone ~28°S

(Astini & Dávila, 2004) the retroarc foreland basins have very different configurations and this major across-

strike tectonic boundary related to indentation seems influential on much of the following Proto-Andean and

later Andean history. Similar assertion can be made to its more obscure southern extent (mostly in subsurface)

close to the northern boundary of Patagonia. Some of these questions are trying to be focused by comparing the

segments aligned with the Precordillera and immediately north, where superb stratigraphic and igneous records

are available.

The Famatina-Precordillera segment

Development of an asymmetric long wave-length retroforeland in the Central Andean Basin contrasts with the

narrow retroforeland that existed in the segment aligned with Famatina (Astini & Dávila, 2004). But this

segment do not only differs in the accommodation of very different sedimentary systems or on the rates of

subsidence, but in the igneous, metamorphic and deformational trajectories. Recent improvement in

geochronology and geological mapping at Famatina and adjacent regions of the Sierras Pampeanas allows

suggesting a cyclic crustal behaviour within this segment. Two stages of extension with pervasive volcanic

activity alternate with two major stages of plutonism and crustal thickening, indicated from both igneous and

detrital ages. Plutonism bracketed at ca. 490-480 Ma and at ca. 470-465 Ma. Both the western Cerro Toro and

the eastern Ñuñorco complexes in Famatina (Aceñolaza et al., 1996) seem to record the two age peaks, whereas

the volcano-sedimentary stages are respectively known as the Cerro Tocino (ca. 480-475 Ma) and the Cerro

Morado/Las Planchadas volcanics (ca. 468-460 Ma). The fact that both volcano-sedimentary stages interact with

sea-level indicates important accommodation space, compatible with extension within the upper-plate arc

setting, a major difference with the present Andean style. Extension is in agreement with the complex array of

volcanic and volcaniclastic facies, which locally describe large variations in recorded thickness. Volcanic stages

are largely acidic and bimodal, but locally, like in northern Famatina (e.g., north of Cazadero Grande), more than

50% of the volcanic volume corresponds to basic lavas. This is largely difficult to reconciliate with crustal

recycling and differentiation processes and might need a mantle component. Some juvenile addition is at least

clear in the 2nd volcanic stage and might have also occurred in the 1st stage. By contrast, most geochemistry and

isotopic work in the granite suites indicates crustal recycling as a major source for plutons, compatible with

stages of important crustal thickening (Pankhurst et al., 2000). Maximum crustal thickening seems to have been

acquired in Famatina after the second plutonic stage. This may relate to the tectonic shortening and angular

unconformity recorded between the Famatina and the Cerro Morado Groups (Astini & Dávila, 2004). Following,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 60-63

62

Late Ordovician arkoses (represented in the La Aguadita Formation), recorded east of the former volcanic stages,

yield two peaks within the detrital zircon population, indicating that granites from both plutonic stages where

being exposed to erosion. Recent work by Dahlquist et al. (in press) shows that detrital zircons in Middle

Ordovician sedimentary units in western Famatina also yield some provenance from the former granite suites.

Unroofing and recycling of the early Ocloyic granites, likely accounted through extension, may have helped

triggering rapid exhumation of plutonic suites and their Cambrian hosts (Collo & Astini, in press). This is

consistent with recent dating of low-grade filonites associated to normal shear zones in Famatina that may relate

to orogenic collapse during termination of the Ocloyic Orogeny (Collo et al, in press). This suggestion allows

understanding provenance as a product of limited detrital dispersal into a narrow retroforeland within this

segment.

An opposite polarity peripheral foreland developed atop of the lower-plate Precordillera terrane (Astini et al.,

1995). Such westward foreland series is unique to the Precordillera and has no other counterparts to the north or

south, hence pointing to a strong difference with the rest of the segments implying absence of symmetry in terms

of convergent basin development.

Discussion

Accretionary orogens are basically protracted subduction orogens (Cawood & Buchan, 2007) wherein due to

the time involved and their areal extent they encompass various sedimentary and igneous addition processes.

These processes seem to have occurred in the case of the Ocloyic segment aligned with the Precordillera and

Famatina. Along-strike segmentation is a natural feature within any orogen influencing igneous activity and

basin development. As it is true in the present day Andes along the South American plate, it has apparently also

been true during the Paleozoic, although related to different triggering mechanisms.

Much debate can be generated on weather or not the Precordillera terrane accretion generated a real collisional

orogen or if, on the other hand, given its relatively small size it contributed as a strong transient coupling

mechanism in the light of the protracted history of the Paleozoic accretionary margin. There is strong geologic

evidence to separate the Precordillera and the Gondwana margin previous to Early Ordovician and also strong

evidence to support their continuity after the Late Ordovician. In fact Late Ordovician glaciation overlaps both.

Basin development and stratigraphy to both sides of this major boundary and also magmatism and

metamorphism is, however, largely different showing dissimilar crustal behaviours related to the complex

tectonothermal Ocloyic orogeny. Our studies, particularly those on the upper plate along Famatina, suggest that

during the Ordovician tectonic switching operated along South America, in a similar manner and in a

comparable time framework than what has been interpreted for Australia (Collins, 2002). Changes in the

convergence mode and tectonic switching processes from dominantly advancing (contractile) to dominantly

retreating (extensional) behaviours can be suggested for the upper plate within the segment of Famatina.

Indirectly, some of the effects (e.g., volcanism), have been recorded within the approaching Precordillera terrane

(Astini et al., 2007). According to that interpretation, some time after subduction initiation (during the Late

Cambrian) tectonic switching to extension may lead to recurrent inboard continental growth; that is both,

magmatic and sedimentary additions in the upper plate. These can take the form of intraarc or backarc extension,

like that apparently recorded in both volcano-sedimentary intervals in Famatina. On the other hand, dominantly

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contractile modes may have lead to crustal thickening, batholite intrusions and recycling. According to Collins

(2002), this switching process may occur at relatively short time intervals (~10 Ma), a resolution we are just

starting to approach. Moreover, this mechanism would have contributed to upper–plate effective growth and

recycling, rather than periods of rifting, separation, and reaccretion like suggested by Ramos (2008).

Because such a cyclic behaviour is not strongly recorded along the northern segment, aligned with the Central

Andean Basin, a different driving mechanism may have operated. However, differences in the convergence

mode responsible for the Ocloyic Orogeny would not be enough to prevent the broad correlation between

segments. Deformed metamorphic terranes like, for example, the pre-Ocloyic Puncoviscana Complex in

northwest Argentina or the Ocloyic Negro Peinado complex along Famatina have seldom been considered

“mobile belts”. This pre-plate tectonic concept, regardless its origin, portrays an unusual inboard situation for

location of deformation, metamorphism and magmatism, where polarity predicted by continental growth related

to accretion does not work. Such a relationship has been observed in ancient crust elsewhere, and although not

fully apparent in the upper crustal levels along the Andean modern analogue for subduction orogens, it is well

preserved in the pre-Andean geology. This means that either uplifting/exhumation processes operating in the

Andes are largely different from those in the past and hence, it is not a good analogue for ancient pre-Andean

cases. An alternative is to consider that the present configuration for the type-subduction orogen is still “too

young” to show effects of the various coupling mechanisms operating and triggering mountain building

processes. However, isotopic evidence showing a clear asthenospheric contribution in the present Andean orogen

is unlikely comparable with the signature during the early Ocloyic Orogeny (Cordani, 2006).

References Aceñolaza, F.G. & Toselli, A.J. 1973. Consideraciones estratigráficas y tectónicas sobre el Paleozoico Inferior del Noroeste

Argentino. 2º Congreso Latinoamericano de Geología, 2: 755-783. Aceñolaza, F.G., Miller, H. & Toselli, A.J. 1996. Geología del Sistema de Famatina. In Aceñolaza, F.G., Miller, H., Toselli,

A.J. (Eds.), Geología del Sistema de Famatina. Münchner Geologische Hefte, Reihe A, 19(6): 412p. Astini, R.A. & Dávila, F.M. 2004. Ordovician back arc foreland and Ocloyic thrust belt development on the western

Gondwana margin as a response to Precordillera terrane accretion. Tectonics, 23: TC4008, doi:10.1029/2003TC001620. Astini, R.A., Benedetto, J.L. & Vaccari, N.E.. 1995. The early Paleozoic evolution of the Argentine Precordillera as a

Laurentian rifted, drifted and collided terrane: a geodynamic model. Geological Society of America Bulletin 107: 253–27 Astini, R.A., Collo, G. & Martina, F. 2007. Ordovician K-bentonites in the upper-plate active margin of Western Gondwana,

(Famatina Ranges): stratigraphic and palaeogeographic significance”. Gondwana Research, 11: 311-325. Cawood, P.A. 2005. Terra Australis Orogen: Rodinia breakup and development of the Pacific and Iapetus margins of

Gondwana during the Neoproterozoic and Paleozoic. Earth Sci. Rev. 69: 249– 279. Cawood P.A. & Buchan C. 2007. Linking accretionary orogenesis with supercontinent assembly. Earth Sci. Rev. 82:217-56. Collins, W.J. 2002. Hot orogens, tectonic switching, and creation of continental crust. Geology, 30: 535–538. Collo, G; & Astini, R.A. (in press). La Formación Achavil: una unidad diferenciable dentro del basamento metamórfico de

bajo grado del Famatina en la región pampeana de los Andes Centrales. Revista Asociación Geológica Argentina. Collo, G.; Astini, R. A.; Cardona, A.; Do Campo, M. D. & Cordani, U. (in press). “Edad del metamorfismo de las unidades

con bajo grado de la región central del Famatina: La impronta del ciclo orogénico oclóyico”. Revista Geológica de Chile. Cordani, U.G., 2006. Neodymium isotopes, accretionary belts, and their bearing on the crustal evolution of South America. V

South American Symposium on Isotope Geology, 211-214. Dahlquist, J.A., Pankhurst, R.J., Rapela, C.W., Galindo, C. Alasino, P., Fanning, C.M., Saavedra, J. & Baldo, E. (in press).

New shrimp U-Pb data from the Fmatina complex: constraining Early–Mid Ordovician Famatinian magmatism in the Sierras Pampeanas, Argentina. Geologica Acta

Pankhurst, R.J., Rapela, C.W. & Fanning, C.M. 2000. Age and origin of coeval TTG, I- and S-type granites in the Famatinian belt of NW Argentina. Transactions of the Royal Society of Edinburgh: Earth Sciences, 91: 151-168.

Ramos, VA. 1986. El diastrofismo oclóyico: un ejemplo de tectónica de colisión durante el Eopaleozoico en el noroeste Argentino. Rev. Inst. Cienc. Geol. 6: 13-28.

Ramos, V.A. 2008. The Basement of the Central Andes: The Arequipa and related Terranes. Annual Review of Earth and Planetary Sciences, 36.

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The 2007 Pisco earthquake (Mw=8.0), Central Peru: Preliminary field investigations and seismotectonic context

L. Audin1, H. Perfettini

1, D. Farber

2, H. Tavera

3, F. Bondoux

1, & J.-P. Avouac

4

1 LMTG, Toulouse, France (laurence.audin@ird.fr)

2 USC, Santa Cruz, California, USA (dfarber@ucsc.edu)

3 IGP, Lima, Peru (hjtavera@geo.igp.gob.pe)

4 Caltech, Pasadena, California, USA (avouac@gps.caltech.ed)

This epicentral area of the 2007 Pisco earthquake marks a major transition in the characteristics of the Nazca

subduction zone: 1) the megathrust dip angle is shallower (10-20°) to the north than to the south (25-30°; Langer

et Spence, 1995) megathrust earthquakes have distinctly smaller magnitudes, recurrence time and are more

fragmented to the north; 3) the distance between the trench and the coastline changes abruptly from ~180km to

the north to ~80km to the south (Figure 1). These variations are likely related to the oblique subduction of the

Nazca ridge - a major bathymetric high - beneath the continental margin.

Figure 1: Spatial dispersion of two weeks of aftershocks along the coastline. In red the focal mecanism showing a 13° dipping plane (IGP, 2007).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 64-66

65

The effect of the subduction of the ridge is anyway obvious in the morphology – river changing course on

Figure 2 - and tectonics of the forearc, in particular, around the Paracas Peninsula where Miocene and Pliocene

marine formations are uplifted and the forearc tectonic regime changes from compression to extension (Figure 2,

Machare et al., 1992; Audin et al., 2008).

Figure 2: Topography and fault systems in Pisco and Paracas Peninsula region.

The geometry of the coastline reflects the sweeping of ridge beneath the margin. The coastline geometry also

seems to mirrors the variation of the downdip edge of the LFZ and some coupling. Just after the earthquake, our

preliminary field survey investigated evidence for uplift or subsidence along the coast and found that the

coastline didn’t experience any significant vertical displacement compared to the tide range (~40cm). The

coastline approximately correspond in general from north to south in Peru to a pivot line marking the transition

from coastal uplift in the south to subsidence in the north, as the distance from the trench increases or decreases.

This model is consistent with the co-seismic slip distribution inferred from waveform modeling (Pritchard and

Fielding., 2008), and with the distribution of aftershocks which suggests that the subduction interface ruptured

mainly updip of the coastline (Figure 1). To place further constraints on the coseismic slip distribution, we have

collected data on the spatial extent of Tsunami waves which hit the coast both south and north of the Paracas

peninsula.

Finally, our field surveys have also revealed evidence for active faulting of the forearc. In particular, the

production of coseismic pressure ridges, with up to 50cm of vertical throw suggests that the east dipping Puente

Huamani thrust fault system was reactivated over a distance of about 20km during this event. This event also

triggered very localized but widely outspread soil liquefaction that lined up with pre existing structures along the

coast in a NS direction. However, we didn’t find evidence for reactivation of any of the normal faults on the

Paracas Peninsula, although some had been reactivated by the 2006 Pisco earthquake (Mw6.4). The main effects

of the tsunami are observed south of the Paracas Peninsula although the main shock occurred north of it. Thus,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 64-66

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the structure and deformation of the Peruvian forearc and coastline seems to contain important information on

lateral variations of seismic and geodetic coupling along the subduction zone.

Figure 3: Damages in Lagunillas Bay ( South of Paracas Peninsula ) after the Tsunami; in Pisco , Plaza de Armas , after the earthquake; in Paracas after the tsunami that lift up the pontoon. Liquefaction evidences, trending NS and lined up with the Huamani Quaternary flexure.

References Audin L., Lacan P. , Tavera H., Bondoux F., Upperplate deformation and seismic barrier in front of Nazca subduction zone:

The Chololo Fault System and active tectonics along the Coastal Cordillera, southern Peru. Tectonophysics. In press. Langer C. J., W. Spence, The 1974 Peru earthquake series. Bulletin of the Seismological Society of America; June 1995; v.

85; no. 3; p. 665-687 Macharé , J., Ortlieb, L., 1992. Plio-Quaternary vertical motions and the subduction of the Nazca Ridge, central coast of

Peru. Tectonophysics 205, 97 ± 108. Pritchard and Fielding, in press. A study of the 2006 and 2007 earthquake sequence of Pisco , Peru, with InSAR and

teleseismic data. GRL.

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The 2006 eruptions of the Tungurahua volcano (Ecuador) and the importance of volcano hazard maps and their diffusion

D. Barba1, P. Samaniego

1, J.-L. Le Pennec

2, M. Hall

1, C. Robin

2, P. Mothes

1, H. Yepes

1, P.

Ramón1, S. Arellano

1, & G. Ruiz

1 Instituto Geofísico, Escuela Politécnica Nacional (IG-EPN), Ap. 17-01-2759, Quito, Ecuador

2 Institut de Recherche pour le Développement, LMV, 5 rue Kessler, 63038 Clermont-Ferrand, France

KEYWORDS : volcanic hazard map, 2006 eruptions, Tungurahua volcano, Ecuador

Introduction

The Tungurahua volcano (5023 m asl) is a steep-sided, andesitic stratovolcano, located in central Ecuador,

ranking as one of the most active volcanoes of the Northern Andes. During historical times Tungurahua

experienced important (VEI 3) pyroclastic flow-forming eruptions in AD 1640, 1773, 1886, and 1918 (e.g.

Hall et al., 1999; Le Pennec et al., 2008).

In October 1999, after about 75 years of quiescence, the Instituto Geofísico of the Escuela Politécnica Nacional

(IG-EPN) registered a renewal of the eruptive activity. During the next six years, this activity was cyclical, with

small to moderate explosions responsible for important ash emissions, the most voluminous of which occurred

on November-December 1999, August 2001, September 2002 and October-November 2003. In 2006, seismic

activity increased dramatically and culminated with the 14-16th July (VEI 2) and 16-17th August 2006 (VEI 3)

explosive eruptions. For the first time since the beginning of this eruptive cycle, Tungurahua volcano produced

pyroclastic flows, which swept over the western half of the cone, as well as giving rise to eruption columns

greater than 15 km in height.

Hazards mitigation during an eruption depends on a continuous monitoring, as well as a reliable hazard map.

The latter is the starting point for develops risk maps, territorial planning and emergencies management. In fact,

the early warning provided by the IG-EPN to the local authorities allowed the evacuation of thousands of people

living in the high-hazard zone. As a result, human loss was limited to 6 fatalities. In this abstract, we will

describe the 2006 eruptions, and the importance of the volcano hazard maps and their diffusion for hazard

assessment and emergency planning.

14-16th July eruption

The seismic activity rapidly increased since 14h30 local time (= GMT-5). At first time, a train-like sound and a

continuous shake were feel around the volcano. The eruption started at 17h33, with strong cannon-like periodical

explosions, which were followed by continuous roars (bramidos), related with a 3-4 km-high eruption column.

An almost continuous lava fountain, reaching up to 300 m-high, produced the first pyroclastic flows at 18h00

(Fig. 1). These flows descended toward the Cusúa and Juive Grande villages. The paroxysmal phase occurred

between 19h40 and 01h00 giving rise to eruption columns greater than 20 km in height. During the paroxysm, at

least 11 pyroclastic flows were generated, which descended on the north-western flank and the Vazcun valley

(Fig. 2).

The activity decreased progressively at 15th July, registered only a few explosions. At least 6 small to moderate

pyroclastic flows were produced at 16th July; all them were associated with vulcanian explosions.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 67-70

68

Fig. 1 A photo of Tungurahua’s vulcanian eruption, with accompanying pyroclastic flows, that occurred

on July 14, 2006. Credits: BBC

Fig. 2 Thermal image (FLIR) obtained from TVO. A pyroclastic flow descending to Cusúa town

during the climatic phase.

16-17th August eruption

Eruptive activity increased from the morning of August 16th. At 14h30, eruptive activity was characterized by a

continuous ash and steam emission, reaching 2-3 km above the crater. First small pyroclastic flows occurred

around 17h00 and descended down the western flank, following the Cusúa and Chontapamba gullies (Fig. 3). An

almost 300 m-high continuous lava fountain, associated with a 3-4 km-high eruption column produced several

small pyroclastic flows, those descended toward Cusúa, Juive Grande and Vascún valleys. Other sporadic, but

probably bigger, pyroclastic flows were generated between 21h00 and 24h00, mostly related to explosions

and/or an increase of the lava fountain, the flows affected the northern and western flanks. The flow most

extensive in the Vascún valley, stopped 1.5 km before Baños city.

Fig. 3 Tungurahua volcano in eruption during August 16, 2006.

Fig. 4 FLIR image showing the lava flow, which marked the end of the eruptive cycle of July-August,

2006.

The paroxysmal phase initiated at 00h15 (August 17) and ended around 40 minutes after. Eruptive activity was

characterized by a powerful lava fountain up to 1000 m above the crater, a 15 km-high eruption column, and the

contemporaneous generation of the most important pyroclastic flows, which descended by 17 ravines on the

north, north-west, west and south-west flanks. The flows reached up to 8.5 km until get the base of the volcano

after a descent of 2600-3000m from the summit crater. The pyroclastic flows of the Rea, Romero and

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69

Chontapamba ravines formed important deltas in the Rio Chambo valley; this was dammed for several hours

after the eruption. The pyroclastic flows that followed in Mapayacu and Juive Grande ravines also dammed the

Puela and Pastaza rivers, respectively.

No pyroclastic flow was witnessed on the eastern flank of the cone and no deposits were observed over this

region during the helicopter observation done by the staff of IG-EPN. After the paroxysmal phase both the

seismic and the volcanic activity rapidly decreased. On the afternoon of August 17th, IG thermal images of the

NW flank confirmed the effusion of an important blocky lava flow which was emitted some hours after the

paroxysmal phase and stopped at 2700 m asl (Fig. 4).

Use of Tungurahua hazard map

The first Tungurahua volcano hazard map was published by IG-EPN in 1988 (Hall et al., 1988). Based on an

extensive study of the volcano by scientific of IG-EPN and IRD, an improved version was published 14 years

later (Hall et al., 2002). Local authorities used these maps for emergency planning during the unrest of the

volcano and during the 2006 crisis, respectively.

Fig. 5 The distribution of pyroclastic flow and surge deposits shows a good agreement with the high-

hazard zone depicted in the 2002 map. The July 14th PF deposits (yellow) are showed on the August 16th

PF (brown) and surges (beige) for a better visualization.

Fig. 6 Third edition of Tungurahua hazard map. Samaniego et al., in press.

Figure 5 shows the distribution of pyroclastic flows and surges for the 14thJuly and August 16th eruptions. A

good agreement exists between these deposits and the high-hazard zone defined by the 2002 map. This

comparison highlights the relevance and validity of this hazard map. The experience obtained during the current

eruption, allow us to incorporate different eruptive scenarios (Fig. 6). This fact is extremely important for

emergency management. Moreover, the well-constrained information from the 2006 eruptions is also being used

to calibrate numerical simulations for pyroclastic flows. This constitutes a first step towards a new generation of

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70

dynamic volcano hazard maps for Ecuadorian volcanoes.

Conclusion

The deployment of a monitoring system by the IG-EPN since 1988, and the installation of the Tungurahua

Volcano Observatory in 1999 at a location close to the volcano, allowed IG scientists to communicate to the

authorities the course of the volcanic events during the seven years process, and finally to successfully issue

early warnings to national and local authorities and to the people before the July and August, 2006 explosive

eruptions.

The relevant use of the hazard map of Tungurahua during the 2006 emergency period is undeniably due to

constant scientific improvement of the hazard map. In parallel, the volcanological information has been

popularized by the publication of a booklet authored by IG-EPN and IRD scientists, and the work with the

community in the framework of an European-funded DIPECHO project.

References Barba D., Arellano S., Ramón P., Mothes P., Alvarado A., Ruiz G., Troncoso L., 2006 Cronología de los eventos erptivos de

Julio y agosto del 2006 del volcán Tungurahua. Resumen extendido de las 6tas. Jornadas en Ciencias de la Tierra. Ecuador, DG-EPN, 177-180.

Samaniego P., J.-L. Le Pennec., Barba D., Hall M. L., Robin C., Mothes P., Yepes H., Troncoso L & Jaya D. (in press). Mapa de los peligros potenciales del volcán Tungurahua. Escala 1:50 000. 3ra Edición. Instituto Geofísico, Escuela Politécnica Nacional. Quito – Ecuador.

Hall M.L., Robin C., Samaniego P., Monzier M., Eissen J.-P., Mothes P., Yepes H., von Hillebrandt C. & Beate B., 2002. Mapa de los peligros potenciales del volcán Tungurahua. Escala 1:50 000. Quito, 2da Edición. Instituto Geofísico, Escuela Politécnica Nacional.

Hall, M.L., Robin, C., Beate, B., Mothes, P., Monzier, M., 1999. Tungurahua volcano, Ecuador: structure, eruptive history and hazard. J. Volcanol. Geotherm. Res. 91, 1-21.

Hall M.L., Beate B. & von Hillebrandt C., 1988. Mapa de los peligros volcánicos potenciales asociados al volcán Tungurahua. Escala 1:50 000. Quito, 1ra Edición. Instituto Geofísico, Escuela Politécnica Nacional.

Le Pennec J.-L., Jaya D., Samaniego P., Ramón P., Moreno Yánez S., Egred J., Submitted manuscript, Journal of Volcanology and Geothermal Research, 2008. Eruptions of Tungurahua volcano, Ecuador from Late Integration to Early Colonial times: evidence from historical narratives, stratigraphy and radiocarbon age determinations.

Le Peneec J.-L., Samaniego P., Eissen J.-P., Hall M.L., Molina I., Robin C., Mothes P., Yepes H., Ramón P., Monzier M. & Egred J., 2005. Los peligros volcánicos asociados con el volcán Tungurahua. Quito, 2da edición. Corporación Editoria Nacional. 113 p.

Samaniego P., Eissen J.-P., Le Pennec J.-L., Hall M.L., Monzier M., Mothes P., Ramón P., Robin C., Egred J., Molina I. & Yepes H., 2003. Los peligros volcánicos asociados con el volcán Tungurahua. Quito, 1ra edición. Corporación Editoria Nacional. 108 p.

Yepes H., Ramón P., Barba D., Arellano S., Samaniego P., Hall M.L., Mothes P., Alvarado A., Le Pennec J.-L., Kumagai H., Rivero D. (2007) Tungurahua Volcano’s 2006 Eruptions, Monitoring and Alert Notifications. Abstract of the Cities on volcano 5; November 19 – 23; Shimabara – Japón.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 71-74

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Control of Mesozoic extensional structures on the Andean deformation in the northern Malargüe fold and thrust belt, Mendoza, Argentina

Florencia Bechis1,2

, Laura Giambiagi1, Daniel Yagupsky

2, Ernesto Cristallini

2, Víctor

García2, & José Mescua

1

1 CONICET - IANIGLA (CRICYT), CC 330 (5500), Mendoza, Argentina (fbechis@lab.cricyt.edu.ar,

lgiambia@lab.cricyt.edu.ar, jmescua@lab.cricyt.edu.ar) 2 CONICET - Laboratorio de Modelado Geológico, Universidad de Buenos Aires, Pabellón 2, Ciudad Universitaria

(1428), Buenos Aires, Argentina (daniely@gl.fcen.uba.ar, ernesto@gl.fcen.uba.ar, victorg@gl.fcen.uba.ar)

KEYWORDS : Neuquén basin, Atuel depocentre, inversion, reactivation, sandbox analogue modelling

Introduction

The study area is located in the northern sector of the Malargüe fold and thrust belt, in the Central Andes of

central-western Argentina (fig. 1). In this sector of the belt, Andean deformation inverted a Late Triassic to Early

Jurassic extensional depocentre of the Neuquén basin (Kozlowski et al., 1993; Manceda and Figueroa, 1995;

Giambiagi et al., in press). The main goal of this contribution is to identify the principal controls exerted by the

extensional structures over the Andean deformation in the northern sector of the Malargüe fold and thrust belt.

Our study is based on data obtained from detailed structural and geological field mapping, integrated with

subsurface information. The orientation, timing and structural style of the Andean structures were compared

with the extensional architecture of the Atuel depocentre, previously addressed by Manceda and Figueroa

(1995), Lanés (2002, 2005) and Giambiagi et al. (2005). In addition, the mapped structures were compared with

the results obtained from scaled sandbox analogue models simulating the deformation of a half-graben oblique to

the shortening direction during the evolution of a fold and thrust belt.

The Atuel depocentre

The extensional structure of the Atuel depocentre is characterized by the presence of two NNW-trending major

faults, the La Manga and Alumbre faults (fig. 2), marked by the distribution of the synrift deposits in outcrops

and from subsurface data (Giambiagi et al., 2005, in press). These faults limited two major half-grabens,

controlling the main subsidence of the sub-basin and the distribution of sedimentary environments and drainage

systems during the synrift phase (Lanés 2002, 2005). Inside the half-grabens, we identified a bimodal

distribution of normal faults with NNW and WNW trends (fig. 2).

Andean structures

The structure of the northern sector of the Malargüe fold and thrust belt was previously studied by Kozlowski

et al. (1993), Manceda and Figueroa (1995), Fortunatti et al. (2004), Turienzo et al. (2004) and Giambiagi et al.

(in press), among others. This belt is characterized by a western thick-skinned sector and an eastern thin-skinned

one (fig. 1). In the western sector the Triassic-Jurassic synrift infill of the Atuel depocentre is involved in the

deformation, whereas in the eastern sector low-angle thrust faults affect Cretaceous to Neogene strata. The limit

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 71-74

72

between these two contrasting structural styles coincides with the eastern border of the Atuel depocentre, which

is controlled by the NNW-trending La Manga normal fault (fig. 2). This fault was strongly inverted during the

Andean deformation, transferring displacement to shallow thrusts located eastwards. Ar-Ar dating of pre-, syn-

and post-tectonic volcanic and subvolcanic rocks showed that the La Manga fault was reactivated early on the

evolution of the fold and thrust belt, localizing great part of the deformation (Giambiagi et al., in press).

Figure 1. Geological and structural map of the study area, modified from Giambiagi et al. (in press).

The thin-skinned structures of the eastern sector have a general N to NNW trend, probably related to a transfer

of slip from the basement-involved inverse faults of the La Manga fault system (fig. 1). On the other hand, in the

western sector most folds and reverse faults show NNE orientation (fig. 1). These NNE compressive structures

are segmented by WNW transfer zones, which were controlled by the presence of Mesozoic second-order

normal faults (e.g. in the Río Atuel and the upper section of the Arroyo Blanco, figs. 1 and 2). NNW orientations

are locally observed, and they are interpreted as being controlled by previous NNW-oriented normal faults (e.g.

the Alumbre fault, fig. 2).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 71-74

73

Figure 2. A) Interpreted extensional architecture of the Atuel depocentre of the Neuquén basin, modified from Giambiagi et al. (2005, in press). See figure 1 for location. B) Results from the analogue modeling of inversion of an oblique half-graben (obliquity angle ~15º). The fault pattern interpretation is shown at each step of progressive shortening (S).

Analogue modelling

A sandbox analogue model was built in order to test the influence of a preexisting obliquely oriented half-

graben over the style of shortening deformation (a complete description of the experimental methodology can be

consulted in Yagupsky et al., in press). The model materials were quartz sand and glass microbeads. In a first

extensional stage a NNW-trending half-graben was created. Later, the model was shortened by displacing a

moving wall oriented at 15º from the half-graben.

The map view evolution of the experiment was registered and interpreted (fig. 2). During the early stage of

contraction the oblique half-graben acted as a nucleation site for thrust faulting, producing the inversion of its

bounding normal fault. This reactivation progressed from NNW to SSE during further shortening. In the next

stage, thrusts branched from the previously developed oblique one, creating a new active deformation front in

the inner side of the system with NNE strikes, cutting through the underlying half-graben. The resulting

structural architecture shows a NNW-trending thrust fault controlled by the underlying structure and two NNE-

trending thrusts affecting the inner sector of the model.

Discussion

We evaluated possible structural explanations for the NNE orientation of most of the compressive structures

that characterize the thick-skinned sector of the fold and thrust belt:

1) Due to the basement involvement in the deformation of the western sector, this orientation could be

reflecting the basement structural grain. Upper Paleozoic structures with a NNE trend have been reported to the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 71-74

74

north of the study area, corresponding to compressive structures related to the San Rafael orogenic phase (Azcuy

and Caminos, 1987; Giambiagi et al., 2008).

2) At these latitudes, in the Chilean side of the Andes, structures with similar orientations to the ones

observed in the Atuel area have been described: WNW and NW faults and lineaments, and NNE thrust faults

(Rivera and Yáñez, 2007). In this case, NNE faults were interpreted as probable Oligocene to Miocene structures

reactivated during Miocene inversion of Tertiary basins. However, there is no record of development of mid-

Tertiary basins in the Argentinean side.

3) The analogue models of inversion of obliquely oriented half-grabens suggest that the observed

structural pattern could be related to a progressive inversion of the Atuel depocentre, assuming a rift trend

oblique to the Andean shortening direction. In the analogue models, an early inverse reactivation of the normal

master fault is followed by the formation of slightly oblique thrusts affecting the synrift infill of the graben (fig.

2). The inversion of the Atuel depocentre had a similar evolution, with an initial reactivation of the NNW-

trending La Manga fault, and a later deformation of the sub-basin interior with development of NNE-oriented

structures.

The close similarity between the results of the analogue models and the mapped area suggest that the third

interpretation is a reliable possibility, taking into account that this similarity does not necessarily imply similar

deformation mechanisms. This option explains the orientation of the structures as a result of inversion of the

Atuel depocentre during the formation of the northern sector of the Malargüe fold and thrust belt. However, the

first possibility has to be taken into account. To test if this option had a role in the Andean deformation, a future

approach considering regional data is needed.

References Azcuy, C.L., & Caminos, R., 1987. Diastrofismo. In S. Archangelsky (ed.): El sistema carbonífero en la República

Argentina, Academia Nacional de ciencias, Córdoba, Argentina: 239-252. Fortunatti, N., Turienzo, M., & Dimieri, L., 2004. Retrocorrimientos asociados al frente de avance orogénico, arroyo Blanco,

Mendoza. Asociación Geológica Argentina, Publicación Especial, Serie D 7: 34-30. Giambiagi, L., Bechis, F., García, V., & Clark, A., in press. Temporal and spatial relationships of thick- and thin-skinned

deformation: a case study from the Malargüe fold and thrust belt, southern Central Andes. Tectonophysics. Giambiagi, L., Bechis, F., Lanés, S., & García, V., 2005. Evolución cinemática del depocentro Atuel, Triásico Tardío –

Jurásico Temprano. 16º Congreso Geológico Argentino, La Plata. Proceedings in CD. Giambiagi, L., Mescua, J., Folguera, A., & Martinez, A., 2008. Pre-andean deformation in the southern Central Andes (32°-

33°S). 7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts. Kozlowski, E., Manceda, R., & Ramos, V. A., 1993. Estructura. In: V. Ramos (Editor), Geología y recursos naturales de

Mendoza. 12º Congreso Geológico Argentino y 2º Congreso de Exploración de Hidrocarburos, Relatorio: 235-256. Lanés, S. 2002. Paleoambientes y Paleogeografia de la primera transgresion en Cuenca Nequina, Sur de Mendoza.

Universidad de Buenos Aires, PhD Thesis (unpublished), 259 pp. Lanés, S., 2005. Late Triassic to Early Jurassic sedimentation in northern Neuquén Basin, Argentina: Tectonosedimentary

evolution of the first transgression. Geologica Acta 3(2): 81-106. Manceda, R., & Figueroa, D., 1995. Inversion of the Mesozoic Neuquén rift in the Malargüe fold-thrust belt, Mendoza,

Argentina. In: A. J. Tankard, R. Suárez and H.J. Welsink (Editors): Petroleum Basins of South America. American Association of Petroleoum Geologists, Memoir 62: 369-382.

Rivera, O., & Yáñez, G., 2007. Geotectonic evolution of the Central Chile Oligocene-Miocene volcanic arc, 33-34ºS: towards a multidisciplinary re-interpretation of inherited lithospheric structures. Geosur, Santiago de Chile, Abstracts: 138.

Turienzo, M., Fortunatti, N., & Dimieri, L, 2004. Configuración estructural del basamento en la confluencia del arroyo Blanco y el río Atuel, Mendoza. Asociación Geológica Argentina, Publicación Especial, Serie D 7: 27-33.

Yagupsky, D.L, Cristallini, E.O., Fantín, J., Zamora Valcarce, G., Bottesi, G., & Varadé, R., in press. Oblique half-graben inversion of the Mesozoic Neuquén Rift in the Malargüe Fold and Thrust Belt, Mendoza, Argentina: New insights from analogue models. Journal of Structural Geology.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 75-76

75

Flat-slab subduction beneath the Sierras Pampeanas in Argentina

Susan Beck1, Patricia Alvarado

2, Lara Wagner

3, Megan Anderson

4, Hersh Gilbert

5, & George

Zandt1

1 Department of Geosciences, University of Arizona, Gould-Simpson Building, 1040 E. 4

th St., Tucson, Arizona,

85721 USA 2 Department of Geophysics and Astronomy, National University of San Juan, San Juan, Argentina

3 Department of Geological Sciences, University of North Carolina at Chapel Hill, 104 South Road, Chapel Hill,

NC 27599, USA 4 Department of Geology, Colorado College, 14 E. Cache La Poudre St., Colorado Springs, CO 80903 USA

5 Department of Earth and Atmospheric Sciences, Purdue University, 550 Stadium Mall Dr., West Lafayette, IN

47907, USA

KEYWORDS : flat slab, Sierras Pampeanas, Juan Fernandez Ridge, broadband seismology

Introduction

One of the intriguing aspects of the subduction of the Nazca plate beneath western South America is the along

strike segmentation of the dip of the descending plate as defined by the slab earthquake distribution. In the south

central Andes, the subducting Nazca slab has a subhorizontal geometry and extends inland over 300 km beneath

the Sierras Pampeanas (SP) near 30°-31°S but returns to a normal dip further south at 33°S. The tectonic

evolution of this region is the result of the interaction between the South American plate and the segment of the

Nazca Plate that contains the Juan Fernandez Ridge. The flat slab region is characterized by an absence of

modern arc volcanism, the Precordillera fold and thrust belt and the presence of the inland basement cored

uplifts of the Sierras Pampeanas. Understanding what causes the flat slab geometry, and its influence on the

overlying lithosphere remains a fundamental goal. In order to study the flat slab region of Argentina we have

done two passive broadband seismic deployments in the region in order to characterize the lithospheric structure.

We combine the results from a range of seismic studies, which used data collected during the Chile Argentina

Geophysical Experiment (CHARGE) to present an up-to-date model of the crustal and upper mantle structure in

central Chile and Argentina. These studies include receiver functions (both regional and teleseismic), earthquake

hypocenter relocations and focal mechanisms, Pn diffraction studies, isostasy studies, regional waveform

modeling, and regional P, S and Vp/Vs tomographic inversions. We are currently collecting additional seismic

data in the SIEMBRA project (40 stations above the flat slab) to do higher resolution imaging of the flat slab

region of Argentina.

Results

We have refined the location of the earthquakes in the slab using a grid-search multiple event location method

to relocate over a 1000 events in the subducted slab (Anderson et al., 2007). The earthquake locations and

resultant Wadati-Benioff zone contours show that the shallowest portion of the flat slab is associated with the

subducting Juan Fernandez Ridge at 31°S. Most of the earthquake focal mechanisms in the slab show

subhorizontal T-axis solutions consistent with slab pull (Anderson et al., 2007). Alvarado et al. (2005) find that

most of the region above the flat slab is in compression with thrust earthquakes in the depth range of 5-25 km.

We have analyzed the crustal thickness across the arc and backarc using several different seismic techniques.

We determined a 2D crustal model beneath Chile and western Argentina along an east west transect at

approximately 30° using both apparent Pn velocities and receiver functions (Fromm et al., 2004; Gilbert et al.,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 75-76

76

2006). This model consists of a thick crustal root beneath the High Cordillera (65 km), a 60 km thick crust under

the Precordillera, thinning to 55 km beneath the western Sierras Pampeanas, and 35-40 km thick crust beneath

the eastern Sierras Pampeanas. Receiver functions and regional waveform modeling indicate that the eastern

Sierras Pampeanas has a crustal thickness of 35 km and a Vp/Vs ratio of less than 1.75 while the western Sierras

Pampeanas crustal thickness increases to 55 km with a Vp/Vs ratio greater than 1.8. The change in crustal

character (both thickness and seismic velocity structure) corresponds to sutures between accreted Precambrian

terrains (Gilbert et al., 2006; Alvarado et al., 2007). The western Sierras Pampeanas has a high velocity lower

crust that may be a higher density material that is partially ecologitized but has not yet been removed (Gilbert et

al., 2006; Calkins et al., 2006). Our seismic results are consistent with the observed composition of the basement

associated with the terrains in the Sierras Pampeanas. The contrasting terrains may be responsible for the

different crustal composition and seismic structure of the eastern and western terranes. The predominant felsic

quartz-rich character of the Eastern Sierras Pampeanas is linked to the collision of a microcontinent, the Pampia

terranes (Alvarado et al., 2005; Rapela et al. 1998), whereas the dominant mafic-ultramafic composition in the

west is associated with oceanic fragments of a crust arc/backarc setting for the western Sierras Pampeanas

(Alvarado et al., 2005; Vujovich & Kay 1998; Ramos et al. 2000).

Regional seismic tomography shows the mantle lithosphere is also very heterogeneous, with low seismic

velocities beneath the volcanic arc region, high velocities directly below the Moho in the backarc, and

anomalous mantle (low Vp/Vs ratio, high Vs) directly above the flat slab (Wagner et al., 2005; 2006). The low

Vp/Vs ratio and high S-wave velocities directly above the flat slab are not consistent with hydrated mantle but

rather with dry conditions and a high percentage of orthopyroxene in the mantle layer above the flat slab

(Wagner et al., 2005; 2006). Taken together the CHARGE results suggest that the lithospheric structure still

reflects Precambrian terrane boundaries and that the present day mantle under the Sierras Pampeanas is dry,

strong and probably able to transmit basal shear from the underlying flat slab.

References Alvarado, P., Beck S. L., Zandt G., 2007, Crustal Structure of the South-Central Andes Cordillera and Backarc Region from

Regional Waveform Modeling, Geophys. J. Int., doi:10.1111/j.1365-246X.2007.03452.x 2007. Alvarado, P., Beck S., Zandt G., Araujo M., and Triep E., 2005, Crustal deformation in the south-central Andes back-arc as

viewed from regional broad-band seismic waveform modeling, Geophys. J. Intl., doi: 10.1111/j.1365-246X.2005.02759. Alvarado, P., Castro de Machuca, B. and Beck, S., 2005, Comparative crustal seismic study of the Western and Eastern

Sierras Pampeanas region, Argentina, (31ºS). Revista de la Asociación Geológica Argentina, Vol. 60 (4), 787-796. Anderson, M., P. Alvarado, G. Zandt, S.L. Beck, 2007, Geometry and brittle deformation of the subducting Nazca plate,

central Chile and Argentina, Geophys. J. Int., doi:10.1111/j.1365-246X.2007.03483.x, 2007. Calkins, J.A., G. Zandt, H.J. Gilbert, and S.L. Beck, 2006, Crustal images from San Juan, Argentina, obtained using high

frequency local event receiver functions. Geophys. Res. Lett., L07309, doi:10.1029/2005GL025516. Fromm, R., P. Alvarado, S. L. Beck, G. Zandt, 2006, The April 9, 2001 Juan Fernandez Ridge outer-rise earthquake (Mw

6.7) and its aftershock sequence, J. of Seismology, 10.1007/s10950-006-9013-3. Gilbert, H., S. Beck, G. Zandt, 2006, Lithospheric and upper mantle structure of central Chile and Argentina; influences of a

flat slab, Geophys. J. Int., 165, 383–398. Ramos, V.A., Escayola, M., Mutti, D.I. & Vujovich, G.I., 2000, Proterozoicearly Paleozoic ophiolites of the Andean

basement of southern South America, in Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program, pp. 181–217, eds Dilek, Y., Moores, E.M., Elthon, D. & Nicolas, A., Geol. Soc. Sp. Publ. 142.

Vujovich, G.I. & Kay, S.M., 1998. A Laurentian? Grenville-age oceanic arc/back-arc terrane in the Sierra Pie de Palo,Western Sierras Pampeanas, Argentina, in The Proto-Andean Margin of Gondwana, eds Pankhurst, R.J. & Rapela, C.W., pp. 159–179, Geological Society, London, Special Publications 142.

Wagner, L. S. Beck, G. Zandt, M. Ducea, 2006, Depleted lithosphere, cold, trapped asthenosphere, and frozen melt puddles above the flat slab in central Chile and Argentina, Earth and Planetary Science Letters, 245 289–301.

Wagner, L., S. L. Beck, G. Zandt, 2005, Upper mantle structure in the South Central Chilean subduction zone, J. Geophys. Res., 110, B01308, doi:10.1029/2004JB003238.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 77-80

77

The November 14, 2007, Mw=7.7 Tocopilla (Chile) earthquake: Preliminary results from InSAR and GPS

M. Béjar-Pizarro1,2

, D. Carrizo2, A. Socquet

2, R. Armijo

2, J.-C. Ruegg

2, J.-B. de Chabalier

2, A.

Nercessian3, O. Charade

3, & S. Bonvalot

4

1 Dpto. Geodinámica, Facultad de Ciencias Geológicas, Universidad Complutense de Madrid, c/José Antonio

Novais s/n, 28940 Madrid, Spain (mbejarpi@geo.ucm.es) 2 Laboratoire de Tectonique et et Mécanique de la Lithosphère, Institut de Physique du Globe de Paris, 4 place

Jussieu, 75252 Paris cedex 05, France 3 Laboratoire de Sismologie, Institut de Physique du Globe de Paris, 4 place Jussieu, 75252 Paris, France

4 Laboratoire des Mécanismes et Transferts Géologique, UMR 5563, IRD, UR154, CNRS, Toulouse, France

KEYWORDS : subduction earthquake, Northern Chile, seismic gap, seismic cycle, InSAR, GPS, elastic models

Introduction

A Mw 7.7 subduction earthquake occurred on November 14, 2007 in Tocopilla (northern Chile). This region

(between 16.5ºS and 23.5ºS) had been identified as major seismic gap (~1000 km length) that had not ruptured

since the occurrence of the South Peru (Mw = 9.1, 16 August 1868) and the Iquique (Mw = 9.0, 10 May 1877)

megathrust earthquakes. This gap was reduced to a length of ~500 km after the occurrence of the Arequipa (Mw

= 8.3, 23 June 2001) and the Antofagasta (Mw = 8.1, 30 July 1995) earthquakes (Figure 1).

Most of the aftershocks following the 2007 event were concentrated in the north of the Mejillones Peninsula

(Figure 1), an inter-segment zone that appears to act both as a barrier arresting rupture of large earthquakes (e.g.

Figure 1. Reference map of our study area. Large subduction earthquakes along the Peru-Chile trench are represented by green rectangles and ellipses. Grey ellipses represent the approximate extent of 1877 and 1868 rupture zones. The Harvard CMT solution for the mainshock and the 3 biggest aftershocks are indicated. Aftershocks are indicated by yellow dots (NEIC catalog). Dashed squares outline the tracks of Envisat data used in this study. Blue points with capital letters represent the location of the GPS permanent stations used here. Red rectangle represents the approximate extent of the 2007 Tocopilla earthquake.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 77-80

78

M 8.8 1877 Iquique earhquake) and as an asperity where large earthquakes nucleates (e.g. Mw 8.1 1995

Antofagasta earthquake, Ruegg et al., 1996).

Important questions arise after this earthquake in relation with the seismic gap and the subduction interface

here: Which part of the subduction interface has ruptured (geometry and location of the rupture plane)? Which

are the slip distribution and the geodetic moment? The end of the rupture, barriers arresting rupture? (Mejillones

Peninsula) Which is the slip deficit after this earthquake? Relations with other events and the seismic gap in the

north?. We address these questions by using data from space geodesy. The Tocopilla earthquake occurred

within a network of continuous GPS stations operated by IPGP, Caltech, DGF and IRD. An array of 21

benchmarks, installed and previously measured periodically by IPGP/DGF, was resurveyed during a postseismic

intervention. Here we combine GPS data from three of these permanent GPS stations and InSAR data from two

descending tracks to determine the geometry and kinematics of the rupture on the subduction interface.

InSAR measurements

We use 4 Envisat ASAR images from two descending tracks (track 96 and track 368, Figure 1) to form two

independent coseismic interferograms. Both interferograms span the date of the earthquake and they include

some days after the mainshock: 10 days in the case of the track 368 interferogram and 26 days in the case of the

track 96 interferogram. It is therefore probable that they include some postseismic deformation together with the

coseismic deformation. Data were processed using the Caltech/JPL repeat-orbit interferometry package, ROI

PAC (http://www-radar.jpl.nasa.gov/roi pac/). The topographic phase contribution was removed using a

3-arc-sec (90-m) digital elevation model from the Shuttle Radar Topography Mission (SRTM). The orbital

information used in the processing was provided by the ESA (DORIS orbits). After processing, the

interferograms presented a linear ramp, probably due to uncertainties in the orbital ephemeris. A plane was

adjusted to this ramp and removed from each interferogram. Figures 2a and 2b show both interferograms with

the observed displacement along the line of sight direction (LOS). Surface deformation pattern is characterized

by two lobes: the western one shows a LOS displacement towards the satellite, with a maximum value of ~30 cm

and the eastern one represents a LOS displacement away from the satellite, with a maximun value of ~15 cm.

GPS measurements

We use 8 continuous GPS stations data from IPG-DGF and IRD-DGF northern Chilean network for calculate

the co-seismic displacement. The GPS-data analysis was done using GAMIT (King and Bock, 1998) and

GLOBK (Herring, 1998) software packages. Daily solutions over 7 days span prior and after the earthquake

were calculated including 13 IGS stations located in South America and 2 local stations from ENS-DGF central

Chili continuous network. Final orbits and antenna phase center corrections from IGS and IERS earth rotation

parameters were used. To estimate the station position every daily GAMIT solutions were combined using

GLOBK referencing the local network to ITRF05 (Altamimi et al., 2007). Our preliminary approximation to co-

seismic displacements was defined here as a difference between the resulting mean coordinates prior and after

the earthquake. We constrain the earthquake zone using 3 stations UAPE, QUIL and PMEJ (figs. 1, 2). The

PMEJ station, located in Mejillones Peninsula, shows displacements of 21 cm to the west, 13 cm to south and

~ 35 cm of uplift. The QUIL station, located ~60 km to inland shows displacement 5 cm to the south, 6 cm to the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 77-80

79

west and 3 cm of subsidence. The UAPF station, located in the coast ~168 km to the north of QUIL, does not

show displacement.

Modelling

We have modelled InSAR data and compare our models with GPS observations. Observed surface

displacements are modeled using Okada’s formulation of a dislocation buried in an infinite elastic half-space

(Okada, 1985). We performed a series of forward models to assess the dislocation parameters that are best fitting

the deformation pattern. The extension of the deformation in the north-south direction constrains the length of

the plane to ~ 150 km (black rectangle in Figures 2a, 2b, 2c, 2f). The plane width (~ 40-43 km) is mostly

constrained by the sharpness of the deformation lobes. Using this fault geometry we test different possible

locations of the plane at depth. As we only have one displacement component (in the LOS direction) the dip of

the fault plane is poorly resolved, so we started by using the CMT Harvard value of 20º. A trade-off between

depth and slip is observed: the deeper the plane is, the more slip is allowed to fit the surface deformation, and the

higher is the equivalent magnitude of the model. he best compromise between seismic moment and depth is

found for a plane that lies between 30-35 and 48-53km depth (Figure 2d), with a total slip of ~1.30 m and a

geodetic moment of ~ 2.75 x1020 N.m. Figure 2c shows the resulting model. It accounts for first order co-seismic

deformation associated with the 2007 Tocopilla earthquake. However the model does not fit very good the

Figure 2. Coseismic surface displacements from a) Interferogram 1 (track 368) b) Interferogram 2 (track 96). c) Synthtetic model, f) Continuos GPS: red and black vectors represent observed and modelled horizontal displacement respectively. Modelled vertical displacements are represented by color contours in figure 2f. Fault plane has been represent in a E-W profile at latitude 22.50º in figure 2d (red line). Dashed line represents the subduction interface deduced by the Ancorp experiment. Global seismicity has been represented by blue dots. 2e) E-W profile to -22.3 latitude with the InSAR LOS displacement for the two co-seismic interferograms and the model.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 77-80

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observations in the southern part of the rupture (north of the Mejillones Peninsula) , due to the irregular spatial

distribution of slip. This can be seen in both interferograms, since more deformation is present in the south than

in the north. Horizontal GPS displacements also show a misfit with this model (Figure 2f) that is increasingly

bigger southward (PMEJ station). This probably means that two distinct patches of slip are needed to fit both

GPS and InSAR data, which is consistent with the first results from seismology (Campos et al., 2008; Peyrat et

al., 2008).

Results and Discussion

Our preliminary results indicate that the Tocopilla rupture extended between 23.30° - 22° S with ~150 km of

longitude. The rupture zone was located between 48-53 and 30-35 km of depth and did not propagate up to the

surface. Our best fit plane is consistent with the subduction plane defined by published seismological data,

indicating that was activated the deeper part of the seismogenic interface, well into the transition zone that was

identified earlier (Chlieh et al., 2004). The southern end of the rupture is clearly defined by the GPS and InSAR

observations to the north of the Mejillones Peninsula confirming this zone as a relevant subduction intersegment

barrier. The misfit of our GPS and InsAR observations in the southern end of the rupture may indicate that the

geometry and the kinematics of the rupture is complex change here and hence cannot be reproduced with a

single dislocation model. Currently we explore the hypothesis of two different dislocation using non linear

inversion models

According to the mean displacement inferred by our models (~1.30 m) the Tocopilla earthquake released a

very small portion of the slip deficit accumulated in the seismic gap during the past 130 years (~ 10m) and may

be regarded as a possible precursor of a much larger subduction earthquake rupturing the 500 km long gap.

Acknowledgments M.B.P acknowledges support of Universidad Complutense de Madrid grant. This work was supported by the ANR-05-CATT-01402 project of the French National Research Agency. We acknowledge the support of the European Agency (ESA) for programming Envisat satellite (research project AO-720). ROI_PAC software was provided by the JPL/Caltech. The GMT program was used to create the figures. References Altamimi, Z., Collilieux, X., Legrand, J., Garayt, B. & Boucher, C. 2007. ITRF2005: A new release of the International

Terrestrial Reference Frame based on time series of station positions and Earth Orientation Parameters , J. Geophys. Res., 112, B09401, doi:10.1029/2007JB004949.

Campos, J., Peyrat, S., Bejar, M.,Socquet, A., Meneses, G., Perez, A., Madariaga, R., Favreau, P, Bernard, P., Barrientos, S., Armijo, R., Ash, G., Sobiesiak, M. & Vilotte, J.P. 2008. The Mw 7.7 Tocopilla, Chile, Earthquake of 14 November 2007: A Comprehensive Study Using Teleseismic, Local and InSAR data, AGU 2008 Joint Assembly Abstract, unpublished material

Chlieh, M., Chabalier, J.B., Ruegg, J.C., Armijo, R., Dmowska, R., Campos, J. & Feigl, K. 2004. Crustal deformation and fault slip during the seismic cycle in the north Chile subduction zone, from GPS and InSAR observations, Geophys. J. Int., 158: 695-711.

Herring, T. A., 1997. GLOBK: Global Kalman Filter VLBI and GPS analysis program, v.4.1, Mass. Inst. of Technol., Cambridge.

King, R. & Bock, Y. 2001. Documentation for the GAMIT GPS software analysis, Tech. rep., Scripps Institution of Oceanography, University of California, San Diego, release 10.33.

Okada, Y. 1985. Surface deformation to shear and tensile faults in a half space, Bull. seism. Soc. Am., 75, 1135–1154. Peyrat et al. 2008. Detailed source process of the 2007 Tocopilla earthquake, AGU 2008 Joint Assembly Abstract,

unpublished material Ruegg, J.C., Campos, J., Armijo, R., Barrientos, S., Briole, P., Thiele, R., Arancibia, M., Cañuta, J., Duquesnoy, T., Chang,

M., Lazo, D., Lyon-Caen, H., Ortlieb, L., Rossignol, J.C. & Serrurier, L. 1996. The Mw =8.1 Antofagasta (North Chile) earthquake July 30, 1995: first results from teleseismic and geodetic data, Geophys. Res. Lett., 23: 917–920.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 81-83

81

Spatial and temporal patterns of exhumation across the Venezuelan Andes from apatite fission-track analysis: Implications for Cenozoic Caribbean geodynamics

Mauricio Bermúdez-Cella1,2

, Peter van der Beek2, & Matthias Bernet

2

1 Laboratorios de Termocronología y Geomatemáticas, Escuela de Geología, Minas y Geofísica. Facultad de

Ingeniería, Universidad Central de Venezuela, Caracas, Venezuela 2 Laboratoire de Géodynamique des Chaînes Alpines (UMR 5025), Université Joseph Fourier, 1381 rue de la

Piscine, 38400, Saint-Martin d'Hères, France (Email: mauricio.bermudez-cella@e.ujf-grenoble.fr)

KEYWORDS : exhumation, thermochronology, apatite, strike-slip faults, Venezuelan Andes

The Venezuelan Andes constitute a northeast trending orogen, extending from the Colombian border in the

south to Barquisimeto in the north of Venezuela (Fig. 1a). This orogen is characterized by five major strike-slip

fault systems, the Boconó, Caparo, Central-Sur Andino, Valera and Burbusay faults, and by foreland thrust belts

to the NW and SE (Fig. 1a). The foremost of these faults is the Boconó that extends 500 km in a NE-SW

direction along the entire Venezuelan Andes. Its morphological appearance is expressed by escarpments and

aligned valleys dividing the Andes almost symmetrically in its central part (Mérida Andes). The Caparo fault is a

dextral strike-slip system parallel to the Boconó fault in the eastern part of Venezuelan Andes. The South-

Andean Central system is located between the Boconó and Caparo faults and does not have the same continuity

as these two systems, being subdivided into a southern and a northern part, seemingly without connection

(Soulas, 1983). The Valera and Burbusay systems are continuous N-S trending faults that locally control the

triangular Trujillo block in the NW of the orogen. These different fault systems, together with the two foreland

thrust belts controlled the Paleocene-Eocene sedimentation in the Maracaibo and Barinas basins (Escalona &

Mann, 2003; Audemard & Audemard, 2002; James, 2000). During the Neogene, they appear to separate several

tectonically active structural domains in the northwest, from structural domains with less tectonic activity to the

southeast (Colleta et al., 1997).

The present-day Venezuelan Andes chain results from a complex geodynamic interaction between the

Caribbean Plate, the Panama Arc and the South American Plate. This triple interaction on a macroscopic scale is

expressed by the convergence of a small continental block, the Maracaibo Block, and the South American Plate

(Aleman & Ramos, 2000; Pindell & Dewey, 1982). The margin of the latter already possessed a series of

tectonic discontinuities of different ages, the Boconó, Caparo, Valera and Burbusay faults systems. Oblique plate

convergence resulted in local thrusting, translation, transtension, extension and rotation that led to exhumation of

individual blocks at different times and rates from the late Eocene to the Pliocene.

We study the spatial and temporal patterns of exhumation across the Venezuelan Andes using apatite fission-

track (AFT) thermochronology. Our database currently consists of 37 AFT ages: 14 reported by Kohn et al.

(1984) and determined mostly using the population method; and 23 new ages determined using the external

detector method (Fig. 1b).

The spatial patterns of AFT ages permit distinguishing at least four different blocks with contrasting

exhumation histories. Two blocks, the Sierra La Culata and Sierra Nevada located in the central part of the

Venezuelan Andes are separated by the Boconó fault system,. The Sierra Culata, to the north of the Boconó

fault, experienced exhumation at 4 ± 2 Ma with rates between 0.7 and 1.5 km/Myr. The Sierra Nevada to the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 81-83

82

south, in contrast, experienced a pulse of exhumation at 11 ± 2 Ma with rates between 2.4 and 5.8 km/Myr. The

AFT data suggest ~8 km of relative uplift of the Sierra La Culata with respect to the Sierra Nevada since

~11 Ma. It has been proposed that the Sierra La Culata was affected by Plio-Quaternary transtension along the

Boconó Fault, but our data rather indicate either a distinct phase of NW-SE compression, causing south-directed

thrusting on the Boconó Fault, or continuous oblique strike-slip across the fault. The data are thus consistent

with models that imply significant transpression and uplift of basement blocks along major pre-existing

discontinuities in the Northern Andes (Cobbold et al., 2007; Cardona-Molina et al., 2006; Mora, 1993).

Figure 1. (a) Location of study area and major system faults (Modified of Audemard et al., 2000). (b) Detailed digital topography of the study area with sample sites and fission track ages (yellow numbers: Kohn et al., 1984; green numbers: data derived of this work).

29.60±2.7 3.3±0.6 24.1±3

20±5.6

22.1±3

23±5.7

11.9±1.5

13±1.9

16.7±1.2

21.5±2.1

6.9±0.6

2.5±0.4

6.1±1.2

2.8±0.4

3±0.5

4.9±0.7

4.7±0.6

3.8±0.7

4.2±0.7

1.8±0.4

2.5±0.4

3.4±0.6

2.8±0.4

2.7±0.5

11.3±0.9

11.2±1.7

10.3±1.7

10.1±1.1

11.0±1.6

10.1±1.5

9.30±2.0

10.1±0.9

10.4±2.2

10.6±1.0

11.5±1.8

11.6±1.6

9.90±0.7

(b)

(a)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 81-83

83

A third block, located to the west is the Caparo Block, characterized by AFT ages ranging from 30 to 17 Ma.

These ages appear to have been influenced by prolonged residence in the partial annealing zone. The fourth

block, to the east, is the Trujillo Block limited by the Valera and Burbusay faults. Within this block, three AFT

ages range from 13 to 21 Ma.

These four blocks are key to a better understanding of the control of Caribbean geodynamics on the collisional

Andean orogen. In this context, our data suggest that intraplate deformation caused by the triple collision of

Panama Arc-Caribbean Plate-South American Plate was active at approximately 30 Ma. This fact is corroborated

by independent AFT data from the Perijá region (Shagam et al., 1984) to the NW of the Andes. From the

beginning of the triple collision up to ~17 Ma the Valera-Burbusay and Caparo-Boconó fault systems apparently

did not yet have their current configuration. As a result of the convergence with the Maracaibo Block, they must

have rotated clockwise around poles located very near the Sierra Nevada block and Santa Marta-Bucaramanga

fault system respectively (Chicangana, 2005).

The coeval timing of exhumation of the Caparo and Trujillo blocks, as recorded by comparable AFT ages,

suggests that they were exhumed together in a context of orthogonal thrusting, before being displaced by sinistral

strike-slip along the Boconó fault. The strong pulse of exhumation recorded at 11 Ma in the Sierra Nevada block

was likely related to major uplift, which may have been the cause for Late Miocene deflection of the Orinoco

River, as inferred from the Neogene sedimentary record (Díaz de Gamero, 1996; Hoorn et al., 1995).

References Aleman, A. & V. Ramos. 2000. Northern Andes. In: Cordani, U.G., Milani, E.J., Thomaz, A., Campos, D.A. (Eds.), Tectonic

Evolution of South America. 31st International Geological Congress, Rio de Janeiro, Brazil, pp. 453-480. Audemard, F.A., M. N. Machette, J.W. Cox, R.L Dart, and K.M. Haller. (2000) Map and Database of Quaternary Faults in

Venezuela and its Offshore Regions. U.S Geological Survey. Open-File Report 00-018 (paper edition) Audemard, F.E. & F.A., Audemard. 2002. Structure of the Mérida Andes, Venezuela: relations with the South America-

Caribbean geodynamic interaction. Tectonophysics 345. 299-327. Cardona-Molina, A., U.G Cordani and W.D MacDonald, 2006. Tectonic correlations of pre-Mesozoic crust from the northern

termination of the Colombian Andes, Caribbean region, Journal of South American Earth Sciences Volume 21, Issue 4, Tectonic evolution of the Colombian Andes, Pages 337-354.

Cobbold, P.R., E. A. Rossello, P. Roperch, C. Arriagada, L.A Gómez & C. Lima (2007). Distribution, timing, and causes of Andean deformation across South America. Geological Society, London, Special Publications; v. 272; p. 321-343.

Chicangana, E. 2005. The Romeral Fault System: A Shear and Deformed Extinct Subduction Zone Between Oceanic And Continental Lithospheres in Northwestern South America. Earth Sci. Res. J. Vol 9, No. 1: 51 -66.

Colletta, B., F. Roure, B. De Toni, D. Loureiro, H. Passalacqua and Y. Gou. 1997. Tectonic inheritance, crustal architecture, and contrasting structural styles in the Venezuelan Andes, Tectonics 16 (5), pp. 777–794.

Díaz de Gamero, M.L. 1996. The changing course of the Orinoco River during the Neogene: a review. Palaeogeography, Palaeoclimatology, Palaeoecology, Volume 123, Issues 1-4, Pages 385-402.

Escalona, A. & P. Mann. 2006. Tectonic controls of the right-lateral Burro Negro tear fault on Paleogene structure and stratigraphy, northeastern Maracaibo Basin. AAPG Bulletin 90: 479-504.

Hoorn,C., J. Guerrero, G.A. Sarmiento, and M.A Lorente. 1995. Andean tectonics as a cause for changing drainage patterns in Miocene northern South America. Geology, 23 (3): 237-240.

James, K.H. 2000. The Venezuelan Hydrocarbon Habitat, Part 1: Tectonics, Structure, Palaeogeography and Source Rocks. Journal of Petroleum Geology, 23 (1), pp 5-53.

Kohn, B., Shagam, R., Banks, P., Burkley, L., 1984. Mesozoic–Pleistocene fission track ages on rocks of the Venezuelan Andes and their tectonic implications. Geological Society of America. Memoir 162, 365– 384.

Mora, J., D. Loureiro and M. Ostos, 1993. Pre-Mesozoic Rectangular Network of Crustal Discontinuities: One of the Main Controlling Factors of the Tectonic Evolution of Northern South America. AAPG/SVG International Congress and Exhibition, Caracas, Abstracts, 58.

Pindell, J. L. & Dewey, J.F., 1982, Permo Triassic reconstruction of western Pangea and the evolution of the Gulf of Mexico/Caribbean region: Tectonics, v. 1, p.179-211.

Shagam, R., Kohn, B., Banks, P., Dasch, L., Vargas, R., Rodrıíguez, G., Pimentel, N., 1984. Tectonic implications of cretaceous–pliocene fission track ages from rocks of the circum-Maracaibo Basin region of western Venezuela and eastern Colombia. Geological Society of America. Memoir 162, 385– 412.

Soulas, J. P., 1983. Tectónica cuaternaria de la mitad Sur de los Andes Venezolanos - Grandes rasgos. XXXIII Convención AsoVAC. Caracas. XXXIV. Acta Científica Venezolana, (1): 525. Resumen.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 84-87

84

Seismicity study of the Ecuadorian margin, using combined inshore-offshore seismological network

Nicole Béthoux1, Bernard Pontoise

1, Viviana Alvarez

2, Mónica Segovia

2, Jean-Yves Collot

1,

Philippe Charvis1, Yann Hello

1, Kevin Manchuel

1, Marc Régnier

1, Yvonne Font

1, Jordi Díaz

3,

Antonio Villaseñor3, & Audrey Gailler

1

1

Géosciences Azur, Université Nice-Sophia Antipolis, IRD, Observatoire Côte d’Azur, Quai de la Darse, 06235

Villefranche-sur-Mer, France (font@geoazur.obs-vlfr.fr) 2

IG, EPN, Ladron de Guevara E11-2534, Quito, Ecuador (msegovia@igepn.edu.ec) 3

Institut Jaume Almeira, Barcelone, Spain (jdiaz@ija.csic.es)

Introduction

Accurate study of offshore earthquakes is a long-time challenge in the scientific community and this problem

is particularly important in subduction regions (Husen et al, 1999). So far, location of offshore events suffers the

lack of seismological marine stations deployed during a long time on the seafloor. Therefore, earthquakes are

only located using land seismological network, often installed far away from the epicentral area. On another part,

OBS (Ocean Bottom Seismometers) are generally used for wide-angle seismic experiment and installed during a

short time, along 2D profiles. These OBS record not only the marine active shots but also the natural seismicity.

We present here results obtained combining active and passive seismology data and/or passive data recorded

both by OBS and land stations

The region under study is the Ecuadorian margin. There, three surveys were already performed. The first one,

was the campaign “SUBLIME”, in 1998. 15 OBSs and 10 land stations were deployed during a three week

period in the region of Esmeraldas. The seismic campaign SISTEUR (August - September 2000) took place

offshore Ecuadorian and south Colombian coasts. In the frame of this experiment, a network of 24 OBS was

deployed on the southern flank of the Carnegie ridge (Fig. 1), distributed in a principal axe perpendicular to the

trench and in two branches parallel to the latter. This marine network was complemented with a land network of

20 stations distributed in two lines: one parallel to the margin and the other perpendicular to it. This land-sea

network recorded the wide-angle shots produced by an air gun and the seismicity of the zone during three weeks.

The third one, the project “ESMERALDAS” was carried on in 2005 from 10 March to 14 June 2005. 26 OBSs

and 33 land stations were deployed during this period.

The scientific objectives of these experiments are included in a international research program, on the

characterization of seismic and gravity hazards of the active subduction margin of Ecuador- Colombia (Collot et

al., 2002). One first aim was to resolve the detailed structure of this convergent margin and to characterize the

interplate contact geometry, mainly by seismic study (Collot et al, 2005; Sage et al., 2007). The objective of the

passive experiment “SUBLIME” was a first trial to better locate the seismic events of the margin. Thanks to the

data of “ESMERALDAS” experiment the active deformation of the margin is going to be characterized by a 3D

tomography both active (Garcia et al, this issue) and passive (Manchuel et al., this issue). Complementary

seismological studies can be lead using the waveform of recorded seismic events in order to obtain information

on seismic sources. Use of broad-band seismometers in OBS stations allows innovant study of marine records.

In this paper, we first focus on the results obtained with data from “SISTEUR” experiment, then we present

some results about “ESMERALDAS” project. Location of events recorded during SUBLIME and

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 84-87

85

“ESMERALDAS” experiments are presented in another study (Manchuel et al., this issue). We present here the

first results about waveform analysis of some seismograms recorded in 2005 by this combined land and marine

network.

“SISTEUR” results

The high quality of shot data allowed to build constrained velocity models perpendicular to the margin on 200

km length and models parallel to the margin along 130 km length. Details of this work are given in two previous

papers (Graindorge et al., 2003, Gailler et al, 2007). One structural characteristics of this area is the presence of a

low-velocity zone at the bottom of the upperplate corner. The main features of velocity structure are known.

oceanic crust velocity on the upper plate, evidence of the sedimentary Manabi basin and slope of the interplate

channel. In order to better locate the seismic events which occurred near the network, we built a 3D grid

obtained with the projection of the 2D velocity models in perpendicular directions extrapolated onshore from one

part to the other of the main perpendicular profile. However, we took into account the small lateral variations

evidenced thanks to the profiles parallel to the margin

Despite realistic velocity models, the configuration of the network forgives reliable location outside the short

range area defined previously. So, we studied the waveform of available records in order to better constrain the

location of some events. Indeed, at regional distances, the waveform is mainly linked with the hypocentral

parameters, and in a second order to the focal mechanism (Bertil et al., 1988). We built synthetic waveform

using the discrete wave-number method implemented by Bouchon and Aki (1977). We used the code modified

by Coutant (1994) who replaced the computation of wave propagation at the interface obtained with Thompson-

Haskell methodology, by a matrix computation of reflection and transmission coefficient at each interface of the

1D velocity model. The so-called AXITRA code computes in the frequency domain the Green solutions,

depending of the hypocentral coordinates, the position of the station respect to this hypocentre, and crustal model

(velocity, density, Q factors and thickness of each layer). These Green solutions are then convolued with the

source function, depending of the focal mechanism and seismic moment M0. Afterwards, The seismogram is

compared with the observed record in the time domain.

In our case study the model is strongly 2D depending of the position in a west-east azimuth. So, in order to

validate a condition of ~1D model between the source and the receptors we limit the computation to ray path

between source and stations, which are roughly parallel to the margin. The crustal model used for this modelling

depends of the position of the ray path respect to the margin.

Even if the period of recorded was very short, some observations can be deduced from these results. The

seismicity is located both in the upperlying plate and in the downgoing margin, with depth increasing from the

trench up to the eastern direction.

Some events are located just at the intraplate boundary, as deduced from the wide-angle modelling. The

minimum depth of these events are ~10km , which can be the upper limit of the ISZ around the latitude of 1.4°S .

Then the hypocenters get deeper up to 35 km at a distance of ~100 km of the trench. However the main result is

the presence of a swarm clearly separated in two parts one in the upperlying plate, the second in the downgoing

crust. This swarm location corresponds to the trace of the Jipijapa fault (Daly, 1989, Deniaud 2000). This major

fault is described in all the geological studies of the Ecuadorian margin. It limits the Manabi basin to the west. It

is described in the litterature as a system of faults defined as a srtike-slip duplex, due to the regional

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 84-87

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transpressive stress field. We could perform three focal solutions compatible, showing a nodal plane compatible

with a dextral fault oriented N20 up to N30. This study shows that this fault is seismically active, and suggests

that this fault is deeply rooted in the crust and reach the interplate boundary. The swarm located beneath the

fault could be the expression of an asperity due to the decoupling of the upper plate by this fault.

13 focal solutions are proposed. A compressive stress field is evidenced in the downgoing plate with a

compressive axis P orientated E-W. In the upperlying plate, normal solutions correspond to rupture planes

orientated NW-SE. Among them, the two largest magnitude recorded during the experiment are located near

the Bahia fault. The other are located in the Manabi basin. The coexistence of these two types of mechanisms

is coherent with a transtensive stress-field due to the convergence of Nazca and South American plates and

escape of the North-Andin block towards the NNE.

Figure 1. a- Location of the three temporary networks deployed in the Ecuadorian margin. b- Some focal solutions deduced from polarity readings and waveform analysis for seismic events between Bahia de Caraquez and Manta region. c- Zoom on the coastal region and other focal solutions for this region.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 84-87

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“ESMERALDAS” results

The same methodology was used for waveform modeling of some seismograms recorded during

“ESMERALDAS” experiment, both by broad-band OBS and by land seismological stations. We present here the

study of events located on the margin.

The discrepancy of waveforms recorded in different azimuthal directions allows to constrain the focal

mechanism and source parameters . We also show the influence on the waveform of the different crustal model

around the receivers.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 88-89

88

Geology of the Río Seco region, Deseado massif (48°35´S), Santa Cruz province, Argentina

Pedro P. Biscayart1, Daniel J. Pérez

1, Leandro E. Echavarría

2, & Maria José Correa

2

1 Laboratorio de Tectónica Andina, Departamento de Ciencias Geológicas, Universidad de Buenos Aires, Ciudad

Universitaria 1428, Buenos Aires, Argentina (daniel@gl.fcen.uba.ar) 2 MHA SA. Emilio Civit 355, Mendoza CP5500, Argentina

KEYWORDS : Chon Aike, Jurassic, Deseado Massif, Santa Cruz, Argentina

Introduction

This work’s purpose is to learn about the structure and tectonic evolution of the Río Seco region, based on the

mapping of the area. The Río Seco region is located southeast of Cerro Vanguardia, in the geological province

known as Deseado Massif ( 48°35´ S and 65°10´ W), belonging to Santa Cruz province, Argentina. The

stratigraphy of this part of the Massif is principally Jurassic volcanic rocks, associated to the evolution of the

meso-atlantic ridge, followed by a series of Mesozoic continental sequences and Cenozoic marine ones.

Stratigraphy and structure

The stratigraphic sequence of the area begins with the continental deposits of El Tranquilo Formation. They

consist mainly of sand and shells, which were referred to the middle-upper Triassic based on their fosiliferous

contents (Panza, 1994). The Triassic sequences are followed by basalts of the Bajo Pobre Formation which

belong to the middle Jurassic. These are intermediate to basic rocks of andesitic composition and calc-alkaline

affinities (Guido et al., 2006).

Acid volcanic rocks from the middle to upper Jurassic age, represented by the Bahía Laura Group, are found

overlaying the previous units (Guido, 2004). These are basically rocks of rhyolitic composition rich in

potassium, which belong to the calc-alkaline trends and are of peraluminous type (Pankhurst y Rapela, 1995).

The Chon Aike and La Matilde formations, which interdigitate laterally, are part of this Group (Panza, 1994).

The Chon Aike Formation is made principally of ignimbrites and tuffs. Due to its extension and importance, it

has been divided into three members: lower, middle and upper Chon Aike. La Matilde Formation consists of

tuffs and tuffites which present planar tabular bedding in some areas. The Jurassic volcanism contains

epithermal mineralization of low sulfuration of Au-Ag that lodges in veins with preferential directions WNW

and NNW (Schamaluk et al., 1999, Echavarría et al., 2005). Over the Bahía Laura Group lay, in angular

unconformity, the tuffs, sand, silt and conglomerates that form the Baqueró Formation. These lower Cretaceous

continental deposits (Panza, 1994) lay horizontally over the Jurassic volcanic units, previously described. The

Cenozoic sequences begin with coquina, sand and silt shells from the Monte León Formation. These sedimentary

rocks known as “Patagoniano”, are the result of an Atlantic marine transgression in both, the San Jorge Gulf and

Austral basins, that covered most of Deseado Massif (Panza, 1994). This extended unit overlays, in erosive

unconformity, the Chon Aike, La Matilde and Baqueró formations. An upper Oligocene age is accepted for this

Formation, based on mega and microfauna studies (Panza, 1994). Over the previous sequences and in erosive

unconformity, gravels and sands accumulated. This accumulation is known as La Avenida Formation, assigned

to the upper Pliocene – lower Pleistocene (Panza, 1994). In some areas, this Formation is covered by La Angelita

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 88-89

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olivinic basalts, aged middle to upper Pleistocene (Panza, 1995). Finally, the quaternary units are widely

distributed. Fluvial terrace deposits, eolic and alluvial deposits and thin sediments from lowlands and lagoons,

have been identified.

The structure of the region is intimately related with the deformation and block faulting, indicating a thick skin

structural style. This is the result of extensional movements during the Mesozoic era. The high angle normal

faults with strike-slip movement might correspond to Permo-Triassic rift reactivations. During Jurassic, the

extensional movements were simultaneous with the eruption of volcanic rocks, suggesting the development of

ENE-WSW grabens (Echavarría et al., 2005).

In the Río Seco valley, a series of fault blocks with NNW orientation can be distinguished. Several evidences

indicate that during lower Cretaceous the passage from an extensional regime to a transtensional-compresional

one, could have occurred (Reimer et al., 1996). However, an extensional reactivation is registered during the

Aptian, based on the eruption of volcanic rocks of the Baqueró Formation (Ramos, 2002).

Several zones with quartz and silicification veins where recognized. Veins show different textures and have

NNW-SSE orientation, controlled by normal faults. Also, different cinematic indicators are found in the veins

with direction N 100° to N 120°, which show that during the upper Jurassic, the extension could have been

accompanied by dextral strike-slip movements. In the area under study, some small folded structures are found.

Generally they present very low inclinations of their sides.

Acknowledgments Thanks to MHA SA for supporting this work, and to UBACYT-X160 for financiation. References Echavarría, L.E., Schalamuk, I.B., Etcheverry, R.O., 2005. Geologic and tectonic setting of Deseado Massif epithermal

deposits, Argentina, based on El Dorado-Monserrat. Journal of South American Earth Sciences 19: 415 – 432. Guido, D.M., 2004. Subdivisión litofacial e interpretación del volcanismo jurásico (Grupo Bahía Laura) en el este del Macizo

del Deseado, provincia de Santa Cruz. Revista de la Asociación Geológica Argentina, 59 (4): 727 – 742. Guido, D., Escayola, M., De Barrio, R., Schalamuk, I., Franz, G., 2006. La Formación Bajo Pobre (Jurásico) en el este del

Macizo del Deseado, Patagonia: vinculación con el Grupo Bahía Laura. Revista de la Asociación Geológica Argentina 61 (2): 187 – 196.

Panza, J.L., 1995. Descripción geológica de la Hoja 4969-II, Tres Cerros, provincia de Santa Cruz, Secretaría de Minería de La Nación. Boletín n° 213, Buenos Aires, Argentina 1995 p. 103.

Pankhurst, R.J., Rapela, C.R., 1995. Production of Jurassic rhyolite by anatexis of the lower crust of Patagonia. Earth and Planetary Science Letters 134: 23 -36.

Ramos, V.A., 2002. Evolución tectonica de la plataforma continental. XIII° Congreso Geológico Argentino y III° Congreso de Exploración de Hidrocarburos (Buenos Aires, 1996). Geología y Recursos Naturales de la Plataforma Continental Argentina, V. A. Ramos y M. A. Turic (Eds.), Velatorio 21: 385 – 404.

Reimer, W., Millar, H., Mehl, H., 1996. Mesozoic and Cenozoic palaeostress fields of the southern Patagonian Massif deduced from structural and remote sensing data. En: Storey, B.C., King, E.C., Livermore, R.A., (Eds.), Weddell Sea tectonics and Gondwana break-up, vol. 108 Geological Society Special Publication, pp. 201 – 263.

Schalamuk, I. B., de Barrio, R. E., Zubia, M., Genini, A. y Echeveste, H., 1999. Provincia auroarentifera del Deseado, Santa Cruz. En: Zappettini, E. (Ed.): Recursos Minerales de la República Argentina, SEGEMAR, Anales 35: 1177-1188. Buenos Aires.

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Electrical conductivity beneath the Payún Matrú Volcanic Field in the Andean back-arc of Argentina near 36.5°S: Insights into the magma source

Aurora Burd1, John R. Booker

1, M. Cristina Pomposiello

2, Alicia Favetto

2, Jimmy Larsen

1,

Gabriel Giordanengo2, & Luz Orozco Bernal

2

1 University of Washington, Box 351310, Department of Earth and Space Sciences, Seattle, Washington 98195,

USA (aburd@u.washington.edu) 2 Instituto de Geocronología y Geología Isotopíca, Pabellon INGEIS, Universidad de Buenos Aires, Ciudad

Universitaria, C1428EHA, Buenos Aires, Argentina (cpomposi@ingeis.uba.ar)

KEYWORDS : Payún Matrú, electrical conductivity, upper mantle, Andean back-arc, Neuquén Basin

Introduction

Southern Mendoza and northern Neuquen Provinces, south of the Nazca flat slab in western Argentina, have

widespread, geologically young basaltic volcanism, but no historic activity. The youngest basalts, erupted in the

vicinity of the large Payún Matrú volcanic center have essentially no arc signature (Bermudez, et al., 1993).

Kay, et al. (2006) and Folguera, et al. (2006) argue that this back-arc igneous province is the result of extension

due to trench roll-back following steepening of a flat slab that existed in the middle to late Miocene.

Magnetotelluric (MT) data collected in 2005 along an east-west profile have been used to probe the source of the

Payún Matrú basalts. These data imply significantly 3D structure. However, preliminary analysis of an arguably

2D region at the center of the profile allows tentative identification of a conductive mantle plume surfacing at

Payún Matrú that rises from below 200 km depth.

Additional MT data are being collected as this abstract is being written. This new work extends the earlier

profile to a spatial array extending from Laguna Llancanelo north of Payún Matrú to beyond the Cortaderas

Lineament that bounds the basaltic province to south. This will add at least 15 new sites with significantly

higher-quality data. The new array will permit 3D interpretation. Data processing in the field suggests that the

deep crust or upper mantle has northwest-southeast striking structure increases in conductivity to the southwest

of Payún Matrú and perhaps also to the northeast. This underlies shallow structure with north-south strike

between Payún Matrú and the Colorado River (the boundary between Mendoza and Neuquen). This complexity

explains our initial difficulty with 2D interpretation.

Geological Background

The Payunia (often called Payenia) volcanic field in the northern Neuquen Basin is located in the back-arc of

the Andes, south of the Chile-Argentina Nazca flat-slab subduction zone. Wadati-Benioff zone earthquakes and

seismic “receiver function” images (Yuan, et al., 2006 and references therein) reveal that the average dip of the

Nazca slab is about 25° from 50 to 100 km depth and about 40° from 100 to 200 km. Its projected depth under

Payún Matrú is thus 300 km or more (Figure 1 below). Volcanism in Payunia, with extension over a steeply-

dipping slab, has been active for roughly the last 5 MA. During this time, the percentage of mantle melting

appears to have increased and the slab influence on the geochemistry has declined (Kay, et al., 2006). For about

15 MA prior, the stress regime was compressional and volcanism showed significant slab influence, suggesting

that the slab was much shallower (same paper). Kay & Copeland (2006) have drawn an analogy between the

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Payunia situation in the middle Miocene (of order 15 MA before the present) and the volcanism at Pocho and in

the Sierras de San Luis over the Nazca flat-slab of about 5 MA age. For at least 5 MA prior to 20 MA, the

Payunia situation was similar to the most recent 5 MA (Kay, et al., 2006). This picture of flat-slab subduction

lasting for about 15 MA preceded and succeeded by steep subduction is compelling.

The Caldera Payún Matrú is responsible for one of the largest flows in the area, creating 5200 km2 of alkaline

basaltic lava at some point in the Pliocene or Quaternary (Ramos and Folguera, 2005). There are approximately

300 volcanic vents in the surrounding field, with no human-recorded eruptions, although a significant portion of

this activity has been designated Late Pleistocene to Holocene (Inbar and Risso, 2000). Measurements of the

degradation of the shapes of cinder cones (ibid.) suggests that the most recent eruption may have been only

about 1,000 years ago. Thus, although the magma production rate is low, it almost certainly continues today.

Intraplate volcanism such as Payún Matrú is geochemically similar to Ocean Island Basalts (OIB) which are

commonly thought to arise from the deep mantle. Berkovici and Karato (2003), however, argue that OIB need

only come from beneath 400 km, the depth at which upper mantle minerals such as olivine transform to higher

pressure phases that are capable of containing a great deal more dissolved hydrogen (i.e. water). Mid-Ocean

Ridge Basalts (MORB), which are clearly the result of partially melting shallow mantle are chemically distinct

from OIB and Payún Matrú basalts (Bermudez, et al., 1993). It is thus reasonable to conclude that the Payún

Matrú source is relatively deep. This is underscored by oil fields producing from Mesozoic sedimentary rocks

very close to Payún Matrú eruptive centers. These eruptions and, in fact, the entire Cenozoic history of

volcanism in this area, has clearly had no influence on the thermal state of the upper crust.

The Magnetotelluric Method

The magnetotelluric (MT) method uses passively recorded electric and magnetic field data at Earth’s surface to

probe electrical resistivity (or its inverse, conductivity) below the surface. Time-series collected at multiple

locations are Fourier transformed and frequency-domain transfer functions between horizontal electric and

magnetic fields are determined. The transfer function between vertical and horizontal magnetic components is

usually also determined at the same time. Inversion of the transfer functions for structure is typically made

unique by minimizing the structure (roughness) of the model, perhaps with constraints from other data.

Measurement of electrical conductivity is useful because it is strongly sensitive to transport properties.

Silicates are mostly very resistive (>10,000 Ohm-m) at the temperatures and pressures of the upper mantle and

crust. Low electrical resistivity (elevated conductivity) is due to minor conductive phases, such as hydrous

fluids, partial melt and less commonly graphite and metals or sulfides. The degree of interconnectivity of the

conducting phase profoundly affects resistivity and bulk resistivity can change by orders of magnitude.

Sediments with saline pore water can be 1 Ohm-m (sea water is 0.3 Ohm-m); sedimentary rocks are usually 3-30

Ohm-m; silicate melts are about 1 Ohm-m, so partial melts will be more resistive (10-100 Ohm-m), but still very

conductive relative to solid silicate rock.

The shallow crust of Payunia should be resistive (resistivity >100 Ohm-m) in the east where crystalline and

metamorphic basement is at or near the surface and less resistive (3-100 Ohm-m) in the west where the basalts

overlie Mesozoic and early Cenozoic sedimentary units of the Neuquen Basin and Rio Grande foreland basin.

Otherwise, the basement and mantle above 400 km should be very resistive (>1,000 Ohm). Active structures

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 90-93

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involving partial melts or hydrous fluids should stand out against this background. Below 400 km, the transition

zone can be as resistive as the upper mantle (>1,000 Ohm-m) or quite conductive (10 Ohm-m) depending on the

amount of dissolved water (Hu, et al., 1998). Below 660 km phase changes result in minerals that are

intrinsically very conductive ( 3 Ohm-m).

Data collection and processing In 2005, we collected an 18 site MT transect along 36.7°S, from 70°W to approximately 67°W. Each site

consists of 5 to 10 days of 4 Hz horizontal electric and 3-component magnetic field time-series recorded with

Narod Intelligent Magnetotelluric System (NIMS) using Pb-PbCl2 electrodes.

Time-series data were processed using the robust multi-station algorithm of Egbert (1997) to determine the MT

and vertical to horizontal magnetic field transfer functions. Considerable improvements were achieved by one of

us (Larsen) using adaptive removal of electric field residuals (due to faulty electrodes) in the time-domain. The

dimensionality of the resistivity structure was then studied using the “phase tensor” technique of Caldwell, et al.

(2004). We concluded that the overall data set requires 3D interpretation, which is not easily done with a single

profile. 2D inversions were never-the-less performed with a range of subsets of the data, assumed strikes, side

conditions and degree of model smoothness. One of these models is shown in Figure 1.

Figure 1. Preliminary 2D electrical resistivity model of Payunia MT data (3.33 s – 213 s) at the 9 sites shown as inverted triangles. Payún Matrú coincides approximately with the left four sites. The resistivity (inverse of conductivity) color scale is logarithmic. The inversion uses the code of Rodi and Mackie (2001) and maximizes smoothness of the model for a given data misfit. It is constrained to be 3 Ohm-m deeper than 660 km and has a Pacific Ocean water conductor to the west. Earthquake hypocenters are shown as small open circles scaled according to magnitude. The estimated position of the subducted Nazca slab is shown as a solid line where it is known with confidence and as a dashed line where it is extrapolated from shallower depths.

Discussion The model in Figure 1 uses periods from 3.3 to 213 s at 9 sites and assumed strike of N 15° W. It is

constrained by a conductive ocean to the west and a resistivity of 3 Ohm-m deeper than 660 km. Both

constraints have a strong effect on the structure, but neither is controversial. Major features of this model are

1 Ohm-m

1,000 Ohm-m

Depth (km)

Distance east of 68.5°W (km)

log10 resistivity

= - log10 conductivity

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only weakly dependent on our choice of data subset or strike and are thus considered “robust”. The most

interesting is the steeply-dipping conductive feature that rises benath the Payún Matrú Volcanic Field (PMVF).

This “plume” is rooted deeper than 200 km and is just above the projected Nazca Slab. One should not trust the

details of this structure at this time. What is known with considerable confidence is that a narrow, near vertical

conductive structure must connect the shallow to deep mantle. It is similar to the conductor associated with the

plunging Nazca flat-slab reported by Booker, Favetto and Pomposiello (2004). This suggests that Payún Matrú

Volcanic Field is sampling mantle deeper than 200 km, perhaps just above the subducted Nazca slab or perhaps

from the mantle transition near where the Nazca slab penetrates into it at 400 km. If transition zone melt is

involved it may explain why the chemistry is like OIB. The deeper part of rising conductive mantle appears to

extend eastward. Its exact relationship to both the Nazca slab and transition zone remains unresolved and is a

major focus of our on-going field work. We hope to be able a preliminary 3D interpretation of these new data at

the meeting.

Acknowledgements This project is supported by the U.S. National Science Foundation (NSF) Grants EAR0310113 and EAR0739116. MT

equipment is from the EMSOC Facility supported by NSF Grant EAR026538. This project also received support from the Agencia de Promocion Cientitica y Tecnologia and CONICET. The first author received support from the Seattle Chapter of the ARCS Foundation, and from Graduate Reseach support and a Anthony Qamar Fellowship provided by the Dept. of Earth and Space Sciences, University of Washington.

References Bercovici, D., & Karato, S. 2003. Whole mantle convection and the transition-zone water filter. Nature 425: 39-44. Bermúdez, A., Delpino, D., Frey, F., Saal, A. 1993. “Los basaltos de retroarco extrandinos”, in: Ramos, V. (ed.) Geologia y

recursos naturals de Mendoza, Relatorio, XII Congreso Gelologico Argentino y II Congreso Exploration de Hydrocarburos, Assoc. Geol Argentina y Inst, Argentino del Petroleo: 161-172.

Booker, J., Favetto, A., & Pomposiello, M.C. 2004. Low Electrical Resistivity associated with plunging of the Nazca flat slab beneath Argentina. Nature, 429: 399 – 403.

Caldwell, G. C., Bibby, H. M. & Brown, C. 2004. The Magnetotelluric Phase Tensor. Geoph. J. Int. 158: 457-469. Egbert, G. 1997. Robust multiple-station magnetotelluric data processing. Geoph. J. Int. 130: 475-496. Folguera, A., Zapata, T. & Ramos, V. 2006. “Late Cenozoic extension and the evolution of the Neuquen Andes”, in: Kay, S.

& Ramos, V. (eds.) Evolution of the Andean Margin: A tectonic and magmatic view from the Andes to the Neuquen Basin (35- 39S lat), Geol. Soc. Am. Special Paper 407: 267-286.

Inbar, M., & Risso, C. 2001. Holocene Yardangs in Volcanic Terrains in the Southern Andes, Argentina. Earth Surf. Process. Landforms 26, 657-666.

Kay, S., Burns, W., Copeland, P. & Mancilla, O. 2006. “Upper Cretaceaous to Holocene magmatism and evidence fro transient Miocene shallowing of the Andean subduction zone under the northern Neuquen Basin”, in: Kay, S. & Ramos, V. (eds.): 19-60.

Kay, S. & Copeland, P. 2006. “Early to middle Miocene backarc magmas of the Neuquen Basin: Geochemical consequences of slab shallowing and the westward drift of South America”, in: Kay, S. & Ramos, V. (eds.): 185-214.

Ramos, V. & Folguera, A. 2005. “Structural and magmatic responses to steepening of a flat subduction, southern Mendoza, Argentina” in: 6th International Symposium on Andean Geodynamics (ISAG 2005 Barcelona), Extended Abstracts: 59-595.

Rodi, W., & Mackie, R. 2001. Nonlinear Conjugate Gradients Algorithm for 2-D Magnetotelluric Inversion. Geophysics 66: 174-187.

Yuan, X., Asch, G., Bataille, K., Bock, G., Bohm, M., Echtler, H., Kind, R., Oncken, O. & Wölber, I. 2006. “Deep seismic images of the Southern Andes”, in: Kay, S. & Ramos, V. (eds.): 61-72.

Xu, Y., Brent, T., Poe, T. Shankland, T., Rubie, D. 1998. Electrical conductivity of olivine, wadsleyite and ringwoodite under upper-mantle conditions, Science 280, 1415-1418.

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Crustal structure and tectonic deformation of the northern Chilean margin, 21-23.5ºS

A. Calahorrano B.1, C. R. Ranero

2, U. Barckhausen

3, C. Reichert

3, & I. Grevemeyer

4

1 Institut de Ciencies del Mar-CMIMA–CSIC, Pg. Maritim de la Barceloneta 37-49, 08003, Barcelona, Spain

(alcinoe@cmima.csic.es) 2 Institució Catalana de Recerca i Estudis Avançats (ICREA), CMIMA, Passeig Maritim de la Barceloneta 37-49,

08003, Barcelona, Spain 3 BGR, Bundesanstalt für Geowissenschaften and Rohstoffe, Stilleweg 2, 30655 Hannover, Germany

4 IFM-GEOMAR and SFB574, Wischhofstrasse 1-3, 24148, Kiel, Germany

KEYWORDS : multichannel seismic data, subduction, mass-wasting erosion, outer rise deformation

This work studies the crustal structure and tectonics of the convergent margin of north Chile, between

Tacopilla and Antofagasta, where the oceanic Nazca Plate (50 Ma) subducts sub-orthogonally below the South

American Plate at ~80-90 mm/yr (DeMets et al., 1990; Clift and Vannucchi, 2004). Here we focus on three

reprocessed multichannel seismic (MCS) reflection lines (SO104-7, SO104-9 and SO104-13) that were acquired

during the CINCA'95 experiment (Figure 1). These lines, acquired using a 3-km-long streamer with 120

channels and a 3,124 cc tuned air gun source, run perpendicular to the coast along ~450 km, imaging the

upper/middle slope of the overriding plate, the trench and some ~300 km of the oceanic downgoing plate (Figure

2A).

Figure 1. Study area and bathymetric map of the northern Chile. White lines correspond to location of multichannel seismic lines SO104-7, SO104-9 and SO104-13. White points represent the location of Antofagasta (A) and Tocopilla (T) cities.

In the oceanic domain, the mcs lines show a seafloor practically deprived of sedimentary coverage. In contrast,

the seafloor is characterised by a strong and continuous reflection that we associate to the top of the oceanic

crust (Figure 2B). This reflection changes from smooth to irregular depending on the mcs line and the distance to

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 94-96

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the trench. Lines SO104-9 and SO104-13 show an irregular reflection that traduces the roughness of the seafloor

resulting from the presence of the NE trending Iquique ridge.

Figure 2. A. CINCA SO104-7 seismic images. A) Time-migrated section of a portion of Line SO104-7. This image shows part of the oceanic Nazca Plate which is subducting below the South American Plate (SAP). In the outer rise is concentrated most of the extensive deformation of the downgoing plate resulting from plate bending prior to subduction. B) Enlarged segment of Line SO104-7. The strong and continuous reflection imaging the seafloor correspond to the top of the oceanic crust. Sediment deposits are practically absent. This image shows landward dipping reflections below the crust-mantle boundary (Moho). Note small normal failure affecting the seafloor in the smoother segments of the ocenic crust. C) Enlargement of the segment corresponding to the margin front. Triangles follow the main reflection associated to the inter-plate contact.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 94-96

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In contrast, line SO1047 is smooth far form the trench, and the influence of the Iquique ridge seems to be

minor. Near the trench, the three lines show irregular bottom morphology, due to a high deformation mainly

resulting from: (1) NW-oriented fabric inherited of spreading during plate generation and (2) the NS horst-and-

graben pattern and normal faulting associated to plate bending nearby the trench (Figures 2A and C).

A second strong reflection, identified around 2s two-way-travel time (stwt) below the seafloor, is associated to

the crust-mantle boundary (Moho). It is quite continuous and clear below soft seafloor reliefs as in Line

SO1047). In line SO104-9, this reflector is locally ~1s deeper twt below the seafloor, at the location of the

Iquique ridge, where an over-thickened crust is expected. Considering a crustal velocity range of 6.5-7 km/s, the

crust would be ~6.5-7 km thick at the segments of normal oceanic crust, and ~10 km at the thickened crustal

segments.

Another feature, which is particularly clear in Line SO104-7, is the presence of landward dipping reflections

pattern below the Moho (Figure 2B). These reflections, which are strong and clear near the trench and disappear

seawards, may be associated to discontinuities, probably compositional, of the asthenosphere.

At a greater scale, it is possible to observe the flexure of the oceanic plate near the trench, produced by bending

prior to subduction. This flexure characterise the outer rise where most of the bending deformation is located.

The top of the oceanic crust reflection preserves the horst-and-graben undulation below the margin, and the

inter-plate contact is observed until ~50 km landward from the trench (Figure 2C). As irregular topography

subducts, basal tectonic erosion is likely to occur below the margin, damaging the bottom of the overriding-plate

basement and removing upper plate material that feeds the subduction channel.

The trench region consists on subducting grabens or half-grabens with some turbiditic desposits (Figure 2C).

The overriding plate shows two slope breaks constraining the upper, middle and low slopes. In the upper slope

the sedimentary coverage is imaged as a finely stratified layer that truncates in the upper/middle slope break. The

top of the overriding-plate basement is interpreted to be a strong reflection that separates the sediment package

and a low-frequency seismic facies body displaying a coarser stratification.

In the upper slope, line SO104-9 images landward dipping normal faults perturbing the basement, sediments

and the seafloor. In the middle and lower slopes sediment are scarce, and normal faulting changes its dip

direction from landward to seaward (lines SO104-9 and SO104-13). These faults indicate that the frontal margin

shoud be dominated by active extension tectonics. Normal faulting and loss of sediment coverage suggest mass

wasting erosion of the frontal margin following the tectonic structure proposed by von Huene and Ranero 2003.

The eroded debris may accumulate at the margin’s toe explaining the presence of an incipient <10 km-wide

sediment prism at the deformation front, but they may also fill the oceanic plate grabens and feed the subduction

channel. The sedimentary prism may eventually be also eroded and be involved into subduction.

References Clift, P. D. and P. Vannucchi. 2004. Controls on tectonic accretion versus erosion in subduction zones: Implications for the

origin and recycling of the continental crust. Reviews of Geophysics 42: 1-31. DeMets, C., R. G. Gordon, D.F. Argus and S. Stein. 1990. Current plate motions. Geophysical Journal International 101:

425-478. von Huene, R. and C. R. Ranero. 2003. Subduction erosion and basal friction along the sediment-starved convergent margin

off Antofagasta, Chile. Journal of Geophysical Research 108(B2): 3-1 3-16.

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Preliminary stratigraphic study of the San Francisco River volcanic sequence, northwestern Purace volcano, Cauca, Colombia

E. Cañola1, S.M. López

1, G. E. Toro

1, & B. Pulgarín

2

1 Universidad EAFIT, Carrera 49 N° 7 Sur -50 Medellin, Colombia (ecanolah@eafit.edu.co;

slopezc1@eafit.edu.co; gtoro@eafit.edu.co) 2 INGEOMINAS, Popayán, Colombia (bpulgarin@ingeominas.gov.co)

KEYWORDS : stratigraphy, Cenozoic volcanism, Purace volcano, magmatic evolution

Introduction

The Purace volcano is part of recent volcanism in the Colombian Andes with 20 actives volcanoes, at the

Northern Volcanic Zone (Hanke and Parodi, 1966; Mendez, 1989). Subduction of the Nazca oceanic plate

beneath the Suramerican plate, generates the magmas that formed the volcanism Cenozoic. The geometric

features at this volcanic zone are: dip of Waddati- Benioff zone is about 25° (e.g. Pennington, 1981, Pilger,

1983), young subducted oceanic plate (10 – 14 Ma, Hardy, 1991) and continental crust with 35- 40 km thick

(e.g. Case et al., 1971; Meissner et al., 1976). According to Aspden et al. (1987), the current activity is probably

a continuation of the magmatic event that generated the Miocene-Pliocene volcanism recorded in the Colombian

Andes.

We have chosen to study the area located in the northwestern flank of the Purace volcano, where several

pyroclastics flows deposit interbedded with lava flows outcrops along the San Francisco River located close to

the Purace town in the Cauca department (Colombia). Our data will contribute to the knowledge of the spatial

and temporal volcanic distribution that originated the volcanism of the Coconucos volcanic chain (CVLC). It is

important to see the change in the volcanic behavior, where explosive events are interbedded with effusive

events.

Geological and tectonic setting

The basement under this volcanic suite is mainly Paleozoic metamorphic rocks, composed by schist grouped in

Cajamarca Complex (Maya y González, 1995). To the western part outcrops the Quebradagrande Complex, this

consists of volcanic and Albian–Aptian sedimentary rocks. The current volcanic activity at this area, is related

with the Coconucos volcanic chain (CVLC), which is formed by 15 eruptive centers aligned NW-SE. (Monsalve

y Pulgarín, 1995). The CVLC is located between two large faults systems: Algeciras and Chusma Faults at east

and Romeral Fault at west (Fig. 1). There are others structures located oblique to the main fault systems

(Velandia et al, 2005). Some authors (e.g. Bohórquez et al., 2003) proposed that the oblique structures had

allowed the emplacement of the Neocene-Quaternary volcanism.

Several volcanic structures have been identified: Paletara caldera (Torres et al, 1999), Chagartón caldera and

Pre-Puracé volcano (Monsalve, 1991; Monsalve, 2000). Paletara caldera is by far the largest caldera structure

with more than 35 km of diameter (Fig. 1). It is believed that the formation of such structure could be associated

to the generation of the ignimbrite deposits of the Popayan and Guacacallo Formations (Torres et al, 1999 and

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Torres in prepared) located to the west and east of the study area. The geological map (fig 1) was realized with

field work and other information (Monsalve y Pulgartín, 1995; Monsalve, 2000).

Figure1. Geological map at northwestern Purace volcano, around San Francisco River. Modified from (Monsalve y Pulgarín, 1995; Monsalve, 2000).

Stratigraphy

The study of a volcanic sequence that outcrops along the San Francisco river, shows different volcanic events

that forms these deposits (fig 2). There is a change in the volcanic behavior where explosive events are

interbedded with effusive events, related with lava flows. At the moment, we have been concentrated in the

description of the pyroclastics flow deposits which are mainly ignimbrites and blast directed pyroclastic flows

deposit (fig 2). The main events are: (1) Coconucos´s ignimbrite: Are located in the bottom of the sequence with

150 m thick. This unit has a high percentage of biotite (25%) which may indicate a hydrous-rich magma source

and represent an intracalderic facies associated to Paletara´s caldera, which is correlated with the huge ignimbrite

found in other places at Cauca department (Torres in prepared) . (2) Unsorted deposits with massive texture,

with 20 m thick, composed by large non-vesicular lithic blocks. It is a block and ash flow deposit and may

represent a dome-collapse associated to the Chagartón´s caldera. (3) Pumice and ash flow deposit, with 60 m

thick and high modal abundance of hornblende (20%). (4) The final event registered is scoria and ash flow

deposits, with 60 m thick and matrix with high percentage of pyroxene crystals (30%). The special composition

of this last event indicates a water-poor magma source which may be different source to the water-rich ones at

the bottom of the sequence. This deposit has been associated with the final stage of Pre-Purace structure

(Monsalve et al, 1991) and a 14C age of 30.000 years BP was obtained for this deposit (Monsalve et al, 1991).

At the moment, the stratigraphic sequence presented here has been determined by field work correlations. A

detailed age dating determination is carrying out using fission track method in zircon grains for the most

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representative volcanic events (López in prepared).

Figure 2. Generalized stratigraphic at the San Francisco river, northwestern Purace volcano.

Conclusions

The stratigraphic sequence along the San Francisco River, had allowed to identify large explosive volcanic

periods interbedded with effusive-type events, which could indicated that the volcanic behavior has been

variable on the time. Variations in the modal composition, the modal abundances of biotite, amphiboles and

pyroxenes from the bottom to the top of the sequence, could be an important evidence for the changing

conditions at the magma chambers which may result in different volcanic styles pyroclastic flow deposits. The

current results are indicating a complex magma process in the formation of this type of highly explosive

intermediate to silicic magmas, where the interaction of fluid-rich mantle-derived magmas assimilates the lower

continental crust (e.g. Marín-Cerón, 2007; Marín-Cerón et al., 2008) in different proportions, producing magmas

highly water saturated, following small or advanced differentiation processes. The complex structural conditions

at this volcanic region may play an important role in the variations in the volcanic behavior. There is a potent

pyroclastic flow deposit (60 m thick) at the top of the sequence with a 14C age of 30.000 year showing a

important hazard volcanic in the area.

Acknowledgments

This work has been realized thanks to the research project Ingeominas and EAFIT University. We thank M. I. Marín-Cerón for suggestions to improve the manuscript and M. L. Monsalve for suggestions in the field’s works.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 97-100

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References Aspden, J.A., McCourt, W.J., Brook, M., 1987. Geometrical control on subduction-related magmatism: the Mesozoic and

Cenozoic plutonic history of Western Colombia. Journal of the Geological Society, London 144: 893–905. Bohórquez, O.P., Monsalve, M.L., Velandia, F., Gil-Cruz, F. 2003. Determinación del marco tectónico regional para la

cadena volcánica más septentrional de la cordillera central de Colombia. Memorias IX Congreso colombiano de geología. Case, J.E., Duran, S.L.G., López, R.A., Moore, W.R., 1971. Tectonic investigations in western Colombia and eastern

Panamá. Geological Society of America Bulletin 82 (10): 2685–2712. Hanke G., Parodi I., 1966. Catalogue of the active volcanoes of the worl including solfatara fields. Part XIX. International

Association of Volcanology, 73 pp. Hardy, N.C., 1991, Tectonic evolution of the easternmost Panama Basin: some new data and interferences, Journal of South

America Earth Science, 4(3): 261-269. Marin-Cerón, M.I., 2007. Major, trace element and multi-isotopic systematics of SW Colombian volcanic arc, northern

Andes: Implication for the stability of carbonate-rich sediment at subduction zone and the genesis of andesite magma. Doctor’s thesis. Graduate School of Natural Science and Technology Okayama University, Japan, 123 p.

Marin-Cerón, M.I, Moriguti, T., Nakamura, E., 2008. Andesite magma generation at the Plio-Quaternary SW Colombian volcanic arc. In this symposium: 7th International Symposium on Andean Geodinamics.

Méndez, R.A., 1989. Catálogo de los volcanes activos de Colombia. Bol. Geol. 30(3), Ingeominas, Bogotá, 75 p. Meissner, R.O., Flueh, E.R., Stibane, F., Berg, E. 1976. Dynamics of the active plate boundary in southwest Colombia

according to recent geophysical measurements. Tectonophysics 35: 115- 136 Maya, M., González, H. 1995. Unidades litodémicas en la Cordillera Central de Colombia. Boletín geológico Ingeominas. 35

(2- 3): 43- 53. Monsalve, M. L., 1991. Geoquímica y episodios de episodios tipo San Vicente en el volcán Puracé. Boletín geológico

Ingeominas 33: 3-16 Monsalve, M.L., Pulgarín, B. 1995. Cadena volcánica de los Coconucos (Colombia) centros eruptivos y productos recientes.

Boletín geológico Ingeominas. (37) 17- 51. Monsalve, M.L, 2000. Catalogo de las vulcanitas Neógenas de Colombia, Fascículo Formación Coconuco. Informe interno

Ingeominas. Murcia, A., Pichler, H. 1981. Geoquímica y dataciones radiométricas de las ignimbritas cenozoicas del Sur de

Colombia.Revita CIAF 6(1-3): 343- 363. Pennington, W.D., 1981. Subduction of the eastern Panama basin and seismotectonics of northwestern South America.

Journal of Geophysical Research 86 10: 753– 770. Torres, M. P., Monsalve, M.L., Pulgarín, B., Cepeda, H., 1999. Caldera de Paletará: Aproximación a la fuente de las

ignimbritas del Cauca y Huila. Boletín geológico Ingeominas. (37) 1- 15. Torres, M.P., Ibañez, D.G., Vasquez, E.J., 1992. Geología y estratigrafía de la Formación Popayán. Informe interno

Ingeominas Popayán. Pilger, R.H. 1983. Kinematics of the South American subduction zone from global plate restructions. Geodynamics of the

plate restructions. Geodynamics of the Eastern Pacific Region, Caribeean and Scotia Arcs. American Geophysical Union Geodynamics Series. 9: 113- 126.

Velandia, F., Acosta, J., Terraza, R., Villegas, H. 2005. The current tectonic motion of the Northern Andes along the Algeciras Fault System in SW Colombia Tectonophysics 399: 313– 329

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 101-104

101

Geochemical characterization of Volatile Organic Compounds (VOCs) in fluid discharges at Copahue volcano (Argentina)

F. Capecchiacci1, F. Tassi

1, O. Vaselli

1,3, A. Caselli2, & M. Agusto

2

1 Department of Earth Sciences, Univ. of Florence, Via G. La Pira, 4, 50121 Florence, Italy

(francesco.capecchiacci@unifi.it; franco.tassi@unifi.it) 2 Departamento Ciencias Geológicas, Facultad de Ciencias Exactas y Naturales, Universidad de Buenos Aires,

Ciudad Universitaria, Pabellón 2, 1428EHA, Buenos Aires, Argentina (acaselli@gl.fcen.uba.ar;

magusto@gl.fcen.uba.ar) 3 CNR-IGG, Institute of Geosciences and Georesources, Via G. La Pira, 4, 50121 Florence, Italy

(orlando.vaselli@unifi.it)

KEYWORDS : VOCs, Copahue volcano, hydrothermal system, fumarolic discharge, fluid chemistry

Introduction

Thermal fluid discharges from geothermal and volcanic systems are characterized by the presence of volatile

organic compounds (VOCs) at relatively low concentrations (up to few tens of μmol/mol) (Capaccioni et al.,

1993; Darling 1998; Tassi, 2004). In these environments hydrocarbon compounds are generally produced by i)

degradation of organic material, mainly buried in sedimentary formations, through bacteria-driven (biogenic)

reactions at low temperature (<150 °C) and ii) thermogenic processes (at 150-350 °C), such as thermal cracking

and catalytic reforming involving both pre-formed organic compounds and inorganic gas species, i.e. CO, CO2

and H2 (e.g. Des Marais et al., 1981; Mango, 2000; Taran and Giggenbach, 2003). Several authors (Capaccioni

and Mangani, 2001; Tassi et al., 2005a) have studied the behavior of reactions regulating the relative

concentrations of the light (C2-C3) alkenes-alkanes pairs at hydrothermal conditions demonstrating their possible

use as geoindicators for geothermal prospection and volcanic sourveillance (Capaccioni et al., 2005; Tassi et al,

2005b; 2007). However, few data are presently available for studies aimed to the understanding of the fate of

more complex organic compounds (>C5) in naturally discharged fluids. In this work, the compositional features

of the organic gas fraction from the fumarolic fluids discharged at the foothill of Copahue volcano (Argentina),

an active system pertaining to the Andean Southern Volcanic Zone (ASVZ), are presented and compared with

those from 1) low-temperature sedimentary environment, 2) geothermal areas and 3) active volcanic systems in

order to investigate the degradation processes of organic matter at different thermodynamic conditions.

Geological setting

The volcanic activity at Copahue (37º45’S-71º10.2’W, 2977 m a.s.l.), nested in the Caviahue-Copahue

Volcanic Complex (CCVC; Argentina-Chile) (Fig.1), started in the Pliocene. In the last 250 years 12 low-

magnitude phreatic and phreato-magmatic eruptions (Naranjo and Polanco, 2004) have marked its volcanic

history. During the last eruptive cycle, opened in July 1992, two major phreato-magmatic eruptions (September

1995 and July-October 2000) have occurred. Presently, the volcano summit consists of nine NE-oriented craters

and the active one hosting a hot (up to 42 °C) hyperacidic lake (Copahue Lake; Varekamp et al., 2001; Caselli et

al., 2005). Several thermal discharges are located in the northern flank of the volcanic edifice, where a thermal

health spa is placed.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 101-104

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Sampling and analytical method

The Copahue fluid discharges include hot mud-, boiling- and bubbling-pools, thermal springs, fumaroles and

areas of diffuse degassing. The gas samples, collected in November 2006 and February 2007, consist of 6

fumaroles, 3 bubbling pools and a 1,241 m deep well of the local (dismissed) geothermal plant, and analyzed for

the determination of the chemical composition of the main constituents using the procedure described in

Montegrossi et al. (2001). Gas samples for the determination of VOC composition were collected into pre-

evacuated 12 mL glass vials equipped with pierceable rubber septum (Labco ExetainerR) after vapor

condensation. The organic volatiles were analyzed by gas-chromatography (Thermo Trace GC Ultra) coupled

with mass spectrometry (Thermo DSQ) for analytical separation and detection of VOCs in the mass range of

40-400 m/z. The pre-concentration and the introduction of the sample were carried out by using a manual SPME

(solid-phase micro-extraction) device (Supelco; Bellefonte, PA, USA), whereas the different compounds were

identified by comparing the mass spectra with those of the NIST-05 library.

Figura 1. Map of the Chaviahue-Copahue volcanic complex (Argentina) (Melnick et al., 2005)

Main gas composition

The Copahue fumarolic discharges and the geothermal well, whose outlet temperatures range between 86 and

135 °C, are mainly composed by water vapour (96-98 % by vol.) and CO2 (2-4 % by vol.), while the bubbling

pools (outlet temperatures comprised between 74 and 83 °C), being affected by water condensation at ground

surface, have relatively low concentrations of water vapour (<34 % by vol.) and dominating CO2 (up to 80 % by

vol.). As far as the dry gas phase is concerned, all the gas discharges are characterized by relatively high

concentrations of H2S, CH4 and N2 (up to 6.7, 2.6 and 2.5 % by vol., respectively) and acidic species, such as

HCl and HF (up to 0.55 and 0.02 % by vol.). Minor but significant concentrations of atmospheric-related

compounds are also present. In the N2excess-CH4-Ne ternary diagram (Fig. 2), where N2excess is non-atmospheric N2

calculated on the basis of Ar concentrations, the chemistry of the Copahue thermal fluids seems to be referred to

a hydrothermal reservoir, with significant contribution from i) the magmatic system (andesitic) and ii) air.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 101-104

103

CH4

N2 excess/5Ne *10000

hydroth

erm

al

andesiticair (asw)

VOC chemical composition

On the whole, 56 different species were identified in almost all the gas discharges, showing concentrations that

vary in a wide range (from 0.1 to 48,743 ppb by vol.). As shown in Fig. 3a, VOCs are composed by comparable

amounts of aromatics (mainly benzene) and alkanes, and relatively low concentrations of alkenes (>C4), cyclics

and S-, Cl- and O-bearing compounds. The total abundances of VOCs are about three orders of magnitude less

than those of methane. This compositional feature seems to suggest that the organic gas fraction has a biogenic

origin (e.g. Whiticar and Suess, 1990) and consequently is produced at shallower depth than that of the

hydrothermal reservoir feeding the gas discharges. This evidence conflicts with the high speciation of the >C2

compounds, commonly characterizing fluids of medium-to-high temperature environments (e.g. Tassi, 2004).

d)c)

b)

0.291%0.131%0.141%0.804%1.1%

58.8%

38.7%

COPAHUE

59%

aromatics

alkanes

39%

alkenes

1%

cyclics

0.8%

S-substituted

0.3% 0.1%

0.1%O-substituted

Cl-substituted

a) HIGH TEMPERATURE

2.82%4.58%

13%

1.41%0.47%8.34%

69.3%

alkanes

aromaticsalkenes cyclics

O-substituted

Cl-substitutedS-substituted

69%

8.3%

13%

4.6%

2.8%

0.5%1.4%

LOW TEMPERATURE

29.3%

0.499%3.87%0%O-substituted

0.5%

Cl-substituted

S-substitutedaromatics

26.9%

alkenes

26.9%

16.9%

22.6%

27%

17%alkanes

22.5%

29%

3.9%

0.874%0.127%2.09%0.0313%

3.47%

49.5%

43.9%

aromatics

49.5%

alkanes44%

HYDROTHERMAL

alkenes3.5%

cyclics

2.1%

0.13%

Cl-substituted0.87%

0.03%

S-substituted

O-substituted

For comparison the compositions of the organic gas fraction of i) high-temperature fluids (Fig. 3b) from

Turrialba volcano (Costa Rica), which are strongly related to magmatic contribution (Tassi et al., 2004), ii)

typical hydrothermal fluids (Fig. 3c) from Yellowstone (USA) (Fournier, 1989), Afar (Ethiopia) (D’Amore et

al., 1997) and Tatun (Taiwan, Cina) (Lee et al., 2005) geothermal areas, and iii) CO2-rich cold gases (Fig. 3d)

from low-enthalpy systems in Tuscany (Italy) (Minissale et al., 1997), are also reported. The relative abundances

of the main groups of organics in Copahue fluids seem to be consistent with those of worldwide hydrothermal

systems (Fig. 3c), whereas significant differences are shown when compared with the composition of the

Turrialba fluids (Fig. 3b), characterized by relatively high concentrations of compounds stable at high-

temperature (i.e. S-substituted species and alkenes; Capaccioni et al., 1995), and low-temperature gases (Fig.

Figure 3a-d. Pie diagram of VOC composition in fluid discharges from a) Copahue volcano (Argentina), b) high-temperature system (Turrialba volcano, Costa Rica), c) hydrothermal systems (Yellowstone, USA; Afar, Ethiopia; Tatun, Taiwan) and d) low-enthalpy systems (Tuscany, Italy).

Figure 2. N2excess-CH4-Ne ternary diagram for gas discharges from Copahue volcano (Argentina).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 101-104

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3d), the latter being enriched in hydrocarbons commonly produced by bacterial activity in shallow environment

(i.e. alkanes and O-substituted species).

Concluding remarks

The organic gas species of the thermal fluids discharged from the Cophaue volcanic system are likely produced

by thermogenic processes occurring in the main hydrothermal reservoir. However, the origin of CH4 seems to be

decoupled from that of the non-methane VOCs, being the former likely related to biogenic processes typical of

low temperature conditions. This means that the uprising fluids from the hydrothermal system feeding the

discharges surrounding Copahue volcano are affected, at some degree, by mixing with shallow aquifers, as also

testified by the presence of atmospheric compounds. On the contrary, the influence on organic gas chemistry of

contributions from the magmatic source can be regarded as negligible. The compositional features of the organic

gas fraction may also be used for geochemical monitoring purposes since they appear to be indicative of deep-

vs. shallow-originated gas compounds, whose ratios could be modified in the case of uprising magmatic masses.

References Capaccioni B., Mangani F., 2001. Monitoring of active but quiescent volcanoes using light hydrocarbon distribution in

volcanic gases: The results of 4 years of discontinuous monitoring in the Campi Flegrei (Italy). Earth Planet. Sci. Lett. 188: 543-555.

Capaccioni B., Martini M., Mangani F., Giannini L., Nappi G., Prati F., 1993. Light hydrocarbons in gas-emissions from volcanic areas and geothermal fields. Geochem. J. 27: 7-17.

Caselli A.T., Agusto M.R., Fazio A., 2005. “Cambios térmicos y geoquímicos del lago cratérico del volcán Copahue (Neuquén): posibles variaciones cíclicas del sistema volcánico”. In XVI° Congreso Geológico Argentino, La Plata, Argentina, 1: 751-756.

D’amore F., D. Giusti And B. Gizaw, 1997. “Tendaho, Ethiopia. Geothermal Project: a Geochemical Assessment”. In 22nd Workshop Geothermal Reservoir Engineering, Stanford, January: 27-29, pp. 435-445.

Des Marais D. J., Donchin J.H., Truesdell A.H., Nehring N.L., 1981. Molecular carbon isotopic evidence for the origin of geothermal hydrocarbons. Nature 292: 826-828.

Fournier R.O., 1989. “Geochemistry and dynamics of the Yellowstone National Park hydrothermal system”. Annu. Rev. Earth Planet. Sci.: 17, 13-53.

Lee H.F., T.F. Yang, T.F. Lan, S.R. Song And S. Tsao, 2005. Fumarolic gas composition of the Tatun Volcano Group, northern Taiwan. Terrestrial, Atmospheric and Oceanic Sciences 16: 843-864.

Mango F.D., 2000. The origin of light hydrocarbons. Geochim. Cosmochim. Acta 64: 1265-1277. Melnick D., Folguera A., Ramos V.A., 2006: Structural control on arc volcanism: The Caviahue–Copahue complex, Central

to Patagonian Andes transition (38°S) Journal of South American Earth Sciences 22: (2006) 66–88. Minissale A., Evans W.C., Magro G., Vaselli O., 1997. Multiple source components in gas manifestations from north-central

Italy. Chem. Geol. 142: 175-192. Montegrossi G., Tassi F., Vaselli O. Buccianti A., Garofano K., 2001. Sulfur species in volcanic gases. Anal. Chem 73: 3709-

3715. Naranjo J.A. & Polanco E., 2004. The 2000 AD eruption of Copahue Volcano, Southern Andes. Revista Geológica Chile 31

(2): 279-292. Taran Y.A., Giggenbach W.F., 2003. “Geochemistry of light hydrocarbons in subduction-related volcanic and hydrothermal

fluids”. In Simmons and I. J. Graham (Eds.):Volcanic, Geothermal, and Ore-Forming Fluids. I - Rulers and Witnesses of Processes Within the Earth. Spec. Publ., 10, S. F. Soc. of Econ. Geol., Littleton, Colorado: 61–74.

Tassi, F., 2004. Fluidi in ambiente vulcanico: Evoluzione temporale dei parametri composizionali e distribuzione degli idrocarburi leggeri in fase gassosa. Ph.D. thesis, Univ. of Florence, Florence, Italy, pp. 292 (in Italian).

Tassi F., Martinez C., Vaselli O., Capaccioni B. And Viramonte J., 2005a. The light hydrocarbons as new geoindicators of equilibrium temperatures and redox conditions of geothermal fields: Evidence from El Tatio (northern Chile). Appl. Geochem. 20: 2049-2062.

Tassi F., Vaselli O., Capaccioni B., Giolito C., Duarte E. Fernandez E., Minissale A., Magro G., 2005b. The hydrothermal-volcanic system of Rincon de la Vieja volcano (Costa Rica): A combined (inorganic and organic) geochemical approach to understanding the origin of the fluid discharges and its possible application to volcanic surveillance. J. Volcanol. Geotherm. Res. 148: 315-333.

Varekamp J., Ouimette A., Hermán S., Bermúdez A., Delpino D., 2001. Hydrothermal element fluxes from Copahue, Argentina: A "beehive" volcano in turmoil. Geology 29 (11): 1059-1062.

Whiticar MJ, & Suess E (1990). Hydrothermal hydrocarbon gases in the sediments of the King-George Basin, Bransfield Strait, Antarctica. Appl. Geochem. 5:135-147.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 105-108

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The lithosphere of Southern Peru: A result of the accretion of allochthonous blocks during the Mesoproterozoic

Víctor Carlotto1,2

, José Cárdenas2, & Gabriel Carlier

3

1 INGEMMET, Avenida Canada 1470, San Borja, Lima 41, Peru (vcarlotto@ingemmet.gob.pe)

2 Universidad Nacional San Antonio Abad del Cusco (UNSAAC), Peru

3 Muséum National d'Histoire Naturelle, Département "Histoire de la Terre", USM 201-CNRS UMR 7160, 61, rue

Buffon, 75005 Paris, France

KEYWORDS : Southern Peru, Altiplano, lithosphere, accretion, Mesoproterozoic

Introduction

Southern Peru exhibits different juxtaposed structural blocks. These blocks have a distinct sedimentary,

tectonic and magmatic evolution. They are bounded by complex, mainly NW-SE fault systems, locally marked

by Cenozoic and Mesozoic magmatic units. The specific Mesozoic and Cenozoic geologic evolution of each

structural block is ascribed to the high heterogeneity of the southern Peruvian depth lithosphere. This lithosphere

results from the accretion of different lithospheric blocks during Laurentia-Amazonia collision at around

1000 Ma.

Structural domains

Southern Peru is characterized by the following morpho-structural domains (Figure 1):

- The Western Cordillera, which exposes siliciclastic and carbonate marine and non-marine formations

correspondining to the filling of a Mesozoic though (the Western Peruvian Mesozoic Basin);

- The Western Altiplano, which acted as a structural high (the Cusco-Puno structural high) during the

Mesozoic times and received more than 10 km of continental red beds during the Cenozoic;

- The Eastern Altiplano and the Eastern Cordillera, which show the Mesozoic sedimentary cover and the

pre-Mesozoic basement of a second mainly marine basin (the Eastern Peruvian Mesozoic Basin),

respectively.

The boundary between these domains is clearly marked by large fault systems that show evidence of activity at

least since the Paleozoic. The boundary between the Western Cordillera and the Western Altiplano is marked by

the NW-trending Cusco-Lagunillas-Mañazo (C-L-M) fault system (Figures 1 and 2, Carlotto, 1998). During the

Mesozoic, SW-dipping faults of this system had normal movements and separate the Western Peruvian Basin

from the Cusco-Puno structural high. They control the marine and continental depositions, which are thicker in

the basin and thinner at the high, respectively (Figure 2). During the Cenozoic, these normal faults acted firstly

as strike-slip faults and then, as reverse faults. Such activity resulted in the uplift of the NE margin of the

Western Peruvian Basin and converted the Cusco-Puno structural high to continental synorogenic foreland basin.

At this time, the strongest deformation and maximum shortening were concentrated along the C-L-M fault

system that represented a NE-verging foreland front (Carlotto, 1998). Further to the north, the C-L-M fault

system joins to a west-trending complex fault system, called Abancay-Andahuaylas (Fig. 1) that coincides with

the boundary between the Arequipa and Paracas blocks (Ramos, 2008).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 105-108

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Figure 1: Morphostructural domains of Southern Peru showing the main fault system. C-L-M: Cusco-Lagunillas Mañazo Fault System, U-S-A: Urcos-Sicuani-Ayaviri Fault System

The boundary between the Western Altiplano and the Eastern Altiplano and the Eastern Cordillera corresponds

to the Urcos-Sicuani-Ayaviri (U-S-A) or Cusco-Vilcanota fault system (Carlotto, 1998; Carlier et al., 2005).

This system behaves similarly to the C-L-M fault system. It separates the Cusco-Puno structural high and the

Eastern Peruvian Mesozoic Basin (Figure 2). During the Mesozoic, it consists of normal, NE-dipping faults.

During the Cenozoic, the system behaved as strike-slip or reverse, but SW-verging structures (Carlotto, 1998).

Figure 2: Mesozoic paleogeographic section viewing the bassin’s boundary and the substrate.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 105-108

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Substrate

The Arequipa Massif

The Arequipa Massif is well exposed along the southern Peruvian coast. It locally preserves semi-grabens that

show a Mesozoic cover unconformably deposited above Precambrian formations indicating that Arequipa

Massif constitutes the basement of the Western Peruvian Mesozoic Basin. Along the Cincha-Lluta Thrust

(Figure 1), this basement overlies the Mesozoic series of the Western Peruvian Mesozoic Basin. Arequipa

Massif had a complex magmatic and metamorphic polycyclic evolution from early Proterozoic to Paleozoic. It

includes 1) rocks displaying protolith ages of 1.9 Ga, affected by metamorphism between 1.9 and 1.8 Ga

(Dalmayrac et al. 1977; Cobbing et al., 1977) and 2) rocks showing Mesoproterozoic protolith and

metamorphism ages (1.2-1.0 Ga; Wasteneys et al., 1995; Loewy et al., 2004). The age of metamorphism

(Martignole & Martelat, 2003), from 1064±45 Ma to 956±50 Ma, confirms that an old protolith of 1900 Ma

underwent rejuvenation around 1000 Ma during a regional high-grade tectonic and metamorphic event related

with the Sunsas or Grenville orogeny. Hence, the geologic history of the Arequipa Massif began with the

collision between Laurentia and Amazonia, when a Paleoproterozoic terrain was trapped during the

Mesoproterozoic times between these two cratonic blocks. The consequence of the collision between the two

cratons is the formation of a mosaic of microblocks along the collisional suture. We suggest that this microblock

mosaic later formed the substrate for the Western Altiplano and Eastern Altiplano.

Western Altiplano - Eastern Altiplano substrates

Recent mineralogical, petrological, geochemical and geochronological studies of Cenozoic magmatism in the

southern Peruvian Altiplano (Carlier et al., 2005) reveal a variety of shoshonitic, calc-alkaline, acid,

peraluminous and metaluminous rocks associated to alkaline potassic (P) and ultrapotassic (UP) rocks. This

variety, together with the spatial distribution of this magmatism, implies that the deep lithosphere beneath the

Andes of southern Peru consists of a mosaic of lithospheric blocks with different origins. In fact, P-UP rocks

mostly derive from partial fusion of lithospheric mantle rocks. Mineralogical, geochemical, isotopic and

geochronological data allow to distinguish three P-UP rock associations (Carlier et al., 2005). The first group,

mostly composed of Oligocene phlogopite lamproites in the Eastern Altiplano (Figure 1), demonstrates the

presence of a Paleoproterozoic to Archaic (TDM = 1130-2485 Ma; Nd = -5.0 to -11.4; 87Sr/86Sri = 0.7100-

0.7159) metasomatized harzburgite mantle beneath this domain. Beneath the Western Altiplano (Figure 1), the

deep lithosphere corresponds to a younger (TDM = 837-1259 Ma; Nd = +0.6 a –6.3; 87Sr/86Sri = 0.7048-

0.7069) metasomatized lherzolithic mantle, as indicated by a second group of Oligocene and Miocene, P-UP,

diopside-rich lavas (leucitites, tephrites with leucite, traquibasalt with olivine and diopside). A more recent

(< 2 Ma) third group crops out at the boundary between both Altiplano domains and is composed of diopside

phlogopite lamproites, kersantites, minettes and augite trachybasalts, showing a mantle source which probably

includes an astenospheric component, apart from material derived from the two lithospheric mantles previously

described (TDM = 612-864 Ma; Nd = -1.1 a -3.5; 87Sr/86Sri = 0.7051-0.7062). This third group, present as

volcanic edifices, dikes, stocks, domes, etc., is located over the fault system, still active fault system of the U-S-

A or Cusco-Vilcanota, and marks the boundary between both parts of the Altiplano.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 105-108

108

Conclusions

The three large morphostructural domains (lithospheric blocs) previously defined evidently exhibits different

kinds of substrate. The first one, beneath the Western Cordillera (Western Peruvian Mesozoic Basin), most

probably corresponds to Arequipa Massif-like material, with ages between 1900 and 600 Ma. It probably reaches

the C-L-M fault system zone. The second one, beneath the Western Altiplano, has a deep lithosphere

corresponding to a metasomatized lherzolithic mantle. It is separated from the Western Cordillera by the C-L-M

fault system and from the Eastern Altiplano by the U-S-A fault system. The third one, beneath the Eastern

Altiplano, corresponds to a depth lithosphere with a Paleoproterozoic to Archaic metasomatized harzburgite

mantle.

Thus, the lithosphere of the western margin of the South American Plate has to be considered as a mosaic of

amalgamated lithospheric blocks (terranes) accreted to Amazonia during the Sunsas orogeny (at 1000 Ma). This

orogeny resulted from the complex collision implicating, in addition of large cratons, several smaller

lithospheric blocks such as the Arequipa Massif. The resulting heterogeneous lithosphere later formed the

basement of the Western Altiplano and Eastern Altiplano. The boundaries of the lithospheric blocks still

constitute weakness zones along which more recent deformations (lateral displacement, overthrusts…) are

concentrated. Thus, during the Cenozoic times, lithospheric microblocks composing the Southern Peru are

apparently displaced by NE-trending and E-trending transform fault systems (Patacancha-Tamburco and

Puyentimari-Rancahua faults and E-W segment of the C-L-M fault system, Figure 1). Some of these structures

like the Abancay-Andahuaylas Fault System extends towards the coast and separate the Arequipa and Paracas

massifs.

References Carlier, G., Lorand, J. P., Liégeois, J. P., Fornari, M., Soler, P., Carlotto, V., Cardenas, J., 2005. Potassicultrapotassic mafic

rocks delineate two lithospheric mantle blocks beneath the southern Peruvian Altiplano. Geology, 33, 601-604. Carlotto, V., 1998. Evolution andine et raccourcissement au niveau de Cusco (13°-16°S, Pérou). Thèse doct., univ. Grenoble,

France, 159 p. Cobbing, E.J., Ozard, J.M., Snelling, N.J. 1977. Reconnaissance geochronoiogy of the crystalline basement rocks of the

Coastal Cordillera of southern Peru. Geol. Soc. Am. Bull. 88:241--46 Dalmayrac, B., Lancelot, J.R., Leyreloup, A. 1977. Two-billion-year granulites in the late Precambrian metamorphic

basement along the southern Peruvian coast. Science 198:49--51 Loewy, S.L., Connelly, J.N., Dalziel, I.W.D. 2004. An orphaned basement block: the Arequipa-Antofalla Basement of the

central Andean margin of South America. Geol. Soc. Am. Bull. 116:171--87 Martignole, J., Martelat, J.E. 2003. Regional-scale Grenvillian-age UHT metamorphism in the Mollendo--Camana block

(basement of the Peruvian Andes). J. Metamor. Geol. 21:99-120 Ramos, V. 2008. The Basement of the Central Andes: The Arequipa and related Terranes, Annual Review of Earth and

Planetary Sciences, Volume 36 (2008): In press Wasteneys, A.H., Clark, A.H., Farrar, E., Langridge, R.J. 1995. Grenvillian granulite-facies metamorphism in the Arequipa

massif, Peru: a Laurentia-Gondwana link. Earth Planet. Sci. Lett. 132:63-73

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Igneous rocks with adakitic-like signature in South America

Silvia I. Carrasquero

Facultad de Ciencias Naturales y Museo-Universidad Nacional de La Plata (UNLP), Paseo del Bosque S/Nª,

(1900) La Plata, Argentina (silviacarrasquero@yahoo.com.ar)

KEYWORDS: adakite, Patagonian, Andes, geochemistry.

Introduction

The geological researches in South American indicate the presence of igneous rocks with adakitic-like

signature. The rocks form two groups: Andean and Patagonic adakites; the Andean adakites are linked with

metallogenic processes (Thiéblemont et al. 1997; Cassard 1999); the patagonian adakites are emplaced in thin

continental crust and no show relation with metallogenic processes.

The aim of this paper is to present the differences and the similarity between the two adakitic-types and if so, to

constrain their origins.

Adakitic magmatism

Adakites are volcanic and intrusive igneous rocks with SiO2 56 wt%, Al2O3 15 wt%, K2O/Na2O typically <

0.6, high La/Yb and Sr/Y ratios coupled with strong depletion in Y and HREE, and typically found in island and

continental arc settings (Defant and Drummond 1990; Peacock et al. 1994; Drummond et al. 1995; Maury et al.

1996).

The presence of adakites are located in the circum-Pacific region: Cascadas, USA (Defant and Drummond,

1993); Ecuador (Monzier et al. 1997, Bourdon et al. 2003; Chiaradia et al. 2004); Philippines and New Guinea

(Richards 1990; Sajona and Maury, 1998; Yumul Jr. et al. 2000); Perú (Coldwell et al. 2005); Bolivia (Bray du

et al. 1995); in porphyry ore deposits “El Abra” and “El Salvador” from Chile (Oyarzún et al. 2001); in Chilean

Central Andean (Sellés and Godoy 2000); in Austral Volcanic Zone -47-54º LS- (Kay et al. 1993; Stern and

Kilian 1996; Hattori et al. 2005), and Argentine Central Andean (Carrasquero, 1998; Carrasquero, 1999).

According to Defant and Drummond (1990), Maury et al. (1996) and Castillo (2006), adakite derive from the

high pressure ( 1-1.2 GPa) partial melting of subducting oceanic crust or underplated basic material. Peacock et

al. (1994) explain the geological conditions under which partial melting of subducting oceanic crust occurs:

subduction zones of young oceanic plate (10 Ma), the adakitic magma derive from the partial melting of young

oceanic plate; active margins thickened and zones of arc/continent collision. In two last cases the adakites would

derive from the partial of a deep crust and basic fusion, with minimum depths of 35 km.

Petrological and chemical characteristics of adakites

The adakites are phenocrysts bearing volcanic rocks. They show intermediate to felsic suites; adakites contain

more Si2O (>56 %), high-Al2O3 (> 15%) and high-Na2O (absence of plagioclase in the restite). Trace elements

allow a better distinction between adakitic and typical arc calc-alkaline magmas: high Sr contents (>400 ppm);

the (La/Yb)N > 20 and the LaN between 40 to 150; the Y content is low (< 19 ppm); the HREE are very low

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 109-112

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(Yb 1.9 ppm) and high ratios Sr/Y > 40 (Castillo, 2006). Adakites present phenocrysts of plagioclases zoned,

amphibole, biotite, apathite, titanite, zircon and titano-magnetite. The characterization of adakites (Table 1) are

proposed by Defant and Drummond, (1990), Kay et al (1993), Peacock et al (1994), Maury et al. (1996) and

Castillo (2006).

Table 1. Characterization of adakites.

Adakite characteristics Implications

phenocrysts zoned plagioclase, amphibole crystallization at high P H2O

andesitic-rhyolitic magma, SiO2 56% partial melting of oceanic basalts

high Sr (400 ppm), high Al2O3 (> 15%) absence of plagioclase in the restite

high La/Yb (>20), depletion in HREE garnet in the source region

low 87Sr/86Sr and high (Nd) MORB signature. No significant component of continental crust or subducted sediments

Ramos et al. (2004) suggest differences in adakitic rocks from South America, according to petrologic,

chemical and isotopic characteristics. There are two groups of rocks with adakitic-like signature: a.- Patagonian-

type and b.-Andean-type adakites; this later type include sectors in Andean Central of Chile and Argentina, the

volcanic region, north of Ecuador, as well as the mining region of Quimsacocha (South of Ecuador), Peru,

Bolivia and the centre-west of Argentina. b.- The Andean adakites (Fig. 1) present high Sr content

(450-900 ppm), although smaller than Patagonian-type; ratio La/Yb (11-48) and Sr/Y high (55-156); as far as the

isotopic ratios, they are less near the MORB that the Patagonian-type.

They are the result of the partial melting of thickened lower crust (eclogitic facies); the region of central

Andean of Argentina and Chile coincide with Flat-slab zone (-28 to -32ºS) where there is a cortical thickening

elder to 45 km; the other variant that presents of formation of Andean-type adakite (Ecuador) is as a result of

Fig. 1. Sr-Y/Y plot (modified from Defant and Drummond, 1990) fo adakites of: El Salvador, Chile (Baldwin and Pearce, 1982). Ecuador (Bourdon, et al. 2003). Paramillos Sur, Argentina (Carrasquero, 1999). Aguilera and Mt. Burney (Stern and Kilian, 1996).

Fig. 2. 87Sr-86Sr/143Nd/144Nd plot for some of the reported adakitic rocks in South America (modified from Castillo, 2006). Andean type adakites: Ecuador (Bourdon et al. 2003); Chile (Bissig et al. 2003; Reich et al. 2003). Paramillos Sur, Argentina (Carrasquero, 2005). Patagonian type adakites: CP Cerro Pampa (Ramos et al. 2004). MB Mt Burney (Stern and Kilian, 1996). Cenozoic adakites, Adak and Cook Is. (Castillo, 2006).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 109-112

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which the subduction caused cortical erosion; the pieces of oceanic crust were taken and introduced in the

astenospheric wedge, where, to high pressures, they underwent partial melting.

Patagonian-type adakites placed in Cerro Pampa, Puesto Nuevo, Chalten and Cook Is. (Volcanic Zone Austral)

are the result of the partial melting of a young and hot oceanic crust in continental crust. These rocks show

chemical characteristics related at partial melting of subducted oceanic crust. The Sr content is high (1300 ppm)

compared with the Andean-type (Table 2); the Y content is minor to 16 ppm; the Patagonian-type rocks have

MORB-like Sr- and Nd- isotopic ratios and more losses than Andean-type adakites (Fig 2); the initial relation 87Sr/86Sr is 0.7028 to 0.7033 whereas the 143Nd/144Nd is 0.51289 (Table 3). Analyses of Hf isotopes (Hattori et al.

2005) from Patagonian-type adakites are correlated with Nd- and Sr- isotopic ratios show compositional and

isotopic variation; the isotopic data suggest the source region contamination.

Finally, this discrimination is interesting since the Andean-type adakites are related with metallogenic

processes, which does not happen to the Patagonian-type. This quality of the Andean-type adakitic magmatism

can be used like an interesting tool in the mining exploration.

Adakites SiO2

%

Sr (ppm) 87Sr/86Sr 143Nd/144Nd La/Yb References

C° Pampa 63 - 68 1330 - 2300 0.7028 to 0.7031

(1)

> 0.5129 30 - 37 Ramos et al. 2004

Puesto Nuevo 65 - 66 1370 - 1440 0.7032 to 0.7033

(1)

0.51289 28 – 30 Ramos et al. 2004

Patagonian type

Cook Islands 59 - 61 1900 - 2000 0.7028 (2)

0.51314 30 - 35 Stern and Kilian, 1996

Ecuadorian margin

61 - 65 450 -500 0.704055 to 0.704065

(1)

0.512882 to 0.512894

13.2 - 16.8 Bourdon et al. 2003

Miocene to Quaternary volcanism,

Ecuador

56 - 69 600 - 1100 0.7040 to 0.7047

(1)

-- 11 - 48.5 Chiaradia et al. 2004

El Indio (Formation

Vacas Heladas) Chile

62 - 71 400 - 650 0.706106 to 0.707159

(2)

0.512369 23 - 33 Bissig et al. 2003

Paramillos Sur Uspallata, Argentina

59 - 64 550 - 900 0.703982 to 0.705614

(1)

0.512523 to 0.512549

34 - 58 Carrasquero 1999, 2005

Andean type

Los Pelambres, Chile

62 - 72 408 - 750 0.70439 to 0.70465

(1)

0.512619 to 0.512635

26 - 113 Reich et al. 2003

Table 2. Essential characteristics of Patagonian-type and Andean-type adakites. (1) The data are corrected with the age. (2) The data are no corrected. References Baldwin, J. A., Pearce, J. A. (1982): Discrimination of productive and nonproductive porphyritic intrusions in the Chilean

Andes. Economic Geology, 77: 664-674. Bissig, T., Clark, A., Lee, K. W., Von Quadt, A. 2003: Petrogenetic and metallogenetic responses to Miocene slab flattening:

new constraints from the El Indio-Pascua Au-Ag-Cu belt, Chile/Argentina. Mineralium Deposita, 38: 844-862.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 109-112

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Bourdon, E., Eissen, J-P., Gutscher, M-A., Monzier, M., Hall M., Cotton, M. 2003. Magmatic response to early aseismic ridge subduction: the Ecuadorian margin case (South American). Earth and Planetary Sciences Letters, 205: 123-138.

Bray du, E. A., Ludlington, S., Brooks, W., Gamble, B. M., Rathe, J. C., Richter, D., Soria Escalante, E. 1995. Compositional characteristics of Middle to Upper Tertiary volcanic rocks of the Bolivian Altiplano. U. S. Geological Survey Bulletin, 219: 1-41.

Carrasquero, S. I. 1998. Volcanismo de arco en el área del pórfido cuprífero Paramillos Sur, Uspallata, Mendoza, Argentina. X Congreso Latinoamericano de Geología y VI Congreso Nacional de Geología Económica. Buenos Aires. I: 95-100.

Carrasquero, S. I. 1999. Porphyry-type and epithermal ore deposits in the Paramillos de Uspallata district, Mendoza, Argentina. Proceedings of the fifth biennal SGA meeting and the tenth Quadrennial IAGOD Symposium, London. In: Mineral deposits: processes to processing. Stanley et al. (Eds.) 1999. Balkema Rotterdam. Actes I: 487-490.

Carrasquero, S. I. 2005. Petrology and geochemistry data of Miocene volcanism of Paramillos de Uspallata, Argentina. Geophysical Research Abstracts 7: 851.

Cassard, D. 1999. GIS Andes: A metallogenic GIS of the Andes Cordillera. Proceedings of the Fourth ISAG, Goettingen, (Germany): 147-150

Castillo, P. R. 2006. An overview of adakite petrogenesis. Chinese Science Bulletin, 51(3): 257-268. Chiarada, M., Fontboté, L., Beate, B. 2004. Cenozoic continental arc magmatism and associated mineralization in Ecuador.

Mineralium Deposita, 39:204-222. Coldwell, B., Petford, N., Murphy, P., Smith, M. 2005. “Adakitic” rocks of the Yungay Formation, Peru: problems with

tectonic setting and origin. Geophysical Research Abstracts 7: 2656. Defant, M. J., Drummond, M. S. 1990. Derivation of some modern arc magmas by melting of young subducted oceanic

lithosphere. Nature, 347: 662-665. Defant, M. J., Drummond, M. S. 1993. Mt. St. Helens: potential example of the partial melting of the subducted lithosphere

in a volcanic arc. Geology, 21: 547-550. Drummond M. S, Defant M. J., Kepezhinskas, P. K. 1995. The petrogenesis of slab derived trondhjemite-tonalite-

dacite/adakite magmas. In “The origin of granites and related rocks” Third Hutton Symposium, University of Maryland (EEUU). Ed. Brown, M.; Candela P. A.; Peck D. L.; Stephens W. E.; Walker R. J. & Zen E. Trans. Royal Soc. Edinburgh: Earth Sci. 87: 205-215

Hattori, K., Hanyu, T., Stern, C., Tatsumi, Y., Nakai, S. 2005. Hafnium isotope data suggesting the contribution of crustal material at the source in the Andean Austral Volcanic Zone. Geophysical Research Abstracts 7: 5935.

Kay, S. M., Ramos, V., Márquez, M. 1993. Dominant slab-melt component in Cerro Pampa adakitic lavas erupted prior to the collision of the Chile rise in Southern Patagonia. Journal of Geology, 101: 703-714

Maury, R., Sajona, F. G., Pubellier, M., Bellon, H., Defant, M. 1996. Fusion de la croûte océanique dans les zones de subduction/collision récentes: l’exemple de Mindanao (Philippines). Bulletin Société Géologique de France, 167 (5): 579-595.

Monzier, M., Robin, C., Hall, M.L., Cotten, J., Mothes, P., Eissen, J.-P.,,samaniego, P. 1997. Les adakites d’Equateur: Modèle preliminaire: Paris, Academie des Sciences Comptes Rendus, 324: 545–552.

Oyarzún, R., Márquez, A., Lilo, J., López, L., Rivera, S. 2001. Giant versus small porphyry copper deposits of Cenozoic age in northern Chile: adakitic versus normal calc-alkaline magmatism. Mineralium Deposita, 36: 794-798.

Peacock, S., Rushmer, T., Thompson, A. B. 1994. Partial melting of subducting oceanic crust. Earth and Planetary Sciences Letters, 121: 227-244.

Ramos, V., Kay, S. M., Singer, B. 2004. Las adakitas de la Cordillera Patagónica: Nuevas evidencias geoquímicas y geocronológicas. Revista de la Asociación Geológica Argentina, 59 (4): 693-706.

Reich, M., Parada, M. A., Palacios, C., Dietrich, A.; Schultz, F., Lehmann, B. 2003. Adakite-like signature of Late Miocene intrusions at the Los Pelambres giant porphyry copper deposit in the Andes of central Chile: metallogenic implications. Mineralium Deposita, 38: 876-885.

Richards, J. P. 1990. Petrology and geochemistry of alkalic intrusives at the Porgera gold deposit, Papua New Guinea. J. Geochem. Exploration. 35: 141-199.

Sajona, F., Maury, R. 1998 Association of adakites with gold and copper mineralization in the Philippines. C. R. Académie Sciences de Paris, Sciences de la Tèrre et des Planètes, Série II a, 326: 27-34.

Sellés, D., Godoy, E. 2000. Residual garnet signatura in Early Miocene subvolcanic stocks from the Andean foothills of central Chile. IX° Congreso Geológico Chileno (Puerto Varas), Actas 1(4): 697-699.

Stern, C. R., Kilian, R. 1996. Role of the subducted slab, mantle wedge and continental crustin the generation of adakites from the Andean Austral Volcanic Zone. Contrib. Mineral. Petrol. 123: 263-281.

Thiéblemont, D., Stein G., Lescuyer J-L. 1997. Giséments épithermaux et porphyriques: la connexion adakite. C. R. Académie Sciences de Paris, Sciences de la Tèrre et des Planètes, Série II a, 325: 103-109.

Yumul Jr. G; Dimalanta C; Bellon H; Faustino D; De Jesus J; Tamayo Jr. R. & Jumawan F. 2000. Adakitic lavas in the Central Luzon back-arc region, Philippines: lower crust partial melting products? The Island Arc, 9: 499-512.

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Long-lived constrictional strain field of the inner part of the Andean Orocline: An example of buttressing effect in oblique subduction curved margin

D. Carrizo1, G. González

2, & T. Dunai

3

1 Laboratoire de Tectonique et Mécanique de la Lithosphere, Institut de Physique du Globe de Paris (IPGP), 4

place Jussieu, 75252 Paris, France (carrizo@ipgp.jussieu.fr) 2 Departamento de Cs. Geológicas, Universidad Católica del Norte, Av. Angamos 0610, Antofagasta, Chile

3 School of Geosciences, University of Edinburgh, Drummond Street, EH9 2DZ Edinburgh, United Kingdom

KEYWORDS : Subduction, curved margin, constriction, buttress

Introduction

Oblique convergence frequently produces arc-parallel migration of the forearc sliver (Fitch, 1972; Beck, 1983).

In some cases the resistance to displacement of forearc sliver arises from a space problem (Beck et al.,1993).

This resistance is called “buttress effect” and came from three major situations: a) margin abrupt geometrical

changes, b) changes in the margin kinematics conditions or/and c) changes in the upper plate reology. Although,

the consequence of buttressing process on the upper plate strain patterns of active continental margins steel

remain poorly characterized. The oblique subduction of the Nazca Plate northeastward beneath South America

has occurred, during the last 25 Ma, along a curve margin. The Coastal Cordillera of northern Chile form part of

the only part of the South American continental crust that is in contact with the Nazca plate, containing the most

relevant geological record of the plates coupling process. These geological record is extraordinarily preserved in

the Atacama Desert, the most hyper arid desert of the Earth.

Despite of relative continuous oblique converge during the Neogene no relevant forearc sliver translation

occurred in the Central Andes (Victor et al., ; Farías et al.,2005). Based on detailed neotectonics study of the

Coastal Cordillera of northern Chile (~20ºS), we document a horizontal constrictional long-term strain field

located to the inner part of the Andean Orocline (Fig.1).

Neotectonics signatures

The fault activity is expressed as a group of fault scarps and fault-bend fold scarps whose orientation defines

three main domains: WNW-ESE-strike reverse faults, N-S-strike reverse faults and NNW-SSE-strike dextral-

reverse faults. The faults kinematics indicates trench-parallel and trench-orthogonal shortening. The faults show

evidences of coexistence activity and strong structural control related to preexistent (Mesozoic) faults system.

Exposure ages using cosmogenics 21Ne shows that the faults disrupt an Oligocene–Miocene and Pliocene (after 4

and 2 Ma) landscape preserved at the Coastal Cordillera. In addition 40Ar/39Ar chronology of displaced volcanic

tuffs and the deformations of Late Pleistocene marine terraces and Coastal Cliff colluvial sediments indicate that

fault activity remain still active during the Quaternary. Recently an Mw 5.7 shallow intraplate earthquake

located in the Coastal Cordillera (~70.15º S) evidence trench parallel shortening indicating that intraplate

deformation processes are still active (Carrizo et al., 2008).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 113-115

114

Figure 1. Geodynamic context of the Andean orocline and this study (Salar Grande area). Is represented the convergence vector and its component at both side of the margin curvature. The grey P-T axes plots represents the long-term strain fault kinematics in the Coastal Cordillera block. In red color is represented the focal mechanism of intraplate earthquakes (http://www.seismology.hardvard.edu/project/CMT/). Also is represented the Atacama Fault System (SFA) and the Precordillera Fault System (PFS).

Discussion

We explain the studied margin deformation signatures as a result of a buttress effect related principally to the

curvature of margin, concave to the ocean (the convergence change from dextral to sinistral sense that prevents

the northward sliver translation). We observe that the Coastal Cordillera block resolve the space problem by

vertical extrusion of blocks along the pre-existent faults describing a constriccional strain field. The scale and

nature of this studied strain field suggest a diffuse internal deformation behavior in the Coastal Cordillera block.

This could be explained considering that the Precordillera Fault System PFS, eastern limit of the called “outer

forearc” or as well the Andean external block can not accommodate efficiently the trench-parallel component of

the convergence as a consequence of both curvature of margin and the orogene geometry that prevents the sliver

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 113-115

115

translation (Fig.1). This deformation process is consistent with a behavior of a rigid and cold external continental

block thats not absorb relavant deformation and could transmit the effort efficiently. This characteristics suggest

that the Andean external block could plays a key role in the Andean orogen building, transmitting efficiently the

efforts from the plate coupling zone to the inland. We propose that this particular tectonics is active al least from

the Later Miocene to the present.

References Beck, M.E. 1983. On the mechanism of tectonics transport in zones of oblique subduction. Tectonophysics, Vol. 93, p. 1-11. Beck, M.E.; Rojas. C.; Cembrano, J. 1993. On the nature of buttressing in margin parallel strike-slip fault systems. Geology,

Vol. 21, p. 755-758. Carrizo, D.; González, G.; Dunai, T. 2008. Constricción neógena en la Cordillera de la Costa, norte de Chile: neotectónica y

datación de superficies con 21cosmogénico Revista Geológica de Chile 35 (1): 1-38. Farías, M.; Charrier, R.; Comte, D.; Martinod, J.; Herail, G. 2005. Late Cenozoic deformation and uplift of the western flank

of the Altiplano: Evidence from the depositional, tectonic, and geomorphologic evolution and shallow seismic activity (northern Chile at 19°S). Tectonics, Vol. 24, TC4001, doi:10.1029/2004TC001667.

Fitch, T.J. 1972. Plate convergence, transcurrent faults and internal deformation adjacent to the southeast Asia and western Pacific. Journal of Geophysical Research, Vol.77, p. 4.432-4.460.

Victor, P.; Onken, O.; Glodny, J. 2004. Uplift of the western Altiplano plateau: Evidence from the Precordillera between 20º and 21ºS (northern Chile). Tectonics, Vol 23, Tc4004, doi: 10.1029/2003TC001519.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 116-119

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The interplay between crustal tectonics and volcanism in the Central and Southern volcanic zones of the Chilean Andes

J. Cembrano1,5

, G. González

1, L.

Lara

2, E. Veloso

1, E.

Medina

1, F. Aron

1, M.

Basso

1,3,

V. Ortega1, P. Pérez

1, & G. Sielfeld

4

1 Depto. de Cs. Geológicas, Univ. Cat. del Norte, Avda.Angamos 0610, Antofagasta, Chile (jcembrano@ucn.cl)

2 Servicio Nacional de Geología y Minería, Avda. Santa María 0104, Santiago, Chile (lelara@sernageomin.cl)

3 Instituto GEA, Univ. de Concepcion, Casilla 160-C Concepción 3, Chile (mabasso@udec.cl)

4 Departamento de Ciencias de la Tierra, Univ. de Concepcion, Chile (gsielfeld@udec.cl)

5 Central Andes Resources, Callao 3785, Santiago, Chile

KEYWORDS : tectonics, volcanism, strike-slip fault, fold-and-thrust-belt

Introduction

One fundamental problem in continental margin tectonics is the nature of the interplay between deformation

processes and magma transport through the lithosphere (e.g. Hutton, 1988, Petford et al. 2000). Fault-fracture

networks have been regarded as efficient pathways through which magma can be transported, stored and

eventually erupted to the earth surface (e.g. Hill, 1977; Shaw, 1980; Clemens and Mawer, 1992). Thus, the state

of stress of the lithosphere at the time of magmatism should somehow control the first and second-order spatial

distribution of plutons, dikes swarms and volcanic centers (e.g. Nakamura, 1977; Takada, 1994, Glazner et al.

1999; Acocella et al., 2007). However, crustal deformation not only plays a significant role in magma migration;

it may also exert a fundamental control on magma differentiation processes that, in turn, can determine the

nature and composition of volcanism along and across continental margins (e.g. Cembrano and Moreno, 1994;

McNulty et al. 1998; Ferrari et al. 2000). The Chilean Andes provides a natural laboratory to assess the link

between tectonics and volcanism. Apart from its well- constrained plate kinematic history, there is a marked

latitudinal segmentation in crustal thickness, upper plate deformation and basement nature upon which the

volcanic arc has developed. Thus, the relative importance of present-day kinematics and inherited crustal

composition and structure in the mechanisms of magma transport and in the nature and composition of

volcanism can be successfully examined along the same orogenic belt. In this contribution, we examine the link

between tectonics and volcanism for two contrasting regions of the Central and Southern Volcanic zones. We

hypothesize that one fundamental, usually overlooked factor controlling the wide variety of volcanic forms and

rock compositions present along a single continental magmatic arc, is the contrasting kinematics of the fault-

fracture networks under which they are transported within the same magmatic arc.

Intra-arc tectonics of the Central and Southern Volcanic zones

New field and structural observations in combination with published seismic data allows a complete

reassessment of the complex relationship between intra-arc long-term/short-term tectonics and the

nature/composition of present day volcanism along the Chilean Andes. A thicker crust and prevailing Pliocene-

Pleistocene east-west shortening within the volcanic arc of northern Chile (22-24°S) are spatially and genetically

associated with several major composite andesitic volcanoes and only a few monogenetic basaltic eruptive

centers. Stratovolcanoes do not exhibit flank vents and clusters of minor eruptive centers are uncommon.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 116-119

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Composite volcanoes and minor eruptive centers, such as Lascar and Tilocalar respectively, are spatially and

temporally linked to the development of a Pliocene-Recent north-south-striking system of blind reverse faults

and fault-propagation folds (see Aron et al, this symposium). Evidence for long-term strike-slip deformation is

weak or absent in this part of the Central Andes Volcanic Zone (CVZ), although arc-parallel, dextral strike-slip

crustal seismicity has been documented to the north, between 18 and 21ºS (e.g. David, 2007). In contrast, the

southern Chilean Andes between 38 and 46°S are built on a much thinner crust (30-40km) that has undergone

intra-arc dextral transpressional tectonics for the last 4 Ma (e.g. Cembrano et al. 2002; Rosenau et al. 2006).

Available data of crustal seismicity consistently shows dextral strike-slip focal mechanisms from ~34º to 46ºS

(e.g. Farías, 2007; Lange et al. 2008). In the Southern Volcanic Zone (SVZ), a wide variety of volcanic forms

and compositions coexist along and across the same volcanic arc. Volcanic forms range from single

monogenetic cones lying on top of master faults to major composite volcanoes organized into either NE- or NW-

trending chains, oblique to the continental margin. Flank vents are common within individual stratovolcanoes

and as elongated clusters of minor eruptive centers. Compositions range from very primitive basalts, particularly

at minor eruptive centers, to highly evolved magmas, found at both mature stratovolcanoes and only at few

minor eruptive centers.

Discussion

Feedbacks between tectonics and volcanism in the CVZ and SVZ of the Chilean Andes can be understood as a

complex set of interactions operating at different space and time scales, ranging from long-term regional to

short-term local. Crustal thickness, nature and structure of the lithosphere, presence of compressive/transcurrent

intra-arc fault systems and magma source largely influence first-order, long-term controls. Second-order controls

include the presence of a faulted volcano-sedimentary cover versus a relatively isotropic plutonic basement, the

existence of deep-seated, seismically active or inactive faults cutting through the lithosphere and the balance

between local tectonic rates and magma input rates.

As a first approximation, a thicker crust favors magma differentiation processes whereas a thinner crust tends

to prevent it. Likewise, whereas bulk intra-arc compression (vertical 3) would tend to enhance longer residence

times of magma stored under the volcanic arc of northern Chile (22-24°S), strike-slip deformation (horizontal

3) in central and southern Chile would provide subvertical pathways for magma ascent and shorter residence

times, which in turn prevents advanced magma differentiation (Figure 1). However, looking more closely within

the strike-slip deformation zone encompassing the whole magmatic arc in southern Chile, transtensional and

transpressional domains can coexist in space and time. On one end of the spectrum, a plumbing system

dominated by NNE-striking subvertical strike-slip faults and ENE-striking tension cracks will favor a rapid

ascent of magmas from the asthenospheric wedge with little crustal contamination. On the contrary, a plumbing

system dominated by NW-striking interconnected, second-order reverse faults and subhorizontal cracks will

favor longer residence times and episodic magma fractionation, which in turn allow eruption of evolved

magmas, similar to what is observed in northern Chile. Whereas the transtensional fault-fracture network does

not require magma/fluid overpressures to operate, the compressional/transpressional does. This is consistent with

the abundant presence of volatiles that accompanies magma fractionation and differentiation as documented in

the more felsic rocks from northern Chile volcanoes and the NW-trending volcanic chains of southern Chile.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 116-119

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W E

LOFZ

?

W E

On the other hand, pre-existing subvertical structures, especially those that cut through the lithosphere, may

serve as channel ways for magma transport regardless of the bulk kinematic regime of the volcanic arc. In

particular, the Liquiñe-Ofqui fault zone (LOFZ) master faults are likely capable to connect the MASH zone or

even the asthenospheric wedge with the surface, by seismic pumping and concomitant magma production by

decompression (Figure 1). The fact that most, if not all volcanic systems that sit on top of the LOFZ are

monogenetic strongly suggests that they resulted from single, geologically instantaneous events. It is then likely

that the architecture of the overall plumbing system is primarily controlled by the nature of the fault-fracture

mesh as formed from different stress regimes and by the inherited basement structure, but more importantly,

these different architectures exert a first-order control in magma differentiation processes, which in turn account

for different volcanic morphologies and rock types along and across the same magmatic arc. Another second-

order factor controlling along-strike differences in the three dimensional architecture of the plumbing system in

the volcanic arc of central and southern Chile is the presence of a thick pre-Quaternary volcano-sedimentary

cover, especially when this cover is folded and faulted. Where such cover is present, between 33º and 37ºS, NE-

striking tension cracks formed under upper crustal dextral strike-slip deformation, may not reach the surface but

merge upwards with high angle presently inactive reverse faults marking major regional contacts between

Mesozoic and Cenozoic sequences as suggested by Diamante-Maipo and Planchón-Peteroa volcanic complexes.

In contrast, south of 38ºS, where volcanic systems are built directly on top of plutonic rocks, NE-trending

tension cracks may reach the surface and then build either a stratovolcano or an elongated cluster of minor

eruptive centers, depending on other factors such as the balance between magma input and strain rate.

A B

Figure 1. Schematic sections showing the possible geometry and kinematics of upper crustal magma plumbing system for the Central Volcanic zone at 23ºS (A) and the Southern Volcanic Zone at 40ºS (B). For the CVZ, interconnected arrays of subhorizontal, sill-like tension fractures and north-south-striking reverse faults may provide channel-ways for magma ascent and emplacement in the upper crust, favoring longer residence times and magma differentiation. In contrast, for the SVZ, magma may ascent directly from the asthenosphere along deep-seated structures such as master faults of the LOFZ, giving rise to primitive monogenetic centers. More commonly, stratovolcanoes and cluster of minor eruptive centers are probably fed by NE-striking subvertical dikes oriented subparallel to the principal stress axis.

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References Acocella, V. Vezzoli, L., Omarini, R., Matteini, M. Mazzuoli, R. 2007 Kinematic variations across Eastern Cordillera at 24°S

(Central Andes): Tectonic and magmatic implications. Tectonophysics 434, 81-92 Cembrano, J.; Moreno, H. 1994. Geometría y naturaleza contrastante del volcanismo Cuaternario entre los 38° S y 46° S:

¿Dominios compresionales y tensionales en un régimen transcurrente? In Congreso Geológico Chileno, No. 7, Actas, Vol. 1, p. 240-244. Concepción, Chile.

Cembrano, J., Lavenu, A., Reynolds, P., Arancibia, G.; López, G., Sanhueza, A. 2002. Late Cenozoic transpressional ductile deformation north of the Nazca–South America–Antarctica triple junction. Tectonophysics 354, 289– 314.

Clemens, J.C., Mawer, C.K., 1992. Granitic magma transport by fracture propagation. Tectonophysics 204, 339–360. David, C. 2007. ”Comportement actuel de l’avant-arc et de l’arc du Coude de Arica dans l’orogenese des Andes Centrales”.

Thèse doct., Univ. Paul Sabatier, Toulouse, France. 284 p. Farías, M. , Comte, D, Charrier, R. 2006. Sismicidad superficial en Chile Central: implicancias para el estado cortical y

crecimiento de los Andes centrales australes. Actas XI Congreso Geológico Chileno, vol. 1, 403- 406. Ferrari, L., Conticelli, S., Vaggelli, Chiara M. Petrone and Manetti, P. 2000. Late Miocene volcanism and intra-arc tectonics

during the early development of the Trans-Mexican Volcanic Belt. Tectonophysics 318, 161-185 Glazner, A.F., Bartley, J.M. and Carl, B., 1999. Oblique opening and noncoaxial emplacement of the Jurassic Independence

dike swarm, California. Journal of Structural Geology 21, 1275-1283. Hill, D.P. 1977. A model for earthquake swarms. Journal of Geophysical Research 82, 347-352. Hutton, D.H.W., 1988. Granite emplacement mechanisms and tectonic controls:inferences from deformation studies. Trans.

R. Soc. Edinburgh, Earth Sci. 79, 245–255. Lange, D., Cembrano, J., Rietbrock, A. Haberland, C. Bataille, K., and Hofmann, S.D. First seismic record for intra-arc

strike-slip tectonics along the Liquiñe-Ofqui fault zone at the obliquely convergent plate margin of the southern Andes. To be published in Tectonophysics.

McNulty B.A. , Farber D.L., Wallace G.S., Lopez R.and Palacios, O., 1998. Role of plate kinematics and plate-slip-vector partitioning in continental magmatic arcs; evidence from the Cordillera Blanca, Peru. Geology 26, 827-830.

Nakamura, K. 1977. Volcanoes as possible indicators of tectonic stress orientation: principle and proposal. J Volcanol Geotherm Res 2, 1-16.

Paterson, S. R., and K. L. Schmidt (1999) Is there a close spatial relationship between plutons and faults?, J. Struct. Geol., 21, 1131 – 1142.

Petford N., Cruden A. R. , McCaffrey K. J. W. & Vigneresse J.L. 2000. Granite magma formation, transport and emplacement in the Earth's crust. Nature 408, 669-673.

Rosenau, M., Melnick, D., and Echtler, H., 2006, Kinematic constraints on intraarc shear and strain partitioning in the Southern Andes between 38°S and 42°S latitude: Tectonics 25, TC4013.

Shaw, H.R. 1980. Fracture mechanisms of magma transport from the mantle to the surface. In Hardgraves, R.B. (editor). Physics of magmatic processes, 201-264. Princeton. Princeton University Press.

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U-Pb geochronologic evidence for the Neoproterozoic – Palaeozoic evolution of the Gondwanan margin of the North-Central Andes

David Chew1, Urs Schaltegger

2, Jan Ko ler

3, Tomas Magna

4, Martin J. Whitehouse

5,

Christopher L. Kirkland5, Aleksandar Mi kovi

2, Agustín Cardona

6, & Richard Spikings

2

1 Department of Geology, Trinity College Dublin, Dublin 2, Ireland

2 Department of Earth Sciences, University of Geneva, Rue des Maraîchers 13, 1205 Geneva, Switzerland

3 Department of Earth Science, University of Bergen, Allegaten 41, N-5007 Bergen, Norway

4 Institute of Mineralogy and Geochemistry, University of Lausanne, CH-1015 Lausanne, Switzerland

5 Laboratory for Isotope Geology, Swedish Museum of Natural History, S-10405 Stockholm, Sweden

6 Smithsonian Tropical Research Institute, Apartado Postal 0843-03092, Balboa, Ancon, Panama City, Republic

of Panama

KEYWORDS : U-Pb, zircon, proto-Andes, Gondwana, Neoproterozoic glaciation

Introduction

The Neoproterozoic – Early Paleozoic evolution of the Gondwanan margin of the north-central Andes has been

investigated by a U-Pb zircon geochronology study. The investigated samples comprise Palaeozoic rocks of the

Eastern Cordilleras of Peru and Ecuador (Fig. 1) and Neoproterozoic glacial sequences which overlie

Precambrian basement gneisses of the Arequipa massif in southern Peru (Fig. 2). LA-ICPMS and ion

microprobe analysis of detrital zircon has been integrated with dating of syn- and post-tectonic Palaeozoic

intrusives by TIMS and ion microprobe.

Neoproterozoic sequences - detrital zircon data

Detrital zircon populations in cover sequences overlying the Arequipa massif basement (an exotic crustal block

to Amazonia) are likely derived from the proto-Andean margin. These cover sequences (the Chiquerío and San

Juan formations in southern Peru) record the only documented Neoproterozoic glacial episode in the Andean

belt. The Chiquerío Formation yields U-Pb detrital zircon ion microprobe data with a restricted age distribution

of 950-1300 Ma. Turbiditic dolomitic sandstones in the overlying San Juan Formation yield a similar

950-1300 Ma peak, but also contain grains dated as 1600-2000 Ma and 700-820 Ma (Chew et al., 2007a). Based

on the presence of a cap carbonate and two negative C isotope excursions the Chiquerío and San Juan formations

probably represent a Sturtian–Marinoan couplet (c. 750 – 635 Ma). The strong link between the Arequipa massif

cover sequences and the proto-Andean margin during the Late Neoproterozoic rules out accretion of the

Arequipa massif during the early Paleozoic Pampean and Famatinian orogenies, and strongly implies accretion

to Amazonia during the 1000–1300 Ma Grenville–Sunsas orogeny (Chew et al., 2007a; cf Loewy et al., 2004).

Palaeozoic sequences – detrital zircon data

In the Palaeozoic metamorphic belts of the Eastern Cordilleras of Peru and Ecuador, the majority of detrital

zircon samples exhibit prominent peaks in the ranges 0.45 - 0.65 Ga and 0.9 - 1.3 Ga, with minimal older

detritus from the Amazonian craton. The detrital zircon data demonstrate that the basement to the western

Gondwanan margin was likely composed of a metamorphic belt of Grenvillian (0.9 – 1.3 Ga) age, upon which

an Early Paleozoic magmatic belt (0.45 – 0.5 Ga) developed in a similar way to the Sierra Pampeanas and

Famatina in northern Argentina (Chew et al., 2007b). These two orogenic belts are interpreted to be either

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buried underneath the present-day Andean chain or adjacent foreland sediments. However the source for detritus

in the 0.55 – 0.65 Ga age range, broadly age-equivalent to the Brasiliano/Pan-African Orogeny in eastern

Amazonia, remains puzzling.

Evidence for a Neoproterozoic active margin in the detrital zircon data?

No obvious source for 0.55 – 0.65 Ga detritus is known in the northern and central Andes. Derivation from

eastern Amazonia is considered unlikely due to the stark paucity of detritus derived from the core of the

Amazonian craton. Instead, we propose that a Late Neoproterozoic magmatic belt is buried beneath the present-

day Andean belt or Amazon Basin, and was probably covered during the Eocene – Oligocene. If this inferred

Neoproterozoic belt was an active margin, it would record the initiation of Proto-Andean subduction and imply

at least partial separation of West Gondwana from its conjugate rift margin of eastern Laurentia prior to

ca. 650 Ma. This separation may be linked to the ca. 770 – 680 Ma A-type magmatism found on eastern

Laurentia in the southern Appalachians (e.g. Tollo et al., 2004) and on the Proto-Andean margin in the Sierra

Pampeanas (Baldo et al., 2006) and in the Eastern Cordillera of Peru.

U-Pb dating of syn- and post-tectonic Palaeozoic intrusives and discussion

Plutons associated with the Early Paleozoic subduction-related magmatic belt have been identified in the

Eastern Cordillera of Peru, and have been dated by U-Pb zircon TIMS and ion microprobe to 474 – 442 Ma

(Chew et al., 2007b). This is in close agreement with the ages of subduction-related magmatism in the Arequipa

– Antofalla Basement (e.g. Loewy et al., 2004). This Early Paleozoic arc is clearly not linear as it jumps from a

coastal location in the Arequipa – Antofalla Basement to several hundred kilometers inland in the Eastern

Cordillera further to the north. This is interpreted as an embayment on the Proto-Andean margin at the time the

arc was initiated; if this is the case the northern termination of the Arequipa-Antofalla Basement in the vicinity

of Lima is an Ordovician or older feature.

The arc magmatism pre- and post dates phases of regional metamorphism in the Eastern Cordillera of Peru.

U-Pb zircon ion microprobe dating of zircon overgrowths in high-grade leucosomes demonstrates the presence

of a metamorphic event at c. 478 Ma, and refutes the previously-assumed Neoproterozoic age for orogeny in the

Peruvian Eastern Cordillera (Chew et al., 2007b; Cardona 2006). The presence of an Early – Middle Ordovician

age magmatic and metamorphic belt in the north-central Andes demonstrates that Famatinian metamorphism and

subduction-related magmatism was continuous from Patagonia (Pankhurst et al., 2006) through northern

Argentina and Chile to as far north as Colombia and Venezuela, a distance of nearly seven thousand kilometres.

The presence of an extremely long Early – Middle Ordovician active margin on western Gondwana invites

comparison with the Taconic – Grampian orogenic cycle of the eastern Laurentia margin (which is of similar age

and strike length) and supports models which have these two active margins facing each other during the

Ordovician.

U-Pb zircon ion microprobe dating of zircon overgrowths in migmatites yields ages of c. 312 Ma, and

represent a previously unreported high-grade Gondwanide event which has affected the Peruvian segment of the

Proto-Andean margin. The original relationship between the Carboniferous and Ordovician metamorphic belts

is uncertain as it has been affected by later Andean (Eocene – Oligocene) thrusting, but overall the pattern of

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crustal growth in the north-central Andes implies that it was dominated by a series of progressive crustal

accretion events, which results in a series of age domains that young away from an old Amazonian core (Chew

et al., 2007b).

Figure 1. Geological map of Peru and Ecuador from Chew et al. (2007b) illustrating the major Palaeozoic metamorphic and magmatic belts along with the Proterozoic gneisses of the Arequipa massif. Inset figures a-f illustrate zircon probability density distribution diagrams for both metasedimentary and magmatic (inherited cores) samples.

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Figure 2. Zircon probability density distribution diagrams from the Chiquerío Formation (SJ-11, SJ-16) and the San Juan Formation (SJ-57) (Chew et al., 2007a).

References

Baldo, E., Casquet, C., Pankhurst, R.J., Galindo, C., Rapela, C.W., Fanning, C.M., Dahlquist, J., & Murra, J. 2006. Neoproterozoic A-type magmatism in the Western Sierras Pampeanas (Argentina): evidence for Rodinia break-up along a proto-Iapetus rift? Terra Nova 18(6): 388-394.

Cardona, A. Cordani, U. G., Ruiz, J., Valencia, V., Nutman, A. P., & Sanchez, A. W. 2006. “U/Pb detrital zircon geochronology and Nd isotopes from Paleozoic metasedimentary rocks of the Marañon Complex: Insights on the proto-Andean tectonic evolution of the Eastern Peruvian Andes”, Fifth South American Symposium on Isotope Geology, April 24–25 2006, Punta del Este, Uruguay: 208–211.

Chew, D.M., Kirkland, C.L., Schaltegger, U. & Goodhue, R. 2007a. Neoproterozoic glaciation in the Proto-Andes: tectonic implications and global correlation. Geology 35(12): 1095-1099.

Chew, D.M., Schaltegger, U., Ko ler, J., Whitehouse, M.J., Gutjahr, M., Spikings R.A. and Mi kovic, A. 2007b. U-Pb geochronologic evidence for the evolution of the Gondwanan margin of the north-central Andes. Geological Society of America Bulletin 119(5/6): 697-711.

Loewy, S.L., Connelly, J.N., & Dalziel, I.W.D. 2004. An orphaned basement block; the Arequipa-Antofalla basement of the Central Andean margin of South America. Geological Society of America Bulletin 116: 171-187.

Pankhurst, R.J., Rapela, C.W., Fanning, C.M., & Márquez, M., 2006. Gondwanide continental collision and the origin of Patagonia. Earth-Science Reviews 76: 235–257.

Tollo, R.P., Aleinikoff, J.N., Bartholomew, M.J. & Rankin, D.W. 2004. Neoproterozoic A-type granitoids of the central and southern Appalachians: intraplate magmatism associated with episodic rifting of the Rodinian supercontinent. Precambrian Research 128(1-2): 3-38.

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Adakitic rocks and their geodynamic significance: Examples from the Andes of Ecuador and Peru

Massimo Chiaradia1, Daniel Merino

1, & Bernardo Beate

2

1 Department of Mineralogy, Rue des Maraîchers 13, 1205 Geneva, Switzerland

(Massimo.Chiaradia@terre.unige.ch) 2

Escuela Politecnica Nacional, Quito, Ecuador (bbeate@uio.satnet.net)

KEYWORDS: Ecuador, Peru, adakitic, magma, isotopes

Introduction

The commonly accepted mechanism that generates subduction-related magmas is the lowering of mantle

wedge solidus by slab-released fluids (flux melting) accompanied or not by decompression melting. On the other

hand, the finding in several modern arc settings of calc-alkaline rocks with peculiar geochemical signatures (e.g.,

Sr/Y>40, high La/Yb>15) has been interpreted as the result of the partial melting of subducted oceanic crust

rather than of the mantle wedge and these rocks have been named adakites (Defant and Drummond, 1990).

Recently there has been an increased scientific interest in the petrogenesis of rocks with adakitic signatures

because of their implications in ancient and modern crustal growth processes and their potential association with

porphyry-related Cu-Au deposits. In the last few years rocks with adakitic signatures have been interpreted

either as the result of slab melting or of evolution of mantle-derived melts at lower crustal levels through

MASH-type processes, involving partial melting of the lower crust, magma mixing, and high-pressure fractional

crystallization.

Figure 1. Geotectonic map of northern South America showing the location of the three investigated areas with adakitic magmatism.

This study presents petrographic, geochemical and isotopic data (Pb, Sr, Nd) from adakitic subduction-related

magmatism of three different areas of the Andes of Ecuador and Peru (Figure 1): (i) the Paleocene(?)-Eocene

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Macuchi island arc (Western Cordillera of Ecuador); (ii) the Miocene world-class Au-district of Yanacocha

(northern Peru); (iii) the frontal volcanic arc of Ecuador (exemplified by the Pululagua volcanic center). The data

presented suggest that in all three cases adakitic signatures result from a deep-seated evolution of “normal”

mantle-derived calc-alkaline melts through high-pressure fractional crystallization accompanied by assimilation

of lower crustal material and/or mixing with its partial melting products.

Geological and geotectonic settings

The Macuchi Unit is an island arc volcanic and volcaniclastic sequence cropping out in the Western Cordillera

of Ecuador (Figure 1) considered Paleocene (?) to Eocene in age, based on fossil associations and sparse

radiometric dating. Chiaradia and Fontboté (2001) have subdivided the Macuchi Unit into a lower (Basal) and an

upper (Main) sequence, based on chemostratigraphic differences. The two sequences would reflect a growing

submarine island arc edifice erupted through a thickened oceanic plateau basement. The marked geochemical

and isotopic differences between the two sequences (Figures 2 and 3) would result from a major geodynamic

rearrangement, such as a subduction jump or reversal (Chiaradia and Fontboté, 2001).

Subduction-related magmatism in the Yanacocha gold district (Figure 1) spans a time interval between 14.5

and 8.4 Ma (Longo 2005) and consists of both eruptive and intrusive rocks evolving from andesite to rhyolite

through time. U-Pb zircon ages of porphyritic rocks investigated in the present study range between 12.6 and

10 Ma. Rosenbaum et al. (2005) suggest that the Yanacocha magmatism and associated mineralization was

broadly coeval with the arrival at the subduction zone of the buoyant (now subducted) Inca plateau (Figure 1).

Quaternary volcanism in Ecuador results from the eastward oblique subduction of the 12-20 My old Nazca

plate beneath the assembled Ecuadorian crust at a regular dip of 25-30° (Guillier et al., 2001). The stress regime

on the overriding plate in Ecuador has been significantly influenced by the Carnegie ridge subduction (Figure 1),

which is causing an increased coupling between the subducted and overriding plate (e.g., Sage et al., 2006).

Figure 2. Correlations between SiO2 and adakitic indices for magmatic rocks of the three investigated areas.

Pululagua is an active volcanic center of the frontal arc of Ecuador (Figure 1). Situated 15 km NNE of Quito, it

contains a 3-km-wide summit caldera narrowly breached to the west and partially filled by a group of dacitic

lava domes. Large explosive eruptions producing pyroclastic flows took place during the late Pleistocene and

Holocene. The latest dated eruption occurred about 2400 years ago and resulted in caldera formation.

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Figure 3. Correlations between radiogenic isotopes and adakitic indices for magmatic rocks of the three investigated areas.

Results

The Macuchi island arc rocks vary from basalt to dacite in composition, Yanacocha investigated rocks range in

composition from andesite to rhyolite, Pululagua eruptive products are andesitic to dacitic. Rocks from the three

investigated areas display typical subduction-related features such as LILE enrichment and depletions in Nb and

Ta. REE spectra are characterized by variably steep LREE to HREE transitions and by the absence of negative

Eu anomalies suggesting that plagioclase fractionation was limited (especially when considering the

intermediate and felsic terms) whereas amphibole, clinopyroxene and, in some cases, garnet fractionation

occurred either in the restitic source of these rocks and/or during magmatic evolution.

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In the Sr/Y versus Y plot the investigated rocks of each area define trends starting in the non-adakitic and

ending in the adakitic field. Adakitic indices (Sr/Y, La/Yb) show strong correlations with both evolution indices

(e.g., SiO2: Figure 2) and various radiogenic isotopes (Figure 3).

Discussion and conclusions

The correlations of adakitic indices (Sr/Y, La/Yb) with magmatic evolution indices and radiogenic isotopes

(Figures 2 and 3) suggest that in the three areas of the Andes here investigated adakitic signatures were acquired

through magmatic evolution processes involving fractional crystallization at depth (outside the plagioclase

stability field but within that of amphibole, clinopyroxene ± garnet) accompanied by assimilation of mid- to

lower crustal rocks and/or mixing with their partial melting products. Such interpretation is in agreement with

the occurrence of disaggregated and rounded lower crustal xenoliths and xenocrysts in the investigated magmatic

rocks and with pressure of crystallization of minerals such as amphibole and clinopyroxene indicating mid- to

lower crustal depths. In all three cases rocks with adakitic signatures follow temporally more or less prolonged

periods of “normal” calc-alkaline magmatism. The switch to the adakitic signature can be broadly correlated

with major geodynamic changes at the subduction zone such as the inception of an aseismic buoyant ridge (e.g.,

Carnegie ridge for Pululagua) or oceanic plateau (Inca plateau for Yanacocha) or subduction jump or reversal

(Macuchi). These geodynamic changes are likely to have caused an increased coupling between overriding and

subducting plates thus provoking a stalling of mantle-derived magmas at depth where they evolved through the

above described processes (see also Chiaradia et al., 2004). Under this point of view rocks with adakitic

signatures may be important indicators of geodynamic changes at subduction zones (essentially increased

compression and thickening), which may also be favorable to the formation of economic mineralization (e.g.,

Macuchi, Yanacocha). Although it is not yet clear whether the slab component is transferred to the mantle wedge

as an aqueous fluid or a hydrous melt, it seems evident that the adakitic signatures in the investigated areas and

probably in many other cases are not the result of slab melting but of magmatic evolution at depth of “normal”

mantle-derived melts.

References Chiaradia, M., & Fontboté, L. 2001. Radiogenic lead signatures in Au-rich VHMS ores and associated volcanic rocks of the

Early Tertiary Macuchi island arc (Western Cordillera of Ecuador). Economic Geology 96: 1361-1378. Chiaradia, M., Fontboté, L. & Beate, B. (2004) Cenozoic continental arc magmatism and associated mineralization in

Ecuador. Mineralium Deposita 39: 204–222. Defant, M.J., & Drummond, M.S. 1990. Derivation of some modern arc magmas by melting of young subducted lithosphere.

Nature 347: 662–665. Guillier, B., Chatelain, J. L., Jaillard, E., Yepes, H., Poupinet, G. & Fels, J. F. 2001. Seismological evidence on the geometry

of the orogenic system in central–northern Ecuador (South America). Geophysical Research Letters 28: 3749–3752. Longo, A. 2005. Evolution of Volcanism and Hydrothermal Activity in the Yanacocha Mining District, Northern Perú. Ph.D

Thesis, Oregon State University, 469 p. Rosenbaum, G., Giles, D., Saxon, M., Betts, P. G., Weinberg, R. F., Duboz, C. 2005. Subduction of the Nazca Ridge and the

Inca Plateau: Insights into the formation of ore deposits in Peru. Earth and Planetary Science Letters 239: 18-32. Sage, F., Collot, J.-Y., & Ranero, C.R. 2006. Interplate patchiness and subduction-erosion mechanisms: evidence from depth-

migrated seismic images at the central Ecuador convergent margin. Geology 34: 997-1000.

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Seismotectonic analysis of the Bucaramanga Seismic Nest, Colombia

Germán Chicangana1, 2

& Carlos A. Vargas2

1 Escuela de Ingenierías, Corporación Universitaria del Meta, Barrio San Fernando, Villavicencio, Colombia

2 Grupo de Geofísica, Universidad Nacional de Colombia, Edificio Manuel Ancizar, Ciudad Universitaria, Bogotá,

D.C., Colombia (gechicanganam@bt.unal.edu.co, cavargasj@unal.edu.co)

KEYWORDS : intermediate seismicity, geodynamics, seismotectonics, lithospheric delamination, Colombia

Introduction

The Bucaramanga Seismic Nest (BSN) is located beneath Central Colombia (Figure 1), between the Central

Cordillera and the Eastern plains or “Llanos orientales” (4° - 8° N, 73° W - 76° W). It represents about 60% of

seismicity recorded annually by the National Seismological Network of Colombia (NSNC).The seismicity of

BSN has intermediate depths (60 < h < 160 km) with magnitudes M 6.0. In this work, we are going to explain

a possible hypothesis about its origin using seismic moment tensors reported by the NSNC and NEIC.

Methodology

In order to visualize the lithospheric and sublithospheric environment beneath central Colombia we use a focal

earthquakes profile from data of seismicity recorded by the NSNC for the period 1993 – 2001. The database

includes 7819 events with 1> ML> 6,8, and 0 < h <200 Km. (Fig. 1 A), then earthquakes solutions were located

using the program HYPOCENTER (Chicangana and Vargas, 2007). Tomographic Vp images (Vargas, 2004)

were used to show the distribution of seismic anomalies in the region framed in 0,7° N to 8,3° N, and 73° W and

78° W (Fig. 1D). These images give us an approach to visualize the lithospheric environment where the BSN is

located.

Results & Conclusions

The BSN sublithospheric environment is a strong intermediate seismicity where the VP distribution allows

observing (figures 2 and 3) important contrast of the % VP indicates a fragile environment around 80 km to

160km depth. In the earthquakes focal profiles is possible to check the increased of seismic activity for this

range of depths suggesting a lithospheric collision zone. The %VP increments between 100 and 120 km

suggesting the presence of a sublithospheric rigid body located beneath Bucaramanga and Los Santos plateau. It

is possible that phenomenon responds to a lithospheric removed fragment that corresponded to the previously

subducted Farallon Plate. This old slab at Present is in destruction process due to their collision with the Nazca

Plate when taking place the lithospheric delamination (Chicangana et al., 2007). Between the 120 and 160 km

gradual %VP increase is observed (Figures 3 & 4).

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Figure 1. Localization of the region of this work. A, Seismicity record of the National Seismological Network of Colombia for the period 1993 - 2001. A - A´ as the profile of the figure 3. B, Seismicity used to obtain the local tomographic of Colombia from inversion of the Model Minimum 1D and the later inversion of the 3D model. After Vargas (2004).

Figure 2. Percentage of velocity of the P wave (VP) relative to the initial 1D model result of the 3D inversion with the sections to different depths of the study area.

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Figure 3. Profile A – A´ located in the figure 1. Above, Tomographic profile after Vargas (2004). Below, Lithospheric and sublithospheric interpretation with focal earthquake profile. CCOP: Colombian – Caribbean Oceanic Plateau. CRPCB: Costa Rica – Panama – Choco – Block. EFFS: Eastern Frontal Fault System. RFS: Romeral Fault System.

Figure 4. Hypothetical 3D model showing the subduction of the Nazca plate beneath northwestern South America, and its collision with old Farallon Plate detached slab by delamination lithosphere effect. This phenomenon gives origin to the seismicity of the BSN from Neogene times. The insert in the left below corner is an image in Google Earth® for the geographical reference. See details in the text.

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A preliminary interpretation of the diverse fault types of the BSN´s seismic moment tensors report by NSNC and

NEIC is indicating a compressional stress field due to the old Farallon Plate slab holding a positive buoyance

because its add to the continental lithosphere when cease the subduction during the Paleogene, when Caribbean

Plate colliding with the northwestern South America margin (Chicangana, 2005). The slab fragment spreads

upward again by this effect but the collision with the Nazca Plate slab impedes it and producing inverse and

normal seismic moment tensors. Probably it is destroying by gravitational effect of the downward Nazca plate

in the superior mantle. As Nazca Plate take an age < 20 Ma it presents negative bouyance producing normal

seismic tensors. In synthesis (Figure 4) the accumulation of tectonic effects as the downward Nazca Plate, the

resistance to the descent of the old removed Farallon Plate slab toward mantle, and the compression for the

effect of the collision is the result of the diverse seismic moment tensors that take place in the BSN.

References Chicangana, G., 2005. Estudio del Sistema de Fallas Romeral (0,5 – 11,5 ° N), a partir de una caracterización

sismotectónica regional. Tesis de Maestría en Ciencias – Geología, Departamento de Geociencias, Facultad de Ciencias, Universidad Nacional de Colombia, Bogotá D.C. 191p.

Chicangana, G., & Vargas, C. A., 2004. Desarrollo y Geometría actual de la litosfera en la Esquina Noroccidental de Suramérica. Memorias del 1er. Congreso Latinoamericano de Sismología, http://olimpia.uan.edu.co/sls/1cls/resumenes/poster/NACIONALES/desarrollo_geometria.pdf

Chicangana, G. & Vargas, C. A.. 2005. Two regions with intermediate seismicity rate increase under Colombian Andes visualizated and interpretated with the Combination of Local Seismic Tomography and Hypocentral Profiles: Regions of Eje Cafetero and the Seismic Nest of Bucaramanga. 6th International Symposium Andean Geodynamics – ISAG. Volume of extended abstracts. 170 – 173.

Chicangana, G., & Vargas, C. A., 2007. Aproximación a la geometría litosférica y sublitosférica bajo los Andes del centro de Colombia, in Memorias del XI Congreso Colombiano de Geología, CD – Room.

Chicangana, G., Vargas, C. A., & Kammer, A., 2006. La Evolución del Centro de Expansión de Galápagos y su Papel en la sismicidad intermedia del occidente colombiano. Memorias del II Congreso Latinoamericano de Sismología, CD – Room.

Kissling, E., Solarino, S., & Cattaneo, M., 1995. Velest Users Guide. Internal report, Institute of Geophysics, ETH Zurich. Lienert, B.R.E. & Havskov, J., 1995. A computer program for locating earthquakes both locally and globally, Seismological

Research Letters, 66, 26-36. Vargas, C. A., 2004. Propagación de Ondas sísmicas y atenuación de ondas Coda en el Territorio Colombiano. Rev. Acad.

Col. Cien. Fis. Exact y Nat. Colección Jorge Álvarez Lleras No. 23, 235p.

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Seismotectonic behavior of the Eastern Frontal Fault System: Seismic hazard for the Villavicencio region, Central Colombia

Germán Chicangana1, 2

, Carlos A. Vargas2, & Andreas Kammer

3

1 Escuela de Ingenierías, Corporación Universitaria del Meta, Barrio San Fernando, Villavicencio, Colombia

2 Grupo de Geofísica, Universidad Nacional de Colombia, Edificio Manuel Ancizar, Ciudad Universitaria, Bogotá,

Colombia (gechicanganam@bt.unal.edu.co, cavargasj@unal.edu.co) 3 Departamento de Geociencias, Universidad Nacional de Colombia, Edificio Manuel Ancizar, Ciudad

Universitaria, Bogotá, Colombia (akammer@unal.edu.co)

KEYWORDS : Eastern Colombian Foothills, seismic hazard, seismotectonic, eastern Frontal Fault System, seismic hazard

Introduction and methodology

The Villavicencio city (500,000 inhabitants), is located on the eastern foothills of the Eastern Colombian

Range, about 60 km toward SE from Bogotá D.C. (8,000,000 inhabitants) (Figure 1). In this area active tectonic

evidences have been previously observed in several faults related to Eastern Frontal Fault System (Paris et al.

2000; Chicangana et al., 2007). At August 31, 1917, a M 6,5 earthquake affected the Villavicencio city and the

region around it with economic and human lives losses (Cifuentes et al., 2006).

Recognition of the main tectonic features of the Eastern Frontal Fault System - EFFS (3º N - 5º N) using

remote sensing, geological maps and field work, allowed us to elaborate a 3D geotectonic model that helps to

understand the potential M 6,0 regional earthquakes and their relationship with the main fault planes. This

model is also optimized with historical and instrumental seismological records. The results of this work

contribute to improve the knowledge about the seismic hazard near to Bogota D.C and Villavicencio cities.

Active tectonics

Active tectonics evidences have been verified near to Villavicencio urban area and as far as 25 km to the north

of this city. These evidences are active front thrust with high load mass in bar - braided drainages, constant big

groundmass slides, and permanent cleaning along strike of main fault scarp in a regional scale. The active

tectonics is product of working of thrust faults due to sliding updip or downdip fault plane, also but less by strike

– slip component (Figures 2 & 3). In this region the Eastern Frontal Fault System is composed by severed

foreland SE vergent thrust faults like Algeciras, Guaicaramo, Guayuriba, Mirador and Servita Faults. Toward

surface high angle ( 90°) predominated in these fault planes, but toward depth ( 10 – 30 km) dip angle

decreases to 20° or less (Figure 2). These tectonic features derived of tectonic positive inversion by Neogene

reactivation of large Mesozoic normal faults; some of these faults were previously affected by a Neoproterozoic

continental collision. The Guaicaramo Fault is a fundamental lithospheric limit between these ancient continents

(Laurentia and Gondwana).

The stress field from microtectonics data for NNE fault planes of the Eastern Colombian range between 3° N

and 4° N indicates a 1 W- E, while at the north (4° N) has a 1 NW – SE (Chicangana et al., 2007). This

stress field was produced by the Pliocene - Pleistocene geodynamic evolution of NW South America, Costa Rica

- Panama – Choco Block and Caribbean plate (Chicangana & Vargas, 2006). The frictional mechanism over

fault thrust planes produced a rupture area toward updip or downdip (Figure 3).

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Figure 1. Localization of the study region and some faults of the Eastern Frontal Fault System.

Figure 2. Left, 3D sketch showing regional tectonic – stratigraphic framework using a DTM. Right, Profiles A - A´, B - B´, and C - C´ of left sketch. GCF: Guaicaramo Fault. FGY: Guayuriba Fault. FM: Mirador Fault. MAF: Manzanares Fault. RChF:Rio Chiquito Fault. GF: Gallo Fault. SF: Servitá Fault. FSJ: San Juanito Fault. AMProtB: Andean Mesoproterozoic Basement.

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Figure 3. DEM showing the stress fields for the central sector Eastern Colombian Range (P: Pressure, T: Tension), with the main kinematics of main thrust faults due to tectonic effect in three different sectors of the Eastern Frontal Fault System exhibited in the cubes that they show the rupture area that would generate the earthquake.

Figure 4. Seismotectonic maps for the study region from historical seismicity data after Chicangana et al. (2007) (left), and seismicity record of National Seismological Network of Colombia (INGEOMINAS, 2001) during the period 1993-2001 (right).

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Historical seismicity and regional seismic hazard

Historical and instrumental seismological records (Figure 4) showing a lower shallow seismicity activity (an

earthquake M 4.0 in a period 10 years) between 4º N - 5º N, and moderate shallow seismicity activity (an

earthquake M 4.0 in a period 10 years) between 3° N and 4° N. The thrust faults extend toward the rigid

Proterozoic basement with lower dip angles, being composed of shortcut planes and potential earthquake

M 6,0. This work proposes that Guaicaramo and Servita thrust faults toward north of 4° N can produce a big

earthquake but equally Guayuriba or Mirador thrust fault toward south or near of 4° N. Also a great seismic

hazard is the Algeciras Thrust Fault due to that is responsible for big historical earthquakes how February of

1967 M = 6.7 earthquake (Ramirez, 1975; Velandia et al., 2005), when Bogota, Neiva and Villavicencio were

strongly affected. The knowledge of seismic cycle for to value the parameters of seismic hazards of these thrust

faults is even poor because the lack of paleosismological studies and the small record of the seismological

network. We conclude that only the realization of future studies with the support of a local seismological

network would contribute to reduce the effects of seismic hazard in Bogota, Villavicencio, and Central

Colombia.

References Chicangana, G., & Vargas, C. A. 2006. Evolución del estilo orogénico actual de los Andes del norte: Resultado de la acreción

del Bloque Costa Rica – Panamá – Chocó (BCRPC) durante el Plioceno Superior. Memorias del II Congreso Latinoamericano de Sismología, En CD – Room.

Chicangana, G., Vargas, C. A., Kammer, A., Hernández Hernández, T. A., & Ochoa Gutiérrez, L.H. 2007. Caracterización Sismotectónica Regional Preliminar de un sector del Piedemonte Llanero colombiano: Corredor San Juan de Arama – Cumaral, Meta: Boletín de Geología – UIS, 29, 61 – 74.

Cifuentes, H. G., Sarabia, A. M., Robertson, K. G., & Dimaté, A. C. 2006. Parámetros Macrosísmicos del Sismo de 1917 en Colombia. Memorias del II Congreso Latinoamericano de Sismología, En CD – Room.

INGEOMINAS. 2001. Boletín de Sismos 1993 –2001. INGEOMINAS - RSNC. París, G., Machette, R., Dart, R. L., & Haller, K. M. 2000. Database and Map of Quaternary faults and folds of Colombia and

its offshore regions, Open – File Report 00 – 0284: http//www.pubs.usgs.gov/of/2003/opf-00-0284. Ramirez, J. E. 1975. Historia de los Terremotos en Colombia, IGAC, 250p. Velandia, F., Acosta, J., Terraza, R., & Villegas, H. 2001. The current tectonic motion of the Northern Andes along the

Algeciras Fault System in SW Colombia: Tectonophysics, 399, 313 – 329.

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The Mw 7.7 Tocopilla earthquake of November 2007: Characteristics of a subduction earthquake that occurred in the brittle-ductile transition zone of the northern Chile seismic gap

Mohamed Chlieh1, Dominique Rémy

2, Bertrand Delouis

1, Sylvain Bonvalot

2, Germinal

Gabalda2, Tony Monfret

1, & Mario Pardo

3

1 Géosciences Azur, Université de Nice, IRD, 250 rue A. Einstein, 06560 Valbonne, France

(chlieh@geoazur.unice.fr) 2 LMTG, Université de Toulouse, CNRS, IRD, OMP, 14 av. E. Belin, 31400 Toulouse, France

3 Departamento de Geofísica, Universidad de Chile, Blanco Encalada 2002, Santiago, Chile

KEYWORDS : subduction earthquake, GPS, InSAR, slip inversion, coseismic, afterslip

Abstract

The Mw 7.7 Tocopilla earthquake of November 14, 2007 is categorized as a subduction underthrusting event

that occurred at the contact of the Nazca and South America plates. This earthquake ruptured the southern-east

tip of the well identified ~500 km-long seismic gap of northern Chile. We determine coseismic and the first-

month postseismic slip distribution associated with this earthquake and its aftershocks from near-field permanent

Global Positioning System (GPS) surveys and InSAR data acquired on two adjacent tracks. The coseismic

model shows that the Nazca subduction megathrust ruptured over a distance of about 150 km and a width of less

than 50 km. Maximum slip of about 1.5-2.5 m occurred around two major asperities between 35 km and 55 km

depth. It releases a total moment of 4.5 1020 Nm, equivalent to a magnitude Mw=7.7. Slip inversion of the InSAR

data that included up to 45 days of postseismic deformation in addition to the coseismic deformation requires

that slip must have continued on the plate interface after the 45s seismic rupture. The postseismic moment

released could have been more than 30% of the coseismic moment release, with significant afterslip between the

two coseismic asperities and off-shore the Mejillones Peninsula.

Introduction

The Mw 7.7 Tocopilla earthquake of November 2007 occurred along the Northern Chile subduction zone,

which absorbs about 60 mm/yr of northeastward motion of the Nazca plate relative to the South American craton

(Figure 1). This event was identified as a subduction underthusting earthquake and occurred in the deep portion

of the seismogenic zone.

Characteristics of the Tocopilla earthquake derived from seismology

A comprehensive view of the slip history and distribution was obtained by combining information from strong-

motion records obtained from a local seismic network and teleseismic data. Delouis et al. [2008] relocated the

mainshock hypocenter with an uncertainty of +/- 4 km at 22.33ºS, 70.16ºW with an hypocentral depth of 45 km

+/- 6km. That study indicates that the rupture took about 45 s to propagate about 150 km southward from the

epicentral area to the northern Mejillones Peninsula with an average rupture velocity of about 2.8 km/s. The slip

distribution inverted in this analysis shows that slip is located in the deeper part of the seismogenic zone,

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between 35 km and 50 km depth (Figure 2a). The slip is mainly concentrated around two asperities, one south of

the epicentral area and one north-east of the Mejillones Peninsula. The seismic moment associated with this

event is Mo=4.5 1020 N.m which corresponds to an earthquake of magnitude Mw=7.7.

Characteristics of the Tocopilla earthquake derived from geodesy

The geodetic data used to derive the slip models come from continuous GPS stations, where the signal is above

the noise level on daily solutions. Five Continuous GPS stations installed in that area through the collaboration

between the Institut de Recherche pour le Developpement (IRD), Institut de Physique du Globe de Paris (IPGP)

and the Departamento de Geofísica of Universidad de Chile (DGF) were operating during the event. Since the

two northern station show nearly no displacement associated with that event, the three southern stations, one

located in the longitudinal valley at the city of Quillagua (QUIL), and two at the costaline Tocopilla (TCPL) and

Mejillones (PMEJ) indicate significant trenchward horizontal displacements. About 22 cm of horizontal

displacements was observed at PMEJ station located about 100 km south from the epicenter and 5 cm at TCPL

and QUIL stations located respectively at about 25 km and 75 km north of the epicenter (Figure 1). PMEJ and

TCPL stations were uplifted respectively by 35 cm and 10 cm, since QUIL station subsided by 4 cm. GPS time

series of Mejillones and Iquique indicated a decreasing trenchward motion in the 45 days that followed the

event. During that period, the GPS station of Mejillones continues moving horizontally trenchward of about 5

cm, but no significant vertical displacements is observed in the time series.

Figure 1. GPS and InSAR measurements associated with of the Mw 7.7 Tocopilla earthquake. The horizontal GPS measurements (in black) and the vertical (shown in grey) have the same scale. InSAR data are unwrapped and presented in the line of sight (LOS) of the ERS-satellite. LOS positive indicate subsidence.

In addition to GPS measurements, InSAR interferograms were constructed using ASAR/ENVISAT images on

two adjacent satellite tracks and covering the epicentral area of Tocopilla to the city of Antofagatsa. The

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displacement in the line of sight of satellite shows that the coastline area moved toward the satellite with a

maximum of 30 cm, since the longitudinal valley moved away from the satellite of 15 cm (Figure 1). The InSAR

data covered as well the coseismic displacements associated with the Tocopilla earthquake but include also the

early postseismic displacements that occurred in the first 45 days after the earthquake. Because the small altitude

of ambiguity between the image pairs used and the precise SRTM DEM, we do not expect much noise

associated to the topography.

Results

The Joint inversion of InSAR data and GPS including postseismic deformation in about the first month

indicates that the seismic moment might be higher than 6.0 x 1020 Nm to fit the geodetic data relatively well. The

discrepancy with the seismic moment deduced from seismological observations may come from the fact that the

geodetic data may contain postseismic deformation up to 45 days after the earthquake. The slip distribution

indicates that coseismic slip may have occurred around two major asperities of ~50-70 km wide (Figure 2). One

asperity appears just north of the epicenter and slips about 1.5 m and the other is northeast of the Mejillones

Peninsula slipping about 2.5 m. Some significant postseismic afterslip may have occurred in between the two

seismic asperities but also off-shore the Mejillones Peninsula where significant aftershocks (with Mw>5.5)

occurred in the days that follow the earthquake (figure 1).

Figure 2. Slip distributions associated to the Mw 7.7 Tocopilla earthquake deduced from a) joint inversion of teleseismic and strong-motion data, b) coseismic GPS data, c) coseismic and postseismic GPS and InSAR data. The difference between b) and c) is presented in d) and may reflects the cumulative afterslip that occurred in the month that follow the Tocopilla earthquake.

From previous study of interseismic strain accumulation in that region, the deep section of the slab interface

located between 35 km and 55 km depth was defined as the brittle-ductile transition of the northern Chile

seismic gap (Chlieh et al. 2004). In the years before and after the 1995 Mw 8.1 Antofagasta earthquake that

ruptured a surface of about 200x100 km of the superficial portion of the subduction interface, < 35km depth

(Ruegg et al. 1996, Delouis et al. 1997), three earthquakes of Mw>7.0 occurred in the brittle-ductile transition

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zone (Pritchard et al. 2002). Then, the occurrence of the Mw 7.7 Tocopilla earthquake in that brittle ductile

transition zone could be suspected as a precursor to the failure of the northern Chile seismic gap.

Even thought its high magnitude, the Tocopilla earthquake did not released much of the seismic energy that is

accumulating along the northern Chile seismic gap. At the scale of that stretch of about 500 km-long by 150 km-

wide, the moment rate deficit estimated from interseismic deformation analysis is about 2.5 1020 N.m/yr. This

indicated that in less than 3-years of plate convergence rate, the seismic moment accumulated over the locked

fault zone of northern Chile seismic gap would be the equivalent of what was released during co- and

postseismic relaxation of the Mw 7.7 Tocopilla earthquake. Since the northern Chile segment did not produce any

great earthquake in 130 years, the moment deficit accumulated there would be enough to produce an event of

Mw 8.7 or higher if only half or more of that moment deficit is released.

References Chlieh, M., de Chabalier, J. B., Ruegg, J. C., Armijo, R., Dmowska, R., Campos, J., and Feigl, K. L., 2004. Crustal

deformation and fault slip during the seismic cycle in the North Chile subduction zone, from GPS and InSAR observations. Geophys. J Int., 158, p. 695-711.

Delouis, B. et al., 1997. The Mw = 8.0 Antofagasta (Northern Chile) earthquake of 30 July 1995: a precursor to the end of the large 1877 gap, Bull. seism. Soc. Am., 87, 427–445.

Delouis, B., Pardo, M., Legrand, D., Monfret, T. 2008. The Mw 7.7 Tocopilla earthquake 0f 14 November 2007 at the southern edge of the northern Chile seismic Gap: Rupture in the deep part of the coupled plate interface. Submitted to BSSA.

Pritchard, M. E., Simons, M., Rosen, P. A., Hensley, S., Webb, F. H. 2002. Co-seismic slip from the 1995 July 30 Mw= 8.1 Antofagasta, Chile, earthquake as constrained by InSAR and GPS observations. Geophys. J Int., 150 (2) , p. 362–376 doi:10.1046/j.1365-246X.2002.01661.x

Ruegg, J.C. et al., 1996. The Mw =8.1 Antofagasta (North Chile) earthquake of July 30, 1995: first results from teleseismic and geodetic data, Geophys. Res. Lett., 23(9), 917–920.

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Progressive avulsion of the Río Pastaza as a response to topographic uplift and backtilt of the Ecuadorian Subandean Zone

F. Christophoul1, C. Bernal

1, J. Darrozes

1, J.-C. Soula

1, & J. D. Burgos

2

1 UMR LMTG Universités de Toulouse-CNRS-IRD-OMP, 14 Avenue Edouard Belin 31 400 Toulouse, France

christo@lmtg.obs-mip.fr 2 Petroecuador, Ecuador

KEYWORDS : Ecuador, Subandean, avulsion, thrust related fold, backtilt

Introduction

Rivers are known to be very sensitive to local slope change (Jones and Schumm, 1999; Schumm et al., 2000;

Schumm, 1977). Response of alluvial rivers to slope change includes changes in fluvial morphology (Schumm et

al., 1987) and avulsions in the case of a downstream slope change (Jones and Schumm, 1999), (Stouthamer and

Berendsen, 2000) or lateral shift (Alexander et al., 1994) (Schumm et al., 2000) and avulsion (Alexander et al.,

1994) in the case of a lateral tilt. In Ecuador, the Rio Pastaza, while it crosses the structures of the Subandean

zone exhibits several sharp curves associated with traces of abandoned channels. This area is known to have

been recently uplifted in response to the growth of the Subandean Front (Bes de Berc et al., 2005).

This article is aimed illustrating a type of fluvial response to tectonics involved in the development of drainage

network on an active orogenic foothill. This study is based on field data, structural cross sections, remote

sensing, GIS and DEM. Data based on historic chronicles and testimonies of the inhabitants of the area were also

included.

Structure of the Ecuadorian Subandean Zone

The Ecuadorian Subandean Zone is bounded by the Western Cordillera to the west and the Amazonian

foredeep to the east (Fig. 1A, 1B), which is filled with deposits ranging from the upper Cretaceous to the

Holocene (Christophoul et al., 2002). The Subandean Zone consists in the Napo and Cutucu antiforms which

deform Jurassic volcanic and sedimentary formations and Cretaceous through Oligocene sedimentary formations

((Baldock, 1982); (Balkwill et al., 1995); (Baby et al., 1999); (Kley et al., 1999); (Bes de Berc et al., 2005)).

These antiforms developed on top of the west dipping Subandean Front. Between these antiforms, the Pastaza

Depression is topped by the Puyo Plateau which consists in a flat surface developed on top of the plio-

pleistocene fluvial sediments (Bes de Berc et al., 2005). A balanced cross-section traversing the front of the

Western Cordillera, the Interandean Depression, the Eastern Cordillera and the Subandean Zone along a line

Latacunga – Arajuno (Fig. 1) shows that the structure of the range in this area results from a complex sequence

of fold and thrusts. The Puyo plateau is known to have been recently uplifted in response to the growth of the

Subandean Front (Bes de Berc et al., 2005). Topographic profiles issued from a DEM derived from an Aster

image reveals that the Puyo Plateau exhibits a 0.4° westward tilt.

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Figure 1. Simplified structural map of Ecuador (W.C.: Western Cordillera, IAD: Interandean Depression, EC: Eastern Cordillera, SZ: Subandean Zone) and structural cross section through the Interandean Depression, the Eastern Cordillera and the Subandean Zone (Burgos, 2006). The white square corresponds to the area represented on Fig.2.

Stream changes

Changes of the courses of the Rio Pastaza downstream from Puyo were reconstructed thanks to the synthesis of

cartographic documents (1906, 1976) aerial photography (1976), Landsat (1987, 1992, 2000) satellite images.

The occurrence of the last floods in the area (2007) is documented by testimonies. Put together, these documents

allow us to trace the evolution of the location and the morphology of the Rio Pastaza over a century. The

topography of the area consist in a plateau (Puyo Plateau), bounded eastward by the steep slopes of the Western

Cordillera and eastward by the scarp of the subandean front (incised by the Rio Pastaza) affected by landslides to

the north. In the center of the Puyo Plateau a rounded shaped hill corresponding to Pliocene volcanics (Burgos,

2006).

1906: The first reliable document concerning the Rio Pastaza in the studied area consists in a map

published in 1906. By these times the Rio Pastaza went round the Pliocene volcanics northward and eastward.

1976: A mission of Aerial photography was led in 1976. It shows the Rio Pastaza as a braided river

made of 2 braid plains, one flowing southward along the western cordillera, the other going round the pliocene

volcanics southarward and eastward. The 1906 path is abandoned.

1987: The Río Pastaza shows the 2 same braid plains. It exhibits 2 bifurcations (X and Y). To the

north, bifurcation X parts the Río Pastaza in two channels. The western one (x') exhibits a braided channel

characterized by a low braiding parameter (BI=1.8). The morphology of the eastern one (x'') varies from single

channel to braiding (with a very low braiding parameter, BI=0.5). To the south, bifurcation Y divides the main

channel in a channels series. The main channel flowing southward (y') presents the same characteristic than

upstream from the bifurcation, (BI=1.6). While the channel series (y'') start exhibiting a single meandering

stream, flowing to the south east.

1992: Figure 2 (14/07/1992) reveals some changes concerning bifurcation X, since the divergence

point is displaced toward the south, ~1.8 Km, and implicates the virtually abandonment of x'' reach, which is

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 140-143

142

reduced to a straight single channel stream. While the braiding parameter, in X' reach, is 1.8. Bifurcation Y

shows a decrease in the braiding parameter (0.9).

Figure 2. Evolution of the Rio Pastaza morphology between 1906 and 2002. This reach of the Rio Pastaza locates on top of the Puyo plateau (white square on Fig. 1), the boundary between the subandean zone and the Eastern cordillera is located a few km upstream from Mera. See text for explanation and data sources. VC: topographic high of Pliocene volcanic rocks.

2000: On figure 2 (09/11/2000), we can see several changes: a) the braiding parameter of the western

reach located downstream of both bifurcation strongly increased (BI: 3.1 and 2, for bifurcation x' and y'

respectively); b) Downstream from bifurcation Y, we can see that series channels y'' disappeared, and we can

quote traces of former parallel channels.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 140-143

143

2002: Figure 2 (12/09/2002) reveals few differences compared to figure 5.c. Channel x'' keep reducing

its width. While, as a consequence of the disappearance of channel y'', the main channel (x'+y') reaches high

braiding parameter (BI:3 and 2 respectively).

2007: a flood of the Rio Pastaza reactivated the eastward channel of the Rio Pastaza. It shows that this

channel, though partially abandoned is still used as a chute channel by the Rio Pastaza.

Discussion and Conclusion

The evolution of the successive paths of the Rio Pastaza over the last century exhibits the characteristics of a

progressive avulsion. This progressive avulsion led the Rio Pastaza to describe a 90° curve by these times. This

avulsion consists in the initiation of an anabranch channel which progressively grows in size. By this time, the

original channel regularly decreases in size an keeps occupied by an underfitted stream. This underfitted stream

remains active in case of flood were it as reactivated as chute channel. This progressive avulsion seems to be the

consequence of the tilt of the Rio Pastaza substratum in response to the growth of the eastward merging

subandean front which tectonic activity is known in the last 20 Ka (Bes de Berc et al., 2005) and in the last years

(Legrand et al., 2005). Such a kind of progressive avulsion has been identified in other modern and ancient

rivers (Bristow, 1999) where was in channel processes triggered avulsions. In the case of the Rio Pastaza,

progressive avulsion is controlled by tectonics.

References Alexander, J., Bridge, J.S., Leeder, M.R., Collier, R.E.L. and Gawthorpe, R.L., 1994. Holocene meander-belt evolution in an

active extensional basin, southwestern Montana. Journal of Sedimentary Research, B64: 542-559. Baby, P., Rivadeneira, M., Christophoul, F. and Barragán, R., 1999. Style and timing of deformation in the Oriente of

Ecuador. In: Orstom (Editor), 4th International Symposium of Andean Geodynamics. ORSTOM, Göttingen, pp. 68-72. Baldock, J.W., 1982. Boletín de Explicación del Mapa Geológico del Ecuador. DGGM, Quito, Ecuador, pp. 80. Balkwill, H.R., Rodrigue, G., Paredes, F.I. and J.P., A., 1995. Northern part of Oriente Basin, Ecuador: reflection seismic

expression of structures. In: A.J. Tankard, R. Suarez Soruco and H.J. Welsink (Editors), Petroleum basins of south America, pp. 559-571.

Bes de Berc, S. et al., 2005. Geomorphic evidence of active deformation and uplift in a modern continental wedge-top - foredeep transition: example of the eastern Ecuadorian Andes. Tectonophysics, 399(1-4): 351-380.

Bristow, C.S., 1999. Gradual avulsion, river metamorphosis and reworking by underfitd streams: a modern example from the Brahmaputra river in Bangladesh and a possible ancient example in the Spanish Pyrenees. In: N.D. Smith and J. Rogers (Editors), Fluvial Sedimentology VI. Special Publication of the International Association of Sedimentologists. Blackwell Science, pp. 221-230.

Burgos, J.D., 2006. Mise en place et progradation d'un cône alluvial au front d'une chaîne active: exemple des Andes équatoriennes au néogène. Phd Thesis, Université Paul Sabatier, Toulouse 3, Toulouse, 373 pp.

Christophoul, F., Baby, P., Soula, J.-C., Rosero, M. and Burgos, J.D., 2002. Les ensembles fluviatiles néogènes du bassin subandin d'Equateur et implications dynamiques. Compte Rendus Géosciences, 334: 1029-1037.

Jones, L.S. and Schumm, S.A., 1999. Causes of avulsion: an overview. In: N.D. Smith and J. Rogers (Editors), Fluvial Sedimentology VI. Special Publication of the International Association of Sedimentologists. Blackwell Science, pp. 171-178.

Kley, J., Monaldi, C.R. and Salfity, J.A., 1999. Along-strike segmentation of the Andean foreland: causes and consequences. Tectonophysics, 301(1): 75-94.

Legrand, D. et al., 2005. The 1999-2000 seismic experiment of Macas swarm (Ecuador) in relation with rift inversion in subandean foothills. Tectonophysics, 395: 67-80.

Schumm, S.A., Dumont, J.F. and Holbrook, J.M., 2000. Active tectonics and alluvial Rivers. Cambridge University Press, Cambridge, 401 pp.

Schumm, S.A., Mosley, M.P. and Weaver, W.E., 1987. Experimental Fluvial Geomorphology. Wiley Interscience, New York, 411 pp.

Schumm, S.A., 1977. The Fluvial System. Wiley & sons, New York. Stouthamer, E. and Berendsen, H.J.A., 2000. Factors Controlling the Holocene Avulsion History of the Rhine-Meuse Delta

(The Netherlands). Journal of Sedimentary Research, 70(5): 1051-1064.

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Superimposed deformational episodes along the migmatitic belt, central portion of the Sierras Pampeanas Septentrionales, Central Andes, Argentina: An example from the Las Cañas Complex

Clara E. Cisterna1,2

, Ricardo Mon1,2

, & Rodolfo Mena2

1 Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET), Miguel Lillo 205, 4000 Tucumán,

Argentina (claracisterna@hotmail.com, monr@satlink.com) 2 Facultad de Ciencias Naturales e Instituto Miguel Lillo, Universidad Nacional de Tucumán (UNT), Miguel Lillo

205, 4000 Tucumán, Argentina (rodmena@gmail.com)

KEYWORDS : migmatites, syntectonic fold, shear, Central Andes, Argentina

Introduction

The northern Argentina Proterozoic - early Paleozoic

basement is composed of an assembly of multiply folded

orogenic belts with different structural characteristics,

amalgamated at about 600 Ma (Mon and Hongn 1996).

Preliminary studies allowed to recognize, at the central

portion of the Pampean belt, (río Las Cañas, cuesta de La

Chilca, La Majada, and others) a crystalline core

composed by gneisses and migmatites, marginated

eastward and westward by low grade schists separated by

west-dipping thrusts (Mon y Hongn 1996) (Fig. 1). This

migmatitic belt is the main object of this study. It

represents a deep magmatic arc, where the relative

synchronicity and/or feedback relations between heat,

deformation, partial melting and regional metamorphism

generated an igneous – metamorphic complex highly

deformed by multiple deformational episodes during a

contractional regime.

The Las Cañas Complex (LCC) represents a portion of

the metamorphic mid-crustal basement migmatitic belt in

the sierra de Aconquija (Sierras Pampeanas

Septentrionales, NW Argentina). A complete range from

folded metatexite to diatexite migmatites has been

produced during a high-grade metamorphic and

deformational events. The structural study reveals a shear

episode at the end of the migmatization. New field, fabric

and mineralogical observations allow constrain pre-, syn-

and post-migmatitic deformational phases and offers a

new insight into the tectonic evolution of the Central Andes crystalline basement in northwestern of Argentina.

Figure 1. Regional geological map (modified after Mon and Hongn 1996).

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Geological setting

The Sierra de Aconquija (NW Tucumán – SE Catamarca) belongs to the Sierras Pampeanas Septentrionales.

This range is mostly composed by Neoproterozoic – early Paleozoic basement rocks and whilst the general

characteristics of the metamorphism and the structure in this area are similar to elsewhere in these ranges

(González Bonorino 1951). Some of the salient features of the local geology will be briefly described here.

The layered schists that crop out along the studied area are formed predominantly of centimeters- to

decimeters- scale layers of fine to medium grained quartz - feldspatic, with intercalations of thin pelitic layers.

They were compared with the Ancasti Formation exposed at the south part of the Sierra de Ancasti. The regional

metamorphisms related to the low- to medium- grade metamorphic rocks are the result of an Early Cambrian

(Rapela et al., 1998, Sims et al., 1998) tectono-thermal event called the Pampean Orogeny. Other one, from

Ordovician to Devonian age (~ 490 - 390 Ma), produced during the Famatinian Orogeny is coincident with the

metamorphism peak in different areas of the Sierras Pampeanas (eg. Sierra de Ancasti). Deformation episodes

along this portion of Sierra de Aconquija were studied by Mon and Hongn (1996) and others. The first one

generated a D1 poorly preserved structure, characterized by the layered schists foliation. Other structural features

throughout most of the area correspond to a D2 deformation phase, which is represented by a crenulation

cleavage (or: the S2 penetrative fabric) oblique to S1 and characterized by mineral segregation and folding.

Migmatites were studied by different authors (Rassmuss 1918) and were defined as the “Complejo de Inyección”

(González Bonorino 1951) along the Aconquija – Ambato -Ancasti mountains.

Las Cañas Complex: field relations, structure and lithotypes

Most of the basement outcrops of this area are represented by migmatites and layered schists. The migmatites

represent the bulk of these outcrops; schists are represented only as resisters bodies included in the migmatitites.

Layered schists. Fine grained schists are the lowest-grade metamorphic rock, showing an S1 foliation,

characterized by the alternation of millimeter - thick grey platy minerals layers with quartz-rich microlithons

and composed by Qtz + Pl + Bt ± Ms and Qtz + Pl + Grt + Ms mineral assemblage (mineral symbols after

Kretz, 1983). The most abundant schists are grey, foliated and/or banded rocks. They show an S2 structures,

characterized by a crenulation cleavage and folds. Banded

fabric is given by quartz - feldespathic rich and biotite rich

layers alternations and are composed by Bt + Qtz + Pl +

Sil ± Ms + Ap + Zrn ± Ilm mineral assemblage. These

schists generally are present as resisters bodies nearly 1m

enclosed in migmatites and show the pre-migmatitic

deformation (Fig.2). Resisters schists commonly enclose

concordant quartzo-feldspathic veins, (10-20 cm thick),

folded and boudinated (Fig. 3).

Migmatites. These rocks display a wide variety of

morphologies. Stromatitic metatexites are the most common ones, characterized by the alternation of layers with

textural and mineralogical features well-developed.

Figure 2. Schists resisters in diatexite, showing the crenulation cleavage (pre-migmatitic structure)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 144-147

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They appear dark grey and have mainly biotitic composition, fine- grained mesosomes with Bt + Qtz ± Pl ± Sil ±

Grt mineralogical assemblages; and white to pink Qtz + Pl + Kfs ± Bt leucosomes. Diatexites have gradational

and sharp contacts with metatexites and are characterized by schlieren enriched in mafic minerals (mainly biotite

and granate). Resisters are common and represented by the schist protolith. Many diatexites outcrops show a

preference for platy minerals and plagioclase orientation. The transition from metatexite to diatexite migmatites

occurs at local scale over a few meters. Diatexite migmatites have coarse - grained equigranular fabric and a Qtz

+ Pl + Kfs ± Bt mineral assemblage characteristic. Migmatites develops flow foliation (S1) and parallel

intrafoliated folding (F1) (Fig. 3, 4). Most leucosomes are concordant or slight discordant to the dominant

structure (S1) and nearly 1 to 5 cm thick. Folds range between a few centimetres to more than one meter wave

length, and they also can be seen at the microscope. The resisters may be folded concordant to (F1) (Fig. 3). A

second folding episode (F2), showing N-S axial plane orientation and perpendicular to (F1) affect these rocks

(Fig. 4). The F2 structures are dipping to the east and folds, varying between one meter or a few centimetres, are

accompanied by an axial cleavage (S2). Refolding structures are common and generate interferences patterns,

examples of them are the partial or complete closures (“eye folds”) indicatives of the effects of the cross-folding.

Finally, over the most outcrops are recognized ductile shear zones (Fig. 5), associated with rotated and rounded

layered blocks and megablasts (eg. garnet).

Figure 3. Outcrop of stromatitic migmatite enclosing a foliated schists as resister. F1 and F2 folds overprinting.

Figure 4. Early fold (F1 ) with axial planar traces curved and F2 folds. Partial or complete clossures (“eye folds”) indicative of the effects of cross-folding.

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Discussion and Conclusions

It is argued that migmatitic rocks were developed during more of one ductile folding episode and shear

structures are related to later phases. A late post-migmatitic ductile-brittle deformation is evidenced by the

development of a NNW-SSE striking vertical spaced crenulation cleavage. The relationships between a large

diversity of migmatitic structures and the progressive production of melt suggest that feedback relations

prevailed as a time-marker during a contractional regime. Deformation, metamorphism and plutonism of the Las

Cañas Complex show that this terrane evolved as a east-verging thrust system with synkinematic metamorphism

and partial melting during the Late Proterozoic – Early Paleozoic. The tectonic history and lithologies of LCC

are compared to the sierra de Ancasti (La Majada Complex, southern to the studied area), showing that they

belong to a same regional belt within the Catamarca portion of the Pampean Ranges, Central Andes (Argentina).

References González Bonorino, F., 1951. Descripción geológica de la Hoja 12 g, Aconquija, Catamarca – Tucumán. Buenos Aires,

Dirección Nacional de Minería, 75, 50 p. Kretz, R., 1983 .. Symbols for rock-forming minerals. American Mineralogist, 68:277-279. Mon, R., Hongn, F. D., 1996 . Estructura del basamento proterozoico y paleozoico inferior del norte argentino. Revista de la

Asociación Geológica Argentina, 51 (1): 3-14. Rapela, C. W., Pankhurst, R. J., Casquet, C., Baldo, E., Saavedra, J., Galindo, C. & Fanning, C. M., 1998.” The Pampean

Orogeny of the southern Proto-Andes: cambrian continental collision in the Sierras de Córdoba”. In Pankhurst, R. J. & Rapela, C. W. (éd.): The Proto-Andean Margin of Gondwana, Geological Society, London, Special Publications 142: 181-217.

Rassmuss, J., 1918. “La sierra de Aconquija”. In: Primera Reunión Nacional de la Sociedad Argentina de Ciencias Naturales, Phycis, 47-69.

Sims, J. P., Ireland, T. R., Camacho, A., Lyons, P., Pieters, P. E., Skirrow, R. G., Stuart-Smith, P. G. & Miró, R., 1998. “U-Pb, Th-U and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin”. In Pankhurst, R. J. & Rapela, C. W. (éd.): The Proto-Andean Margin of Gondwana, Geological Society, London, Special Publications 142: 259-281.

Figure 5. Folds associated with a ductile shear deformation.

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Where is the evidence for Oligocene rifting in the Andes? Is it in the Loncopué Basin of Argentina?

Peter R. Cobbold1, Eduardo A. Rossello

2, & Fernando O. Marques

3

1 Geosciences (UMR6118, Université de Rennes 1 et CNRS), 35042 Rennes, France

(peter.cobbold@univ-rennes1.fr) 2 CONICET y Universidad de Buenos Aires, 1428 Buenos Aires, Argentina (rossello@gl.fcen.uba.ar)

3 Dep. Geologia e CGUL, Univ. Lisboa, Edifício C6, Piso 2, 1749-016 Lisboa, Portugal (fomarques@fc.ul.pt)

KEYWORDS : Andes, Oligocene, extension, Loncopué Basin, Argentina

Introduction

According to several authors, the Andes went through a phase of rifting in the Oligocene. Jordan et al. (2001)

described evidence for this from the Loncopué Basin of Argentina. This is between the volcanic arc of the main

Andean Cordillera and the uplifted western edge of the Neuquén Basin (Figure 1). The latter formed as a

composite rift basin during the early Mesozoic. Then Andean compression inverted it, in various stages, from the

middle Cretaceous onwards (Ramos, 1998; Cobbold and Rossello, 2003). The Loncopué Basin contains

Oligocene to Miocene continental strata. However, we have found in it little or no evidence for coeval extension.

Instead, we have found growth strata around folds and reverse faults. This makes us question the idea of

Oligocene extension. Our doubts extend to other parts of the Andes as well.

Figure 1. Left: simplified geological map of Neuquén Basin (modified after Cobbold and Rossello, 2003). Large rectangle indicates northern Loncopué Basin. Right: simplified geological map of northern Loncopué basin.

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Loncopué Basin

The Loncopué basin (37° - 38° S) is a topographic and structural depression, lying between the main volcanic

cordillera in the W and the Cordillera del Viento in the E (Figure 1). Some authors have claimed that this

depression is a rift basin (Ramos, 1988), whereas others have argued that it is a compressional foreland basin

(Lesta et al., 1985; Cobbold and Rossello, 2003). Neogene volcanic rocks and alluvial sediment cover much of

the depression, except where four river valleys cut through a buried N-S-trending fold-and-thrust belt. At the

eastern edge of this belt, Pliocene lava flows have been sharply folded into an eastward verging monocline. The

lavas are in strong angular unconformity upon intensely deformed strata of the Lileo Fm (Leanza et al., 2002;

Rovere, 2004), which is equivalent to the Cura-Mallín Fm in Chile. The Lileo Fm consists of mudstone,

limestone, sandstone and tuff, which accumulated in lacustrine and fluvial environments, during the late

Oligocene to early Miocene. The age range (28 Ma to 22 Ma) is well constrained by 39Ar-40Ar and fission-track

dating of volcanic rocks (Jordan et al., 2001; Burns et al., 2006), fossil spores (Leanza et al., 2002), and fresh-

water mussels (Diplodon sp.; Uliana, 1979; Burns et al., 2006). According to various authors, the Lileo Fm

accumulated in a rift basin, during a phase of Oligo-Miocene extension, and this was later inverted in

compression (Folguera and Ramos, 2000; Jordan et al., 2001; Burns et al., 2006; Folguera et al., 2006).

However, the evidence for extension in the Loncopué basin is poor. Structural data are lacking, stratigraphic

arguments are inconclusive, seismic data are of bad quality, and there are no well data to constrain them at depth.

In the Lileo valley, outcrops of the Lileo Fm are continuous and of good quality (Figure 2). The entire sequence,

some 3000 m thick, has been folded. Burns et al. (2006, their figure 9) and Folguera et al. (2006, their figure 4)

have drawn sections, in which the strata are of equal thickness around folds. The authors conclude that the folds

are post-sedimentary and due to a Miocene phase of rift inversion. However, a panoramic view (Figure 2) shows

that in fact the bed thicknesses vary significantly around the folds, and the dips vary through 35°, 45° and 55° on

the limbs. Because many of the beds are lacustrine, we exclude the possibility of large initial dips. Thus the folds

are due to syn-sedimentary deformation. We infer that the Lileo Fm accumulated, not during extension, but

during horizontal shortening and vertical thickening. This begs two questions. Was the shortening thin-skinned

or did it involve the basement? Was the shortening of local or regional significance? Unfortunately, existing

seismic data (Jordan et al., 2001, their figure 7) are of poor quality. Neighbouring outcrops provide better clues.

The Cordillera del Viento (Figure 1), nearly 3000 m high, has a carapace of Permo-Triassic volcanic rocks

(Choiyoi Fm). Its deeply eroded western scarp provides a window into an underlying Palaeozoic sequence of

Lower Carboniferous tuffs, Upper Carboniferous marine shale, and early Permian tuffs (Zöllner and Amos,

1973). This sequence repeats above a flat-lying westward-verging thrust, which has been cut by Permian granite.

The thrust and the deep-seated intrusion are relicts of Palaeozoic orogenesis. The Choiyoi Fm lies in strong

angular unconformity upon the deformed and eroded Palaeozoic rocks. Above the Choiyoi Fm is a more

conformable sequence of mainly marine strata, which are typical of the Neuquén basin and go from early

Jurassic to middle Cretaceous. At some moment between the Hauterivian and the Eocene, the Mesozoic

sequence was further folded and back-thrust to the west, probably by reactivation of the underlying Palaeozoic

thrust system. Thus the Cordillera del Viento is a large eroded antiformal culmination, above a back-thrust

(dashed trace, Figure 1). Next to this fault, andesitic rocks of the Cayanta Fm are unconformable upon tilted

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 148-151

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strata of various ages (from early Carboniferous, through Permian, Triassic and Jurassic, to middle Cretaceous).

The Cayanta Fm has been dated as Eocene by the 39Ar-40Ar method (Jordan et al., 2001).

Figure 2. Panoramic view of growth strata in Lileo Fm, along middle reaches of Lileo river valley. View is to south. Field of view is about 4 km, angle of view is about 60°, relief is about 500 m. Bedding traces (enhanced by white or black lines) are continuous over large areas. On limbs of syncline and anticline, true dips vary through angles of 35°, 45° and 55° (sectors between dashed lines).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 148-151

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Eocene andesites erupted through a series of vents that lie mainly along the back-thrust or close to it. To the W

of Andacollo, the Eocene andesites have been folded and uplifted by another back-thrust, which is active today

(full black trace, Figure 1). Probably it was active also during deposition of the Lileo Fm. In outcrops near the

confluence of the rivers Lileo and Neuquén, the Lileo Fm onlaps tilted Eocene andesites. On this basis, we

suspect that the Lileo Fm accumulated in a compressional or transpressional foreland basin. At one quarry in the

lower Lileo valley (Leanza et al., 2002, their figure 3), we found one syn-sedimentary strike-slip fault, left-

lateral and trending NW-SE, within mudstones of the Lileo Fm. Otherwise, we found no evidence for strike-slip

deformation in the area.

Conclusions

In the northern Loncopué basin, there are compressional growth strata within late Oligocene to early Miocene

continental strata of the Lileo Fm. Although there is no guarantee that the shortening involved basement,

neighbouring outcrops point to episodes of basement-involved shortening in the Palaeozoic, late Cretaceous to

Eocene, late Oligocene to early Miocene, and Quaternary.

Nowhere in this area (nor anywhere in the Neuquén Basin) have we found firm evidence for a phase of thick-

skinned extension in the late Oligocene to early Miocene, as claimed by Jordan et al. (2001) and Burns et al.

(2006). The only well-substantiated rifts are those that formed in the early Mesozoic. We cannot rule out

Tertiary rifting, nor do we exclude strike-slip motions (for which we have little evidence). However, we urge

prudence in inferring phases of rifting in the Andes, especially if the arguments are purely stratigraphic or the

structural evidence is incomplete.

References Burns, W.M., Jordan, T.E., Copeland, P. & Kelley, S.A. 2006. “The case for extensional tectonics in the Oligocene-Miocene

Southern Andes as recorded in the Cura Mallín basin (36°-38°S)”. In Kay S.M., Ramos V.A. (eds): Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°- 39°S lat), Geological Society of America Special Paper 207: 163-184.

Cobbold, P.R. & Rossello, E.A. 2003. Aptian to recent compressional deformation, foothills of the Neuquén Basin, Argentina. Marine and Petroleum Geology 20: 429-443.

Folguera, A. & Ramos, V.A. 2000. Control estructural del volcán Copahue (38°S-71°O): implicancias tectónicas para el arco volcánico cuaternario (36-39°S). Asociación Geológica Argentina Revista 55: 229-244.

Folguera, A., Ramos, V.A., González Díaz, E.F. & Hermanns, R. 2006. “Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt in the Neuquén Andes between 37° and 37° 30’S”. In Kay S.M., Ramos V.A. (eds): Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°- 39°S lat), Geological Society of America Special Paper 207: 247-266.

Jordan, T.E., Burns, W.M., Veiga, R., Pángaro, F., Copeland, P., Kelley, S. & Mpodozis, C. 2001. Extension and basin formation in the southern Andes caused by increased convergence rate: A mid-Cenozoic trigger for the Andes. Tectonics 20: 308-324.

Leanza, H.A., Volkheimer, W., Hugo, C.A., Melendi, D.L. & Rovere, E. 2002. Lutitas negras lacustres cercanas al limite Paleógeno-Neógeno en la region noroccidental de la provincia del Neuquén: Evidencias palinológicas. Asociación Geológica Argentina Revista 57: 280-288.

Lesta, P.J., Digregorio, J. & Mozetic, M.A. 1985. Presente y futuro de la exploración de petróleo en las cuencas subandinas, Argentina. II Simposio Bolivariano, Exploración Petrolera en las Cuencas Subandinas, Bogotá, Publicaciones 3: 1-35.

Ramos, V.A. 1998. Estructura del sector occidental de la faja plegada y corrida del Agrio, Cuenca Neuquina, Argentina. X Congreso Latinoamericano de Geología y VI Congreso Nacional de Geología Económica, Buenos Aires, Actas 2: 105-110.

Rovere, E.I. 2004. Hoja Geológica 3772-IV, Andacollo, Provincia del Neuquén. SEGEMAR, Programa Nacional de Cartas Geológicas de la República Argentina, 1:250.000, Boletín 298: 1-104.

Uliana, M.A. 1979. Geología de la region comprendida entre los ríos Colorado y Negro, Provincias de Neuquén y Rio Negro. Doctoral thesis, Universidad Nacional de La Plata, Argentina.

Zöllner, W. & Amos, A.J. 1973. Descripción geológica de la Hoja 32b, Chos Malal (Provincia del Neuquén). Servicio Nacional Minero Geológico, Buenos Aires, Boletín 143: 1-91.

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Burial history and estimation of ancient thermal gradients in deep synorogenic foreland sequences: The Neogene Vinchina Basin, south-Central Andes

Gilda Collo & Federico M. Dávila

Laboratorio de Análisis de Cuencas, CICTERRA-Universidad Nacional de Córdoba, Av. Vélez Sársfield 1611,

2º piso, of. 7, X5016GCA Córdoba, Argentina (gcollo@efn.uncor.edu, fmdavila@efn.uncor.edu)

KEYWORDS : Neogene Andean foreland, Vinchina basin, burial history, illitization progress, paleogeothermal gradients

Introduction

The Miocene to Pliocene Vinchina basin (Vinchina and Toro Negro formations) constitutes one of the thickest

foredeep sequences of the Central Andes, between the High Cordillera and the Sierras Pampeanas province

(~ 28° SL). In spite of its great thickness, locally >10 km, strata lacks of diagenetic signatures, even in the basal

section of the stratigraphic column, where it is expected to be identified (assuming middle geothermal gradients

of 20-30ºC/km) rocks with evidences of low-grade metamorphism.

In order to unravel the burial history of the Vinchina basin, and to estimate a Mio-Pliocene paleogeothermal

gradient for the Central Andean region, we evaluated the progression of the illitization processes on fine-grained

rocks that affected this sandy-silty dominated alluvial succession. Relations between interstratified clay mineral

distribution and temperature (eg., Arostegui et al., 2006; Srodon, 2007) allowed to estimate the diagenetic

history and the maximun burial conditions. Clay mineral associations were identified by X-ray analysis in <2μm

fractions of 5 samples from bottom to top in the Quebrada de los Colorados section (28°41'S, 68°16'W; La Rioja

Province). Relative proportions of interstratified illite/semectite (I/S) within the neoformed phases were

established from decomposition of the XRD diagrams (cf. Lanson, 1997).

Progression of illitization process

In the analyzed samples of the Vinchina basin, the dominant neoformed clay mineral phases are illite and

interstratified illite/smectite (I/S), with lesser amounts of chlorite. Detailed analysis of expandable I/S allowed

establishing the coexistence of interstratified with R0, R1 and R3 orderings. In the shallowest sample (uppermost

section of the Toro Negro Formation, Figure 1) the clay mineral assemblage is dominated by R0 (~70%), R1 and

illite phases, with absence of R3 ordering. The appearance of R3 takes place at ~5 Km depth within the Vinchina

Formation and, likewise R1 and illite, shows an increment towards the base of the unit. Although randomly,

mixed-layered R0 clearly decreases to the deeper levels, significant proportions (~30%) in the I/S phases are still

present in the lowermost analyzed sample (depth of ~7 Km).

Burial history and paleo-geothermal gradient estimation

Distribution of I/S interstratified phases through the succession allow establishing a progressive smectite-

illitization process (R0 R1 R3 I) related to the sedimentary burial history of the Tertiary sequence, as

shown by the increment in the I/S ordering and illite content from top to base of the units and a strong

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correlation between the degree of illitization and the stratigraphic age of the rocks. The presence of R0 randomly

mixed-layered I/S even in the deepest levels (~7 Km) evidences that the base of the basin fill did not exceed the

diagenetic field (cf. Frey and Robinson, 1999). Given that R0 becomes unstable over temperatures of 120ºC a

maximum paleogeothermal gradient of 17ºC/Km can be estimated. This value is coherent with

thermochronological studies that suggested the sequences would not have exceeded temperatures of ~90ºC

(Coughlin, 2000; Carrapa et al. 2006). Our values are also consistent with the coldest geothermal records

reported in the modern Andean foreland (see Hamza et al., 2005).

An Early Miocene-Pliocene age (between 19 - 3.4 Ma) for the Vinchina stratigraphy is interpreted from new

geochronological data, (Re y Barredo, 1993, Dávila et al., this congress) and is coincident with the onset of the

flat subduction at these latitude (Kay et al., 1988; Dávila et al., 2004). Within this context, crustal refrigeration

could associate to modifications in the thermal structure by reduction of the astenospheric wedge. Supporting

this hypothesis, similar geothermal gradients (18-20ºC/Km) were calculated for 5-km depth oil boreholes in the

Bermejo Valley (Precordillera de San Juan), which also above the modern flat-slab segment. Exhumation ages

for the Tertiary package are, however, not well constrained. The presence of subhorizontal Pleistocene(?)

conglomerates (Santa Florentina Fm) unconformably lying above the Vinchina and Toro Negro Formation on

the same thrust sheet allows interpreting that exhumation would have occurred between the youngest age of

Toro Negro Fm (3.4 Ma) and the deposition of this coarse conglomeratic succession. Given that the age of

deposition of the Santa Florentina Fm is unknown, but considered Pleistocene sensu lato, the maximum

residence interval for the Tertiary under extreme burial conditions would be 3.4 my.

Although the estimated maximum paleogeothermal gradient (~17ºC/Km) is consistent with those from other

coldest foreland basins, with the available thermocronologic data we cannot discard the influence of other factors

in the persistence of R0 I/S over the 7 km depth. “Effective K+ concentration” (Cuadros, 2006) and “residence

time” of the sequences at maximum burial conditions may have interfered in the evolution of clay mineralogy,

retarding the progression in the I/S ordering. Modeling of smectite-illitization process and comparisons between

correlative exhumed and buried successions will allow further evaluation of the influence of these factors.

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Figure 1: I/S Interstratified distribution through the estratigraphic column of the Vinchina Basin (Modified from Ramos, 1970). To the right the decomposition of XRD Diagrams of each sample show the clear diminishing, but without desapearing, of R0 I/S from top to bottom of the sequence.

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References Arostegui, J., Sangüsa, F.J., Nieto, F. and Uriarte, J.A. 2006. Thermal models and clay diagenesis in the

Tertiary-Cretaceous sediments of the Alava block (Basque-Cantabrian basin, Spain). Clay Minerals 41: 791-809.

Coughlin, T.J., 2000. Linked origen-oblique fault zones in the Central Argentine Andes: The basis of a new model for Andean orogenesis and metallogenesis. PhD Theses, University of Queensland, Dep. Earth Sciences.

Cuadros, J. 2006. Modeling of smectite illitization in burial diagenesis environments. Geochimica et Cosmochimica Acta 70: 4181-4195.

Dávila, F.M., Astini, R.A., Jordan, T.E., and Kay, S.M., 2004. Early Miocene andesite conglomerates in the Sierra de Famatina, broken foreland region of western Argentina, and documentation of magmatic broadening in the south-central Andes. Journal of South American Earth Sciences, 17: 89-101.

Dávila, F.M., Collo, G., Astini, R.A., and Gehrels, G., 2008. U-Pb detrital ages on a tuffaceous sandstone sheet in the Vinchina Formation, La Rioja, Argentina: Deposition and exhumation implications. XIII Congreso Geológico Argentino.

Frey, M. y Robinson, D., 1999. Low grade metamorphism. Blackwell Science. Cambridge. 313 p. Hamza V.M., Silva Dias, F.J.S., Gomes, A.J.L., and Delgadilho Terceros, Z.G. 2005. Numerical and functional

representations of regional heat flow in South America. Physics of the Earth and Planetary Interiors 152: 223–256.

Kay, S.M., Maksaev, V., Moscoso, R., Mpodozis, C., Nasi, C., and Gordillo. C.E., 1988. Tertiary Andean magmatism in Chile and Argentina between 28ºS and 32ºS: correlation of magmatic chemistry with a changing Benioff zone. Journal of South American Earth Sciences, 1: 21-38.

Lanson, B., 1997. Decomposition of experimental X-Ray diffraction patterns (Profile fitting) a convenient way to study clay minerals. Clays and Clay Minerals, 45 (2): 132-146.

Ramos, V.A., 1970. Estratigrafía y estructura del Terciario en la sierra de los Colorados (Provincia de La Rioja), República Argentina. Revista de la Asociación Geológica Argentina, 25 (3): 359-382.

Re, G.H. and Barredo, S.P., 1993. Esquema de correlaciones de las formaciones terciarias aflorantes en el entorno de las Sierras Pampeanas y la Precordillera Argentina. XII Congreso Geológico Argentino y II Congreso de Exploración de Hidrocarburos. 2: 172-179.

Srodon, J. 2007. Illitization of smectite and history of sedimentary basins. Euroclay. 74-82

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Coeval subduction erosion and underplating associated with a crustal splay fault at the Ecuador-Colombia convergent margin

Jean-Yves Collot1, Alessandra Ribodetti

1, Boris Marcaillou

2, & William Agudelo

3

1

Géosciences Azur, Université de Nice Sophia-Antipolis, IRD, CNRS, Université Pierre et Marie Curie,

Observatoire de la Côte d’Azur, BP 48, 06235 Villefranche-sur-mer, France (collot@geoazur.obs-vlfr.fr) 2 IFREE/JAMSTEC, 2.15 Natsushima-cho, Yokosuka, Kanagawa, 237-0061, Japan

3 ECOPETROL ICP.Km 7 Via Piedecuesta, Piedecuesta, Santander, Colombia

KEYWORDS : subduction zone, seismic reflection, subduction channel, erosion, underplating

Introduction

Subduction erosion and underplating are major processes governing the structural evolution of convergent

margins. Subduction erosion is required at many margins by large-scale, long-term margin subsidence and is

likely to be driven by over pressured fluids that disaggregate the underside of the margin basement (von Huene

et al., 2004). In contrast, underplating has been invoked to thicken accretionary margins by duplex formation at

the base of the accretionary complex (Park et al., 2002; Bangs et al., 2004). Based on seismic reflection,

refraction and swath bathymetric data, we show that both subduction erosion and underplating occur

simultaneously at an erosive segment of the North-Ecuador-South Colombia margin. The margin consists of an

accreted oceanic terrane overlain by thick fore-arc basin deposits (Jaillard et al., 1995), and underthrust eastward

at 5.4 cm/yr by. the Neogene Nazca plate (Trenkamp et al., 2002)(Fig. 1).

Data

In 2000, the SISTEUR cruise onboard the French R.V. Nadir acquired deep marine multichannel seismic

reflection (MCS) and wide angle seismic data using Ocean Bottom Seismometer across the Ecuador and south

Colombia margin to investigate its yet poorly-known deep structures (Collot et al., 2002). In 2001, the Salieri

cruise onboard the German R.V. Sonne acquired complementary wide-angle seismic data and multibeam

bathymetric data to explore crustal and seafloor structures in the same region (Flueh et al., 2001). In 2005, the

AMADEUS cruise onboard the French R.V. L’Atalante collected 55000 km2 of contiguous swath bathymetry

coverage and underway geophysics, sedimentary cores, dredged rocks, and heat flow data at 12 core- and heat

probe-sites between 0° and 3°30N (Collot et al., 2005).

Results and Interpretation

These data have allowed discovering the seafloor trace of the trench-parallel, ~ 90 km-long, Ancon fault

system, which extends across and north of the Esmeraldas canyon along the north Ecuador-south Colombia

margin (Fig. 2). The fault system separates a shallow outer basement high from the Manglares fore-arc basin,

and is segmented along strike. The N57°-trending, southern fault segment deforms the seafloor by extensional

faulting associated with an anticline. The N25°-trending northern fault segment is characterized by a high-angle

crustal reverse fault that fans out northward into a horsetail pattern.

Reprocessing multichannel seismic reflection (MCS) line SIS-44 (Fig. 2) through Prestack Depth Migration,

using vertical reflection and wide-angle data to construct a refined velocity model (Agudelo, 2005), shows the

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deep margin and interplate structures across the southern Ancon fault segment. A landward dipping reflector,

coupled with a remarkable lateral velocity contrast between 4-5 km/s outer wedge basement rocks and

5-6.2 km/s inner wedge basement rocks, supports the existence of a crustal splay fault (Fig. 3) associated with

the summit graben and anticline described along the Ancon Fault southern segment. The splay fault soles out on

the plate interface near a 12-15-km-depth. The fault is associated with a several km-wide, low velocity shear

zone interpreted as a major conduit for fluid flows. Fluids migrating from the subduction channel have altered

outer wedge basement rocks and lowered their velocities and mechanical strengths (Fig. 3).

Downdip, the subduction channel shows two segments with different seismic characters that reflect contrasting

processes acting on the plate interface. The updip segment of the subduction channel extends to a 9 km-depth, is

poorly reflective, 1.0-1.3 km thick, and was assigned a ~3.5 km/s velocity based on modeling wide-angle data.

The subduction channel poor reflectivity may be indicative of weak porosity contrast and suggests that fluids

transported with underthrust sediment have pervasively invaded and altered the overlaying basement, thus easing

basal erosion. Subduction erosion is further substantiated by thinning of the outer wedge basement associated

with clear trenchward tilt of the westernmost part of the fore-arc basin. The deep segment of the subduction

channel, from ~ 9 to 15 km depths, decreases irregularly in thickness from 1.3 km to less than 0.6 km, and is

characterized by a ~ 3.5 to 3.8 km/s low velocity zone relative to overlaying 4.5-5.5 km/s basement rocks. Near a

depth of 11-15 km, the SC shows strong internal, sigmoid reflectors that form toplaps beneath a continuous

landward-dipping surface. These reflectors are compatible with imbricated layers that dip locally seaward and

Figure 1: Bathymetric map (km) of the Nazca plate and adjacent North Andean margin derived from satellite altimetry data (Smith and Sandwell, 1997). Location of the study area offshore Ecuador and Colombia. Plate convergence after Trenkamp et al., (2002).

Figure 2: Swath bathymetry (contours 100 m) of the North Ecuador-South Colombia Margin (Collot et al., 2005) with location of the combined wide-angle (WA) and multichannel seismic reflection (MCS) line SIS-44. Black circles indicate the position of Ocean Bottom Seismometers (OBS). Barbed line is the deformation front and black arrow is Nazca-South America relative plate motion after Trenkamp et al., (2002). SS is seafloor swell.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 156-159

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are truncated by a roof thrust. According to the model proposed by Boyer and Elliott, (1982), we interpret the

imbricated layers as forward-dipping duplexes and antiform stacks that developed from, and structurally above

the subducting mélange.

Model and Conclusions

The erosive segment of the North Ecuador-South Colombia margin is characterized by basal erosion and

underplating, two processes that occur synchronously along two segments of the plate boundary. Along the

shallow segment of the plate boundary, fluid-altered rocks from the underside of the outer wedge basement are

torn away piecemeal. Debris are incorporated into the subducting mélange and dragged down dip in the

subduction channel. Near the junction between the splay fault and the plate interface, the décollement thrust is

forced to step down, and horses are interpreted to detach from the subducting mélange. Horses are thrust seaward

between a roof and a décollement thrust, and stack at the back of previously underplated duplexes. One-km thick

underplated material may be responsible for a swell in the fore-arc basin seafloor (Fig.2 and 3), where recent

sediment are being truncated. The underplated material is progressively driven seaward, counter-subduction,

above the descending subducting mélange, up to the region where underplated material is eroded and re-

incorporated to the subduction channel (Fig. 3). Coeval erosion and underplating control the margin mass

budget, which, in this case, comes out negative because of the seaward tilt and subsidence of the outer wedge.

These observations suggest that underplating is a transient process, and that the subducting mélange ultimately

recycles deeper in the subduction zone, when part of the mélange passes beyond a critical point, here defined as

the root zone of the splay fault. Both basal erosion and underplating appear to be facilitated by pre-existing

crustal splay faults.

Figure 3: Interpreted cross-section along line SIS-44 showing coeval basal erosion beneath the outer wedge and underplating beneath the inner wedge. Downgoing plate sediments are dragged in the subduction channel together with debris removed from the base of the outer wedge forming the subducting mélange. The margin wedge basements are separated by a major splay fault. Part of the subducting mélange cannot bypass the junction between splay fault and plate interface, and is underplated as seaward-dipping duplexes, promoting a reverse flow of material that propagates trenchward prior to be eroded at the base of the outer wedge. Hatched areas are sheared and/or dominantly fluid-altered basement rocks interpreted from seismic reflectivity and relatively low Vp velocity.

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References Agudelo, W., Imagerie sismique quantitative de la marge convergente d'Equateur-Colombie, 2005. PhD thesis, Université

Pierre et Marie Curie, Paris, 203 p. Bangs, N., T.H. Shipley, S.P.S. Gulick, G.F. Moore, S. Kuromoto, and Y. Nakamura, 2004. Evolution of the Nankai trough

décollement from the trench into the seismogenic zone : inferences from three-dimensional seismic reflection imaging, Geology, 32 (4), 273-276.

Boyer, S.E. and D. Elliott, 1982. Thrust systems, The American Association of Petroleum Geologists Bulletin, 66 (9), 1196-1230.

Collot, J.-Y., P. Charvis, M.A. Gutscher, and S. Operto, 2002. Exploring the Ecuador-Colombia active margin and inter-plate seismogenic zone, EOS Transactions, American Geophysical Union, 83 (17), 189-190.

Collot, J.-Y., S. Migeon, G. Spence, Y. Legonidec, B. Marcaillou, J.-L. Schneider, F. Michaud, A. Alvarado, J.-F. Lebrun, M. Sosson, and A. Pazmiño, 2005. Seafloor margin map helps in understanding subduction earthquakes, EOS Transactions, American Geophysical Union, 86 (46), 464-466.

Flueh, E.R., J. Bialas, P. Charvis, and Salieri scientific party, 2001. Cruise report SO159 SALIERI, in Report 101, pp. 256, Geomar Research center, Kiel, Germany.

Jaillard, E., M. Ordoñez, S. Benitez, G. Berrones, N. Jimenez, G. Montenegro, and I. Zambrano, Basin development, 1995. In an accretionary, oceanic-floored fore-arc setting: southern coastal Ecuador during late cretaceous-late eocene time, in Petroleum basins of South America, edited by A.J. Tankard, R. Suarez, and H.J. Welsink, pp. 615-631.

Park, J.-O., T. Tsuru, S. Kodaira, P.R. Cummins, and Y. Kaneda, 2002. Splay fault branching along the Nankai subduction zone, Science, 297, 1157-1160.

Smith, W.H.F., and D. T. Sandwell, 1997. Global seafloor topography from satellite altimetry and ship depth soundings, Science, 277, 1957-1962.

Trenkamp, R., J.N. Kellogg, J.T. Freymueller, and P. Mora, H. 2002. Wide plate margin deformation, southern Central America and northwestern South America, CASA GPS observations, Journal of South American Earth Sciences, 15, 157-171.

von Huene, R., C.R. Ranero, and P. Vannucchi, 2004. Generic model of subduction erosion, Geology, 32 (10), 913-916.

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Active tectonics in the Central Chilean Andes: 3D tomography based on the aftershock sequence of the 28 August 2004 shallow crustal earthquake

Diana Comte1, Marcelo Farías

1,2, Reynaldo Charrier

2, & Andrea González

2

1 Dept. Geofísica, Universidad de Chile, Santiago, Chile (dcomte@dgf.uchile.cl)

2 Dept. Geología, Universidad de Chile, Santiago, Chile (mfarias@dgf.uchile.cl)

KEYWORDS : crustal seismicity, tectonics, Central Andes, 3D tomography

Introduction

Most of the seismological research in the Andes has been mainly oriented to the detection and understanding

of the seismicity associated with the large thrust-fault earthquakes that characterize the subduction environment

that governs tectonics in this region. However, the growing number of stations in the permanent seismological

network and the deployment in the last years of temporary networks in different regions of the country have

allowed the detection of intense crustal seismicity beneath the Chilean forearc-arc region.

For instance, the temporary seismic network deployed along the Las Leñas and Pangal river valleys (34°25'S),

between January and May 2004, permitted to better constrain the abundant shallow intra-continental seismicity

previously detected in that region. Although most of the seismicity is randomly distributed in the region, several

microearthquakes occur along the trace of the major El Diablo - El Fierro fault-system. This fault, recognized

between 33°30' and 35°15’S, is located at or close to the eastern contact between Mesozoic and Cenozoic

deposits in the Principal Cordillera and, locally, below active volcanoes , and is considered to have participated

in the development (extension) and tectonic inversion of a widely extended (>600 km long) Cenozoic

extensional basin along the Principal Cordillera. The associated seismic activity implies that this structure is still

active and participates in the present-day adjustments of the Andean crust. Further south, at 35°S, a Mw=6.5

strike-slip shallow (<10 km) earthquake occurred on August 28, 2004, generating moderate damage in the

region, reaching a maximum intensity VI MM.

The Seismological Service of the University of Chile deployed a local network to monitoring the aftershock

sequence. A 3D detailed Vp and Vs velocities determination was obtained along the aftershock area of the 2004

earthquake; results show an essentially NS distribution reaching depths lower than 15 km. This behaviour is in

agreement with that observed further north, in the Las Leñas - Pangal region. The 2004 shallow earthquake is the

second one recorded by local networks in Chile, the previous one occurred in the northern Chile forearc in 2001

(Mw=6.2) (Farías et al., 2005). The 2004 shallow earthquake is similar to the major intraplate Las Melosas

earthquake (Mw=6.9) occurred on September 4, 1958, possibly associated with the El Diablo - El Fierro fault-

system. The occurrence of the 2004 earthquake offers the possibility to analyze this seismicity from a

seismotectonic point of view, in order to understand the present-day crustal adjustments.

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Geological setting

Figure 1. Simplified structural map of the Central Andean region between 32º30'S and 36ºS. Epicentral location and focal mechanisms of the 28 August, 2004 (Mw=6.7) and the 12 September, 2004 (Mw=5.3) earthquakes.

Four main morpho-structural units compose the continental region of the southern Central Andes between 33ºS

and 36ºS (Figure 1): The Coastal Cordillera, the Central Depression, the Principal Cordillera, and the Frontal

Cordillera. The Coastal Cordillera is a smooth relief constituted mainly by Paleozoic and Jurassic plutonics

rocks in the western part, and east dipping Cretaceous volcanics and sediments. The Central Depression is an

erosional feature partially filled with late Neogene clastics and ashy deposits. The Principal Cordillera is mainly

composed by volcanic and sedimentary rocks deposited in an extensional basin during Late Eocene to Late

Oligocene times, inverted during the Early Miocene that forms the bulk of the west versant of the cordillera

(Charrier et al., 2002, 2005). West-vergent and east-vergent thrusts delimitate the western and eastern border of

the former basin, respectively. The eastern border corresponds to the westernmost structures of the Aconcagua

(north of 33º45'S) and the Malarguüe (south of 33º45'S) fold-and-thrust belts that accommodated accommodated

most of the shortening in this part of the Andes (Giambiagi et al., 2003).

The 3D tomography

Farías et al. (2008) analyzed the crustal-scale structural architecture of the Central Andes and its implications

in the mountain building in subduction zones. He used the seismologic data recorded by the Seismological

Service of the University of Chile between 1980 and 2004, and the data obtained by the temporary network

deployed from January to April 2004. The permanent network has 24 seismologic stations in the study region

and the temporary network consisted of 7 short period 3-component stations, their final database includes 23,444

events, with 212 shallow (<20 km depth) crustal events recorded by the temporary network. In this work we add

about 500 shallow aftershocks recorded by 8 short period seismological stations deployed by the Seismological

Service to monitor the aftershock sequence of the 2004 curstal earthquake.

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The hypocenters were first estimated using the HYPOINVERSE program (Klein, 1978) with a 1D P-wave

velocity model based on Thierer et al. (2005). Each earthquake was located with different trial depths in order to

minimize the effect of the initial conditions on the final hypocentral determination. Trial depths were varied

between 0 and 100 km with an increment of 5 km. The location with the lowest root mean square misfit and with

the maximum number of body-waves first-arrivals was selected for each event.

Figure 2. Epicentral distribution of the aftershocks of the 28 August 2004 (Mw=6.7)

earthquake, white circles represent earthquakes with depths between 0 and 15 km, and yellow circles represent earthquakes with depths between 15 and 30 km. The red square shows the study area for the tomography.

From the preliminary hypocenters and seismic wave arrival times, a 3D velocity structure was calculated (see

details in Roecker et al. (1993)). Inversion was made on a region divided into 6 7 blocks with a grid spacing of

10 10 km2 and 12 layers of 5 km thick for the shallower layers (depth < 20 km) and each 10 km for the deeper

ones (Figure 2). Because P-wave and S-wave velocities were inverted independently, this procedure ended with

815 final blocks with 609 blocks considered as reliable (those having >20 rays hits, however most of block are

hit by >1000 rays); 48,856 and 40,010 P and S arrivals, respectively were used for this inversion. The resulting

velocity models were used to relocate the hypocenters, which were classified and filtered. Filtered hypocenters

were used for a new inversion. This procedure was repeated iteratively until the changes in velocities became

very small (2%), being five iterations required.

The August 28, 2004 (Mw = 6.7) earthquake occurred near of the river Teno, close to the Planchón volcano.

The aftershocks are distributed along a trace of NNE-SSW direction with depths lower than 15 km, which is

consistent with one of the focal mechanism solutions given by the Harvard CMT (Figure 1). This mechanism has

a solution oriented NNE-SSW dextral strike-slip. Likewise, the hypocentral distribution suggests that the rupture

occurred along one branch of the El Fierro fault system, located westward of the main fault of this system.

According with Farías (2007) the current state of the Main Cordillera would have a kinematic predominantly

dextral course with a forearc toward advancing to the north, although there determinations of focal mechanisms

showing other solutions (both normal and reverse) for earthquakes much lower magnitude (Barrientos et al.,

2004; Pardo et al., 2002). Considering that strike-slip earthquakes have magnitudes Mw between 5.2 and 6.5,

the energy related with them should exceed by several orders of magnitude the energy accumulated by the

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microseismicity that could have other focal mechanism solutions. Therefore, it seems to be valid to say that the

kinematics in the Principal Cordillera responds predominantly to transform movements, while other mechanisms

are likely to represent post- and inter-seismic accommodations. Indeed, this agrees well with geological

observations. According to Giambiagi and Ramos (2002) and Giambiagi et al. (2003), shortening in the chain

migrated westward since approximately four million years ago. Moreover, according with morphological

evidence, the range ceased to rise at about the same time after a fast and massive uprising event, which raised the

Cordillera over 2 km in 2 to 4 million years (Farías et al., 2008).

Acknowledgements

This work was funded by FONDECYT grant Nº 1030965 and Nº1070279, Bicentennial Program in Science and Technology grant ANILLO ACT Nº 18. The authors particularly recognize the labor made by the Seismologic Service at the University of Chile. We acknowledge Steven Roecker for providing the SPHREL90/SPHYPIT programs. References Barrientos, S., Vera, E., Alvarado, P., Monfret, T., 2004. Cristal seismicity in central Chile. J. South Amer. Earth

Sci. 16, 759-768. Charrier, R., Baeza, O., Elgueta, S., Flynn, J.J., Gans, P., Kay, S.M., Muñoz, N., Wyss, A.R., Zurita, E., 2002.

Evidence for Cenozoic extensional basin development and tectonic inversion south of the flat-slab segment, southern Central Andes, Chile (33º-36ºS.L.). J. S. Am. Earth Sci. 15, 117-139, doi:10.1016/S0895-9811(02)00009-3.

Charrier, R., Bustamante, M., Comte, D., Elgueta, S., Flynn, J.J., Iturra, N., Muñoz, N., Pardo, M., Thiele, R., Wyss, A.R., 2005. The Abanico Extensional Basin: Regional extension, chronology of tectonic inversion, and relation to shallow seismic activity and Andean uplift. Neues Jahrb. Geol. P-A. 236, 43-47.

Farías, M., Charrier, R., Comte, D., Martinod, J., Hérail G., 2005, Late Cenozoic deformation and uplift of the western flank of the Altiplano: Evidence from the depositional- tectonic- and geomorphologic evolution and shallow seismic activity (northern Chile At 19° 30’ S), Tectonics, 24, 0278-7407

Farías M., Comte, D., Charrier, R., Martinod, J.,Tassara, A., Fock, A., Crustal-scale structural architecture of the Central Chile Andes based on 3D seismic tomography, seismicity, and surface geology: Implications for mountain building in subduction zones, submitted to Earth and Planetary Science Letters, 2008

Giambiagi, L.B., Ramos, V.A., Godoy, E., Alvarez, P.P., Orts, S., 2003. Cenozoic deformation and tectonic style of the Andes, between 33° and 34° south latitude. Tectonics, 22, 1041, doi:10.1029/2001TC001354.

Giambiagi, L. B., Ramos, V. A., 2002. Structural evolution of the Andes between 33°30 and 33°45 S, above the transition zone between the flat and normal subduction segment, Argentina and Chile. J. S. Am. Earth Sci. 15, 99–114, doi:10.1016/S0895-9811(02)00008-1.

Giambiagi, L.B., Ramos, V.A., Godoy, E., Alvarez, P.P., Orts, S., 2003. Cenozoic deformation and tectonic style of the Andes, between 33° and 34° south latitude. Tectonics 22, 1041, doi:10.1029/2001TC001354.

Klein, F.W., 1978. Hypocenter location program HYPOINVERSE. U.S. Geol. Surv., Open-File Rep. 78-694. Pardo, M., Comte, D., Monfret, T., 2002. Seismotectonic and stress distribution in the central Chile subduction

zone. J. S. Am. Earth Sci. 15, 11-22, doi:10.1016/S0895-9811(02)00003-2. Roecker, S.W., Sabitova, T.M., Vinnik, L.P., Burmakov, Y.A., Golvanov, M.I., Mamatkanova, R., Munirova, L.,

1993. Three-dimensional elastic wave velocity structure of the western and central Tien Shan. J. Geophys. Res. 98, 15779-15795.

Thierer, P.O., Flüh, E.R., Kopp, H., Tilmann, F., Comte, D., Contreras, S., 2005. Local earthquake monitoring offshore Valparaiso, Chile. Neues Jahrb. Geol. P-A., 236, 173-183.

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Seismic structure of the continental margin offshore the southern Arauco Peninsula, Chile, at ~38°S

Eduardo Contreras-Reyes1, Ingo Grevemeyer

1, Ernst R. Flueh

1, Christian Reichert

2, & Martin

Scherwath1

1 Leibniz-Institute of Marine Sciences (IFM-GEOMAR), Wisschofstr.1-3, Kiel, Germany

(econtreras-reyes@ifm-geomar.de, igrevemeyer@ifm-geomar.de, eflueh@ifm-geomar.de,

mscherwath@ifm-geomar.de) 2 Federal Institute for Geosciences and Natural Resources, BGR, Stilleweg 2, Hannover, Germany

(Christian.Reichert@bgr.de)

KEYWORDS : subduction, accretionary wedge, continental slope, backstop region and seismic tomography

Introduction

The formation of accretionary wedges is mainly controlled by two factors: (1) slow convergence rate and (ii)

thick trench fill sediments (von Huene and Scoll, 1991; Clift and Vannuchi, 2004). Convergent margins such as

Cascadia (Hyndman et al., 1990), Nankai (Moore et al., 1990), and Makran (Kopp et al., 2000) are characterized

by a convergence rate slower than 6.5 cm/yr and a sequence of trench-fill sediments thicker than 1 km (Clift and

Vannuchi, 2004, and references therein). All these convergent margins carry large accretionary wedges >50 km

wide, and are usually classified as typical accretionary margins (e.g., von Huene and Scoll, 1991). The southern

central Chile margin (34°-45.5°S) is characterized by a filled trench confined between two main oceanic

features: the Juan Fernandez Ridge and the Chile Rise (Figure 1). Here, large volumes of terrigeneous sediments

sourced from the Andes have been transported via diverse canyon systems deposited in the trench during

Cenozoic times (e.g., Thornburg et al., 1990). Currently, the oceanic Nazca plate approaches the continent with a

covergence rate of 6.6 cm/a (Angermann et al., 1999), carrying a thin blanket of pelagic/hemipleagic sediments

(0-400 m) to the plate boundary (e.g., Contreras-Reyes et al, 2007). The total thickness of the sedimentary

sequence (pelagic and terrigeneous sediments) at the trench axis ranges between 1.5 and 2.5 km between 34° and

45°S (e.g., Grevemeyer et al 2003). In this manner, the southern central Chilean margin displays the typical

features characteristic of an accretion-dominated subduction zone. Nevertheless, marine seismic images (Bangs

and Cande; 1997) have shown that this margin exhibits only a small accretionary prism <30 km wide, which

abuts the truncated paleo accretionary complex that extends seaward from beneath the shelf. The small amount

of sediments accumulated here is not compatible with a continuous history of accretion, which implies episodic

history of accretion, nonaccretion and erosion of the southern central Chilean margin (Bangs and Cande; 1997).

Furthermore, Melnick and Echtler (2006) argued that the Glacial age trench fill and the steady decrease in plate

convergence rate had shifted the margin from erosive to accretionary during the Pliocene. The small size of the

accretionary wedge, the subduction rate history, and the deformational style along the margin suggest that the

current rapid rate of accretion cannot have lasted more than 1-2 Ma (Bangs and Cande, 1997).

The main aim of this study is to investigate the seismic structure of this young and active accretionary wedge

offshore southern Arauco Peninsula, in particular the seismic character of the transition zone between the

accretionary wedge and paleo accretionary complex. The paleo accretionary complex is made of sequences of

deformed sedimentary rock much older than, and not tectonically part of a presently growing accretionary mass

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(von Huene and Scoll, 1991). The sediment wedge grows up either by frontal accretion or underplating against

the seawardmost part of the margin's mechanical backstop (von Huene and Scoll, 1991). Seismically, this should

correpond to a rapid landward increase of seismic velocity, which might represent the backstop, a region within

the forearc that is much stronger than the region trenchward of it and thus it is able to support larger deviatoric

stresses (von Huene and Scholl, 1991; Kopp and Kukowski, 2003).

Figure 1. Geodynamic setting of Nazca, Antarctic, and South America plates; plates join at the Chile Triple Junction (CTJ), where the Chile Rise is currently subducting at ~46.4°S. The southern central Chilean margin is heavily sedimented and lies between the Juan Fernandez Ridge (JFR) and Chile Rise spreading center. Square denotes the study area offshore Arauco peninsula.

Results

In order to better understand the processes of accretion off south central Chile, a joint interpretation of swath

bathymetric, seismic refraction, wide-angle reflection and multi-channel seismic data was used to derive a

detailed tomographic image of the margin wedge offshore southern Arauco Peninsula, Chile at ~38°S. The data

were acquired during RV Sonne cruise SO161 of SPOC (Subduction Process off Chile) and SO181 of TIPTEQ

(from the Incoming Plate to mega-Thrust EarthQuake processes) projects (Krawczyk and SPOC Team, 2003;

Scherwath et al, 2006). The derived tomographic model (Figure 2) reveals two prominent velocity transition

zones characterized by steep lateral velocity gradients, resulting in a seismic segmentation of the marine forearc.

The margin is composed of three main domains; (1) a ~20 km wide frontal prism below the continental slope

with Vp 3.5 km/s, (2) a ~50 km area with Vp= 4.5-5.5 km/s, interpreted as a paleo accretionary complex, and

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(3) the seaward edge of the Paleozoic continental framework with Vp 6.0 km/s (Contreras-Reyes et al., 2008).

Frontal prism velocities are noticeably lower than those found in the northern erosional Chile margin (Sallares

and Ranero, 2005), confirming recent accretionary processes in south central Chile.

Figure 2. (a) Detailed tomographic image of the marine forearc complex, showing the seismic segmentation of the accretionary margin wedge. (b) Extracted velocities along the thick black line shown in (a), which landward of the lowermost slope corresponds to the uppermost basement velocities below the slope and shelf sediments. Note the strong horizontal velocity gradients at the slope break and seaward edge of the Paleozoic continental framework, which may suggest a change in rock type. References Angermann, D., Klotz, J., & Reigber, C. 1999. Space-geodetic estimation of the Nazca-South America Euler vector, Earth

Planet. Sci. Lett. 171, 3, 329-334. Bangs, N.L., & Cande, S.-C. 1997. Episodic development of a convergent margin inferred from structures and processes

along the southern Chile margin, Tectonics 16(3), 489–503. Clift, P., & Vannucchi, P. 2004. Controls on tectonic accretion versus erosion in subduction zones: Implications for the origin

and recycling of the continental crust. Rev. Geophys 42: RG2001, doi:10.1029/2003RG000127. Contreras-Reyes, E., Grevemeyer, I., Flueh, E.-R., Scherwath, M., & Heesemann, M. 2007. Alteration of the subducting

oceanic lithosphere at the southern central Chile trench–outer rise, Geochem. Geophys. Geosyst. 8, Q07003, doi:10.1029/2007GC001632.

Contreras-Reyes, E., Grevemeyer, I., Flueh, E.-R., & Heesemann, C. 2008. Upper lithospheric structure of the subduction

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zone of southern Arauco Peninsula, Chile at ~38°S, J. Geophys. Res. (in press). Grevemeyer, I., Diaz-Naveas, J.-L., Ranero, C.-R., Villinger, H., & Ocean Drilling Program Leg 202 Scientific Party. 2003.

Heat flow over the descending Nazca plate in Central Chile, 32°S to 41°S: evidence from ODP Leg 202 and the occurrence of natural gas hydrates, Earth Planet. Sci. Lett. 213, 285–298.

Hyndman, R. D., Yorath, C. J., Clowes, R.-M., & Davis, E.E. 1990, The northern Cascadia subduction zone at Vancouver Island: Seismic structure and tectonic history, Can. J. Earth Sci. 27(3), 313–329.

Kopp, C., Fruehn, J., Flueh, E.-R., Reichert, C., Kukowski, N., Bialas, J., & Klaeschen, D. 2000. Structure of the Makran subduction zone from wide-angle and reflection seismic data, Tectonophysics 329, 171-191.

Kopp, H., & Kukowski, N. 2003. Backstop geometry and accretionary mechanics of the Sunda margin, Tectonics 22(6), 1072, doi:10.1029/2002TC001420.

Krawczyk C.-M, SPOC Team .2003. Amphibious seismic survey images plate interface at 1960 Chile earthquake. Eos Trans. AGU 84(32):301, 304-305.

Melnick, D., & Echtler, H. 2006. Inversion of forearc basins in south-central Chile caused by rapid glacial age trench fill, Geology 34 (9), 709–712.

Moore, G..-F., Schipley, T.-H., Stoffa, P.-L., & Karig, D.-E.1990. Structure of the Nankai Through Accretionary Zone from Multichannel Seismic Reflection Data, J. Geophys Res. 95(B6), 8753-8765.

Sallares, V., & Ranero, C.-R. 2005. Structure and tectonics of the erosional convergent margin off Antofagasta, north Chile (23°30'S), J. Geophys. Res. 110, B06101, doi:10.1029/2004JB003418.

Scherwath, M., Flueh, E.-R., Grevemeyer, I., Tilmann, F., Contreras-Reyes, E., & Weinrebe, W. 2006. Investigating Subduction Zone Processes in Chile, Eos Trans. AGU 87(27), 265.

Thornburg, T.-M., Kulm, D.-M., & Hussong, D.-M .1990. Submarine-fan development in the southern Chile trench: a dynamic interplay of tectonics and sedimentation, Geol. Soc. Am. Bull 102, 1658-1680.

von Huene, R., & Scholl, D.-W. 1991. Observations at convergent margins concerning sediment subduction, subduction erosion, and the growth of continental crust, Rev. Geophys 29(3), 279-316, doi:10.1029/90JB00230, 1990.

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Fractures in the Mejillones Peninsula triggered by the Tocopilla Mw=7.7 earthquake

J. Cortés1, D. Rémy

2, G. González

3, J. Martinod

2, & G. Gabalda

2

1 Programa de Doctorado en Ciencias Mención Geología, Universidad Católica del Norte, Avenida Angamos

0610, Antofagasta, Chile (jcortesa@ucn.cl) 2

Laboratoire des Mécanismes et Transferts en Géologie (LMTG), 14, Avenue Edouard Belin, 31400 Toulouse,

France (remy@lmtg.obs-mip.fr, martinod@lmtg.obs-mip.fr, gabalda@ird.fr) 3 Departamento de Ciencias Geológicas, Universidad Católica del Norte, Avenida Angamos 0610, Antofagasta,

Chile (ggonzale@ucn.cl)

KEYWORDS : Tocopilla earthquake, fracture development, intraplate deformation, seismic cycle, Mejillones Peninsula

Introduction

The coastal margin of northern Chile and Southern Peru is characterized by the convergence between Nazca

and South American plates, at a velocity of 6.5 cm/yr (Angermann et al., 1999). This process is responsible for

large subduction earthquakes that historically have affected this margin, such as the 1877 Iquique Mw=8.8, the

1995 Antofagasta Mw=8.1, the 2001 Arequipa Mw=8.2-8.4 and recently the 2007 Tocopilla Mw=7.7

Earthquake. A key problem in northern Chilean forearc is to establish the relationship between these large

subduction earthquakes and the intraplate deformation observed at the topographic surface. Several authors have

postulated that subduction earthquakes trigger normal faulting in different parts of the forearc (e.g. Delouis et

al., 1998; González and Carrizo, 2003). As a matter of fact, seismological data did not report any active

superficial deformation in the early nineties in the Mejillones area, the period during which the subduction zone

of that region was in pre-seismic state (Delouis et al., 1996). On the other hand, fracture formation has been

postulated as secondary superficial process related to coseismic extension produced by this type of earthquake.

By using high resolution satellite images, Loveless et al. (2005) and González et al. (In Press) mapped mesh of

fractures in different parts of the Coastal Cordillera. These authors distinguished fractures orientated both

parallel (Loveless et al. 2005) and normal to the trench (González et al. In Press). By using numerical models,

Loveless (2008) postulated that fracture parallel to the trench can be formed by successive subduction

earthquakes, whereas Gonzalez et al. (In Press), based on numerical modeling, demonstrated that fractures

orientated normal to the trench would be produced by intraplate faulting.

On November 14th 2007 a Mw=7.7 earthquake occurred along the Coastal Region of Northern Chile. The

seismic rupture propagated over 200 km from north to the south ending below the Mejillones Peninsula. In this

place, we documented the formation of open fractures a few days later the earthquake. In order to understand the

origin of these fractures we performed numerical modeling (Elastic Dislocation Model) and InSAR analyses of

the coseismic displacements.

Long term deformation of the Mejillones Peninsula

The Mejillones Peninsula is characterized by the occurrence of normal faults which affect Miocene to

Pleistocene marine sediments (Armijo and Thiele, 1990; Niemeyer et al. 1996; Marquardt 2005). Normal faults

are spectacularly expressed by prominent fault scarps that control the morphology of this peninsula. Fractures

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are a very common structural feature which can be observed in different parts of the Mejillones Peninsula. For

example, fractures can be observed affecting Pliocene-Pleistocene marine sediments along the coastal cliff of the

Mejillones Bay. In the same way alluvial sediments near the Mejillones Fault are also profusely affected by

fractures. These fractures show a characteristic infill given by detritical material with internal layering parallel to

their borders. Fault slip data and the orientation of the fractures indicate that the Mejillones Peninsula in the long

term scale is under an E-W extension.

Marine terraces of late Pleistocene are notably preserved in the Mejillones Peninsula (Figure 1). The

occurrence of these terraces indicates that the Mejillones Peninsula has experienced a continuous uplift at least

during the last 400 ky (Marquardt, 2005). The uplift rate varies with the location of the marine terraces relative

to the normal faults. For example, the terraces located in the footwall of the Mejillones Fault have higher uplift

rates than those located in the hangingwall. In fact, 0.5-0.7 mm/y has been determined in the footwall and 0.2-

0.5 mm/y for the hangingwall. The difference in the uplift rates of these two blocks indicates that the Mejillones

Fault has a slip rate of 0.2-0.3 mm/y.

Figure 1. Shaded relief 1:400000 of the Mejillones Peninsula. Red lines represent Caleta Herradura Fault (CHF), Mejillones Fault (MF) and other extensional structures identified in the area. Black lines correspond to paleocoastal lines preserved mainly in the Pampa Mejillones (PM). MM is referred to Morro Mejillones, the footwall of Mejillones Fault. Yellow circles are the sites where fractures were observed (A-G). To the right, rose diagrams showing the fracture orientation in each point. The coloured zones in the MM correspond to the main marine terraces identified by Marquardt (2005).

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The Tocopilla Mw=7.7 Earthquake

According to the USGS, the Tocopilla Earthquake occurred at 40 km depth, and the epicenter was located at

22°19’S-69°84’W, 40 km east-southeast of Tocopilla. The rupture plane has an orientation N5W and the dip is

of 16NE. The surface rupture is 180 km along strike, and 120 km downdip. The aftershocks registered occurred

mainly in the northern part of the Mejillones Peninsula.

Fracture documentation

A few days later the Tocopilla Earthquake, we visited the epicenter area defined by USGS and the Mejillones

Peninsula. In the epicenter area we found several open fractures affecting alluvial deposits. Because in this area

we did not find recent anthropogenic features cut by the fractures, we could not estimate whether if the fractures

were produced during this event. On the contrary, in the Mejillones Peninsula we found fresh fractures affecting

alluvial fan deposits. Because we are currently working on the Mejillones Peninsula, our own vehicle tracks of

previous field campaigns were profusely cut by fractures. This observation clearly shows that these fractures

formed during the Tocopilla Earthquake. In general, the fractures are disposed close to the main traces of Caleta

Herradura and Mejillones Faults. In the case of Caleta Herradura Fault, we found a significant number of cracks

affecting alluvial fans disposed east of the main scarp of this fault (Figure 1).The length of the fractures varies

between 3-46 meters, with apertures of 5-10 mm. The orientation of these fractures is closely parallel to the

strike of the main scarp of Caleta Herradura Fault. Local variation in fracture orientation is controlled by the

local strike of Caleta Herradura Fault. In this area, some fractures constitute reopened cracks; it is inferred by

occurrence of an open central portion affecting the infill of pre-existing fractures. Close to the Mejillones Fault

we identified similar fractures affecting alluvial and eolian deposits. These fractures are less abundant than those

close to the Caleta Herradura Fault. The length of these fractures varies between 1.3 to 26 m and the aperture is

close to 0.5-10 mm. In this case, fractures are nearly parallel to the main scarp of Mejillones Fault (site A Figure

1) and parallel to a secondary fault eastward of the Mejillones Fault (sites B and C, Figure 1). The strain related

to the fracture aperture is extremely low (~0.01%), indicating that extension is diffuse at the kilometric scale.

InSAR and Coulomb Stress Change Analysis

We used ASAR radar images acquired between August 2007 and December 2007, from the satellite

descending track numbers 368 and 96, covering the eastern and western part of the study area, respectively.

Calculated interferograms show long-wavelength deformation related to the main shock, with a maximum 19 cm

of range increase (i.e. ~ subsidence) located in the eastern part of the study area South Maria Elena, and up to 26

cm of range decrease (i.e.~ uplift) located in the north of the Mejillones Peninsula. Interferograms also show a

small wavelength signal located SW Mejillones that cannot be explained by slip occurring on the main rupture

plane. This small-wavelength anomaly is situated on a narrow (<10 km-wide), N-S oriented area located on the

hangingwall of to the Mejillones Fault. It is precisely in this zone, eastward of the main scarp of the Mejillones

Fault, where we observe fractures (sites A to C in figure 1).

We use the radar data to calculate the slip that occurred on the main fault during the Tocopilla Mw=7.7

Earthquake. We check that the long-wavelength ground deformation results from thrust motion on a N-S deep

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fault. The calculated seismic moment is 5.8 x 10 20 N.m equivalent to Mw = 7.7 (assuming μ=3.1 x 1010 Pa),

slightly higher than the one deduced by Harvard CMT (4.9 x 10 20 N.m). We calculate the Coulomb Stress

Change (CSC) induced at the surface by the main shock, for normal fault planes striking NS and dipping 55°

towards the east, i.e. similar to the major normal faults of the Mejillones Peninsula. In Caleta Herradura and

Mejillones Faults, positive values of CSC were found, indicating potentially for triggering fault reactivation. In

these sites where the presence of cracks was observed, CSC values vary between +1.5 and +2 bars.

Conclusions

The formation of fractures in the Mejillones Peninsula during or shortly after the Mw= 7.7 Tocopilla

Earthquake shows that superficial extension is partly related to subduction earthquakes. In particular, the

documented fractures are disposed directly above the southern termination of the seismogenic rupture related to

the Tocopilla Earthquake. The Mejillones Peninsula was the focus of several aftershocks which indicate that

stress release was concentrated in this southern termination. Our fracture documentation shows that the

Mejillones Peninsula was under diffuse superficial extension during the Mw= 7.7 Tocopilla Earthquake. CSC

analysis indicates that the Mw= 7.7 Tocopilla Earthquake resulted in tensional stress field on N-S faults. The

short term deformation following the Tocopilla Earthquake correlates with the long term extensional

deformation observed in the Mejillones Peninsula. It suggests that the long term deformation of the Mejillones

Peninsula is related to subduction earthquakes.

References Angermann, D.; Klotz, J. & Reigber, C. 1999. Space-geodetic estimation of the Nazca-South America Euler vector. Earth

Planet Sci Lett, 171: 329-334. Armijo, R. & Thiele, R. 1990. Active faulting in northern Chile: ramp stacking and lateral decoupling along a subduction

plate boundary? Earth Planet Sci Lett, 98: 40-61. Delouis, B., Cisternas, A., Dorbath, L., Rivera, L. & Kausel, E. 1996. The Andean subduction zone between 22 and 25°S

(northern Chile) : precise geometry and state of stress. Tectonophysics, 259, 81-100. Delouis, B., H. Philip, L. Dorbath & Cisternas, A. 1998. Recent crustal deformation in the Antofagasta region (northern

Chile) and the subduction process. Geophys. J. Int., 132: 302 – 338. González, G., Gerbault, M., Martinod, J., Cembrano, J., Carrizo, D., Allmendinger, R. & Espina, J. In press. Crack formation

on top of propagating reverse faults of the Chuculay Fault System northern Chile: Insights from field data and numerical modelling. Journal of Structural Geology.

González, G. & Carrizo, D. 2003. Segmentación, cinemática y cronología relativa de la deformación tardía de la Falla Salar del Carmen, Sistema de Fallas de Atacama, Cordillera de la Costa de Antofagasta. Revista Geológica de Chile, 30 (2): 223-244.

Loveless, J. 2008. Extensional tectonics in a convergent margin setting: Deformation of the northern Chilean forearc. Ph.D Thesis, Cornell University, 311 p.

Loveless, J., Hoke, G., Allmendinger, R., González, G., Isacks, B. & Carrizo, D. 2005. Pervasive cracking of the northern Chilean Coastal Cordillera: New evidence for forearc extension. Geology, 33: 973-976.

Marquardt, C. 2005. Déformations Néogènes le long de la cotê nord du Chili (23°-27°S), avant-arc des Andes Centrales. Thèse doct., univ. Toulouse-III, 212 p.

Niemeyer, H., González, G. & Martinez-de Los Rios E. 1996. Evolución tectónica cenozoica del margen continental activo de Antofagasta, norte de Chile. Revista Geológica de Chile, 23 (2): 165–186.

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Analyse of the Tarapaca paleolandslide (North Chile) using generalized Newmark approach and implications on paleoseismicity, and on paleoclimate changes

J. Darrozes1, J.-C. Soula

1, J. Ingles

2, R. Rilqueme

3 , & G. Herail

4

1 LMTG, OMP, CNRS, University Paul Sabatier, 14 avenue Edouard Belin, 31400 Toulouse, France

(corresponding author: darrozes@lmtg.ups-tlse.fr) 2 Department of Civil Engineering, University Paul Sabatier, 118 route de Narbonne, 31400 Toulouse, France

3 Departamento de Ciencias Geológicas, Universidad Católica del Norte, Avenida Angamos 0610, Antofagasta,

Chile 4 IRD UR 154 - LMTG, Lima, Peru

KEYWORDS : paleolandslide, sismicity, thrust fault-propagation fold, generalized Newmark analysis, Tarapaca

(northern Chile)

Introduction

Landslides much larger than today’s are not uncommon in the geological record, and are observed in arid or

hyperarid areas where a good preservation allow detailed analyses of their geometrical and mechanical

characteristics. Should such large-scale landslides be possible today, much larger devastations than caused by

the present-day’s greatest registered events could be expected.

Figure 1. Tectonic sketch map modified after Hérail (1996), red box identified the studied Tarapaca landslide.

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One way of checking this possibility is to determine the magnitude and site-to-source distance of the seismic

event necessary for the landslide to form using Newmark permanent displacement analysis (see e.g. Jibson &

Keefer, 1993). One of the limitations of the conventional Newmark’s method is to consider ground acceleration

as parallel to the slope, and downslope.

Figure 2. Morphological and structural features of the Tarapaca landslide : A) Aster VNIR image (res.: 15m); black dotted line = location of the B) schematic profile of the landslide note the free edge favouring the landslide and installation of an immature drainage network on the landslide foot; white doted line secondary landslide located at the foot of the main landslide.

A generalized Newmark analysis (Ingles et al., 2006) has been developed where ground acceleration is not

slope-parallel and the ratio of vertical to horizontal acceleration depends on the seismic situation of the slope

(magnitude, earthquake source distance, style of faulting). As a consequence, the seismic horizontal critical

acceleration is in most cases lower than Newmark’s critical acceleration and the displacement greater than

calculated from Newmark analysis. The difference may be considerable and result in significant consequences,

particularly for low-angle slopes with potential deep shear surfaces that are located close to the source of large

earthquakes. Therefore, Arias intensity and the magnitude of earthquakes necessary for large-scale landslides to

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occur may be lower than expected from the traditional Newmark approach. This approach has been applied to a

large-scale paleo-landslide, the Tarapaca landslide, observed in the Atacama Desert, northern Chile.

Morphological, geological and geotechnical characteristics of Tarapaca paleo-

landslide

The Tarapaca landslide (Fig. 2) is one of a series of large-scale landslides formed on the forelimb of a ~400 km

long west-vergent thrust fault-propagation fold known as the Moquella flexure (Darrozes et al., 2002). The

landslides postdate the initiation of a well-preserved drainage network dated of 4.5 Ma (Naranjo et Paskoff,

1985). Its dimensions are ~9 km wide by ~4 km wide, with a total volume close to 7 km3. The triggering

mechanism of the Tarapaca paleo-landslide was studied by constraining characteristics of rocks, slope value,

climatic changes and paleoseismic activity. The clay cohesion, which constitute the slip surface, varying from 90

kPa to 30 kPa under unsaturated to saturated conditions (Pinto et al., 2007). The friction angle is very stable and

close to 40° and the rock mass unit weight varying from 20 for strongly weathered rocks to 25 kN/m3 not

weathered rocks The static safety factor (eq. 1) show all the classical parameters which trigger landslide like

load, weathering, slope and ground water pressure are not sufficient to initiated the paleo-landslide.

eq. 1

The aseismic safety factor predicting slope stability even for the most severe possible conditions, earthquake

shaking is now to be considered. The characteristics of the Tarapaca landslide: large volume, low-angle and

deep-seated basal shear surface, materials having a relatively high strength strongly suggest that the landslide

was triggered by strong and probably long duration earthquake shaking (high magnitude and large Arias

intensity). Arias intensities and moment magnitudes have been calculated for the Tarapaca landslides using

Inglès et al.'s (2006) model, the characteristics of the materials being those defined in the calculations of the

aseismic safety factor. The critical displacement was assumed to be 10 cm. These calculations indicate Arias

intensity of 16.84 m/s and moment magnitudes of Mw= 6.98 at 15 km to the landslide, Mw= 8.26 at 50 km to

the landslide and Mw= 9.2 at 120 km to the landslide. As a comparison, the conventional Newmark analysis

(1965) would have given moment magnitudes of Mw= 7.65; Mw= 9.23 (exceptional) and Mw= 10.38

(unrealistic) and unrealistically high Arias intensities of 32.49 m/s (Fig. 3).

Discussion and conclusions

If ones looks only the Tarapaca landslide two scenarii are possible either a seism of Mw~7 of subsurface (at a

distance close to 15km from the landslide) or a very strong one, Mw upper than 9, within the subduction plane

(distance close to 110 km). If ones look, now, the whole area we can observe the presence of many landslides of

comparable size. We can note that they take again the same characteristics as Tarapaca i.e. the landslides are

located on the forelimb of Moquella flexure, the slided mass is formed by the same ignimbritic layer and the

slope is lesser than 10°. But for these two assumptions it is necessary to add a climatic aspect which allows

existence of an important underground water table (m>0.5). These important water tables may result from the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 172-175

175

vicinity of rivers. The two limiting rivers which have, in the past, a channel close to the rupture plane can

effectively increase drastically the water table and allowed the initiation of the slip.

Figure 3. Variation of moment magnitude Mw versus earthquake source-distance determined : the critical displacement DN is considered as 10 cm. For Ingles et al. model (2006) (solid line) we used a ratio of vertical to horizontal seismic ground acceleration k1 = 1 and a resulting Arias Intensity of 16.84 m/s. For Newmark model (1965) (dashed line) we used k1 =

-tan( ) and a resulting Arias intensity of 32.49 m/s.

References Darrozes, J., Pinto, L., Inglès, J., Soula, J.C., Maire, E., Courjault-Radé, P., Hérail, G., 2002 - Origin of the paleolandslide of

Tarapaca (North Chile, Andean belt)- Geophysical Research Abstract, EGS02-A-03 136. Inglès, J., Darrozes, J., Soula, J.-C., 2006 - Effects of vertical component of ground shaking on earthquake-induced

landslides displacements using generalized Newmark's analysis - Engineering Geology 8,134-14. Jibson, R.W., Keefer, D.K., 1993 -Analysis of the seismic origin of landslides: Examples of the New Madrid seismic zone -

Geol. Soc. Am. Bull. 105, 421-436. Naranjo, J.A., Paskoff, I., 1985 - Evolution Cenozoica del piemonte andino en la Pampa del Tamagural, norte de Chile (18°-

21° S).- IV Congreso Geologico Chileno, 4, 149-165, 1985. Newmark, N.M., 1965 -“Effects of earthquakes on dams and embankment s- Geotechnique 15, 139-159. Pinto L, Herail G, Sepulveda SA, Krop P, Darrozes J, 2007 -The Lataguella megalandslide, Tarapaca region, Northern Chile

an example of Cenozoic instability of Andean arc the Bolivian orocline -, AVH2-A-00205, 2007. Pinto, L, Herail, G, Rinaldo, C, 2004 - Sedimentacion sintectonica asoiciada a las estructuras neogenas en la Precordillera de

la zona de Moquella, Tarapaca, Rev. Geol. Chile, 31,1, 19-44.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 176-179

176

Dynamic topography during flat-slab subduction: A first approach in the south-Central Andes

Federico M. Dávila1 & Carolina Lithgow-Bertelloni

2

1 Laboratorio de Análisis de Cuencas, CICTERRA-Universidad Nacional de Córdoba, Av. Vélez Sársfield 1611,

2º piso, of.7, X5016GCA Córdoba, Argentina (fmdavila@efn.uncor.edu) 2 Department of Earth Sciences, University College London, Gower St., London WC1E 6BT, United Kingdom

(c.lithgow-bertelloni@ucl.ac.uk)

KEYWORDS : dynamic topography, flat subduction, basin analysis, Sierras Pampeanas, Pampean plain

Introduction

In the flat-slab segment of the south-Central Andes, recent stratigraphic reconstructions have proposed that

long-wavelength and high-amplitude accommodation spaces have controlled the alluvial sedimentation in the

foreland during Miocene to Present (Dávila et al., 2005, 2007). Given that the tectonic and sedimentary loads do

not yield the large magnitude of these spaces by conventional flexural models (Dávila et al., 2005), alternative

explanations are needed. Eastward of the Andean foreland system, within the pericratonic Pampean Plains (to

>700 km from the Chilean trench, where the slab descends again at high angles; Booker et al., 2005), Cenozoic

subsurface sequences show thicknesses of >0.5 km. But this region is even more distant from the High Andes

loads. The closest topography is the modest Sierras de Cordoba range (easternmost range of the Sierras

Pampeanas), which only records a shortening <5%, i.e. ~5 km within the ~100-km orogen wide. This undersized

tectonic load can explain <<50% of the accommodation spaces in the Plains (see Dávila, 2008).

Dávila et al. (2005, 2007) suggested “hidden sub-lithospheric loads” (like dynamic topography, lithospheric

mantle densification or eclogitization of the lower crust) might overlap the tectonic and sedimentary loads and

also explain the recorded load deficits. These mechanisms, likewise, would provide reconciliation with other

geophysical and geomorphological features of the foreland (see Dávila et al., 2005). But a question arises: which

of these controls occur in the Andes and what triggered it? With the onset of subduction, Earth’s surface deforms

by vertical stresses induced by mantle flow. This deformation is called “dynamic topography” because the

buoyancy forces driving the surface deflections are actively moving. Thus, whenever subduction occurs,

dynamic topography will be present in a foreland. However, at >700 km from the trench, a second question

arises: what is the magnitude and importance of the dynamic topography?

We test and quantify the degree to which dynamic topography explains how much of the total subsidence

(negative vertical deflection) in the south-Central Andean foreland is controlled by non-isostatic loads originated

in the astenospheric mantle during flat subduction. We specifically focus on dynamic topography during flat

subduction, as it seems clear that the synorogenic accommodations began in the Pampean regions in the last

10-7 my, coevally with the arrival of the Juan Fernandez Ridge to these latitudes, coincident with the initiation

of slab flattening and basement thrusting in the Sierras Pampeanas (or broken foreland). We compare predicted

values of dynamic topography with maximum deflections and maximum accommodation spaces estimated by

flexural analysis and stratigraphic approaches, respectively.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 176-179

177

Methodology

We calculated dynamic topography based on mantle flow models that assume a three-dimensional density

distribution inside the Earth and solve the viscous flow induced by such density heterogeneity. The model is

based on the history of 120 m.y. of subduction (Ricard et al., 1993) and assumes that cold subducted slabs are

the main source of thermal buoyancy in the mantle, and therefore of mantle density heterogeneity. This is a

reasonable approximation for a mantle largely heated internally, although it neglects the role of active

upwellings. The model calculates predicted geoids, which can be compared with observed geoids. Both geoid

and dynamic topography are sensitive to the mantle’s convective pattern and therefore to the mantle’s viscosity

structure. But, the model assumes subvertical subduction from the trenches. Therefore, we had to modify the

slab geometry in our region to simulate flat subduction from the trench to ~64° WL, where the slab submerges

again vertically (Booker et al. 2005).

Results

As expected, the wavelength of dynamic topography is long and smooth at low (10-30) spherical harmonic

degrees (Fig. 1a). However, to analyze a 1000 - 500 km length region, higher degrees (>30) were required.

Dynamic topography in these cases is more complex and tends to adjust to the geometry of the subducting slab

(Fig. 1b). According to previous calculations (see Lithgow-Bertelloni and Richards, 1998 and references

therein), a lithosphere that is 10 times more viscous than the upper mantle and a lower mantle that is 50 times

more viscous than the upper mantle can account for >80% of the present-day geoid. Testing different viscosity

contrasts (lithospheric mantle / upper mantle / transitional zone / lower mantle), good correlations between

predicted geoids and observed geoids (coefficient correlations of 0.65-0.90) were obtained. For a lithospheric

mantle more viscous than the lower mantle (1000 / 1 / 1 / 150) dynamic topography above the subducting cells is

-731.2 m. Keeping similar viscosity constrasts for the lithosphere and lower mantle (50 / 1 / 1 / 50), dynamic

topography above the subducting cells is -1135.9 m. When lower mantle is much more viscous than lithosphere

(1 / 1 / 1 / 50), dynamic topographic is -580.2 m. During flat subduction and with variations in viscosity in the

mantle wedge, dynamic topography increases and varies on much shorter wavelengths (Burgess et al., 1997,

Billen et al. 2001)

(a) (b)

-400

-300

-200

-100

0

100

200

300

400

500

600

-120 -100 -80 -60 -40 -20 0

-700

-600

-500

-400

-300

-200

-100

0

100

200

300

400

-120 -100 -80 -60 -40 -20 0

Figure 1. Profiles at 31° SL depicting the dynamic topography variation above the flat subduction of the south-Central Andes depicting results of experiments with: (a) low spherical harmonic degrees (10), using a lithosphere that is 10 times more viscous than the upper mantle and a lower mantle that is 50 times more viscous than the upper mantle; and (b) high spherical harmonic degrees (50), in the latter case using a viscosity contrast simulating a rigid lithosphere with a viscous lower mantle (50 times more). X-axis is longitude (coordinates) and Y-axis is meters respect to sea level.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 176-179

178

Discussion

The accommodation spaces, estimated by negative values of dynamic topography, are in the order of hundreds

of meters in amplitude (between 500-1100 m) and of thousands of kilometers in wavelength (between 700 km

for 1 / 1 / 1 / 50 and 3500 km for 1000 / 1 / 1 / 100). This indicates mantle-driven forces during subduction are

large enough to reproduce the subsidence in foreland and pericratonic areas even using higher spherical

harmonic degrees, and other loads would not be needed. However, comparing these results with flexural models

and stratigraphic studies (e.g., Dávila et al. 2005, 2007), nearly all predictions overestimate the magnitude and

wavelength of the subsidence. A rigid lithosphere and a viscous lower mantle (e.g., 1 / 1 / 1 / 50) reproduce

nearly proper wavelengths and amplitudes (~700 km and ~500 m, respectively). This result is consistent with

geological studies and with low viscosities in the astenospheric wedge during slab flattening (due to low

concentrations of water in the wedge, Billen and Gurnis, 2001). However, and contrary to previously proposed

(Dávila et al., 2005), the modification of the slab geometry, to simulate flat subduction, does not reproduce

negative dynamic topographies in the flattest part of the slab. Instead, it favors the generation of positive

“relieves”. Thus, when the viscous shear generated by the subducting slab is cancelled laterally, upwarping

surfaces develop close to the trench (Fig. 2). This is in apparent contradiction with previous results (e.g., Burgess

et al., 1997) that suggested subsidence even along the flattest part of the slab. However, taking into account the

dynamic subsidence is triggered by mantle flows, it is reasonable that a reduction in the astenospheric wedge, by

slab flattening, produces a decrease in the negative values of dynamic topography.

Figure 2. Dynamic topography in the flat slab segment of the Central Andes with a lithosphere 5 times more viscous than upper mantle and lower mantle 40 times more viscous than upper mantle (spherical harmonic degree 40, correl. coef. = 0.88). Note the flat slab (between 290° and 295° WL in the map) does not contribute on the subsidence and the sinking part of the subduction cratonward (at ~295° WL, darker blue) is the responsible of the major negative dynamic topography values.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 176-179

179

Assuming flat subduction is a result of ridge interaction in the Andes and the slab flattening is not static and

shifted southward during the Miocene (Yañez et al., 2001), subsidence driven by dynamic topography should

have also migrated in this direction tracking the leading edge of the Juan Fernandez Ridge. If flat subduction

favors surface uplift, the exhumation and denudation of the northern Sierras Pampeanas broken foreland may

have occurred when the slab flattening passed across this area in the middle-late Miocene. The location of the

flat slab at ~31° SL since the late Miocene-Pliocene (Kay and Mpodozis, 2002) would allow to suggest

exhumation/denudation of the western broken foreland and the subsidence cratonward in pericratonic areas, out

of the influences of tectonic topographies, are mainly a result of the sublithospheric processes.

Based on these new results, some aspects might be re-evaluated, like the origin of “subduction erosion” in flat-

slab segments and of the crustal overcompensations in the Sierras Pampeanas evidenced by seimic velocities

studies (Fromm et al., 2004).

References Billen, M.I. and Gurnis M., 2001. A low viscosity wedge in subduction zones. Earth and Planetary Science Letters 193(1):

227-236. Billen, M.I. and Gurnis M., Simons, M., 2003. Multiscale dynamics of the Tonga–Kermadec subduction zone. Geophys. J.

Int.: 153: 359–388 Booker, J., Favetto, A., Pomposiello, C. & Xuan, F., 2005. The role of fluids in the Nazca flat slab near 31ºS revealed by

electrical resistivity structure. 6º International Symposium on Andean Geodynamics, Extended Abstract: 119-122. Barcelona.

Burgess, P. M., Gurnis, M., & Moresi, L., 1997. Formation of sequences in the cratonic interior of North America by interaction between mantle, eustatic and stratigraphic processes, Bulletin of the Geological Society of America 108: 1515-1535.

Dávila, F.M., 2008. The Modern Pampean Plain foreland basin system at 31º SL: Depozones controlled by crystalline basement thrusting. Congreso Geológico Argentino, San Salvador de Jujuy, Argentina.

Dávila, F.M., Astini, R.A. & Jordan, T.E., 2005. Cargas subcorticales en el antepaís andino y la planicie pampeana: Evidencias estratigráficas, topográficas y geofísicas. Revista de la Asociación Geológica Argentina 60: 775-786.

Dávila, F. M., R.A. Astini, T E. Jordan, G. Gehrels, & M. Ezpeleta, 2007. Miocene forebulge development previous to the broken foreland partitioning in the southern Central Andes, west-central Argentina. Tectonics 26: TC5016, doi:10.1029/2007TC002118.

Fromm, R., Zandt, G. & Beck S.L., 2004. Crustal thickness beneath the Andes and Sierras Pampeanas at 30°S inferred from Pn apparent phase velocities. Geophysical Research Letters 31: L06625, doi: 10.1029/2003GL019231.

Kay, S.M. and Mpodozis, C., 2002. Magmatism as a probe to the Neogene shallowing of the Nazca plate beneath the modern Chilean flat-slab. Journal of South American Earth Sciences 15: 39-57.

Lithgow-Bertelloni, C., & Richards, M.A., 1998. The dynamic of Mesozoic and Cenozoic plate motion. Reviews of Geophysics 36: 27-78.

Ricard, Y., Richards, M. Lithgow-Bertelloni, C. & LeStunff, Y., 1993. A geodynamical model of mantle density heterogeneity. J. Geophys. Res. 98: 21,895–21,909.

Yañez, G.A., Ranero, C.R., von Huene, R., & Diaz, J., 2001. Magnetic anomaly interpretation across the southern central Andes (32°–34°S): the role of the Juan Fernández Ridge in the late Tertiary evolution of the margin. Journal of Geophysical Research 106: 6325–6345.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 180-183

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Dynamics of the November 3, 2002 eruption of El Reventador volcano, Ecuador: Insights from the morphology of ash particles

S. Delpit1, J.-L. Le Pennec

1, P. Samaniego

2, S. Hidalgo

2, & C. Robin

1,2

1 IRD, UMR 163, Laboratoire Magmas et Volcans, 5 rue Kessler 63038 Clermont-Ferrand, France

(s.delpit@opgc.univ-bpclermont.fr, lepennec@ird.fr) 2 Instituto Geofisico, Escuela Politécnica Nacional, Ap. 17-01-2759, Quito, Ecuador

KEYWORDS : Reventador, Ecuador, tephra, VEI, magmatic fragmentation

Introduction

The Ecuadorian Volcanic Arc occurs in the Northern

Volcanic Zone of the Andes and results from the

subduction of the Nazca plate, which supports the

Carnegie Ridge, beneath the South American plate

(fig.1). The andesitic El Reventador stratovolcano,

about 100 km east of Quito city, belongs to the Back

Arc lineament. It is composed of an old edifice that

suffered two collapse events, and a young frequently

active stratocone whose summit rises about 3500 m

above sea level (Aguilera et al., 1998). On November

3, 2002, after 26 years of quiescence and little

precursory warning, El Reventador volcano erupted suddenly and violently, becoming the most important

eruption in Ecuador since the 1886 eruption of Tungurahua (the volcanic explosivity index, VEI, ranks at 4). The

paroxysmal phase of the eruption, started at 9:12 AM (local time), was a short event (around 48 minutes) during

which a 13 km-high eruptive column rose above the crater. The mainly westward wind drifted the volcanic cloud

above populated areas of the Interandean Valley and deposited a fine grained tephra fall layer (Hall. et al., 2004).

Many pyroclastic flows were also emplaced during the event, which was followed by the emplacement of

andesitic lava flows (Hall. et al., 2004; Ridolfi et al., 2008; Samaniego et al., 2008). Based on preliminary data,

Le Pennec et al. (2003) estimated the uncompacted tephra fall layer volume to ~0.28 km3 while new estimates

obtained in this work using the method of Fierstein and Nathenson (1992) is about 0.15 km3. The eruption style

has been described as subplinian with a VEI of 4. As this eruption style remains poorly known from the

literature, we present in this note new constraints on the eruption style by studying the typology and morphology

of tephra particles especially volcanic ash to infer eruption parameters such as the nature of the magmatic

fragmentation mechanism and the nature of the transport process (e.g. Heiken et al., 1985, Riley et al., 2003;

Ersoy et al., 2007).

Methods and results

After the eruption 19 tephra samples were collected at distances of 55-95 km downwind from the volcano.

These were sieved at the “Laboratoire Magmas et Volcans” in Clermont-Ferrand, France and ~600 particles for

Figure 1. Geodynamic map of the Ecuadorian Volcanic arc showing the location of the Reventador volcano.

Reventador

volcano

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 180-183

181

a c d e

j i h g f

b

Figure 3. Binocular images. a,b,f,g: juvenile (moderate to highly) vesicular grain; c,d,h,i: non-vesicular juvenile glass; e,j: lithic grains.

each three fractions (90 to 125μm, 125 to 180μm and 180 to 250μm) of each sample have been identified under

a binocular. To acheive this study some samples were selected for Scanning Electron Microscopy (SEM)

allowing to characterize the surface texture of particles. Below we summarize some results obtained for three

samples (fig. 2) at different distances (proximal part ~55 km, intermediate part ~75 km and distal part ~85 km)

from the vent.

Observations of these particles under the binocular reveals 3 textural classes (Fig. 3): (1) juvenile glass (with

mainly whitish grains and dark grains in lesser quantity) divided into 3 subclasses (1a) dense (1b) moderately

vesicular (1c) highly vesicular (2) free crystal (principally plagioclase crystals and few pyroxenes) (3) lithic

grains (fragments of the basement, old lavas from the volcano, etc) divided into 2 subclasses (2a) altered grains

with a characteristic reddish color (2b) other xenolithic grains.

Figure 4 shows some selected SEM images. Two types of vesicular grains were identified (Fig 4. a,b,c). The

first type (Fig 4. a,b) comprises highly vesicular grains with elongated shapes, sharp and irregular contours, and

the vesicles form tubular to capillary structures with narrow diameter (few μm). The second type (Fig 4. c) are

moderate to highly vesicular grains with an equant shape, a low irregular contour and mainly subspherical

vesicles with a variable diameter (few μm to tens of μm). Many juvenile glass particles are non-vesicular and

Figure 2. Evolution of the grain’s proportion for 3 samples from proximal, intermediate and distal parts; A: highly (black and white) juvenile glass; B: moderate (black and white) juvenile glass; C: non-vesicular juvenile glass; D: crystal (plagioclase and pyroxene); E: lithic grains (altered grain and others).

> 180 m

distal part

> 180 m

> 90 m

%

> 125 m > 125 m

Proximal part

> 180 m

> 90 m

% Intermediate part

> 125 m

> 90 m

%

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 180-183

182

show blocky to platy morphologies with local curviplanar and peculiar stepped fractures (Fig. 4, d,e) but no

quenching cracks features were observed. Free crystals are commonly observed (Fig. 4, e) and many of these

exhibit a glass coating. Some altered lithic grains with a subrounded to subangular shape are also recognized (not

shown in Fig. 4).

Discussion and conclusion

The proportion of each class of grains in the selected samples (Fig. 3) shows that the deposit is heterogeneous

in terms of component concentration. The abundance of juvenile glasses, especially the moderately to highly

vesicular particles, is high in relatively coarse-grained fractions (180 to 250 μm). Free crystals are more

abundant in tephra samples collected closer to the source (~10% at about 55 km from the vent, Fig. 2) and tend

to concentrate in the fine-grained fractions (~20%). The lithic particles represent from 2% to 10% of the samples

and, as for crystals, their concentration increases toward the source, especially in the fine-grained fractions of the

samples. On the whole, our component analyses indicate that ample density fractionation took place in the plume

during transport to the west: highly vesicular grains were concentrated in distal areas and in large grain-size

fractions whereas dense particles as non-vesicular glass shards and free crystals accumulated in the proximity of

the source. The wide range morphology of the juvenile particles suggests complex fragmentation processes at the

vent. On one hand, the highly vesicular grains with frothy to tubular textures support a dominant magmatic

fragmentation mode. On the other hand, the abundance of dense vitric glass shards, as well as the surface

textures (Wohletz, 1983), and also xenolithic grain point to some magma-water interactions, probably in the

conduit. These preliminary results lead us to portray the Reventador eruption of the November 3, 2002 as a

powerful VEI 4 event characterized by bimodal magmatic/phreatomagmatic fragmentation processes at the vent.

Work in progress aims to investigate the tephra fall deposits in more details to better characterize the eruption

dynamics.

Figure 4. SEM images. a, b : vesicular grain with tubular structure and elongated shape ; c : vesicular grain with an equant shape and an irregular contour; d: glass with peculiar stepped surface; e: platy glass with no quenching cracks features; f: crystal of plagioclase.

a b c

d e f

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 180-183

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References Aguilera, E., Almeida, E., Balseca, W., 1988. El Reventador: an active volcano in the sub-Andean zone of Ecuador.

Rendiconti della Società Italiana di Mineralogia e Petrologia, 43: 853-875. Ersoy, O., Gourgaud, A., Aydar, E., Chinga, G., Thouret, J.-C., 2007. Quantitative scanning-electron microscope analysis of

volcanic ash surfaces: Application to the 1982-1983 Galunggung eruption (Indonesia), Bull. Geol. Soc. of Amer., 119: 743-752.

Fierstein, J., Nathenson, M., 1992. Another look at the calculation of fallout tephra volumes. Bull. Volcanol., 54: 156-167. Heiken, G., Wohletz, K., 1985. Volcanic ash, 246 pp., Univ. of Calif. Press, Berkeley. Hall, M., Ramón, P., Mothes, P., Le Pennec, J-L., García, A., Samaniego, P., Yepes, H., 2004. Volcanic eruptions with little

warning: the case of volcán Reventador’s surprise November 3, 2002 eruption, Ecuador. Revista Geológica de Chile, 31: 349-358.

Le Pennec, J-L., Hidalgo, S., Samaniego, P., Ramos, P., Yepes, H., Eissen., J-P., 2003. Magnitud de la Erupcion del 3 de Noviembre del 2002 del Volcán Reventador, Ecuador. Escuela Politécnica Nacional, Terceras Jornadas en Ciencias de la Tierra, Abstract, 94-96.

Samaniego, P., Eissen, J-P., Le Pennec, J-L., Robin, C., Hall, M.-L, Mothes, P., Chavrit, D., Cotten, J., 2008, in Press. Pre-eruptive physical conditions of El Reventador volcano (Ecuador) inferred from the petrology of the 2002 and 2004-05 eruptions. J. Volcanol. Geotherm. Res.

Ridolfi, P, PhD., Puerini, M., Renzulli, A., Menna, M., Toulkeridis, T, 2008, in press. The magmatic feeding system of El Reventador volcano (Sub-Andean zone, Ecuador) constrained by texture, mineralogy and thermobarometry of the 2002 erupted products. J. Volcanol. Geotherm. Res.

Riley, C. M., Rose, W. I., Bluth, G. J. S., 2003. Quantitative shape measurements of distal volcanic ash. J. Geophy. Res., 108, B10.

Wohletz, K.H., Mechanisms of hydrovolcanic pyroclast formation: grain-size, scanning electron microscopy, and experimental studies. J. Volcanol. Geotherm. Res., 17: 31-63.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 184-187

184

Three-dimensional P- and S-wave seismic attenuation models in central Chile - western Argentina (30°-34°S) from local recorded earthquakes

P. Deshayes1, T. Monfret

1, M. Pardo

2, & E. Vera

2

1 Géosciences Azur, Université de Nice, IRD, Sophia-Antipolis, Valbonne, France

2 Departamento de Geofísica, Universidad de Chile, Santiago, Chile

KEYWORDS : flat slab, tomography, velocity, attenuation, mineralogy

Introduction

Beneath the Chilean coast, the Nazca plate subducts at a constant convergence velocity of 6.4cm/year

(Angermann et al, 1999). One of the first subduction zone segments to be identified as a low-angle subduction

zone, or flat slab, has been observed between 28°S and 33°S (Figure 1A). This region has coincided with an

absence of Quaternary active volcanoes since 9-10 Ma. To the north and south of that region, the subduction

zone dips “normally” with an angle of 30° (Figure 1B) (Barazangi and Isacks, 1976). Pardo et al (2004) and

Anderson et al. (2007) performed in the Central Chile and Western Argentina precise hypocenter location of

microearthquakes locally recorded by temporary seismic networks deployed within the zone for several months

and also by the Chilean permanent seismic network. They found out that the distribution of the seismicity, apart

to confirm that the flat geometry of the slab occurs at around 100 km depth with a slight western dipping of 5°,

forms at that depth a clear cluster, limited by a 30-km thick, 30-km wide and 160-km long “finger” shape, along

the expected Juan Fernandez Ridge (JFR) track (Yañez et al, 2002).

In a normal-angle subduction zone, volcanism is commonly generated by partial melting into the mantle if

temperature and water content conditions are fulfilled, and then by migration of the melting through the

continental crust till the surface. However, temperature variations, partial melting or water content may produce

similar seismic velocity signature, and therefore it is often difficult to associate observed velocity anomaly with

one of these causes (Wiens and Smith, 2003). Seismic attenuation responses to partial melting and temperature

may differ, which helps to discriminate between these causes and hence, to better characterize the medium.

In this study, we determine three dimensional attenuation models for P- and S-wave in Central Chile-Western

Argentina (30-34°S), based on frequency analysis of seismic displacement of local earthquakes, initial layered

attenuation models and the three-dimensional velocity models proposed by Pardo et al (2004).

Hypocenter location and velocity models

The three-dimensional velocity structures and precise hypocenter location of Pardo et al (2004) were jointly

obtained by tomography of P and S arrival times of earthquakes of magnitude between 1 and 6, locally recorded

during the OVA99 and CHARSME experiments. In the flat slab segment, Pardo and al (2004) put in evidence a

clear double seismic zone till 90 km depth while the oceanic crust is mainly characterized by positive P- and S-

wave velocity anomalies (Figure 2). In these velocity models, JFR track coincides with low Vp/Vs ratios (Pardo

et al, 2004). Moreover, at 33.5°S where the subduction zone is “normal”, a strong negative P- and S-wave

velocity anomaly is discerned in the continental crust, underneath the volcanoes: this anomaly might be

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 184-187

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correlated either with partial melt or high temperatures (Figure 2).

In the zone, using CHARGE data, Wagner et al (2005) determined a three-dimensional velocity models of P

and S waves, mainly of the upper continental mantle while Alvarado et al (2007) proposed a continental crustal

model of the flat slab region, based on forward modelling of crustal parameters.

Figure 1. Local seismicity (black dots) recorded at temporary seismic stations (gray squares), during the OVA99 (November 1999-January 1999) and CHARSME (November 2002-March 2003) experiments (Pardo et al., 2004). Quaternary active volcanoes are represented by gray triangles. Isodepths of the slab are from Pardo et al (2004) and spacing between two isodepths is 20 km. East-West cross sections showing the local seismicity (black dots) along two transects at 31°S (A) and 33.5°S (B). Topography is shown on top of the cross-sections with an exaggerated vertical scale. A black reverse triangle indicates the location of the Peru-Chile Trench.

Attenuation models

The three-dimensional attenuation structures (Figure 3), represented by the Q quality factor anomalies for P-

( Qp) and S-wave ( Qs), were determined by inverting t* parameter through an adapted version of the TLR3

algorithm commonly used in velocity tomography (Latorre et al , 2004; Monteiller et al, 2005). t* (=t/Q) is

achieved by fitting the spectral displacement amplitude of P and S waves (Abercrombie, 1995) and t, the

estimated seismic wave travel-time between the source and the seismic station, is calculated in the three

dimensional velocity models of Pardo et al (2004). In this study, around 10700 t*-values were available to carry

out the attenuation tomography, however azimuthal coverage and number of ray paths were not as spatially well

distributed as for the velocity tomography of Pardo et al (2004).

While in velocity tomography the initial velocity models for P and S waves are generally well constrained

(multilayered models), for seismic attenuation tomography studies, the initial model is commonly a half-space

homogeneous constant Q-model, badly constrained. So, in order to have a more realistic initial Q-model, we

establish well-constrained one-dimensional Qp and Qs layered models through the probabilistic Metropolis-

Hastings method (Metropolis et al, 1953). The best one-dimensional Q models for P and S waves increase with

depth till 140 km where a low Q-value zone is found before augmenting with depth.

In the flat slab region (31.5°S) and at depths shallower than 20 km, the seismic attenuation (1/Q) is highly

laterally heterogeneous, showing Qp and Qs anomalies of opposite signs for the same positions (Figure 3),

may be indicating complexity in fluid content and/or fluid flow circulation. Moreover, these anomalies do not

seem directly correlated with specific velocity perturbations. On the other hand, the lower parts of continental

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crust as the upper continental mantle are almost homogeneous in attenuation, with mainly positive anomalies in

Qp and Qs, meaning that seismic waves are less attenuated than they were expected in the initial attenuation

models. Although temperature, water presence and partial melt have similar influence on seismic velocity,

attenuation (or Q) anomalies are more sensitive to temperature variations than seismic velocities, in particular

attenuation increases (Q decreases) when temperature increases. In the “normal” subduction zone region (south

to 33°S), crustal attenuation variations extend in all the continental crust and show a different behaviour than in

the flat slab region. Moreover, seismic waves are less attenuated in the continental crust beneath the volcanoes

(69.5°W) than in the surroundings areas, whereas seismic velocities strongly decrease, which could be done

mainly by partial melting (see Figure 2 and Figure 3). Nevertheless, the ray path coverage is poorer there (gray

zones in Figure 3), which might affect the quality of the attenuation models.

Figure 2. East-West cross-sections of final P-, and S-velocity models. Color scale indicates percentage velocity deviation in the flat slab region (31.5°S) and in the dipping slab region (33.5°S). Zones with poor resolution are in gray. Moho depths (white line) are from Tassara (2005) and white dots indicate earthquakes. Quaternary active volcanoes are represented by black reversed triangle. Exaggerated topography is added on top of each cross-section.

Figure 3. East-West cross-sections of final P-, and S-velocity attenuation models. Color scale indicates percentage attenuation deviation in the flat slab region (31.5°S) and in the dipping slab region (33.5°S). Zones with poor resolution are in gray. Moho depths (white line) are from Tassara (2005) and white dots indicate earthquakes. Quaternary active volcanoes are represented by black reversed triangle. Exaggerated topography is added on top of each cross-section.

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Conclusions

Attenuation tomography is complementary to velocity tomography to characterize the medium. The

simultaneous study of the attenuation and velocity perturbations in the medium, allow to differenciate the

temperature and melting origin of the anomalies, while the velocity tomography models, single-handed, can not

do it.

In order to improve the coverage of Qp and Qs attenuation models south to 33°S (Figure 3), we should use in a

future work, additional station-source ray paths from CHARGE data and/or the permanent Chilean seismic

network for example.

References

Abercrombie, R. E., 1995a, Earthquake source scaling relationships from -1 to 5 ML, using seismograms recorded at 2.5 km depth, J. Geophys. Res., 100: 24015-24036.

Alvarado, P., Beck, S., and Zandt, G., 2007. Crustal structure of the south-central Andes Cordillera and backarc region from regional waveform modelling, Geophys. J. Int., 170: 858-875.

Anderson, M., Alvarado, P., Zandt, G., and Beck, S., 2007. Geometry and brittle deformation of the subducting Nazca Plate, Central Chile and Argentina, Geophys. J. Int., 171: 419-434.

Angermann, D., Klotz, J. and Reigber, C., 1999. Space-geodetic estimation of the Nazca-South America Euler vector, Earth Planet. Sci. Lett., 171: 329-334.

Barazangi, M., and Isacks, B., 1976. Spatial distribution of earthquakes and subduction of the Nazca Plate beneath South America. Geology, 4: 686-692.

Latorre, D., Virieux, J., Monfret, T., Monteillet, V., Vanorio, T., Got, J.-L., and Lyon-Caen, H., 2004. A new seismic tomography of Aigion area (Gulf of Corinth, Greece) from the 1991 data set, Geophys. J. Int., 159: 1013-1031.

Metropolis, N., Rosenbluth, A. W., Rosenbluth, M. N., Teller, A. H., and Teller, E., 1953. Equation of State Calculations by Fast Computing Machines, Journal of Chemical Physics, 21: 1087-1092.

Monteillet,V., Got,J. L., Virieux,J. and Okubo,P., 2005. An efficient algorithm for double-difference tomography and location in heterogeneous media, with an application to the Kilauea volcano, J. Geophys. Res., 110: doi:10.1029/2004JB003466.

Pardo, M., Monfret, T., Vera, E., Yanez, G., and Eisenberg, A., 2004. Flat-Slab to Steep Subduction Transition Zone in Central Chile-Western Argentina: Body Waves Tomography and State of Stress. AGU Fall Meeting Abstracts, B164.

Tassara, A., 2005. Structure of the Andean continental margin and causes of its segmentation, PhD thesis, Freie Universität Berlin Institut für Geologische Wissenschaften, 165p.

Wagner, L. S., Beck, S., and Zandt, G., 2005. Upper mantle structure in the south central Chilean subduction zone 30°S to 36°S. J. Geophys. Res., 110, doi :10.1029/2004JB003238.

Wiens, D.A., and Smith, G.P., 2003, Seismological constraints on structure and flow patterns within the mantle wedge: in Eiler, J., ed., Inside the subduction factory: American Geophysical Union Geophysical Monograph 138: 59–81 doi: 10.1029/138GM05.

Yáñez, G., Cembrano, J., Pardo, M., Ranero, C. and Selles, D., 2002. The Challenger-Juan Fernandez-Maipo major tectonic transition of the Nazca-Andean subduction system at 33-34°S: geodynamic evidence and implications, J. of South Am. Earth Sc., 15: doi:10.1016/S0895-9811(02)00004-4.

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Magnetotelluric study of the Parinacota and Lascar volcanoes

D. Díaz, D. Brändlein, & H. Brasse

Freie Universität Berlin, Fachrichtung Geophysik, Malteserstr. 74-100, 12249 Berlin, Germany

(d.diaz@geophysik.fu-berlin.de)

KEYWORDS : electromagnetic methods, magnetotellurics, electrical properties

Introduction

Electromagnetic methods allow to detect zones with different electrical properties. Among them, the

magnetotelluric sounding method, with a large range of frequencies, allows to measure the electrical properties

of rocks at considerable depths.

In the study of volcanoes and hydrothermal systems, magnetotellurics has been widely used considering the

different electrical properties expected in these structures, due to hydrothermal fluids, gas or melt in contrast

with the surrounding rocks (see, e.g., Heise et al., 2008; Müller et al., 2004).

This investigation considers two regions of interest, the first one includes the zone around Parinacota (6362 m,

18°09’S, 69°08’W), a subduction related stratovolcano situated in the limit of Bolivia and Chile. The second

zone is more to the south, around Lascar volcano (5592 m, 23°22’S, 67°44’W), located on the eastern side of the

Salar de Atacama basin in northern Chile. It has been one of the most active volcanoes of the central Andes in

the last years. Its recent activity is characterized by repetitive dome growth and subsidence, accompanied by

degassing and explosive eruptions of various magnitudes (Pavez et al., 2006).

Preliminary data

Data acquisition

During October and November 2007, magnetotelluric and audio magnetotelluric sites were built in the area

close to Lascar and Parinacota volcanoes. While the AMT data could reach periods between 0,01 to 1000 s,

which is appropriate for a more shallow view, these sites were installed in the proximities of the volcanoes, the

MT sites, which can reach longer periods and larger depths, were installed on a profile south of Lascar, and as an

outer ring in the Parinacota region (Figure 1).

First results

To obtain the apparent resistivity curves from the time series, the real and imaginary parts of the impedance

tensor were calculated with the robust Egbert processing program (Egbert, 2002). With this data, the first steps

were to see the changes of the apparent resistivity with the period, the phase with the period, and also to

calculate induction vectors and phase tensor ellipses.

Induction vectors are functions of the ratio of the vertical and horizontal components of the magnetic field,

which in turn are functions of period and the horizontal resistivity gradient (Wiese, 1962). According to “Wiese

convention”, these vectors should point away from conductive zones.

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The phase ellipse is a graphical representation of the phase tensor, which expresses how the phase relationships

change with polarization, independent of the galvanic distortion produced by heterogeneities in the near surface

(Caldwell et al. 2004).

Figure 1. Study zone and measurement sites, upper near Parinacota and lower near Lascar.

Induction vectors and phase ellipses have been calculated for every measured period of the AMT sites, from

0.00391 s, as the lowest period, until 1024 s.

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From a first analysis of the apparent resistivity curves for the different sites, deep large conductive bodies seem

to be absent in these zones.

Besides this, topographic corrections have been developed for these zones, so the next step will be inversion of

the data, and constructing 2D and 3D models which could fit with these results.

First results of conductivity models will be presented.

References Brasse, H., Lezaeta, P., Rath, V., Schwalenberg, K., Soyer, W. & Haak, V. 2002. The Bolivian Altiplano conductivity

anomaly. Journal of Geophysical Research, 107 (B5), doi:10.1029/2001JB000391. Caldwell, T.G., Bibby, H.M. & Brown, C. 2004. The magnetotelluric phase tensor. Geophysics Journal International, 158:

457-469. Echternacht, F., Tauber, S., Eisel, M., Brasse, H., Schwarz, G. & Haak, V. 1997. Electromagnetic study of the active

continental margin in northern Chile. Physics of the Earth Interiors, 102: 69-87. Egbert G.D. 2002. Processing and Interpretation of Electromagnetic Induction Array Data. Surveys in Geophysics, 23: 207-

249. Heise, W., Caldwell, T.G., Bibby, H.M. & Bannister S.C. 2008. Three-dimensional modeling of magnetotelluric data from

the Rotokawa geothermal field, Taupo Volcanic Zone, New Zeland. Geophysical Journal International, doi:10.1111/j.1365-246X.2008.03737.x

Mackie, R., Smith, J.T. & Madden, T.R. 1994. Three-dimensional electromagnetic modeling using finite difference equations: The magnetotelluric example. Radio Science, Vol. 29, n.4, 923-935.

Müller, A., Haak, V. 2004. 3-D modeling of the deep electrical conductivity of Merapi volcano (Central Java): integrating magnetotellurics, induction vectors and the effect of steep topography. Journal of Volcanology and Geothermal Research, 138: 205-222.

Pavez, A., Remy, D., Bonvalot, S., Diament, M., Gabalda, G., Froger, J-L., Julien, P., Legrand, D. & Moisset, D. 2006. Insight into ground deformations at Lascar volcano (Chile) from SAR interferometry, photogrammetry and GPS data: Implications on volcano dynamics and future space monitoring. Remote Sensing of Environment, Volume 100, Issue 3: 307-320.

Wiese, H. 1962. Geomagnetische Tiefentellurik Teil II: Die Streichrichtung der Untergrundstrukturen des Elektrischen Widerstandes, Erschlossen aus Geomagnetischen Variationen. Pure and Applied Geophysics, 52: 83-103.

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Volcán Llaima (38.7ºS, Chilean Southern Volcanic Zone): Insights into a dominantly mafic and ‘hyperactive’ subduction-related magmatic system

M. A. Dungan1, C. Bouvet de Maisonneuve

1, D. Sellés

2, J. A. Naranjo

3, H. Moreno

4,

C. Langmuir2, O. Reubi

5, S. Goldstein

6, J. Jweda

6, S. Escrig

2, O. Bachmann

7, &

B. Bourdon5

1 Section des Sciences de la Terre, Université de Genève, 13 rue des Maraîchers, 1205 Genève, Switzerland

(michael.dungan@terre.unige.ch; caroline.bouvet@terre.unige.ch) 2 Department of Earth and Planetary Sciences, Harvard University, 20 Oxford Street, Cambridge MA 02139, USA

(Selles@fas.harvard.edu; langmuir@fas.harvard.edu; escrig@fas.harvard.edu) 3 SERNAGEOMIN, Avenida Santa Maria 0104, Providencia, Santiago, Chile (jnaranjo@sernageomin.cl)

4 OVDAS-SERNAGEOMIN, Cerro Nielol, Sector antennas, Casilla 641, Temuco, Chile

(hmoreno@sernageomin.cl) 5 Institute of Isotope Geology and Mineral Resources, ETH-Zentrum, 8092 Zürich, Switzerland

(olivier.reubi@bristol.ac.uk; bourdon@erdw.ethz.ch) 6 Lamont-Doherty Earth Observatory, Columbia University, Palisades NY 10964, USA

(steveg@ldeo.columbia.edu; jjweda@ldeo.columbia.edu) 7 Department of Earth and Space Sciences, University of Washington, Seattle WA 98195, USA

(bachmano@u.washington.edu)

KEYWORDS : Andean arc magmatism, Chilean Southern Volcanic Zone, Llaima, conduits and chambers, magma

replenishment

Introduction

Holocene Volcán Llaima, one of Chile’s historically most active volcanoes, is broadly typical of Holocene

frontal arc centers of the Andean Southern Volcanic Zone (SVZ) that lie between the latitudes of 38.4° S

(Lonquimay) and 41° S (Osorno). It is located in proximity to the northern Liquiñe-Ofqui Fault Zone, it is

dominated volumetrically by evolved basaltic and basaltic andesitic magmas (<6.5 wt% MgO; 51-54 wt% SiO2),

and magma evolution is primarily by fractional crystallization without large contributions from assimilated upper

crust. Holocene volcanism began with eruption of the Curacautin Ignimbrite (13.5 ka; >10 km3; 52-58 wt%

SiO2), and early Holocene activity featured several large explosive eruptions (52-69 wt% SiO2) separated by

relatively long intervals (>1500 years?). During the last ~4000 years, repose times have shortened, eruptive

products with >60 wt% SiO2 are absent or extremely rare, and eruptive centers have alternated between main-

cone vents that define a N–S orientation and transverse fissural vents located on the lower NE, NW, and SW

flanks. Our understanding of the late Pleistocene eruptive products of the Llaima magmatic system is limited by

edifice destruction related to caldera collapse and glacial erosion, plus extensive burial by Holocene eruptive

products, but pre-Holocene Llaima magmas were primarily mafic andesite, rather than basalt.

Ongoing studies of historic and prehistoric eruptive products, plus observations during eruptive activity in

2008, suggest that late Holocene activity of Volcán Llaima has been characterized primarily by: (1) strombolian

eruptions of spatter and lava from summit craters or centers on the upper flanks of the edifice that were triggered

by magma replenishment, which in part remobilized crystal-rich residual products of previous eruptions by gas

sparging (degassing of impeded magma leading to gas-streaming through and mobilization of resident mush),

and (2) the generation of large and far-traveled lahars by larger eruptions due to interactions with voluminous

glacial ice on the steep upper slopes of the edifice. Lahars are the primary volcanic hazard at Llaima, and their

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ability to deeply erode canyons on the lower slopes of the volcano has led to extensive and rapid ‘resurfacing’ of

the edifice.

The goals of our investigation are to: (1) characterize Llaima in sufficient geological and geochemical detail to

place constraints on the relationships among magma evolution, conduit geometry and location, and eruptive

behavior, (2) use these data and insights to address questions about the nature of the asthenospheric mantle and

magma generation processes beneath Llaima, and (3) enhance the value of regional-scale, along-arc assessments

of magma genesis and evolution by creating a ‘reference volcano’ data-set for calibration of trends that are

defined in large part by reconnaissance sampling. This contribution is based on geologic mapping and 14C

chronology (Naranjo and Moreno, 2005), plus preliminary interpretations of results obtained in 2007-2008 (150

major and trace element XRF analyses in hand, and ~400 analyses in progress). First-order characterization of

whole-rock samples will be followed by mineral and melt inclusion chemistry, trace element determinations by

ICP-MS, and isotopic measurements (Sr, Pb, Nd, Os, Hf, O, U-series disequilibrium, and 10Be) on a select suite

of samples.

Structural control of conduit and vent geometry

Volcán Llaima lies in a half-graben, wherein the main cone is offset ~9 km to the west from a N10°E-trending

segment of the northern Liquiñe-Ofqui Fault Zone (LOFZ). Llaima’s deeply eroded early Pleistocene

predecessor, Sierra Nevada, lies directly on the LOFZ to the NNE of Llaima. The alignment of Llaima main-

cone vents parallel to the LOFZ implies that there is a buried, secondary splay that plays an important role in

magma ascent. The localization of frontal arc volcanoes along or near the LOFZ in the Lonquimay to Osorno

portion of the SVZ may be one factor in minimizing crustal contamination in this part of the SVZ; i.e. ease of

passage through the crust. Fissural flank vents of volumetrically secondary importance (<4000 ka) are present on

the lower NE, NW, and SW flanks of Llaima. The NE fissural system, which is the most long-lived of the three,

has an E-W orientation (normal to LOFZ) and it lies directly to the south of a major E-W dike swarm in Sierra

Nevada’s western flank. This implies a transtensional stress component along this part of the LOFZ and a

regional rather than an edifice-related control on these structures. The NW and SW Llaima fissures have an

intermediate orientation (~N40°E). Preliminary results suggest that young main-cone lavas tend to be less

evolved and less diverse (~51-55 wt% SiO2, dominantly <53.5% SiO2) than broadly contemporaneous products

of three episodes of fissural eruptions (~51-59 wt% SiO2); this inference will be tested rigorously on the basis of

the chemistry of samples collected in 2008.

First-order chemical signature of mafic Llaima magmas

Mafic magma compositions at Llaima are not significantly different from those at Villarrica or Puyehue, in that

they have low abundances of Rb, K, Th, U, and Cs in combination with low La/Yb (2.5-3.7), Zr/Y (2.5-4.9) and

Nb/Y (0.10-0.16). These narrow ranges in ratios contrast with the much greater diversity at the Tatara-San Pedro

complex (36° S: La/Yb = 4.5-20, Zr/Y = 3.5-9.5; Nb/Y = 0.1-0.6), which apparently is a manifestation of mantle

heterogeneity at a volcano located at a greater distance from the trench and on thicker lithosphere. The apparent

lack of significant heterogeneity in mantle magma sources at Llaima suggests that source region parameters may

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be reduced to variable partial melt fraction, contributions from diverse slab-derived components, and the

relationships between these two.

Importance and role of mafic magma replenishment

The long-term temporal evolution of the Holocene Llaima edifice appears to converge on a higher eruption-

frequency and a tendency for the eruption of compositionally similar evolved basalts (5.8-6.3 wt% MgO). A

corollary is that the ‘hyperactivity’ of the last four, or more, centuries is driven by a high frequency of mafic

magma replenishment into the upper crustal conduit system. Llaima has erupted 50 times since 1640. Although

only a few of these produced sufficient volumes of lava and tephra to render their products currently accessible

for study (many very short and volumetrically minor strombolian events within summit craters), the average

repose period during this period is 7.4 years and the only repose lasting longer than 35 years is the interval

between the 1640 and 1751 eruptions (Table 2 in Naranjo and Moreno, 2005). The latter two are among the four

large historic events (including 1780-90 & 1955-57). The historic alternation between mafic basaltic andesite

(e.g., 1640, 1927, 1994) and evolved basalt (e.g., 1780-90, 1751, 1945, 1955-57, 2008) is evidence that large

eruptions at Llaima are immediate responses to the quantity of new mafic input. The sudden onset of the January

1-2, 2008 eruption of Llaima, without significant seismic precursors during preceding days, suggests that such

inputs ascend rapidly to shallow depths. The 2008 eruption terminated in late February with voluminous

degassing, following six weeks of intermittent and weak strombolian activity.

Information derived from lavas is consistent with these observations: (1) Llaima mafic magmas are crystal-rich

(20-35 vol%) and the phenocryst populations are dominated by plagioclase (>80%). (2) Almost every large

plagioclase crystal in almost all main-cone lavas is characterized by sieve-textured resorption, and (3) The onset

of the large 1780-90 eruption, 30 years after the large 1751 eruption, was characterized by crystal-rich pahoehoe

followed by aa flows. The high effective viscosity of dry, crystal-rich, evolved basaltic magma is impossible to

reconcile with the fluidity of pahoehoe. We infer that the effective viscosity of a melt-solid-gas mixture may be

sufficiently lowered by a sudden injection of hot, water-rich gas into a ‘rheologically stiff’ crystal mush. This

scenario could be created by voluminous gas release from recharged mafic magma that stalled against partially

solidified, residual magma which previously had undergone low-P decompression crystallization. Some or all of

the magma erupted during such events would be the inflated and remobilized products of immediately preceding

episodes rather than new input, hence the crystal cargoes of resorbed plagioclase. This eruptive mechanism is

likely to be a factor in highly active systems subject to frequent replenishment by water-rich mafic magmas

which have not substantially degassed during ascent through the crust.

Volcán Llaima: an unusually laharogenic volcano

Lahars have been observed during only a small fraction of historic Llaima eruptions, but geologic evidence

suggests that most significant eruptions during the last ~4000 years from main-cone vents have generated large

lahars. The distal deposits of the vast majority of these lahars are buried or eroded, but they have left a record on

the lower flanks of the main cone. The short and modest strombolian eruption of January 1-2, 2008 triggered

glacier melting and a lahar that eroded deeply into the flank, exposing older canyon-filling lavas and clastic

deposits that reflect repeated incision and infilling of precursor canyons that followed similar courses. Such

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canyons and scoured lavas are widespread on the flanks of the main cone, and such lahars have interacted in

vastly different ways with the flank fissural units as a function of the proximity of the fissural cones to the

summit: i.e., those which are located high on the steep north flank of Llaima have been extensively eroded,

whereas the distal chains of cones on the west flank have served as barriers and formed a basin in which lahar

accumulation has dominated (local escape through erosional gaps in the fissural ridge). The factors favoring

lahar generation are: (1) A high Holocene edifice (~3100 m) built on pre-Holocene edifice remnants. (2)

Dominantly mafic strombolian eruptions of spatter and lava fed by summit craters. (3) Climatic and topographic

conditions which lead to rapid accumulation of glacial ice on steep flanks.

Implications

Volcán Llaima provides an opportunity to assess magma generation and evolution, eruption dynamics, and

edifice modification. As all these phenomena are linked, and bear on the assessment of volcanic risks, it is

crucial to approach volcanoes as integrated systems from the magma source to the surface.

Reference Naranjo, J.A., Moreno, H., 2005 – Geología del Volcan Llaima, Región de la Araucanía. Carta Geológica de Chile, Serie

Geología Basica, No. 88, Servicio Nacional de Geología y Minería – Chile, Subdirección Nacional de Geología.

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Sedimentologic, paleontologic, and ichnologic evidence for deep-marine Miocene deposition in the present Intermediate Depression of south-central Chile (38°30’-41°30’S)

Alfonso Encinas1, Kenneth L. Finger

2, & Luis Buatois

3

1 Departamento de Ciencias de la Tierra, Universidad de Concepción, Barrio Universitario s/n, Concepción, Chile

(aencinas@udec.cl) 2 University of California Museum of Paleontology, 1101 Valley Life Sciences Building, Berkeley, CA 94720-4780,

USA (kfinger@berkeley.edu) 3 Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, SK S7N 5E2,

Canada (luis.buatois@usask.ca)

KEYWORDS : Miocene, south-central Chile, deep-marine, subsidence

Introduction

Neogene marine strata occur at different localities along the Chilean coastline (Cecioni, 1980). Previous

sedimentological studies generally refer to these deposits as shallow marine (e.g., Cecioni, 1978). However,

more recent sedimentological, ichnological and paleontological studies indicate that they are deep-marine,

deposited at bathyal depths (~500-2000 m) during an interval of major Miocene subsidence that took place along

the Chilean forearc (Encinas et al., 2008 and references therein).

Although most exposures of these deposits are along the coast, outcrops also occur inland in the region located

between Temuco and Puerto Montt (38°30’-41°30’S) (Figure 1), where they also crop out in the eastern Coastal

Cordillera and the Intermediate.

Former studies have shown significant differences in the sedimentology and paleobathymetry of these deposits.

In the Temuco area (Figure 1), Osorio and Elgueta (1991) determined that sedimentation took place in at lower-

bathyal depths (~2000 m) during the middle-early to late Miocene, based on their study of foraminifera

recovered from ENAP’s Labranza 1 and Cunco 1 boreholes. In contrast, Le Roux and Elgueta (2000) interpreted

that sedimentation in the Valdivia and Osorno-Llanquihue area (Figure 1) occurred in a much shallower

environment. According to them, estuarine sedimentation during the late Oligocene-early Miocene was followed

by deposition in deeper coastal embayments during the middle Miocene. However, listed foraminiferal

assemblages from Neogene deposits from this area (i.e., Martínez-Pardo and Pino, 1979; Marchant and Pineda,

1988; Marchant, 1990) include lower bathyal species. Deep-water deposition is also suggested by the presence

of the trace fossil Chondrites isp. (Covacevich et al., 1992), and turbiditic facies (Le Roux and Elgueta, 2000)

that, although do not occur exclusively in these environments, are common in deep-marine settings. These data

suggest that Miocene deposition in the Valdivia and Osorno-Llanquihue basins took also place in a deep-marine

environment.

To try to unravel the sedimentary environment of Neogene deposits and the tectonic history of the area located

in the present Coastal Cordillera an Longitudinal Depression between Temuco and Puerto Montt (38°30’-

41°30’S) during that period we carried out a thorough sedimentological, ichnological and micropaleontological

study of both outcrops and ENAP well sections.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 195-198

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Figure 1. Map of the study area showing locations of outcrops of the Neogene marine successions (gray color) and the most important boreholes drilled by ENAP in the Intermediate Depression (open circles).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 195-198

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Results

The Neogene succession in our study area discordantly overlies the metamorphic basement and coal-bearing

deposits of the Pupunahue-Catamutun Formation. It includes a basal breccia overlain by gray sandy siltstone and

minor medium to coarse-grained sandstone and breccia. Sedimentary features typical of gravity flows are

present, such as convolute lamination and rhythmic alternating sandstones and siltstones. Fine siltstones show

the presence of abundant Chondrites isp. and minor Zoophycos isp.

Foraminiferal analyses of ENAP well samples reveal the presence of two very different units. The upper unit,

which is the focus of this study, consists of sedimentary rock bearing the rich foraminiferal fauna characteristic

of the regional Miocene. In stark contrast, the lower unit known as the Pupunahue-Catamutún Formation

includes only internal molds of Globigerinatheka sp., a planktic genus restricted to the late early to late Eocene

(Pearson et al., 2007).

The upper Miocene unit yielded mixed associations of littoral, neritic, and bathyal species of benthic

foraminifera. The deepest-dwelling taxa in the majority of samples indicate minimum depths of deposition in the

lower-bathyal (2000–4000 m) zone. Among the lower-middle and lower-bathyal indicators in the Chilean

Miocene are species of Bathysiphon, Melonis, Osangularia, Pleurostomella, and Siphonodosaria that are similar

or identical to those Van Morkhoven et al. (1986) classified as cosmopolitan deep-water taxa.

A refined age span of the studied succession remains somewhat elusive, as it contains few planktic

foraminifers and index species are very scarce. Of the samples studied in this work, only three localities yielded

planktic forams with concurrent ranges that confine the age to the middle to late Miocene interval. In all other

respects, the foraminifer fauna is very similar to that we recovered from late Miocene-early Pliocene outcrops in

the coastal area of south-central Chile (Finger et al., 2007).

Discussion

Sedimentological studies show the occurrence of a basal, shallow marine breccia, overlain by a succession of

sandstone and siltstone with sedimentary facies characteristic of gravity flows that are typical of deep marine

settings. Ichnological studies indicate the presence of abundant Chondrites isp. and Zoophycos isp., which are

characteristic of slope and apron settings (Frey and Pemberton, 1984). Paleontological studies reveal bathymetric

mixing of littoral, neritic, and bathyal species of foraminifers indicating downslope transport and deposition at

minimum water depths of approximately 2000 m.

These findings indicate that the area corresponding to the present Coastal Cordillera and Intermediate

Depression between Temuco and Puerto Montt was subjected to major subsidence and marine transgression

during the Miocene, followed by uplift and emergence, probably during the Pliocene. This matches the results

we attained from our work on the coastal region at the same latitudes (Finger et al., 2007; Encinas et al., 2008),

and thereby expands the paleogeography of the region known to have been affected by the Neogene events

previously described. Major subsidence is attributed to an important event of tectonic erosion that took place

along the Chilean margin during the Neogene (Encinas et al., 2008).

Acknowledgments. This research was supported by Proyecto Fondecyt No. 3060051 of Conicyt and the IRD (Institut de Recherche pour le Développment) to which gratefully thank their financial help. LB was supported by a NSERC Discovery

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 195-198

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Grant 311726-05. We also thank ENAP (The National Petroleum Chilean Company) for kindly allowing us to study their borehole microfossils. References Cecioni, G. 1978. Petroleum possibilities of the Darwin's Navidad Formation near Santiago, Chile. Publicación Ocasional

del Museo Nacional de Historia Natural de Chile 25, 3-28. Cecioni, G. 1980. Darwin´s Navidad embayment, Santiago region, Chile, as a model of the southeastern Pacific shelf.

Journal of Petroleum Geology 2-3, 309-321. Covacevich, V., Frassinetti, D., & Alfaro, G. 1992. Paleontología y condiciones de depositación del Mioceno marino en las

nacientes del río Futa, Valdivia, Chile. Boletín del Museo Nacional de Historia Natural de Chile 43, 143-154. Encinas, A., Finger, K., Nielsen, S., Lavenu, A., Buatois, L., Peterson, D & Le Roux, J.P. 2008. Rapid and major coastal

subsidence during the late Miocene in south-central Chile. Journal of South American Earth Sciences 25, 157-175. Finger, K.L., Nielsen, S.N., DeVries, T.J., Encinas, A. & Peterson, D.E. 2007. Paleontologic evidence for sedimentary

displacement in Neogene forearc basins of central Chile. Palaios 22, 3-16. Frey, R.W. & Pemberton, S. 1984, “Trace fossils Facies Models”. In Walker, R.G., (ed.): Facies Models. Geoscience Canada

Reprint Series: 189-207. Le Roux, J.P. & Elgueta, S. 2000. Sedimentologic development of a late Oligocene-Miocene forearc embayment, Valdivia

complex, southern Chile. Sedimentary Geology 130, 27-44. Marchant, M. & Pineda, V. 1988. Determinación de la edad del miembro superior marino de los estratos de Pupunahue,

mediante foraminíferos. V Congreso Geológico Chileno (Actas), Tomo 2, p. C311-C325, Santiago de Chile. Marchant, M. 1990. Foraminíferos Miocénicos de los Estratos de Pupunahue (Provincia de Valdivia: X Región):

Determinación de la edad probable y paleoambiente. Segundo Simposio sobre el Terciario de Chile, p. 177-188, Concepción, Chile.

Martínez-Pardo, R. & Pino, M. 1979. Edad, paleoecología y sedimentología del Mioceno marino de la Cuesta Santo Domingo, Provincia de Valdivia, X Región. II Congreso Geológico Chileno (Actas), p. H 103-H 124, Arica, Chile.

Morkhoven, F.P.C.M. van, Berggren, W.A. & Edwards, A.S. 1986. Cenozoic cosmopolitan deep-water benthic foraminifera. Bulletin des Centres de Recherches Exploration-Production Elf-Aquitaine, Memoire 11, 421 p.

Osorio, R. & Elgueta, S. 1990. Evolución paleobatimétrica de la Cuenca Labranza documentada por Foraminíferos. II Simposio sobre el Terciario de Chile, p. 225-233, Concepción, Chile.

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Kinematics evolution of the Camisea Subandean thrust belt from apatite fission-track thermochronology, Peru

N. Espurt1, J. Barbarand

2, S. Brusset

3, P. Baby

3, M. Roddaz

3, & W. Hermoza

4

1 IFP, 1 et 4 Av. de Bois-Préau, 92852 Rueil-Malmaison cedex, France (nicolas.espurt@ifp.fr)

2 Univ Paris Sud, UMR CNRS 8148 IDES, Bâtiment 504, Orsay cedex, F-91405, France

3 LMTG, Université de Toulouse, CNRS, IRD, OMP, 14 av. E. Belin, F-31400 Toulouse, France

4 REPSOL-YPF, Paseo de la Castellana 280, 1ª Pl., 28046 Madrid, Spain

KEYWORDS : apatite fission-track thermochronology, balanced cross section, thrust tectonic, Camisea basin, Peru

Introduction

The propagation of the deformation into a foreland basin system is controlled by the advance of the adjacent

thrust wedge (DeCelles and Giles, 1996). The Peruvian Subandean zone is considered as a foreland thrust belt

which started to develop since the Middle Miocene (Hermoza et al., 2005) and is still active today (Dorbath,

1996). The Camisea basin (Fig. 1a) belongs to the southern edge of the Ucayali basin and includes the giant

Camisea gas/condensate province (Chung et al., 2006). The structural architecture of the Camisea basin has been

the site of several studies (Bellido, 1969; Dumont et al., 1991; Mathalone and Montoya, 1995; Shaw et al., 1999;

Espurt et al., 2008) but the temporal evolution of the deformation remains poorly constrained. In this study, we

have used apatite fission-track data to constrain the exhumation and the structural evolution of the Camisea

Subandean zone.

Figure 1 : (a) Geological map of the Camisea basin and location of the Pongo de Mainique. (b) Balanced cross section of the Camisea basin and (b) restoration (modified from Gil, 2002). The restoration shows a shortening of 56 km. PMBT: Pongo de Mainique back-thrust.

Structural setting, strategy sampling and results

The frontal structural architecture of the Camisea basin (Fig. 1b) corresponds to thrust related long wavelength

anticlines (Espurt et al., 2008) which branch to the décollement developed at the base of the Devonian shales

(Gil, 2002). The Pongo de Mainique back-thrust and frontal structures accommodate the shortening of an

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 199-202

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internal triangle zone. The floor thrust of this duplex is located at the Ordovician–basement interface while its

roof thrust corresponds to the shales of the base of the Devonian. Behind the duplex, syn-tectonic Pliocene

sediments progressively sealed the recent deformation of the Shihuayro syncline. The restoration of the balanced

cross section shows a total shortening of 56 km (Fig. 1c).

Apatite fission-track thermochronology is commonly used to determine the timing of rock uplift and

magnitude of the cooling from shallow crustal levels (Fitzgerald et al., 1995). We collected 8 fine- to coarse-

grained sandstone samples for apatite fission-track analyses (Fig. 2) along vertical in different stratigraphic

levels of pre-foreland strata of the Pongo de Mainique canyon at the water level of the Urubamba River.

Considering that the strata were horizontal at deposition, apatite fission-track samples cover a thickness of

~2940 m from the Late Carboniferous to Paleogene rocks. The sample profile covers the largest paleo-depth

range of the hanging wall of the Pongo de Mainique back-thrust and gives information in terms of exhumation of

the southern limb of the Shihuayro syncline.

Apparent cooling ages of the 8 apatite fission-track samples spread between 128±12 Ma and 5.4+2.0/-1.5 Ma (Fig.

2). The confined track lengths range between 13.05 and 10.65 μm with standard deviations between 1.05 and

2.59 μm. The three stratigraphically deepest samples have been reset. They have recorded an apparent tectonic

cooling at ~6 Ma in response to the Pongo de Mainique back-thrust activity. In contrast, the four shallowest

samples have been partially reset in the fossil partial annealing zone as confirmed by the dispersion of the

component ages of these samples (between 128 and 11.6 Ma) (Fig. 2). Thrust stacking of the internal triangle

zone generated topography and induced tectonic burial. The restoration of the southern flank of the Shihuayro

syncline (Fig. 2a) shows that the paleo-temperatures of the Paleozoic to Miocene sedimentary series of the

Pongo de Mainique have been maximal during the Neogene burial. The Neogene burial history is related to the

reset of the apatite fission-tracks ages of the three deepest samples of the profile (Fig. 2a).

Kinematics evolution of the Pongo de Mainique back-thrust

The restoration of the Pongo de Mainique back-thrust, from initial and final states, coupled with apatite

fission-track data allows us to reconstruct the two dimensional thermal history of the southern limb of the

Shihuayro syncline. Isochronous apatite fission-track ages of the three deepest samples indicate that these

samples have simultaneously crossed the 110°C closure isotherm while the upper samples were maintained in

the fossil partial annealing zone. Subsequently, we propose a geometrical model where a south-verging fault

thrust with an angle value of ~50° northwards induced rotation and flattening of the deepest samples (Fig. 2b).

The high dip angle of the thrust fault is probably related to the development of the internal triangle zone. In this

structural model, we supposed that the isotherms have not been deformed during thrusting, which implies that

the erosion rate is the same as the one the thrusting rate. Therefore, a northward minimum horizontal

displacement of ~3 km would be essentially accommodated by the Pongo de Mainique back-thrust to allow the

three deeper samples to cross the 110°C closure isotherm. This shortening has plausibly archived by the first

Miocene conglomeratic sequences of the Shihuayro syncline (Fig. 2b). However, the initiation of the growth

strata deposits of the Timpia formation in the axis of the Shihuayro syncline (Fig. 2c) attests of strong

morphological modifications of the depositional architecture related to thrust imbricate growth of the triangle

zone.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 199-202

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Figure 2 : Sequential restoration of the Pongo de Mainique back-thrust. Apatite fission-track samples are located by red squares and ages in Ma are shown. PAZ: Partial annealing zone.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 199-202

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Shortening rate and consequences for growth of the northern Altiplano Plateau

The ~3 km of shortening necessary to produce the exhumation of the Pongo de Mainique back-thrust is greatly

inferior to the total shortening of 56 km calculated on the whole of the Camisea basin cross section. If we

consider that (1) this displacement can be neglected vis-à-vis of the total shortening and (2) the deformation in

the Camisea basin began with the initiation of the Pongo de Mainique back-thrust at ~6 Ma, we obtain a mean

shortening rate of about 8.8 mm/an for the Camisea Subandean zone. This shortening rate is close to the present-

day shortening rate (9 mm/an) from the southern Peruvian zone calculated by Bevis et al. (2001) from GPS data.

In addition, the recent work of Garzione et al. (2006) shows that the growth and the eastward propagation of the

Subandean zone would be related to the exhumation of the Altiplano Plateau between 10.3 and 6.8±0.4 Ma.

Consequently, our results from the Camisea Subandean zone located on the northern edge of the Altiplano

Plateau may permit to constrain its exhumation at ~6 Ma.

References Bellido, B.E. 1969. Sinopsis de la geología del Perú. Servicio de Geología y Minería Boletín 22: 54 p. Bevis, M., Kendrick, E., Samlley Jr, R., Brooks, B., Allmendigger, R., & Isacks, B. 2001. On the strength of interplate

coupling and the rate of back arc convergence in the central Andes: An analysis of the interseismic velocity field, Geochemistry, Geophysics, Geosystems: 2.

Chung, J., Arteaga, M., Davis, S. & Seminario, F. 2006. Impacto de la sismica 3D en el desarrollo de los yacimientos de Camisea. Bloque 88 – Cuenca Ucayali – Peru. Bol. Soc. Geol.. Peru 101: 73-89.

DeCelles, P. G. & Giles, K. A. 1996. Foreland basin systems. Basin Research 8: 105-123. Dumont, J.F., Deza, E. & Garcia, F. 1991. Morphostructural provinces and neotectonics in the Amazonian lowlands of Peru.

J. South Am. Earth Sci. 4: 373-381. Dorbath, C. 1996. Velocity structure of the Andes of central Peru from locally recorded earthquakes. Geophys. Res. Lett. 23:

205-208. Espurt, N., Brusset S., Baby P., Hermoza W., Roddaz M., Bolaños R., Uyen D., & Déramond J. 2008. Paleozoic structural

controls on transfer of Subandean shortening in a foreland thrust system, Ene and southern Ucayali basins, Peru. Tectonics: doi:10.1029/2007TC002238, in press.

Fitzgerald, P.G., Sorkhabi, R.B., Redfield, T.F., & Stump, E. 1995. Uplift and denudation of the central Alaska Range: A case study in the use of apatite fission-track thermochronology to determine absolute uplift parameters, Journal of Geophysical Research 100: 20175-20191.

Garzione, C.N., Molnar, P., Libarkin J.C., & MacFadden B.J. 2006. Rapid late Miocene rise of the Bolivian Altiplano: Evidence for removal of mantle lithosphere. Earth and Planetary Science Letters 241: 543-556.

Gil, R.W. 2002. Evolución lateral de la deformación de un frente orogénico: ejemplo de las cuencas subandinas entre 0° y 16° S. Sociedad Geológica del Perú, Publicación especial 4, 146 p.

Mathalone, J.M.P., & Montoya R.M. 1995. Petroleum geology of sub-Andean basins of Peru, in Tankard, A., Soruco, R. S., and Welsink, eds., Petroleum basins of South America, AAPG Memoir 62: 423-444.

Hermoza, W., Brusset, S., Baby, P., Gil, W., Roddaz, M., Guerrero, N. & Bolaños, R. 2005. The Huallaga foreland basin evolution: Thrust propagation in a deltaic environment, northern Peruvian Andes. J. South Am. Earth Sci. 19: 21-34.

Shaw, J.H., Bilotti, F. & Brennan P.A. 1999. Patterns of imbricate thrusting, Geological Society of American Bulletin 111: 1140-1154.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 203-205

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Tectonic and glacial forcing of motion along an active detachment fault

Daniel L. Farber1,2

& Gregory S. Hancock3

1 Atmospheric and Earth Sciences Division, University of California, Lawrence Livermore National Laboratory,

Livermore, California, U.S.A. (farber2@llnl.gov) 2 Earth Sciences Department, University of California, Santa Cruz, Santa Cruz, California, U.S.A.

(dfarber@ucsc.edu) 3 Department of Geology, College of William and Mary, Williamsburg, VA 23187, U.S.A. (gshanc@wm.edu)

KEYWORDS : cosmogenic, uplift, detachment faulting, climate, isostasy

The Cordillera Blanca of central Peru form the highest topography of Peru and contain the tallest Andean peak

north of 32° S. The range is bounded on the west by the spectacular ~200 km long Cordillera Blanca

Detachment Fault (CBDF) and widespread tectonic indicators for extension in this portion of the Andes have

lead a number of authors to cite the CBDF as a type example of gravitational collapse of high

topography(Dalmayrac and Molnar, 1981; Richardson and Coblentz, 1994). Exposures of 5-8 Ma granodiorite

with ~4 Ma apatite fission track ages at ~7000 m in the footwall of the CBDF, document extremely rapid

exhumation of the range (McNulty and Farber, 2002; Perry and Garver, 2004). This exhumation has been

interpreted as rapid Miocene – Pliocene surface uplift, leading to the apparent paradox of extensional lowering

of the topography with significant topographic uplift. In order to understand the nature of the CBDF we have

determined the rates of Quaternary motion along the fault and modeled the isostatic effects of the dissection of

the footwall.

The broad regions of accordant summits at 4200 to 4300 m.a.s.l. extending for ~100 km east of the northern

Cordillera Blanca are the remnants of the now partially dissected Puna plateau(Cobbing et al., 1981; Wilson et

al., 1995). To the west across the Callejon de Huaylas, the crest of the Cordillera Negra is a partly-dissected

northward extension of the Puna plateau(McLaughlin, 1924) (at ~4400 to 5000 m.a.s.) that truncates Tertiary

volcanic and folded Mesozoic sedimentary rocks. Along ~150 km of north-south transects, the topography of the

Cordillera Negra summits and the high elevations of the plateau east of the Cordillera Blanca are nearly identical

(Fig. 2) suggesting that the present topography represents a formerly continuous northward extension of the

Puna plateau.

To calculate slip rates along the CBDF, we have measured topographic profiles and the ages of offset

moraines and tectonically-generated fluvial terraces(Bierman et al., 1995; Van der Woerd et al., 2000) at four

locations along the CBDF, from north to south: Huaytapallana, Cojup, Queroccocha and Tuco. The vertical slip

rates decrease monotonically from north to south, and are 5.1±0.8 mm/yr at Huaytapallana, 2.9±0.3 mm/yr

Cojup, 0.77±0.1 mm/yr at Queroccocha, and 0.59±0.2 mm/yr at Tuco. Both the maximum elevation and the

relief of the Cordillera Blanca are strongly correlated with offset rates along the CBDF. Maximum elevations

decrease from ~6700 m (Huaytapallana) to ~6200 m (Cojup) and ~5600 m (Queroccocha), and maximum relief

decreases from ~2900 m to ~1000 m, reflecting the diminishing depth of glacial erosion in the south. Assuming

the initial topography of the Puna surface was at an elevation of ~4500 m prior to onset of the CBDF motion,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 203-205

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these maximum elevations provide minimum estimates of footwall uplift of ~2000 m, ~1600 m, and ~1000 m at

Huaytapallana, Cojup, and Queroccocha. This is however, a lower bound, as an unknown amount of material

may have been exhumed from on top of the present peaks since the initiation of fault motion.

In the northern and central part of the Callejon de Huaylas, a distinct, gently rolling paleovalley floor incised

by the Rio Santa is preserved on the hanging wall of the CBDF. The modern slope of this paleovalley parallels

that of the Rio Santa (Fig. 2) suggesting little downward to the north tilting of the paleovalley deposits and,

therefore, of the hanging wall of the CBDF. Thus, the concordance of the thermochronologic data taken from the

footwall block together with our Quaternary offset rates and the correlation between the uplift rates and range

geomorphology, suggests that relative to sea level, motion along the CBDF is largely confined to the footwall.

To quantify the relative contributions of tectonics and erosional unloading to generation of the topography, we

have calculated the isostatic component of the uplift by estimating the mass removed and the resulting flexural-

isostatic response since onset of motion along CBDF. To do so, we have used the continuity of the Puna plateau

across the region now occupied by the Cordillera Blanca as an initial condition of the topography prior to onset

of motion along the CBDF. The calculated flexurally-driven uplift patterns predict the present elevation of much

of the plateau remnants, with the exception being the crest of the Cordillera Blanca. There, peaks extend to

elevations well in excess of that predicted by the model thus requiring a significant additional tectonic forcing.

The localization of the excess topography along the CBDF, together with the correlation of the excess

topography with the relief and slip rates along the CBDF suggests that much of this uplift is accommodated

through tectonically driven footwall motion. Our calculations imply that tectonically driven extensional footwall

uplift along the CBDF is substantial, generating 60% to 70% of the total uplift since fault initiation. However, at

least a portion of the more rapid fault motion in the north is plausibly generated by extensive glacial erosion

allowed by the higher topography. Indeed, this growing topography likely facilitated the growth of larger and

more erosive glaciers, accelerating the rate of footwall uplift in the north by isostatically-driven footwall uplift

superposed on the tectonically-driven fault motion.

The style of deformation we document along the CBDF has important tectonic implications. The increase in

mean elevation (relative to the initial plateau topography) requires a mechanism other than erosional unloading

or extension and crustal thinning accommodated by the CBDF, both of which would produce an overall lowering

of the mean topography. Thus, the previously proposed(Dalmayrac and Molnar, 1981; Richardson and Coblentz,

1994) models calling on gravitational collapse of high topography to account for the presence of extension in

this portion of the high Andes cannot explain our observations. We consider the most likely explanation to be

additions of material at the base of the lithospheric section. The location of the Cordillera Blanca directly above

the Peruvian flat slab section suggests that this may indeed be the case. In central Peru, the onset of flat slab

subduction since ~5 Ma likely accommodated replacement of dense lithospheric material beneath the Cordillera

Blanca with the buoyant oceanic slab, increasing the thickness of the crustal section below central Peru(Gutscher

et al., 2000). In contrast to the Altiplano section studied by Ghosh et al. (Ghosh et al., 2006) where uplift was

largely complete by 5 Ma, in this portion of the Andes, topography has not yet reached steady state and is likely

still increasing today.

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References Bierman, P.R., Gillespie, A.R., and Caffee, M.W., 1995, Cosmogenic Ages For Earthquake Recurrence Intervals And Debris

Flow Fan Deposition, Owens-Valley, California: Science, v. 270, p. 447-450. Cobbing, E., Pitcher, W., Wilson, J., Baldock, J., Taylor, W., McCourt, W., and Snelling, N., 1981, The geology of the

Western Cordillera of northern Peru: London. Dalmayrac, B., and Molnar, P., 1981, Parallel Thrust And Normal Faulting In Peru And Constraints On The State Of Stress:

Earth And Planetary Science Letters, v. 55, p. 473-481. Ghosh, P., Garzione, C.N., and Eiler, J.M., 2006, Rapid uplift of the Altiplano revealed through C-13-O-18 bonds in paleosol

carbonates: Science, v. 311, p. 511-515. Gutscher, M.A., Spakman, W., Bijwaard, H., and Engdahl, E.R., 2000, Geodynamics of flat subduction: Seismicity and

tomographic constraints from the Andean margin: Tectonics, v. 19, p. 814-833. McLaughlin, D.H., 1924, Geology and physiography of the Peruvian Cordillera, Departments of Junin and Lima: Bulletin of

the Geological Society of America, v. 35, p. 591-632. McNulty, B., and Farber, D., 2002, Active detachment faulting above the Peruvian flat slab: Geology, v. 30, p. 567-570. Perry, S.E., and Garver, J.I., 2004, Onset of tectonic exhumation of the Cordillera Blanca, northern Peru based on fission-

track and U+Th/He dating of zircon: Abstracts with Programs - Geological Society of America, v. 36. Richardson, R.M., and Coblentz, D.D., 1994, Stress Modeling In The Andes - Constraints On The South-American Intraplate

Stress Magnitudes: Journal Of Geophysical Research-Solid Earth, v. 99, p. 22015-22025. Van der Woerd, J., Ryerson, F.J., Tapponnier, P., Meriaux, A.S., Gaudemer, Y., Meyer, B., Finkel, R.C., Caffee, M.W.,

Zhao, G.G., and Xu, Z.Q., 2000, Uniform Slip-Rate along the Kunlun Fault: Implications for seismic behaviour and large-scale tectonics: Geophysical Research Letters, v. 27, p. 2353-2356.

Wilson, J., Reyes, L., and Garayar, J., 1995, Geologia de quadrangulos de Pallasca, Tayabamba, Corongo, Pomabamba, Carhuaz, and Huari: Lima, Insituto Geologico and Minero y Metalurgico.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 206-209

206

No subsidence in the development of the Central Depression along the Chilean margin

Marcelo Farías

1,2, Sébastien Carretier

3, Reynaldo Charrier

1, Joseph Martinod

3, Andrés

Tassara2, Alfonso Encinas

4, & Diana Comte

2

1 Departamento de Geología, Univ. de Chile, Plaza Ercilla # 803, Santiago, Chile (mfarias@dgf.uchile.cl,

rcharrie@cec.uchile.cl) 2 Departamento de Geofísica, Univ. de Chile, Blanco Encalada #2002, Santiago, Chile (andres@dgf.uchile.cl)

3 LMTG. IRD, CNRS, Univ. Toulouse, 14, av. Edouard Belin, Toulouse, France (carretier@lmtg.obs-mip.fr,

martinod@lmtg.obs-mip.fr). 4 Departamento de Ciencias de la Tierra, Univ. de Concepción, Barrio Universitario s/n, Concepción, Chile

(aencinas@udec.cl)

KEYWORDS : Central Depression, subsidence, forearc tectonics, subduction effects, Chile

Introduction

Longitudinal valleys enclosed by a coastal range and a fold-and-thrust belt are common features in convergent

margin. Particularly, the Chilean forearc is characterized by the Central Depression located between the main

belt and the Coastal Cordillera. Pioneer works on the Chilean geology (e.g., Brüggen, 1950; Carter & Aguirre,

1965) proposed that this valley would be a graben, i.e., the result of subsidence; this is consistent with physical

modeling on subduction zones, which suggests that forearc subsidence is a consequence of slab pull effects (e.g.,

Hassani et al., 1997; Gerbault et al., 2005).

In this contribution, we examine different aspects on the late Cenozoic evolution of the Chilean forearc. As we

will put forward, there is no evidence for a subsidence origin of the Central Depression. On the contrary, we will

show that either differential erosion because of surface uplift (central-south Chile) or stationary topography

(north Chile) occurs.

Northern Chile

The northern Chile forearc consists of 4 morphostructural units: Coastal Cordillera, Central Depression,

Precordillera and Western Cordillera. In spite of pioneer geological works that suggested a “Basin and Range”-

like model for this region, the E-flank of the Central Depression is bounded by the West-vergent Thrust System

(WTS), which has accommodated 2-3 km of surface uplift since 30 Ma (Victor et al., 2004; Farías et al., 2005a).

This uplift is well recorded by the pre-30 Ma Choja Pediplain (Galli, 1967), which appears beneath most of the

Oligoce-Neogene cover and whose relicts can be observed on the summits of the Coastal Cordillera (Tosdal et

al., 1984; Dunai et al., 2005) and beneath the Altos de Pica Formation along the Precordillera (Farías et al.,

2005a; Muñoz et al., this symposium) and beneath the Central Depression in seismic images (see images in

Victor et al., 2004). These images, as well as drill core data (Mordojovich, 1965), show that this buried pediplain

is very flat (excepting for some inselbergs) and located at 100-300 m a.s.l.

The development of flat and extended erosive surfaces requires a minimal slope to inhibit incision. Therefore,

the present-day elevation of the Choja Pediplain beneath the Central Depression proves that no subsidence has

been accumulated since the Oligocene in the zone, and that its development is related to differential uplift of the

Coastal and the Western Cordillera. Hence, the Central Depression has remained tectonically stationary.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 206-209

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Central Chile

Similarly as in northern Chile, most of the works have considered the Central Depression as a graben in which

the Coastal Cordillera and the Principal Cordillera are its footwalls (e.g., Brüggen, 1950). However, neither flank

is a normal fault. The W-flank of the Principal Cordillera corresponds to the San Ramon-Pocuro thrust and there

is no significant faulting in the limit between the Coastal Cordillera and Central Depression. On the other hand,

the disposition of the Mesozoic sequences in the Coastal Cordillera (E-dipping) suggests a priori that tilting of

the Coastal Cordillera could have produced the subsidence necessary to from the Central Depression (Farías et

al., 2005b). However, junctions between the Coastal and Principal Cordillera, and inselberg on the Central

Depression do not support this model. In fact, the presence of Miocene peneplains on the summits of the Coastal

Cordillera extending to the Principal Cordillera (but being about 1 km offset by the San Ramon-Pocuro fault,

Fig. 1) shows that this zone underwent a major surface uplift of about 2 km during the late Miocene (Farías et

al., 2008). Therefore, the only mechanism capable to decrease the elevation of the Central Depression is the

erosion. That implies (1) differential erosion and (2) the development of a longitudinal drainage. Differential

erosion would be the result of: (1) knickpoint migration was retained by hard lithologies (Late Cretaceous

granitic bodies), and (2) fast knickpoint retreat in zones in which these rocks are absent, which allowed the

capture of the headwaters of the retained rivers (Farías et al., 2008). It is widely recognized that granitic rocks

can resist several orders of magnitude more the erosion that other rocks (e.g., Stock and Montgomery, 1999).

The morphologic correlation between both cordilleras shows a decrease in elevation along the axis of the

Central Depression (Fig. 1). That could be the sole evidence of subsidence. Anyway, the maximum relative

subsidence would only reach ~400 at 33º30’S and <200 m at 34º20’S, and the rock uplift would be more than 1

km and 0.5 km, respectively. This probable relative subsidence would have favored the development of a

longitudinal drainage, which has been the most important process during Central Depression development in this

region.

Southern Chile

According to Encinas et al. (this symposium), between Temuco and Puerto Montt, several outcrops containing

Neogene marine sediments related to bathyal facies evidence that uplift occurred in the Pliocene in both Central

Depression and Coastal Cordillera after a major Miocene subsidence. Near Temuco, these deposits are hundred

of meter above the sea level, thus eustatic variation cannot explain their present elevation. Moreover, at ~38ºS,

the Coastal Cordillera joins the Principal Cordillera through some hill belts having flat summits as high as 700 m

a.s.l. (relict uplifted peneplains). Although it is likely local uplift occurring there, the bedrock corresponds to a

Paleozoic granitic belt; as previously mentioned, this rocks can resist more than 1-2 orders of magnitude the

erosion, thus they are relicts of the decrease in elevation given by differential erosion in response to uplift.

The “Norte Chico”

Between Vallenar (28º30’S) and Santiago (33ºS), coinciding with the flat-slab region, the literature shows that

there is no Central Depression. However, immediately W of the cordilleran front, main rivers joins with

longitudinal drainages, which diminish the maximum elevation in WE transects (Fig. 1). The morphology

resulted of these longitudinal valleys can be considered as “proto-Central Depression” in the sense of what is

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 206-209

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observed further south. On the other hand, some rivers abandoned their sea outlet because of Plio-Pleistocene

coastal uplift (as occurs further north). Both situations suggest that this region is a transitional between the North

and South types of Central Depression. Here, uplift has been larger than S-ward as shown by the decrease in

peneplain elevations. Likewise, this zone present two standing out features: (1) north of 33ºS precipitation rates

largely decreases and (2) the bedrock is predominantly granitic, thus erosion has been largely resisted by these

rocks and by climate than that observed southward. Therefore, climate, lithology and tectonics would have

favored the preservation of high peneplains and resisted the development of an erosive Central Depression.

Discussion

We exposed evidence for a no-subsidence

origin of the Central Depression along the

Chilean forearc, at least since the Pliocene. It is

likely that in southern Chile subsidence would

have occurred in the Miocene (Encinas et al.,

this symposium). However, subsidence has been

the major cause invoked to explain the Central

Depression until now. Moreover, subsidence is

predicted by numerical and analogical modeling.

How is this possible in one of the most long and

continuous subduction region on the world?

Among others possibilities, two reasons can

explain the lack of continental forearc

subsidence in the late Cenozoic. The first one

consists in the very high rigidity of the forearc

lithosphere (Tassara, 2005). Thus subsidence

due to slab-pull effects should have a very great

wavelength (>100 km) and a low amplitude.

According to Billen & Gurnis (2001), slab pull

would be favored by a low viscosity mantle

wedge, which only occurs beneath the main

mountain belt (Tassara, 2005). Thus, rheology

does not allow the development of forearc

subsidence as mentioned by modeling. The

second possibility consists in the particularities

of plate convergence. According to Heuret &

Lallemand (2005), slab pull is minor in young slabs (as the Nazca plate) and in fast continents (Nazca and South

American plates have had similar velocities in the late Cenozoic). Gripp & Gordon (2002) proposed that E-ward

motion of slabs would favor minor subduction angle (as actually occurs), thus slab pull would be minor (because

of slab anchoring; Heuret & Lallemand, 2005).

Fig. 1. Swath profiles showing maximum, minimum, and mean elevations along each transect. (a, b): “Norte Chico” transects. (c, d): Central Chile transects. (e): Southern Chile transect.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 206-209

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To sum up, geological and morphological data evidence no-subsidence in the development of the Central

Depression along the Chilean forearc, in contradiction with numerical and analogical modeling. This would be a

consequence of the particularities of the Chilean subduction margin given by (1) the high rigidity of the forearc

lithosphere and (2) fast trenchward velocity of the continent, young slab, and low subduction angle that inhibit

the slab pull effects. We propose that our conclusions could contribute to constrain better subduction modeling.

Acknowledgements This work has been financed by the Anillo ACT-18 Project, the Institut de Recherche pour le Développement (IRD), and INSU GRANT “Reliefs de la Terre. Impact du climat sur la dynamique du relief des Andes: quantification et modélisation”.

References Carter, W. & Aguirre, L., 1965. Structural geology of Aconcagua Province and its relationship to the Central Valley graben.

Geol. Soc. Am. Bull., 76: 651-664. Billen, M.I., & Gurnis, M., 2001. A low viscosity wedge in subduction zones. EPSL, 193: 227-236. Dunai, T.J., Lopez, G.A.G. & Juez-Larre, J., 2005. Oligocene-Miocene age of aridity in the Atacama Desert revealed by

exposure dating of erosion-sensitive landforms. Geology, 33: 321-324. Encinas, A., Finger, K.L., & Buatos, L., this symposium. Sedimentologic, paleontologic, and ichnologic evidence for deep-

marine Miocene deposition in the present Intermediate Depression of south-central Chile (38º30’-41º30’S). Tectonic implications.

Farías, M., Charrier, R., Comte, D., Martinod, J. & Hérail, G., 2005a. Late Cenozoic deformation and uplift of the western flank of the Altiplano: Evidence from the depositional, tectonic, and geomorphologic evolution and shallow seismic activity (northern Chile at 19º30'S). Tectonics, 24: TC4001.

Farías, M., Charrier, R., Fock, A., Campbell, D., Martinod, J., & Comte, D., 2005b. Rapid late Cenozoic surface uplift of the central Chile Andes (33º-35ºS), 6th ISAG. IRD, Barcelona.

Farías, M., Charrier, R., Carretier, S., Martinod, J., Fock, A., Campbell, D., Cáceres, J., & Comte, D., 2008. Late Miocene high and rapid surface uplift and its erosional response in the Andes of Central Chile (33º-35ºS). Tectonics, 27: TC1005.

Galli, C., 1967. Pediplain in northern Chile and the Andean uplift. Science, 158: 653-655. Gerbault, M., 2005. Normal faulting in a forearc submitted to slab pull: Numerical models and insight on the structure of the

Chilean forearc, 5th ISAG. IRD, Barcelona. Gripp, A.E. & Gordon, R.G., 2002. Young tracks of hotspots and current plate velocities. Geophys. J. Int., 150: 321-361. Hassani, R., Jongmans, D. & Chéry, J., 1997. Study of plate deformation and stress in subduction processes using two-

dimensional numerical models. J. Geophys. Res., 102: 17952-17964. Heuret, A. & Lallemand, S., 2005. Plate motions, slab dynamics and back-arc deformation. Physics of the Earth and

Planetary Interiors, 149: 31-51. Mordojovich, C., 1965. Reseña sobre las exploraciones de la ENAP en la zona norte, años1956 a 1962. Minerales, 20: 30pp. Muñoz, V., Hérail, G., & Farías, M., this symposium. Origin and age of a topographic highs into the Tarapacá Pediplain. Stock, J.D. & Montgomery, D.R., 1999. Geologic constraints on bedrock river incision using the stream power law. J.

Geophys. Res., 104: 4983-4993. Tassara, A., 2005. Interaction between the Nazca and South American plates and formation of the Altiplano-Puna plateau:

Review of a flexural analysis along the Andean margin (15º-34ºS). Tectonophysics, 399: 39-57. Tosdal, R.M., Clark, A.H. & Farrar, E., 1984. Cenozoic polyphase landscape and tectonic evolution of the Cordillera

Occidental, southernmost Peru. Geol. Soc. Am. Bull., 95: 1318-133 Victor, P., Oncken, O. & Glodny, J., 2004. Uplift of the western Altiplano plateau: Evidence from the Precordillera between

20° and 21°S (northern Chile). Tectonics, 23: TC4004, doi:10.1029/2003TC001519.

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Southern Andean (34º-46ºS) tectonic evolution through the inception of Cretaceous to Neogene shallow subduction zones: A south to north trend?

Andrés Folguera1 & Víctor A. Ramos

1

1 Laboratorio de Tectónica Andina, Universidad de Buenos Aires, and CONICET, Buenos Aires, Argentina

(andresfolguera2@yahoo.com.ar, andes@gl.fcen.uba.ar)

KEYWORDS : shallow subduction, Southern Andes, orogenesis, orogenic collapse

Introduction

Andean topography has been constructed by stacking of crustal sheets and was as well as destroyed by cycles

of orogenic collapses several times through its evolution. Both processes are relevant in understanding final

structure of the Andean edifice. Shallow subduction processes through the Andes have acquired an important

relevance during the last years in explaining sudden phases of orogenic construction temporally related to arc

expansions, followed by crustal collapse and eruption of intraplate volcanic provinces during arc retreat periods

(James and Sacks, 1999; Kay et al., 2006). In intervals no longer than 10 to 20 million years, the Andean crust

absorbs critical amounts of shortening that pushes low density materials to depth in the search of isostatic

compensation, becoming eventually eclogitized and therefore unstable. Increase in sublithospheric thermal

gradient, derived from broadening of the asthenospheric wedge after slab steepening, produces orogenic collapse

coupled with lower crustal looses. Modern examples have allowed to document the effects of ongoing subducted

slab shallowing and to understand the relation between shifting of the arc front and development of brittle-

ductile transitions associated with crustal stacking and therefore shortening (Ramos et al., 2002). The study of

ancient shallow subduction settings have complimented the model showing that redefinition of the arc front

nearer the trench is a fast phenomena, almost instantaneous in the geological record, and that crustal weakening

develops synchronously producing extension behind the arc (Ramos and Folguera, 2007).

Most of the past flat and shallow subduction proposals along western Gondwana have been made for Andean

orogenic times and particularly for the last 50 Ma (James and Sacks, 1999; Kay et al., 2006), with a few

exceptions for Late Paleozoic-Early Mesozoic times (Dalziel et al. 2000; Ramos and Folguera, 2007).

This study proposes that main orogenic phases in northern Patagonia and southern Central Andes were linked

to shallow subduction processes of the Farallón plate and then of the Nazca plate beneath the South American

margin, probably describing a south to north trend during the last 110 millions of years.

Past shallow subduction zones

Bernárdides shallow subduction zone (late Early Cretaceous to Late Cretaceous)

The expansion of arc-related sequences since 130 to 110 Ma between 42º and 48ºS representins an eastward

migration of more than 300 km from arc-related plutons (140-120 Ma) located at the Pacific coast to the present

eastern slope of the Andes (Pankhurst et al., 1999). Those mesosilicic sequences are separated from upper

ignimbritic packages dated at 110 Ma by an angular unconformity (Folguera and Iannizzotto, 2004). Next to the

Present Pacific coast, areas uplifted during Cretaceous times corresponding to the western part of the orogen

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 210-213

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(44º-47ºS), suffered a phase of tectonic collapse ending in the formation of the Paleogene Traiguén basin (Figure

1; Hervé et al., 1995). At the far retroarc area, 80 Ma basaltic flows with tholeiitic affinities as well as Paleogene

intraplate mafic rocks are unconformably covering Cretaceous uplifts (Pezzutti and Villar, 1979; Baker et al.,

1981).

Figure 1. Proposed shallow subduction episodes since Late Cretaceous to Neogene times affecting the Southerrn Central and Northern Patagonian Andes. These are based on arc expansions followed by arc retractions contemporaneous to orogenic collapse and intraplate magmatic emplacement. The Payenia shallow subduction segment is based on Kay et al. (2006).

These orogenic systems involve inverted Permian-Early Cretaceous extensional faults (47º-48ºS), located more

than 600 km from the trench (Homovc and Constantini, 2001); and Early to Middle Jurassic depocenters (43º-

46ºS), more than 500 km from the trench (Peroni et al., 1995). Then Early-Late Cretaceous expansion of arc-

related series was coetaneous with thick-skinned deformation at the retroarc, and then eruption of crustal melts

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 210-213

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and mantle derived materials and orogenic collapse occurred during Latest Cretaceous to Paleogene, presumably

linked to a slab retreat stage.

Somuncura and Palaoco shallow subduction zones (Late Cretaceous to Paleocene)

Eastward expansion of Late Cretaceous to Paleocene (75-50 Ma) arc-related sequences (36º-43º30´S) (Ramos

and Folguera, 2005) describes two separate areas centered at 36º-39ºS and 40º-43º30´S respectively, where arc

migration was maximum up to 400 km. Both zones correlate with Late Oligocene to Early Miocene intraplate

volcanic plateaus at the retroarc area (De Ignacio, 2001), and with the latitudinal extent of the Cura Mallín and

Ñirihuao extensional basins developed at the western part of the orogen (Figure 1). Basement uplifts, located

more than 600 km from the trench, were emplaced since 100 Ma in coincidence with the area of arc expansion.

This scenario points to a shallow episode of subduction between Latest Cretaceous and Eocene, followed by the

emplacement of Late Oligocene to Early Miocene intraplate series, and orogenic collapse, previously to

reestablishment of normal subduction in the area.

Payenia shallow subduction zone (Middle to Late Miocene)

Arc related rocks were emplaced more than 550 kilometers away from the trench in the eastern slope of the

Andes, during Late Miocene times (13-4 Ma) from 34º30´to 37º45´S (Kay et al. 2006). At the retroarc zone

basement blocks cannibalized the foreland basin in Late Miocene times associated with the Malargüe fold and

thrust belt to the west (Desanti 1956; Soria 1984; Yrigoyen 1994), whose main phase of contraction has been

constrained in 13-10 Ma (Giambiagi et al. 2007). This indicates a genetic relationship between the arc

expansion, uplift of the Andes, sedimentation in the foreland basin, and the breaking of the foreland area. This

stage changed to an extensional regime since Latest Miocene-Early Pliocene times. Extensional troughs were

developed in the area that previously recorded arc expansion until late Quaternary (Bermúdez et al. 1993;

Hildreth et al., 1999), controlling the emplacement of crustal melts and poorly differentiated mantle products

(Rossello et al. 2002; Kay et al. 2006).). Presently this area is associated with crustal attenuation as well as

anomalous sublithospheric heating inferred by teleseismic, tomographic and gravimetric analysis (Gilbert et al.

2006; Yuan et al. 2006; Folguera et al. 2007). This scenario during the last 15 Ma between 34º30´to 37º45´S

points out to a scenario in which shallow subduction (15-5 Ma) was followed by subducted slab steepening, and

related orogenic collapse.

Conclusion

Three discrete areas from 34º to 46ºS have experimented eastward expansion of arc-related rocks

diachronically since Late Cretaceous to Neogene time, describing an apparent south to north trend of episodes of

shallow subduction, where Present flat subduction zone (Pampean subduction zone) develops at their northern

end. Those areas explain fairly well main stages of orogenic development, time and setting of intraplate volcanic

associations, main pulses of orogenic collapse, erosional stage, and Present tensional regime.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 210-213

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References

Baker, P.E., Rea, W.J., Skarmeta, J., Caminos, R., Rex, D.C., 1981 – “Igneous history of the Andean Cordillera and patagonian Plateau around latitude 46°S”. Phil. Trans. R. Soc. Lond. A 303, 105-149. Gran Bretaña, Boulder, Colorado, Geological Society of America, Special Paper 265: 1-12.

Bermúdez, A., Delpino, D., Frey, F., Saal, A., 1993 – “Los basaltos de retroarco extraandinos”. In Ramos, V.A. (ed.). Geología y Recursos Naturales de Mendoza. Actas del XII° Congreso Geológico Argentino y II° Congreso de Exploración de Hidrocarburos, Relatorio I-13, Buenos Aires: 161-172,.

Dalziel, I., Lawver, L., Murphy, J., 2000 - Plumes, orogenesis, and supercontinental fragmentation. Earth and Planetary Science Letters, 178 (2000): 1-11.

Desanti, R., 1956 - Hoja Cerro Diamante.1 sheet 1: 250,000. Provincia de Mendoza. Servicio Nacional Minero Geológico, Boletín.

De Ignacio, C., López, I., Oyarzun, R., Márquez, A., 2001 - The northern Patagonia Somuncura plateau basalts: a product of slab-induced, shallow asthenospheric upwelling? Terra Nova, 13: 117-121.

Folguera, A., Iannizzotto, N., 2004 - The Lagos La Plata and Fontana fold and thrust belt. Long lived orogenesis at the edge of western Patagonia. Journal of South American Earth Sciences, 16 (7): 541-566.

Folguera, A., Introcaso, A., Giménez, M., Ruiz, F., Martínez, P., Tunstall, C., García Morabito, E., Ramos V.A., 2007 - Crustal attenuation in the Southern Andean retroarc determined from gravimetric studies (38º-39º30´S): The Lonco-Luán astenospheric anomaly. Tectonophysics, Doi: 10.1016/j.tecto.2007.04.001

Gilbert, H., Beck, S., Zandt, G., 2006 - Lithospheric and upper mantle structure of central Chile and Argentina. Geophysical Journal International, 165 (1), 383. doi: 10.1111/j.1365-246X.2006.02867.x.

Giambiagi L., Bechis, F., García, V., Clark, A., 2007 – “Temporal and spatial relationships of thick- and thin-skinned deformation: a case study from the Malargüe fold and thrust belt, Southern Central Andes”. In Sempere, T., Folguera, A., Gerbault, M. (ed.): Tectonophysics Special Issue-New insights into Andean evolution ISAG 2005. In Press.

Hervé, F., Pankhurst, R.J., Drake, R., Beck, M.E., 1995 - Pillow basalts in a mid-tertiary extensional basin adjacent to the Liquiñe Ofqui fault zone: The Isla Magdalena area, Aysén, Chile. Journal of South American Earth Science, 8 (1): 33-46.

Hildreth, W., Fierstein, J., Godoy, E., Drake, R., Singer, B., 1999 - The Puelche volcanic field: extensive Pleistocene rhyolite lava flows in the Andes of Central Chile. Revista Geológica de Chile, 26 (2): 275-309.

Homovc, J.F., Constantini, L.A., 2001 - Hydrocarbon exploration potential within intraplate shear-related depocenters, Deseado and San Julián basins, southern Argentina. American Association of Petroleum Geologists, Bulletin, 85 (10): 1795-1816.

James, D.E., Sacks, S., 1999 – “Cenozoic formation of the Central Andes: A geophysical perspective”. In Skinner B. et al. (eds.): Geology and Mineral Deposits of Central Andes-Society of Economic Geology, Special Publication, 7: 1-25.

Kay, S.M., Burns, M., Copeland, P., 2006 – “Upper Cretaceous to Holocene Magmatism over the Neuquén basin: Evidence for transient shallowing of the subduction zone under the Neuquén Andes (36°S to 38°S latitude)”. In Kay, S.M. and Ramos, V.A. (eds.). Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén basin (35º-39ºS)-Geological Society of America, Special Paper, 407: 19-60.

Pankhurst, R., Weaver, S., Hervé, F., Larrondo, P., 1999 - Mesozoic-Cenozoic evolution of the North Patagonian Batholith in Aysén, southern Chile. Journal of the Geological Society of London, 156: 673-694.

Peroni, G.O., Hegedus, A.G., Cerdan, J., Legarreta, L., Uliana, M.A. Laffitte, G., 1995 – “Hydrocarbon accumulation in an inverted segment of the Andean Foreland: San Bernardo belt, Central Patagonia”. In Tankard, A.J., Suárez, R., Welsink, H.J. (eds.). Petroleum Basins of South America-AAPG Memoir, 62: 403-419.

Pezzutti, N., Villar, L.M., 1979 – “Los complejos alcalinos en la zona de Sarmiento, provincia de Chubut”. Actas del VIIº Congreso Geológico Argentino. Neuquén, 2: 511-520. Buenos Aires.

Ramos, V., Folguera, A., 2005 – “Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc and foreland deformation”. In Veiga, G., Spalletti, L., Howell J. and Schwarz E. (eds.). The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics-Geological Society of London, Special Publication, 252: 15-35.

Ramos, V.A., Folguera, A., 2007 – “Andean flat subduction through time”. In Murphy B. (ed.). Ancient orogens and moder analogues-Geological Society of London, Special Publication. In Press.

Ramos, V.A., Cristallini, E., Pérez, D.J., 2002 - The Pampean flat-slab of the Central Andes. Journal of South American Earth Sciences, 15 (1): 59-78.

Rossello, E., Cobbold, P., Diraison, M. & Arnaud, N., 2002 - Auca Mahuida (Neuquén Basin, Argentina): a Quaternary shield volcano on a hydrocarbon-producing substrate. Vº International Symposium on Andean Geodynamics, Extended Abstracts, 549-552.

Soria, M., 1984 - Vertebrados fósiles y edad de la Formación Aisol, provincia de Mendoza. Revista de la Asociación Geológica Argentina, 38: 299-306.

Yrigoyen, M., 1994 - Revisión estratigráfica del Neógeno de las Huayquerías de Mendoza septentrional, Argentina. Ameghiniana, 31 (2): 125-138.

Yuan, X., Asch, G., Bataile, K., Bohm, M., Echtler, H., Kind, R., Onchen, O., Wölbern, I., 2006 – “Deep seismic images of the Southern Andes”. In Kay, S.M. and Ramos, V.A. (eds.). Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén basin (35º-39ºS)-Geological Society of America, Special Paper, 407: 61-72.

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Hypocentral determinations of earthquakes in a 3D heterogeneous velocity model, Ecuador and Northern Peru: Preliminary results

Yvonne Font1, Mónica Segovia

2, & Hernando Tavera

3

1 Géosciences Azur, Université Nice-Sophia Antipolis, IRD, Observatoire Côte d’Azur, Quai de la Darse, 06235

Villefranche-sur-Mer, France (font@geoazur.obs-vlfr.fr) 2 IG, EPN, Ladron de Guevara E11-2534, Quito, Ecuador (msegovia@igepn.edu.ec)

3 IGP, 169 Calle Badajoz, Lima, Peru (hjtavera@geo.igp.gob.pe)

Introduction

Great earthquakes mostly generate on plate interface of subduction zones. In Ecuador, the subduction of the

Nazca plate, that carries the 200km long Carnegie Ridge, beneath the North Andean Block triggered during the

last century, 4 major earthquakes of magnitude greater than 7.7. However, in the offshore area where these major

earthquakes occurred, hypocentral determinations based on local seismological observation are usually poorly

resolved. Consequently, the subduction thrust fault zone – producer of the most destructive earthquakes and

tsunamis - is badly imaged and seismic hazard inefficiently evaluated.

Hypocentral determination uncertainties depend mainly on 3 parameters: (1) poor azimuthal coverage, (2) the

use of a 1D-velocity model in a region that is highly heterogeneous, (3) the choice of the location technique. In

Ecuador, seismic stations are densely distributed on the volcanic chain and sparsely on the coastal area.

Furthermore their performance in time also presents disparities. Consequently, the azimuthal coverage is not

constant in time and azimuthal gap is poor for subduction earthquakes. In these conditions, it is difficult to

process seismic tomography in order to obtain a coherent regional velocity model. Earthquake locations are thus

classically determined in a 1D velocity by minimizing residual rms (Hypo71 technique).

This study aims to improve the absolute earthquake hypocenter locations by performing the location process

within a 3D velocity model and using the 3D MAXI technique.

3D velocity model

The first step of this study consists in the construction of a 3D velocity model in the Ecuadorian-North Peru

region. Even though seismic tomography cannot provide satisfactory Vp model, many local crustal studies have

been conducted in the area given sparse information on the different geological structures. We thus combined

that information to construct the first 3D model of the area.

In this construction, we consider:

1. The oceanic Nazca plate and sedimentary covertures. The subducting plate dip is constrained with local

geophysical studies near the surface (multichannel seismic reflection and wide-angle data) and in depth, using

local and teleseismic seismological information.

2. The North Andean Block margin composed of accreted oceanic plateaus. The Moho depth is approximated

using gravity modeling.

3. The metamorphic volcanic chain (oceanic for the occidental chain and the inter-andean valley, continental

for the oriental one). The velocity structure is constrained from local refraction studies.

4. The continental Guyana shield and sedimentary basins.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 214-215

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The resulting 3D velocity model covers an area from 2°N to 6.5°S and 283°E to 277°E and reaches a depth of

200km. It is discretized in constant velocity block of 12 x 12 x 3 km in x, y and z, respectively.

Data

Ecuadorian seismic network is mainly installed on the Andean chain (5 stations on the coastal area versus 30

on the Andes). This network recorded, since 1994, more than 40000 earthquakes. Until now, the total P-arrivals

set was used to locate subduction earthquakes. However, the disequilibrium of the station density between

coastal and volcanic region, associated to the travel time propagation errors, contribute to the bad performance of

the earthquake location procedure. In this preliminary work, we first selected a sub-set of seismic station

regarding their geographic location and performance in time. In a second step, we define 5 sub-geographic area

(or volume) in function of the azimuthal coverage. These geographical areas are: 1. The coast and marine

earthquakes. 2. The North-Andean margin. 3. The volcanic region (all of three are north of 2°S), 4. The Northern

Peru-Southern Ecuador region (where Ecuadorian and Peruan networks are combined), and 5. a “deep” region

(deeper than 50km).

In each sub-region, we sort the earthquake quality regarding the number of phases from each sub-set network.

We thus provide consistent set of earthquake location in terms of quality (azimuthal gap).

Method

In this study, we use the 3D MAXI technique to improve earthquake location because the technique is well

adapted to velocity model presenting strong lateral Vp heterogeneities.

The MAXI method determines, within a 3D velocity model, the absolute location of each earthquake

independently based on measurements of arrival times. The algorithm used for this study is fundamentally

different from classical determination method because it is based on the concept of Equal Difference Time

surfaces (EDT surfaces - Zhou, 1994) that are established from P-wave measurement differences at pairs of

stations. Basically, the method can be summarized into 3 steps. First, the algorithm seeks for the spatial node of

the velocity model that is crossed by the maximum number of EDT surfaces, i.e., the spatial node that better

satisfies the arrival time differences computed at all station pairs. This node is called PRED, standing for

predetermination solution. The great advantage of this search mode is that it depends neither on the origin time

estimate nor on any residual minimization. Second, thanks to the PRED characteristics, residual outliers can be

objectively detected and are cleaned out from the original dataset, without any iterative process or weighting.

Then, in a third step, a statistical minimization is conducted in a small domain around the PRED node, which

results in a unique FINAL solution.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 216-218

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Determination of effective elastic thickness of the Colombian Andes using satellite-derived gravity data with admittance technique

Remy A. Galán & Iván F. Casallas

1 Universidad Distrital Francisco José de Caldas, Cr 8 No 40 – 62, Bogotá, Colombia

(remyalexander@gmail.com, ivancasallas@gmail.com)

KEYWORDS : elastic thickness, isostasy, Colombian Andes, satellite gravity, admittance

Abstract. Gravity anomaly values derived from Global Gravity Models (calculated from the CHAMP and

GRACE satellite missions), are compared with free air terrestrial gravity data to find the best representation of the surface data. Using these values and values of topographic heights, applies the isostatic response function (admittance) to a collection of profiles, to find an average of elastic thickness for the Colombian Andes.

Gravity Data

The global models of gravity offer a uniform coverage of the study area, so it is possible to obtain the required

term in the field of gravity for a particular study. Data from terrestrial gravity not have a uniform distribution,

despite the generation of maps of gravity anomaly is possible thanks to the different spatial methods exist.

The selection of the final model that best represents the field of terrestrial gravity, takes place through the

analysis of the parameters of correlation between the map of terrestrial gravity and each of the different maps

resulting from the models. The following table shows the correlations obtained final:

According to Table concluded that the model that best represents the field of terrestrial gravity in Colombian

territory is the EIGEN-CG03C, following figure shows the map obtained with this data.

--- GGM02C TEG4 EIGEN-CG03C

Terrestrial 0.57244 0.58255 0.58921

Table 1. Correlation between terrestrial gravity and GGM gravity maps.

Figure 1. EIGEN-CG03C Free air gravity map. The Models EIGEN (from 01 to 04) are derived from the observations CHAMP and GRACE and presented two versions, accompanied by a S only has the satellite component (n = 120) and the accompanying a C has two components (Satellite and Terrestrial), the latter includes the same terrestrial data contained in the model EGM96 and presents until n = 360; these models are maintained and updated by the GFZ (GeoForschungsZentrum).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 216-218

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Isostatic Response Function

In large time scales Earth's lithosphere exhibits a behavior regional therefore tends to experiment flexure,

because of the load that supports them. Can assume that the lithosphere presents the behaviour of a filter which

removes long amplitudes, ie, short wavelengths that are associated with local isostasy models and allows the

pass of small amplitude, or long wavelengths that are associated with flexural models. (Watts, 2001).

In the internal structure of the Earth, the part in which deflection occurs is called Efective Elastic Thickness

(EET), which is defined as the thickness of the crust that behaves elastic and support some or all of the

topographic load. (Burov and Diament, 1995). To calculate the Efective Elastic Thickness, there are various

methods which mostly are based on spectral and spatial relations between the topography and gravity, which are

obtained through the use of maps or profiles. Gravitational Admittance, is the wavenumber parameter that

modifies the topography so as so produce the gravity anomaly. (Watts, 2001).

The admitance function Z(k) is defined as:

)(

)()(

kH

kGkZ =

Where k is wavenumber, and G(k) and H(k), are Fourier transforms of gravity and topography respectively.

Figure 2. Profiles used in the Admittance Analisys. These seven (7) profiles crossing perpendicularly the Colombian Andes, and each one of them have a longitude among 400 and 700 km. These was drawed for gravity anomaly an topography heights data.

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The value obtained for the average of Admittance observed curve is compared to a set of Theoretical

Admittance curves for different values of Effective Elastic Thickness. By obtaining an root mean square

between observed curve and the overall theoretical curves, get the final value of EET.

In the present case the observed admittance fits best for the curve with EET = 20 km (Figure 3).

Conclusions

We demonstrate that the model EIGEN-CG03 it best represents surface gravity data for the Colombian Andes,

which was consistent with the result obtained by Tassara et al (2007), who used the same model for all South

America.

The obtained result here, compared with studies that employ other methods, eg, Coherence is a bit low, however

this within the expected values when using the admittance technique.

Acknowledgements We are very grateful to M.Sc. Laura Sánchez Rodríguez (Deutsches Geodätisches Forschungsinstitut) who with its concepts and suggestions enriched and allowed the work to take the right direction.

References Watts, A. B. 2001. Isostasy and Flexure of the Listhosphere, First Edition, Cambridge University Press, Cambridge, 458 pag. Tassara, A., Swain, C., Hacknet, R., & Kirby, J. 2007. Elastic Thickness structure of South America estimated using

wavelets and satelite-derived gravity data. Earth and Planetary Science letters. 253, 17 – 36. Burove, B., & Diament, M. 1995. The Effective Elastic (Te) of the continental lithosphere: What does it really mean?.

Journal of Geophysical Research. Vol. 100. No. B3. 3895-3904.

Figure 3. Admittance analysis plot showing relationship between theoretical and observed admittance with error bars.

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Numerical modeling of interplay between growth folds and fluvial-alluvial erosion-sedimentation processes: Application to the Mendoza Precordillera orogenic front (32º30’S)

Víctor Hugo García1,2

& Ernesto Osvaldo Cristallini1,2

1 Laboratorio de Modelado Geológico, Departamento de Ciencias Geológicas, Facultad de Ciencias Exactas y

Naturales, Universidad de Buenos Aires. Pabellón II, Ciudad Universitaria, C1428EGA, Ciudad Autónoma de

Buenos Aires, Argentina (victorg@gl.fcen.uba.ar) 2 Consejo Nacional de Investigaciones Científicas y Técnicas

KEYWORDS : numerical modeling, tectonic geomorphology, growth folds, fluvial-alluvial processes, mountain front

Introduction

The numerical modeling of fluvial erosion-transport-sedimentation processes has been object of strong

research in the last years. The stream power models have demonstrated to be the most efficient for simulating

fluvial processes in mountainous environments (see Whipple (2004) for a synthesis). Bedrock-rivers are the

dominant ones in these regions (Howard, 1980). These rivers are characterized by bedrock exposures along

almost all the river path, with some sectors covered by transient sediments.

Once the rivers reach the piedmont its erosion and transport capacity is rapidly reduced by the slope decrease.

The construction of alluvial fans in the mountain front is controlled by discharge events repeated along time

(Harvey et al., 2005).

In the piedmont of the active mountainous chains is frequent to find neotectonic features (fault scarps, growth

folds, etc.) as direct evidence of deformation propagation and recent tectonic activity. The geomorphic

characteristics of those features are used to establish the level of activity of the system (i.e. Burbank and

Anderson, 2001). The growth strata deriving from the interaction between the growing structures and alluvial

processes represents an important tool to analyze the temporal evolution of the structures and the deformation

rates (i.e. Burbank and Vergés, 1994).

A hypothetical fold growing in the piedmont of the Andean orogenic front was analyzed for a 10000 years

forward evolution. The results show that the presented numerical modeling platform, named ERSEDE, (García,

to be published) is a useful tool to analyze the interactions between neotectonic and surface processes.

Study area

The study area is located close to the orogenic front of the Andes at the eastern flank of the Precordillera

Mendocina between 68º 51’ and 69º 09’ LW and 32º 30’ and 32º 41’ LS (Figure 1). The hypothetical growth

anticline has been located in the Rodeo Grande pampa, this pampa is the present bajada of the mountain range.

The Canota and del Toro rivers are the main rivers draining the mountainous area and cross the Rodeo Grande

pampa up to the Las Higueras river. The Silurian-Devonian greenish sandstones and mudstones of the

Villavicencio Group (Cuerda et al., 1988) are the dominant lithology of this sector of the Precordillera

Mendocina. These rocks are deformed and can have low-grade metamorphism.

The climate of the region is arid with annual precipitations lower than 350-400 mm. The rivers are ephemerals

being actives only during the rain epoch. These rivers can transport great amount of materials during floods. The

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intense seismic activity (INPRES, 1995) and the presence of neotectonic structures as the La Cal range fault

scarp (Mingorance, 2006) and the Borbollón-Capdeville growth folds (Costa et al., 2000) are markers of

Quaternary tectonic activity in the area.

Figure 1: Location map of the study area. The approximate location of the growth fold is indicated by the anticline symbol.

Methodology

Modeling of fluvial-alluvial processes

The numerical models that simulate the landscape evolution are designed with empirical relationships and

simplifications derived from engineering transport laws (Whipple, 2004). The erosion capacity (E) of a river at

one point of its path can be determined from the follow algorithm

E = (Sm Qn Ke) – (Qs + T) (1)

where, S is the local slope, Q is the water discharge, Ke is the rock erodibility, Qs is the sediment charge in

transport and T is a threshold for fluvial erosion. The exponents m and n have been obtained from previous

works (Whipple and Tucker, 1999, Clevis et al., 2004) and have values of 0,66 y 0,33 respectively.

The water discharge (Q) is obtained from the next formulae:

Q = A P0.65 (2)

where, A is the upstream drainage area, and P is the precipitation. A straight corollary from the equation (1) is

that erosion will take place only when the result being positive.

The erodibilities (Ke) for different litologies have been calibrated using denudation rates at short temporal

scales in the Bolivian Andes (Aalto et al., 2006).

When the transport capacity is surpassed by the sediment charge the river has to reduce it and release some

material. The quantity of material to release is function of the space available. This material added to the old

topography can not exceed the actualized topography of the immediately previous point in the river path. The

limit will be controlled by the slope of the point of interest with respect of the next point in the river path. The

sedimentation can not generate depressions or increase slopes. This process continues up to the recovery of the

transport-erosion capacity, or up to the end of the river path.

Modeling of growth folds

The program ERSEDE includes a module to simulate the growth of an anticline structure on a previous

topography, real (DEM’s) or artificial. The growing of the fold is controlled by the fault-parallel flow model.

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The input data requested includes the fault geometry (dip, depth of detachment, vergence), the shortening rate

and the recurrence interval.

Model parameters

A model of 100 steps has been executed. An erodibility of 3000 x 10-7 m2/a has been assigned for all the pre-

deformation topography not taking into account the presence of Quaternary cover. This value is an average of

the erodibilities determined for metasedimentary and weak sedimentary rocks (Aalto et al., 2006).

The geometry of the fault is in agreement with others folds in the orogenic front (García et al., 2005). To

obtain a double vergent anticline the displacement over the fault plane has been distributed from 2 mm/a in the

center of the fault up to 0 mm/a at both ends of the fault. The shortening rate is coherent with rates obtained for

similar structures in the region (Vergés et al., in press).

Results and discussion

The progressive growing of the anticline can be observed in the topographic evolution of the region. Two N-S

scarps are cutting the piedmont. The western scarp represents the backlimb of the anticline and the eastern scarp

the forelimb. The area between both scarps is the hinge of the fold (Figure 2a). The presence of straight scarps

related to fold kinks is not very common in the field, which could reflect the limitation of the deformation model

applied.

In the hinge zone the pre-deformation surface is partially preserved and progressively degraded (Figure 2d).

This kind of surfaces (or pediments) is a common feature in the piedmont of the Andes at these latitudes (i.e.

García et al., 2005).

Figure 2: Temporal evolution of the study area showing: a) growing of the fold, b) variability in drainage network, c) denudation patterns, d) progressive degradation of pre-deformation surfaces (pediments) and e) profiles growth strata in the scarp zones. The location of the growth strata profiles is marked with black bars in the topographic scenes.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 219-222

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The path of the rivers changes in the piedmont along the model evolution, but without abrupt bends (Figure

2b). The path changes could be controlled by alluvial autocyclic avulsion processes. The uplift rate could be too

low to establish barriers to the rivers.

The erosion concentrates in the mountainous areas with high slopes. The sedimentation in the mountain

bedrock channels is transient. There are not great accumulations in these rivers (Figure 2c). The sedimentation

dominates in the piedmont showing distributary character along the model evolution (Figure 2c). This character

can be correlated with the processes responsible for the construction of alluvial fans (Harvey et al., 2005)

In the anticline zone the sedimentation occurs on both flanks of the structure. Erosion and no deposition

dominate in the hinge (Figure 2e). The growth strata geometry obtained is different of those modeled for longer

time spans and can be useful to analyze growth patterns in Quaternary deposits.

Conclusions

The interaction between the growth anticline and the fluvial-alluvial processes has been successfully modeled.

Geomorphic markers of neotectonic activity (pediments and scarps) have been obtained along the model

evolution. The alluvial sedimentation in the piedmont is controlled by avulsion and abandon of individual

channels. The construction of alluvial fans can be simulated with ERSEDE.

The shortening rate and the recurrence interval applied are not enough to modify the general design of the

drainage network, at least for the time span analyzed in this paper. The interaction between deformation and the

alluvial processes in the scarp zones produces growth strata geometries. The program ERSEDE is a valid tool to

simulate the interplay between neotectonic and fluvial-alluvial erosion-sedimentation processes.

References Aalto, R., Dunne T. and Guyot, J.L. 2006. Geomorphic controls on Andean denudation rate. J. Geol., 114: 85-99. Burbank, D.W. and Anderson, R.S. 2001. Tectonic geomorphology. Blackwell Science, Malden. 274 p. Burbank, D.W. and Vergés, J. 1994. Reconstruction of topography and related depositional systems during active thrusting.

J. Geophys. Res., 90: 20281-20297. Clevis, Q., de Boer, P. and Wachter, M. 2003. Numerical modelling of drainage basin evolution and three-dimensional

alluvial fan stratigraphy. Sediment. Geol., 163: 85-110. Costa, C., Gardini, C., Diederix, H. and Cortes, J. 2000. The Andean thrust fromt at Sierra de las Peñas, Mendoza, Argentina.

J. S. Am. Earth Sci., 13: 287-292. Cuerda, A.J., Lavandaio, E., Arrondo, O. and Morel, E. 1988. Investigaciones estratigráficas en el Grupo Villavicencio,

Canota, prov. de Mendoza. Rev. Asoc. Geol. Argentina, 43 (3): 356–365. García, V.H. (to be published). Modelado numérico y análogo de procesos de erosión y sedimentación fluvial y su

interacción con estructuras neotectónicas. PhD Thesis, Universidad de Buenos Aires (in preparation). García, V.H., Cristallini, E.O., Cortés, J.M. and Rodríguez, C. 2005. Structure and neotectonics of Jaboncillo and del Peral

anticlines. New evidences of Pleistocene to Holocene? Deformation in the Andean piedmont. 6º International Symposium on Andean Geodynamics, Extended Abstracts: 301-304, Barcelona.

Harvey, A.M., Mather, A.E. and Stokes, M. 2005. Alluvial Fans: Geomorphology, Sedimentology, Dynamics. Geological Society, London, Special Publications, 251.

Howard, A.D. 1980. Thresholds in river regimes. In: Coates, D.R. and Vitek, J.D. (Eds.): Thresholds in Geomorphology, 227–258. Boston.

INPRES. (1995). Microzonificacion sismica del gran Mendoza: Executive abstract, Technical publication, 19. Mingorance, F. (2006). Morfometría de la escarpa de falla histórica identificada al norte del cerro La Cal, zona de falla La

Cal, Mendoza. Rev. Asoc. Geol. Argentina, 61 (4): 620-638. Vergés, J., Ramos, V. A., Meigs, A., Cristallini, E. O., and Cortes, J. M. (in press). Crustal wedging triggering recent

deformation in the andean thrust front between 31ºS and 33ºS: Sierras Pampeanas-Precordillera interaction: J. Geophys. Res.

Whipple, K. (2004). Bedrock rivers and the geomorphology of active orogens. Annu. Rev. Earth Pl. Sc., 32:151-185. Whipple, K. and Tucker, G. (1999). Dynamics of the stream-power river incision model: implications for height limits of

mountain ranges, landscape response, and research needs. J. Geophys. Res., 104: 17661-17674.

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3D structure of the subduction zone at the Colombia – Ecuador border

Lina Constanza García-Cano1, Audrey Galve

1, Philippe Charvis

1, Audrey Gailler

1,2,

Jean-Xavier Dessa1, Bernard Pontoise

1, Yann Hello

1, Alain Anglade

1, & Ben A. Yates

1

1 Géosciences Azur, University of Nice Sophia Antipolis, IRD, CNRS, University Pierre and Marie Curie, BP 48,

Villefranche-sur-mer, 06235, France (garcia@geoazur.obs-vlfr.fr) 2 Now at IFREMER, Marine Geosciences Department, BP 70, 29280 Plouzané, France

KEYWORDS : subduction zone, 3D tomography, earthquakes, rupture zone

Introduction

At the Ecuador-Colombia border, the 500 km long rupture zone of the 1906 event (Mw= 8.8) was partially

reactivated, from south to north, by a sequence of 3 thrust events in 1942 (Mw = 7.8), 1958 (Mw = 7.7) and 1979

(Mw = 8.2), as consequence of the subduction of Nazca plate below South America plate. The rupture zones of

these seismic events abuts betweeen them (Kelleher, 1972) Fig.1. Bathymetric, passive and active seismic data,

collected off Ecuador and southwestern Colombia suggest that the interplate earthquake and the extension of

their rupture zone are at least partly controlled by structures on the downgoing and upper plates. Collot et al.

(2002) suggest that the coseismic slip zones are limited by crustal transverse faults that segment the margin in

this area, then the limit between the 1942 and 1958 rupture zones could be the Esmeraldas Fault, and the

Farallon Fault could be the boundary between ruptures zones of 1958 and 1979 seismic events.

The Esmeraldas experiment was designed to obtain a 3D lithospheric image of the Ecuador-Colombia margin

to constrain lateral variations of structures observed previously in 2D studies, and discuss their possible role in

the regional seismic cycle.

The Esmeraldas experiment

The wide-angle seismic survey was conducted from R/V Atalante (IFREMER, France) and R/V Providencia

(DIMAR, Colombia) in March-April 2005. It comprised 25 crossing seismic lines, a total of 19,300 shots

triggered at a 150 m interval using seismic source composed of 8 16-liter airguns. Thirty-one 3-component

portable stations on land and twenty-five 4-component Ocean Bottom Seismometers (OBS) offshore were

deployed to record nearly 2,890 km of seismic lines (Fig.2c).

For OBS record-sections, the good quality of the data leads us to identify P-waves arrivals up to offsets of

146 km applying only a butterworth filter which bandpass is 5 to 15 Hz. For on-land stations, we apply a

frequency-dependent phase coherence filter (Schimmel and Gallart, 2007) to enhance signal coherency at far

offsets and attenuate incoherent noise. Therefore, we were able to pick arrivals up to 195 km. To perform our 3D

tomography, we picked the first arrivals on both OBS and land-stations close to the shoreline. For each pick an

uncertainty ranging from 0.02 s to 0.10 s was assigned based on the signal to noise ratio.

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Figure 1. Geodynamic context of the South Colombia-North Ecuador. Rupture zones of the 1906, 1942, 1958 and 1979 earthquakes (yellow star are epicenters) are outlined respectively in violet, dark blue, red and light green contours (Beck and Ruff, 1984). In dark green, the 1958 earthquake aftershocks (mb > 4.8) relocated by Mendoza and Dewey (1984), in light blue by Engdahl and Villaseñor (2002). The arrow shows Nazca plate motion vector relatively to South America plate (Trenkamp et al., 2002). The black rectangle represents our tomographic inversion box.

3D Tomographic calculations

For our tomographic inversion we delimited a tomographic box delimited by the region covered by the

seismometer network. Therefore the box dimension is 332 km x 254 km x 30 km. The bathymetry was slightly

smoothed and the coordinate system transformed to UTM and rotated 50° from 1.40°N/84.4°W to have a

tomographic box in distance and parallel to the seismic lines.

The FAST code (Zelt and Barton, 1998) was used to obtain the 3D velocity model. This tomographic approch

requires a starting model to solve the direct problem. We tested two kinds of models: one model associated with

the oceanic crust and the other one with the continental crust. These models were a 3D extrapolation of 1D

models determined by trial-and-error forward modeling of an average traveltime curve calculated from several

OBS, using the least squares inverse method of Zelt and Smith (1992) and taking account the arrival times of

rays that travel in each type of crust. In order to solve the direct problem, we used a grid of equidimensional

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distance nodes of 0.5 km and a cell size of 2 km x 2 km x 0.5 km for the inverse problem.

The 3D velocity structure

The 3D image highlights the oceanic plate having a thickness of around 6 to 8 km that plunges towards the east

and a variation of the slab dip that seems to increase from south to north.

First of all, in order to check our 3D inversion, we compare our results with those of Gailler et al. (2007). To

do so, we extract a 2D slice from our 3D velocity model along the 2D tomographic profile SAL-6 of Gailler et

al. (2007). We observe similar structures in spite of a coarser spatial sampling of wavefields, such as the

presence of a low velocity zone in the upper plate.

Our 3D tomographic inversion allows us to follow the variation of these structures laterally to see if they really

play a role in the triggering and propagation of large earthquakes.

We can follow the variation of the low velocity zone structure seen in the upper plate at depth between 5 to

10 km (Fig.2). Gailler et al. (2007) showed that the top of the zone of velocity inversion is coincident with a

reflector seen on MCS data. This reflector was interpreted on MCS data from Collot et al. (2004) as a splay fault

that would decouple the bulk of the margin basement from its frontal part during the great earthquake rupture.

From the 2D profile SAL-6 location, the low velocity zone can be seen down to 30 km to the south, in the region

of the Esmeraldas fault.

However, we also detected an along parallel trench oriented "high velocity" zone in the upper plate that might

denote a change in lithology. It is centered on the 1958 rupture zone. This "high velocity" zone may be to extend

southward of the 1958 rupture zone determined by Beck and Ruff (1984) from waveform modelling, that means

southward of the Esmeraldas fault. If there is a relation in between this “high velocity” zone and the 1958

rupture zone, our results are in favor of a larger 1958 rupture zone consistent with aftershocks location.

In addition, the seaward limit of the 1958 earthquake rupture zone defined by Beck and Ruff (1984) coincides

with the occurrence of a rapid lateral variation of seismic velocities in the first 10 km (Fig. 2).

Figure 2. a and b) 2D vertical cross-section through the final 3D velocity model obtained by first arrival time inversion tomography. Note the presence of the low velocity zone between 110-120 km distance at depth of 5 to 10 km on both slices. The black line on top indicates the extension of the 1958 earthquake rupture zone. c) Our tomographic box in UTM coordinates. OBS positions are represented by red circles, the on-land station locations by yellow circles. The red lines

correspond to the seismic lines. Orange lines a and b represent the position of extracted 2D vertical cross-section.

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References Beck, S.L., Ruff, L.J., 1984 — The rupture process of the great 1979 Colombia earthquake: Evidence for the asperity

model. J. Geophys. Res., Vol. 89, B11: 9281-9291. Collot, J.-Y., Marcaillou, B., Sage, F., Michaud, F., Agudelo, W., Charvis, P., Graindorge, D., Gutscher, M.-A., Spence, G.,

2004 — Are rupture zone limits of great subduction earthquakes controlled by upper plate structures? Evidence from multichannel seismic reflection data acquired across the northern Ecuador–southwest Colombia margin, J. Geophys. Res., 109, B11103, doi:10.1029/2004JB003060.

Engdahl, E.R., Villaseñor, A., 2002 — “Global seismicity 1990-1999” . In Lee, W., Kanamori, H., Jennings, P., Kisslinger, C. (éd.): International handbook of earthquakes and engineering seismology (part A), Hardbound, Academic Press: 665-690.

Gailler, A., Charvis, P., Flueh, E.R., 2007 — Segmentation of the Nazca and South American plates along the Ecuador subduction zone from wide-angle seismic profiles, EPSL, 260, doi: 10.1016/j.epsl.2007.05.045.

Kelleher, J.A., 1972 — Rupture zone of large South American earthquakes and some predictions. J. Geophys. Res., 77: 2087-2103.

Mendoza, C., Dewey, J.W., 1984 — Seismicity associated with the great Colombia-Ecuador earthquakes of 1942, 1958, and 1979: Implications for barrier models of earthquake rupture. Bull. Seismol. Soc. Am., Vol.74, (2): 577-593.

Schimmel, M., Gallart, J., 2007 — Frenquency-dependent phase coherence for noise suppression in seismic array data. J. Geophys. Res., 112, B04303, doi: 10.1029/2006JB004680.

Trenkamp, R., Kellogg, J.N., Freymueller, J.T., Mora, H., 2002 — Wide plate margin deformation, southern Central America and northwestern South America, CASA GPS observations, Journal of South American Earth Sciences 15: 157-170.

Zelt, C.A., Smith, R.B., 1992 — Seismic traveltime inversion for 2-D crustal velocity structure. Geophys. J. Int., 108: 16-34.

Zelt, C.A., Barton, P.J., 1998 — Three-dimensional seismic refraction tomography: A comparison of two methods applied to data from the Faeroe Basin. J. Geophys. Res., 103, B4: 7187-7210.

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Block uplift and intermontane basin development in the northern Patagonian Andes (38º-40ºS)

Ezequiel García-Morabito1,2

& Víctor A. Ramos1,2

1 Laboratorio de Tectónica Andina, Departamento de Cs. Geológicas, Facultad de Ciencias Exactas y Naturales,

Universidad de Buenos Aires, Argentina (ezegm@gl.fcen.uba.ar) 2 CONICET (Consejo Nacional de Investigaciones Científicas y Técnicas)

KEYWORDS : Northern Patagonian Andes, intermontane basins, block uplift, Miocene compression, syntectonic sequences

Introduction

The northern Patagonian Andes are a relatively low relief mountain chain with an attenuated crustal thickness

(~ 30 Km) (Yuan et al., 2006) as the result of crustal collapse related to the steepening of the subducted Nazca

Plate after a period of shallow subduction during the Late Miocene (Kay, 2002). Superimposed structural styles

associated with alternating tectonic regimes derived from this process can be recognized in a segment between

the 38ºS and 39º30’S (García Morabito & Folguera, 2005). Most of the compressive structures recognized in this

segment in the inner retro-arc area were active during Middle – Upper Miocene times. The Quechua orogeny

(Miocene to Recent) produced N and NW trending folds and thrusts, reactivation of Triassic and Cretaceous

structures and the conditions for the development of several Miocene depocenters related to basement block

uplift and west-verging thrusts. The recognition and interpretation of the main structures and spatial and

temporal distribution of Tertiary sequences allowed us to establish a tectonic model in which the uplift of a N-

NW-trending block during Miocene times, originated a series of small intermontane depocenters in the inner

retro-arc area of the Northern Patagonian Andes. As a result of that syntectonic and synorogenic deposits in

some cases represented by 500 meters of volcaniclastic and clastic sequences (Mitrauquen, Chimehuin, Collon

Cura Formations) accumulated in a compressive regime. These depocenters can be integrated in a narrow basin

developed in association with the western margin of a structural high called Copahue - Catan Lil High. This

block has a good expression between the 38º and 39º30’S. It extends for over 180 km from North to South and

constitutes the drainage divide at these latitudes. It also presents along-strike differences in style, magnitude and

distribution of the deformation, controlled by pre-existing variations in the rift geometry, in the basement

structures, as well as in the position of the depocenters of a Oligocene – Lower Miocene extensional basin (Cura

Mallín Basin). We differentiate three domains in the base of these differences, separated by transverse features;

each domain is associated with an intermontane depocenter (Lonquimay, Kilca and Catan Lil (Rossello et. al,

2007) Depocenters).

Segmentation of the Copahue – Catan Lil block

Northern Segment (Lonquimay depocenter)

The western margin of this block is controlled here by the N-S-trending Pino Seco thrust (Suárez & Emparán

1997) which overlaps the Late Miocene Mitrauquén Formation with Jurassic turbidites that can be assigned to

the Cuyo Group. The Mitrauquén Formation is formed by ignimbrites, andesitic and basaltic lavas and

conglomeradic sequences interpreted as deposited in a braided river system (Suárez & Emparán, 1997). This

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 227-230

228

sequence lies above the Cura Mallin Formation; K-Ar ages between 9.5 ± 2.8 and 8.0 ± 0.3 Ma were obtained by

Suérez y Emparan, 1997. It was interpretated as a syntectonic unit deposited in the forelimb of a fault

propagation fold related to the Pino Seco thrust (Melnick et al., 2006).

We have recognized as Lonquimay a small narrow depocenter developed in the western border of a west -

verging uplifted block which concentrates about 400 meters of volcanic products and clastic sediments.

Figure 1. Mayor structures and morphotectonic units of the Andes between 37°30’S and 40°S. In grey the CPCL Block, in dark grey retroarc volcanism. Pointed areas indicate the position of Miocene depocenters, yellow arrows indicate sediments supply areas and sense.

Central Segment (Kilca depocenter)

The Kilca depocenter lies parallel to the course of the Aluminé River and extends for about 80 Km from north

to south concentrating over 500 meters of ignimbrites, conglomerates and sandstones of the Chimehuin

Formation. It is limited to the east by the Catan Lil Range, a N-NW-trending block which produces a

topographic break along an W-E transect as well as a change in the amplitude of the orogen at these latitudes.

This block has higher altitudes than the Principal Cordillera and it is separated by tens of kilometers from the

volcanic arc. The western slope of this range is controlled by the Kurumil fault system (Fig.2A), represented by

N-S, NW and NE-trending west verging thrusts which overlies the lower members of the Chimehuin Formation

with pre-Mesozoic rocks. Some of these faults represent compressional reactivation of pre-existing normal faults

associated to the Mesozoic Catan Lil depocenter.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 227-230

229

Figure 2. A : Simplified geological map of the inner retro-arc area between 38°45’-40°S showing the principal structures and the position of the Kilca and Catan Lil depocenters. B : Interpretation of a seismic line (lbv 94-152) located north of Las Coloradas town showing the deep strutures at the eastern margin of the Catan lil depocenter.

We can correlate this unit with the Mitrauquen Formation based on similar lithological characteristics and

times of deposition. Ages between 13.8 ± 0.9 and 6.2 ± 0.3 Ma were obtained from volcanic rocks intercalated in

the sequence by Vattuone y Latorre (1998) and Re et al., 2000. Progressive intraformational unconformities and

internal unconformities were also observed along the eastern margin of the basin, suggesting a syntectonic

sedimentation, with the coarse grained conglomerates of this unit representing the synorogenic deposits related

to the uplift of the Catan Lil Range during middle?-upper Miocene times.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 227-230

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Southern Segment – Catan Lil Depocenter

The genesis of this narrow basin is related to the inversion of a Triassic N-NW trending half graben system. An

asymmetric fault propagation fold was formed as the result of a west verging thrust developed by compressional

reactivation of a preexisting normal fault of the Mesozoic rift phase (Fig. 2B). The basement was involved in

the deformation as well as the Mesozoic strata which were strongly folded to the west as we can observe in a

seismic line located north of Las Coloradas town, generating the conditions for the deposition of volcanic and

sedimentary strata of the Chimehuin Formation in onlap relation with the frontal limb of the anticline. Next to

the contact, at the eastern margin of the Tertiary depocenter, this unit contents growth strata (Leanza et al.,

2003), which indicates that the deposition of the lower member of the Chimehuin Formation was simultaneous

with the last inversion phase. Farther to the west of this contact, the dip of the Chimehuin Formation increases

from sub horizontal up to 30°E, conforming a wide syncline associated to a westward propagation of the

deformation. This depocenter was passively carried by the fold belt after the deposition of the lower members of

the Chimehuin Formation when the fault system propagated to the west.

Concluding remarks

Thrust and belt loading produced several internal depocenters by fault reactivation and flexural subsidence.

The volcanic, volcaniclastic and fluvial sequences that constitutes the infill of this depocenters represents

syntectonic and synorogenic deposits associated to the uplift of the Copahue – Catan Lil Block during middle? -

upper Miocene times, and can be integrated in a NW-trending basin associated with its western margin.

References Garcia Morabito, E., Folguera, A., 2005. El alto de Copahue – Pino Hachado y la fosa de Loncopue: un comportamiento

tectónico episódico, Andes Neuquinos (37°-39°S). Revista de la Asociación Geológica Argentina, 60 (4): 742-761. Kay, S.M., 2002. Tertiary to Recent transient shaloww subduction zones in the Central and Southern Andes. XV Congreso

Geologico Argentino (El Calafate). 3: 282-283. Leanza, H.A., Repol, D., Escosteguy, L., Salvarredy Aranguren, M., 2003. Estratigrafia del Mesozoico en la comarca de

Fortin 1 de Mayo, cuenca Neuquina sudoccidental, Argentina. – Geologia, 1: 1-22. SEGEMAR (Servicio Geológico Minero Argentino), Serie de Contribuciones Técnicas.

Melnick, D., Rosenau, M., Folguera, A., Echtler, H. 2006. Late Cenozoic tectonic evolution, western flank of the Neuquén Andes between 37º and 39º south latitude: in S.M. Kay and V.A. Ramos (eds.), Late Cretaceous to Recent Magmatism and Tectonism of the Southern Andean Margin at the Latitude of the Neuquen Basin (36°-39°S). Geological Society of America, Special Paper, 407: 73-95.

Rossello, E.A., Cobbold, P.R., Marques, F.O., 2007. Late Oligocene to Miocene growth strata in two Andean intermontane basins of Neuquen province, Argentina (37º-40ºS). 20th Colloquium on Latin American Earth Sciences, Kiel. Actas: 55.

Suárez, M. y Emparán, C., 1997. Hoja Curacautín. Regiones de la Araucanía y del Bio-Bio. Carta Geológica de Chile, 1:250.000, Servicio Nacional de Geología y Minería de Chile. 71: 1-105.

Yuan, X., Asch, G., Bataile, K., Bohm, M., Echtler, H., Kind, R., Oncken, O., Wölnbern, I., 2006. Deep seismic image of the Southern Andes: in S.M Kay and V.A Ramos (eds.), Late Cretaceous to Recent Magmatism and Tectonism of the Southern Andean Margin at the Latitude of the Neuquen Basin (36°-39°S). Geological Society of America, Special Paper, 407: 61-72.

Vattuone, M.E., Latorre, C.E., 1998. Caracterización geoquímica y edad K/Ar de basaltos del Terciario superior de Aluminé. Neuquén. 10° Congreso Latinoamericano de Geologia y 6° Congreso Nacional de Geología Económica, Buenos Aires. 2: 184-190.

Re, G.H., Geuna, S.E., Lopez Martinez, M., 2000. Geoquímica y geocronología de los basaltos neógenos de la región de Aluminé (Neuquén – Argentina). 9° Congreso Geológico Chileno, Puerto Varas. 2: 62-66.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 231-234

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Pre-Andean deformation in the southern Central Andes (32°-33°S)

Laura Giambiagi1, José Mescua

1, Alicia Folguera

2, & Amancay Martínez

3

1 Instituto Argentino de Nivología, Glaciología y Ciencias Ambientales (IANIGLA)_CCT-CONICET, Parque San

Martín s/n, 5500 Mendoza, Argentina (lgiambia@lab.cricyt.edu.ar, jmescua@lab.cricyt.edu.ar) 2 Servicio Geológico Minero Argentino, Instituto de Geología y Recursos Minerales. Julio A. Roca 651, piso 10.

Buenos Aires, Argentina (alifolguera@yahoo.com) 3

MAPaS. Lavalle 409 1ºB. San Luis, 5700 San Luis, Argentina (amancaymartinez@yahoo.com)

KEYWORDS : early Late Paleozoic compressional deformation, Permo-Triassic extension, Precordillera, Cordillera Frontal

Introduction

The present day structure of the Andes between 32° and 33°S, is characterized by different N-S trending

morphostructural units which are, from west to east (Fig. 1): Cordillera Principal, Cordillera Frontal and

Precordillera. The Precordillera forms a north-south trending mountain chain composed of thick metamorphic

and sedimentary sequences of Cambrian to Neogene age. This paper presents a detailed investigation of the

structure and evolution of the southern Precordillera, focusing on the probable geometry of the Pre-Andean

structures and their control on the subsequent Cenozoic Andean deformation.

Figure 1. Location map of the study area, showing the morphostructural units that composed the Andean Mountains between 30° ans 34°S. Inferred suture zones between Gondwana and Cuyania and Chilenia terranes are outlined.

Deformational events and related structures

The prolonged history of convergence against the Pacific edge of Gondwana resulted in several episodes of

shortening, extensional and strike-slip deformation. Overprinting relationships between different structures in the

Precordillera and Cordillera Frontal preserve evidence for at least four deformational events occurred between

Early Paleozoic and Cenozoic times. We discriminated major structures into Eopaleozoic, Neopaleozoic, Permo-

Triassic, Middle Triassic and Cenozoic structures on the basis of ages of affected units, chronological criteria,

fault orientation and sense of displacement, and analysis of mechanical consistency (Fig. 2).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 231-234

232

The Early Paleozoic tectonic history was mainly controlled by a subduction process below the western margin

of Gondwana. During this time, the western Gondwana continental edge was located east of the Andes at the

western border of the Precordillera. Folding, faulting and accompanying metamorphism, are the result of a W-E

to NW-SE crustal compressional process in Silurian to Devonian times, known as Precordilleran and Chanic

tectonic phases. These tectonic phases are thought to be related to the collision between Cuyania and Chilenia

terranes (Ramos et al., 1984; Astini, 1996). Faults identified as Early Paleozoic, affecting the Cambro-

Ordovician rocks, present N-S to NNE trends and westward vergence (von Gosen, 1995; Cortés et al., 1997;

Folguera et al., 2001). The Neopaleozoic deformation generated a broad NNE-trending belt which structures

consist of concentric folds and low- and high-angle reverse faults with NNE to N-S trends and double vergence.

In the eastern part of the Precordillera a foreland thin-skinned thrust belt developed, where sheets composed of

Silurian to Carboniferous strata were thrust eastward by low-angle NE- to NNE-trending faults. In the western

part of the belt, thick-skinned faults with westward vergence affected the Cambro-Ordovician to Carboniferous

rocks. The inconsistency between regional NNW-trending and local NNE- to NE-trending Late Paleozoic

structures can be explained by clockwise block rotations inferred to have taken place between 280 and 265 Ma,

before the extrusion of Late Permian volcanics (Rapalini and Vilas, 1991). These crustal block rotations have

been attributed by Rapalini and Vilas (1991) to dextral strike slip movement parallel or subparallel to the

continental margin.

Figure 2. Geological map of the southern Precordillera and eastern sector of the Cordillera Frontal.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 231-234

233

The Permo-Triassic evolution of southwestern South America was characterized by the development of a great

amount of volcanism under extensional conditions. This extensional regime continued during Triassic times and

led to the formation of a series of rift depocentres, with an overall NNW trend. Permo-triassic volcanic rocks are

affected by normal and oblique-slip normal faults with WNW to NW trends (Fig. 2). Middle Triassic

sedimentary rocks were deposited during the formation of NW to NNW normal faults which controlled the

formation of several hemigrabens. The Cenozoic Andean chain was formed through interaction of the Nazca and

South American plates. The pre-existing structures were oriented oblique to the direction of plate convergence

during the Neogene, and some of them were reactivated and subjected to oblique compression.

Kinematic analysis of the deformational events

Detailed outcrop-scale field analysis was conducted at 56 sites throughout the Southern Precordillera and

eastern sector of the Cordillera Frontal. At each outcrop, we measured fault orientation, slip direction, average

displacement or fault width and sense of displacement of the structures. We also measured and studied fold

attitudes, the angular relationship between bedding and tectonic foliations and asymmetric folds in shear zones

to obtain the vergence of Paleozoic structures. Fault-slip data were acquired by measurements on minor faults

and were considered in terms of incremental strain (Cladouhos and Allmendinger, 1993). We used the kinematic

hypothesis proposed by Marrett and Allmendinger (1990) and Twiss and Unruh (1998) to determine constraints

on the orientation and magnitudes of the principal strain rates from a large set of fault-slip data. Principal strain

axes have been computed using the moment tensor summation method (both unweighted and weighted by

measured displacement) as implemented in FaultKin 2.1.1 stereonet program of Almendinger et al. (2001).

Early Paleozoic rocks are affected by folds and faults which kinematic analysis indicates an E-W

compressional direction and vergence toward the west. Late Paleozoic faults indicate a NW-SE to NNW-SSE

compressional direction. We rotated both Early and Late Paleozoic structures 80° counterclockwise in order to

reconstruct their orientation previous to Late Permian vertical rotation, and obtained a N-S compressional

direction with southward vergence during the Early Paleozoic and a SW-NE compressional direction with

double vergence for the Late Paleozoic. The kinematic axes of the Permo-Triassic faults indicate that this

deformational phase was characterized by NNE-SSW oriented extension (N23ºE stretching direction).

Cenozoic structures appear to be due to two interfering processes: a regional E-W shortening direction and

sinistral strike-slip movements along preexisting NW trending crustal weakness zones. These Cenozoic strike-

slip faults are affecting only the western sector of the Precordillera and the eastern sector of the Cordillera

Frontal. Crosscutting relationships of Cenozoic structures in the western part of the Precordillera indicate that

the E-W shortening event occurred first and thrusts and reverse faults were afterward cut by strike-slip faults.

The shortening event is interpreted to be related to a compressional tectonic regime which evolved into a strike-

slip regime at the time the compressional process migrated progressively further to the east, towards eastern

Precordillera. This change from a compressional stress regime to a strike slip one in western Precordillera could

have been due to a change in the vertical stress axes from 3 to 2.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 231-234

234

Figure 3. Paleozoic to Cenozoic episodes of deformation in the Precordillera. Fault-slip data and kinematic solutions for fault arrays.

References Allmendinger, R. W., Marrett, R. A., Cladouhos, T., 2001. Faultkin 2.1.1. A program for analyzing fault slip data. Absoft

Corp-. Astini, R. A., 1996. Las fases diastróficas del Paleozoico medio en la Precordillera del oeste Argentina – evidencias

estratigráficas. 13° Congreso Geológico Argentino y 3° Congreso de Exploración de Hidrocarburos, Buenos Aires, Actas 5: 509-526.

Cladouhos, T., Allmendinger, R., 1993. Finite strain and rotation from fault-slip data. Journal of Structural Geology 15: 771-784.

Cortés, J. M., González Bonorino, G., Koukharsky, M., Pereyra, F., Brodtkorb, A., 1997. Hoja 3369-09, Uspallata. Servicio Geológico Minero Argentino. Boletín inédito, 210 p.

Folguera, A., Etcheverria, M., Pazos, P., Giambiagi, L. B., Cortés, J. M., Fauqué, L., Fusari, C., Rodríguez, M. F., 2001. Descripción de la Hoja Geológica Potrerillos (1:100.000). Subsecretaría de Minería de la Nación, Dirección Nacional del Servicio Geológico, 262 p.

Gosen, W. von, 1995. Polyphase structural evolution of the southwestern Argentine Precordillera. Journal of South American Earth Sciences 8, 377-404.

Marrett, R., Allmendinger, R. W., 1990. Kinematic análisis of fault-slip data. Journal of Structural Geology 12, 973-986. Ramos, V., Jordan, T., Allmendinger, R., Kay, S., Cortés, J., Palma, M., 1984. Chilenia: un terreno alóctono en la evolución

paleozoica de los Andes Centrales. 9° Congreso Geológico Argentino. Actas 1: 84-106. Rapalini, A. E., Vilas, J. F., 1991. Tectonic rotations in the Late Palaeozoic continental margin of southern South America

determined and dated by palaeomegnetism. Geophysical Journal International 107, 333-351. Twiss, R. J., Unruh, J. R., 1998. Analysis of fault slip inversions: Do they constrain stress or strain rate?. Journal of

Geophysical Research 103, 12,205-12,222.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts : 235-237

235

Origin of flat subduction zones: Numerical application to central Chile – western Argentina between 29°S and 34°S

Gaelle Gibert1, Riad Hassani

2, Emmanuel Tric

1, & Tony Monfret

1

1 Laboratoire Géosciences Azur, 250 rue A. Einstein, 06560 Valbonne, France (Gibert@geoazur.unice.fr,

Tric@geoazur.unice.fr, Monfret@geoazur.unice.fr) 2 Laboratoire de Géophysique Interne et de Tectonophysique, Campus scientifique Université de Savoie, 73376

Le Bourget-du-Lac, France (Riad.Hassani@univ-savoie.fr)

KEYWORDS : subduction, thermomechanical 3D numerical modelisation, Nazca and South American plates, Juan Fernandez

Ridge, equivalent elastic thickness, finite element

Adeli is a 2 or 3-dimentional finite element mechanical software (Hassani, 1994). We added a thermal module

to the initial mechanical 3-dimensioned approach. Temperature of the mantle is applied to the envelope of the

slab which heat up by thermal transfer. Thermal contact between two bodies is resolved by a double

interpolation point by point. Possibilities of adding thermal (initial or imposed along the simulation) flux,

internal heating or initial or imposed temperature are now provided in Adeli. Several tests have been performed

to valid this new thermal module. Using this new tool, an application to South America have been attempted, as

thermal conditions might be very important in the special context of central Andes (Gutscher, 2002).

Central Chile and Argentina Andes are an almost linear orogenic belt formed at the convergent plate margin

where Nazca plate plunge under South America. Between 29°S and 34°S, a flat subduction is underplating the

continental plate of South America (figure 1). Previous studies aiming to understand this phenomenon are

numerous. Literature in this domain is various, and lecturer can find paper about seismicity in the crust

(Barrientos et al, 2004) and in the slab (Cahill and Isacks, 1992), tectonism in the orogen (Smalley et al, 1993;

Ramos, 1999), geometry of the slab (Araujo and Suarez, 1993) and so on.

Figure 1: Synthetic map of South America and eastern part of Nazca plate between 26°S and 37°S. Solid lines on the map represent isodepth contours of the slab, depth in km. Dotted line symbolized south separation between north flat slab and south 30° dipping slab. Profile on the left show general trend of the slab north and south (ie. flat and not flat slab) of the dotted line. After Cahill and Isacks (1992).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts : 235-237

236

If history of the flattening is now well known (Kay and Abbruzzi, 1996; Litvak et al, 2007), several reasons

attending to explain the existence of a flat slab in this region are suggest, yet none of them is making consensus.

Numerous authors, like Kopp et al (2004), Yanez et al (2002) propose numerous models (respectively

serpentinization and buoyancy of the Juan Fernandez Ridge, slab detachment and buoyancy of the Juan

Fernandez Ridge). This numerical work aim to search which reason(s) are effectively responsible of the flat

subduction and the extent of each.

Continental South American and oceanic Nazca plates are modeled by respectively 45 km and 25 km

equivalent elastic thickness plate (see Figure 2, equivalent elastic thickness from Burov and Diament, 1996). The

Juan Fernandez Ridge is modeled by a second volume incorporated in the oceanic plate, with its specific density

(see Figure 3, after Kopp et al, 2004). We modeled various examples, and compared them.

Starting with some known elements (density and effective elastic thickness of continental and oceanic plates,

friction, …) we attempt to obtain underplating. We propose different configurations, varying in (1) size and

orientation of the Juan Fernandez Ridge (according to Yanez et al (2002) work), (2) presence or not of a rupture

in the equivalent continental elastic thickness (see Perez-Gussinyé et al, 2008), (3) thermals proprieties of both

plates. Very first results will be presented here.

Figure 2: Starting model of subduction generated with Adeli. (1): Nazca plate, (2): South American plate, black arrows: direction of velocity imposed to the plates, length is proportional to modulus, white line: profile of the contact zone where friction is imposed.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts : 235-237

237

Figure 3: (a): starting model, (b): model with a perpendicular volcanic chain.

References Araujo, M., & Suarez, G. 1993. Geometry and state of stress of the subducted Nazca plate beneathh central Chile

and Argenitna: evidence from teleseismic data. Geophysical Journal International. 116: 283-303. Barriento, S., Vera, E., Avarado, P., & Monfret, T. 2004. Crustal seismicity in central Chile. South America

Earth Sciences. 16: 759-768. Burov, E. 1996, Isostasy, equivalent elastic thickness, and inelastic rheology of continents and oceans. Geology.

24: 419-422. Cahill, T., & Isacks, B. 1992. Seismicity and shape of the subducted Nazca plate. Journal of Geophysical

Research. 97: 17,503-17,529. Gutscher, M. 2002. Andean subduction styles and their effect on thermal structure and interplate coupling. South

American Earth Sciences. 15: 3-10. Hassani, R. 1994. Modélisation numérique de déformation des systèmes géologiques. PhD Thesis document. 139

pages Kay, S., & Abbruzzi, J. 1994. Magmatic evidence for Neogene lithospheric evolution of the centreal Andean

« flat-slab » between 30°S and 32°S. Tectonophysics. 259: 15-28. Kopp, H., Flueh, E., Papenberg, C., & Klaeschen, D. 2004. Seismic investigations of the O’Higgins Seamount

Group and Juan Fernandez Ridge : aseismic rige emplacement and lithosphere hydration. Tectonics. 23. Litvak, V., Poma, S., & Kay, S. 2007. Paleogene and Neogene magmatism in the Valle del Cura region : new

perspectives on the evolution of the Pampean flat slab, San Juan province, Argentina. South American Earth Sciences. 24: 117-137.

Perez-Gussinyé, M., Lowry, A., Phipps Morgan, J., & Tassara, A. 2008. Effective elastic thickness variations along the Andean margin and their relationship to subduction geometry. Geochemistry Geophysics Geosystems. 9: 2.

Ramos, V. 1999. Plate tectonic setting of the Andean cordillera. Episodes. 22: 3. Smalley, R., Pujol, J., Regnier, M., Chiu, JM., Chatelain J.-L., Isacks, B., Araujo, M., & Puebla, N. 1993.

Basement seismicity beneath the Andean precordillera thin-skinned thrust belt and implications for crustal and lithospheric behavior. Tectonics. 12: 63-76.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 238-241

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The active upper plate deformation of the Central Andes forearc, northern Chile

G. González1, R. Allmendinger

2, T. Dunai

3, J. Cembrano

1, J. Martinod

4, D. Rémy

4,

D. Carrizo5, J. Loveless

6, E. Veloso

1, F. Aron

1, & J. Cortés

1

1 Departamento de Ciencias Geológicas, Universidad Católica del Norte, Chile (ggonzale@ucn.cl)

2 Department of Earth and Atmospheric Sciences, Cornell University, USA (rwa1@cornell.edu)

3 School of Geosciences, University of Edinburgh, United Kingdom (tibor.dunai@ed.ac.uk)

4 Laboratoire des Mécanismes de Transfert en Géologie (LMTG), Université de Toulouse, France

(martinod@lmtg.obs-mip.fr). 5

Geodesy Laboratory, Institut de Physique du Globe de Paris, Paris, France (carrizo@ipgp.jussieu.fr) 6 Department of Earth and Planetary Sciences, Harvard University, USA (loveless@eps.harvard.edu)

KEYWORDS : active deformation, Central Andes, forearc, northern Chile

Introduction

Along the Central Andes, subduction of oceanic crust beneath the South American Plate corresponds to a first

order process that controls the accumulation of bulk strain in the overriding plate. Despite this general statement,

the precise understanding of the present day strain accumulation in the forearc is still matter of debate. Several

questions related to this topic – such as strain distribution, compatibility between long- (5 Ma to Present) and

short-term deformation (decennial scale), and fraction of convergence velocity accumulated as permanent

deformation in the overriding plate – are still not solved issues. During the last years, we have performed

detailed field studies aimed to unravel these main issues. Main goals were determination of the kinematics, the

age and the slip rate of upper-plate faults at four key localities: 1) the Peninsula de Mejillones, 2) Coastal

Cordillera near Antofagasta (23°30’S), 3) Coastal Cordillera close to the Salar Grande area (21°30’S), and 4) the

south-eastern part of the Salar the Atacama Basin (21°30’S-68°20’W). Methodology considered fault

characterization (Allmendinger et al., 2005), Ar40-Ar39 age determinations and exposures ages based on

cosmogenic nuclide dating (González et al., 2005; Carrizo et al., 2008). Also, we performed numerical

modelling to understand the deformation processes that operate at the scale of a single structure (Loveless 2007;

González et al., 2008) and at the scale of the convergent margin (Loveless, 2007; Cortés et al., this issue).

In this contribution we would like to present a large-scale overview of the superficial distribution of the strain

field in the Central Andes forearc of Northern Chile, particularly focusing on the main factors that control this

distribution. By using the slip rate of the faults we estimate the velocity of deformation of the whole forearc

system.

The long-term strain accumulation in the Central Andes Forearc

The entire portion of the Coastal Cordillera close to Antofagasta, including the Mejillones Peninsula, is

characterized by NS-trending (parallel to the margin) normal faulting. The most conspicuous normal faults are

the Caleta Herradura, Mejillones and Salar del Carmen faults. The first two structures – exposed in the

Mejillones Peninsula – affect Mio-Pleistocene marine sediments, which represent an emerged portion of the

marine shelf (Niemeyer et al., 1996; Delouis et al., 1998). Radiometric data and micropaleontological dating of

the graben infill show that these faults are active, at least, since 24 Ma to Present. Recent normal faulting is

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expressed by deformation of Late Pleistocene-Holocene alluvial fan sediments as well as near-shore marine

sediments. By using the vertical offsets and the maximal age of the sedimentary infill of the graben structures it

is possible to estimate a long-term slip-rate of 0.03 mm/yr. On the contrary, a faster slip-rate of about 0.2 mm/yr

was determined by Marquardt (2005) by using the numerical age of Late Pleistocene alluvial fan sediments cut

by the Mejillones Fault. This discrepancy between long-term constrained slip-rate and short-term slip-rate can be

explained by the occurrence of quiescence periods of fault activity. These periods can be identified when

considering a time window of million years. In the Salar del Carmen area, the main branch of the Atacama Fault

System forms spectacularly well preserved fault scarps which deform several alluvial fans (Armijo and Thiele,

1990; González et al. 2003). At this locality 21Ne exposures ages – obtained from the surface of the alluvial fans

– indicate that fans became inactivate ca. 400 ka before Present. By using these exposure ages it is possible to

obtain a minimal slip-rate of 0.01 mm/yr (González et al., 2005). Fault-slip data show that the Mejillones

Peninsula and the Coastal Cordillera close to Antofagasta is affected by E-W extension.

At the Mejillones Peninsula the occurrence of Late Pleistocene marine terraces evidences a long-term uplift

that started, at least, 400 ka before Present. According to Marquardt (2005), marine terraces formed during

maximal interglacial stages. Uplift rates in the northern part of the Mejillones Peninsula are strongly controlled

by the activity of the normal faults (Marquardt, 2005); for example, marine terraces located in the footwall of the

Mejillones Fault show uplift rates of 0.5-0.7 mm/yr whereas those located in the hangingwall exhibit uplift rates

of 0.2-0.5 mm/yr.

Particulary, on November 14th, 2007, a strong earthquake occurred along the coupling zone between the Nazca

and the South American plates. The seismogenic fracture propagated nearly 200 km from north to the south

stopping at the Mejillones Peninsula. The southern ending of this rupture coincided with the northern terminus of

the Mw=8.1 June 1995 Antofagasta earthquake (Neic catalogue USGS). Coseismic deformation, registered by

InSar Data, shows that during the November earthquake the Mejillones Peninsula experienced an uplift of 25 cm

(see details in Cortes et al., this issue). In addition to the co-seismic uplift, interferograms show a small

subsidence of 1.5 cm at the eastern part of the Mejillones Fault. The coseismic displacement of the Mejillones

area is compatible with the long term uplift registered in the marine terraces whereas the small subsidence close

to the Mejillones Fault is compatible with reactivation of this structure.

The Coastal Cordillera near the Salar Grande shows several structures that have clear topographic expressions.

These structures are: 1) NW-SE trending dextral faults, 2) E-W reverse faults, and 3) N-S reverse faults

(Allmendinger et al., 2005; Carrizo et al., 2008). All of these structures deform Miocene alluvial surfaces and

several younger drainages incised within the alluvial surfaces. Ar40-Ar39 dating of tuff layers deformed by the E-

W trending reverse faults indicates that fault activity is younger than 300 ka; which implies that slip-rate for

these faults is close to 0.01 mm/yr. Because the other sets of structures (NW-SE and N-S trending faults)

displaced very old alluvial surfaces we cannot calculate an accurate slip rate for them. Exposures ages for these

old surfaces (determined by 21Ne) indicate that these latest faults have minimal slip-rates of ca. 1x10-3 mm/yr

(Carrizo, 2007). Kinematic analyses based on fault-slip data indicate that this part of the Coastal Cordillera is

affected by subhorizontal constriction, in where N-S shortening is dominant. During the last years, three crustal

earthquakes, with P axes oriented N-S (Neic catalogue of USGS), affected the coast of southern Peru and

northern Chile, indicating that N-S contraction in the forearc is an active process.

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On the contrary, the inner part of the forearc – the Salar the Atacama Basin – shows a more regular strain

pattern characterized by E-W shortening affecting several Pliocene ignimbrites and the sedimentary infill of the

Salar de Atacama Basin. Deformation here is represented by several subparallel N-S trending anticlinal ridges

which are the superficial expressions of blind reverse faults. We have calculated that the long-term slip-rates in

this area vary between 1 x10-3 to 1 x 10-4 mm/yr. Faster slip rate of 1.7 mm/yr was obtained for the Holocene

activity of reverse fault beneath the Salar de Atacama Basin (Jordan et al. 2002). Henceforth, we suggest that

folding mechanism of the ignimbrite layers should be a slow process mainly controlled by creeping of the layers

above the underlying reverse fault. Kinematic data show pure E-W contraction. We did not find evidences of

trench parallel strike-slip deformation in this part of the forearc. This contrasts with reports of dextral faults in

the magmatic arc close to the Salar de Atacama Basin affecting several valleys incised in Mio-Pliocene

ignimbrites.

Discussion

The above described structures show that the present day active strain in the forearc of Northern Chile has been

and it is heterogeneously accumulated. Upper plate extension is restricted to the western border of the forearc

whereas shortening is concentrated in the internal eastern portion of the forearc. Both types of strain regimes

started during the Miocene and have remained active until Present. Constrictional deformation is a local process

restricted to the inner part of the curved portion of the Central Andes forearc. Upper plate extension is related to

the coseismic phase of the subduction earthquake cycle whereas shortening is related to the interseismic phase.

Slip rates documented in the forearc range between 2x10-1 mm/yr and 1x10-4 mm/yr. When using younger

stratigraphic markers faster-slip rates are obtained. This indicates that at the long-term scale (millions of years)

upper plate faults experience quiescence periods which reduce the calculated slip-rates, in contrast to those based

on a short-term scale. The presently available slip-rates show that a minor fraction (<< 10%) of the present day

convergence is absorbed by distributed faulting on the upper plate. In contrast, a larger portion of the

convergence velocity (>40%) is taken by the seismic coupling zone during large subduction earthquakes.

The absence of active trench-parallel strike-slip faults in the forearc indicates that this part of the Central

Andes is rheologically homogeneous at the million and millennium time scales. Interseismic GPS velocities

measured in the forearc are lightly oblique to the N-S trending margin and subparallel to the convergence

velocity (Bevis et al., 1999). The decrease in the interseismic velocity field from the Coastal Cordillera to the

magmatic arc suggests that a fraction of this velocity is accumulated in the forearc, which is consistent with our

observation. In order to produce the observed local N-S contraction, the forearc needs to move northward. GPS

velocities show that the forearc -during interseismic periods of the subduction earthquake cycle – tries to move

obliquely with respect to the margin. However, the curved geometry of this margin does not allow it. In

consequence the curved portion of the forearc acts as a buttress structure, allowing that the stress-loading on

trench orthogonal reverse faults to be effective. Similarly, the slightly oblique movement of the forearc should

result in minor strike-slip displacements in the magmatic arc.

References Allmendinger, R., Gonzalez, G., Yu, J., Hoke, G., Isacks, B., 2005. Trench-parallel shortening in the northern Chilean

forearc: Tectonic and climatic implications: Geological Society of America Bulletin 117, 89–104.

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Armijo, R., Thiele, R. 1990. Active faulting in northern Chile: ramp stacking and lateral decoupling along a subduction plate boundary? Earth Planet Sci Lett, 98: 40-61.

Bevis, M., Kendrick, E., Jr., Smalley, R., Brooks, B.A.,Allmendinger, R.W., and Isacks, B.L., 2001, On the strength of interplate coupling and the rate of back arc convergence in the Central Andes; An analysis of the interseismic velocity field: Geochemistry, Geophysics,Geosystems, G3, v. 2, doi:10.129/2001GC000198.

Carrizo, D. 2007. Procesos de deformación neogena en el antearco externo del norte de Chile (20,5-21°S): La subducción oblicua en un margen curvo. Tesis de doctorado, Universidad Católica del Norte, 261 p.

Carrizo, D., González, G, Dunai, T. 2008, Constricción neógena en la Cordillera de la Costa, norte de Chile: neotectónica y datación de superficies con 21Ne cosmogénico Revista Geológica de Chile 35 (1): 1-38.

Delouis, B., H. Philip, L. Dorbath & Cisternas, A. 1998. Recent crustal deformation in the Antofagasta region (northern Chile) and the subduction process. Geophys. J. Int., 132: 302 – 338.

González, G. Dunai, T, Carrizo, D. and Allmendinger, R. 2006. Young displacements on the Atacama Fault System, northern Chile from field observations and cosmogenic 21Ne concentrations. Tectonics. Vol.25,No.3,TC3006.

González, G., Cembrano, J., Carrizo, D., Macci, A. & Schneider, H. 2003. The link between forearc tectonics and Pliocene-Quaternary deformation of the Coastal Cordillera, northern Chile. Journal of South American Earth Sciences, 16: 321-342.

González, G., Gerbault, M., Martinod, J., Cembrano, J., Carrizo, D., Allmendinger, R., Espina, J. 2008. Crack formation on top of propagating reverse faults of the Chuculay Fault System northern Chile: Insights from field data and numerical modelling. Journal of Structural Geology.

Jordan, T. N. Muñoz, N., Hein, M., Lowenstein, T., Godfrey, L., Yu J. 2002. Active faulting and folding without topographic expression in an evaporite basin, Chile GSA Bulletin; November 2002; v. 114; no. 11; p. 1406–1421;

Loveless, J. 2008. Extensional tectonics in a convergent margin setting: Deformation of the northern Chilean forearc. Ph.D Thesis, Cornell University, 311 p.

Loveless, J., Hoke, G., Allmendinger, R., González, G., Isacks, B., Carrizo, D. 2005. Pervasive cracking of the northern Chilean Coastal Cordillera: New evidence for forearc extension. Geology, 33: 973-976.

MARQUARDT, C. 2005. Déformations Néogènes le long de la cotê nord du Chili (23°-27°S), avant-arc des Andes Centrales. Thèse doct., univ. Toulouse-III, 212 p.

Niemeyer, H., González, G., Martinez-de Los Rios E. 1996. Evolución tectónica cenozoica del margen continental activo de Antofagasta, norte de Chile. Revista Geológica de Chile, 23 (2): 165–186.

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Modern geodata management — A tool for interdisciplinary interpretation and visualization

H.-J. Götze, T. Damm, & S. Schmidt

Institut für Geowissenschaften, Abtl. Geophysik, Otto-Hahn-Platz 1, 24118 Kiel, Germany

(hajo@geophysik.uni-kiel.de)

KEYWORDS : geophysics, geoinformatics, 3D visualization, Central America, 3D modeling

Introduction

In the last years new methods of data acquisition and processing in geosciences have increased the amount of

data, inspired by the growing computational power available. In this paper we present the conception and

technical realization of a Web Portal of a big interdisciplinary research group. The combination of geodata

management as a metadata catalogue together with web mapping technology is presented. Furthermore future

aims like implementing common standards to simplify data exchange will be pointed out, their impact on

geoscientific work and the benefit for other collaborative research centers in particular will be discussed.

The Kiel Collaborative Research Centre “SFB 574 - Volatiles and Fluids in Subduction Zones: Climate

Feedback and Trigger Mechanisms for Natural Disasters” is an interdisciplinary geoscientific research project.

As over fifty researchers are working on different geoscientific aspects of subduction processes, data

management and presentation using internet technologies like web mapping is crucial for interdisciplinary

cooperation. Also efforts are made to strengthen the intensive collaboration and data exchange with partners

from the participating countries of Central America and colleagues from the US Margins program.

Technical aspects

The data bank consists of geophysical data (seismic reflection, receiver function and earthquake data and

surface, areo- and satellite potential field data sets), high resolution topography, geochemical and tectonic data,

and geological maps in digitized formats. Data sets cover the territories of Costa Rica and Nicaragua, the

offshore Pacific and Caribbean. Coupled to the catalog is a web mapping solution based on the “UMN

MapServer”2 project of the Minnesota University, which dynamically plots datasets from the catalog. These two

parts interact with the content management system “Conpresso”3 for the static pages of the website and a

seamlessly integrated web portal has been formed. We like to mention that no commercial software was used.

The “Conpresso Content Management System” (CMS) was improved by integrating “phpDBrelations”5, a

system, which was developed by one of authors (T. Damm). Essentially it is a web based tool for creating and

modifying database tables and relations. Generalized export routines for web content have been programmed, for

example, the queried content can be simply listed, put into tables or send as XML; refer also to Fig 1.

In the last years Geosciences have developed - guided by advanced in computer technology - from 2D map

production on paper towards 3D and 4D modeling of the reality. We use the data of our data bank for case

studies and present here the possibilities of stereoscopic projection. Clearly related with the storage and retrieval

of the different datasets is the need of visualization. Nowadays more and more geoscientific disciplines have to

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interpret not only one- or two-dimensional data. Hence we will present a 3D visualization system, which helps to

conduct multi-parameter interpretation on e.g. GravMag data.

Figure 1. Using the apache web server and a MySQL database, we are using the “Conpresso Content Management System”

(CMS) for static content. Dynamic content is managed using the phpDBrelations toolkit with it is export routines for HTML

and XML. For web mapping, the UMN Mapserver is used via the PHP/Mapscript framework

Perspective visualization

3D visualization in geoscientific interpretation is a useful tool, if numerous, heterogenic datasets have to be

visualized at the same time - not just for purely three-dimensional datasets. As soon as 1D and 2D data is

georeferenced correctly, it can be shown together e.g. together with 3D topography or bathymetry, or with

modeled or measured underground structures, for example from density modeling or 3D seismic results like

tomography or receiver function analysis. For the stereo visualization we use two mainstream beamer with XGA

resolution and 2500 ANSI Lumens, polarizing filters, a 200x150cm silver screen and polarizing glasses (see

Fig. 5). The polarizing filters are mounted in front of the two roof mounted projectors. They let transmit the light

orthogonally polarized to each other, beamer A emits just horizontally polarized light, beamer B just vertically

polarized. The silver coated screen preserves this polarization states and hence using the polarizing glasses, a

channel separation is achieved, as each eye just observes the picture of one beamer. 3D visualization in

geoscientific interpretation is not only a very helpful tool, if numerous, heterogenic datasets have to be

visualized at the same time but it is very popular amongst software developers, with highly sophisticated 3D

tools being especially fashionable. The profit for geo-science is indisputable: only a few years ago, one had to

construct geological bodies (salt domes, subduction zones etc.) using foamed or transparent plastics in order to

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model ideas. Today, software tools allow visualization of complicated geometries on the screen, and enable the

interpreter/observer to interact with the model through rotation, change of viewpoint and/or illumination.

Figure 2. Bouguer Anomaly onshore and Free-Air Anomaly offshore is visualized for the area of Costa Rica and Nicaragua.

Most prominent is the positive anomaly in the area of the Nicoya Peninsula and the negative gravity west of the Talamanca

belt.

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Reflection seismic imaging of the Chilean subduction zone around the 1960 Valdivia earthquake hypocenter

Kolja Groß, Stefan Buske, Serge A. Shapiro, Peter Wigger, & the TIPTEQ Research Group,

Seismics Team*

Free University Berlin, Department of Geophysics, Malteserstr. 74-100, 12249 Berlin, Germany

(kolja@geophysik.fu-berlin.de)

* TIPTEQ Research Group, Seismics Team: Groß, K., FU Berlin, Germany (a); Micksch, U., GFZ Potsdam,

Germany (b); Araneda, M., SEGMI, Santiago, Chile (c); Bataille, K., Universidad de Concepción, Chile (d);

Bribach, J. (b); Buske, S. (a); Krawczyk, C.M. (b); Lüth, S. (a,b); Mechie, J. (b); Schulze, A. (b); Shapiro, S.A. (a);

Stiller, M. (b); Wigger, P. (a); Ziegenhagen, T. (b).

KEYWORDS : reflection seismics, subduction zone processes, subduction channel, convergent margins, South America

Introduction

With a quarter of the worldwide seismic energy in the last century having been released in the Chilean region

alone (Scholz 2002), the Andean subduction zone is a natural laboratory for our seismogenic zone studies. The

overarching purpose of project TIPTEQ (from The Incoming Plate to mega-Thrust EarthQuake processes),

which comprises 13 subprojects, is to investigate processes active at all scales in the seismogenic coupling zone

in south central Chile (Fig. 1), which hosted the rupture plane of the 1960 Valdivia earthquake (Mw =9.5)

(Cifuentes 1989; Barrientos & Ward 1990).

In this paper we present a structural image and a interpretation of the reflection seismic data set across the

Chilean subduction zone at 38.2° S. Figure 1 shows the location of the approx. 100 km long reflection seismic

profile running from west to east across the Chilean subduction zone. For details on the experiment design and

acquisition parameters see Groß et al. (2008).

Figure 1. Location map of the active-source seismological experiment of project TIPTEQ . Black line – receiver line, black ticks – shot locations, blue line – CDP line, red star – hypocentral area of the 1960 earthquake (Krawczyk & the SPOC Team 2003). The yellow and magenta lines mark the CDP line of the SPOC onshore seismic reflection profile and the SPOC wide-angle refraction profile, respectively (Krawczyk et al. 2006). The green line marks the eastern end of the marine seismic reflection profile SO161-038/42 (Reichert & Schreckenberger 2002). The red line maps the surface trace of the Lanalhue fault zone (LFZ; after Melnick & Echtler 2006). NAZCA and SAM in the inset label the Nazca and South American plates, respectively.

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Data processing

The SPOC wide-angle refraction experiment provided us with a P-wave velocity model (Krawczyk & the

SPOC Team 2003) along the same latitude (Fig. 1). This velocity information was used to produce a prestack

depth migrated section of the vertical component of the Near Vertical Reflection (NVR) data set. Kirchhoff pre-

stack depth migration (KPSDM) was performed in a 3-D volume using the true source and receiver coordinates,

thereby taking into account the actual crooked line geometry and the topography along the profile using the

method of Buske (1999). The single shot gathers were migrated separately using the true phase information.

Then trace envelopes of all migrated single shots were calculated and stacked to form a 3-D image. An analysis

of the 3-D migration volume showed almost no structural dip perpendicular to the survey line. That allowed us

to further increase the signal-to-noise ratio up to a factor of 2 by summing of the W–E oriented depth slices. The

resulting 2-D depth section is shown in Figure 2.

Figure 2. Kirchhoff prestack depth migration of the NVR data set. The figure shows a stack of the envelopes of migrated single shots. Intensity increases from blue to red. The two triangles on the horizontal axis mark the beginning of the profile (at the coast x 0km) and the end of the profile (in the Central Valley at x 95 km).

The migration of all three components considering P- and S-wave traveltimes (PP, SS), as well as converted

waves (PS, SP) shows S-wave energy from the plate interface down to a depth of approx. 35 km. There is almost

no converted energy observed. We characterized the subduction zone further using two innovative imaging

techniques based on KPSDM: Reflection Image Spectroscopy (RIS) and Fresnel Volume Migration (FVM).

Results and discussion

The seismic section (Fig. 2) clearly shows the subducted oceanic Nazca plate below the segmented forearc and

a highly reflective overriding South American plate. We associate the high reflectivity at the plate interface with

the existence of a subduction channel with a varying thickness of 2-5 km down to a depth of at least 38 km. It

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Figure 3. Sections through the Chilean subduction zone at 38.2° S. The coast is at x = 0 km for all sections. Above: P-wave velocity model from local earthquake tomography (Haberland et al. 2008). Center: Distribution of electrical resistivity (Brasse et al. 2008). Below: preliminary interpretation. In all three figures the Kirchhoff prestack depth migration (Fig. 2) is superimposed.

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might continue towards depth. The continental Moho is not clearly imaged. The reflectivity east of the

hypocenter shows horizontal structures at various depths, which give rise to different eastward continuations of

the continental Moho. The position and extent of the continental mantle wedge changes accordingly. Major

forearc features such as the crustal Lanalhue fault zone and a strong west-dipping reflector perpendicular to the

plate interface, can be observed.

Figure 3 shows the reflectivity together with the interpretation and superimposed on other geophysical results

obtained within project TIPTEQ. The comparison with the P-wave velocity model from local earthquake

tomography (Haberland et al. 2008) shows good correlation between high velocities and high reflectivity in the

continental crust (and vice versa) and the comparison to the distribution of electrical resistivity (Brasse et al.

2008) shows a correlation of low resistivity and high reflectivity (and vice versa).

The area around the 1960 Valdivia earthquake hypocenter is characterized by high electrical resistivity and low

reflectivity. Migration of lowpass filtered seismic data (RIS), however, images the plate interface as a

continuous linear feature, that shows no reduced reflectivity around the hypocenter.

For a more detailed discussion of the seismic section see Groß et al. (2008).

Acknowledgements

This work was part of the R&D-Program GEOTECHNOLOGIEN funded by the German Ministry of Education and Research (BMBF) (Grant 03G0594) and the German Research Foundation (DFG). The project benefited from grants of the Freie Universität Berlin and the GFZ Potsdam; seismic stations were provided by the Geophysical Instrument Pool Potsdam and the Freie Universität Berlin. We thank all participants in the field and the Chilean inhabitants for having made this survey possible.

References Barrientos, S. & Ward, S. 1990. The 1960 Chile earthquake: inversion for slip distribution from surface deformation.

Geophys. J. Int., 103, 589–598. Brasse, H., Kapinos, G., Li, Y., Mütschard, L., Soyer, W. & Eydam, D. 2008. Structural electrical anisotropy in the crust at

the South-Central Chilean continental margin as inferred from geomagnetic transfer functions. PEPI, submitted. Buske, S. 1999. Three-dimensional pre-stack Kirchhoff migration of deep seismic reflection data. Geophys. J. Int., 137, 243–

260. Cifuentes, I. 1989. The 1960 Chilean earthquake. J. geophys. Res., 94(B1), 665–680. Groß, K., Micksch, U. & TIPTEQ Research Group, Seismics Team 2008. The reflection seismic survey of project TIPTEQ -

the inventory of the Chilean subduction zone at 38.2° S. Geophysical Journal International 172 (2) , 565-571 doi:10.1111/j.1365-246X.2007.03680.x

Haberland, C 2008. Structure of the seismogenic zone of the South-Central Chilean margin revealed by local earthquake travel time tomography. JGR, in preparation.

Krawczyk, C. & the SPOC Team 2003. Amphibious seismic survey images plate interface at 1960 Chile earthquake. EOS Trans. Am. geophys. Union, 84(32), 301, 304–305.

Krawczyk, C., Mechie, J., Lüth, S., Ta árová, Z., Wigger, P., Stiller, M., Brasse, H., Echtler, H.P., Araneda, M. & Bataille, K. 2006. “Geophysical Signatures and Active Tectonics at the South-Central Chilean Margin”. In The Andes—Active Subduction Orogeny. Frontiers in Earth Sciences, Vol. 1, pp. 171–192, eds Oncken, O., Chong, G., Franz, G., Giese, P., Götze, H., Ramos, V., Strecker, M. & Wigger, P., Springer Verlag, Berlin.

Melnick, D. & Echtler, H., 2006. “Morphotectonic and geologic digital map compilations of the south-central Andes (36–43° S)”. In The Andes—Active Subduction Orogeny. Frontiers in Earth Sciences, Vol. 1, pp. 565–568, eds Oncken, O., Chong, G., Franz, G., Giese, P., Götze, H., Ramos, V., Strecker, M. & Wigger, P., Springer Verlag, Berlin.

Reichert, C. & Schreckenberger, B., 2002. Cruise report SO-161 leg 2 & 3, SPOC (Subduction Processes Off Chile). Tech. rep., BGR Hannover, pp. 142.

Scholz, C., 2002. The Mechanics of Earthquakes and Faulting. Cambridge University Press, Cambridge, UK, 471 pp.

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Chile Triple Junction migration, mantle dynamics and Neogene uplift of Patagonia

B. Guillaume1, J. Martinod

1, & L. Husson

2

1 LMTG, Université de Toulouse-CNRS-IRD-OMP, 14, Avenue Édouard Belin 31400 Toulouse, France

(bguillau@lmtg.obs-mip.fr, martinod@lmtg.obs-mip.fr) 2 Géosciences Rennes, UMR CNRS 6118 - Université Rennes-1, Campus de Beaulieu, 35042 Rennes cedex,

France (laurent.husson@univ-rennes1.fr)

KEYWORDS : subduction, Patagonia, dynamic topography, uplift, geomorphology

Introduction

Geologists often consider that the topography of the earth essentially results from isostasy, topographic highs

being balanced by crustal roots and/or hot lithospheric mantle. Mantle dynamics, however, also induce forces

that deflect the earth topography, with vertical motions that can reach several hundreds of meters (Hager and

Clayton, 1989; Le Stunff and Ricard, 1997; adek and Fleitout, 2003). Dynamic topography reaches its

maximum amplitude above subduction zones (Husson, 2006). In continental domain, the dynamic component of

topography is difficult to discriminate, because the altitude is largely controlled by lithospheric mass and

temperature heterogeneities. Continental margins are nevertheless affected by long-wavelength surface

deflections that can be recorded by the geological imprint.

During the last 14 Myr, the Chile Triple Junction (CTJ) has migrated from 54°S to its present-day position at

about 46°30’S, as different segments of the Chile spreading ridge successively entered the subduction zone. In

order to investigate the impact of the associated mantle flow on the vertical surface motion, we analyze the

evolution of sedimentation, erosion, and tectonic features during the Neogene. We focus our study on the mild-

deformed central Patagonian basin, between 44°S and 48°S, from the thrust front of the Cordillera to the Atlantic

coast, because this area is poorly affected by Neogene tectonics and also remained ice-free during glaciations,

therefore showing a pristine morphology, preserved from the erasure of the glaciers.

Upper Oligocene to Holocene geological evolution of eastern Patagonia

From upper Oligocene to early Miocene, a widespread transgression occurred in the Patagonian basin. This

transgression is marked by the deposition of near-shore marine strata. These marine series are replaced by fluvial

deposits of the Santa Cruz Formation and its lateral equivalents (Ramos, 1989; Suárez et al., 2000), which have

been dated, south of the CTJ, between 22 and 14 Ma (Blisniuk et al., 2005). These deposits display syn-

contractional structures (Flint et al., 1994; Suárez and de la Cruz, 2000). The Patagonian transgression and the

deposition of the overlying continental series coincide with the break-up of the Farallon plate at ~24 Ma

(Lonsdale, 2005). This period is also marked by a more trench-perpendicular and faster convergence between the

oceanic and South America plates (Pardo-Casas and Molnar; 1987; Somoza, 1998; Lonsdale, 2005). The

development of the basin and deposition of the Santa Cruz Formation would result from the overfilling of a

subsiding basin responding to the uplift of the Cordillera and to the rapid subduction of the newly formed Nazca

plate beneath the continent.

In the middle Miocene, subsidence stopped and poorly consolidated conglomerates, known as “Rodados

Patagonicos” deposited, forming widespread terraces from the Andean foothills to the Atlantic coast. Two

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generations of surfaces have been mapped: (1) surfaces associated to different pulses of piedmont aggradation,

and (2) surfaces corresponding to major fluvial terraces (Río Senguerr, Río Deseado, Cañadon Salado-Cañadon

del Carril) (Figure 1).

Figure 1. Miocene to Holocene geologic map of the Patagonian basin.

A long-wavelength uplift can be detected looking at the present-day slope of terraces. If the initial downstream

slope of terraces is difficult to constrain, it is reasonable to consider that the terraces displayed horizontal

profiles in the direction perpendicular to paleovalleys. Figure 2 shows that the older levels of the Río Deseado

terraces system are tilted southward, with slopes that range between 0.05% and 0.11%. Levels T8De to T12De

are not tilted. The evolution of the slope for each terrace shows that a gentle southward tilting occurred between

the deposition of T3De and T8De, following a northward regional tilting that developed between the deposition

of T1De and T2De. North of the CTJ, the topography of the aggradation deposit levels along with the

longitudinal profiles of the Río Senguerr terraces and the captures of Cañadon Salado and Río Senguerr suggest

that this area has continuously been tilted northward, until recent times.

Figure 2. Valley-perpendicular topographic profile of Río Deseado terraces (see Fig. 1 for location) and values of tilting for each terrace. 5.28 : age in Ma from Gorring et al. (1997).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 249-252

251

Causes of the post-middle Miocene uplift of eastern Patagonia

The growth of southern Andes in the Oligo-Miocene resulted from crustal shortening, which, in turn,

controlled subsidence in the foreland and deposition of the Santa Cruz molasse. Using standard elastic

parameters and an elastic thickness for the Patagonia continental lithosphere between 20 and 30 km (Tassara et

al., 2007), the distance between the chain and the forebulge would range between 150 and 210 km. The posterior

flexural uplift of the foreland resulting from the diminution of the Andean load should be restricted to the same

area close to the chain. It cannot explain the regional continental-scale uplift registered from the Andes to the

Atlantic coast. In contrast, dynamic topography over subduction zones result in long-wavelength downward

deflections of the overriding topographic surface as far as 1000 km away from the trench.

Figure 3. (A) Map of the uplift of the overriding plate resulting from the episodic subduction of 4 ridge segments (SCR) below South America accompanying the northward migration of the CTJ. Light and dark gray dots mark the position of the CTJ before and after the subduction of each ridge segment, respectively. (B) Trench-parallel uplift profiles at 300 km, 500 km, and 700 km from the trench for each triple junction migration increment. The boundary between regions of northward and southward tilting for each longitudinal profile shifts northward, delineating sectors with different tilting histories.

The peculiar geodynamic evolution of Patagonia during the last 14 Myr can be responsible for the regional-

scale uplift and tiltings recorded by the post middle-Miocene terraces. South of the CTJ, the Antarctic plate is

slowly subducting below the continent whereas north of the CTJ, the Nazca plate subducts rapidly, inducing the

downward deflection of the continental plate. We propose that this downward deflection has been progressively

cancelled in southern Patagonia as the CTJ was migrating northward, resulting in the uplift of that part of the

continent. We computed the dynamic deflection induced by the Nazca slab, using a simple model based on the

Stokeslet approximation (Husson, 2006). Vertical deflections appear as far as 1600 km from the trench and reach

a maximum value close to 1000 meters at a distance of ~325 km from the trench. Figure 3 reproduces more

specifically the vertical motions resulting from mantle dynamics as a consequence of the episodic northern

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 249-252

252

migration of the CTJ. The tilting values obtained in our models correspond to those observed in central

Patagonia.

Conclusions

Middle Miocene is marked by a switch from subsidence to uplift of eastern central Patagonia. We propose that

this change results from the episodic northward migration of the CTJ, inducing the opening of a slab window

below southern Patagonia that cancels the dynamic downward deflection of the continental plate above the

subduction zone and, in turn, causes the observed uplift and associated N-S trending tilting.

References

Blisniuk, P.M, Stern, L.B., Chamberlain, C.P., Idleman, B., & Zeitler, P.K. 2005. Climatic and ecologic changes during Miocene surface uplift in the Southern Patagonian Andes. Earth and Planetary Science Letters 230: 125-142. adek, O., & Fleitout, L. 2003. Effect of lateral viscosity variations in the top 300 km on the geoid and dynamic topography. Geophys. J. Int. 152: 566-580.

Flint, S.S, Prior, D.J., Agar, S.M., & Turner, P. 1994. Stratigraphic and structural evolution of the Tertiary Cosmelli Basin and its relationship to the Chile triple junction. Journal of the Geological Society, London 151: 251-268.

Gorring, M.L., Kay, S.M., Zeitler, P.K., Ramos, V.A., Rubiolo, D., Fernandez, M.I., & Panza, J.L. 1997. Neogene Patagonian plateau lavas: Continental magmas associated with ridge collision at the Chile Triple Junction. Tectonics 16(1): 1-17.

Hager, B.H., & Clayton, R.W. 1989. Constraints on the structure of mantle convection using seismic observations, flow models and the geoid. In Peltier, W.R. (ed.): Mantle Convection, New-York, Gordon and Breach: 657-763.

Husson, L. 2006. Dynamic topography above retreating subduction zones. Geology 34(9): 741-744. Le Stunff, Y., & Ricard, Y. 1997. Partial advection of equidensity surfaces: A solution for the dynamic topography problem?

Journal of Geophysical Research 102: 24,655-24,667. Lonsdale, P. 2005. Creation of the Cocos and Nazca plates by the fission of the Farallon plate. Tectonophysics 404: 237-264. Pardo-Casas, F., & Molnar, P. 1987. Relative motion of the Nazca (Farallon) and South American plates since Late

Cretaceous time. Tectonics 6: 233-248. Ramos, V.A. 1989. Andean Foothills Structures in Northern Magallanes Basin, Argentina. AAPG Bulletin 73(7): 887-903. Somoza, R. 1998. Updated Nazca (Farallon)-South America relative motions during the last 40 Ma. Implications for

mountain building in the central Andean region. Journal of South American Earth Sciences 11: 211-215. Suárez, M., & De La Cruz, R. 2000. Tectonics in the eastern central Patagonian Cordillera plutons, Chile (45°30’-47°30’S).

Journal of the Geological Society, London 157: 995-1001. Suárez, M., De La Cruz, R., & Bell, C.M. 2000. Timing and origin of deformation along the Patagonian fold and thrust belt.

Geological Magazine 137: 345-353. Tassara, A., Swain, C., Hackney, R., & Kirby, J. 2007. Elastic thickness structure of South America estimated using wavelets

and satellite-derived gravity data. Earth and Planetary Science Letters 253: 17-36.

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253

The dynamic forearc of southern Peru

Sarah R. Hall1, Daniel L. Farber

2, Laurence Audin

3, & Robert C. Finkel

4

1 University of California, Santa Cruz, 1156 High St., Santa Cruz, CA 95060, USA (shall@pmc.ucsc.edu)

2 Lawrence Livermore National Laboratory, LLNL, Livermore, CA 94550, USA (farber2@llnl.gov)

3 Institut de Recherche pour le Développement, LMTG – UMR 5563, Observatoire Midi-Pyrénées, Toulouse,

31400 France (laurence.audin@ird.fr) 4 Lawrence Livermore National Laboratory, LLNL, Livermore, CA 94550, USA (finkel1@llnl.gov)

KEYWORDS : cosmogenic, neotectonic, erosion, incision, pediment

Recently, there has been a renewed interest in models of active tectonic and climatologic processes along the

Andean margin. While many new studies in the forearc regions of southern Peru and northern Chile have

presented data constraining morphotectonic chronologies as well as the rates of surface process, there has yet to

be a complete synthesis of these new data. As recently as ~5 years ago, the preferred working model was that

most of the low-relief surfaces within the Atacama Desert were ancient relict surfaces abandoned >7Ma due to

incision caused by periods of intense surface uplift (Tosdal et al., 1984), and that the western limb of the

Altiplano is a passive monocline with no significant Neogene structures accommodating deformation (Isacks,

1988). Until recently, documented active deformation was limited to major strike-slip and normal faults in the

Precordillera, respectively that are related to oblique subduction and gravitational collapse of the western margin

of the Altiplano (Wörner et al., 2002). Extensional features, oriented both perpendicular and normal to the coast,

were also mapped (Hartley et al., 2000), however very little was known about the slip history, kinematics or

rates of motion along these structures.

Using the combination of remote sensing with high-resolution data, in situ cosmogenic isotope concentrations

and thermochronology, in recent years the community has made important advances in addressing the rates,

timings, styles, and locations of active deformation within the forearc of the Andean margin. Specifically, we

see 1) ancient surfaces reflecting erosion rates as low as <0.1m/Ma (Kober et al., 2005; Nishiizumi et al., 2005;

Hall et al., to be published) are well preserved in the forearcs of both Peru and Chile, 2) the existence of young

(30ka-1Ma) low-relief pediment surfaces due to recent landscape modifications (Hall et al., in press), 3) active

structures accommodating compressional, tensional, and shearing stresses in numerous localities within the

forearc (Allmendinger et al., 2005a; Gonzalez et al., 2006; Hall et al., in press; Audin, et al., in press), 4) a

consistent rate of river incision of ~0.3mm/yr along exoreic rivers (Hall et al., to be published), 5) uplift rates

that been variable in time and space with pulses throughout the last 10Ma (Schildgen et al., 2007; Saillard et al.,

to be published) and 6) instantaneous modern forearc rotation rates are similar to time integrated rates over the

past 10Ma (Allmendinger et al., 2005b).

To first order, we find that the Andean forearc during the last 10Ma has been quite a dynamic region, both in

terms of tectonics and climate. The coastal Atacama Desert is situated in a zone that has been hyperarid for at

least the last 3My and this has contributed to the high degree of geomorphic surface preservation in this region

(Hartley, 2003). Indeed, in an area spanning of over 11 degrees of latitude, erosion rates based on cosmogenic

isotope concentrations are consistently less than 0.1m/Ma (Kober et al., 2005; Nishiizumi et al., 2005; Hall et

al., to be published). On shorter timescales, changes in precipitation may enhance or dampen incision rates (i.e.

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254

during a glacial-interglacial transition the increased discharge and tools may enhance valley incision), however

ultimately the amount of potential base-level change is set by surface uplift or sea level change. Thus, the exact

timing of periods of more intense incision may correspond with climate events, but the total amount of incision

over time is useful for tectonic interpretations. Based on zircon and apatite (U-Th)/He ages, Schildgen et al.

(2007) interpreted periods of more intense canyon incision along the Rio Majes in southern Peru (16˚S).

Specifically, from the period 5.1-2.3 Ma, 1.4 km of incision occurred yielding an incision rate of ~0.5mm/yr and

an additional older period of 1 km of incision from 9-5.1 Ma yielding a rate of ~0.25mm/yr. Thouret et al.

(2007) suggest incision rates of ~0.2mm/yr since 9Ma in a similar area of southern Peru. Our recent work

suggests that these incision rates are very similar to measured time integrated rates since the Pleistocene on the

major exoreic rivers (Hall et al., in press; Hall et al., to be published). Along the Rio Sama, Rio Locumba, and

the Rio Osmore of southern Peru, we have mapped sequences of well-preserved strath terraces and dated (along

the Rio Sama and Rio Locumba) these using cosmogenic 10Be. Our work yields a consistent set of incision rates

of ~0.3 ± 0.1mm/yr (Figure 1). Further, where these rivers are cut by active structures, the local incision rate

determined near the knick-point reaches 0.8 mm/yr.

Figure 1. The forearc of southern Peru. Structures active during the Pleistocene include the Purgatorio Fault, the Incapuquio Fault System, and the Calientes Fault, in addition to some of the normal faults along the coast including the Chololo Fault (Audin, et al., in press) near the town of Ilo. Incision rates based on cosmogenic 10Be concentrations are 0.2-0.4 mm/yr in major drainages and up to 0.8mm/yr near active structures (Calientes Fault). The vast incised late Pliocene and Pleistocene pediment surfaces north of the Purgatorio Fault suggest surface uplift has driven incision and abandonment of these surfaces during the past ~2Ma.

Along the three major drainages of southernmost Peru, we have mapped multiple flexures trending sub-parallel

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to the coast and the Western Cordillera. In many cases, these flexures correspond to abrupt changes in river

incision and topography. Along the Rio Sama, the largest of these flexures is produced by a propagating

hanging-wall anticline above a blind thrust (Figure 1). The youthfulness of this feature is suggested by the

deflection of active channels around the propagating tip of the anticline and by young (~30-500ka) surface

exposure ages on terraces along those active channels (Hall et al., to be published).

Schildgen et al. (2007) conclude that the pattern of apatite and zircon U-Th/He ages along the Rio Majes

(16˚S) either supports the Isacks (1988) monocline hypothesis assisted mechanistically by lower crustal ductile

flow or reflects distributed forearc deformation along multiple non-surface breaking, or un-mapped faults. While

our work highlights role of contractile deformation in the southern Peruvian forearc, contractile structures

trending sub-parallel to the range front have also previously been observed in Northern Chile (Oxaya anticline;

Victor et al., 2004; Garcia and Hérail, 2005; Kober et al., 2005). While based on our mapping and chronologic

data we cannot rule out a role for lower-crustal ductile flow in the forearc of southern Peru and northern Chile

(Husson and Sempere, 2003; Hoke et al., 2007; Schildgen et al., 2007), our observations of surface breaking and

blind reverse faults as well as active footwall anticlines shows that a significant amount of uplift is

accommodated in contractile structures in the Precordillera of southernmost Peru. In this light, any additional as

of yet unmapped active contractile structures reduce the need to call on lower-crustal ductile flow to

accommodate surface uplift in this area. Given the limited number of field sites that have been studied in detail,

it is not unreasonable to suggest there is a high likelihood that more active contractile structures exist in this

region of the Peruvian forearc.

In summary, the geomorphic and structural features in this region of southern Peru provide strong evidence of

distributed crustal deformation along range-sub-parallel contractile and strike-slip structures. The observation

that Pleistocene incision rates are comparable with Late Miocene and Pliocene rates suggests to us that the rates

and style of surface uplift within the forearc of southern Peru has been ongoing and consistent (on the timescale

of 1 Myr) during the past 10Ma.

References

Allmendinger, R.W., Gonzalez, G., Yu, J., Hoke, G. and Isacks, B., 2005a — Trench-parallel shortening in the Northern Chilean Forearc: Tectonic and climatic implications. Geological Society of America Bulletin., 117(1-2): 89-104.

Allmendinger, R.W., Smalley, R., Bevis, M., Caprio, H. and Brooks, B., 2005b — Bending the Bolivian orocline in real time. Geology., 33(11): 905-908.

Audin, L., David, C., Hall, S.R., Farber, D.L., Hérail, G., in press — Geomorphic evidences of recent tectonic activity in the forearc, southern Peru., RAGA, 61, 545-554.

Garcia, M., and Hérail, G., 2005 — Fault-related folding, drainage network evolution and valley incision during the Neogene in the Andean Precordillera of Northern Chile. Geomorphology., 65: 279-300.

Gonzalez, G., Dunai, T., Carrizo, D. and Allmendinger, R., 2006 — Young displacements on the Atacama Fault System, northern Chile from field observations and cosmogenic Ne-21 concentrations. Tectonics., 25(3).

Hall, S.R., Farber, D.L., Audin, L., Finkel, R.C., and Mériaux, A.-S., in press — Geochronology of pediment surfaces in southern Peru: Implications for Quaternary deformation of the Andean forearc. Tectonophysics.

Hall, S.R., Farber, D.L., Audin, L., Finkel, R.C., to be published — Contractile deformation in the forearc of southern Peru Hartley, A.J., May, G., Chon, G., Turner, P., Kape, S.J., and Jolley, E.J., 2000 — Development of a continental forearc: A

Cenozoic example from the Central Andes, northern Chile. Geology., 28(4): 331-334. Hartley, A.J., 2003 — Andean uplift and climate change. Journal of the Geological Society., London, 160: 7-10. Hoke, G.D., Isacks, B.L., Jordan, T.E., Blanco, N., Tomlinson, A.J., and Ramenzani, J., 2007 — Geomorphic evidence for

post-10 Ma uplift of the western flank of the central Andes 18˚30’-22˚S. Tectonics., 26, TC5021, doi: 10.1029/2006TC002082.

Husson, L. and Sempere, T., 2003 — Thickening the Altiplano crust by gravity-driven crustal channel flow. Geophysical Research Letters., 30(5).

Isacks, B.L., 1988 — Uplift of the Central Andean plateau and bending of the Bolivian Orocline. Journal Of Geophysical

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Research-Solid Earth And Planets., 93(B4): 3211-3231. Kober, F., 2005 — Quantitative analysis of the topographic evolution o the Andes of Northern Chile using cosmogenic

nuclides, Ph.D. Thesis, ETH, Zürich, 131 pp. Nishiizumi, K., Caffee, M.W., Finkel, R.C., Brimhall, G. and Mote, T., 2005 — Remnants of a fossil alluvial fan landscape

of Miocene age in the Atacama Desert of Northern Chile using cosmogenic nuclide exposure age dating. Earth Planetary Science Letters. 237(3-4): 499-507.

Saillard, M., Hall S.R., Audin, L., Hérail, G., Farber D.L., Finkel, R.C., Martinod, J., Bondoux, F., and Regard, V., to be published — Pleistocene marine terrace development and non-steady long-term uplift rates along the Andean margin of Chile (31°S).

Schildgen, T.F., Hodges, K.V., Whipple, K.X., Reiners, P.W., Pringle, M.S., 2007 — Uplift of the western margin of the Andean plateau revealed from canyon incision history, southern Peru. Geology., 35(6), 523-526.

Thouret, J.C., Wörner, G., Gunnell, Y., Singer, B., Zhang, X., and Souriot, T., 2007 — Geochronologic and stratigraphic constraints on canyon incision and Miocene uplift of the Central Andes in Peru. Earth and Planetary Science Letters., 263: 151-166.

Tosdal, R.M., Clark, A.H. and Farrar, E., 1984 — Cenozoic polyphase landscape and tectonic evolution of the Cordillera Occidental, southernmost Peru. Geological Society of America Bulletin., 95(11): 1318-1332.

Victor, P., Oncken, O., and Glodny, J., 2004 — Uplift of the western Altiplano plateau: Evidence from the Precordillera between 20˚ and 21˚S (northern Chile). Tectonics., 23, TC4004.

Wörner, G., Uhlig, D., Kohler, I. and Seyfried, H., 2002 — Evolution of the West Andean Escarpment at 18 degrees S (N. Chile) during the last 25 Ma: uplift, erosion and collapse through time. Tectonophysics., 345(1-4): 183-198.

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Oxygen isotopes evidence for crustal contamination and mantle metasomatism in the genesis of the Atacazo-Ninahuilca magmatic suites, Ecuador

Silvana Hidalgo1,2

, Marie-Christine Gerbe2,3

, Hervé Martin2, Gilles Chazot

4, & Jo Cotten

4

1 Instituto Geofisico, Escuela Politécnica Nacional, Ap. 17-01-2759, Quito, Ecuador

2 Laboratoire Magmas et Volcans, UMR 6524, Université Blaise Pascal, Clermont-Ferrand, France

3 Université Jean Monnet, UMR 6524 "Magmas et Volcans", Saint Etienne, France

4 Université de Bretagne Occidentale, BP 809, 29285 Brest, France

KEYWORDS : crustal contamination, mantle metasomatism, oxygen isotopes, Atacazo-Ninahuilca, Ecuador Introduction

Two main geochemical signatures characterize the magmas from the Quaternary Ecuadorian volcanic arc: a

“classic” calc-alkaline signature and an adakitic one (Bourdon et al., 2002; Samaniego et al., 2002; Samaniego

et al., 2005; Hidalgo et al., 2007). Even though the temporal transition from calc-alkaline to adakitic features

seems well established, geochemical processes that lead to adakite suites in this region are still debated. Based

on major and trace element geochemistry, the discussion is particularly focused on the relative roles of slab-

melts, mantle metasomatism (by fluids and silicic melts) and crustal contamination.

In this paper, oxygen isotopes are used to constrain the participation of these processes at different levels: in

the mantle on the magma sources and within the continental crust during magma evolution. Interaction of

magmas with the arc crust on ascent is considered to play a limited role to produce the andesites and dacites in

Ecuador, whereas it as been considered as a major process in other continental margin settings where a thick

crust is present (e.g. central Andean volcanic zone, (James and Murcia, 1984). Furthermore, the role of source

contamination involving subducted sedimentary components has not been evaluated for Ecuadorian lavas,

whereas it was demonstrated to participate to the mantle-wedge metasomatism in island arc settings (Shimoda et

al., 1998; Bindeman et al., 2005). This study is focused on the Atacazo-Ninahuilca volcanic complex (ANVC),

located in the Western Cordillera of Ecuador, 10 Km South of Quito (Fig. 1a). This complex consists of three

Pleistocene andesitic edifices, 1 Ma to 80 Ka old (Carcacha, Atacazo and Arenal I), and several younger dacitic

domes {Hidalgo, in press. #1448}. Two of these pre-Holocene extrusions, La Viudita and Gallo Cantana, are

located outside the caldera, whereas the other Holocene five domes were emplaced within this depression

(Arenal II, La Cocha I and II, and Ninahuilca Chico I and II) (Fig. 1b).

Geochemical data

One hundred thirty-three samples from lava flows, domes and pyroclastic deposits were analysed for major and

trace elements at the Université de Bretagne Occidentale. Analytical procedure is described in detail in Cotten et

al., (1995). Oxygen isotope analyses of 23 whole rock samples and 16 mineral separates (plagioclase,

clinopyroxene and amphibole) were determined in the “Laboratoire Transferts Lithosphériques” at the Université

JeanMonnet in Saint Etienne (France) following analytical protocol by Clayton and Mayeda, (1963). O-isotope

ratios of mineral pairs (plag-cpx or plag-amph) were used to assess the equilibrium at magma temperature and to

verify that the lavas did not suffer sub-sequent O-isotopes re-equilibration, so that the whole-rock 18O can be

approximated to approach the magma 18O.

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Sr- and Nd-isotope analyses of 17 samples from ANVC were performed at the “Laboratoire Magmas et

Volcans”, Université Blaise Pascal in Clermont Ferrand (France) following the procedures described in Pin et al.

(1994) and Pin and Santos Zaldegui (1997).

Figure 1a. Geodynamic framework of Ecuador and location of ANVC in the volcanic arc. b. Sketch map of ANVC showing main volcanic units.

Characteristics of magmatic suites

Two different magmatic suites have been identified in the ANVC. The first calc-alkaline suite corresponds to

andesites of the Carcacha, Atacazo and Arenal I (CAA) edifices, while the second one is clearly an adakitic

series represented by the dacitic domes, which mostly developed during the last 12 ky (Hidalgo et al., in press.).

The older part of the complex (CAA) consists of two pyroxenes andesites, with subordinated amphibole. The

SiO2 contents (57-62 wt.%) show a positive correlation with Na2O and K2O and are negatively correlated with

MgO, CaO, TiO2 and FeO (Fig. 2a, b). All data plot on a single differentiation trend. Oxygen isotopic ratios for

these lavas (8.0 to 8.9‰; Fig. 2c) are very high compared to typical arc magmas which commonly range

between 5.7 and 6.9 ‰ (Harmon and Hoefs, 1995). Sr and Nd isotopic ratios show mantle-like homogeneous

values (0.704096-0.704372; 0.512822-0.512887). Both ratios are negatively correlated: 87Sr/86Sr increases while 143Nd/144Nd decreases (Fig. 2d).

Figure 2. a. SiO2 vs. MgO. b. SiO2 vs. K2O : note the different evolution trends for CAA lavas and the younger domes ones. c. SiO2 vs. 18O‰ wr VSMOW. d. 87Sr/86Sr vs. 143Nd/144Nd for ANVC products. e. 87Sr/86Sr vs. 18O‰ wr VSMOW.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 257-260

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The Holocene domes, bearing pl+amph+opx+mag, determine a narrow compositional trend (62 to 67 SiO2

wt.%). 18O values display a larger range than that from the CAA andesites, varying 6.1 to 9.0‰, with most

values above 7.7‰ (Fig. 2c). 87Sr/86Sr and 143Nd/144Nd are quite homogenous (0.704155-0.704318; 0.512840-

0.512894) and contrary to the former CAA magmatic suite, only weak 87Sr/86Sr variations are observed and 143Nd/144Nd is almost constant (Fig. 2d).

Discussion

Low-pressure crustal contamination

Given the mafic composition of the local basement (basalts of oceanic origin with intercalated granodioritic

intrusions), Sr and Nd isotopic systems became useless to reveal crustal contamination because there is no

isotopic contrast between the accreted basement basalts and the Quaternary lavas (Fig. 2e). Nevertheless, the

exceptionally high 18O values and the large variation ranges in the CAA lavas (8.0 to 8.9‰) suggest a

contamination process. Indeed, the strong increase in 18O (more than 0.9‰) for Carcacha-Atacazo andesites,

compared to weak SiO2 increase in the composition (7 wt%) suggests assimilation of 18O-rich materials from the

crust, in addition to a fractional crystallization process. Modelling of such an AFC process is show in Figure. 3a.

Contaminant composition is the average of those of the Miocene intrusions which display high 18O values (7.9

– 13.7 ‰). On the other hand, the dacitic domes display a large range of 18O values which cannot be related to

the previous lavas by an AFC process (Fig. 2c). Interestingly, no mafic compositions forming a clear suite with

the domes analyses have been encountered.

Mantle metasomatism

AFC models partly explain the high 18O values of CAA lavas. Nevertheless, an important enrichment in 18O

(>8.0‰) characterizes the less evolved magmas from this series (Mg#>50 and SiO2 ~ 57 wt%), compared with

mantle-derived primitive basaltic magmas of continental subduction zones (average 18O values of 6.2±0.7‰

(Harmon and Hoefs, 1995). This suggests that 18O-rich materials have contributed to the magma source.

Michaud et al. (2005) have shown that a 400-500 m-thick pile of carbonated sediments overlies the subducting

oceanic crust under the Ecuadorian margin. Such sediments have high 18O values (25-32‰, Eiler, 2001).

Futhermore hydrothermally altered oceanic basalts have highly variable 18O values, from 5 to 25 ‰ (Eiler et

al., 1998; Eiler, 2001; Schulze et al., 2003).

Based on these data the petrogenesis of the ANVC rocks seems controlled by different processes:

1) The adakitic geochemical characteristics of ANVC rocks require the addition of melts or supercritical fluids

from the subducting slab to the mantle and/or an early garnet fractionation.

2) The high O-ratios of mafic lavas of the CAA series suggest a contamination of the mantle source by 18O-rich

materials. These materials are certainly issue of partial melting or dehydration of the subducting carbonated

sediments or the altered oceanic crust.

3) The younger domes show an important variability in 18O and homogenous major elements compositions.

These characteristics indicate that these lavas could not be related to the CCA series by simple CF or AFC

processes. Thus, the high and variable 18O values may also account for source processes, or early garnet

fractionation from more primitive terms which have not been encountered during this study.

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Figure 3a. AFC model for Carcacha-Atacazo lavas. This model has been performed using major, trace elements and oxygen isotopic ratios. The evolution trend is explained by ~ 35% fractional crystallization of a cumulate composed of 43% plg + 32% amp + 16% opx + 9% Ti-mt and 7% assimilation rate of average dioritic Miocene intrusions (MI, 18O = 10,8 ‰) located in the ANCV area basement. Note that younger domes do not follow the same evolution trend. b. Curves representing mantle metasomatism by carbonated sediments (CS) and oceanic crust melts (CRB= Carnegie Ridge Basalts) in different proportions. CAA basic terms isotopic characteristics are explained by the low degrees of partial melting of a previous metasomatized source by 5 to 10% of carbonated sediments and ~10% of oceanic crust dacitic melts.

References Bindeman, I.N., Eiler, J.M., Yogodzinski, G.M., Tatsumi, Y., Stern, C.R., Grove, T.L., Portnyagin, M., Hoernle, K. and

Danyushevsky, L.V., 2005. Oxygen isotope evidence for slab melting in modern and ancient subduction zones. Earth and Planetary Science Letters, 235(3-4): 480-496.

Bourdon, E., Eissen, J.-P., Gutscher, M.-A., Monzier, M., Samaniego, P., Robin, C., Bollinger, C. and Cotten, J., 2002. Slab melting and slab melt metasomatism in the Northern Andean Volcanic Zone: adakites and high-Mg andesites from Pichincha volcano (Ecuador). Bulletin de la Société Géologique de France, 173(3): 195-206.

Clayton, R. and Mayeda, T., 1963. The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochimica et Cosmochimica Acta, 27: 43-52.

Cotten, J., Le Dez, A., Bau, M., Caroff, M., Maury, R.C., Dulski, P., Fourcade, S., Bohn, M. and Brousse, R., 1995. Origin of anomalous rare-earth element and Yttrium enrichments in subaerial exposed basalts: Evidence from french Polynesia. Chemical Geology, 119: 115-138.

Eiler, J.M., 2001. Oxygen isotope variations of basaltic lavas and upper mantle rocks. In: J.W. Valley and D.R. Cole (Editors), Stable Isotope Geochemistry. Reviews in Mineralogy and Geochemistry, Blacksburg, Virginia, pp. 319-359.

Eiler, J.M., Mc Innes, B., Valley, J.W., Graham, C.M. and Stolper, E., 1998. Oxygen isotope evidences for slab-derived fluids in the sub-arc mantle. Nature, 393: 777-781.

Harmon, R.S. and Hoefs, J., 1995. Oxygen isotope heterogeneity of the mantle deduced from global 18O systematics of basalts from different geotectonic settings. Contributions to Mineralogy and Petrology, 120: 95-114.

Hidalgo, S., Monzier, M., Almeida, E., Chazot, G., Eissen, J.-P., van der Plicht, J. and Hall, M.L., in press. Late Pleistocene and Holocene activity of the Atacazo-Ninahuilca Volcanic Complex (Ecuador). Journal of Volcanology and Geothermal Research.

Hidalgo, S., Monzier, M., Martin, H., Chazot, G., Eissen, J.P. and Cotten, J., 2007. Adakitic magmas in the ecuadorian volcanic front: Petrogenesis of the Iliniza Volcanic Complex (Ecuador). Journal of Volcanology and Geothermal Research, 159(4): 366-392.

James, D.E. and Murcia, L.A., 1984. Crustal contamination in northern Andean volcanics. Journal of the Geological Society of London, 141: 823-830.

Michaud, F., Chabert, A., Collot, J.-Y., Sallarès, V., Flueh, E.R., Charvis, P., Graindorge, D., Gustcher, M.-A. and Bialas, J., 2005. Fields of multi-kilometer scale sub-circular depressions in the Carnegie Ridge sedimentary blanket: Effect of underwater carbonate dissolution? Marine Geology, 216: 205-219.

Pin, C., Briot, D., Bassin, C. and Poitrasson, F., 1994. Concomitant separation of strontium and samarium–neodymium for isotopic analysis in silicate samples, based on specific extraction chromatography. Analytica Chimica Acta, 298: 209-217.

Pin, C. and Santos Zalduegui, J.F., 1997. Sequential separation of light rare-earth elements, thorium and uranium by miniaturized extraction chromatography: Application to isotopic analyses of silicate rocks. Analytica Chimica Acta, 339: 79-89.

Samaniego, P., Martin, H., Monzier, M., Robin, C., Fornari, M., Eissen, J.P. and Cotten, J., 2005. Temporal evolution of magmatism at Northern Vocanic Zone of the Andes : The geology and petrology of Cayambe vocanic complex (Ecuador). Journal of Petrology, 46: 2225-2252.

Samaniego, P., Martin, H., Robin, C. and Monzier, M., 2002. Transition from calc-alkalic to adakitic magmatism at Cayambe volcano, Ecuador: insights into slab melts and mantle wedge interactions. Geology, 30(11): 967-970.

Schulze, D.J., Ben Harte, John W. Valley, James M. Brenan and Channer, D.M.D.R., 2003. Extreme crustal oxygen isotope signatures preserved in coesite in diamond. Nature, 423(6935): 68-70.

Shimoda, G., Tatsumi, Y., Nohda, S., Ishizaka, K. and Jahn, B.M., 1998. Setouchi high-Mg andesites revisited -geochemical evidence for melting of subducting sediments. Earth and Planetary Science Letters, 160: 479-492.

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Incipient tectonic inversion in a segmented foreland basin: From extensional to piggyback settings

Diego Nicolas Iaffa1, F. Sábat

1, O. Ferrer

1, R. Mon

2, & A. A. Gutiérrez

2

1 Dpt. de Geodinàmica i Geofísica, Facultat de Geologia UB, Martí i Franqués s/n, 08028 Barcelona, Spain

(diegoiaffa@ub.edu) 2 Dpt. de Geología, Facultad de Ciencias Naturales e Instituto Miguel Lillo, Universidad Nacional de Tucumán,

Miguel Lillo 205 4000 Tucumán, Argentina

KEYWORDS : Argentina, Tucumán, Cretaceous rift, foreland, tectonic inversion

Introduction

The Tucumán Basin is located in the foothills of the Andes, to the East of the northern end of the Pampean

Ranges and South of the Eastern Cordillera and Santa Bárbara System (Ramos, 1999). All of them basement

uplifts bounded by high angle reverse faults (González Bonorino, 1950). The Subandean thin skinned fold and

thrust belt is located farther North where detached thick cover Paleozoic series are present (Allmendinger et al.,

1983). Several factors influenced the present structure of the foothills of the central Andes in NW Argentina.

Among these factors: 1) the regime and intensity of several tectonic events, 2) the geometry of the Wadati –

Benioff zone, 3) structural discontinuities in the basement and 4) the architecture of the sedimentary cover

(Allmendinger et al., 1983; Jordan and Alonso, 1987).

A foreland stage evolved from a distal to a proximal position when Sierra de Aconquija and Ambato block

started to uplift 5 Ma ago (Sobel and Strecker, 2003). At the present is becoming a piggy back basin on top of

the active thrust fault that elevates the Guasayán Range.

In this communication we aim to identify the different deposits form Cretaceous rift to foreland and show

evidences of tectonic inversion.

Depositional history

Basement is represented by low to medium grade metamorphic rocks with strong foliation and by intrusive

granitoids (Battaglia, 1982; Mon and Hongn, 1996). This basement shows major structural discontinuities

produced during the accretion of terranes up to Early Paleozoic (Ramos, 1988). Eastwards of the Rosario Fault

there are thick series of Paleozoic rocks, absent in the Tucumán underground (Cristallini et al., 2004). These

Paleozoic rocks correspond to Silurian and Devonian detritic rocks deposited in a foreland basin during the

Ocloyic Orogeny (Ramos, 1988).

The Salta rift Basin developed during Cretaceous times, had three major arms and several sub-basins (Turner,

1959; Salfity and Marquillas, 1999). Continental red beds syn-rift deposits were overlied by onlapping lacustrine

and shallow marine post-rift deposits that infill a thermal sag basin (Bianucci et al., 1981). Post-rift deposits are

in turn overlied by Andean foreland deposits. During Middle Miocene the Atlantic Parana marine transgression

covered the area and deposited a gypsum rich sandstones (Battaglia, 1982). These deposits are outcroping along

the periphery of the Tucumán Basin (at the southern part of the Medina Range, in the western slope of Guasayán

Range and in the eastern foothills of the Aconquija Range). Are also present in the western side of the Aconquija

Range (in the El Cajón and Santa Maria Valleys, fig. 1) suggesting that uplift of the Aconquija Range postdated

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the Parana transgression (Sobel and Strecker, 2003). The stratigraphic record is completed by Pliocene to

Quaternary clastic taphrogenic sequences which are coeval to the uplift of the Pampean Ranges.

Figure 1. A) Topographic relief shape of South America, with Gansser classifaction (1973) and study area remarked. B)

Salta Rift syn-extensional deposits, with the location of the Tucumán Basin (modified from Salfity and Marquillas, 1999). C) Geological map of the study area over a Landsat Tm and Srtm topography, with the main structural features. The main isotime lines of the basement have been distinguished.

Tucumán Basin structure

The basin has a triangular shape (Fig. 1), bounded to the East by the N-S trending Guasayán Range. To the

northwest is limited by the Aconquija Range and Cumbres Calchaquíes with a NE-SW trending and to the

southwestern by the Ambato block with NNW-SSE trending. The basin is asymmetric with a more steeped

western border and a gentle eastern slope (Porto et al., 1982).

We present four seismic sections. They shown some features related to tectonic inversion of Cretaceous rift

structures that involve foreland deposits (Fig. 2).

Seismic section 2527 (Fig. 2A) display foreland deposits gentle dipping to the East and folded by a syncline –

anticline pair. Syn-rift strata are affected by a tree-like fault arrangement. When flattened to the bed interpreted

as the top of sag basin the geometry of the syn-rift deposits, the extensional faults and the semigraben rift

depocenter can be clearly distinguished.

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Seismic section 2542 (Fig. 2B) is a cross-section parallel to the longest axis of the basin and shows the deepest

part of the basin in the center of the section. In the right side of the image a gentle anticline is visible, located in

the hangingwall of a high angle thrust fault that during rift episode was an extensional fault and generated

subsidence.

Seismic section 1559 (Fig. 2C) is located East of Tucumán city and close to the South end of Ramada Range

(Fig. 1). The syn-rift deposits are thinner and the width of the basin is smaller than in the other sections

suggesting that this area corresponds to the North margin of the Tucumán Basin.

Seismic section 2481 (Fig. 2D) shooted on the eastern margin of the basin where foreland deposits are gently

dipping to the West and post-rift deposits are onlapping to the East (Cristallini et al., 2004). Syn-rift deposits

shows strong lateral thickness variations related to an extensional fault system and to a semigraben

configuration. Tectonic inversion slightly affected the two most eastern extensional faults but is almost null on

the main extensional faults.

Figure 2. Four seismic lines, with structural and depositional interpretation.

Evolution of the Tucumán Basin and conclusions

The Tucumán Basin experienced several stages of development. The absence of Paleozoic deposits suggests

that at this time the area has been an structural high. During Cretaceous times extension affected the area related

to the the Salta rifting event (Porto et al., 1982, Salfity and Marquillas, 1999). During Paleogene post-rift related

thermal subsidence generated a sag basin. Later on, the area became a foreland related to Andes uplift. This

foreland basin was segmented when surrounding rangers began to uplift, becoming a piggy-back basin.

Shortening produced inversion of previous extensional faults and folding of the sedimentary cover.

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Basement, syn-rift, post-rift and foreland strata can be differentiated in the seismic sections by its structural

architecture. Extensional and compressional features are clearly visible, evidences of slight reactivation of older

extensional faults during posterior compression can be observed in the four examples (Fig. 2). Andean orogeny

partially inverted Cretaceous rift extensional faults which are recognized by high angle geometries and more

thickness of the syn-rift deposits in the hangingwall. Incipient tectonic inversion is documented from seismic

section and constitutes a relevant feature of the Tucumán Basin. Moreover, inversion structures are also clearly

noticed in surface, as in the Medina Range (Fig. 1) where syn-rift deposits are thrusting over post-rift and

foreland deposits.

Acknowledgements

This research is supported by the following projects: 2005-00397SGR de la Generalitat de Catalunya and Consolider-Ingenio 2010 programme (CSD2006-004 “Topo-Iberia”) and CGL2007-66431-C02-01/BTE (Modelización Estructural 4D) del Ministerio de Educación y Ciencia. The AlBan scholarships program, for funding the first author. Yanina Basile, Tomás Zapata from Repsol-YPF and Ernesto Cristallini of Universidad de Buenos Aires for facilitate the seismic information. References Allmendinger, R. W., Ramos, V. A., Jordan, T. E., Palma, M., and Isacks, B. L., 1983, Paleogeography and Andean

structural geometry, northwest Argentina: Tectonics 2, 1-16. Battaglia, A. C. (1982), Descripción geológica de las Hojas 13f, Río Hondo, 13g, Santiago del Estero, 14g, El Alto, 14h,

Villa San Martín, 15g, Frías. Provincias de Santiago del Estero, Catamarca y Tucumán. Serv. Geol. Nac., Buenos Aires, Argentina. 569 186, 80.

Bianucci, H. A., O. M. Acevedo, and J. J. Cerdán, 1981, Evolución tectosedimentaria del Grupo Salta en la subcuenca Lomas de Olmedo (Provincias de Salta y Formosa): VIII Congreso Geológico Argentino (San Luis) Actas 3, 159– 172.

Cristallini, E.O., Comínguez, A., Ramos, V., Mercerat, E.D. (2004). Basement Double-wedge Thrusting in the Northern Sierras Pampeanas of Argentina (27ºS) - Constraints from Deep Seismic Reflection. In: K.R. McClay, (ed.): Thrust tectonics and hydrocarbon systems. AAPG Memoir 82: 1-26.

Gansser, A. (1973), Facts and theories on the Andes, Journal of the Geological Society London 129, 93– 131. González Bonorino, F., 1950. Algunos problemas geológicos de las Sierras Pampeanas. Revista de la Asociación Geológica

Argentina 5(3), 81-110. Jordan, T.E.; Alonso, R.N. (1987). Cenozoic Stratigraphy and Basin Tectonics of the Andes Mountains, 20º-28º South

Latitude. The American Association of Petroleum Geologists Bulletin .71: 49-64. Porto, J., C. Danieli, and O. Ruíz Huidobro, 1982, El grupo Salta en la provincia de Tucumán, Argentina: 58 Congreso

Latinoamericano de Geología (Buenos Aires) 4: 253–264. Ramos, V. A., 1988, Tectonics of the Late Proterozoic– Early Paleozoic: a collisional history of Southern South America:

Episodes 11: 168–174. Ramos, V. A., 1999, Las provincias geológicas del territorio argentino, In R. Caminos, (ed.): Geología Argentina: Instituto

de Geología y Recursos Minerales, Anales 29(3): 41–96. Salfity, J.A. and Marquillas, R.M., 1994. Tectonic and sedimentary evolution of the Cretaceous–Eocene Salta Group Basin,

Argentina. In: Salfity, J.A. (ed.), Cretaceous Tectonics of the Andes. Braunschweig/Wiesbaden, Earth Evolution Sciences, Friedr. Vieweg and Sohn, 266–315.

Sobel, E., and Strecker, M.R., 2003, Uplift, exhumation and precipitation: Tectonic and climatic control of late Cenozoic landscape evolution in the northern Sierras Pampeanas, Argentina: Basin Research 15 (4): 431–451.

Turner, J.C.M., 1959. Estratigrafía del cordón de Escaya y de la sierra de Rinconada (Jujuy). Revista de la Asociación Geológica Argentina, 13: 15-39.

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Mesozoic backarc basins in western Peru: A brief summary

Javier Jacay1, Verónica Alejandro

1, Adán Pino

2, & Thierry Sempere

3

1 Universidad Nacional Mayor de San Marcos, EAP Ingeniería Geológica, Av. Venezuela Cd. 34 s/n., apartado

3973, Lima 100, Peru (jjacayh@unmsm.edu.pe; j_jacay@yahoo.com) 2 Universidad Nacional Jorge Basadre Grohmann, EAP Ingeniería Geológica-Geotecnia, Tacna, Peru

3 LMTG, Université de Toulouse, CNRS, IRD, OMP, 14 avenue Edouard Belin, 31400 Toulouse, France

KEYWORDS : Mesozoic, backarc basin, sedimentology, paleogeography, volcanic arc.

Introduction

The history of the western edge of the Andean margin during Mesozoic times is recorded by stratigraphic

successions that were deposited in extensional backarc basins (Fig.1), such as the Arequipa basin (~11-18°S,

Early Jurassic to Early Cretaceous), Chicama basin (~7-12°S, latest Jurassic to Early Cretaceous) and Casma

basin (~8-15°, Albian-Turonian). Within these basins, the record is dominantly volcanic and volcanodetritic in

the west, due to the proximity of the volcanic arc, whereas in the east it mainly consists of clastic, carbonate, or

mixed deposits; this eastern part of the backarc behaved as a somewhat stable shelf.

Arequipa Basin

The oldest unit recorded in this basin is the dominantly volcanic Chocolate Formation (Jenks, 1948) and is best

known in the Arequipa area. The unit consists of an accumulation of andesitic, basaltic, and trachytic flows, tuffs

and agglomerates, and includes interbedded shales, sandstones, conglomerates, and carbonate rocks. Total

thickness can be over 1500 m. The uppermost part of the Chocolate Formation is intercalated by Sinemurian

limestones and is sharply overlain by Toarcian-Bajocian shelf limestones (Socosani Formation) (Vicente, 1981).

On the coast south of Arequipa, the same name is assigned to a succession consisting of flows, agglomerates

and breccias of andesitic and dacitic composition, which unconformably overlies older strata or Precambrian

gneisses (Bellido & Guevara, 1963). This succession is correlated with the Chocolate Formation of the Arequipa

area due to its stratigraphic position and lithology, and its paleontologic and isotopic ages (Roperch & Carlier,

Figure 1. Location of the Arequipa, Chicama, and Casma basins in Peru. Geographic distribution and timing of the Arequipa and Chicama basins suggest that the latter may represent a late northern extension of the former, likely to have been created through northward propagation of extensional tectonics.

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1992; Romeuf et al., 1993 and 1995). The fact that it is intruded by Hettangian to Toarcian plutons (Clark et al.,

1990; Romeuf et al., 1993) suggests that it possibly includes Triassic or older deposits (Pino et al., 2004).

The upper part of the Socosani Formation commonly displays evidence of synsedimentary extensional

tectonics in the Arequipa region (Vicente et al., 1982; Salinas, 1986). Aalenian and early Bajocian strata display

progressively deeper carbonate facies. These are sharply overlain by turbidite and other resedimented deposits

(Puente Formation, late Bajocian to Bathonian; Cachíos Formation, Callovian); large turbiditic fans accumulated

in a trough parallel to the margin, with a source in the NW and paleocurrents towards the SE (Vicente et al.,

1982). Submarine fan facies are characteristically siliciclastic. Turbidites are thicker in the lower part of the

Puente Formation. In southernmost Peru (Tacna area), deep facies are recorded as early as the Toarcian and

include radiolarites and bedded chert (Salinas, 1985) with extremely low depositional rates (Pino et al., 2004).

Facies grade upwards into distal lobe facies and in turn into proximal distributary channels (basal part of the

Cachíos Formation) (Vicente et al., 1982). The shallowing-upward Cachíos Formation exhibits a higher

proportion of shales, includes slides and olistolites, and was deposited on a continental slope. Facies include

black shales and progressively grade into prodelta or distal platform facies. The Cachíos Formation grades into

the also shallowing-upward Labra Formation (Oxfordian-Kimmeridgian), which consists of prograding

quartzitic sandstones with subordinate shale intercalations that were deposited in a variety of deltaic sub-

environments. The unit is thick and grades upwards into the Murco Formation, which was deposited in a deltaic

plain environment in the Early Cretaceous.

Chicama Basin

No stratigraphic record older than Tithonian can be observed in the Chicama basin. The basin deepened

markedly in the Tithonian, as recorded by black shales and associated conglomeratic beds in the uppermost

Simbal Formation (the oldest outcropping unit) (Jacay, 1992). The overlying Chicama Group includes

submarine-fan facies of late Tithonian age (Geyer, 1983, Enay et al., 1996), which consist essentially of

volcaniclastic turbidites (Punta Moreno Formation) that reworked the nearby Colán magmatic arc. Paleocurrents

are toward the SE or SW, with turbiditic lobes prograding southwards, in a longitudinal trough deepening to the

SSE.

In the north (Cascas-Ascope-Compatición), thick conglomerates and volcaniclastic sediments were deposited

in proximal submarine fan systems. In the south (Simbal, Tanguche, Santa river) facies are typical of middle fan

- suprafan lobes, grading to distal turbidites. Slope facies occur in the uppermost Punta Moreno Formation and

are characterized by alternations of black shales and volcaniclastic sandstones containing plurimetric limestone,

sandstone and gabbro olistoliths, in association with submarine slides, channels and contourites; facies are

overall suggestive of sedimentation on an unstable slope maintained by tectonic activity. The overlying Sapotal

Formation consists of black shales interbedded with fine-grained, thinly-bedded sandstones that were deposited

in prodelta or confined environments. The overlying latest Tithonian-Berriasian Tinajones Formation consists of

shales and shallow-marine to fluvial sandstones, and records the progradation of deltaic and coastal facies fed by

continental erosion to the east.

Coeval activity of the volcanic arc is documented by the Puente Piedra Group, which includes the Ancón,

Piedras Gordas, Puente Inga and Ventanilla formations, and mainly consists of agglomerates, lava flows, tuffs

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and pyroclastic breccias. The products of erosion of the corresponding volcanic arc were deposited near the arc,

prograding into the basin: the Puente Inga and Ventanilla formations consist of volcaniclastic sandstones and

tuffaceous shales that form turbiditic sequences, as well as limestones and intercalated andesitic flows. Fossils

reported throughout this sequence are more abundant within the tuffaceous shales of the Puente Inga Formation

and are assigned a Tithonian age (Bulot & Jacay, in press); however a Berriasian age cannot be ruled out for the

upper levels of the Puente Piedra Group.

Casma Basin

The western margin of Peru underwent major subsidence starting in the early Albian (e.g., Atherton & Webb,

1989). The resulting Casma Basin is characterized by volcanic rocks in the west and platform deposits in the

east. Relatively abundant fossils occur throughout the stratigraphic succession in both regions, indicating that the

basin was active from the early Albian to the Turonian. It is interesting that initiation of the Casma basin

coincided or closely followed the basal Albian sedimentary discontinuity (Bulot & Ferry, 2007).

The Casma Group forms a dominantly volcanic succession, which best crops out west of the Coastal Batholith

(e.g., Guevara, 1980; Atherton et al., 1985; Atherton & Webb, 1989; Aguirre et al., 1989; Soler, 1991). Volcanic

facies include pillow-lavas and volcanic breccias with few sedimentary interbeds. The middle part of the unit,

however, includes finely-bedded shales, that are intercalated with volcaniclastic sandstones in turbiditic

sequences; laterally these facies are associated with thick chert beds which suggest that the basin reached

significant depths. Features such as volcanidetritic and/or carbonate resedimentation, as well as progressive

unconformities, suggest high unstability, possibly due to recurrent extensional tectonics.

To the east, platform and slope facies can be recognized and are widely distributed among 7°S to 15°S. The

former include the Chulec, Pariatambo and Jumasha formations (but stratigraphic nomenclature varies as it

depends on the area). The Arahuay Formation was deposited on the slope of this carbonate platform. The early

Albian Chulec Formation consists of a monotonous marl and marly-limestone succession that is very

homogeneous in the north and center of the Peruvian Andes and has yielded a wealth of open-sea fauna; it is rich

in organic matter and some levels have a strong hydrocarbon odour. The middle Albian Pariatambo Formation

represents an Albian anoxic event: facies consist of black, ammonite-rich, laminated, bituminous limestones that

were deposited in an euxinic platform environment of regular depth (Jaillard, 1990). The mid-Albian to late

Turonian Jumasha Formation (Benavides, 1956; Jaillard, 1990) conformably but sharply overlies the Pariatambo

Formation, marking a change in lithology and sedimentary environment; this thick gray carbonate succession

can be divided into three sequences: the Lower Jumasha consists of 0.5 to 1 m-thick limestone beds that are

interbedded with thin black marl levels associated with chert; the Middle Jumasha consists of a thick limestone

unit which locally includes mass-wasting bodies; the Upper Jumasha consists of thin limestone beds interbedded

with grey marl.

Summary

Observation of the current distribution and timing of the Arequipa and Chicama basins (Fig. 1) suggests that

they probably formed one same backarc basin, the difference in age being only apparent due to the fact that no

strata older than Tithonian crop out in the latter. The initiation of the Arequipa Basin involved voluminous

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volcanism (Jenks, 1948; Sempere et al., 2002; Pino et al., 2004), and it evolved over 70 Myr. Maximum

depths were reached at about 180 Ma (Tacna) to about 170 Ma (Arequipa), and at about 148 Ma (Chicama),

suggesting that extension along the margin was markedly diachronous and progressed from south to north.

Similarly, the basin was invaded by siliciclastic material from the east from about 160 Ma in the Arequipa

region, but only from about 145 Ma in the Chicama region.

Development of the Casma Basin represented the most considerable episode of lithospheric thinning recorded

along the Peru margin (e.g., Atherton & Webb, 1989; Sempere et al., 2002). It was accompanied by voluminous

volcanism (Soler, 1991) but lasted only about 20 Myr.

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Atherton M. P. & Webb S. (1989) Volcanic facies, structure and geochemistry of the marginal basin rocks of central Peru, J. South Amer. Earth Sci., v. 2, p. 241-261.

Bellido, E. y Guevara, C. (1963) Geología de los Cuadrángulos de Punta Bombón y Clemesi. Carta Geológica Nacional, Lima, p. 92.

Bulot L.G. & Ferry S. (2007). La discontinuité albienne à l’echelle globale et ses implications paléobiogéographiques et biostratigraphiques. In: Bulot L.G., Ferry S. & Grosheny D. (eds), Relations entre les marges septentrionale et méridionale de la Téthys au Crétacé. Carnets de Géologie, Brest, mémoire 2007/2, résumé 11, p. 56-59.

Clark, A.H., Farrar, E., Kontak, D.J., Langridge, R.J., Arenas, M.J., France, L.J., McBride, S.L., Woodman, P.L., Wasteneys, H.A., Sandeman, H.A. & Douglas, D.A. (1990) Geologic and geochronologic constraints on the metallogenic evolution of the Andes of southeastern Perú. Economic Geology, v. 85, p. 1520-1583.

Enay R., Barale G., Jacay J., Jaillard E. (1996) Upper Tithonian ammonites and floras from the Chicama basin, northern Peruvian Andes. GeoResearch Forum, v. 1-2, 221-234, Transtec Publ., Switzerland.

Guevara C. (1980) El Grupo Casma del Perú Central Entre Trujillo y Mala. Bol. Soc. Geol. Perú, v. 67, 73-83. Geyer O. (1983) Obertithonische Ammoniten-fauna von Peru. Zblatt Geol. Palaeont., v. 1, 335-350. Jacay J. (1992) "Estratigrafía y sedimentología del Jurásico Curso medio del Valle del Chicama y esbozo Paleogeográfico de

Jurásico-Cretáceo del Nor Perú (6 30'-8 Latitud Sur)". Tesis Ing. Geol. UNMSM, 180p. Jaillard, E. (1990) Evolución de la margen andina en el norte del Perú desde el Aptiano superior hasta el Senoniano. Bol. Soc.

Geol. Perú, v. 81, 3 - 13. Jenks, W. (1948) Geología de la hoja de Arequipa, al 1/200.000. Boletín del Instituto Geológico del Perú, v. 9, p. 104. Pino, A., Sempere, T., Jacay, J. Fornari, M. (2004) Estratigrafía, paleogeografía y paleotectónica del intervalo Paleozoico

superior-Cretáceo inferior en el área de Mal Paso-Palca (Tacna). In: J. Jacay & T. Sempere (eds.), Nuevas contribuciones del IRD y sus contrapartes al conocimiento geológico del sur del Perú, Sociedad Geológica del Perú, Publicación Especial 5, 15-44.

Romeuf N., Aguirre L., Carlier G., Soler P., Bonhomme M., Elmi S. & Salas G. (1993) Present knowledge of the Jurassic volcanogenic formations of southern coastal Perú. II International Symposium on Andean Geodynamics, Oxford, p. 437-440.

Romeuf N., Aguirre L., Soler P., Féraud G., Jaillard E. & Ruffet G. (1995) Middle Jurassic volcanism in the Northern and Central Andes. Revista Geológica de Chile, v. 22, p. 245-259.

Roperch P. & Carlier G. (1992) Paleomagnetism of Mesozoic rocks from the Central Andes of southern Perú: Importance of rotations in the development of the Bolivian Orocline. Journal of Geophysical Research, v. 97, B12, p. 17233-17249.

Salinas, E. (1985) Evolución paleogeográfica del sur del Perú a la luz de los métodos de análisis sedimentológicos de las series del departamento de Tacna. Universidad Nacional San Agustín de Arequipa, Tesis de grado, 205 p.

Sempere T, Carlier G, Soler P, Fornari M, Carlotto V, Jacay J, Arispe O, Néraudeau, Cárdenas J, Rosas S, Jimenez N (2002) Late Permien-Middle Jurassic lithospheric thinning in Peru and Bolivia, its beraing on Andean-age tectonics. Tectonophysics, v. 345, p. 153-181.

Soler P. (1991) El volcanismo Casma del Perú Central: cuenca marginal abortada o simple arco volcánico? Volumen de resúmenes del VII Congreso Peruano de Geología, p. 659- 663.

Vicente J.-C., Beaudoin B., Chávez A., León I. (1982) La cuenca de Arequipa (Sur Perú) durante el Jurásico-Cretácico inferior. 5th Cong. Latinoamer. Geol., Buenos-Aires 1981, v. 1, 121-153.

Vicente J.-C. (1981) Elementos de la Estratigrafía Mesozoica Sur-Perúana. In: W. Volkheimer & E.A. Musacchio (eds.), Cuencas sedimentarias del Jurásico y Cretácico de América del Sur. Comité Sudamericano del Jurásico y Cretácico, Buenos Aires, v. 1, p. 319-351.

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Geometric reconstruction of fault-propagation folding: A case study in the western Cordillera Principal at 34º15’S-34º30’S

Pamela Jara1,*, Reynaldo Charrier

1, Marcelo Farías

1,2, & César Arriagada

1

1 Departamento de Geología, FCFM, Univ. de Chile, Plaza Ercilla 803, Santiago, Chile

2 Departamento de Geofísica, FCFM, Univ. de Chile, Blanco Encalada 2002, Santiago, Chile

* presenting author (pajara@ing.uchile.cl)

KEYWORDS : Central Chile Andes, structural reconstruction, volcanic rocks, Trishear models, Cenozoic evolution

Introduction

The Central Chile Principal Andes is formed by Cenozoic and Mesozoic sequences intruded by several

Miocene-Pliocene plutonic bodies. The Cenozoic series mainly corresponds to volcanic and volcano-clastic

deposits with some sedimentary intercalations grouped in the Abanico Fm. (late Eocene to late Oligocene/Early

Miocene) and the Miocene Farellones Fm. The Mesozoic sequences mainly correspond to marine and

continental sedimentary deposits (Fig. 1). The structural features observed in the Cenozoic units have been

attributed to thick-skinned deformation controlled by the inversion of normal faults associated with the

development of the late Eocene to late Oligocene extensional Abanico Basin (Charrier et al., 2002, 2005).

Fig. 1: Simplified geological map of Central Chile and Western-Central Argentina. Compiled by Farias et al. (2008).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 269-272

270

Structural models and balanced cross-sections in the Andean mountain belt have been mostly constructed in

regions where shortening was accommodated in predominantly sedimentary successions in the backarc. Such

successions allow detailed mapping and further geometrical reconstructions. Almost no balanced cross-sections

and geometric models have been attempted in the mainly volcanic Cenozoic rocks present in the forearc because

of the difficulty to apply these methods to deposits lacking continuous key-layers and the presence of large

batholits. In this contribution, we present a structural model based on field observations, presenting six W-E

oriented geological cross-sections made on deformed Cenozoic volcanic rocks of the western Principal

Cordillera that record the late Cenozoic mountain building in Central Chile.

Structure: Geometric reconstruction and Trishear model

The structure framework in the study region corresponds to a 2 km-wavelength syncline (“Alto de Los

Peñascos” syncline in Fig. 2), which is bounded by faults. We reconstructed the described structure by the

prolongation in depth of the surface data using the kink-band method (Fig. 2).

Considering that both sides of the

syncline correspond to anticlines, the

fault-propagation folding should have

occurred with a considerable

displacement capable of transport the

anticlines along the faults and cutting

their crests along the axial plane as an

anticlinal breakthrough deformation

mode (Suppe and Medwedeff, 1990).

The “broken crests” would correspond

to the faults bounding the described

syncline (Western and Eastern Fault,

hereafter WF and EF, respectively),

thus the presently layers dip is

interpreted as the axial orientation of

the broken anticline (Fig. 2). The axial angle was obtained by the means of the graphs for fault-propagation

folding of Suppe and Medwedeff (1990). After the geometric reconstruction the faults were interpreted with an

eastern vergence and with a cut-off angle of ~60ºW for the WF and ~20º for the EF. Although the resulting

shortening is difficult to quantify because of the lack of key-horizons and the unknown total thickness or the

initial length of the deformed section, a minimum of shortening could be measured by considering the difference

between the horizontal length and the inclined flanks of the interpreted folds using the axial angles (20-30%

shortening, i.e., ~2 km)

Using the results for the geometrical model (Fault angle in Fig. 2), we constructed a Trishear model (Erslev,

1991; Hardy and Ford, 1997; Zehnder and Allmendinger, 2000) with the TRISHEAR 4.5.4 Program

(Allmendinger, 1997, 2003). Based on the construction of a non-deformed “trishear” section with the observed

proportional thickness and utilizing the cut-off angles for the fault obtained using the geometric model of

Fig. 2: Geometric reconstruction of structure using the fault-propagation folding mechanism.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 269-272

271

deformation (WF = 60º and EF = 20º), we developed a forward model varying the propagation, slip, and

Trishear angle in order to arrive to the best-fit parameters explaining the deformation caused by the WF. After a

comparative analysis between the resulting trishear models and the observed deformation, we choose a cut of

angle of 60º, a Trishear angle of 30º, a slip of 150 pixels (that corresponds to 1.5 km) and a P/S of 1 as the best-

fitting parameters. Using the resulting geometry of the model with these parameters, we superimpose the activity

of the EF; the better resulting geometry corresponds to a P/S of 0.5, a Trishear angle of 60º and a slip of 200

pixels (2 km). The resulting geometry was superimposed to the faults geometric model reconstruction in order to

compare them (Fig. 3). The total shortening of the deformed trishear section correspond to 30% (~3 km), which

is consistent with the shortening obtained using the geometric reconstruction.

Fig. 3: Superimposed faults geometric model reconstruction and Trishear model results.

Discussion and conclusions

The results for the geometrical and Trishear models suggest that folding in the study region would be at least

controlled by two E-vergent faults (WF and EF). The most important one, the WF, has a higher dip angle

(~60ºW) than the EF, being mechanically compatible with the inclination expected for inverted normal faults.

Moreover, this is also consistent with the regional geology interpreted for the faults affecting the Cenozoic units

in the western Principal Cordillera (Charrier et al., 2002). Based on the resulting geometric model, we interpret

the EF (with a dip angle near 20ºW) as a neo-formed short-cut, which would have facilitated shortening

accommodation.

The estimated shortening using these models is about 30% (2-3 km for the study region) which would have

been accommodated during the basin inversion stage (22 and 16 Ma) because the upper levels of the Farellones

Formation remain regionally non-deformed. In the northernmost part of the study region, the axial plane of the

WF develops as an anticline. This fold extends further north, thus the WF would represent a major structure of

the western flank of the Principal Cordillera. Indeed, this anticline is the western boundary of the Farellones

Formation (FW in Fig. 4) in which its lower portion develops growth strata (Fock et al., 2006; Farías et al.

2007). According to Farías et al. (2007), this east-vergent fault has a deep origin associated with a regional

ramp-flat structure connecting the subduction zone with the tectonic front of the Andes in Argentine territory

(Fig. 4). In this structural context, the WF would be pass-by thrust of the ramp segment in the zone where this

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 269-272

272

crustal-scale structure evolved to a detachment fault eastward at a depth close to 20 km beneath the western

Principal Cordillera.

Finally, the use of geometrical and numerical models has allowed proposing a structural kinematic model for

the deformation affecting volcanic rocks in the forearc with a very-well correlation with geophysical and

regional structural data of this Andean segment.

Fig. 4: Structural cross section at 34ºS. Modified after Farías et al. (2008).

Aknowledgements This work was supported by the FONDECYT Project Nº1030965 and the Bicentennial Program in Science and Technology Grant ANILLO ACT-18. References Allmendinger, R., 1998. Inverse and forward numerical modeling of trishear fault-propagation folds. Tectonics, Vol. 17, Nº

4, p. 640-656. Allmendinger, R. 1997-2003. FaultFold 4.5.4 (formerly TRISHEAR, Windows version). Charrier, R., Baeza, O., Elgueta, S., Flynn, J.J., Gans, P., Kay, S.M., Muñoz, N., Wyss, A.R., Zurita, E., 2002. Evidence for

Cenozoic extensional basin development and tectonic inversion south of the flat-slab segment, southern Central Andes, Chile (33º-36ºS.L.). Journal of South American Earth Sciences, Vol. 15, p. 117-139.

Charrier, R., Bustamante, M., Comte, D., Elgueta, S., Flynn, J.J., Iturra, N., Muñoz, N., Pardo, M., Thiele, R. and Wyss, A.R., 2005. The Abanico Extensional Basin: Regional extension, chronology of tectonic inversion, and relation to shallow seismic activity and Andean uplift. Neues Jahrbuch für Geologie und Paläontologie Abh. 236, (1-2), p. 43-47.

Erslev, E., 1991. Trishear fault-propagation folding. Geology, v 19, p.617-620. Farías, M., Comte, D., Charrier, R., Martinod, J., Tassara, A., and Fock, A. (2007). Ramp-flat crustal-scale structure as the

first order feature of the Andean margin: Seismologic, surface structural and rheological evidence for Central Chile. GEOSUR 2007, Libro de Resúmenes, Santiago de Chile, p. 59.

Farías, M., R. Charrier, S. Carretier, J. Martinod, A. Fock, D. Campbell, J. Cáceres, and D. Comte (2008), Late Miocene high and rapid surface uplift and its erosional response in the Andes of central Chile (33°–35°S), Tectonics, 27, 17 January 2008.

Fock, A., Charrier, R., Farías, M., Muñoz, M. (2006). Fallas de Vergencia Oeste en la Cordillera Principal de Chile Central (33º S- 34º S). Asociación Geológica Argentina, Serie: Publicación especial, 6, 48-55.Hardy, S. y Ford, M., 1997. Numerical modeling of trishear fault propagation folding. Tectonics, Vol. 16, Nº 5, p. 841-854.

Suppe, J and Medwedeff, D., 1990. Geometry and kinematics of fault-propagation folding. Eclogae geol. Helv. 83/3:409-454.

Zehnder, A.T. and Allmendinger, R.W., 2000. Velocity field for the trishear model. Journal of Structural Geology 22.1009±1014

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273

Organization and evolution of a segmented deformation front: Llanos foothills, Eastern Cordillera of Colombia

Andreas Kammer, Antonio Velásquez, Alejandro Beltran, Alejandro Piraquive, & Wilmer

Angel Robles

Universidad Nacional de Colombia, 14490 Bogota, Colombia (akammer@unal.edu.co)

Introduction

The Eastern Cordillera of Colombia evolved since the Upper Eocene within the Cretaceous to Paleocene

retroarc basin of the Early Mesozoic Northandean plate margin and designates, by its isolated position, an

intracontinental chain. Its initial folding closely correlates to the onset of convergence between Caribbean and

Southamerican plates and its still ongoing evolution marks a long lasting collisional event. On its western side

this mountain chain is embraced by the gently tilted western flank of the Central Cordillera, while on its eastern

side it abuts against the rigid lithosphere of the Guayana shield. By its clear-cut deformation fronts it defines a

weak crustal welt between converging rigid blocks as conceptualized in the experimental vice model (Ellis et al.,

1998). This same belt was previously the site of various rift events, among which an Upper Paleozoic and an

Early Cretaceous one bear a particular importance on the evolution of its deformation fronts. Its attenuated crust

was therefore predestined to accrue much of the Tertiary contraction by the reactivation of inherited faults that

originated at its boundaries and a pure-shear deformation of its inner parts. On its eastern border the Tertiary,

especially Neogene contraction and exhumation produced uplift in excess of 10’000 m with pre-Cretaceous

basement units reaching altitudes of 3000 to 4000 m. Uplift in the central high plain of Bogotá, on the other

hand, is constrained to less than 2500 m. These contrasting structural reliefs pose the question about deformation

mechanisms intervening in the formation of the Eastern Cordillera. In his contribution we focus on the evolution

of the deformation front east of the High plain of Bogotá comprised between latitudes of 4oN and 4.5oN.

Deformation front east of the High Plain of Bogotá

The oblique view on the study area (Fig. 1) depicts a synthetic view on the Guavio and Tauramena segments of

the Llanos foothills and illustrates the composite nature of its deformation front. The folded and principal

deformation front is represented by the Farallones Anticline that is breached by a normal fault in its hinge. This

giant anticlinal structure is tightly coupled to the blind Servitá fault that limits to its W a rift basin of Late-

Paleozoic red beds. East to the Farallones Anticline Late Cretaceous platform and Tertiary foreland sediments

are transported in a piggy-back manner to the E, as attested by the emergent Guaicaramo fault, which limits

folded foothills against undeformed shield area. The Guavio segment lacks a seismic activity and, considering

the high activity of adjacent segments, an aseismic creep may be anticipated for the slip on the Guaicaramo

thrust. At the southern border of the Tauramena segment the folded deformation front steps E to the thrust front

of the Guaicaramo fault. The Guacaramo fault is here constrained to a dip angle of 50o by an aftershock

distribution of the 1995 Tauramena earthquake (Dimaté et al., 2003). How can this uncommon situation of the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 273-276

274

merging of a high-angle and a thrust fault be explained? To this end we analyze the structural evolution of the

two segments by means of serial cross-sections (Fig. 2).

Figure 1. Synthetic map and DEM of the study area with main structural elements and traces of vertical sections of Fig. 2.

Figure 2. Serial cross sections.

Structures of the Guavio segment

The generation of about 6 km of displacement on the Guavio thrust has been accounted for by various

scenarios, among which a thin-skinned solution (1) explains the excess length of the cover as a consequence of a

strain-partitioning between highly deformed basement and little deformed cover. In solutions focusing on an in-

850

1080

1400

Aguacalara – La Mesa Bridge

– San Agustinera Creek

Guavio Anticline

Pescana Creek

Chameza

Guavio Segment

Tauramena

Segment

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 273-276

275

situ deformation (2), the Guavio Anticline is interpreted as a fault-bend fold (Rowan and Linares, 2000) or a

fault-bounded pop-up (Branquet et al., 2002) and the Guaicaramo thrust roots thus in the basement underlying

the Llanos foothills. Finally, in an in-sequence scenario (3) the Guaicaramo thrust represents, among other

possible reverse faults, a front-most splay that ultimately merges into a single mid-crustal master fault, which

transported the Farallones Anticline to the E (Mora et al., 1979).

Shortcomings of these scenarios are the unrecognized change in structural style that would localize a

detachment (1), missing criteria for the recognition of a (reactivated) fault that would justify the supposition of a

fault-related anticline in the foothills (2) and the steeply inclined attitude of the Servitá fault which precludes its

correlation with a low-angle fault (3). Instead, we posit an out-of-sequence thrust which displaced the eastern

flank of the Farallones Anticline during a late deformation stage. This model is substantiated by gravimetric

surveys which identified positive anomalies coinciding with the Farallones and the Guavio anticlines, but a

negative intermediate anomaly shifted to the W with respect to the intervening Nazareth syncline. This

discrepancy can be modeled by a thrust sedimentary wedge, as required by the hypothesis of a breached

deformation front. (Fig. 3).

Figure 3. Modeled structural section of the Guavio segment. Observed (resp. calculated) gravity is indicated by green (resp. red) line.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 273-276

276

Structures of the Tauramena segment

The right-stepped relay from the Guavio to the Tauramena segment is smooth and manifests itself in the

hangingwall block by a loss of structural relief of the Farallones Anticline and an amplification of the easterly

Tierra Negra Anticline which supersedes the Guavio Anticline to the N. Most important, however, is the tightly

folded tip of the hangingwall block at the fault intersection. These small-scale buckle folds may be fault-

bounded along their synclinal axial planes. They designate deformation zones that predate the final throw on the

Guaicaramo fault, as they are partially delaminated from the hangingwall block array by its frictional interaction

with the footwall block.

Discussion

The Pliocene Andean paroxysmal deformation phase is characterized by an accelerated uplift and an

exhumation rate estimated at 1mm/year (Mora2007) that focused particularly on the eastern deformation front. In

the Guavio segment amplification and outward growth of the folded deformation front are antagonistic

processes. Fold growth guided by the steeply inclined Servitá fault requires material to be added along steep

flow paths, perhaps linking the marginal high to a mid-crustal flow channel. Fold growth implied a steepening of

the eastern flank which evolved toward a critical angle, before it was breached by the Guavio fault. This

Pliocene (?) out-of-sequence thrust fossilized this folded deformation front by a displacement transfer to an

easterly more position. Contrary to this big-scale folding the deformation front of the Tauramena segment did

not become arrested at the Servitá fault and affected progressively (?) more foreland oriented areas, until it

became blocked at the deformation zone at the Guaicaramo fault.

The Guavio Anticline marks a transitional behavior between arrested and outward migrating deformation

front. It did not attain the stage of a faulted forelimb, as further north in the Tierra Negra Anticline. Collapse of

the folded deformation front attenuated the compressional regime.

References Branquet Y., Cheilletz A., Cobbold P. R., Baby P., Laumonier B., and Giuliani G., 2002, — Andean deformation and rift

inversion, eastern edge of Cordillera Oriental (Guateque-Medina area), (Colombia). Journal of South American Earth Sciences., v. 15, no. 4, p. 391-407.

Dimaté, C., Rivera, L. A., Taboada, A., Delouis, B., Osorio, A., Jiménez, E., Fuenzalida, A., Cisternas, A., and Gómez, I., 2003 — The 19 January 1995 Tauramena (Colombia) earthquake: geometry and stress regime: Tectonophysics., 363, no. 3-4, 159-180.

Ellis, S., Beaumont, C., Jamieson, R.A., Quinlan, G., 1998 — Continental collision including a weak zone; The vice model and its application to the Newfoundland Appalachians. Canadian Journal of Earth Sciences, 35, 1323-1346.

Mora, A., 2007 — Inversion tectonics and exhumation processes in the Eastern Cordillera of Colombia. Doctoral thesis, University of Potsdam, 133 p, Potsdam.

Rowan, M., and Linares, R., 2000 — Fold evolution matrices and axial-surface analysis of fault-bend folds: Application to the Medina Anticline, Eastern Cordillera, Colombia: American Association of Petroleum Geologists Bulletin, v. 84, no. 6, p. 741-764.

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The PUNA passive seismic array in the southern Puna: Tests of lithospheric delamination in the region of the Cerro Galán ignimbrite

S. Mahlburg Kay1, B. S. Heit

2, B. L. Coira

3, E. Sandvol

4, X. Yuan

2, N. A. McGlashan

1,

D. Comte5, L. D. Brown

1, R. Kind

2, & D. Robinson

4

1 Dept. Earth. Atm. Sci., INSTOC, Cornell Univ. Ithaca, NY, 14853, USA (smk16@cornell.edul)

2 GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany (yuan@gfz-potsdam.de)

3 Inst. Geología y Minería, Univ. Nac. de Jujuy, 4600 S.S. de Jujuy, Argentina (bcoira@idgym.unju.edu.ar)

4 Dept. Geoi. Sci., Univ. of Missouri, Columbia, MO, 65211-1380, USA (sandvole@missouri.edu)

5 Dept. Geofísica, Universidad de Chile, Santiago, Chile (dcomte@dgf.uchile.cl)

KEYWORDS : southern Puna, delamination, Central Andes, Cerro Galán, seismic data

Introduction

A passive seismic array is deployed at present in the southern Puna of the Central Andean plateau between

25°S to 28°S latitude in Argentina and Chile to address fundamental questions on the processes that form,

modify and destroy continental lithosphere and control lithospheric dynamics along Andean-type continental

margins (Fig. 1). The southern Puna is important in this regard as this is the region where the delamination

hypothesis for removal of thickened eclogitic continental crust and mantle lithosphere was initially suggested

(Kay and Kay, 1993). The case for young delamination of the lithosphere in the southern Puna was built on

magmatic patterns, geochemical signatures and evolutionary models for mafic lavas and silicic ignimbrites

(particularly Cerro Galán), changing and mixed deformational styles, high topography accompanied by

insufficient crustal shortening, an underlying slab with a gap in intermediate depth seismicity and evidence for

Sn attenuation (e.g., Kay et al. 1994; 1999; Whitman et al. 1996). Since then, the crustal delamination process

has gained popularity as it can explain features like formation of giant ignimbrites and the near absence of a

mafic crustal root in many orogenic regions (e.g., Beck and Zandt, 2002; Yuan et al., 2002). On another level,

the delamination model provides a way to explain the bulk andesitic composition of the continental crust (e.g.,

Kay and Kay, 1993). The southern Puna, which overlies a down-going slab with a transitional dip between a

steeper dipping segment to the north and a flat-slab to the south, is a benchmark for comparative studies with

orogenic systems like those in the western US and Tibet where changing subduction angles, lithospheric

delamination and crustal shortening have been used to explain lithospheric evolution and plateau formation.

The seismic experiment

The seismological data being acquired in the southern Puna are designed to test if the current lithospheric and

mantle structure is compatible with latest Miocene and younger crustal and mantle lithospheric delamination and

to fill a critical gap in along strike geophysical coverage in the central Andes. The seismological data

complement those of Yuan et al. (2002), Heit (2005), Schurr et al (2006) and others. They should provide the

first detailed constraints on crustal thicknesses, slab dip and mantle heterogeneity from receiver function and

tomographic images under the southernmost Puna. The region being studied underlies the latest Miocene to

Quaternary southern Puna mafic volcanic centers, the 6 to 2 Ma Cerro Galán ignimbrites, and < 1 Ma Cerro

Blanco caldera and associated ignimbrites (e.g., Sparks et al., 1985; Kay et al., 1994; 1999; Siebel et al., 2001).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 277-280

278

The project involves 73 three-component broad and intermediate band seismic stations deployed in arrays

designed to image the seismic velocity structure of the crust and mantle as well as local seismicity patterns. The

instruments are from NSF IRIS PASSCAL and the Universities of Missouri and St. Louis in the USA, and the

GeoForschungZentrum in Potsdam, Germany. The deployment, which took place in late 2007 and early 2008, is

for approximately 18 months. The first major data recovery from the array will take place in the middle of 2008.

The location of the seismic array is shown in Figure 1.The short period stations are deployed along the north-

south (25°30’S, 67°15’W to 27°55’S, 67°38’W) and east-west (26°55’ S, 68°53’W to 27°S, 65°15’W) oriented

lines with a spacing of 10 to 15 km to optimize the imaging of crustal and upper mantle properties. The broad

band stations are deployed across the region in a 2-D array with an approximate spacing of 50 km to optimize

determination of mantle properties in three dimensions. Mantle dynamics and mantle flow associated with slab

lithosphere and geometry are to be interpreted through a combination of surface wave, body wave, and

attenuation tomographic methods in combination with existing geologic and geochemical studies.

Fundamental questions regarding Nazca-South American plate dynamics and the origin of the central Andean

Puna plateau to be addressed include: (1) What are the geometry, extent and fate of thinned or removed

lithosphere beneath the Puna plateau? (2) What controls crustal thickness beneath different parts of the Puna

plateau and surrounding regions? (3) What is the rheology of the mantle wedge and subducting slab in a region

of changing slab dip over a seismically quiet Benioff zone? (4) What is the history of deformation and

magmatism within the mantle wedge? and (5) What are the factors controlling along strike similarities and

differences in continental plateaus along Andean type margins?

The Cerro Galán ignimbrite and mantle/crustal magma generation models

being tested from seismic studies in the southern Puna

A number of the characteristic features of the Central Andean Puna plateau have been proposed to reflect

delamination and other deep crustal and mantle processes that are being investigated through the seismic

experiments. Among these features are the voluminous dacitic ignimbrites that are now commonly linked with

delamination processes as was suggested for the Cerro Galán ignimbrites by Kay et al. (1994). The volumes and

chemistry of the ignimbrites reflect the crustal and mantle conditions required to generate large magma systems.

Geochemical models based on 87Sr/86Sr ratios, 18O/16O and trace element contents show that the Cerro Galán

ignimbrites are best modeled as approximately 50:50 mixtures of mantle-wedge derived basaltic composition

magmas and crustal melts. Erupted mafic magmas and geochemical modeling studies in the Cerro Galán region

show that the mantle-derived magmas are isotopically enriched (87Sr/86Sr ~ 0.7055; Nd ~ -2) most likely though

lithospheric recycling due to forearc subduction erosion and delamination. Bulk crustal contaminants in the

Cerro Galán ignimbrites need to have 87Sr/86Sr ratios ranging from 0.720 to 0.735 at 300 to 125 ppm Sr and 18O/16O values of 11.5 to 12. Trace element evidence for residual garnet and Sr systematics in crustal

contaminants indicate that crustal melting and contamination occurred in the deep crust and that the

contaminated melts accumulated in magma chambers like those seen near 20 to 30 km depth on seismic images

further north (e.g., Beck and Zandt, 2002; Yuan et al. 2002: ANCORP, 2003) and at the northern limit of the

current array (Heit, 2005). Negative Eu anomalies superimposed on high pressure REE patterns and calculated

bulk distribution coefficients are best explained by feldspar fractionation in mid-crustal magma chambers at

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these levels. Trends from higher to lower La/Yb and Sm/Yb ratios in the 6 to 2 Ma Cerro Galán ignimbrites

indicate a changing role for garnet as a controlling residual phase as the Cerro Galán system evolved. Single

crystal 40Ar/39Ar sanidine and biotite ages for pumice and whole rock samples from three flows show that the

terminal eruptions of the giant Cerro Galán ignimbrite consisted of a series of flows that had begun erupting by

at least 2.13 Ma and continued to erupt until at least 2.06 Ma (Kay and Coira, 2008). The volume of the

ignimbrites and their compositions require major heat input from mantle-derived magmas generated in the

mantle wedge above the slab. The estimated volumes for the Puna ignimbrites coupled with a 90 km3/km/Ma

arc magma production rate suggest a plutonic/volcanic ratio of about 4:1 and up to 5 kilometers of thickness of

new crust spread under the plateau.

The tomographic images of Heit (2005) from the northern region of the current seismic deployment near

25.5°S show significantly more negative P and S wave velocity anomalies at mantle depths below 100 km than

for seismic images from beneath the northern Puna. The anomalies in the southern Puna are consistent with the

younger age of the Cerro Galán ignimbrites compared to the northern Puna ignimbrites and with recent

delamination processes in the southern Puna. The seismic results from the current seismic array will allow the

seismic character of the mantle to be traced to the south and aid in correlations of the generation of the

distinctive young southern Puna ignimbrites from Cerro Galán and Cerro Blanco and the young mafic flows with

mantle and slab processes.

Figure 1. Satellite Radar Topography Mission (SRTM) image of the southern Puna showing the location of the broadband array (in the boxed region) and the north-south and east-west seismic lines (dashed white lines). Also shown are the locations of the large northern Puna ignimbrite field near 22°S and the southern Puna Cerro Galán and Cerro Blanco ignimbrites.

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References ANCORP Working Group 2003. Seismic imaging of a convergent continental margin and plateau in the central Andes

(Andean Continental Research Project 1996 ANCORP’96). Journal of Geophysical Research 108: 2328, doi:10.1029/2002JB001771.

Beck, S. & Zandt, G. 2002. The nature of orogenic crust in the central Andes. Journal of Geophysical Research 107: 2230, doi:10.1029/2000JB000124.

Heit, B.S., 2005. Teleseismic tomographic images of the Central Andes at 21 °S and 25.5 °S: an inside look at the Altiplano and Puna plateaus. GeoForschungsZentrum Potsdam, Scientific Technical Report STR06/05 139 (Potsdam).

Kay, R. W. & Kay, S.M. 1993. Delamination and delamination magmatism. Tectonophysics 219: 177– 189. Kay, S. Mahlburg, Coira, B. & Singer, B. 2008. “Single crystal sanidine and biotite 40Ar/39Ar ages for the Cerro Galán

intracaldera and extracaldera ignimbrite flows.” In: VI South American Symposium on Isotope Geology, San Carlos de Bariloche, Argentina, 14-16 April 2008.

Kay, S. Mahlburg, Coira, B. & Viramonte, J. 1994. Young mafic back-arc volcanic rocks as indicators of continental lithospheric delamination beneath the Argentine Puna plateau, Central Andes. Journal of Geophysical Research 99: 24323-24339.

Kay, S. Mahlburg, Mpodozis, C. & Coira, B. 1999. Magmatism, tectonism and mineral deposits of the Central Andes (22°-33°S latitude). In Skinner, B. J. (ed) : Geology and ore deposits of the central Andes, Society of Economic Geology Special Publication 7: 27-59.

Schurr, B., Rietbrock, A., Asch, G. Kind, R. & Oncken, O. 2006. Evidence for lithospheric detachment in the central Andes from local earthquake tomography. Tectonophysics 415: 203-223.

Siebel, W., Schnurr, W., Hahne K, Kraemer, B, Trumbull, R.B., van den Bogaard, P & Emmermann, R. 2001. Geochemistry and isotope systematics of small- to medium-volume Neogene-Quaternary ignimbrites in the southern central Andes: evidence for derivation from andesitic magma sources. Chemical Geology 171: 213-237.

Sparks, R. S. J., Francis, P. W., Hamer, R. D., Pankhurst, R.J., O'Callaghan, L.L., Thorpe, R. S. & Page, R. S. 1985. Ignimbrites of the Cerro Galán Caldera, NW Argentina. Journal of Volcanology and Geothermal Research 24: 205-248.

Whitman, D., Isacks, B. L. & Kay, S. Mahlburg 1996. Lithospheric structure and along-strike segmentation of the central Andean Plateau. Tectonophysics 259: 29-40.

Yuan, X., Sobolev, S.V. & Kind, R. 2002. Moho topography in the central Andes and its geodynamic implications. Earth and Planetary Science Letters 199: 349-402.

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Incision and erosion of the deepest Andean canyons in southern Peru, based on ignimbrites, remote sensing, and DEM

A. de La Rupelle1, J.-C. Thouret

1, F. Albino

1, T. Souriot

1, T. Sempere

2, & Y. Gunnell

3

1 Laboratoire Magmas et Volcans, UMR 6524 CNRS, OPGC and IRD, Université Blaise Pascal, 5 rue Kessler,

63038 Clermont-Ferrand cedex, France (a.delarupelle@opgc.univ-bpclermont.fr) 2 LMTG - Observatoire Midi-Pyrénées 14, avenue Edouard Belin 31400 Toulouse, France

3 LGP - UMR 8591 CNRS and University Denis Diderot Paris 7, France

KEYWORDS : Central Andes, valley erosion, incision rate, uplift, ignimbrite

Introduction

In the northern Central Volcanic Zone in southern Peru, the deepest Andean valleys cut perpendicularly

through the WNW-ESE-trending western Altiplano, the Western Cordillera and the Coastal Cordillera. The Rios

Ocoña – Cotahuasi canyons expose a 2 to 3.5 km-deep N-S cross section 200 km long in pre-volcanic bedrock

and its ignimbrite and lava cover (Fig. 2). These old valleys in a young active orogen are used as a tool to

compute erosion volumes and incision rates, therefore contributing to the debate regarding the Central Andes

uplift history. The uplift, estimated to be 3 to 4 km over the past 24 Ma (Thouret et al., 2007) has led to a large-

angle NE-SW bulge of the range after 14 Ma. Based on low-temperature thermochronology, FT and (U/Th)/He

ages on zircon and apatite (Gunnell et al., 2008), the low erosion rate of the high plateau, estimated to be in the

order of 1 km over 60 Ma, contrasts with the high erosion rate of the canyons, estimated to be in the range

between 0.8 and 1.5 km over 13 Ma. This is based on the estimated volumetric erosion rate of 220 to

420 km3/Ma over a catchment of ca. 3400 km2. This leads to the following unsolved issue: the low erosion rate

of the high plateau (a peneplain formed between c.60 Ma and c.19 Ma) corroborates the arid climate that

prevailed in the region (Hartley, 2005), whereas the high erosion rate in valleys since 13 Ma implies a

combination of surface uplift and climate change towards less aridity (Fig. 1) reflected by a change in river

incision power.

Figure 1. Location of the study area and the present rainfall distribution (Allmendinger & the Cornell Andes Project, 2007).

Figure 2. Main geological units and faults, and active

drainage network.

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Ignimbrites as tools for estimating erosion and incision

Pre- and post-incision ignimbrites

Among the five sheets of ignimbrites that erupted in the region since ca. 20 Ma, three are used here as

stratigraphic markers of the canyon incision. The pre-incision 19.4–18 Ma-old Alpabamba ignimbrites (Thouret

et al., 2007) and the 14.3 –12.7 Ma-old Huaylillas ignimbrites overlay the paleosurface of, and form, the high

plateaus at 4000-4500 m asl. The ca 9 Ma-old Caraveli ignimbrites mantle shallow wide valleys cut in these high

plateaus, pointing therefore to the initiation of downcutting some time before 9 Ma. Two post-incision lower and

upper Sencca ignimbrites (4.9 – 3.6 Ma and 2.3 – 1.4 Ma) crop out near the valley bottom or cover terraces 400

to 600 m above the river. Based on this chronostratigraphy and on the fact that lower Sencca ignimbrites

3.76 Ma are exposed near the bottom of the valley of the Río Cotahuasi, the valley erosion started some time

before 9 Ma but the deep canyon downcutting occurred well after 9 Ma and before 3.76 Ma.

Erosion volumes and incision rates based on DEM and remote sensing

To compute the volume of eroded material during the valley formation, we use the Huaylillas ignimbrites,

which predate the valley formation, and the Sencca ignimbrites, which post-date the major phase of valley

incision. The contour base of each ignimbrite sheet has been drawn on a Landsat image (Fig. 3) and mapped on a

DEM, and is used as a paleo-surface, accounting for the paleo-topography of the valley at each given time span

of the stepwise incision.

Figure 3. Contours of bases of each ignimbrites and volcanic rocks, drawn on Landsat image of study area.

N

10 km

Figure 4. DEM (20 m- resolution), with shaded relief, from SPOT images.

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Using ENVI, the contours of volcanic units (Fig. 3) and a DEM (Fig. 4), the paleosurfaces at each ignimbrite-

marker are compared with the present valley bottom (Fig.5), and we compute two data sets: the volume of

eroded material and the vertical erosion rate (Table 1).

Figure 5. Slice of removed rock since 13 Ma from the Ocoña - Cotahuasi confluence to the Pacific Ocean (Albino, 2006). The volume of eroded material from the canyon watershed is shown as inverted relief. Maximum eroded material (= maximum depth of the canyon) is color coded as dark red. Minimum eroded material (paleosurface prior to the c. 40 Ma – 14 Ma peneplain) is color coded as black.

Also, using ignimbrites 40Ar/39Ar ages and the difference in altitude between two dated markers, we estimate

incision rates at which the canyon was cut down (Table 1): the incision rate increased from 140 m/Ma between

14 and 9 Ma to 450 m/Ma between 9 and 4 Ma, reflecting therefore the major incision phase during Late

Miocene-Earliest Pliocene.

Dated ignimbrite taken as stratigraphic marker

Volume of eroded material for each time slice

Erosion volumetric rate as volume / time slice considered

Incision rate from height difference between 2 levels

Huaylillas ign., base = 13 Ma 3300 km3 820 km3 / Ma over 4 Ma 140 m/Ma

Caraveli ign., base = ~9 Ma 880 km3 200 km3 / Ma over 5 Ma 450 m/Ma

Sencca ign., base = ~4 Ma 1320 km3 340 km3 / Ma between 4 Ma and

present bottom

500 m/Ma

TOTAL 5500 km3 420 km3 / Ma over 13 Ma 170 m/Ma

Table 1. Results of computed eroded volumes and calculated erosion and incision rates for three time slices.

These results display spatial and time variations in incision rate. First, spatial variations are observed between

the Cotahuasi canyon, cut after 3.76 Ma at a rate of 750 m/Ma, and the Ocoña canyon where rocks have been

removed since 4 Ma at a rate of 250 m/Ma. Second, incision rates have also fluctuated through time: the 3.76-1.6

Ma valley fill of the Cotahuasi canyon was cut again after 1.6 Ma at a faster incision rate of 1300 m/Ma (Fig. 6).

The vertical incision is responsible of almost 80% of the overall valley downcutting between 9 and 4 Ma. This

is related to the work expenditure of the river, reflecting either change in channel gradient or in stream power

linked to the rainfall amount and distribution. The estimated incision rate (170 m/My) averaged on 13 Ma has

been compared with incision rates computed in similar catchments, i.e. cut in Mid-Late Cenozoic volcanic rocks,

under arid climate and in a convergent plate setting. As an example, the incision rate of the Rio Cotahuasi-Ocoña

Figure 6. Diagram comparing incision rates in the Rio Cotahuasi, northern part of the canyon system with the Rio Ocoña in the southern canyon system. The temporal evolution of incision rates is shown in both canyons. The two canyons join in a confluence site shown in Fig 3.

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system lies in the upper range of the incision rate (10 – 200 m/My) computed by Montgomery and Lopez-Blanco

(2003) in the Sierra Madre Occidental, Mexico. This demonstrates that either a change in water discharge and

stream power or a change in slope gradient may have occurred in the canyon area.

Hypotheses on the causes of the major incision phase

The ultra-deep valleys under study show two geomorphic sections across transverse profiles: a wide (10-30

km), shallow (<12 km) valley above a narrow (1-2 km), ultra-deep (2-3.5 km) canyon. The wide valley was

shaped between 13 and 9 Ma (Middle-Late Miocene), whereas the deep canyon was cut well after 9 Ma (Late

Miocene-Early Pliocene). This geomorphic contrast is attributed to a change in erosion processes, i.e lateral

“pediment” erosion (driven by rill and sheet wash) vs. linear fluvial downcutting (driven by river water discharge

and stream expenditure in hard bedrock). Both categories of erosion processes are governed by rainfall intensity

and seasonal distribution. The climate must have changed from hyperarid to semi-arid conditions if we assume

that the erosion style has probably changed between Late Miocene and Early Pliocene. This leads us to

investigate what critical parameters can sustain high incision rates under arid conditions, and what processes are

responsible for initiating the vertical incision.

Physical properties, (facies and welding grading of the ignimbrites may be called upon to explain the distinct

rates of incision and erosion between the initial wide valley formation (13-9 Ma) and the subsequent canyon

downcutting (<9-4 Ma). Most of non-welded, loose Sencca ignimbrites sheets have favoured high-rate vertical

incision due to their ‘soft’ lithology as opposed to the ‘hard’ welded Huaylillas ignimbrites. In contrast, the Late

Cretaceous diorite batholiths of the Western Cordillera underwent weathering and lateral erosion that formed

pediments and enlarged the valley (e.g. Río Ocoña segment between Iquipi and the confluence with Río

Cotahuasi). Thus, bedrock lithologies display distinct responses to slope erosion processes and river incision.

Also, incision rate can increase if water discharge is rapidly added to the drainage network, hereby increasing the

river flow and stream power. A capture of the drainage network from catchments located in the western

Altiplano may have occurred t Late Miocene time. This could have been induced by tectonic vertical movements

related to the uplift of western Altiplano and to the flexure of the Western Cordillera from NE to SW.

Another process contributing to increase incision may be linked to landsliding and debris avalanches. Unstable

cliffs cut in overhanging non-ignimbrites (Rio Cotahuasi river) or in hydrothermally altered cores of eroded

volcanoes (e.g. Nevado Solimana) have failed and triggered huge landslides. Debris-avalanche deposits

repeatedly dammed the river channel (e.g. Cotahuasi village). Subsequent dam breakouts trigger debris flows

and floods downvalley, which increase stream power and eventually accelerate the incision process downvalley

over a short period.

References Albino F., 2006. Contribution au problème d’érosion dans les Andes centrales à l’aide de modèles numériques : exemple des

canyons péruviens. Travail d’Etude et de Recherche, L.M.V. Université Blaise Pascal., Unpublished 20p. Gunnell Y., J.-C. Thouret, S. Brichau and A. Carter, 2008. A low-temperature thermochronology of denudation, crustal uplift

and canyon incision in the Western Cordillera of southern Peru. Geoph. Res. Abs., vol.10, EGU2008. Hartley A.J.,2005. What caused Andean uplift? Extended Abstract, 6th ISAG, Barcelona 2005: 824-827. Montgomery D. R. and J. Lopez-Blanco, 2003. Post-Oligocene river incision, southern Sierra Madre Occidental, Mexico.

Geomorphology, 55: 235-247. Thouret J.-C., G. Wörner, Y. Gunnell, B. Singer, X. Zhang and T. Souriot, 2007. Geochronologic and stratigraphic

constraints on canyon incision and Miocene uplift of the Central Andes in Peru. Earth Plan. Sci .Letters, 263: 151-166.

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Holocene submarine volcanoes in the Aysén fjord, Patagonian Andes (44ºS): Relations with the Liquiñe-Ofqui Fault Zone

Luis E. Lara

Servicio Nacional de Geología y Minería, Chile. Av. Santa María 0104. Santiago. Chile (lelera@sernageomin.cl)

KEYWORDS : submarine volcano, Liquiñe-Ofqui Fault Zone, 2007 Aysén seismic swarm, Patagonian Andes

Introduction

During the CIMAR 13 cruise to the Aysén fjord (45ºS), basaltic rocks were recovered from a submarine

volcanic cone located few kilometers west of an active intra-arc seismic zone. Since the middle of January 2007

an unusual intraplate seismic swarm has been taking place in the Aysén region. At the time of writing, minor yet

perceptible shakes are still being felt in the area nearby. The epicentral region was initially centered in the

middle of the Aysén fjord (Barrientos et al., 2007), coinciding with the master fault of the Liquiñe-Ofqui Fault

Zone (LOFZ; e.g., Hervé, 1994; Cembrano et al., 2007). The latter is a margin-parallel ca. 1200 km long

structural system, active at least since the Miocene (Cembrano et al., 1996; 2000; Thomson, 2002; Rosenau et

al., 2006). The overall stress regime along the arc domain is a bulk dextral transpression which determines

mostly strike-slip displacement along the NS-striking faults with some local vertical movements. The short-term

kinematic history is poorly constrained but growing evidence is in agreement with the long-term picture

provided by the tectonic analysis of fault populations at mesoscale from 38º-46ºS (Lavenu and Cembrano, 1999).

A damaging Mw: 6.3 shallow earthquake occurred on April 21st causing huge landslides and debris flows (ca.

23 km3) along the high slope walls of the fjord. Rock masses fell into the water and triggered tsunami waves that

shattered the coast killing 10 peoples. Although the tectonic scenario was favorable to shallow earthquakes and

the consequent slope failure, the near real-time hazard assessment of the ongoing crisis was at the beginning

focused on a possible submarine eruption above the epicentral area. Because several pyroclastic cones are

present along the LOFZ in the surrounding area, this hypothesis seemed to be plausible although did not account

for the entire geological process which would be, at a regional scale, tectonic in origin (e.g., Cembrano et al.,

2007). The presence of a bathymetric high across the Aysén fjord, few kilometers west of the epicentral area,

supported the idea of submarine volcanism in the recent past and thus the exploration of that place became a

target for the scientific assessment strategy. This article gives the first news of the finding of submarine volcanic

rocks in the Aysén fjord and briefly discusses their implications.

Quaternary tectonics and volcanism in Southern Andes

The most significant Neogene regional structure in the Patagonian Andes is the LOFZ. In the Aysén region, the

LOFZ consist of two overstepping NNE-striking master faults joined by a series of en échelon NE-striking

subsidiary faults forming a strike-slip duplex structure. Tectonic analysis combined with thermochronology

document a dextral transpression ductile deformation at ca. 4 Ma, followed by brittle compressional to strike-slip

deformation after 3.8 Ma and 1.6 Ma, respectively (Hervé, 1994; Cembrano et al., 1996; 2000; 2002; Lavenu

and Cembrano, 1999; Thomson, 2002). Quaternary volcanoes are spatially related to the strike-slip structural

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system and a probable causative relation has been proposed for the entire Southern Andean Volcanic Zone

(Cembrano and Moreno, 1994; López et al., 1995a; Lara et al., 2006a; Rosenau et al., 2006). Some clusters of

monogenetic volcanoes as Caburgua (39ºS), Anticura (40ºS), Ralún-Cayutue (41ºS) and Puyuhuapi (44ºS) are

exactly on top of the master faults and seem to show a contrasting magmatic evolution with respect to the main

arc volcanoes (López et al., 1995a; 1995b; Hickey-Vargas et al., 2002). In the Aysén fjord, a NNE-striking

alignment of pyroclastic cones extends along the Pescado River whereas another NS-striking cluster is located

near the Cuervo River, both along the main trace of the LOFZ. In the middle of them, and just above an

inflection of the fault trace, the epicentral area of the ongoing seismic swarm is located. A few kilometers west,

the sampled submarine volcano emerges from the flat bottom of the fjord.

Volcanoes of the Aysén fjord

A bathymetric positive anomaly across the Aysén fjord was already detected on the nautical 1:50,000 scale

chart No. 8106 by the Hydrographic and Oceanographic Survey of the Chilean Navy (SHOA) where a NNE-

trending ridge shows minimum depths of ca. 60 m.b.s.l above a flat floor that averaged 200 m.b.s.l to the east

and 330 m.b.s.l. to the west, respectively (Fig. 1). This ridge is also located above a vertical disruption of the

fjord floor which could be tectonic in origin. Given the regional geologic setting, the presence of a submarine

volcano was considered as a possibility during the hazard assessment of the seismic crisis. However, other

options as basement highs or glacial morphologies cannot be precluded. In fact, Araya-Vergara et al. (2007)

described such submarine elevations as frontal moraines in the Reloncaví fjord (42ºS) and that seem to be

common in other similar environments, as for example in the Norwegian fjords (Laberg et al., 2007).

As a complement to the hazard evaluation, a multibeam batimetric survey was done in April 2007 by the PSH

Cabrales vessel from SHOA institution (Fig. 1). This survey did not detected significant bathymetric changes on

the epicentral area but improved the morphological knowledge of the western ridge. Considering the already

scheduled CIMAR 13 cruise to the Aysén region, a dredging procedure was conducted on board of the R/V Agor

Vidal Gormaz. The third sampling attempt allowed to collect several fragments of massive dark vesicular basalts

embedded in fine-grained gray sediments. Scarce biological remnants were recovered and they mostly consist of

annelids, sea urchins or some adhered cirripedial crustaceous. Neither hydrothermaly altered samples nor

pyroclastic material were recovered. The fine sediments seem to have formed a thin layer above the basalts. In

fact, the low-frequency echosounder control showed a sharp acoustic response on the volcano flank that permits

to image mostly lava flows.

What impact placed this finding to the 2007 Aysén seismic crisis?. Maybe directly none because no thermal

anomalies were measured along the fjord even above the epicentral area where gas outputs would have been

detected. The sampled volcano is, of course, inactive at present. However, the steep slope, the absence of

alteration or vesicles infilling and the thin sediment cover suggest its young, probably late Holocene age. A few

kilometers further south, the base of a pyroclastic sequence related to an emerged and partially eroded

pyroclastic cone of the Pescado River yield an age of ca. 6 ka. Thus, although the 2007 seismic swarm did not

show clear evidence of an ongoing magma ascent, the presence of this young volcano enhances the hypothesis of

volcanism in the fjord but taking part of a more regional process controlled by the fault kinematics and the

regional stress regime (Cembrano et al., 2007).

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Several lines of research derive from this finding. For instance, geochemistry of these lava flows will allow a

comparison with other LOFZ-related monogenetic volcanoes. As already mentioned, small eruptive centers on

top of the LOFZ show key features that could be interpreted as isolated magma batches perhaps not related with

the dominant flux melting process that imprint the arc signature to the magmas erupted from the stratovolcanoes.

The precise role of the LOFZ in both the magma genesis and the subsequent ascent should be investigated in

detail. This finding along with additional data could improve our knowledge of these coupled processes and thus

become a real contribution to the hazard assessment in the Southern Andes.

Figure 1. (a) Nautical chart No. 816 by SHOA. (b) Shaded relief image based on a multibeam survey by R/V Agor Vidal Gormaz vessel of SHOA institution (unpublished). Blue ellipse shows the approximate 2007 (pre-April 21th) epicentral area.

References Araya-Vergara, J.; Vieira, R.; Suárez, M. 2007. El sistema submarino Relocaví (Norpatagonia): análisis morfoacústico,

batimétrico y fundamentos sedimentológicos. Revista Ciencia y Tecnología del Mar. (en prensa). Barrientos, S.; Bataille, K.; Aranda, C.; Legrand, D.; Báez, J.C.; Agurto, H.; Pavez, A.; Genrich, J.; Vigny, C.; Bondoux, F.

Complex sequence of earthquakes in Fjordland, Southern Chile. In Geosur 2007, Libro de Resúmenes, p. 21. Cembrano, J.; Moreno, H. 1994. Geometría y naturaleza contrastante del volcanismo Cuaternario entre los 38ºS y 46ºS:

¿dominios compresionales y extensionales en un régimen transcurrente? In Congreso Geológico Chileno, N° 7, Actas 1: 240-244, Concepción.

Cembrano, J.; Lara, L.; Lvenu, A.; Hervé, F. 2007. Long-term and short-term kinematic history of the Liquiñe-Ofqui fault zone: a review and implications for geologic hazards assessment. In Geosur 2007, Libro de Resúmenes, p. 30.

Cembrano, J.; Hervé, F.; Lavenu, A. 1996. The Liquiñe-Ofqui fault zone: a long-lived intra-arc fault Zone in southern Chile. Tectonophysics 259: 55-66.

Cembrano, J.; Shermer, E. ; Lavenu, A. ; Sanhueza, A. 2000. Contrasting nature of deformation along an intra-arc shear zone, the Liquiñe-Ofqui fault zone, southern Chilean Andes. Tectonophysics 319: 129-149.

Cembrano, J.; Lavenu, A.; Reynolds, P.; Arancibia, G.; López, G.; Sanhueza, A. 2002. Late Cenozoic transpressional ductile

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deformation north of the Nazca–South America–Antarctica triple junction. Tectonophysics 354: 289– 314. Hervé, F., 1994. The southern Andes between 39° and 44°S latitude: the geological signature of a transpressive tectonic

regime related to a magmatic arc. In Tectonics of the Southern Central Andes (Reutter, K.J.; Scheuber, E.; Wigger, P.J.; editors). Springer, Berlin, pp. 243-248.

Hickey-Vargas, R.; Sun, M.; López-Escobar, L.; Moreno-Roa, H.; Reagan, M.; Morris, J.; Ryan, J. 2002. Multiple subduction components in the mantle wedge: Evidence from eruptive centres in the Central Southern Volcanic Zone, Chile. Geology 30: 199-202.

Laberg, J.S.; Eilertsen, R.S.; Salomonsen, G.R.; Vorren, T.O. 2007. Submarine push moraine formation during the early Fennoscandian Ice Sheet deglaciation. Quaternary Research 67: 453-462.

Lara. L.E., Lavenu, A., Cembrano, J.; Rodríguez, C. 2006. Structural controls of volcanism in transversal chains: resheared faults and neotectonics in the Cordón Caulle-Puyehue area (40.5ºS), Southern Andes. Journal of Volcanology and Geothermal Research 158: 70-86.

Lara, L.E.; Moreno, H.; Naranjo, J.A.; Matthews, S.; Pérez de Arce, C. 2006b. Magmatic evolution of the Puyehue-Cordón Caulle Volcanic Complex (40º S), Southern Andean Volcanic Zone: from shield to unusual rhyolitic fissure volcanism. Journal of Volcanology and Geothermal Research 157: 343-366.

Lavenu, A.; Cembrano, J. 1999. Compressional and traspressional-stress pattern for Pliocene and Quaternary brittle deformation in fore arc and intra-arc zones (Andes of Central and Southern Chile). Journal of Structural Geology 21: 1669-1691.

López-Escobar, L.; Cembrano, J.; Moreno, H. 1995a. Geochemistry and tectonics of the Chilean Southern Andes basaltic quaternary volcanism (37-46ºS). Revista Geológica de Chile 22 (2): 219-234.

López-Escobar, L.; Kempton, P.D.; Moreno, H.; Parada, M.A.; Hickey-Vargas, R.; Frey, F.A. 1995b. Calbuco volcano and minor eruptive centers distributed along the Liquiñe-Ofqui fault zone, Chile (41°-42°S): contrasting origin of andesitic and basaltic magma in the Southern Volcanic Zone of the Andes. Contributions to Mineralogy and Petrology 119: 345-361.

Rosenau, M.R.; Melnick, D. ; Echtler, H. 2006. Kinematic constraints on intra-arc shear and strain partitioning in the Southern Andes between 38°S and 42°S latitude. Tectonics 25: TC4013, doi:10.1029/2005TC001943.

Thomson, S.N. 2002. Late Cenozoic geomorphic and tectonic evolution of the Patagonian Andes between latitudes 42° and 46°S: An appraisal based on fission-track results from the transpressional intra-arc Liquiñe-Ofqui fault zone. Geological Society of America Bulletin 114: 1159-1173.

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Determination of an arc-related signature in Late Miocene volcanics over the San Rafael block, Southern Central Andes (34º30´-37ºS), Argentina: The Payenia shallow subduction zone

Vanesa D. Litvak, Andrés Folguera, & Víctor A. Ramos

Laboratorio de Tectónica Andina, FCEyN, Universidad de Buenos Aires, Argentina (vane@gl.fcen.uba.ar)

KEYWORDS : volcanic arc, shallow subduction, geochemistry, Payenia

Introduction

New dates and geochemical studies in Chachahuén volcanic complex (Kay 2001; Kay et al. 2006a, 2006b),

emplaced far away from the actual volcanic arc (Figure 1), evidence the development of a shallow subduction

zone during the Late Miocene at these latitudes. The development of this zone was restricted to the time interval

of the Chachahuén volcanic complex corresponding to 7.3-4.9 Ma. The uplift of San Rafael block would have

been one of the consequences of this shallow subduction, as a result of the inversion of previous –Late Triassic–

structures in the foreland zone (Figure 1) (Ramos and Folguera, 2005) in a similar way to the Sierras Pampeanas

to the north directly related to the Present flat subduction Pampean zone. The main objective of this work is to

evaluate the chemical signature of San Rafael Block mesosilicic volcanism in order to document Late Miocene

eastward expansion in the area.

Middle to Late Miocene volcanic sequences

A series of isolated volcanic centers can be located over the San Rafael block 300 km away of the Present arc

front and approximately 550-600 km from the trench (Figure 1). They correspond to highly eroded

stratrovolcanoes of Late Miocene age. Two groups of volcanic events with particular spatial distributions and

age can be defined: a) The oldest sequences (15-10 Ma) are located in the central and western area of the San

Rafael Block while b) The youngest sequences (8-3.5 Ma) are located in an eastern position relative to the

previous group and even to the north and south of it (Figure 1).

Both groups are constituted by porphiritic andesites with felty groundmass. The main phenocryst phase is

plagioclase. Hornblende is the common mafic mineral for both groups, while pyroxene is much common in

andesites of the first one. Tridimite is present in the more differentiated varieties of both groups of volcanic

rocks. Petrographical features characterize typical calk-alkaline volcanic andesites. Major elements were

analyzed by fusion-ICP while REE and trace elements were analyzed by fusion- ICP MS, according to the

Activation Laboratories standards and methodology. Geochemical classifications correlate with the petrographic

one being the first group formed by andesites with 60-62 % of SiO2, while the second group has a wider range of

differentiation (54 to 67 % of SiO2), being basandesites to dacites. All of them correspond to high-K lavas and to

a calk-alkaline signature in the SiO2 vs. FeO/MgO diagram. Trace elements ratios, such as Ba/La > 20 and La/Ta

> 25 evidence an arc-like signature for their magma sources (Kay, 1977; Gill, 1981). The same chemical

signature is shown by their trace elements mantle normalized diagrams, which show depletion of HFSE relative

to LILE –regardless their silica content– as a result of subducted slab component contributions (Hildreth and

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Moorbath, 1988). Rare earth chondrite normalized diagrams are essentially concave-up, showing that pyroxene

and amphibole were the residual mineral phases in the source; this is reflected in their La/Yb, La/Sm and Sm/Yb

ratios. One particular difference between samples from younger and older volcanic centers is that the first ones

have smaller Sm/Yb ratios (2.4-2.5) than the second ones (2.7-3.2), which presumes an increase of residual

mineral phases that retains heavy rare earth elements –such as amphibole– for younger lavas. This difference is

consistent with the mafic mineral assemblage present of both groups of lavas; while older andesites have mainly

pyroxene, younger ones have higher amounts of amphibole. Negative Eu anomalies are not present in both

groups of rare earth diagrams, which imply that plagioclase was not an important residual mineral phase in

equilibrium with the magmas at their sources.

Figure 1. Location of San Rafael Block and distribution of mesosilicic Late Miocene volcanic centers. Dashed line separates two groups of studied volcanic stages: Older San Rafael Block volcanics and younger San Rafael Block volcanics. Isopachs indicate thickness of related Neogene sinorogenic sequences.

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Discussion and concluding remarks

The San Rafael andesitic to dacitic stratovolcanoes represent Late Miocene volcanic arc activity. Evidence for

their arc-like pattern comes from their petrographical characteristics and from their major and trace element

chemical signatures. The recognition of these volcanic sequences, and their affinity with an arc-like signature,

allow understanding real dimensions of arc-expansion at that time over the San Rafael block.

The Chachahuén Volcanic Complex is located further south (Figure 1) and includes three main volcanic events

with an age range from 7.3 to 4.9 Ma (Kay et al. 2006b). Andesitic volcanism of Chachahuén shows similar

chemical behavior than other San Rafael Block volcanics: all of them are high-k calk-alkaline rocks with arc-like

affinities. Main differences are seen in REE behavior between younger San Rafael Block volcanics and

Early/Late Chachahuén volcanic rocks. While the latest share similar Sm/Yb ratios with older San Rafael Block

volcanics, they differ from younger in having smaller Sm/Yb ratios. The increase of Sm/Yb ratios in the first

volcanic stage of San Rafael Block and Chachahuén Volcanic Complex towards the youngest San Rafael Block

volcanics shows an increase in pressure conditions at the melt generation site. The absence of plagioclase

fractionation in younger San Rafael volcanics also evidences a relatively deeper site of magma generation.

Higher pressure conditions are consistent with tectonic shortening and crustal thickening as a result of subducted

slab flattening during the development of the Payenia shallow subduction zone.

References Gill, J.B., 1981 - Orogenic Andesites and Plate Tectonics. Springer-Verlag, Berlin Heidelberg New York, 392 p. Hildreth W., Moorbath, S., 1988 - Crustal contributions to arc magmatism in the Andes of Central Chile. Contributions to

Mineralogy and Petrology. 98: 455-489. Kay, R.W, 1977 - Geochemical constrains on the origin of Aleutian magmas. In Talwani, M., Pitman W.C.III (eds.). Islands

arcs, deep sea trenches and back-arcs basins. AGU Ewing Ser. 1: 229-242. Kay, S., 2001 - Tertiary to recent magmatism and tectonics of the Neuquén basin between 36°05´ and 38°S latitude: Repsol-

YPF Buenos Aires. Unpublished Report, 77 p. Kay, S.M., Burns, M., Copeland, P., 2006a – “Upper Cretaceous to Holocene Magmatism over the Neuquén basin: Evidence

for transient shallowing of the subduction zone under the Neuquén Andes (36°S to 38°S latitude)”. In Kay, S.M. and Ramos, V.A. (eds.). Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén basin (35º-39ºS)-Geological Society of America, Special Paper, 407: 19-60.

Kay, S.M., Mancilla, O., Copeland, P. 2006b – “Evolution of the Backarc Chachahuén volcanic complex at 37°S latitude over a transient Miocene shallow subduction zone under the Neuquén Basin”. In Kay, S.M. and Ramos, V.A. (eds.). Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén basin (35º-39ºS)-Geological Society of America, Special Paper, 407: 215-246.

Ramos, V., Folguera, A., 2005 – “Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc and foreland deformation”. In Veiga, G., Spalletti, L., Howell J. and Schwarz E. (eds.). The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics-Geological Society of London, Special Publication, 252: 15-35.

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Sedimentary constraints on the tectonic evolution of the paired Tumaco–Borbón and Manglares forearc basins (southern Colombia - northern Ecuador) during the Late Cenozoic

Eduardo López1, Jean-Yves Collot

1, & Marc Sosson

1

1

Université de Nice Sophia-Antipolis, IRD, Université Pierre et Marie Curie, CNRS, Observatoire de la Côte

d’Azur, Geosciences Azur, BP 48, Port de la Darse, 06235, Villefranche s/mer, France

(lopez@geoazur.obs-vlfr.fr, collot@geoazur.obs-vlfr.fr, sosson@geoazur.unice.fr)

KEYWORDS : duplex, forearc, basin, interplate

Abstract Based on seismic reflection and drilling data we analyze the temporal evolution of theTumaco – Borbon and Manglares forearc basins that lay on the North Ecuador-South Colombia margin. This evolutionconstrains the uplift of the Remolino high and the development of a distinctive double fore-arc basin setting to the Early to Late Miocene. The development of the Remolino high is compatible with lower crust thickening by duplexing during a period of increased plate convergence.

Introduction

The Tumaco – Borbon and Manglares forearc basins are located respectively in the coastal range and

submarine part of the South Colombia - Northern Ecuador margin (Figure 1a). These basins, which are separated

by the Remolino High and contain thick siliciclastic sediment sequences, are underlain by oceanic terranes

Figure 1. a) Location and tectonic setting of Tumaco – Borbon basin. b) Geological sketch of the South western Colombia and Northern Ecuador region showing the location of structural and seismic sections, well data (Remolino Grande 1 – RG1, Majagua 1 – MJ1, Chagüi 1 CH1, Camarones 1 – CM1) and places mentioned in this work.

accreted against the western border of the South American plate, during the Late Cenozoic to Early Paleogene

(Kerr et al, 2002), (Marcaillou, 2003; Escovar et al, 1992; Evans and Whittaker, 1982). Despite the extensive

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land and marine geophysical data, regional geological cartography (Evans and Whittaker, 1982; IGAC –

INGEOMINAS, 2006), and sparse well log data, the two connected fore arc basins have been studied separately,

and their sedimentary and tectonic evolution has remained poorly i known. A detailed sedimentary study of

their stratigraphic records based on seismic stratigraphy and electrofacial analysis of well logs (Figure 1b),

allows constraining the age, of the basins separation, and analyzing their deformation in relation with the

subduction processes during the last 30 My.

Data set and methods

Based on electrosequential analysis of more of 3 km of well logs at the Remolino Grande 1 well, we

reconstructed the temporal variations of sedimentary environments in the Remolino high. Their abrupt changes

are correlated with seismic surfaces on both onshore and offshore seismic reflection profiles (isochronous

surfaces), to constrain the ages of the seismic units in both the Tumaco – Borbon and Manglares forearc basins.

Regional faults and the deep crustal geometry of the margin were constructed by kink band methods (ref) and

seismic refraction data (Suppe and Chang, 1983).

Figure 2. a) Structural cross section through Southern Colombia - Northern Ecuador arc – trench system (subduction dip plate based on Agudelo 2003). b) Depth-converted seismic line across the Tumaco – Borbon basin (see Figure 1b for location), showing Growth unconformities (GU), Onlaps (Ol), Down laps (Dl), Clinoforms (Cf). c) seismic line through the Manglares off shore basin, showing the shale diapir deformation.

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Analysis and results

Well log electrosequential analysis, and seismic stratigraphy interpretation reveal five seismic units in both

basins (Figure 2b): From bottom to top, these units are dated: 1) Late Oligocene, 2) Early Miocene, 3) Middle

Miocene, 4) Late Miocene, 5) Pliocene to Pleistocene. The complete sedimentary sequence is characterized by 1)

a gradual up ward coarsening grain size, and 2) an upward shallowing of the paleodepth sediment environments.

This shallowing sedimentation is characterized by the following successive environments : basin floor fan, slope

fan with high volcanic supply, outer shelf, inner shelf and continental fans with high volcanic supply. The

SONIC well log data reveal that the seismic unit dated Late Oligocene is a geologic layer including

overpressured shales, thus producing diapir structures interpreted in both basins (Figure 2c). Middle Miocene to

Pleistocene sedimentary sequences are thicker in the Tumaco – Borbon basin than in the Manglares basin. A

synthetic cross section reveals that the Remolino high is the expression of a partial overlap ramp anticline,

possibly produced by a lower crust duplex system, with 45 km of shortening. The duplex rests directly over the

interplate zone, (Figure 2a), and may be associated with the Eastern extension of the high velocity basement

zone defined offshore by Agudelo (2005). These structures separate the landward Tumaco-Borbon forearc basin

from the seaward Manglares forearc basin.

Discussion and conclusion

The Tumaco – Borbon and Manglares forearc basins were separated from each other by the development of the

Remolino high during the late Cenozoic time. Regional basin floor fan sedimentation occured in a single initial

basin during the Early Paleocene to Late Oligocene times. During the Late to Middle Miocene, slope fan

sediments developed concurrently with the uplift of the Remolino high, thus separating the Tumaco – Borbon

basin from the Manglares basin. During the Late Miocene to Quaternary, shale diapirs rose into both basins.

References Agudelo, W., (2005). Imagerie sismique quantitative de la marge convergente d’Equateur-Colombie : Application des

méthodes tomographiques aux données de sismique réflexion multitrace et réfraction-réflexion grand-angle des campagnes SISTEUR et SALIERI. Thèse de doctorat de l’Université Paris 6. 203 p.

Escovar, R., Gomez, L. A., and Ramirez, J. R., 1992. Interpretacion de la Sismica Tumaco 90 y evaluacion preliminar del area. Informe final proyecto Tumaco 90 Empresa Colombiana de Petroleos, Gerencia de Exploracion. 58 p.

Evans, C. D. R., and Whittaker, J. E., 1982. The geology of the Borbon Basin, Nortwest Ecuaador, in Trench-forearc geology, edited by J. K. Legget, Geological Society of London Special Publication, pp. 191 – 198.

IGAC – INGEOMINAS, 2006. Investigacion integral del Anden Pacifico Colombiano. Tomo 1 Geologia. 165 p. Kerr, A. C., Aspden, J. A., Tarney, J., and Pilatasig L., F., 2002. The nature and provenence of accreted oceanic terranes in

Western Ecuador: geochemical and tectonic constrains. Journal of the Geological Society, London, v. 159, p. 577 – 594. Marcaillou, B., 2003. Régimes tectoninques et thermiques de la marge Nord Equateur – Sud Colombie (0° - 3,5°N° -

Implications sur la sismogènese. Phd thesis, Université Pierre et Marie Curie, Paris. 197 p., 10 anexes. Suppe, J. and Chang, Y. L., 1983. Kink method applied to structural interpretation of seismic sections, western Taiwan:

Petroleum Geology of Taiwan, n° 19, pp. 29 – 47.

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Compressive active fault systems along the Central Andean piedmont

J. Macharé1, L.

Audin

2, C. Benavente

1, M.

Saillard

2, V. Regard

2, & S. Carretier

2

1 INGEMMET, Lima, Peru (jmachare@ingemmet.gob.pe)

2 LMTG-IRD, Toulouse, France (laurence.audin@ird.fr)

KEYWORDS : Central Andes, active tectonics, forearc deformation, geomorphology

Introduction

It’s now established that Andean forearc is not concentrating as much tectonic shortening as the foreland since

Middle Miocene. GPS measurements are neither available to inform on the long-term deformation across the

Andes in Peru and anyway rather describe the elastic response of the Andean forearc to the Nasca-South

American Plate convergence. Few neotectonic studies focuses on the Western side of the Andes and little is

known about the active deformation in the Central Andes Pacific lowlands (Sébrier et al., 1988). Recent

publications mainly improved the description of geomorphic surfaces (Thouret et al. 2007) and cosmogenic

dating of the latter show much younger ones than expected (Hall et al., 2008). The topographic gradient on the

western side of the Peruvian Andes is quite high as the trench (-7000m) lies only 200km away from the highest

point (6000m). Moreover, authors still question the fact that the Andes build through a giant focused monocline

or normal fault and demonstrate doing so the need of further mapping of the fault systems on the western side of

the Central Andes (Schildgen et al., 2007).

Geomorphic evidences of recent tectonic activity are observed from the Coastal Cordillera to the piedmont of

the Western Cordillera (Audin et al., 2008). We present here evidences of newly mapped compressional fault

system, together with evidences of their activity since at least the Pliocene in the southern Peruvian forearc, near

Tacna. Examination of aerial photographs , satellite data, and focused field work not only confirms that there is

recent tectonic activity but also revealed the presence of additional active structures that should be taken into

account in the description of Andean deformation. In response to active tectonics, these fault systems affected

very young terraces and Quaternary pediments along the piedmont of the Central Andes (Figure 1). We present

Figure 1: Google earth 3D image on the Calientes Fault system from South to North.

Figure 2: Zoom on white dot, figure 1; with details of the recent scarplets on the main fault.

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some of the geomorphic signatures, such as active fault traces, scarplets (Figure 2), sag ponds, river terraces and

some major and minor landslides, which are demonstrative of active tectonics in this area. Mapping of fault trace

geometry and identifying recent surface offsets are used to determine the kinematics of the Calientes active

thrust fault.

Figure 3: Topographic profile for the offset of the Pachia Fm ( 2,8Ma; Flores et al., 2002 and 2005).

Discussion and Conclusion

The purpose of this work is to show the existence of previously undescribed crustal fault systems in the forearc

of the Central Andes in Southern Peru (Figure 1). Some of these markers are robust enough to allow us to

characterize the kinematics of Quaternary active faults (Figure 2). The main active faults identified along the

Central Andean Piedmont in Peru are trending parallel to the trench and aare part of compressional or

transpressional fault systems (see the Incapuquio Fault System). At the scale of a single structure, even being

part of a segmented fault system, the deformation is comparatively small with respect to the Andean uplift that

accomodates the building of the mountain range, but at a larger scale the fault system could be responsible of

quicker uplift rates (than those proposed here on Figure 3 on one segment). We propose that despite the large

degree of segmentation that is observed along those fault systems, some crustal seismic events can occur in this

area of the Andean forearc, on the Calientes Fault system (Figure 2). Many of these faults we have identified are

capable of generating earthquakes, some small and local ( as the October 2005 one, Ml 5.7), others major and

capable of impacting human activities. Even if today we do not calculate a recurrence interval, we can at least

place bounds on this and we argue, that it should be less than historical times (~1000 yr). Moreover, both the

piedmont of the Western Cordillera in its lower parts and the central basin experienced extremely low

denudation rates (Kober et al., 2005), much of which is likely accommodated by mass movements triggered by

active tectonics or subduction earthquakes (Figure 1).

Our morphological data suggests an interpretation that differs from the GPS measurements and models which

report that no active deformation is observed in the forearc of southern Peru (Khazaradze and Klotz, 2003).

Some major tectonic structures ( that belongs to the Incapuquio Fault system for exemple) shows Quaternary

activity, mainly compresionnal or transpressional. Although there is only one permanent GPS station;

segmentation of the faults, small displacements and long recurrence times are probably the cause of the

uncomplete mapping of active faults in Southern Peru.

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References Audin L., P. Lacan, H. Tavera and F. Bondoux. Upperplate deformation and seismic barrier in front of Nazca subduction

zone: The Chololo Fault System and active tectonics along the Coastal Cordillera, southern Peru. Tectonophysics. In Press. Available online 4 April 2008.

Flores, A., Jacay, P., Roperch, P., Sempere, P. (2002). Un evento volcánico de edad Plioceno superior en la región de Tacna: la ignimbrita Pachía. XI Congreso de Geología del Perú. Lima, pp.199-205.

Flores, A., Jacay, P., Roperch, P., Sempere, P. (2005). Oligocene-Neogene tectonics and sedimentation in the forearc of southern Peru, Tacna area (17.5º-18.5ºS). 6éme ISAG. Barcelona. pp.269-272.

Hall S.R., D.L. Farber, L. Audin, R.C. Finkel and A-S. Mériaux. Geochronology of pediment surfaces in southern Peru: Implications for Quaternary deformation of the Andean forearc Tectonophysics. In Press. Available online 4 April 2008.

Khazaradge, G., and Klotz, J. (2003). Short and long-term effects of GPS measured crustal deformation rates along the South-Central Andes. Journal of geophysical research, vol. 108. pp. 1-13.

Kober F., S. Ivy-Ochs, F. Schlunegger, H. Baur, P.W. Kubik and R. Wieler 2007. Denudation rates and a topography-driven rainfall threshold in northern Chile: Multiple cosmogenic nuclide data and sediment yield budgets. Geomorphology, Volume 83, Issues 1-2: 97-120.

Schildgen, TF, Whipple, KX, Hodges, KV, Reiners, PW, Pringle MS, 2007, Uplift of the western Altiplano from canyon incision history, southern Peru, Geology, v. 35, no. 6., p. 523-526; doi: 10.1130/g23532A.1.

Sébrier, M., Lavenu, A., Fornari, M., & Soulas, J.-P. 1988. Tectonics and uplift in the Central Andes (Peru, Bolivia and Northern Chile) from Eocene to Present. Géodynamique 3: 85-106.

Thouret J.-C., G. Wörner, Y. Gunnell, B. Singer, X. Zhang and T. Souriot 2007. Geochronologic and stratigraphic constraints on canyon incision and Miocene uplift of the Central Andes in Peru, Earth and Planetary Science Letters, Volume 263, Issues 3-4: 151-166.

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Tracing a major crustal domain boundary based on the geochemistry of minor volcanic centres in southern Peru

Mirian Mamani1, Gerhard Wörner

2, & Jean-Claude Thouret

3

1 Georg-August University, Goldschmidstr. 1, 37077 Göttingen, Germany (mmirian@gwdg.de,

gwoerne@gwdg.de) 2 Université Blaise Pascal, Clermont Ferrand, France (J.C.Thouret@opgc.univ-bpclermont.fr)

KEYWORDS : minor volcanic centres, crust, tectonic erosion, Central Andes, isotopes

Introduction

Geochemical studies of Tertiary to Recent magmatism in the Central Volcanic Zone have mainly focused on

large stratovolcanoes. This is because mafic minor volcanic centres and related flows that formed during a single

eruption are relatively rare and occur in locally clusters (e.g. Andagua/Humbo fields in S. Peru, Delacour et al.,

2007; Negrillar field in N. Chile, Deruelle 1982) or in the back arc region (Davidson and de Silva, 1992). These

studies showed that the "monogenetic" lavas are high-K calc-alkaline and their major, trace, and rare elements,

as well as Sr-, Nd- and Pb- isotopes data display a range comparable to those of the Central Volcanic Zone

composite volcanoes (Delacour et al., 2007). It has been argued that the eruptive products of these minor centers

bypass the large magma chamber systems below Andean stratovolcanoes and thus may represent magmas that

were derived from a deeper level in the crust (Davidson and de Silva, 1992; Ruprecht and Wörner, 2007). This

study represents a continuation of our work to understand the regional variation in erupted magma composition

in the Central Andes (Mamani et al., 2008; Wörner et al., 1992). Here we concentrate on the northern boundary

of the Arequipa Pb-domain (Mamani et al., 2008) in the Colca and Cotahuasi valley regions.

Distribution of minor volcanic centres

Minor centres of late Pleistocene to Historical age (< 1 Ma, Delacour et al., 2007) are found in southern Peru

in the Andahua, Huambo, Llauce, Caylloma fields and outcrops in Auquihuato, Iquipi and Yura area (Fig. 1).

We also include lavas of Llauce valley in the Ocoña Cañon, which have Pliocene ages (2.27 ± 0.05 Ma, Thouret

et al., 2007; 2.261 ± 0.046 Ma, Schildgen et al., 2007). Cinder and scoria cones of the younger fields are all well

preserved and cones have a typical height of 200-300 m and are 500-650 m across. Apparently most of the cones

lie on valley floors. However, this may be an artifact and result from preferential accumulation into the valleys

and enhanced erosion by glaciers at high altitudes. Some lava flow associated with the cones extents as far as 4

to 8 km and thick lavas cover the floor of Andahua and Llauce valleys and act as natural dams. Within the

Llauce valley lava dams are associated to large outburst-floods. Thinner lava dams cover the Huambo valley,

Auquihuato and Sibayo area (Fig. 1a). Petrographic types are basaltic andesites, andesites and dacites (Fig. 2a).

The most mafic sample is from Nicholson centre with SiO2 52.3%. Plagioclase is the prominent mineral phase

and clinopyroxene and Fe-Ti oxides are present in all lavas. Where Plagioclase is less abundant, olivine and

clinopyroxene occur higher but in equal proportion. Hornblende and orthopyroxene appear in andesites and

biotite phenocrysts are found only in dacites (Delacour et al., 2007). According to the Pb-isotope domain map of

Mamani et al. (2008), the Iquipi, Huambo and Yura centres occur within the Arequipa domain whereas the

Llauce lavas, Auquihuato, Andahua and Caylloma centres occur within the northern Cordillera domain.

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Isotopic composition

A striking characteristic of the minor centres in this region is their systematic variation Sr-, Nd-, Pb- isotopic

data with an abrupt change (within 60 km) in isotopic compositions between Arequipa and northern Cordillera

domain (Fig. 1). The eNd values (-2.5 to -4.5), 87Sr/86Sr ratios (0.7055 to 0.7065) and 206Pb/204Pb ratios (18.56 to

18.78) of minor volcanic centres within the northern Cordillera domain encompass the entire range recorded

from the large stratovolcaneos in this domain (e.g. Sara Sara, Coropuna, Solimana and Antapuna volcanoes).

Only dacite sample from Puca Mauras centre have 206Pb/204Pb ratios (18.53 to 18.57) like lavas from Arequipa

domain and eNd values around -5 that plot between both domains (Fig. 2b). Equally, eNd values (-5 to -6.3), 87Sr/86Sr ratios (0.7065 to 0.707) and 206Pb/204Pb ratios (18.23 to 18.58) of minor volcanic centres in the Arequipa

domain cover most of the same range observed in stratovolcanoes to the S of the domain boundary in southern

Peru (e.g. Sabancaya, Chachani, Misti, Ubinas, Huaynaputina, Yucamane, Tutupaca, Ticsani volcanoes). An

andesite sample from Marbas Chico has 206Pb/204Pb ratios of 18.58 like lavas from the northern Cordillera

domain (i.e. basaltic andesite of Llauce valley). This implies that the isotopic signatures are really different

between minor centers and large stratovolcanoes within a given region, but both change their geochemical

character when crossing the boundary between crustal domains.

Fig. 1. a) Present-day 206Pb/204Pb ratios map and location of the minor volcanic centres and related lava flows. Thick black lines are the main faults in the study area. Gray arrows are the directions of plate convergence vector according to Norabuena et al. (1998). b) Fig. 3. Schematic cross section showing interpretation of the northern boundary of Arequipa domain. Red line is the Ichupampa fault. 4-Puca Mauras, 5-Angahua, 8-Tischo, 9-Ninamama, 10-Chilcayoc, 17-Marbas Chico, 18 Huambo, 20-Nicholson.

Crustal contamination of minor volcanic centres

The amount of crustal contamination in typical andesite is 16% according to EC-AFC modeling (Chang,

2007). However, the composition of the assimilated crustal component is variable in the two domains (Fig. 1, see

Mamani et al., 2008 for a full discussion). TDM ages for lavas in the Cordillera domain vary between 0.8 and 1.1

Ga., and contaminated magmas in the Arequipa domain have TDM ages from 1.3 to 1.5 Ga.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 298-301

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Loewy et al. (2004) published TDM ages from 1.9 to 2.3 Ga. of the Arequipa basement. Interestingly, the Nd

model ages correlate nicely with the Pb-isotopic composition of contaminated magmas (Fig. 2d) and this

suggests that minor volcanic centres derive their assimilated component from crust of different age and

composition (Fig. 1b). The fact that we highlighted above, i.e. that minor centres and large stratovolcanoes are

not distinct in their isotopic character, allows to define the boundary between the domains of different

assimilated crust with much better local spatial constraint because these minor centers happen to be particularly

abundant in this region (Fig. 1). Therefore, we demonstrate that the domain boundary is surprisingly abrupt

(within 60 km laterally for a crust that is > 60 km thick), which suggest that the boundary most likely is

relatively steep. If so, the boundary probably represents a major, crustal suture between distinct crustal blocks.

As the isotopic difference is very large, this implies that these blocks must have had a long (> 1Ga) distinct

geochemical history. It is therefore surprising to find that this region shows a system that runs along the crustal

domain boundary (Iquipi fault, Roperch et al., 2006). It has been argued also that the eruptions of minor centers

were controlled by regional scale faults (Huanca and Uchupampa faults, Antayhua et al., 2001). If so, then these

eruptions indeed are fed from deeper level magma storage areas, which implies that both, minor centers and

large stratovolcanoes receive their crustal imprint equally at depth and that shallow crustal assimilation is not a

major process in determining the isotopic composition of Central Andean magmas.

Lower crustal assimilation or mantle source contamination?

Lower crustal assimilation may occur in MASH or "Hot Zones" (Hildreth and Morbath, 1988; Annen et al.,

2006) and there is no doubt to us that a major portion of the crustal signature in Central Andean Arc magmas

derives from crustal assimilation. As the Peru-Chile trench is almost free of sediments and no accretionary prism

is observed (von Huene et al., 1999) the subduction of sediments into the mantle wegde source region for

Central Andean magmas is not expected. However, tectonic erosion of the forearc region in northern Chile and

southern Peru is a well-established process (von Huene et al., 1999; Stern, 1991a) and has more recently been

emphasized again for affecting magma genesis in the Central Andes (Kay et al., 2005) and was quantified in

more detail by Clift and Hartley (2007). However, the question remains whether such tectonically eroded forearc

Fig. 2. a) Classification of calc-alkaline series. b) Plot of eNd values versus 87Sr/86Sr ratios showing the isotope range for the minor volcanic centres of the Cordillera (CD) and Arequipa (AD) domains. c) Pb isotope composition of the minor volcanic centres. The upper crust (U), orogen (O) and mantle (M) evolution curves are from Zartman and Doe (1981). d) Diagram of TDM ages versus Pb-isotopes of minor and mayor volcanic centres of the CD and AD. Figs. (a) and (b) are compared to the North Volcanic Zone (NVZ, Bourdon et al., 2002), South Volcanic Zone (SVZ, Kay et al., 2005) and Austral Volcanic Zone (AVZ, Stern and Killian 1996). Minor volcanic centres (MiVC), Major volcanic centres (MaVC),

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 298-301

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material is actually subducted to >100 km depth into the magma generation, or whether the eroded material is

quantitatively underplated below the forearc region (Clift and Hartley, 2007). New studies of O isotopes and U-

Th isotopes now show that limited source contamination of 1-2% for lavas of El Misti and possibly other Central

Andean volcanoes (Chang, 2007; Kiebala, 2008) in addition to lower crustal assimilation. Our study has

significant implication for this discussion. If subduction of tectonically eroded material from the forearc would

be the main process controlling the isotopic composition of the erupted magmas (i.e. no or limited crustal

assimilation, Stern 1991a), then the isotopic composition of forearc rocks would directly project downward

parallel to the plate convergence vector. In this case, Pb-isotope domain boundaries in the erupted magmas

should all be parallel to the plate vector motion. This is in fact what we observe (Fig. 1). However, plate vectors

have changed through time and domain boundaries have remained constant through time. This is shown by the

fact that young and old (>30 Ma) rocks all show he same domain distribution. Thus we conclude that

assimilation in the deep crust still is the main process that determines the isotopic composition of Central

Andean magmas and that defines the domain boundaries. The effect of limited tectonic erosion, however, cannot

be excluded.

References Annen, C., Blundy, J.D. & Sparks, R.S.J., 2006. The Genesis of calcalkaline intermediate and silicic magmas in deep crustal

hot zones. J. Petrol., 47, 505-539. Antayhua, Y., Tavera, H., & Bernal, I., 2001. Analisis de la actividad sismica en la region del volcán Sabancaya (Arequipa).

Bol. Soc. Geol. Perú, 92, 78–79. Bourdon, E., Eissen, J., Monzier, M., Robin, C., Martin, H., & Hall, M.L., 2002. Adakite-like lavas from Antisana Volcano

(Ecuador); evidence for slab melt metasomatism beneath the Northern Andean Zone. J. Petrol, 43-2, 199-217. Chang, Y. 2007, O-isotopes as Tracer for Assimilation Processes in Different Magmatic Regimes (El Misti, S.Peru and

Taapaca, N. Chile). Master Thesis, Göttingen University. Clift, D., & Hartley, A.J., 2007. Slow rates of subduction erosion and coastal underplating along the Andean margin of Chile

and Peru. Geology, 35, 503–506, doi: 10.1130/G23584A.1. Davidson J P, & de Silva S L, 1992. Volcanic rocks from the Bolivian Altiplano: insights into crustal structure,

contamination, and magma genesis in the central Andes. Geology 20: 1127-1130. Delacour, A., Gerbe,M.C., Thouret, J. C., Wörner, G., & Paquereau-Lebti, , 2007. Magma evolution of Quaternary minor

volcanic centres in Southern Peru, Central Andes. Bull. Volcanol. 69: 581–606. Deruelle B (1982) Petrology of Plio-Quaternary volcanism of the south central and meridional Andes. J. Vol. Geo. Res. 14:

77–124. Hildreth, W., & Moorbath S., 1988. Crustal contributions to arc magmatism in the Andes of central Chile. Contr. Min. Petrol. 98: 455-489.

Kay, S.M., Godoy, E., & Kurtz, A., 2005. Episodic arc migration, crustal thickening, subduction erosion, and magmatismo in the south-central Andes. Geol. Soc. Am. 117: 67-88.

Kiebala, A., 2008. Magmatic Processes by U-Th disequilibria comparison of two Andean magmatic systems: El Misti (S. Peru) and Taapaca (N. Chile). PhD thesis, Göttingen University.

Loewy, S.L., Connelly, J.N., & Dalziel, I.W.D., 2004, An Orphaned Basement Block: The Arequipa-Antofalla Basement of the Central Andean margin of South America: Geol. Soc. Am. Bull. 116: 171-187.

Mamani, M., Tassara, A., & Wörner G., 2008. Composition and structural control of crustal domains in the central Andes, Geochem. Geophys. Geosyst. 9, doi:10.1029/2007GC001925.

Norabuena, E., Leffler-Griffin, L., Mao, A., Dixon, T., Stein, S., Sacks, I., Ocala, L., & Ellis, M., 1998. Space geodetic observation of Nazca–South America convergence across the Central Andes. Science 279: 358–362.

Roperch, , Sempere, T., Macedo, O., Arriagada, C., Fornari, M., Tapia, C., & Laj, C., 2006. Counterclockwise rotation of late Eocene–Oligocene fore-arc deposits in southern Peru and ist significance for oroclinal bending in the central Andes. Tectonics 25: TC3010.

Ruprecht, & Wörner, G., 2007. Variable regimes in magma systems documented in plagioclase zoning patterns: El Misti stratovolcano and Andahua monogenetic cones. J. Vol. Geo. Res. 165 (3): 142-162.

Stern, C.R., 1991a. Role of subduction erosion in the generation of Andean magmas. Geology 19: 78-81 Stern, C.R., & Kilian R., 1996. Role of the subducted slab, mantle wedge and continental crust in the generation of adakites

from the Andean Austral Volcanic Zone. Contrib. Mineral. Petrol. 123: 263-281. Wörner, G., Moorbath, S., & Harmon, R.S. 1992. Andean cenozoic volcanics reflect basement isotopic domains. Geology, 20: 1103-1106.

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Seismicity and structural implications in northern Ecuador from the Esmeraldas experiment

K. Manchuel1, B. Pontoise

2, N. Béthoux

1, M. Régnier

3, & the ESMERALDAS team

1 UMR Géosciences Azur, Port de la Darse, BP48, 06235 Villefranche sur Mer, France

2 UMR Géosciences Azur / IRD, IRD-LGTE Case 119, Université Pierre et Marie Curie, 4, place Jussieu, 75252

Paris, France 3 UMR Géosciences Azur / IRD, 250 rue Albert Einstein, Sophia-Antipolis, 06560 Valbonne, France

The North-Andean margin is being deformed by the subduction of the Nazca plate (5-7 cm/y) along a N80°

direction. The Nazca plate carries the Carnegie ridge, a 200 km-wide buoyant ridge (figure 1), which subducts

under the Ecuadorian central margin involving major crustal deformation. The northern flank of the Carnegie

ridge divides the Ecuador-Colombian margin in two seismically and tectonically contrasted segments [Collot, et

al., 2002; Gutscher, et al., 1999; Pontoise and Monfret, 2004]:

a) the northern segment (north of 0.5°N) which is subsident and where 4 megathrust earthquakes occurred in

the couple zone during the 20th century. The 500km long rupture zone of the 1906 event (Mw=8.8) was

partially reactivated by 3 thrust events occurring in 1942 (Mw=7.8), 1958 (Mw=7.7) and 1979 (Mw=8.2).

Almost, all centroïd moment tensor solutions from Harvard catalog, during the 1976-2001 period are of

thrust-type mechanisms.

b) the southern segment (south of 0.5°N) which is undergoing a general uprising and had no registered large

earthquake during the last century, but presents a seismicity organized in earthquake swarms.

Figure 1: Geodynamic sketch of the Ecuadorian active margin and networks location on the field. Dashed areas represent the surface rupture of the four great subduction earthquakes that occurred during the 20th century. Large black arrow is the Nazca plate motion vector [Trenkamp, et al., 2002]. DGM= Dolores-Guayaquil-Megashear. The North Andean Block is being displaced to the northeast along the DGM. Yellow triangles are for ESMERALDAS stations (OBSs and land stations).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 302-305

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Some marine seismic experiments conducted during the last decade along the Ecuador-Colombia active

margin and yeldied a detailed knowledge of the margin structure [Calahorrano, 2005; Collot, et al., 2002;

Gailler, et al., 2007; Graindorge, et al., 2004]. Inland, the IG-EPN (Instituto Geofisico de la Escuela Politecnica

Nacional) seismological network (rensig network), mainly located on the cordillera in order to survey the

volcanic activity, allows to constrain the cordillera seismicity but gives poor determination of the fore-arc

seismic activity. So far, global catalogs provide a diffuse image of the coastal seismicity, and only few

seismological studies were conducted in the Esmeraldas region [Guillier, et al., 2001; Pontoise and Monfret,

2004]. The scarcity of seismicity data between the coast and the cordillera led authors to propose different slab

geometries in northern Ecuador. Some of them postulate a slab dipping eastward with an angle of 25° to 40°

[Guillier, et al., 2001; Pontoise and Monfret, 2004; Taboada, et al., 2000] and others suggest a 100km deep flat

slab [Gutscher, et al., 1999; Gutscher, et al., 2000].

Figure 2: a) Map of epicenters. Red dots present ESMERALDAS experiment locations. Black line show trace of the projection plane on the surface and the brackets show width of the projection. b) Cross section along the north-Ecuadorian margin.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 302-305

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Using data of a local temporary network, deployed during the ESMERALDAS experiment our aim is to

obtain a better definition of the seismic zones in the Esmeraldas region, between the margin and the Andes. This

experiment was carried out from 10 march to 14 june 2005. We used land and marine seismological networks

allowing a good azimuthal coverage (figure 1). 26 OBS (Ocean Bottom Seismometer) and 31 land stations were

deployed simultaneously.

We use the SEISAN [Havskov and Ottemöller, 2000] tool to create a data base, read seismograms and locate

events, with the Hypocenter code used routinely in SEISAN and which allow the use of one single 1D velocity

model. Then, in order to perform locations we locate again events with the Hypoellipse code. This location

technique allows the use of several 1D velocity models, assigned to the stations located in different seismogenic

regions. two velocity models were used. One for the OBS, derived from wide-angle seismic profile across the

margin of Esmeraldas [Agudelo, 2005; Gailler, et al., 2007]. The second one, deduced from CRUST2.0 (a global

earth velocity model specified from 2*2 degree [Bassin, et al., 2000]), is for the land stations. 1091 events were

studied, only 363 could be located and 282 of them exhibit rms better than 1 second. These 282 events are

presented in the figure 2.

From the trench, up to ~40km east of it, no seismicity is detected. At this distance of the trench, the shallowest

seismicity observed is located at a ~10km depth. This observation shows that, east of the trench, there is very

low seismic activity in the shallowest few kilometres of the interplate zone. We interpret this shallow

earthquakes distribution as an indication of the depth of the Updip Limit (UdL) of the seismogenic zone. This is

in agreement with previous results in the same area [Pontoise et Monfret, 2004], and with seismic events

distribution in Manta area, constrain by an active seismic image [Bethoux et al., submitted].

Concerning the slab geometry, we propose a slab dipping at ~35° from the trench up to, at least, 110 km depth.

This dip was already proposed by Taboada et al. [2000] and Guiller and Chatelain [2001]. The hypothesis of a

100 km deep flat slab [Gutscher, et al., 1999; Gutscher, et al., 2000] seems to be unproved by the presence of

earthquakes deeper than 150 km located in the continuity of the slab geometry defined under the coastal block.

Due to scarce onland data, some authors did not observed crustal upper plate seismicity in the Esmeraldas

region [Guillier, et al., 2001; Pontoise and Monfret, 2004]. So they propose that the coastal block of Ecuador,

composed of several accreted oceanic blocks [Cediel, et al., 2003], acts as an undeforming body. Now, the dense

ESMERALDAS network allows evidencing of crustal seismicity beneath the coastal block and the western slope

of the Andes, organised along structures dipping eastward and westward and reaching a 40km depth. The

geometry of structures and focal mechanisms implie a compressional stress regime across the coastal block of

North Ecuador, up to the western slope of the Andes.

We note the presence of earthquakes immediately west of the trench. This seismicity might be due to the

bending of the slab. Because this seismicity is deep (down to 50 km) we rather suggest that it also correspond to

the compressional stress regime observed in the coastal block which is extended west of the trench. We therefore

interpret the organisation of this seismicity as the presence of a new thrust zone west of the trench, suggesting

the accretion of a new oceanic block. The use of focal mechanisms and a local earthquake tomography will allow

us to constrain this interpretation.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 302-305

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References Agudelo, W. (2005), Imagerie sismique quantitative de la marge convergente d'Equateur-Colombie: Application des

méthodes tomographiques aux données de sismique réflexion multitrace et réfraction-réflexion grand-angle des campagnes SISTEUR et SALIERI, Thèse de doctorat thesis, 203 pp, Pierre et Marie Curie, Villefranche sur mer.

Bassin, C., et al. (2000), The Current Limits of Resolution for Surface Wave Tomography in North America, Eos Transactions American Geophysical Union, 81, 897.

Calahorrano, A. (2005), Structure de la marge du Golfe de Guayaquil (Equateur) et propriétés physiques du chenal de subduction, à partir de données de sismique marine réflexion et réfraction., PhD Thesis, Université P. et M. Curie, Paris VI, 221 p.

Cediel, F., et al. (2003), Tectonic assembly of the Northern Andean Block, The Circum-Gulf of Mexico and the Caribbean: Hydrocarbon habitats, basin formation, and plate tectonics: AAPG, 815-848.

Collot, J. Y., et al. (2002), Exploring the Ecuador-Colombia Active Margin and Interplate Seismogenic Zone, Eos Transactions American Geophysical Union, 83, 185.

Gailler, A., et al. (2007), Segmentation of the Nazca and South American plates along the Ecuador subduction zone from wide angle seismic profiles, Earth and Planetary Sciences Letters, 260, 444-464.

Graindorge, D., et al. (2004), Deep structures of the Ecuador convergent margin and the Carnegie Ridge, possible consequence on great earthquakes recurrence interval, Geophysical Research Letters, 31.

Guillier, B., et al. (2001), Seismological evidence on the geometry of the orogenic system in central-northern Ecuador (South America), Geophysical Research Letters, 28, 3749-3752.

Gutscher, M.-A., et al. (1999), Tectonic segmentation of the North Andean margin: impact of the Carnegie Ridge collision, Earth and Planetary Sciences Letters, 168, 255-270.

Gutscher, M.-A., et al. (2000), Geodynamics of flat subduction: Seismicity and tomographic constraints from the Andean margin, Tectonics, 19, 814-833.

Havskov, J., and L. Ottemöller (2000), SEISAN earthquake analysis software, Seismological Research Letters, 70, 532-534. Pontoise, B., and T. Monfret (2004), Shallow seismogenic zone detected from an offshore-onshore temporary seismic

network in the Esmeraldas area (northern Ecuador), Geochemistry, Geophysics, Geosystems, 5, 22. Taboada, A., et al. (2000), Geodynamics of the northern Andes: Subductions and intracontinental deformation (Colombia),

Tectonics, 19, 787-813. Trenkamp, R., et al. (2002), Wide plate margin deformation, southern Central America and northwestern South America,

CASA GPS observations, Journal of South American Earth Sciences, 15, 157-171.

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Influence of trench sedimentation rate on heat flow and location of the thermally-defined seismogenic zone in the North Ecuador – South Colombia margin

B. Marcaillou1, G. Spence

2, K. Wang

3, J.-Y. Collot

4, & A. Ribodetti

4

1 IFREE/JAMSTEC, Japan (boris.marcaillou@gmail.com)

2 School of Earth and Ocean Sciences, University of Victoria, Canada

3 Pacific Geoscience center, Geological Survey of Canada, Canada

4 Géosciences Azur, Université de Nice Sophia Antipolis, IRD, Université Pierre et Marie Curie, CNRS,

Observatoire de la Côte d’Azur, France

Introduction

The limited region of interplate contact that can generate thrust earthquakes, known as the seismogenic zone,

is at least partially temperature-dependant [Hyndman et al., 1997]. The updip and downdip limits of this

seismogenic zone are commonly relate to 60-150°C and 350-450°C, respectively [Hyndman and Wang, 1993;

Saffer and Marone, 2003]. Thermal modelling in convergent margin aims at estimating the temperature

distribution along interplate contact in order to propose a location for these temperatures ranges and thus for the

seismogenic zone.

The amount of trench sediments supplied to the subduction system is widely known to impact the tectono-

structural, mechanical and seismological framework of convergent margins. Numerous authors claimed the

influence of trench sediments on the tectonic regime at the deformation front [Clift and Vannucchi, 2004;

Lallemand et al., 1994], on the mechanical interplate friction [Calahorrano et al., 2008] and the onset of the

stick-slip behaviour along mega-thrust faults [Cloos and Shreve, 1996]. However thermal modelling usually

neglect the sediment loading in the trench by considering that the trench fill thickness have homogeneously

deposited over the incoming plate through times since the oceanic crust formation.

In Ecuador – Colombia, depth-migrated multichannel seismic reflection (MCS) data allows to assess the

evolution of the sedimentation rate over the oceanic crust as it approaches the trench. By performing at various

latitude along the margin, thermal models that include the sediment loading and compaction in the trench we

investigate the impact of trench sedimentation variation on the temperature distribution along the interplate

contact and thus on the seismogenic zone location.

Structural settings: sedimentation rate in the trench

The North Ecuador – South Colombia (NESC) Margin (1-4°N) divides into three segments, named Patía,

Tumaco and Manglares (Fig. 1), with different structure and tectonic regime [Marcaillou, 2003]. Among other

features, the trench-fill thickness varies along-strike with sediments accumulation three-time thicker in the

central Tumaco segment than in the northern Patía segment. In contrast, 30 km to the west, in the abyssal plain

the hemipelagic sedimentary layer is homogeneous along-strike as substantiated by seismic lines parallel to the

trench. This implies along-strike variations in sedimentation rate in the vicinity of the trench estimated to be

~height-time higher in the Tumaco segment than in the Patia segment.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 306-309

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Seismogenic zone location

Along the North Andean margin, four great subduction earthquakes occurred in 1906, 1942, 1958 and 1979

[Beck and Ruff, 1984; Swenson and Beck, 1996] (fig. 1). The 3-month aftershocks beneath the overriding plate

are distributed in an area extending seaward to near the deformation front for the 1979 main shock, but restricted

to 40 km landward for the 1958 event [Mendoza and Dewey, 1984]. The distribution of thrust events in the

Harvard University Centroid Moment Tensor (CMT) archive shows similarly varying patterns along the NESC

margin (fig. 1). Moreover, The coseismic rupture defined by inversion of seismic data for the 1979 event appears

to have extended very close to the trench in the Tumaco segment, whereas in 1958 it stopped ~30 km landward

of the trench in the Manglares segment [Kanamori and McNally, 1982] (fig. 1). These data consistently suggest

that the seismogenic zone extends seaward to near the trench beneath the Patia and Tumaco segments but is

restricted farther landward beneath the Manglares segment.

Figure 1: bathymetric map Figure 2: Heat flow map

Figure 3: thermal model along line SIS-40 Thermal model along line SIS-37

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 306-309

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Data and Method

On the margin and in the trench, the heat flow at the surface was measured during the AMADEUS experiment

(2005) and derived, using a classical calculation [Yamano et al., 1982], from Bottom Simulating Reflectors

(BSRs) in MCS lines recorded during the AMADEUS and SISTEUR (2000) experiments. Both data-set are

consistent within a 10% estimated uncertainty in calculation. These data results in a heat flow map highlighting

along-strike thermal variations with a heat flow in the Tumaco segment (55-65 mW m-2) 50% lower than in the

Patía segment (100-110 mW m-2) (fig. 2).

We investigate the thermal structure of every margin segment with a 2-D steady-state finite element method

[Wang et al., 1995] that takes into account for plate ages, convergence rate, margin structure, sedimentation and

compaction rates in the trench and isoviscous corner flow in the continental mantle.

Results

These models suggest that:

1/ The temperature range from 60-150°C to 350-450°C, commonly associated with the updip and downdip

limits of the seismogenic zone, extends along the plate interface over a distance of 160 to 190 ± 20 km (fig. 3).

2/ The updip limit of the seismogenic zone for the great subduction earthquakes during the 20th century is

associated with low-temperature (60-80°C) processes.

3/ 60-70% of the two-fold decrease in measured heat flow from the Patia to the Tumaco segment is related to

the abrupt southward increase in sedimentation rate in the trench. Such a change may induce a landward shift of

the 60-150°C isotherms, and thus the updip limit of the seismogenic zone, by 10-20 km. As a result, the

sedimentation history of the oceanic plate prior to subduction is a key-parameter of the thermal structure for

convergent margins and should not be neglected in thermal modelling.

References Beck, S. L., and L. J. Ruff (1984) The rupture process of the great 1979 Colombia earthquake: evidence for the asperity

model, J. Geophys. Res., 89, 9281-9291. Calahorrano, A., et al. (2008) Nonlinear variations of the physical properties along the southern Ecuador subduction channel:

results from depth-migrated seimic data, Earth Planet. Sci. Lett., 267 (3-4), 453-467. Clift, P. D., and P. Vannucchi (2004) Controls on tectonic accretion versus erosion in subduction zones: Implications for the

origin and recycling of the continental crust, Rev. Geophys., 42, RG2001, doi:2010.1029/2003RG000127. Cloos, M., and R. L. Shreve (1996) Shear-zone thickness and seismicity of Chilean- and Marianas-type subduction zones,

Geology, 24 (2), 107-110. Hyndman, R. D., and K. Wang (1993) Thermal constraints on the zone of the major thrust earthquakes failure: the Cascadia

subduction zone, J. Geophys. Res., 98 (2039-2060). Hyndman, R. D., et al. (1997) The seismogenic zone of subduction thrust faults, Island Arc, 6 (3), 244-260. Kanamori, H., and K. C. McNally (1982) Variable rupture mode of the subduction zone along the Ecuador-Colombia coast,

Bull. Seis. Soc. Am., 72 (4), 1241-1253. Lallemand, S. E., et al. (1994) Coulomb theory applied to accretionary and non accretionary wedges: Possible causes for

tectonic erosion and/or frontal accretion, J. Geophys. Res., 99, 12,033-012,055. Marcaillou, B. (2003) Régimes tectoniques et thermiques de la marge Nord Équateur- Sud Colombie (0°- 3,5°N) -

Implications sur la sismogenèse, Phd thesis, 220 pp, Université de Pierre et Marie Curie, Paris. Mendoza, C., and J. W. Dewey (1984) Seismicity associated with the great Colombia-Ecuador earthquakes of 1942, 1958

and 1979: implications for barrier models of earthquake rupture, Bull. Seis. Soc. Am., 74 (2), 577-593. Saffer, D. M., and C. Marone (2003) Comparison of smectite- and illite-rich gouge frictional properties: application to the

updip limit of the seismogenic zone along subduction megathrusts, Earth Planet. Sci. Lett., 215, 219-235.

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Swenson, J. L., and S. L. Beck (1996) Historical 1942 Ecuador and 1942 Peru subduction earthquakes, and earthquake cycles along Colombia-Ecuador and Peru subduction segments, Pageoph, 146 (1), 67-101.

Wang, K., et al. (1995) Thermal regime of the southwest Japan subduction zone: effects of age history of the subducting plate, Tectonophysics, 248, 53-69.

Yamano, M., et al. (1982) Estimates of heat flow derived from gas hydrates, Geology, 10, 339-343.

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Andesite magma generation at the Quaternary volcanic arc of southwest Colombia

M. I. Marín-Cerón1,2

, T. Moriguti1, & E. Nakamura

1

1 Pheasant Memorial Laboratory (PML), Institute for Study of the Earth’s Interior (ISEI), Okayama University at

Misasa, Yamada 827, Misasa Tottori, Japan 2 Present adress: EAFIT University, Department of Geology, Medellin, Colombia (mmarince@eafit.edu.co)

KEYWORDS : NVZ, Andes, lower crust, mantle-derived magmas, andesites

Introduction

There is a general consensus about the formation of arc magmas by the addition of slab-derived fluids,

triggering the partial melting of the mantle wedge (e.g. Hawkesworth et al., 1993a; Pearce & Parkinson, 1993;

Poli & Schmidt, 1995; Tatsumi & Eggins, 1995; Davidson, 1996). Once primary magmas are formed in the

mantle wedge, an open question appears: which processes control the intermediate and silicic arc magma

generation? Two main processes have been identified: (1) differentiation of primary magmas by crystallization

within the crust or uppermost mantle (e.g. Gill, 1981; Grove & Kinzler, 1986; Musselwhite et al., 1989) and (2)

partial melting of older crustal rocks (e.g. Smith & Leeman, 1987; Petford & Atherton, 1996; Chappell & White,

2001). A combination of the above mentioned processes is also possible in which the interaction of mantle-

derived magmas can trigger the crustal melting and enhance the assimilation fractional crystallization of crustal

rocks (AFC) proposed by DePaolo (1981) or mixing, assimilation, storage and hybridization (MASH) proposed

by Hildreth & Moorbath (1988).

Andean volcanic zones are among the most important regions to study in order to better understand andesite

magma generation at convergent margins because the volcanic activity is extended along more than 4000 km of

Andes Cordillera. The volcanism in the Andes can be subdivided into four zones (Fig. 1a): the Southern

Volcanic Zone (SVZ), Central Volcanic Zone (CVZ) and Northern Volcanic Zone (NVZ) result from subduction

of Nazca plate beneath the Andean block, and the Austral Volcanic Zone (AVZ), which is related to the

subduction of the Antartic plate (e.g. Thorpe et al., 1982; Simking and Siebert, 1994).

Several studies have been undertaken at the Andes, mainly at the SVZ and the CVZ. However, the NVZ,

especially in Colombia, is poorly studied, and the available data is insufficient to understand the magmatic

processes and the relation of geochemical variations of those volcanoes with the spatial distribution. In order to

understand the Andes volcanic zones from a global perspective, a systematic study of Colombian arc volcanism

in the NVZ is indispensable. In this study we have chosen to study the Southwestern Colombian arc (Fig. 1b),

which is an arc with a simple geophysical structure characterized by relatively constant Moho depth (35-40 km)

across the arc and small differences in the dip of the seismic zone (25°- 30°), with volcanoes lying at 120 to 200

km above the Wadati-Benioff Zone (WBZ). In this region, Marín-Cerón (2007) has proposed on the basis of

multi-isotopic systematics that due to the higher biogenic activity at this region of the Pacific Ocean, fluids from

carbonate-rich sediments have been introduced into the mantle wedge beneath the study area affecting mainly

the volcanic front primary magmas.

In this study we propose a model for andesite magma generation in the SW Colombian volcanic arc on the

basis of petrography, major and trace element, and Sr, Nd, Pb, Hf isotopic systematics. Available data from

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lower crustal xenoliths erupted by volcanic tuffs during the Tertiary (Weber et al., 2002) provide additional

constraints.

Figure 1. (a) Volcanic zones distribution in the Andes Cordillera. (b) Tectonic map of the SW Colombian volcanic arc. Tectonic setting of the NVZ in the Andes showing the distribution of active and inactive ridges, min fault systems and the distribution of active volcanoes. Waddati-Benioff Zone (WBZ) contours represent the isobaths for the top of the deep seismic caused by the subduction of Nazca plate beneath the North Andean block. (Modified from Gustcher et al. (1999); Droux and Delaloye (1996) and reference there in).

Samples

Samples were collected from four southwestern Plio-Quaternary volcanoes: Azufral, Galeras, Dona Juana and

Purace-Coconucos; they are located 140 km, 160 km, 170 km and 190 km to the depth of Wadati-Benioff zone

respectively. The calc-alkaline andesites and dacites of southwestern Colombian volcanoes are generally

porphyritic (modal phenocryst up to 50%). Andesites at this region can be divided petrographically and

geochemically into two groups: (1) volcanic front (VF) formed by andesites from Azufral and Galeras volcanoes

and (2) rear arc (RA) andesites from Doña Juana and Purace-Coconucos volcanic complexes. The main

mineralogical differences in both groups are related to the modal abundances of Ca-rich pyroxenes, small

amounts of amphibole (~1%) and the absence of olivine and quartz at the VF compared to the RA.

Geochemically, all analyzed sixty samples are quite evolved, with silica contents > 53%. The lavas belong to

medium-K in the volcanic front area to high-K in the back-arc, with enrichment in total alkalis across-arc. In

primitive normalized pattern diagram, positive anomalies are observed in fluid mobile elements such as B, Pb, Sr

and Li. On the other hand, high field strength elements show negative anomaly. These features and the clear

across-arc variation in the Ba/Nb ratio may indicate that studied samples were generated by fluid related

processes. The primary magmas at this arc are defined as mantle-derived magmas metazomatized by the

subduction component which were decreasingly added to the mantle wedge with the depth of the Waddati-

Benioff zone (Marín-Cerón, 2007; Marin-Cerón et al., 2008). The generated magmas are fluid-rich basalts

isotopically and geochemically different at the VF compared to the RA related with the amount of fluids derived

from the slab dehydration and decarbonation (Marín-Cerón, 2007; Marín-Cerón et al., 2008).

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Petrological constraints for andesites formation

Our available petrographical and geochemical data indicates that the andesites at the study area cannot be

derived only by fractional crystallization of fluid rich mantle-derived basalts. The several disequilibrium features

above mentioned, the lack of basalts in the study area and the clear binary mixing trends in Pb-isotopic

systematics between primary magmas and lower continental crust (LCC) which is Pb-radiogenic, suggest the

importance of assimilation of LCC materials by mantle-derived arc-magmas.

Experimental data for the formation of intermediate to silicic magmas (Annen et al., 2006 and reference

therein), is supportive to assume that the main mineral phases (pl+amp+cpx+opx) of the andesites in the SW

Colombian arc were coexisting the LCC. Moreover, it is necessary to invoke assimilation and/or mixing with the

surrounding radiogenic-Pb-bearing lower crust rocks (Weber et al., 2002), to explain the isotopic and

geochemical signature of these rocks compared to other andesites in the Andes.

The main petrological constraints are: (1) Clinopyroxene under high-H2O conditions in the lower-crustal

environment, and at lower pressure conditions, indicates that when the temperature of the melts drops,

clinopyroxene becomes unstable, and reacts with the melt to form amphibole, resulting in the evolved melt being

more siliceous (Foden & Green, 1992). (2) Amphibole is stable only for H2O contents 4 wt% (Eggler, 1972)

and temperatures below ~1050°C (e.g. Muntener et al., 2001). (3) Orthopyroxene is confined to relatively low

pressures and temperatures over 920°C whereas garnet appears only above ~1.1GPa. (4) Plagioclase stability

decreases and An content increases with increasing H2O. Based on the experimental data of Kawamoto (1996)

and Pichavant et al. (2002b) it is possible to infer that the maximum H2O content of andesite melt in equilibrium

with plagioclase (An>80) is ~ 10 wt% H2O.

After the main mineral phases started to crystallize at the lower crust, the intermediate to silicic magma is

developed depending of the gradients of P, T and H2O content. The produced magmas at this region may get the

geochemical flavour of the lower crust which is rich in Pb-radiogenic due to the interaction of primary mantle

derived magmas and lower crust materials. In the way to the surface the rising magmas may be stored in shallow

magma reservoirs where mainly crystallization occurs. In such a case, the disequilibrium features observed in the

SW Colombian arc, such as resorbed and/or pseudomorphed amphibole by anhydrous reaction products,

generally replaced by oxides, may indicate that the amphiboles crystallized in the lower crust become unstable at

pressures less than ~0.1 GPa (Rutherford & Hill, 1993). Similarly, the complex zoning in the mineral phases

may be a response of the magma supply from the lower crust, which creates P-T changes in the magma that are

recorded during the complex growth of the plagioclase and pyroxenes. To clarify the above mentioned hypotesis

a detailed multi-isotope analysis in minerals phases following the zoning patterns is much needed.

Conclusion

In a global perspective of the understanding of the volcanic zones at the Andes cordillera, we can conclude that

assimilation of lower crust is a common process in this mature continental arc, and it is not only related to the

thickness of the crust but primarily related to the gradients of temperature, pressure, H2O content and melt

fraction developed at the upper-mantle and lower-crust region. Thus, assimilation of different crust domains in

terms of Pb isotopic composition beneath the Andes cordillera (Figure 2, e.g. Cretaceous domain at the NVZ;

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Proterozoic domain at the north portion of CVZ; Paleozoic to early Mesozoic domain at the south portion of the

CVZ, SVZ and AVZ), may explain the along-arc variation in the Andean volcanic suites. However, the different

subduction components, together with the various thermal regimes at each zone, may explain the variability of

primary magmas across the Andean arc.

Figure 2. Plots of 207Pb/204Pb vs 206Pb/204Pb for the Andean volcanic zones and the pre-andean basement. Arequipa and Barroso volcanics and basement gneisses of South Peru from Tilton & Barreiro (1980). Pacific sediments (Dasch, 1981; White et al., 1985); Precambrian basement (Worner et al., 1992b); Paleozoic basement (Chiaradia et al., 2002); Metalliferous sediments from DSDP leg 92 (Barret et al., 1987); Cretaceous Domain (Keer, 2002); Lower crust xenoliths (Weber, 2002); ACC from Hole 504 (Pedersen and Furnes, 2001).

References Annen, C., Blundy, J. D. & Sparks, R.S, The genesis of intermediate and silicic magmas in deep crustal hot zones. Journal of Petrology

47, 3, p. 505-539, (2006). Barret T. J., Taylor P. N. and Lugowski J., Metalliferous sediments from DSDP leg 92: the East Pacific Rise transects. Geochim.

Cosmochim. Acta 46, pp. 651-666 (1987). Chappell, B. W. & White, J. R., Two contrasting granite types: 25 years later: Australian Journal of Earth Sciences 48, 489–499, (2001). Chiaradia M. & Fontbote L., 2002. Lead isotope systematics of Late Cretaceous – Tertiary Andean arc magmas and associated ores

between 8°N and 40°S: evidence for latitudinal mantle heterogeneity beneath the Andes. Terra Nova, 14 (5), p.337–342 Dasch, E.J, Lead isotopic composition of metalliferous sediments from the Nazca plate. Geol Soc Am Mem 154: 199-200, (1981). Davidson, J. P., Deciphering mantle and crustal signatures in subduction zone magmatism In: Bebout, G. E., Scholl, D. W., Kibry, S. H.,

& Platt, J. P. (eds). Subduction: Top to Bottom. American Geophysical Union 96, p. 251–262, (1996). DePaolo, D. J., Trace-element and isotopic effects of combined wallrock assimilation and fractional crystallisation. Earth and Planetary

Science Letters 53, 189–202, (1981). Droux, A. & Delaloye, M. Petrography and Geochemistry of Plio-Quaternary Calc-AlkalineVolcanoes of Southwestern Colombia.

Journal of South America Earth Sciences. 9, No. 1-2, p. 27-41 (1996). Eggler, D. H., Amphibole stability in H2O-undersaturated calc alkaline melts. Earth and Planetary Science Letters 15, 38–44, (1972). Foden, J. D. & Green, D. H., Possible role of amphibole in the origin of andesite: some experimental and natural evidence.

Contributions to Mineralogy and Petrology 109, 479–493, (1992). Gill, J. B., Orogenic Andesites and Plate Tectonics, Heidelberg: Springer, (1981). Grove, T. L & Kinzler, R. J. Petrogenesis of andesites. 14, 417-454, (1986). Hawkesworth, C. J., Gallagher, K., Hergt, J. M. & McDermott, F., Mantle and slab contributions in arc magmas. Annual Review of

Earth and Planetary Sciences 21, p. 175–204, (1993a). Hildreth W. & Moorbath S., Crustal contamination to arc magmatism in the Andes of Central Chile. Contributions to Mineralogy and

Petrology (1988) 98: p. 455-489, (1988). Kawamoto, T. Experimental constraints on differentiation and H2O abundance of calc-alkaline magmas. Earth and Planetary

Science Letters 144, 577–589, (1996). Kerr, A. C. 2003. Oceanic Plateaus. Treatise On Geochemistry, ISBN (set): 0-08-043751-6 Volume 3; (ISBN: 0-08-044338-9); pp.

537–565.

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Marin-Ceron, M.I. Major, Trace element and multi-isotopic systematics of SW Colombian volcanic arc, northern Andes: Implication for the stability of carbonate-rich sediment at subduction zone and the genesis of andesite magma. Unpublished PhD thesis, Okayama Universtiy, Japan. (2007).

Marín-Cerón, M.I, Moriguti, T. and Nakamura, E. Slab decarbonation and CO2 recycling in the Southwestern Colombian arc, The Misasa 3rd COE-21 International symposium, (2008).

Müntener, O., Kelemen, P. B. & Grove, T. L., The role of H2O during crystallisation of primitive arc magmas under upper most mantle conditions and genesis of igneous pyroxenites: an experimental study. Contributions to Mineralogy and Petrology141, 643–658, (2001).

Musselwhite, D. S., De Paolo, D. J. & McCurry, M. The evolution of a silicic magma system—isotopic and chemical evidence from the Woods Mountains Volcanic Center, Eastern California. Contributions to Mineralogy and Petrology 101, p. 19–29, (1989).

Pearce, J. A. & Parkinson, I. J., Trace element models for mantle melting: application to volcanic arc petrogenesis. In: Prichard, H. M., Alabaster, T., Harris, N. B. W. & Neary, C. R. (eds) Magmatic Processes and Plate Tectonics. Geological Society, London, Special Publications, 76, p. 373–403, (1993).

Pedersen R. & Furnes, H., Nd- and Pb-isotopic variations through the upper oceanic crust in DSDP/ODP Hole 504B, Costa Rica Rift. Earth and Planetary Sc. Lett. 189, p. 221-235, (2001).

Petford, N. & Atherton, M., Na-rich partial melts from newly underplated basaltic crust: the Cordillera Blanca Batholith, Peru. Journal of Petrology 37, 1491–1521, (1996).

Pichavant, M., Martel, C., Bourdier, J. L. & Scaillet, B., Physical conditions, structure, and dynamics of a zoned magma chamber: Mount Peleé (Martinique, Lesser Antilles Arc). Journal of Geophysical Research 107, article number 2093, (2002b).

Poli, S. & Schmidt, M. W., H2O transport and release in subduction zones: experimental constraints on basaltic and andesitic systems. Journal of Geophysical Research 100B, p. 22299–22314, (1995).

Rutherford, M. J. & Hill, P. M., Magma ascent rates from amphibole breakdown—an experimental study applied to the 1980–1986 Mount St. Helens eruptions. Journal of Geophysical Research 98,19667–19685, (1993).

Simking T. & Siebert L., Volcanoes of the world. Geosciences Press, Tuscon, p. 1-349, (1994). Smith, D. R. & Leeman, W. P., Petrogenesis of Mount St. Helens dacitic magmas. Journal of Geophysical Research 92, 10313–10334.

Smith, R. L. (1979). Tatsumi, Y. & Eggins, S., Subduction Zone Magmatism. Oxford: Blackwell Scientific, (1995). Tilton, G.R., and Barreiro, B.A., Origin of lead in Andean calc-alkaline lavas, southern Peru: Science, v. 210, p. 1245-1247. (1980) Thorpe, R. S., Francis, P. W, Hammill M. & Baker M.C.W., The Andes. Andesites. Ed Thorpe, R.S., pp. 187-205, (1982). Weber, M.B.I., Tarney, J., Kempton, P.D. & Kent, R. W., Crustal make-up of the northern Andes: evidence based on deep crustal

xenolith suites, Mercaderes, SW Colombia. Tectonophysics 345, p. 49–82 (2002). White, W. M.; Hofmann, A. W.; and Puchelt, H., Isotope geochemistry of Pacific mid-ocean ridge basalts. J. Geophys. Res. 92:4881–

4893 (1987). Wörner, G.; Moorbath, S.; Harmon, R.S. 1992b. Andean Cenozoic volcanic centers reflect basement isotopic domains. Geology, Vol.

20, p. 1103-1106.

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Estimating building and infrastructure vulnerability in the city of Arequipa, Peru, from volcanic mass flows: A challenge

Kim Martelli1, Jean-Claude Thouret

1, Cees van Westen

2, Denis Fabre

3, Michael Sheridan

4, &

Rubén Vargas2

1 Laboratoire Magmas et Volcans, Université Blaise Pascal, 5, rue Kessler, 63000 Clermont-Ferrand, France

(k.martelli@opgc.univ-bpclermont.fr) 2 International Institute for Geoinformation Science and Earth Observation, P.O. Box 6, 7500AA Enschede, The

Netherlands 3 Conservatoire National des Arts et Métiers (CNAM), 2 rue Conté, 75003 Paris, France

4 Department of Geology, Unviersity of Buffalo, SUNY, Buffalo, NY 14260, USA

KEYWORDS : Arequipa, El Misti, Peru, volcanic mass flow, vulnerability, hazard

Introduction

Rapid population growth and urban expansion has led to an increase in the vulnerability of communities living

within close proximity to an active volcano. Arequipa, the second largest city in Peru with a population

exceeding 860,000 is no exception, and is much like the city of Naples in Italy is exposed to Vesuvius. Arequipa

has experienced rapid population growth since the 1940s, and from 1970 onwards the urban area grew

substantially due to social unrest and related migration from rural areas, mainly in the form of poorly designed

suburbs and illegal settlements. Settlements have now expanded onto the southwest flank of the volcano, the R o

Chili River terraces and adjacent to tributaries within 9 km of El Misti summit. Studies of the type, extent, and

volume of Holocene pyroclastic and lahar deposits have concluded that future eruptions of El Misti, even if

moderate in magnitude, will pose a serious threat to Arequipa (Thouret et al., 1999; Delaite et al., 2005). Here

we discuss computer simulation of mass flows, classification of buildings and infrastructure and the challenges

we are faced with while assessing building and infrastructure vulnerability within Arequipa.

Geologic setting and volcanic mass flow hazards

El Misti is one of the seven active volcanoes within the Central Volcanic Andean Zone (CVZ) of southern

Peru. Arequipa is located 17 km SW of and 3 km below the summit of El Misti. The city is situated upon

volcaniclastic fans of pyroclastic-flow and lahar deposits from El Misti that are less than 10,000 years old.

Three possible hazard scenarios have been proposed for El Misti volcano (Thouret et al. 1999; Delaite et al.

2005). Scenario 1 is described as the most probable type of future activity with a VEI2 and a recurrence interval

of 300 to 1000 years; Scenario 2 is a moderate magnitude (VEI3) / frequency (1600 to 5000 years) eruption; and

Scenario 3 is the maximum expected pyroclastic eruption with a VEI>3 and a recurrence interval of 10,000 to

20,000 years. All eruption scenarios result in the formation of volcanic mass flows. These could include; dam-

break floods, pyroclastic flows, block-and-ash flows; lahars; and pumice flows, surges and high energy directed

blasts. In addition, lahars and flash floods can occur in the R o Chili River and Quebradas without an eruption

(occurring on average once every ten years from El Misti) from rainfall, snow meltwater, and a dam break flood.

The challenge of modelling

Stinton et al. (2004), Delaite et al. (2005) and Vargas et al. (2007) have attempted the delineation of lahar

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prone areas on El Misti flanks, ring plain, and in the city of Arequipa using LaharZ and Titan2D (single and two-

phase models). The “single” phase model is used to model dry flows, and the “two phase” model allows the

definition of a solid and a liquid fraction for the simulated flow. Lahars ranging from 0.01x106 m3 to 11x106 m3

in volume were simulated down the R o Chili valley, and the Quebradas San Lazaro and Huarangal. Solid

fractions of 0.3 to 0.5 were incorporated into the simulated flows.

The sensitivity of the DEM in Titan2D was analysed using changes in the simulation parameters and the

impact of flow features such as starting point, internal and bed friction angles, solid fraction, runout, super-

elevation, ponding,, flow divergence and convergence were examined. Previous simulations performed on a

30 m DEM, based upon digitising 1:25,000-scale topographic maps and on radar interferometry, compared

Titan2D with LaharZ and highlighted discrepancies between the two models. The largest flow volume simulated

by Titan2D (11.0 x 106 m3) did not reach further than the smallest volume of a flow (1.5 x 106 m3) modelled for

LaharZ (Delaite et al., 2005), refer to figure 1. Discrepancies may be explained by the differing models; LaharZ

is a statistically based method for delineating lahar-prone zones (Schilling, 1998), while Titan2D is a depth-

averaged, thin-layer Computational Fluid Dynamics program (Pitman et al., 2003).

To investigate the effect of the DEM on simulated results a 10 m DEM was computed using DGPS data, aerial

photographs and stereophotogrammetry. Detailed topographical data was acquired from a DGPS survey

undertaken on the four main terraces of the R o Chili River, an area of approximately 5km2, from the Military

Camp (approximately 15 km from the summit) downstream to the Bolognesi Bridge (city centre). Characteristics

such as overbanking, uphill flow, lateral spreading and flow divergence and convergence were identified in

Titan2D simulation outputs, and the results are more realistic than flow features identifies from LaharZ

simulations. At abrupt changes in channel direction, particularly where the R o Chili canyon opens out from a

steep sided gorge to a wide river valley (near the Chacani hydro-electric dam) the flows form temporary ponds

or even cease moving altogether. The DEM will be further refined with additional DGPS data to be collected this

year, and compared with previous DEMs. Lahar-prone areas and eruption scenarios will be now used to establish

the impact of mass flows on buildings and infrastructure within the city of Arequipa.

Figure 1. A– Map showing the difference between Titan2D runout and LaharZ runouts. B – example of flow divergence and convergence with Titan2D modelling on Qda. San Lazaro. This closely resembles reality as the flow moves around an obstacle in the channel. The dark black outline is the LaharZ simulated flow outline (Stinton et al., 2004).

A

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Assessing the physical vulnerability of buildings and infrastructure:

Classification of construction and landuse

Damage caused by volcanic mass flows has been observed during historic eruptions such as Vesuvius in

AD 79 and 1631, Nevado del Ruiz in 1985, and Soufrière Hills Volcano in 1997 (Spence et al., 2004). Damage

resulting from lahars and floods can include burial, foundation failure, debris impact with forces as high as 104-

106 kgm-2, transportation, excessive wall or roof loads, collapse, undermining and corrosion. A survey was

carried out on the characteristics of buildings and infrastructure that may be vulnerable to flow impact, with the

aim to define the probability of a building being in a particular damage state, given the intensity level of the

particular hazard concerned. The descriptive survey using a method adapted from Chevillot (2000) and Spence et

al. (2004) was conducted at street and city block level and where permitted, within the boundaries of the land-

owners property. Building types were defined according to the dominant building material; number of floors;

building reinforcement; roof type and style; opening type and quantity; and overall building structural integrity.

Figure 2. Top left – location map of the study area in Arequipa city centre in relation to El Misit. Bottom left – examples of two Types of buildings surveyed. Right – Classification map of the landuse and building construction classification in a section of the city centre.

Nineteen land-use patterns and ten construction types were identified (Figure 2). Most new construction

comprised un-reinforced masonry panels (perforated red brick and mortar) with cast-in-situ reinforced concrete

frames (horizontal and vertical), and flat or pitched reinforced concrete slab roofs. Large glass windows are

present throughout with aluminium or wood framing and often secured with steel bars. Doors are solid and

wooden with steel security screen/bars. Type A buildings represented 30% of those surveyed. Conversely, Type

I construction comprised old stone/ignimbrite base with unreinforced masonry panels (ignimbrite, brick or

adobe, with poor quality mortar). The walls were not confined by either reinforced horizontal or vertical cast-in-

situ concrete, and in most cases appear unstable. Wooden rafters support corrugated iron roofs which are secured

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by rocks and wood. These buildings represented 5% of those surveyed, and were more commonly situated on the

lower terraces of the Rio Chili and associated with agricultural lifestyle blocks. Less than 50% of the population

surveyed reside in dwellings less than Type C. Housing of poorer quality was often situated in the most

vulnerable areas upstream of the city and within the river channels (apart from type A housing located on the

confluence of the Río Chili and Qda. San Lázaro).

Bridges located within the Río Chili River are susceptible to debris accumulation due to their low height.

During the floods of February 1992 and 1997 debris accumulated behind the bridges, subsequently overtopping

and flooding nearby homes and businesses. If bridges are destroyed, access from one side of the city to the other

will be severely limited. Water pipes and power lines are often located at bridges, and are thus equally

susceptible to damage. Services such as the Egasa power station and the hydro-electric dams are vulnerable to

inundation from even small volume lahars, possibly resulting in the severe disruption of power supply to the

city, and having a flow-on effect for lifeline services such as hospitals and other emergency services.

Discussion

At El Misti challenges have arisen due to the considerably different results of two codes, Titan2D and LaharZ.

While Titan2D models the behaviour of debris flows adequately, the runout of LaharZ simulations is more

comparable to mapped deposits. Even on our enhanced DEM, these discrepancies arise; thus the need for

additional DEM refinement and research into the input parameters for geophysical flow modelling. Our results

highlight that simulations must only be used as a tool alongside geological mapping to aid the delineation of

inundation zones. The results of the building and infrastructure survey identified a range of construction types,

and often within the same city block. The poorest quality houses (and not structurally sound) are often located

closest to the river channels and in many cases could provide additional debris for the flow. Bridges, which link

two sides of the city, are vulnerable due to their low height and narrow spans, acting as a dam for flow debris.

Further modelling would aid the characterisation of building and infrastructure vulnerabilities by redefining

likely lahar prone zones, and therefore expected deposit thicknesses and flow velocities; all of which are of

importance when defining the likely damage states of buildings inundated by volcanic mass flows.

References Chevillot B. 2000. Rapport de mission a Arequipa (sud Pérou) 26 julliet – 15 août 2000. Objet: mise en œuvre d’un S.I.G.

appliqué aux risques hydrologiques et volcaniques. Rapport., Lab. De traitement de données géographiques ENITA de Clermont-Ferrand, 44 p.

Delaite, G., Thouret, J.-C., Sheridan, M. F., Stinton, A., Labazuy, P., Souriot, T., and van Westen, C., 2005. Assessment of volcanic hazards of El Misti and in the city of Arequipa, Peru, based on GIS and simulations, with emphasis on lahars: Zeitschrift für Geomorphology N.F., suppl.- vol. 140, p 209-231.

Pitman, E.B., Patra, A., Bauer, A., Nichita, C., Sheridan, M. and Bursik, M. 2003. Computing debris flows. Physics of Fluids 15: 3638-3646

Schilling, S.P., 1998. LAHARZ: GIS program for automated mapping of lahar inundation hazard zones: U.S. Geological Survey Open-File Report 98, 638. p.

Spence R.S.J., Baxter P.J., Zuccaro, G. 2004. The resistance of buildings to pyroclastic flows: analytical and experimental studies and their application to Vesuvius. Journal of Volcanology and Geothermal Research 133: 321-343.

Stinton, A., Delaite, G., Burkett, B., Sheridan, M., Thouret, J.-C., and Patra, A., 2004. Titan2D simulated debris flow hazards: Arequipa, Peru: Abstracts, International Symposium on Environmental Software Systems (ISESS), U.S.A.

Thouret, J.-C., Finizola, A., Fornari, M., Legueley-Padovani, A., Suni, J., Frechen, M. 2001. Geology of El Misti volcano near the city of Arequipa, Peru, Bull. Geol. Soc. Amer., 113: 1593-1610.

Vargas F.R, Thouret J.-C., Delaite G., Van Westen C., Sheridan M.F., Siebe C., Mariño J., Souriot T., and Stinton A. 2007. Mapping and assessing volcanic and flood hazards and risks, with emphasis on lahars, in the city of Arequipa, Peru. Geol. Soc. Amer. Spec. publ. (accepted).

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Metamorphic P-T constraints for the low-temperature assemblages overimposed on metamorphic and igneous rocks nearby Ñorquinco Lake, Aluminé, North-Patagonian Andes

C. I. Martínez-Dopico

Cátedra de Mineralogía, Departamento de Ciencias Geológicas, Universidad de Buenos Aires, Ciudad

Universitaria, C1428EHA, Buenos Aires, Argentina (carmen@gl.fcen.uba.ar)

KEYWORDS : subgreenschists facies, geothermobarometry, polymetamorphism, Aluminé, northern Patagonian Andes

Introduction

The area of the Ñorquinco-Pulmarí valley, 50 km west Aluminé city, North-Patagonian Cordillera (Fig. 1) is

distinguished by the presence of isolated outcrops of medium to high metamorphic grade rocks accredited by

amphibolites and gneises. These rocks compone the Upper Paleozoic Colohuincul igneous-metamorphic

Complex (Dalla Salda et al., 1991; Varela et al., 2005). This basament is intruded by the Paso de Icalma

Granodiorite (Cucchi et al., 2005; Latorre et al., 2001), a local Jurassic to Upper Cretacic igneous episode of a

mayor Jurassic to Miocene event known as North-Patagonian Batolith. These rocks are covered by a Tertiary to

Quaternary Andean thick andesitic to basaltic pile (Auca Pan Formation, Rancahue Basalt, Hueyeltué Basalt and

Lanín Basalt). Vattuone et al. (2005), among other authors, have characterized a low to very low grade

metamorphism in the Eastern Andean volcanic pile. Through the fieldwork, the petrographical study, EDAX on

amphibols, pumpellyite and zeolites crystals (Martínez Dopico, 2007; Gallegos, 2007), and the use of an

internally consistent thermodynamic data as the one proposed by Berman (1988, 2007), the mineral assemblages

and metamorphic facies are diagnosed to characterize and establish a sequence of paragenesis from the medium

to very low grade metamorphism overimposed on the andean southern basement and volcanic cover.

Figure 1: a) Geodynamic framework for Patagonia and location of the concern area within the North Patagonian Cordillera b) Studied localities in the Ñorquinco-Pulmary valley.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 319-321

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Metamorphic mineralogy and P-T constraints

Four assemblages are found for the amphibolites of the Colohuincul Complex. The higher grade ones in

amphibolite- greenschists facies are developed by the associations: opx + cpx + pl+ tsc (equilibrated at 550°C

and 4,7 kbar) and ed + act + ab + chl + ep. Subsequently, there are two overimposed re-equilibrated facies, the

first under greenschist P-T conditions set up as act + ep + chl ± ab (450°C and 2,3 kbar) and the latter, under a

prehnite-pumpellyite facies conditions defined by prh + pmp + chl ± ab + ep (300°C and 2,4 kbar). The proposed

protolith for these amphibolites has a mafic to ultramafic affinity.

Two paragenesis are recognised in the paragneises, the higher grade one in sillimanite facies is composed of kfs

+ sil + and+ crd + bt, its equilibrium point was found at 1,8 kbar and 630°C, the other association, steady at

temperatures below 250 °C, in the biotite zone is detected from chl + ms ± bt ± ab ± cb ± kln.

The tonalites compiled under the Granodiorita Paso de Icalma have a slight secondary overprint evidenced by

chl/ smectites + act + prh in quartz veins. These features indicate non-coaxial deformation operating under

prenhite-actinolite facies conditions of metamorphism, probably at temperatures between 300 and 200°C. This T

interval is compatible with the observed mineral assemblages in the other protoliths.

The metamorphic assemblage in metadacites of the paleocene Auca Pan Formation consist of ab+ ep + chl +

phl, formed under zeolite facies PT conditions (underneath 250°C), evidenced by the presence of phillipsite. In

the Upper Miocene Rancahue Basalt the secondary assemblage observed in amygdales consists of thomsonite-

Ca, faujasite-Ca and smectites, in the matrix phillipsite, scolecite and epidote were found (Gallegos, 2007).

These associations reach their equilibrium under 250°C. Slightly secondaries processes are represented in minor

veins and fractures by cb + act in the Middle Pleistocene Hueyeltué olivinic basalt.

Discussion

These data allow us to relate these metamorphic events to three historical pulses barothermically different. The

higher grade metamorphism (>550°C) in amphibolite- greenschists facies and its local reversions, are assigned to

an Upper-Paleozoic to Jurassic pulse, associated with the emplacement at different crustal levels of plutonic

episodes. South of the studied area, in San Martín de los Andes, this igneous activity could be represented by the

igneous fraction of the Colohuincul Complex and, in the concern area, by the intrusion of the Paso de Icalma

Granodiorite The lower grade event in prehnite-pumpellyite facies (300-350°C). could play as an overimposed

metamorphism linked to the early upper Cretaceous metamorphic ages associated with an extensional regime,

crustal attenuation and subsidence developed within the Andes as proposed by Aguirre et al.(1999). The very

low grade event (<250°C), in zeolite facies is able to be subdivided in two stages, Paleogene- Miocene

represented in the volcanic rocks of the Auca Pan Formation and Rancahue Basalt and an upper Miocene to

Pleistocene stage associated with the secondary formation of actinolite in the Hueyeltué Basalt and Granodiorite

Paso de Icalma. This last event in prehnite- actinolite facies was associated with the proximal Andean Miocene

granitoids and dated by Ar40/Ar39 in actinolite crystals in 8 Ma of the Rio Damas metabasites in the western

Andean margin by Oliveros et al., (2008). Regarding the tectonic enviroment of the Neuquén Andes it is

considered that these changes in the steady mineralogy are consistent with a polymetamorphic evolution

according to the variation in the angle of steepening of the Wadati- Benioff zone that generates the progressive

stages of compression and extension at this latitude of the Andes as discussed by Folguera et al. (2007).

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Acknowledgements This work was supported by the grants UBACyT X238. We would like to thank Dr. Leal and Dr Vattuone at University of Buenos Aires for their constructives reviews in the degree thesis “Geología y Geotermobarometría de las rocas metamórficas e ígneas de los lagos Ñorquinco y Nompehuén. Cordillera Patagónica Septentrional, Provincia del Neuquén”.

References

Aguirre, L., Feraud, G., Morata, D., Vergara, M. & Robinson, D. 1999. Time interval between volcanism and burial metamorphism and rate of basin subsidence in a Cretaceous Andean extensional setting. Tectonophysics, 313:433-447.

Berman, R.G., 1988. Internally consistent thermodynamic data for stoichiometric minerals in the system Na2O K2O CaO MgO FeO Fe2O3 Al2O3 SiO2 TiO2 H2O CO2. Journal of Petrology, 29:515-522.

Berman, R.G., 2007. WinTWQ (version 2.3): A software package for performing internally-consistent thermobarometric calculations. Geological Survey of Canada, Open File 5462.

Cucchi, R., Leanza, H.A., Repol, D., Escosteguy, I., González, R. y Danieli, J.C., 2005. Hoja geológica 3972-IV, Junín de los Andes. Provincia de Neuquén. Instituto de Geología y Recursos Minerales, Servicio Geológico Minero Argentino. Boletín 357, 102 pp., Buenos Aires.

Dalla Salda, L.H., Cingolani, C. y Varela, R., 1991. El Basamento Pre- andino ígneo- metamórfico de San Martín de los Andes, Neuquén. Revista de la Asociación Geológica Argentina, 46:223-234, Buenos Aires.

Folguera, A., Introcaso, A., Gimenéz, M., Ruiz, F., Martínez, P., Tunstall, C., García Morabito, E. & Ramos, V.A., 2007. Crustal attenuation in the Southern Andean retroarc (38°-39°30´S) determined from tectonic and gravimetric studies: The Lonco-Luán asthenospheric anomaly. Tectonophysics, 439: 129-147.

Gallegos, E., 2007. Geología del basamento del Valle del río Pulmarí. Trabajo Final de Licenciatura, Universidad de Buenos Aires, inédito, 121pp.

Latorre, C.O., Vattuone, M.E., Linares, E. & Leal, P.R. 2001. K-Ar ages of the rocks from the Lago Aluminé, Rucachoroi and Quillén, North Patagonian Andes, Neuquen, República Argentina. III° Simposio Sudamericano de Geología Isotópica: 577-580, Pucón.

Martínez Dopico, C.I., 2007. Geología y Geotermobarometría de las rocas metamórficas e ígneas de los lagos Ñorquinco y Nompehuén. Cordillera Patagónica Septentrional, Provincia del Neuquén. Trabajo Final de Licenciatura, Universidad de Buenos Aires, (inédito), 149 p., Buenos Aires.

Oliveros, V., Aguirre, L., Morata, D., Simonetti A., Vergara, M., Belmar, M. & Calderon, S., 2008. Geochronology of very low-grade Mesozoic Andean metabasites; an approach through the K-Ar, 40 Ar/ 39 Ar and U-Pb LA-MC-ICP-MS methods. Journal of the Geological Society of London, 165(2):579-584

Varela, R., Basei, M.A.S., Cingolani, C.A., Siga, O.Jr y Passarelli, C.R., 2005. El basamento cristalino de los Andes Norpatagónicos en Argentina: geocronología e interpretación tectónica. Revista Geológica de Chile, 32(2): 167-187.

Vattuone, M.E., Latorre, C.O. and Leal, P.R., 2005. Polimetamorfi smo de muy bajo a bajo grado en rocas volcánicas jurásico – cretácicas al sur de Cholila, Chubut, Patagonia Argentina. Revista Mexicana de Ciencias Geológicas, 22 (3):315-328.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 322-325

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Dynamic topography into the Amazonian basin: Insights from 3-D analogue modelling

J. Martinod1, N. Espurt

2, S. Brusset

1, F. Funiciello

3, C. Faccenna

3, & P. Baby

1

1 LMTG, Université de Toulouse, CNRS, IRD, OMP, 14 av. E. Belin, F-31400 Toulouse, France

(martinod@lmtg.obs-mip.fr) 2 IFP, 1 et 4 av. de Bois-Préau, Rueil-Malmaison cedex, France (nicolas.espurt@ifp.fr)

3 University Roma Tre, Dip. Scienze Geologiche, L.S. Leonardo Murialdo 1, 00146 Rome, Italy

KEYWORDS : Nazca ridge, dynamic topography, Amazonian basin, analogue model

Introduction and geodynamic setting

Subducting slab-induced mantle flow is a prominent control on the topographic signal within the overriding

lithosphere (Husson, 2006). The topographic signal can be also correlated with the dip of the subducting slab

(Mitrovica et al., 1989) which, in turn, controls the flexure of the overriding lithosphere (Catuneanu et al., 1997).

The subduction of buoyant aseismic ridge controls the dip of the slab and may generate flat slab subduction

(Gutscher et al., 2000). The Nazca Ridge is one of the major oceanic ridges subducting below South America.

This ridge has an average bathymetric relief of 1.5 km above the adjacent sea floor of the Nazca Plate, a

maximum width of 200 km at its base, and an average crustal thickness of 18±3 km (Woods and Okal, 1994)

(Fig. 1). The ridge migrates southward below the South American Plate because the ridge segment is N45°E

trending, oblique to the N78° present-day plate convergence (Gripp and Gordon, 2002). The subduction of this

ridge controls the morphology and tectonic of the forearc area since the Miocene (Macharé and Ortielb, 1992)

and constitutes the southern edge of the Peruvian flat slab segment (Gutscher et al., 1999).

Figure 1 : Geomorphic map of the northern South America with the Andean backbone and the subducting Nazca plate on the left and the Amazonian basin on the right (NASA SRTM Gtopo 30 data). The Fitzcarrald Arch constitutes a major relief which divides the western Amazonian foreland basin into two parts: the northern-Amazonian foreland basin and the southern-Amazonian foreland basin. To the east, the Arch is bounded by the eastern-Amazonian basin. Plate convergence vector is from Gripp and Gordon (2002). The projection of the Nazca Ridge beneath the South American plate is from Hampel (2002) and has been draped of the Nazca slab geometry (modified from Espurt et al., 2007).

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From geomorphical, sedimentological and geophysical data, the recent study of Espurt et al. (2007) shows that

the Nazca ridge flat slab also controls the present-day morphology of the Amazonian foreland basin, producing

the uplift of the Fitzcarrald Arch (Fig. 1). This Arch divides the western Amazonian basin into two main

subsiding retroforeland basins: the northern Amazonian retroforeland basin and the southern Amazonian

retroforeland basin (Roddaz et al., 2005). To the east, the Fitzcarrald Arch uplift is bounded by the anomalous

subsiding eastern-Amazonian basin. A simple examination of the morphologic map of the Amazonian basin and

Nazca ridge flat slab eastwards extension reveals that the Fitzcarrald arch uplift maintains at more 400 km

eastwards of the present-day flat slab segment whereas the slab plunges vertically into the asthenosphere (Fig. 1).

Manifestly, the Amazonian basin shows dynamic topography evidence probably related to the western

subduction process. Using lithospheric scale analogue experiments, the paper aims to explore the effects of the

subduction of an oblique buoyant ridge on (1) the mantle flow process and (2) the dynamic topography evolution

into the Amazonian basin.

Model set-up

The experimental setting adopted here is close to the one used in Funiciello et al. (2004). We use Newtonian

viscous materials within a Plexiglas tank to reproduce the subduction of lithospheric plates within the upper

mantle. Lithospheric plates are modelled, using high-viscosity silicone putty. We vary the density of silicone

putty to take into account the different lithosphere buoyancies. The upper mantle is modelled, using a Newtonian

low-viscosity glucose syrup solution. Plate convergence is modelled, using a piston advancing at constant

velocity. The evolution of the overriding plate topography is monitored using a three dimensional laser

stereoscopic technique (Real Scan USB model 300) during experiments and digital elevation models have been

performed.

Results

Here are the results of two experiments: experiment 1 (Fig. 2) models the subduction of a negatively buoyant

oceanic plate with an oblique ridge The relative buoyancy of this ridge is positive and it may constitute a good

analogue of the Nazca ridge. Experiment 2 (Fig. 3) is similar to the previous experiment but a pushed continental

plate is placed above the subduction zone.

In experiment 1, steady-state subduction is essentially governed by trench retreat. When the tip of the ridge

reaches the trench, the velocity of subduction decreases in front of the ridge. In contrast, in the lower part of the

plate, we do not measure any significant change in the velocity of subduction where the subduction process is

essentially governed by the negative buoyancy of the dense oceanic plate. When a large length of ridge has been

subducted, the dense subducted slab located above the ridge pulls again the lithospheric plate toward the

subduction zone. Subduction velocity increases again. In contrast, the oblique ridge locked the subduction in the

lower part of the plate. The partitioning of the subduction velocity between the ridge and the rest of the plate

produces an arched shape of the trench. Experiment 1 shows that the oblique buoyant ridge controls (1) the

geometry of the slab and (2) the kinematic of the subduction.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 322-325

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Figure 2 : (a): top views of experiment 1 simulating the subduction of a dense oceanic plate with an oblique buoyant ridge without upper plate; (b): velocity of subduction and trench retreat vs. time during experiment along three cross sections. The oblique geometry of the ridge generated a partitioning of the subduction process. Dynamic subsidence is correlated with the decrease of the subduction velocity. In contrast, dynamic uplift is observed when the velocity of the subduction decreases, i.e., above the ridge.

Figure 3 : Digital elevation models of experiment 2 during run. The migration of the oblique ridge below the advancing plate is associated with vertical topographic motion within the upper plate (uplift above the ridge and subsidence above the ocean). The topographic evolution of the upper plate is partly related to the flow in the syrup glucose induced by the slab retreat.

In experiment 2 (Fig. 3), the trench retreat is essentially controlled by the advance of the upper plate. The

subduction of the ridge below the upper plate results in a horizontal subduction only if a large amount of buoyant

segment is forced into subduction as described by Espurt et al. (2008). Part of the buoyant ridge is incorporated

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 322-325

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in the steep part of the slab to balance the negative buoyancy of the dense oceanic lithosphere. The analysis of

the upper plate topography during the experiment from the digital elevation models shows that (1) the sweeping

of the oblique ridge induces vertical motions on the chain and in the backarc basin. It results an uplift above the

flat slab segment, followed by accelerated subsidence, inducing an asymmetrical shape of the backarc basin with

two pronounced and asymmetrical subsiding basins.

Discussion: dynamic topography in the Amazonian basin

The Amazonian basin is an unconventional foreland basin system because in this central part the flexure is

overcompensated by the Nazca ridge flat slab segment (Espurt et al., 2007). However, the only flat slab segment

cannot explain the observed wide Fitzcarrald Arch uplift which develops more than 400 000 km2. Analogue

experiments show that the subduction of a buoyant segment may perturb the mantle flow and, in turn, may

control the dynamic topographic signal in the continental lithosphere (Mitrovica et al., 1989; Pysklywec and

Mitrovica, 2000). The dynamic subsidence can be correlated with the high subduction velocity on the both sides

of the ridge. In contrast, the ridge subduction cancels the vertical component of the slab velocity and dynamic

uplift is observed. Consequently, the subduction of the Nazca ridge may cancel the subduction process below the

western Amazonian basin and the broad Fitzcarrald Arch uplift could be dynamically maintained by the stop of

the mantle flow eastwards of the present-day flat slab segment.

References Catuneanu, O., Beaumont, C., & Waschbusch, P. 1997. Interplay of static loads and subduction dynamics in foreland basins:

Reciprocal stratigraphies and « missing » peripheral bulge. Geology 25: 1087-1090. Espurt, N., Funiciello, F., Martinod, J., Guillaume, B., Regard, V., Faccenna, C., & Brusset, S. 2008. Flat subduction

dynamics and deformation of the South American plate. Insights from analogue modelling. Tectonics doi:10.1029/2007TC002175, in press.

Espurt, N., Baby, P., Brusset, S., Roddaz, M., Hermoza, W., Regard, V., Antoine, P.-O., Salas-Gismondi, R., & Bolaños, R. 2007. How does the Nazca Ridge subduction influence the modern Amazonian foreland basin? Geology 35: 515-518.

Funiciello, F., Faccenna, C., & Giardini, D. 2004. Role of lateral mantle flow in the evolution of subduction systems: Insights from laboratory experiments. Geophys. J. Int. 157: 1393-1406.

Gripp, A.E., & Gordon, R.G. 2002. Young tracks of hotspots and current plate velocities. Geophys. J. Int. 150: 321-361. Gutscher, M.A., Olivet, J.L., Aslanian, D., Eissen, J.P., & Maury, R. 1999. The “lost Inca Plateau”: Cause of flat subduction

beneath Peru? Earth and Planetary Science Letters 171: 335-341. Hampel, A. 2002. The migration history of the Nazca Ridge along the Peruvian active margin: a re-evaluation. Earth and

Planetary Science Letters 203: 665–679. Husson, L. 2006. Dynamic topography above retreating subduction zones. Geology 34: 741-744. Macharé, J., & Ortlieb, L. 1992. Plio-Quaternary vertical motions and the subduction of the Nazca Ridge, central coast of

Peru. Tectonophysics 205: 97-108. Mitrovica, J.X., Beaumont, C., & Jarvis, G.T. 1989. Tilting of continental interiors by the dynamical effects of subduction.

Tectonics 8: 1078-1094. Pysklywec, R.N., & Mitrovica, J.X. 2000. Mantle flow mechanisms of epeirogeny and their possible role in the evolution of

the Western Canada Sedimentary Basin. Canadian Journal of Earth Sciences 37: 1535-1548. Roddaz, M., Viers, J., Brusset, S., Baby, P., Brusset, S., & Hérail, G. 2005. Sediment provenances and drainage evolution of

the Neogene Amazonian foreland basin. Earth and Planetary Science Letters 239: 57–78. Woods, M.T., & Okal, E.A. 1994. The structure of the Nazca Ridge and Sala y Gomez seamount chain from dispersion of

Rayleigh waves Geophys. J. Int. 117: 205–222.

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Tectonic control on the 1960 Chile earthquake rupture segment

Daniel Melnick1, Marcos Moreno

2, Dietrich Lange

1,3, Manfred R. Strecker

1, & Helmut P.

Echtler2

1 Institut für Geowissenschaften, Universität Potsdam, 14415 Potsdam, Germany (melnick@geo.uni-potsdam.de)

2 GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany

3 Present address: Bullard Laboratory, University Cambridge, UK

KEYWORDS : subduction earthquakes, seismotectonic segmentation, forearc microplates, crustal faults

Introduction

Understanding the principles that govern triggering of great subduction earthquakes and finite rupture length

and consequently magnitude is an utmost important challenge (e.g., Satake and Atwater, 2007). In principle, two

major conditions are required to generate a giant-magnitude subduction earthquake (M>8.5): (1) the forearc of

the upper plate has to accumulate enough elastic strain to burst rupture and fault slip, and (2) rupture has to

propagate unstalled for hundreds of kilometers (e.g., Kanamori, 1977). The great 1960 Chile earthquake (Mw of

9.5) was such a megathrust event that ruptured ~1000 km of the Nazca-South America plate boundary, involving

up to 40 m of fault slip resulting in up to 5.7 m of vertical coastal uplift (Barrientos and Ward, 1990, Plafker and

Savage, 1970). This event was the largest earthquake instrumentally recorded by modern seismology (Engdahl

and Villaseñor, 2002). The 1960 earthquake started at 38.2°S (Engdahl and Villaseñor, 2002) and propagated

southward until the Nazca-Antarctic-South America Triple Plate Junction at 46°S (Barrientos and Ward, 1990,

Plafker and Savage, 1970). An historical earthquake of similar magnitude (M~9.5) and rupture length (38-46°S)

occurred in 1575, and paleoseismic records document similar events with a recurrence of ~300 years (Cisternas

et al., 2005). Here we present geologic, geodetic, and seismologic data to address the tectonic processes that

control strain accumulation and rupture propagation for the 1960 earthquake segment.

Regional tectonic setting & glacial-age sedimentation in the South Chile trench

The Chile margin is formed by oblique subduction of the Nazca plate below the South American continent at

~80 mm/a. Between the Juan Fernández Ridge (33°S) and the Chile Triple Junction, the trench has been filled

with over 2 km of sediments eroded from the high Andes by Patagonian glaciations since Pliocene time leading

to an accretionary margin (Bangs and Cande, 1997), whereas north of 33°S the trench is virtually depleted of

sediments and the margin has been erosive over the entire Cenozoic (e.g., Clift and Vannucchi, 2004).

Subduction of a coherent sedimentary sequence deposited in the trench smoothes the seismic strength of the

plate interface allowing larger earthquake-rupture propagation (Ruff, 1989), and explains differences in the long-

term evolution of the central and southern Andes (Lamb and Davis, 2003).

Decoupling of the Chiloé microplate by the Liquiñe-Ofqui fault zone

The 1960 rupture segment (38-45°S) is coincident with the extent of the Liquiñe-Ofqui fault zone (LOFZ), a

major strike-slip system that straddles the volcanic arc accommodating oblique plate convergence (e.g., Lavenu

and Cembrano, 1999, Rosenau et al., 2006). The LOFZ decouples a forearc sliver the Chiloé block from

stable South America as evidenced by Pliocene-Recent fault kinematics and paleomagnetic rotation patterns

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 326-329

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(Rosenau et al., 2006), clusters of shallow seismicity and focal mechanisms (NEIC, 2007, Lange et al., 2007,

Lange et al., 2008), and space geodetic data (Wang et al., 2007, Moreno et al., 2008). In April 2007, the LOFZ

generated a shallow Mw 6.2 earthquake near Aysén with a dextral focal mechanism (NEIC, 2007).

Figure 1. Plate-tectonic setting of the south-central Andes. Black lines denote major Quaternary faults (Melnick and Echtler, 2006), triangles active volcanoes. Trench-fill thicknesses from Bangs and Cande (1997). Historical earthquake ruptures from Lomnitz (2004) and Cisternas et al. (2005). Stars denote epicenters. Note the spatial correlation between the Chiloé microplate and the 1960 earthquake segment.

Collision of the Chiloé microplate, uplift of the Arauco Peninsula and

segmentation of the 1960 rupture segment

The northern boundary of the Chiloé sliver is the Arauco Peninsula, a major anomaly of South America’s

Pacific shore. At Arauco the entire Andean margin including the volcanic arc bends ~10° eastward. The bending

axis is coincident with the Nahuelbuta Mountains, an abnormally-high segment of the Coastal Ranges, and with

a major transition in fault kinematics and structural styles along the forearc, intra-arc, and foreland regions

(Melnick et al., 2006), which defines the Arauco Orocline (Melnick et al., 2008). The major structure of the

Arauco-Nahuelbuta region is the Lanalhue fault, which includes a Permian milonitic shear zone (Hervé, 1988,

Glodny et al., 2006), subsequently reactivated cutting Pliocene-Quaternary deposits. Clusters of crustal

seismicity have been registered below the surface expression of the Lanalhue fault (Haberland et al., 2006).

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Focal mechanisms and distribution of hypocenters are consistent with a steeply-dipping fault, as imaged by a

deep seismic-reflection profile (Groß et al., 2008), and expected from linear fault traces (Melnick and Echtler,

2006). GPS data show shortening across the Lanalhue fault and counterclockwise rotation of its western, coastal

block, consistent with collision of the Chiloé microplate (Moreno et al., 2008). The 1960 earthquake sequence

started the 21 of May with an Mw 8.2 foreshock followed 25 hours later by the Mw 9.5 mainshock and during the

next months and years by several Mw 6 to 7.9 aftershocks (Engdahl and Villaseñor, 2002). All these events

occurred adjacent to the Lanalhue fault zone, independent of relative relocation uncertainties (Engdahl and

Villaseñor, 2002). The cumulative seismic moment released here during the 60s and 70s suggests that

abnormally-high magnitudes of elastic strain had accumulated in this region prior to 1960.

Conclusions

We find that a combination of microplate behavior of the forearc and a ~2-km-thick pile of trench sediments

entering the subduction zone provides mechanical homogeneity to the upper plate and plate interface,

respectively, smoothing the seismic strength and allowing ruptures to propagate over 1000 km. Upper plate

contraction in the 1960 earthquake segment is enhanced at the northern sector of its rupture segment by collision

of the Chiloé microplate against the Arauco-Nahuelbuta buttress. The northern, leading edge of the Chiloé

microplate is marked by the Lanalhue fault. Enhanced strain in this region might have facilitated nucleation of

the 1960 seismic sequence, which included several high-magnitude (M 7-9.5) events that nucleated in a

relatively small, localized area. Collision of the Chiloé microplate has been ongoing over the past ~6 m.y.,

resulting in uplift of the Arauco Peninsula and Nahuelbuta Coastal ranges as well as bending of the entire

Andean orogen across the Arauco Orocline. This spatial correlation between tectonic segmentation at the million

year time scale and the transient, but recurring earthquake ruptures suggests that the 1960 earthquake segment

might be stable in space and time. However, the processes that control the temporal recurrence of high-

magnitude events that achieve rupture of the entire segment (Cisternas et al., 2005) still remains to be fully

understood (Satake and Atwater, 2007).

References Bangs, N.L. & Cande, S.C., 1997. Episodic development of a convergent margin inferred from structures and processes

along the southern Chile margin, Tectonics, 16: 489-503. Barrientos, S.E. & Ward, S.N., 1990. The 1960 Chile earthquake: inversion for slip distribution from surface deformation,

Geophysical Journal International, 103: 589-598. Cisternas, M., Atwater, B.F., Torrejón, F., Sawai, Y., Machuca, G., Lagos, M., Eipert, A., Youlton, C., Salgado, I., Kamataki,

T., Shishikura, M., Rajendran, C.P., Malik, J.K., Rizal, Y. & Husni, M., 2005. Predecessors of the giant 1960 Chile earthquake, Nature, 437: 404-407.

Clift, P. & Vannucchi, P., 2004. Controls on tectonic accretion versus erosion in subduction zones: Implications for the origin and recycling of the continental crust, Reviews of Geophysics, 42: RG2001.

Engdahl, E.R. & Villaseñor, A., 2002. Global Seismicity: 1900–1999. in International Handbook of Earthquake and Engineering Seismology, pp. 665-690, eds. Lee, W. H., Kanamori, H., Jennings, P. C. & Kisslinger, C. Academic Press.

Glodny, J., Echtler, H., Figueroa, O., Franz, G., Gräfe, K., Kemnitz, H., Kramer, W., Krawczyk, C., Lohrmann, J., Lucassen, F., Melnick, D., Rosenau, M. & Seifert, W., 2006. Long-term geological evolution and mass-flow balance of the South-Central Andes. in The Andes - Active Subduction Orogeny, pp. 401-428, eds. Oncken, O., Chong, G., Franz, G., Giese, P., Götze, H.-J., Ramos, V. A., Strecker, M. & Wigger, P. Springer-Verlag, Berlin, Heidelberg, New York.

Groß, K., Micksch, U., Araneda, M., Bataille, K., Bribach, J., Buske, S., Krawczyk, C.M., Lüth, S., Mechie, J., Schulze, A., Shapiro, S.A., Stiller, M., Wigger, P. & Ziegenhagen, T., 2008. The reflection seismic survey of project TIPTEQ-the inventory of the Chilean subduction zone at 38.2° S, Geophysical Journal International, 172: 565-571.

Haberland, C., Rietbrock, A., Lange, D., Bataille, K. & Hofmann, S., 2006. Interaction between forearc and oceanic plate at the south-central Chilean margin as seen in local seismic data, Geophysical Research Letters, 33: L23302.

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329

Hervé, F., 1988. Late Paleozoic subduction and accretion in southern Chile, Episodes, 11: 183-188. Kanamori, H., 1977. The energy release in great earthquakes, Journal of Geophysical Research, 82: 2981-2987. Lamb, S. & Davis, P., 2003. Cenozoic climate change as a possible cause for the rise of the Andes, Nature, 425: 792-797. Lange, D., Cembrano, J., Rietbrock, A., Haberland, C., Dahm, T. & Bataille, K., 2008. Seismic activity of the intra-arc

Liquiñe-Ofqui shear zone constraining current strain partitioning in southern Chile, Tectonophysics, under review. Lange, D., Rietbrock, A., Haberland, C., Bataille, K., Dahm, T., Tilmann, F. & Flueh, E.R., 2007. Seismicity and geometry

of the south Chilean subduction zone (41.5°S-43.5°S): Implications for controlling parameters, Geophysical Research Letters, 34.

Lavenu, A. & Cembrano, J., 1999. Compressional- and transpressional-stress pattern for Pliocene and Quaternary brittle deformation in fore arc and intra-arc zones (Andes of Central and Southern Chile), Journal of Structural Geology, 21: 1669-1691.

Melnick, D., Bookhagen, B., Strecker, M. & Echtler, H., 2008. Segmentation of subduction earthquakes from forearc deformation patterns over hundreds to millions of years, Arauco Peninsula, Chile, Journal of Geophysical Research, in review.

Melnick, D., Charlet, F., Echtler, H.P. & De Batist, M., 2006. Incipient axial collapse of the Main Cordillera and strain partitioning gradient between the Central and Patagonian Andes, Lago Laja, Chile, Tectonics, 25: TC5004.

Melnick, D. & Echtler, H.P., 2006. Morphotectonic and geologic digital map compilations of the south-central Andes (36°-42°S). in The Andes - Active Subduction Orogeny, pp. 565-568, eds. Oncken, O., Chong, G., Franz, G., Giese, P., Götze, H.-J., Ramos, V. A., Strecker, M. & Wigger, P. Springer-Verlag, Berlin Heidelberg New York.

Moreno, M.S., Klotz, J., Melnick, D., Grund, V., Echtler, H. & Bataille, K., 2008. Active faulting and forerarc block rotation in the south Chile from GPS-derived deformation, Geochemistry, Geophysics, Geosystems, under review.

NEIC, 2007. National Earthquake Information Center. www.neic.usgs.gov Plafker, G. & Savage, J.C., 1970. Mechanism of the Chilean earthquake of May 21 and 22, 1960, Geological Society of

America Bulletin, 81: 1001-1030. Rosenau, M., Melnick, D. & Echtler, H., 2006. Kinematic constraints on intra-arc shear and strain partitioning in the

Southern Andes between 38°S and 42°S latitude, Tectonics, 25: TC4013. Ruff, L.J., 1989. Do trench sediments affect great earthquake occurrence in subduction zones?, Pure and Applied

Geophysics, 129: 263-282. Satake, K. & Atwater, B.F., 2007. Long-term perspectives on giant earthquakes and tsunamis at subduction zones, Annual

Review of Earth and Planetary Sciences, 35: 349-374. Wang, K., Hu, Y., Bevis, M., Kendrick, E., Smalley Jr., R., Barriga-Vargas, R. & Lauría, E., 2007. Crustal motion in the

zone of the 1960 Chile earthquake: Detangling earthquake-cycle deformation and forearc-sliver translation, Geochemistry, Geophysics, Geosystems, 8: Q10010.

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Late Jurassic extensional tectonics in the southwestern Mendoza province, Argentina

José F. Mescua, Laura Giambiagi, & Florencia Bechis

Instituto Argentino de Nivología, Glaciología y Ciencias Ambientales (IANIGLA), CCT-CONICET, Parque San

Martín s/n, 5500 Mendoza, Argentina (jmescua@lab.cricyt.edu.ar, lgiambia@lab.cricyt.edu.ar,

fbechis@lab.cricyt.edu.ar)

KEYWORDS : Kimmeridgian, extension, Neuquén basin, reactivation, Cordillera Principal

Introduction

We present a study of the Kimmeridgian tectonic setting of the Andean region of south-western Mendoza

province, Argentina. The study area is located between 34º and 35º S (fig. 1). At these latitudes the Mesozoic

Neuquen embayment, developed south of 37º S, turns into a narrow longitudinal basin (90 km wide) that

continues to the north into the San Juan province. The Kimmeridgian fill of the Neuquen basin corresponds to

the continental deposits of the Tordillo Formation. This unit is one of the most conspicuous of the basin fill,

consisting of sediments deposited in different fluvial systems with associated playa lakes and eolian fields

(Legarreta and Uliana, 1999). Towards the west, these sedimentary deposits interfinger with ocoitic lavas and

andesitic breccias of the Río Damas Formation. This unit, developed mainly in Chilean territory, corresponds to

Kimmeridgian retroarc volcanic activity in an extensional basin (Charrier, 2007).

Different methodologies were used to study the Kimmeridgian tectonic setting in the study area. A

reconstruction of the thickness variations of the Tordillo Formation was obtained from outcrop and subsurface

data. A provenance analisys is being carried out on sandstones of this unit, the first results of which are

presented here. These lines of evidence suggested that the deposition of the sediments of the Tordillo Formation

was contemporaneous with extensional tectonics. The recognition of syn-sedimentary normal faults in some

localities supports this interpretation

Data analysis and interpretation

A thickness reconstruction of the Tordillo Formation is presented in figure 1. Although in some places the

isopachic contours had to be interpreted, given the sparse outcrops of the Tordillo Formation in the High Andes,

abrupt thickness changes are observed. A thickness increase from 150-300 m to 700 m is found superposed to

structures recognized as the main normal faults of the Lower Jurassic extensional event in the Río Atuel area (La

Manga fault, Giambiagi et al., 2005). In this area, the WNW-trending Río Atuel lineament seems to have

controlled the deposition of the Tordillo Formation as well. Thickness increase is even larger west of the

“Tordillo basement high” in the Las Leñas/Valle Hermoso area, where the Kimmeridgian deposits reach more

than 5000 m. This high element was already recognized by Legarreta and Gulisano (1989) as a tectonically

active zone in Kimmerdigian times in which the basement uplift controlled the facies and geometry of

sedimentary deposits. In the eastern piedmont, drilling data of YPF oil company boreholes show the presence of

isolated depocenters filled with sediments of the Tordillo Formation. These depocenters have the same

orientation of the Lower Jurassic structures and thicknesses up to 110 m. These data suggest that reactivation of

the Lower Jurassic normal faults could have taken place during the Kimmeridgian, controlling the deposition of

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 330-333

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the Tordillo Formation. This reactivation is also suggested by Charrier (2007) for the contemporaneous Río

Damas Formation westward of the study area.

Figure 1. Study area and localities mentioned in the text, with outcrops of the Tordillo and Río Damas Formations, partial isopaquic map for these units and inferred major extensional faults which controlled its deposition.

Furthermore, syn-sedimentary normal faults were observed in some of the studied localities of the Tordillo

Formation. One of these localities is Arroyo Colorado, where normal faults affecting the Tordillo fluvial facies

are associated with a fanning geometry in the strata. In another locality, Arroyo Pincheira (35º30’S), south of the

study area, normal faults with displacements of tens of meters lose slip upward within the Tordillo Formation

and are unconformably covered by beds of the same unit. In other areas such as Paso de las Damas, small-scale

normal faults with a NNE to ENE trend (Az=20º to Az=65º, n=6) are observed, with displacements in the order

of tens of centimeters to one meter and related thickness variations in the beds. In this locality, ocoitic lavas are

interbedded with red sandstones and shales. Ocoites are porphyritic basaltic andesites, which in this locality

present large tabular phenocrists of plagioclase of up to 5 cm. They could be related to an extensional setting, as

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proposed for the ocoites of the Lower Cretaceous in the Coastal Range of Chile (Morata and Aguirre, 2003). No

geochemical analyses of the Kimmeridgian volcanics are available to test this hypothesis.

The first results of a provenance study undertaken in sandstones and conglomerates of the Tordillo Formation

have revealed that rhyolites and granites of the Permian-Triassic Choiyoi Group are the main components of the

deposits in the localities of Río Borbollón, Cerro Amarillo, and Arroyo La Manga, showing that a large amount

of sediments from this source was fed to the basin in Kimmeridgian times. This implies the existence of large

exposure areas of the basement to the east of the basin, with a high relief margin to account for large clast sizes

in the conglomerates of the eastern outcrops of the Tordillo Formation. On the other hand, in exposures in the

western sector of the basin, in the localities of Paso de las Damas and Río del Cobre, where the Tordillo

Formation is interfingered with the andesites of the Río Damas Formation, conglomerates are composed almost

exclusively of andesite clasts from this last unit. Therefore, both margins of the basin provided sediments from

high relief areas during the Kimmeridgian. Taking into account the extensional tectonic setting described in

northern Mendoza and in Chile (Sanguinetti and Ramos, 1993; Cegarra and Ramos, 1996; Giambiagi et al.,

2003; Vergara et al., 1995; Charrier, 2007; see below), we interpret this fact as further proof of extensional

activity during this period.

Discussion and concluding remarks

Several studies have addressed the control of pre-existing structures in the Andean deformation in south

western Mendoza (e.g. Manceda and Figueroa, 1995; Giambiagi et al., 2003). They have shown that knowledge

of the pre-Cenozoic geologic history is needed to understand the processes involved in the formation of the

Andes.

Based on the existence of an unconformity between the Oxfordian Auquilco Formation and the Kimmeridgian

Tordillo Formation in some localities, early studies of the Neuquen Basin suggested an orogenic phase

(Stipanicic and Rodrigo, 1970; Davidson and Vicente, 1973). Later studies in northern Mendoza (Sanguinetti

and Ramos, 1993; Cegarra and Ramos, 1996; Giambiagi et al., 2003) proposed an extensional setting for this

period, whereas in Chile regional extension is well documented (Vergara et al., 1995; Charrier, 2007).

Nevertheless, it is generally assumed that extension in southern Mendoza was over by the late Early Jurassic

(Legarreta and Gulisano, 1989; Gulisano and Gutiérrez Pleimling, 1995; Vergani et al., 1995; Ramos, 1999;

Legarreta and Uliana, 1999).

As shown above, observations in southern Mendoza support the existence of an extensional tectonic setting for

the Kimmeridgian in this region. The region affected by Kimmeridgian extensional processes in Argentina

would extend, at least, from 32º30’ S (Aconcagua region) to 36º S. The unconformities observed between

Oxfordian and Kimmeridgian deposits in some localities of southwest Mendoza (e.g. Río Salado, Dajczgewand,

2002) could be related to tilting and erosion associated with this extension.

Vergani et al. (1995) interpreted that extensional fault-controlled subsidence was restricted to the Late

Triassic- Early Jurassic in the Neuquén Basin. Based on data collected south of 36ºS, they suggest that the high

rates of subsidence in Late Jurassic and Early Cretaceous times are due to relaxation of in-plane stresses with a

NW-directed 3. On the other hand, in south-western Mendoza, activity of normal faults would have taken place

at least during the Kimmeridgian. Available data from limited measurements of minor faults coeval with the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 330-333

333

deposition of the Tordillo Formation suggest that this extension had a NW direction. Major Early Jurassic NNW-

trending normal faults are interpreted to be reactivated with an oblique normal sense.

References Cegarra, M. I., Ramos, V. A., 1996. “La faja plegada y corrida del Aconcagua”. In V. A. Ramos (ed.), Geología de la región

del Aconcagua, provincias de San Juan y Mendoza. Dirección Nacional del Servicio Geológico, Subsecretaría de Minería de la Nación, Anales 24: 387-422, Buenos Aires.

Charrier, R., 2007. “Kimmeridgian backarc extensional reactivation and magmatism in the Northern and Central Chilean Andes (21º-36º LS)”. Geosur 2007, Libro de Resúmenes: 32, Santiago de Chile.

Dajczgewand, D. M., 2002. Faja Plegada y corrida de Malargüe: estilo de deformación en la región de Mallín Largo, departamento de Malargüe, provincia de Mendoza. Trabajo Final de Licenciatura, Universidad de Buenos Aires.

Davidson, J., Vicente, J. C., 1973. “Características paleogeográficas y estructurales del área fronteriza de las nacientes del Teno (Chile) y Santa Elena (Argentina), Cordillera Principal (35º a 35º15’ de Latitud Sur)”. V Congreso Geológico Argentino, Actas 5: 11-55, Buenos Aires.

Giambiagi, L. B., Alvarez, P., Godoy, E., Ramos, V. A., 2003. The control of pre-existing extensional structures on the evolution of the southern sector of the Aconcagua fold and thrust belt, southern Andes. Tectonophysics, 369: 1-19.

Giambiagi, L. B., Bechis, F., Lanés, S., García, V., 2005. “Evolución cinemática del depocentro Atuel, Triásico tardío- Jurásico temprano”. XVI Congreso Geológico Argentino, Actas CD-ROM, Artículo n° 169, 6 pp., La Plata.

Gulisano, C. A., Gutiérrez-Pleimling, A.R., 1995. Guía de campo: El Jurásico de la Cuenca Neuquina, b) Provincia de Mendoza. Asociación Geológica Argentina, Serie E nº 3, Buenos Aires.

Legarreta, L., Gulisano, C. A., 1989. “Análisis estratigráfico secuencial de la Cuenca Neuquina” (Triásico superior-Terciario). In Chebli, G. y Spalletti, L. (eds.), Cuencas Sedimentarias Argentinas. Facultad de Ciencias Naturales, Universidad Nacional de Tucumán, Correlación Geológica Serie 6: 221-243, Tucumán.

Legarreta, L., Uliana, M. A., 1999. “El Jurásico y Cretácico de la Cordillera Principal y la Cuenca Neuquina”. In Caminos, R. (Ed.), Geología Argentina, Servicio Geológico Minero Argentino, Anales 29: 399-416, Buenos Aires.

Manceda, R., Figueroa, D., 1995. “Inversion of the Mesozoic Neuquén rift in the Malargüe fold and thrust belt, Mendoza, Argentina.” In A. J. Tankard, R. Suárez, H. J. Welsink, Petroleum basins of South America. AAPG Memoir 62: 369-382.

Morata, D., Aguirre, L., 2003. Extensional Lower Cretaceous volcanism in the Coastal Range (29º20’- 30ºS), Chile: geochemistry and petrogenesis. Journal of South American Earth Sciences, 16: 459-476.

Ramos, V. A., 1999. “Evolución tectónica de la Argentina”. In Caminos, R. (Ed.), Geología Argentina, Servicio Geológico Minero Argentino, Anales 29: 715-759, Buenos Aires.

Sanguinetti, A. S., Ramos, V. A., 1993. “El volcanismo de arco mesozoico”. In V. A. Ramos (ed.), Geología y Recursos Naturales de Mendoza. XII Congreso Geológico Argentino y II Congreso de Exploración de Hidrocarburos, Relatorio: 115-122, Buenos Aires.

Stipanicic, P. N., Rodrigo, F., 1970. “El diastrofismo jurásico en Argentina y Chile”. IV Jornadas Geológicas Argentinas, Actas 2: 353-368, Buenos Aires.

Vergani, G. D., Tankard, A. J., Bellotti, H. J., Welsink, H. J., 1995. “Tectonic evolution and paleogeography of the Neuquén Basin, Argentina”. In A. J. Tankard, R. Suárez, H. J. Welsink, Petroleum basins of South America. AAPG Memoir 62: 383-402.

Vergara, M., Levi B., Nyström, J. O., Cancino, A., 1995. Jurassic and Early Cretaceous island arc volcanism, extension, and subsidence in the Coast Range of central Chile. GSA Bulletin, 107 (12), 1427–1440.

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Moment-tensor inversion of explosion events recorded on the Ubinas volcano, Peru

J.-P. Métaxian1, V. Monteiller

1, O. Macedo

2, G. S. O’Brien

3, E. Taipe

2, & C. J. Bean

3

1 LGIT, Université de Savoie, 73376 Le Bourget du lac cedex, France (metaxian@univ-savoie.fr)

2 Instituto Geofisico del Peru, Arequipa, Peru (omacedo@geo.igp.gob.pe)

3 UCD School of Geological Sciences, University College Dublin, Ireland (gareth.obrien@ucd.ie)

KEYWORDS : volcano, Ubinas, seismology, broadband, moment-tensor

Volcanic settings

The Ubinas (Peru) stratovolcano (5672 m), located 60 km east from Arequipa city, forms part of the range

resulting from the subduction of the Nazca plate under the South American plate. It is located 200-250 km east

of the trench and 120-150 km above the Benioff-Wadati zone defining the slab. This is historically the most

active volcano in Peru. Ubinas volcano begun erupting once more on March 25th 2006. The Geophysical

Institute of Peru (IGP) with the cooperation of the Institut de Recherche pour le Developpement (IRD-France)

has carried out the monitoring of seismic activity associated to this eruptive process. Seven broadband stations

equipped with CMG-40T seismometers were setup around the volcano during several weeks (Figure 3). About

30 explosions were recorded during this period with more than 4 stations. An example of explosion is shown in

Figure 1. These explosions correspond to the destruction of a magmatic dome formed at the bottom of a crater

situated in the south part of the sommital caldera. The dome is contiguous with the southern wall of the crater.

Some explosions are preceded by the occurrence of a LP swarm (Figure 2). The LP events are monochromatic

with a dominant frequency between 2 and 4 Hz. They have a similar shape and identical frequency content,

suggesting a unique source area. These swarms can last several minutes to more than one hour. The frequency of

the LP events is increasing with time while approaching the explosion. A few tens of minutes before the

explosion, the events are close enough in time to constitute a tremor (Figure 2).

Figure 1. Example of explosion (Energy, waveform, LP events).

Figure 2. Zoom of several successive LP events.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 334-336

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Method

The displacement field generated by a seismic source is described by representation theorem, which, for a

point source, may be written in frequency domain,

s

n

u ( ) =np ,q

s

G ( )pqM +

np

s

G ( )pF ( ) , n,p,q=x,y,z

where s

n

u is the n-th component of displacement at the stations s, np ,q

s

G and np

s

G are the Green’s functions,

pqM and pF are respectively the force-couples and single forces applied at the source. The waveform

calculations are performed over a grid of 20mx20mx20m, yielding a 3-D mesh with 576x550x240 nodes. The

synthetic seismograms are calculated in an homogeneous elastic medium including that takes the topography

into account. The Green’s functions are computed for each station and each component over a 3-D grid situated

under the volcano. The grid spacing is 200mx200mx100m, with a total of 8977 nodes.

Preliminary results

We have performed a waveform inversion for several explosions. Figure 4 shows the result obtained with the

explosion recorded the June 23th of 2006 by 7 stations. The computed position is located exactly at the bottom

of the crater (Figure 3). As shown in Figure 5, the fit is acceptable for most of the stations. Worse results were

obtained with the other explosions. The sources are positioned apart from the crater although the waveforms are

rather well adjusted for part of the stations.

This could be explained by too large grid spacing used for Green’s function sampling. A thinner sampling will

improve the spatial derivatives of the Green’s functions.

Using an heterogeous velocity model may also be able to improve the wave form fits.

Figure 3. Topographic map showing the station position and the source location.

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Figure 4. Waveform fits. Blue and red lines denote respectively real data and synthetics. The three components are plotted for each station (E,N,Z). The relative error between data and synthetics is 15%.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 337-338

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Tectono-magmatic evolution and crustal growth along west-central Amazonia since the late Mesoproterozoic: Evidence from the Eastern Cordillera of Peru

Aleksandar Mi kovi & Urs Schaltegger

Earth Sciences Section, University of Geneva, 13 rue des Maraîchers, CH 1205 Geneva, Switzerland

Whereas the Cretaceous to recent orogenic cycle is well characterised (Ramos and Aleman; 2000), the

knowledge of the early Phanerozoic and Proterozoic evolution of the Andes is increasingly fragmentary with age

due to paucity of exposed lithologies. The problem is less pronounced along the Peruvian segment of the orogen

where a lacuna in the ubiquitous Cenozoic volcanic cover is interpreted to have resulted from the flat slab

subduction of the Nazca ridge (Jaillard et al., 2000). Batholiths of the Eastern Cordillera of Peru which straddle

the tectonic boundary between the allochthonous western Amazonian tectonic provinces of San Ignacio (1.57-

1.24 Ga) and Sunsás (1.19-0.92 Ga) on one side and comparatively few parautochthonous to allochthonous

crustal domains (1.9-1.8 Ga Arequipa-Antofalla; 150 Ma Olmos-Amotape terrane) on the other, thus provide an

optimal record of the nature and rate of crustal growth at a long lived, non-accretionary cratonic margin. Despite

its fortuitous setting however, the timing of magmatism in the central Andes is relatively poorly understood with

most of the geochronological work so far relying heavily upon whole rock Rb-Sr and K-Ar techniques, both of

which are known to yield ambiguous dates thanks to low retention temperatures and a possibility of isotopic

disturbance by subsequent tectono-thermal episodes. This is a particularly acute problem in Peru considering

~150 Ma of uninterrupted compressive tectonism of the last Andean cycle (Benavides, 1999).

We use a combination of in situ U-Pb geochronology and Lu-Hf isotopic tracing of plutonic zircons along the

strike of the Eastern Cordillera of Peru to construct a detailed geochronological framework and identify sources

of consecutive magma pulses in order to define cratonic domains and track crustal evolution of the proto-Andean

margin of Amazonia. By relating the secular changes in magma sources to the tectono-magmatic cycles of

continental assembly and breakup over the last 1.1 Ga, we can test both the current geodynamic scenarios for the

evolution of the western Amazonian shield with particular focus on the poorly understood break up of Rodinia

(Meert and Torsvik, 2003; Loewy et al., 2003; Cordani et al., 2003; Fuck et al., 2008; Li et al., 2008) and the

models constraining the relative contributions of Phanerozoic and Neoproterozoic arc magmatism in the

formation of the continental crust (Condie, 2001; Davidson and Arculus, 2005).

The results of a laser ablation ICPMS U-Pb isotopic study on zircons from 60 Eastern Cordilleran intrusives of

Peru reveal 1.15 Ga of magmatic activity along the central western Amazonian margin that is largely dominated

by mid-Phanerozoic plutonism related to the assembly and break up of Pangea. A Carboniferous-Permian (340-

285 Ma) continental arc is identified along the orogenic trend from Ecuadorian border (6oS) to the inferred

inboard extension of the Arequipa –Antofalla terrane in the southern Peru (14oS). The widespread crustal

extension and thinning which affected the western Gondwana throughout Permian and Triassic resulted in the

central late to post orogenic La Merced-San Ramón-type anatectites dated between 275 and 220 Ma while the

emplacement of the southern Cordillera de Carabaya peraluminous granitoids in the late Triassic to early

Jurassic (220-190 Ma) represents, temporally and regionally, a separate tectono-magmatic event likely related to

re-suturing of the Arequipa-Antofalla block. Alkaline volcano-plutonic complexes and stocks associated with the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 337-338

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onset of the modern Andean cycle in southeastern Peruvian Andes cluster between 170-180 Ma. A

volumetrically minor intrusive pulse of Oligocene age (~30 Ma) is detected near the SW Cordilleran border with

Altiplano, and only one remnant of the late Ordovician intrusive belt is recognised in the Cuzco batholith (446.5

± 9.7 Ma) indicating that the Famatinian arc system previously identified in Peru only along the north-central

Cordillera Oriental and the coastal Arequipa terrane had also developed inboard of this para-autochthonous

crustal fragment. Both post-Gonwanide and Precambrian plutonism are restricted to isolated occurrences

spatially comprising less than 15% of the Eastern Cordillera intrusives. Hitherto unknown occurrences of the late

Mesoproterozoic and middle Neoproterozoic granitoids from the south central cordilleran segment define

magmatic events at 691 ± 13, 751 ± 8, 985 ± 14, and 1071 to 1123 ± 23 Ma that are broadly coeval with the

Braziliano and Grenville-Sunsás orogenies. Our data suggest the existence of a contiguous orogeny > 3800 km

along western Amazonia during the formation of Rodinia and its “early” fragmentation prior to 690 Ma.

In addition to dating the emplacement of plutonic rocks, we performed an in situ the LA MC-ICPMS survey of

the Hf isotope systematics on magmatic zircons from the Eastern Cordillera batholiths. These are invariably

characterised by a range in the initial 176Hf/177Hf compositions for a given intrusive event suggesting mixing of

material derived from the Paleoproterozoic crustal substrate and variable Neoproterozoic to recent juvenile

sources. The periods of well documented compressive tectonics correspond to negative mean eHfi values of -

6.73, -2.43, -1.57 for the Ordovician Famatinian, Carboniferous-Permian and late Triassic respectively,

suggesting the minimum crustal contribution between 74% and 44% by mass. The average initial Hf systematics

from granitoids associated with intervals of regional extension such as the middle Neoproterozoic, Permian-

Triassic and Cenozoic Andean back arc plutonism are consistently shifted toward the positive values (mean eHfi

= -0.7 to +8.0) indicating systematically larger inputs of juvenile magma (22% to 49%). In the absence of

evidence for lateral accretion of exotic crust, the time integrated Hf record from the central proto-Andean margin

of western Amazonia suggests crustal reworking as the dominant process during episodes of arc magmatism and

implies that most of continental growth took place vertically via crustal underplating of isotopically juvenile,

mantle derived magma during intervals of crustal attenuation.

References Benavides, V., 1999. Orogenic evolution of the Peruvian Andes: The Andean Cycle. Geology and ore deposits of the Central

Andes, SEG Spec. Pub., 7, 61-107 Condie, K. C., 2001. Rodinia and continental growth. Gondwana Research, 4, 154-155. Cordani, U. G., Brito-Neves, B. B., D’Agrella-Filho, M, S., 2003. From Rodinia to Gondwana: a review of the available

evidence from South America. Gondwana Research, 6, 275-283. Davidson, J. P., Arculus, R. J., 2005. The significance of Phanerozoic arc magmatism in generating continental crust. In:

Evolution and differentiation of the continental crust (eds.) M. Brown & T. Rushmer, Cambridge Univ. Press, 135-172. Fuck, R. A., Brito, B. B., Schobbenhaus, C., 2008; Rodinia descendants in South America. Prec. Research, 160, 108-126. Jaillard E., Hérail, G., Monfret, T., Díaz-Martínez, E., Baby, P., Lavenu, A., Dumont, J.F., 2000. Tectonic evolution of the

Andes of Ecuador, Peru, Bolivia and northernmost Chile. In: Tectonic Evolution of South America. (Eds.) U. Cordani, E.J. Milani, A. Thomaz Filho, & M.C. Campos Neto, Rio de Janeiro, 635-685.

Li, Z. X., Bogdanova, S.V., Collins, A.S., Davidson, A., de Waele, B., Ernst, E.E., Fitzsimons, I.C.W., Fuck, R.A., Gladkochub, D.P., Jacobs, J., Karlstrom, K.E., Lu, S., Natapov, L.M., Pease, V., Pisarevsky, S.A., Thrane, K., Vernikovsky, V., 2008. Assembly, configuration and break-up history of Rodinia: a synthesis. Precambrian Research, 160, 179-210.

Loewy, S. L., Connelly, J. N., Dalziel, I. W. D., Gower, C. F., 2003. Eastern Laurentia in Rodinia: constraints from whole-rock Pb and U/Pb geochronology. Tectonophysics, 375, 169-197.

Meert, J.G., Torsvik, T.H., 2003. The making and unmaking of a supercontinent: Rodinia revisited. Tectonophysics, 375, 261-288.

Ramos, V. A., Aleman, A., 2000. Tectonic Evolution of the Andes. In: Tectonic Evolution of South America. (Eds.) U. Cordani, E.J. Milani, A. Thomaz Filho, M.C. Campos Neto, Rio de Janeiro, 635-685.

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Seismic tomography of the Cotopaxi volcano, Ecuador

V. Monteiller, J.-P. Métaxian, B. Valette, & S. Araujo

LGIT, Université de Savoie, 73376 Le Bourget du lac cedex, France (vadim.monteiller@univ-savoie.fr)

KEYWORDS : volcano, Cotopaxi, seismology, tomography

Volcanic settings

The Cotopaxi volcano (5897m) is situated in the Cordillera Real of Ecuador, 60km south of Quito. This active

andesitic volcano, with a base diameter of 25 km and almost 3000m of relief is covered by an icecap on the

uppermost 1000m of the cone. The seismic monitoring is performed by the Instituto Geofisico of the Escuela

Politécnica National since 1989 using a network of short-period seismic stations set up on the volcanic shield.

The local set of data comes from an experiment performed in 1996-97 on Cotopaxi volcano. The array was

composed of 4 classical short period stations employing L4-3D or L4C seismometers and 8 telemetered stations

divided in two groups of four stations comprising sub-arrays which had separate reception and acquisitions units.

The experiment was carried out in two phases. First some stations were installed on the volcanic cone warying

azimuths and distances from the crater and other stations in a wide area around the volcano, up to 20 km from

the summit. In a second phase, part of the equipment was moved closer to the crater in order to record in greater

detail the volcanic activity concentrated below the summit area. One station was set up along the edge of the

crater on a rock base at the elevation of 5820m with the aim to better constraint the structure and the localization

on the uppermost 1000m of the cone occupied by the glacier. In total, 16 different sites were occupied during

this experiment, additionally to the 4 permanent sites of the IGEPN array.

The tomographic inversion was performed using 6425 P arrivals times from 1147 earthquakes.

Results

We used the tomographic algorithm of Monteiller et al. (2005). Travels times are computed by solving the

Eikonal equation using a finite-difference approach (Podvin and Lecomte 1991). The inverse problem is solved

by using a probabilistic approach (Tarantola Valette 1982). Figure 2 display the 3-D P-wave velocity model. We

Figure 1. Seismic network. Red triangles indicate the position of the seismic stations.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 339-340

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show horizontal and vertical slices at different depth, latitude and longitude. The viewpoint position is South-

West.

The result clearly show the presence of high velocity anomaly surrounded by low velocity ring interpreted as

effusive products associated to the construction of the Cotopaxi. The high velocity anomaly is interpreted as a

solidified intrusive magma body.

Figure 3 display slices of 3-D structure velocity model. The yellow patch delimit the iso-velocity surface

corresponding to 3350 m/s. The gray patch delimit the topography of Cotopaxi. The viewpoint position is South-

East.

Figure 2. P-wave velocity model cross sections.

References Monteiller, V., Got J.-L., J. Virieux and P. Okubo, An efficient algorithm for double-difference tomography and location in

heterogeneous media, with an application to Kilauea volcano, J. Geophys. Res., 110, B12306, doi:10.1029/2004JB003466 Podvin, P., & Lecomte, I., 1991. Finite difference computation of traveltimes in very contrasted velocity models: a massively

parallel approach and its associated tools., Geophys. J. Int., 105, 271-284. Tarantola, A. and Valette, B., 1982, Inverse problems = quest for information., J. Geophys., 50, 159-170.

Figure 3. 3-D P-wave velocity model.

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Analysis of the January 23, 2007 Aysén swarm using joint hypocenter determination

Cindy Mora1, Diana Comte

1, Ray Russo

2, & Alejandro Gallego

2

1 Dept. Geofísica, Universidad de Chile, Santiago, Chile (cindy@dgf.uchile.cl).

2 Department of Geological Sciences, University of Florida, Gainesville, FL, USA (rrusso@ufl.edu).

KEYWORDS : Aysen swarm, crustal seismicity, Southern Andes, Liquiñe-Ofqui fault

Introduction

It is well known that the Chile margin triple junction is one of the few good examples of where an active

spreading oceanic ridge is colliding with an active continental margin. In this location the Chile Rise intersects

the sediment-filled Peru-Chile Trench (Figure 1). North of the Triple Junction, the Nazca plate is subducting

below the South American plate at 7 cm/yr, whereas to the south, the Antarctic plate is subducting below the

South American plate at about 2 cm/yr. For the last 15 m.y. the directions of plate motions between the South

American, Nazca, and Antarctic plates have remained relatively constant (Cande, 1983). As a result, three

segments of the Chile Rise spreading center, separated by transform fault offsets, have converged and collided

with the southern margin of Chile. A fourth ridge segment, separated by the Taitao and Darwin fracture zones,

has recently been buried under the sedimentary fill of the Peru-Chile Trench. As collision progressed, the Triple

Junction has migrated northward to its present position at 46°S. The Golfo de Penas and Taitao Peninsula, just

south of the present position of the Triple Junction, are parts of the western margin of the South American plate

which have overridden short segments of the Chile Rise about 5-6 and 2.5-4 Ma, respectively.

Successive Chile Rise segment collision generated changes in surface geology of South American plate such

as volcanic gaps and younger plateau lavas in Southernmost Chile, due to magma ascending through slabs

windows formed as the subducting ridge continue the spreading, and acts as well as an indent and main force in

the regional tectonism (Thompson, 2002). Such forces and the difference in subduction velocities cause stress

partition along the trench between 38° and 42°S, the detachment of the sliver west to the Liquiñe-Ofqui fault and

clockwise and counter clockwise rotations east and west of the fault (Rosenau et al, 2006).

Geological Setting

The Liquiñe –Ofqui Fault Zone (Figure 1) is located between 38°S -48°S in the Austral Andes, presenting two

~1000 km long main parallel to the trench lineaments, between 39°-44°S and 44°-47°S, connected by N-E

trending echelon lineaments and concave to the ocean lineaments of tens of kilometers (Cembrano et al, 1996).

The volcanoes from the Austral Andes Volcanic Zone are distributed along the main N-S faults and on the

duplex centers, 250 - 300 km east from coast, also North Patagonian Batholit (NPB), the old volcanic arc,

contains the main fault at its central part. Field studies of the fault zone reveal both ductile and brittle

deformation in rocks from the NPB (Cembrano et al, 1996) that are consistent with actual dextral movement,

although older mylonitic deformation associated with the fault may indicate that the original sense was left-

lateral and took places at deeper levels in the crust.

The paleomagnetic data of blocks in the Liquiñe-Ofqui fault are consistent with vertical axis rotation in a

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 341-343

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buttressed fault system (Cembrano et al, 2002; Beck et al, 1993), showing rotations up to 31±4° and 9±1°

respectively. The fault zone is believed to have formed together with the volcanic arc during the Eocene-

Miocene due to partitioning of deformation in the forearc in South American plate, in a transpresional process

owning to oblique subduction of Nazca plate (Pardo-Casas and Molnar, 1987) and changes in the sense of

motion and angle (Cembrano et al, 1996).

Seismicity along the Liquiñe-Ofqui fault has been poorly studied, mainly because teleseismic events have been

few (Cifuentes, 1989), and almost no local or regional seismicity associated with the fault has been recorded

because of a lack of local seismic networks in the proximity of the fault.

Through a joint project between Universidad de Chile and University of Florida a temporary network (Fig. 1)

was deployed during two years, between 43°-49°S and 71°-76°W, comprising island and inland, with 60

seismometers and 4 fat stations consisting in broadband STS-2 Streckeisen stations accompanied by closely

spaced short period seismometers. At the end of this period, on January 23, a seismic swarm began in the Aysen

region, near Puerto Chacabuco with a mainshock Mw=5.2, presenting a dextral strike slip focal mechanism. This

swarm had its major activity during January until April of the same year, recording more than 1000

microearthquakes per month, decreasing the activity towards December, 2007 (Chilean Seismological Service).

The Joint Hypocentral Determination

During the January 1-27, 2007 period the temporary network recorded more than 300 events in the Aysen

fjord, where the majority of them occurred between January 23 and 27, 2007. Data was read and located with

SEISAN program and relocated with the Joint Hypocentral Determination (JHD, Dewey, 1972, Douglas, 1967)

in five different groups corresponding to the events of each day in the swarm period, using the mainshock as

master event. Dewey’s method assumes that travel times anomalies of P waves are identical to all common

Figure 1. Distribution of temporary network near Liquiñe-Ofqui Fault Zone (thick black line). Inverted red triangles correspond to seismic stations and red squares correspond to fat stations. Fracture Zones are represented by thinner white lines; ridge segments in thin white lines. The Antartic and Nazca plate vector movement are shown with red arrows. The red lined square denotes the study zone.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 341-343

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station in a group of events, then the difference in arrival times is due to differences in the hypocenter only.

Relative to the master event location, latitude, longitude, depth and origin time differences on each after- and

fore-shocks are iterated until solution converge (Marshall and Russo, 2005).

After relocating the events, changes in depth are significant, showing that the events are finally constrained to

depths between 0 to 7 km (before the JHD procedure, this distribution in depth was between 0 to 15 km depth)

(Figure 2.). The results show an area confined in about 11 km in latitude and 33 km in longitude. The new

locations suggest that this segment of the Liquiñe-Ofqui fault is deeping towards west and trending south, which

is in agreement with the main trending of the Liquiñe-Ofqui fault.

The evidence of seismicity associated with the Liquiñe-Ofqui fault system clearly indicates that seismic

hazard estimations in this region must be re-estimated.

Acknowledgement This work was funded by National Science Foundation-USA N°0126244 and CONICYT-CHILE grant Nº1050367 . The authors particularly acknowledge to Universidad de Chile for the scholarship “Becas de Estadias cortas de Investigación Destinadas a Estudiantes Tesistas de Doctorado y Magíster de la Universidad de Chile” to work on this investigation.

References Beck, M.E., C. Rojas, J. Cembrano, 1993. On the nature of buttressing in margin-parallel strike-fault systems, Geology, 21,

755-758. Cande, S.C., 1983. Nazca-South America Plate Interactions 80 my BP to present. EOS. Cembrano, J., Hervé, F., Lavenu, A., 1996. The Liquine–Ofqui fault zone: a long-lived intra-arc fault system in southern

Chile. Tectonophysics 259, 55– 66. Cembrano, J., Lavenu, A., Reynolds, P., Arancibia, G., Lopez, G., Sanhueza, A., 2002. Late Cenozoic transpressional ductile

deformation north of the Nazca-South America-Antarctica triple junction. Tectonophysics, 354, 289-314. Cifuentes, I.L., 1989. The 1960 Chilean earthquakes. Journal of Geophysical Research 94 B1, pp. 665–680. Dewey, J.W., 1972. Seismicity and tectonics of western Venezuela. Bull. Seismol. Soc. Am. 62, 1711 –1751. Douglas, A. Joint Epicentre Determination. Nature,vl. 215, is 5096, p. 47, July 1967 Marshall, J.L., Russo, R.M.,2005. Relocated aftershocks of the March 10, 1988 Trinidad earthquake: Normal faulting, slab

detachment and extension at upper mantle depths. Tectonophysics, 398, 2005, 101-114. Pardo-Casas, F. and Molnar, P., 1987. Relative motion of the Nazca (Farallón) and South American Plates since Late

Cretaceous time. Tectonics 6, 233–248. Rosenau, M., Melnick, D. and Echtler, H., 2006. Kinematic constraints on intra-arc shear and strain partitioning in the

southern Andes between 38ºS and 42ºS latitude. Tectonics, 25, TC4013, doi:10.1029/2005TC001943. Thomson, S., 2002. Late Cenozoic geomorphic and tectonic evolution of the Patagonian Andes between latitudes 42 degrees

S and 46 degrees S; an appraisal based on fission-track results from the transpressional intra-arc Liquiñe-Ofqui fault zone. Geological Society of America Bulletin, v.114; no.9; p.1159–1173.

Figure 2. Latitude and longitude versus depth profile for the relocated January 23-27 Aysen swarm data. Colours denote depth.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 344-347

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Further evidences of Quaternary activity of the Maradona faulting, Precordillera Central, Argentina

S. M. Moreiras & A. L. Banchig

CONICET – IANIGLA (CCT), Av Ruiz Leal s/n, Parque Gral. San Martín, CP 5500, Mendoza, Argentina

(moreiras@lab.cricyt.edu.ar, abanchig@lab.cricyt.edu.ar, abanchig@iinfo.unsj.edu.ar)

KEYWORDS : morphotectonic features, active fault, landslides, Quaternary, Precordillera

Introduction

The Maradona faulting is located in the Pampa of Maradona (San Juan province), an NS elongated valley that

runs parallel to the Eastern Precordillera and the Central Precordillera (31º 46´ LS - 68º 51´) (Figure 1)

continuing towards the south as the Pampa of Bachongo. This faulting, with a total extension of 32 km, was

initially identified by photo-interpretation and was considered as a Lower-Middle Pleistocene age fault (Bastías

et al., 1984). Further research deduced that this faulting is an expression of the “Matagusanos-Maradona-

Acequión” tectonic belt (Perucca, 1990; Perucca et al., 1990) located between the Eastern Precordillera and the

Central Precordillera. Both, geological provinces has a completely different structural behaviour. Whereas, the

former is characterised by a skinned fold - thrust belt with occidental vergence where mainly Cambric-

Ordovician carbonatic rocks outcrop, the latter corresponds to overthrustings with eastern vergence and its

outcroppings are Palaeozoic rocks of talus or outer marine platform.

Even through, the Maradona faulting is reported as active during Quaternary (Amos, 1981) detailed

geomorphological studies have not been carried out in this area. Hence, this research is focussed to amend this

lacking analysing preserved geomorphological expressions of this faulting that may be engaged to evaluate the

regional seismic hazard.

Morpho-tectonic features of Maradona faulting

The Maradona faulting is easily identified by remote sensing due to the undoubtedly alignment generated by

the fault scarpe on the eastern edge of the valley. However, detailed field examinations let to identify another

eastern fault (F2) associated with this compressive system.

The fault (F1), indirectly identified by previous authors, was recognised in the field southern of the Maradona

farm, where Tertiary outcrops of Mogotes formation (173º, 53º E) overlap Quaternary alluvial deposits. This

reverse fault has a plane with N-S trend (12º) dipping 44º to the West and a a vertical offset of 5 m was

measured. Moreover, this faulting also exposes Tertiary rocks over alluvial fans in Bachongo place, where

direction measured on the fault plane is 335º and it dip of 32º to 42º to the West.

Whereas, the second faulting (F2) was recognised in the surroundings of the Papagallos farm, where this

reverse fault affects a 20 m-thick sequence of alluvial fans. There, Quaternary deposits are deformed dipping 20º

to the West and they are covered by a sequence of barreal-lake deposits generated as a consequence of dam of

the Papagallos River resulting from the western block up-lift.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 344-347

345

Figure 1: Map of the study area showing Maradona faulting (F1 and F2) and landslides distribution.

Geomorphic markers of Quaternary activity

The F1 is evidenced by a 9 km-long scarpe extending from the Papagallos River to Bachongo locality. This

morphology exposes its free face towards the east as a result of western block lift. It generated elongated hills

with N-S trending in the Eastern edge of the Pampa of Maradona (Figure 1). Likewise, the F1 offsets distal

piedmont sediments of alluvial-fan northward the Papagallos River where ephemeral streams are interrupted or

displaced 300 to 500 m towards the north.

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In Bachongo place, Tertiary levels and alluvial fans are folded just on the fault plane of F1 (335º, 32º to 42º

W). The latter show an asymmetrical anticline (335º, 30º W and 82º E); while alluvial fans are forming an

anticline with a vertical western flank and an eastern flank dipping 25º E.

Then, the presence of the 5 to 8 m-thick lake sequence indicate the existence of a pond related to the F2

faulting, being a clear marker of strike slip component of this faulting. The extension of this paleo-lake is

uncertain. Similar disconnected fine deposits were observed up-stream of the Papagallos River, reason why they

could not be properly correlated. Even though, the lake sequence is thinner in the Maradona farm, as it is

expecting in lake borders, the maximum elevation of these fine deposits are higher than in the Papagallos farm.

The Maradona faulting has favoured water springs; hence numerous of them are markedly aligned to N-S fault

traces. The trace of the F1 coincides with the springs of Maradona, Agua Raja, Bachongo and Pampa Amarilla

farms; while the F2 runs over those existing at Papagallos and La Ciénaga farms.

These places located in this arid region used to be important farms since the beginning of XVIIIth century due

to their water resource.

Landslide occurrence

At the latitude of study area, nine slumps are recognised in the eastern slope of the Sierras de las Osamentas

(Fig. 1). The total debris volume mobilised by these rotational slides reaches 1,2x109 m3.

These landslides are spatially distributed close to traces of Maradona faulting, approximately 8.5 km and

16 km far from the F1 and the F2, respectively. Likewise, they are located in an area of 12 km2, reason why a

seismic shaking of magnitude >5 is proposed. The earthquake recorded on January 24th 1978 (Ms 5,7) was

related to the Maradona faulting (Inpres, 1982) and at least four epicentres of Ms=4.2 earthquakes are located in

this region.

Moreover, recurrence of paleo-earthquakes is assumed because at least three of these slumps correspond to

multiple events. In these cases, ancient relit deposits are overlaid by younger deposits. Still numerical dating of

these events is missing.

These triggering earthquakes activating previous landslide scarps could be more recently. The M1 landslide

located close to the Agua Pinto farm have been reactivated in the last century affecting the route connecting the

Pampa of Maradona valley with Barreal valley that was also used in the past by indigenous (Huarpes).

Conclusion

The present geomorphological study reveals that Maradona faulting system has more than one fault trace what

is fundamental for future seismic hazard evaluation as a partial earthquake record could be obtained studying

only the F1. According to morpho-tectonic features observed in the field, this faulting shows two dextral oblique

reverse faults. Moreover, many evidences about recent tectonic activity of this faulting could be found. The

offset of Quaternary alluvial surfaces (Gaudemer et al. 1989; Audin et al, 2003, 2006) and gathered distribution

of landslides triggered by earthquakes (Keefer, 1984, 1987) have been successfully used to identifying active

faults systems.

In Argentina seismic hazard lacks on paleo-seismic assessment (Costa et al., 2000), hence our work let

advance in this problem, still numerical dating is missing.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 344-347

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Referencias Amos, A.J., 1981. Fallas activas en la Republica Argentina Actas VIII Congreso Argentinote Geología. Audin L., Herail G., Riquelme R., Darrozes J., Martinod J., Font E., 2003. Geomorphological markers of faulting and

neotectonic activity along the western Andean margin, northern Chile. Journal of Quaternary Science, 18 (8), pp. 681-694. Audin L., David C., Hall S., Ferber D., Hérail G., 2006. Geomorphic evidence of recent tectonic activity in the forearc of

Perú. RAGA 61(4): 545-554. Bastias, H., Weidmann, N., y Perez, M., 1984, Dos zonas de fallamiento Plio-Cuaternario en la Precordillera de San Juan.

Actas IX Congreso Geologico Argentino, Bariloche. Vol. II: 329-341. Costa C., Machette M.N., Dart R.L., Bastias H.E., Paredes J.D., Perucca L.P, Tello G.E:, Haller K.M., 2000. Map and

database of Quaternary faults and folds in Argentina. A project of the International Lithosphere Program Task Group II-2, Major Active Faults of the World. United States Geological Service open file 00-0108.

Gaudemer, Y., Tapponnier, P., Turcotte, D.L. 1989. River offsets across active strike-slip faults. Ann. Tecton, 3 (2), pp. 55-76.

INPRES, 1982, Microzonificacion sismica del valle de Tulum, Provincia de San Juan: Resumen Ejecutivo, San Juan,120 p. Keefer, D.F., 1984. Landslides caused by earthquakes. Geol. Soc. America Bulletin, 95: 406-21. Keefer, D.F., 1987. Landslides as indicators of prehistoric earthquakes. Directions in paleosismology. U.S. Geol. Survey

Open File report 87-673, pp: 178-180. A. Crone and E. Modal (editors). Perucca, L., 1990. Sistema de fallamiento La Dehesa-Maradona-Acequión, San Juan, Argentina. Actas XI Congreso

Geológico Argentino, San Juan. Vol. II: 431-434. Perucca, L., Sanches, A., and Uliarte, E., 1990, Morfoneotectónica en la zona norte del corredor tectónico Matagusanos-

Maradona-Acequión, San Juan, Argentina. Actas XI Congreso Geológico Argentino, San Juan. Vol. II: 435- 438. Ramos V., 1999. Las provincias geológicas del territorio Argentino. En: Geología Argentina (Ed. R Caminos). Instituto de

geología y recursos minerales. Anales 29 (12): 41-96.

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Contemporary forearc deformation in south-central Chile from GPS observations (36-39°S)

Marcos Moreno1, Jürgen Klotz

1, Daniel Melnick

2, Helmut P. Echtler

1-2, & Klaus Bataille

3

1

GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany 2 Institut für Geowissenschaften, Universität Potsdam, Potsdam, Germany

3 Departamento de Ciencias de la Tierra, Universidad de Concepción, Chile

Introduction

Present-day surface deformation along active forearc regions primarily responds to the earthquake cycle phases

[e.g., Savage, 1983; Thatcher, 1984] and locally to upper plate structures [e.g., McCaffrey and Goldfinger,

1995]. Mega-earthquake cyclic deformation is a transient process and is conditioned by the mechanical coupling

between both plates. Local upper plate structures like active forearc crustal-scale faults may accommodate and

partly release the elastic strain accumulation during the earthquake cycle [e.g., Park, et al., 2002]. Similarities

between syntectonic sediments and differential coastal uplift and tilting by repeated slip on crustal faults suggest

that some of these structures have controlled the surface deformation in some forearc regions over 106 years

[e.g., Melnick, et al., 2006]. Active crustal faults can delimit forearc slivers, which may rotate or translate with

respect to a stable continental frame, producing different patterns of surface deformation [e.g., Fitch, 1972;

Jarrard, 1986].

Regional GPS survey over the last 4 years analyses the present-day deformation in the forearc of the south-

central Chile margin, especially on the Arauco-Nahuelbuta forearc block [Hackney et al., 2006]. The Arauco-

Nahuelbuta block defines the overlapping area of two mega-thrust earthquake rupture zones: the Valdivia 1960

and Concepción 1835. There, active crustal faults have been identified and mapped based on deformed coastal

geomorphic features, seismic-reflection profiles, and micro-seismicity [Melnick and Echtler, 2006]. GPS data

and finite-element models are presented to gain insight into forearc kinematics and particularly to address the

role of crustal faults in the seismotectonic segmentation.

Results and discussions

Forearc deformation varies markedly from north to south in the Arauco-Nahuelbuta block. Present-day

deformation in the forearc segment defined by the rupture zone of the 1835 earthquake clearly represents strain

accumulation due to the ongoing interseismic phase. Whereas, surface deformation in the adjacent southern

forearc segment, which in turn is part of the 1960 earthquake rupture zone, manifests the effects of protracted

and still ongoing postseismic mantle rebound in addition to locking of the seismogenic zone.

GPS observation and modeling results suggest that the coupling zone narrows southward. The change in

downdip depths between the 1835 and 1960 earthquake-rupture segments may be a result of: (a) differential age

gradient of the incoming oceanic plate parallel to the trench; (b) or the effect of the lower plate Mocha fracture

zone, which in its projection below the active margin corresponds to our limits between the northern and

southern domains (Fig.1).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 348-350

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Finite-element models better reproduce the GPS velocities by including the Santa María fault, which is rooted

in the plate interface, and accommodates about 30 % of the margin-parallel plate convergence. The effect of the

fault deformation reaches to 30 km from the fault, and may be responsible for the velocity gradient observed in

the GPS observation [Moreno et al., 2008].

A rotating forearc block is identified in the northern limit of the 1960 rupture zone, bounded by the Lanalhue

crustal-scale fault [Moreno et al., 2008]. Maximum rotation at the edge of this block is accommodated by diffuse

deformation across the Lanalhue fault, which seems to be actually locked. We explain this block rotation as a

result of convergence between the Chiloé forearc sliver and a buttress formed by the Arauco-Nahuelbuta block

(Fig. 1). Geological and geomorphic data also support margen- parallel shortening interpreted as micro-plate

collision [Melnick et al., 2008].

Our results emphasize the importance of crustal-scale faults in the present-day surface deformation in active

forearcs. Crustal faulting can produce forearc block segmentation and distributed stress and strain during the

interseismic phase. These structures thereon have an important role for the seismotectonic segmentation of

subduction zones and control the geometry of the rupture area of the mega-thrust earthquakes. In consequence

these results have major implications for the evaluation of the state of stress and the earthquake recurrence.

Figure 9. Schematic map showing the main tectonic features in the south-central Chile forearc. The convergence between the Chiloé forearc sliver and the Arauco-Nahuelbuta block may explain the counterclockwise rotation in the northern limit of the 1960 rupture zone.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 348-350

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References Fitch, T.J. (1972), Plate convergence, transcurrent faults, and internal deformation adjacent to southeast Asia and the western

Pacific, J. Geophys. Res., 77, 4432-4460. Hackney, R., H. Echtler, G. Franz, H. Götze, F. Lucassen, D. Marchenko, D. Melnick, U. Meyer, S. Schmidt, Z. Ta árová, A.

Tassara, and S. Wienecke (2006), The Segmented Overriding Plate and Coupling at the South-Central Chilean Margin (36-42°S), The Andes - Active Subduction Orogeny, Frontiers in Earth Sciences, vol.1, edited by Oncken, O., G. Chong, G. Franz, P. Giese, H. Götze, V. Ramos, M. Strecker, and P. Wigger, pp. 355-374, Springer-Verlag, Berlin.

Jarrard, R.D. (1986), Terrane motion by strike-slip faulting of forearc slivers, Geology, 14, 780-783. McCaffrey, R., and C. Goldfinger (1995), Forearc deformation and great subduction earthquakes: Implications for Cascadia

offshore earthquake potential, Science, 267, 856-859. Melnick, D., B. Bookhagen, M. Strecker, and H. Echtler (2008), Segmentation of subduction earthquakes from forearc

deformation patterns over hundreds to millions of years, Arauco Peninsula, Chile, J. Geophys, Res., under review. Melnick, D., B. Bookhagen, H. Echtler, and M. Strecker (2006), Coastal deformation and great subduction earthquakes, Isla

Santa María, Chile (37°S), Geol. Soc. Am. Bull., 118(11), 1463–1480, doi:10.1130/B25865.1. Melnick, D., and H. Echtler (2006), Inversion of forearc basins in southcentral Chile caused by rapid glacial age trench fill,

Geology, 34, 709–712. Moreno, M.S., J. Klotz, D. Melnick, V. Grund, H. Echtler, and K. Bataille (2008), Active faulting and forerarc block rotation

in the south Chile from GPS-derived deformation, Geochem., Geophys., Geosys., under review. Park, J.-O., T. Tsuru, S. Kodaira, P.R. Cummins, and Y. Kaneda (2002), Splay fault branching along the Nankai subduction

zone, Science, 297, 1157-1160. Savage, J.C. (1983), A Dislocation Model of Strain Accumulation and Release at a Subduction Zone, J. Geophys. Res., 88,

4984-4996. Thatcher, W. (1984), The earthquake deformation cycle, recurrence, and the time- predictable model, J. Geophys. Res., 89,

5674-5680.

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Regional tephro-stratigraphic correlation in the Ecuadorian coastal region

Patricia A. Mothes, Silvia Vallejo, & Minard L. Hall

Instituto Geofísico, Escuela Politécnica Nacional, casilla 1701-2759, Quito, Ecuador (pmothes@igepn.edu.ec)

KEYWORDS : volcanic tephra, regional correlations, Holocene uplift, Ecuadorian coast

In continental Ecuador about 30 volcanoes are considered potentially active (Barberi et al., 1991; Hall and

Beate, 1991). Most of them have produced eruptions during the Holocene, some in the range of VEI = 4 to 6.

Plinian and co-plinian ash sequences related to these eruptions are found in abundance in sedimentary layers in

the litoral zone, the product of direct aerial fall and via minimal water transport in local streams.

This study is creating a regional correlation of the ash sequences that we have collected from distinct sites

between northern Esmeraldas to southern Manabí. Ash and mineralological characteristics of each layer are

being determined by examination under binolcular microscope. Granulometric, chemical and radiocarbon

analises are in progress. Based on the results, and by comparing these distal ash samples with the character of

known eruptive products in the Sierra, parent volcanoes will be assigned.

The presence of tephra layers along the Ecuadorian litoral have been noted by various researchers, particularly

archaeologists, (Zeidler and Pearsall, 1994) and (Valdez, 2006), and geomorphologists (Dumont et al., (2006)

and (Usselmann, 2006) however until now no regional correlations exists between the various ash layers found

in the litoral zone.

Fig. 1: Map showing location of some sampling sites, coastal Ecuador, in province of Manabí.

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Fig. 2: Stratigraphic column in the Rio Muchacho drainage, near Canoa, coastal Ecuador

Application of the regional tephra correlation may contribute to constraining rates of recent coastal uplift from

active subduction processes, be used in determining rates of coastal subsidence and within an archaeological

context, possibly infer societal disruptions by heavy ashfalls onto early inhabitation sites.

Fig. 3: Cut in the Rio Muchacho channel (UTM 05662/99553). Observe the 5 layers of volcanic ash.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 351-353

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Fig. 4: Stratigraphic column from near San Isidro, west of Canoa.

References Barberi, F., Coltelli, M. Ferrara, G., Innocente, F., Navarro J.M. and Santacroce, R., 1988. Plio-Quaternary Volcanism in

Ecuador. Geol. Mag., 125 (1) 1-14. Dumont, J. F., Santana, E., Valdez, F., Tihay, J.P., Usselmann, P., Ituralde, D., and Navarette, E., 2006. Fan beheading and

drainage diversion as evidence of a 3200-2800 BP earthquake event in the Esmeraldas-Tumaco seismic zone: A case study for the effects of great subduction earthquakes. Geomorphology. Vol. 74 (1), p. 100-123.

Hall, M. L. and Beate, B. 1991. Pliio-Quaternary Volcanism in the Ecuadorian Andes. In: Mothes, P. (ed) The Volcanic Landscape of the Ecuadorian Sierra. Studies in Geography. Vol. 4, p. 5-18. Coorperación Editora Nacional, Quito (In Spanish).

Usselmann, P., 2006. Dinámica geomorfológica y medio ambiente en los sitios arqueológicos Chirije y San Jacinto/Japoto (costa del Manabí central, Ecuador) in: Boletín del Institute Francés de Estudio Andinos, Vol. 35 (No. 3) p. 257-264.

Valdez, F., (Editor), 2006. Agricultura ancestral, camellones y albarradas: Contexto social, usos y retos del pasado y del presente. Colección "Actas & Memorias" del IFEA. Quito, Ecuador p. 361.

Zeidler, J. A. and Pearsall, D. M. (Editors), 1994. Regional Archaeology In Northern Manabí, Ecuador: Environment, Cultural Chronology, and Prehistoric Subsistence In the Jama River Valley. University of Pittsburgh, USA. p. 224.

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Interactions between block rotations and basement tectonics in the Copiapó-Vallenar region, northern Chile: Preliminary results

Constantino Mpodozis1, César Arriagada

2, Pierrick Roperch

3, & Esteban Salazar

2

1 Antofagasta Minerals, Ahumada 11, Oficina 602, Santiago, Chile (cmpodozis@aminerals.cl)

2 Departamento de Geologia, Universidad de Chile, Casilla 13518, Correo 21, Santiago, Chile

(cearriag@cec.uchile.cl) 3 IRD, LMTG & Géosciences Rennes, campus de Beaulieu, 35042 Rennes, France (pierrick.roperch@ird.fr)

KEYWORDS : Central Andes rotation pattern, forearc, block rotations, paleomagnetism

Introduction

One of the most prominent tectonic features of the Andes is the Central Andean Rotation Pattern (CARP,

Taylor et al., 2005) which is characterized by, paleomagnetically determined, clockwise rotations in northern

Chile and counterclockwise rotations in southern Peru (Arriagada et al., 2006; Roperch et al., 2006; Taylor et

al., 2007). Recent studies by Somoza et al. (1999), Arriagada et al. (2006) and Roperch et al. (2006) suggest that

most of the CARP rotations were probably acquired during the Paleogene. A remaining problem, however, is the

still poor knowledge of the distribution and magnitude of the tectonic rotations in north-central Chile (Copiapó-

Vallenar region, 28º-31ºS, Figure 1) and its relation with the structural evolution of the Andean range at this

latitude. Along this region, placed over the modern “Pampean” flat-slab segment of the Andes, the Coastal

Region is formed by Jurassic- Cretaceous volcano-sedimentary sequences deformed by long wavelength folds

and intruded by a series of east-younging Mesozoic-Early Cenozoic plutonic complexes. By contrast, the Main

Andean Range (Frontal Cordillera) to the east, is formed by a series of large crystalline (Late Paleozoic) thick-

skinned basement blocks, including basement-cored anticlines bounded by east and west verging reverse faults

(Moscoso and Mpodozis, 1988). In plan view, a 30 km eastwards displacement of the basement front occurs

between ~28º45’-29ºS (“Vallenar Discontinuity”, VD. Figures 1 & 2). South of the VD the main basement-

bounding faults are NS oriented while to the north they trend to the NNE.

New paleomagnetic and structural data

We have recently obtained new structural and sampled more than 60 new paleomagnetic sites from this region.

Preliminary results from some of these sites added to already published paleomagnetic data, are shown in the

structural map of Figure 1. These data demonstrate that clockwise tectonic rotations are, consistently large (30º-

45º) north of the VD (Taylor et al., 2005; Arriagada et al., 2006; Taylor et al., 2007), but south of the VD

rotations seems to be much smaller and almost negligible south of 30ºS (Figure 1). At a regional scale, rotation

vectors are parallels to the orientation of the major faults bounding the basement blocks showing that the VD

seems to represent a regional “kink” both in the orientation of the paleomagnetic vectors and the trend of the

regional basement faults (Figure 2). Several anomalous tectonic features occur along the VD with the most

conspicuous being the termination of the NS and NNE basement faults along an EW structural corridor hosting

along its western edge an elliptical, NE oriented, Paleocene intrusive complex (Los Morteros Pluton, c.a 64 Ma,

Figure 2). An anomalous large rotation value (50º) was obtained from a paleomagnetic site sampled in Jurassic

red beds outcropping at the VD (Figure 1). Rapid facies changes in Mesozoic sedimentary and volcanic

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 354-356

355

sequences cropping out immediately to the west of the offset basement blocks suggest that the VD also coincide,

at least in part, with a soft - linked accommodation zone connecting overlapping rift basins that during the

Triassic and Jurassic were filled by thick terrigenous and carbonate marine sequences. Further to the east, in

Argentina, the VD coincides with a system of EW normal faults bounding a narrow basement horst that separate

the Oligocene Macho Muerto and Río de la Sal extensional basins (Mpodozis and Kay, 2003).

Figure 1: Main tectonic features and paleomagnetic database at Copiapó-Vallenar (north-central Chile, ~27º-30ºS).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 354-356

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Discussion

The Vallenar Discontinuity, that behaves as a complex EW “accommodation” zone during diverse periods of

Mesozoic and Cenozoic extension, probably corresponds to an inherited Paleozoic feature. The analysis of the

new structural and paleomagnetic data collected in the Copiapó-Vallenar region show that the VD separates two

tectonic domains North of the VD, the faults bounding the basement cored blocks of the Andean Range trend to

the NNE, while to the south the faults trend NS. Paleomagnetic vector parallels the trend of the major faults,

NNE to the north, NS-N10ºE to the south. Overall, the VD seems to act a ”hinge” or pivot zone between two

mayor rotation domains. Further studies are needed to understand the significance of this until now

underestimated tectonic feature.

References

Arriagada, C., Roperch, P., Mpodozis C. and Fernández, R. 2006. Paleomagnetism and tectonics of the southern Atacama Desert region (25-28ºS) Northern Chile. Tectonics, 25: TC4001, doi:10.1029/2005TC001923.

Moscoso, R. and Mpodozis, C., 1988. Estilos estructurales en el Norte Chico de Chile (28-31ºS), regions de Atacama y

Coquimbo, Revista Geológica de Chile, 15/2, 151-166.

Mpodozis, C. and Kay, S. M, 2004. (Abs) Neogene tectonics, ages and mineralization along the transition zone between the

El Indio and Maricunga mineral belts (Argentina and Chile 28°-29°S) X Congreso Geológico Chileno Actas (CD ROM),

Concepción

Roperch, P., Sempere, T., Macedo, O., Arriagada, C., Fornari, M., Tapia, C., García, M., and Laj, C., 2006.

Counterclockwise rotation of late Eocene – Oligocene fore-arc deposits in southern Peru and its significance for oroclinal

bending in the central Andes, Tectonics, 25: TC3010, doi:10.1029/2005TC001882.

Somoza, R., Singer, S., and Tomlinson, A, 1999, Paleomagnetic study of upper Miocene rocks from northern Chile: Implications for the origin of late Miocene-Recent tectonic rotations in the southern Central Andes: Journal of Geophysical Research, v. 104, p. 22923-22936.

Taylor, G.K., Dashwood, B., Grocott, J., 2005, Central Andean rotation pattern: Evidence from paleomagnetic rotations of an anomalous domain in the forearc of northern Chile: Geology, v. 33, p. 777-780.

Taylor, G.K., Grocott, J., Daswood, B, Gipson, M., Arévalo, C., 2007, Implications for crustal rotation and tectonic evolution in the Central Andes forearc: New paleomagnetic results from the Copiapó region of northern Chile, 26º to 28ºS: Journal of Geophysical Research, v. 112, B01102, doi:10.1029/2005JB003950.

Figure 2: Simplified geological map of the Vallenar Discontinuity region, drapped over a Landsat image of the Main Andean Range. Location in Figure 1.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 357-360

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Tracing petrogenetic crustal and mantle processes in zircon crystals from rocks associated with the El Teniente porphyry Cu-Mo deposit in the high Andes of central Chile: Preliminary results

Marcia Muñoz1, Reynaldo Charrier

1, Víctor Maksaev

1, & Mark Fanning

2

1 Departamento de Geología, Universidad de Chile, Casilla 13518, Correo 21, Santiago, Chile

(marmunoz@cec.uchile.cl) 2 Research School of Earth Sciences, Australian National University, Canberra ACT 0200, Australia

KEYWORDS : El Teniente, porphyry Cu-Mo deposit, zircon, in-situ analyses, Hf-isotopes

Introduction

The origin of the igneous rocks associated with porphyry copper mineralization is a controversial issue. The

critical components of a mineralizing intrusive pulse (e.g., metals, water and sulphur contents) may be features

acquired either from the source of the magma or from its subsequent evolution during ascent throughout the

crust. Therefore, an accurate petrogenetic framework is essential

for understanding the evolution of mineralizing porphyry

magmas. The enormous porphyry Cu-Mo deposits of the high

Andes of Central Chile are not an exception. They are

characterized by multiple, superposed mineralization,

hydrothermal alteration and brecciation events, which have

greatly contributed to their extraordinary metallic concentrations,

but also obscured most of the primary textural, mineralogical,

and chemical characteristics of the ore-bearing intrusive rocks.

This fact complicates the application of traditional analytical

methods for the petrological study of these igneous rocks. We

have applied penetrative micro-analytical techniques on intrusive

rocks from El Teniente Cu-Mo deposit for getting insight into

their primary chemical characteristics, despite the consequences

of pervasive hydrothermal alteration that characterizes these

rocks. Preliminary results are presented here, which are part of

the Ph.D. thesis project of the first author developed under the

research framework of the Anillo ACT-18 project.

Geological background

The El Teniente porphyry Cu-Mo deposit is located in the high Andes of Central Chile (34°23’S-71°35’W) 70

km southeast of Santiago city (Fig. 1). It occurs within late Miocene extrusive and intrusive igneous rocks,

which are part of the Farellones Formation (e.g.: Skewes et al., 2005) and is the southernmost economic

porphyry deposit of the extensive Miocene to early Pliocene Cu belt of the Andes. The resources plus production

totals 94.35 Mt Cu, which makes El Teniente the largest copper deposit in the world (Camus, 2003).

Fig. 1: Main tectonic features of the Chilean continental margin. The location of Juan Fernandez Ridge, the volcanic gap separating the Central Volcanic Zone (CVZ) and the Southern Volcanic Zone (SVZ) are shown. Additionally, the location of the Los Pelambres, Río Blanco-Los Bronces and Teniente porphyry Cu-Mo deposits which are part of the Miocene to early Pliocene Cu belt of the Andes are indicated with black squares.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 357-360

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This supergiant Cu-Mo deposit is genetically related to

late Miocene-early Pliocene magmatic-hydrothermal

processes (e.g.: Howell & Molloy, 1960; Cuadra, 1986;

Skewes et al., 2005; Maksaev et al., 2004). Most of the

ore-bearing rocks correspond to altered basaltic and

andesitic host rocks and gabbroic sills that are currently

referred as El Teniente Mafic Complex, and a number of

felsic stocks and dikes (e.g.: Skewes et al., 2005;

Maksaev et al., 2004; Fig. 2). The whole lifespan of

igneous activity at El Teniente area can be roughly

traced from at least 12.0 ± 0.7 Ma, which is the oldest

K-Ar age obtained for the Farellones Formation (Kay et

al., 2005) and up to 3.85 ± 0.18 Ma which is the

hornblende 40Ar/39Ar age for a postore andesite dike

(Maksaev et al., 2004; Fig. 2); whereas, hydrothermal

mineralization developed from 6.4 to 4.3 Ma according

to combined U-Pb, 40Ar/39Ar, Re-Os and fission-track data (Maksaev et al., 2004).

Analytical techniques

In-situ geochronological, isotopic and chemical analyses on single zircon crystals

Zircon is a common accessory mineral phase of felsic igneous rocks

which concentrates significant amounts of trace elements, including

radiogenic elements, and is chemically and physically highly resistant

(e.g.: Watson, 1996). Its remarkable resistance to high temperature

diffusive re-equilibration allows it to preserve unaltered its primary

chemical signature and isotopic systems (e.g.: Watson, 1996; Watson

& Cherniak, 1997). These characteristics make of zircon crystals

sensitive tools for tracing petrogenetic crustal and mantle processes of

their host igneous rocks and are particularly well suited for rocks that

were subjected to one or more superposed hydrothermal alteration

events that modified other primary igneous characteristics of the

rocks.

We have applied a combination of SHRIMP (sensitive high resolution ion microprobe), and LA-ICP-MS (laser

ablation – inductively coupled plasma – mass spectrometry) analytical techniques for in-situ determination of

trace element (REE, Y, and Hf) and Ti contents, along with Hf and O isotopic signatures for individual zircon

crystals from igneous rocks of El Teniente Cu-Mo deposit (Fig. 3). These same zircon crystals have been

previously analyzed by back scattered electron–cathode luminescence (BSE-CL) and their U-Pb ages were

determined by SHRIMP analyses (Maksaev et al., 2004; Fig. 3). The samples correspond to five mineralized

intrusive bodies: the A porphyry, the Sewell stock, the northern and central quartz-diorite tonalite bodies, the

Fig. 2: Distribution of main lithologic units within the El Teniente Cu-Mo deposit at 4 LHD level (2,354 m). Location of samples analyzed in this study, which were previously dated by U-Pb SHRIMP (Maksaev et al., 2004), are indicated with white diamonds.

Fig. 3: Cathode luminescence (CL) and reflected light (RL) images for analyzed zircon crystals from the Dacite ring dike.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 357-360

359

Teniente Dacite Porphyry and a dacite ring dike (Fig. 2).

Preliminary results

All the analyzed zircons from the five studied intrusive bodies

show normalized REE concentration patterns typical of igneous

zircon. These are characterized by an enrichment of HREE

relative to LREE, with a steeply rise from LaN to LuN, and a

strong positive Ce-anomaly and a slight negative Eu-anomaly

(Fig. 4). The Eu-anomaly values are above 0.4 for zircons from

all felsic intrusive units, similar to the values reported for zircons

from intrusions associated with porphyry copper mineralization in

northern Chile (Ballard et al., 2002; Fig. 5). Additionally, the Ce-

anomaly appears to evolve towards higher values from the oldest

to the youngest intrusive unit of El Teniente (Fig. 5). This pattern could be associated with an increase of the

oxygen fugacity with time of the overall magmatic system; but a more detailed approach is still required to

further evaluate this hypothesis, similar to that applied by Ballard et al. (2002) taking in consideration the actual

valence state of Ce.

Fig. 5: Average Ce and Eu anomalies, and Tº per unit sampled. Numbers inside error ellipses correspond to: (1) A porphyry, (2) Central quartz-diorite tonalite, (3) Sewell stock, (4) Northern quartz-diorite tonalite, (5) Teniente Dacite Porphyry, (6) Dacite ring dike. Error ellipses are constructed over 1 error level from the mean in the y-axis and 2 error level from the U-Pb age obtained for each unit. Ce* = (LaN*PrN)1/2; Eu* = (SmN*GdN)1/2.

Geothermometry based on Ti concentration in zircon crystals (Watson et al., 2006) indicates a temperature of

the igneous solidus between 770º-580ºC (Fig. 5). The data scattering within individual units produces some

overlap in the temperatures obtained for each intrusion. However, there is a global decrease in this parameter

from the oldest unit represented by the A porphyry to the youngest ore-bearing dacite ring dike (Fig. 5),

consistent with progressive waning of igneous activity within the orebody.

All samples have high initial 176Hf/177Hf ratio and positive a Hf(i) whose values range from 6.2 to 8.5 and fall

between the silicate earth and depleted mantle (Fig. 6). There are no distinct differences among the analyzed

igneous units (Fig. 6). These characteristics, along with the almost null presence of inherited zircon or zircon

cores, are consistent with a common magmatic system originating the different intrusive pulses for which the

high Hf(i) values indicate a strong mantle signature with little or no interaction of the magmas with upper crustal

evolved rocks. Crustal residence time estimated through Hf isotopic compositions have a minimum limit of

Fig. 4: Chondrite normalized REE

concentration diagram for all analysis. Colored

area shows the field covered by all the

analyses, and the individual patterns obtained

from zircon crystals of the Teniente Porphyry

are shown in individual black lines as an

example.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 357-360

360

around 300 my (TDM) and a crustal model age of around 550 my (TDMc; average felsic continental crust

176Lu/177Hf = 0.015; Fig. 6). These characteristics of the Hf isotopic system point to the subcontinental

lithospheric mantle or the lower crust as the source of the intrusive pulses of the El Teniente deposit.

References Ballard, J., Palin, M., & Campbell, I., 2002. Relative oxidation states of magmas inferred from Ce(IV)/Ce(III) in zircon:

application to porphyry copper deposits of northern Chile. Contributions to Mineralogy and Petrology 144: 347-364. Camus, F., 2003. Geología de los Sistemas Porfíricos en los Andes de Chile. Servicio Nacional de Geología y Minería, 267 p. Cuadra, P., 1986. Geocronología K-Ar del yacimiento El teniente y áreas adyacentes. Revista Geológica de Chile 27: 22-54. Howell, F.H., & Molloy, J.S., 1960. Geology of the Braden orebody, Chile, South America. Economic Geology 55: 863-905. Kay, S.M., Godoy, E., & Kurtz, A., 2005. Episodic arc migration, crustal thickening, subduction erosion, and magmatism in

the south-central Andes. Geological Society of America Bulletin 117: 67-88. Maksaev, V., Munizaga, F., McWilliams, M., Fanning, M., Mathur, R., Ruíz, J., & Zentilli, M., 2004. New Chronology for

El Teniente, Chilean Andes, from U-Pb, 40Ar/39Ar, Re-Os, and Fission-Track Dating: Implications for the Evolution of a Supergiant Porphyry Cu-Mo Deposit. In Sillitoe, R.H., Perelló, J., Vidal, C.E. (eds): Andean Metallogeny: New Discoveries, Concepts and Updates, Society of Economic Geologists, SEG Special Publication 11: 15-54.

Skewes, A., Arévalo, A., Floody, R., Zuñiga, P., & Stern, Ch., 2005. The El Teniente Megabreccia Deposit, the world´s largest cooper deposit. In Porter, T.M. (ed): Super Porphyry Copper & Gold Deposits: A Global Perspective, PGC Publishing, Adelaide 1: 83-113.

Watson, E.B., 1996. Dissolution, growth and survival of zircons during crustal fusion: kinetic principles, geological models and implications for isotopic inheritance. In Transactions of Royal Society of Edinburgh: Earth Sciences 87: 43-56.

Watson, E.B., & Cherniak, D.J., 1997. Oxygen diffusion in zircon. Earth and Planetary Science Letters 148: 527-544. Watson, E.B., Wark, D.A., Thomas, J.B., 2006. Cristallization thermometers for zircon and rutile. Contributions to

Mineralogy and Petrology 148: 471-488.

Fig. 6: Left image shows the mean Hf(i) obtained for each unit versus age. Error ellipses and symbols as in Fig. 5. Right

image shows as a green circle the field defined by all initial 176

Hf/177

Hf ratios obtained in the analyzed samples versus

age and the models for the evolution of the CHUR (blue) and the depleted mantle (red).

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The brittle/ductile transition in the lithosphere of the Andes region and its relationship with seismogenesis

Miguel Muñoz

Jorge Matte 2005, Santiago, Chile (mmhGEO@hotmail.com)

KEYWORDS : heat flow, rheology, earthquakes, Wadati-Benioff zones, oceanic ridges

Introduction

Thermal and gravity data, together with rheological properties of rocks, are used to determine the brittle/ductile

(B/D) transition in the lithosphere along the Andes region by comparing a linear frictional fracture criterion

describing a brittle regime and a nonlinear flow stress corresponding to a ductile rheology (Ranalli, 1991). This

is an approach that involves basic physics and microphysics of the Earth’s interior, and hence it should be taken

as a sound understanding of a major geophysical process – the genesis of earthquakes in the continental

lithosphere. In this contribution the estimates of the B/D transitions in the lithosphere are compared with

seismological observations in terranes having very different heat flow.

Northern Andes

A background heat flow (Q) could be set between 60 and 80 mWm-2 for northeastern Venezuela (Hamza and

Muñoz, 1996). The southern area has values as low as 40 mWm-2, increasing to the northwest to about

50-60 mWm-2. In the area of the Oca-Ancon fault system in northern Venezuela, temperature in the crust/mantle

boundary (CMB) ranges between about 500 °C and 720 °C. In this area the thickness of the brittle crust should

not be larger than 14-18 km; for a dry olivine mantle there is indication of a thin brittle layer in the uppermost

mantle -from the CMB at 27 km to about 32 km depth. Maximum and minimum depths of seismicity associated

to this area according to seismological studies are of about 23 km and 11 km, respectively (Audemard & Singer,

1996), which means that in some zones the lower crust should be in the brittle regime. For acceptable thermal

and rheological model parametrizations, it is not possible to obtain a brittle layer of 23 km thickness. A seismic

event located close to this area at the northern termination of the Bocono fault system and at a depth of about

42 km (Dewey, 1972) with depth solution according to the International Seismological Center (ISC) was

relocated by Kafka & Weidner (1981) at 6 km depth.

For the southern termination of the Bocono fault system, where Q is of about 50 mWm-2, the temperature at the

CMB ranges from about 570 °C to about 600 °C and the crustal seismogenic layer is not larger than 20 km.

Some models describe a brittle layer beneath the CMB down to about 55-60 km depth, in accordance with the

analysis of seismic events of Kafka & Weidner (1981) that gives a depth solution of 55 km, and where the events

were interpreted to be within the continental lithosphere in a zone of brittle deformation. This seismic activity

does not appear to be related to the subducted ancient Farallon plate as proposed by Pennington (1981).

In the Central Cordillera of Colombia, and in the central northernmost area of Ecuador, an estimate of heat flow

is at least 70 mWm-2. The temperature at the CMB is close to 1000 °C. Only tectonic seismic activity of low

magnitude can be expected to be generated in the upper 13 km of the crust. High heat flow in Ecuador seems to

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 361-364

362

be characteristic only in small areas; low heat flow –between 40 and 60 mWm-2– can be expected in most of the

region (Hamza & Muñoz, 1996). In the sub-Andes of Ecuador, in the eastern slopes of the Cordillera, for a heat

flow of 50 mWm-2 and surface radiogenic heat generation Ao of 2.0 μWm-3, temperature at the CMB is about

450 °C. The seismogenic layer in the upper crust is of about 20 km thickness, and a thin brittle layer is also

described in the upper mantle. In northeastern Ecuador, in the eastern slopes of the Andes, where a relatively

higher heat flow density could be expected, Kawakatsu & Cadena (1991) adopting focal depths for several

earthquakes suggest a seismogenic zone of 15 km thickness. Suárez et al. (1983) reexamined the focal depth of

an earthquake in the high Andes of Ecuador relocating it at 16 km (the ISC determination was 33 km) –the

epicentre is close to the epicentres of shocks studied by Kawakatsu & Cadena (1992); other two earthquakes in

the sub-Andes were relocated at 16 km and 10 km depth.

For Ecuador –with Q = 40 - 45 mWm-2 and Ao = 1.0 - 1.2 μWm-3– the maximum thickness of the crustal

seismogenic layer ranges from about 20 to 35 km, and the mantle is seismogenic down to 50-86 km for multiple

thermal and rheological parameters. Studying the seismicity of northern central Ecuador, Guillier et al. (2001)

have used the global relocation of hypocentres carried out by Engdahl et al. (1998) on the ISC database and data

from a temporary network within the Ecuadorian network of telemetred stations. Earthquakes foci in the

continental mantle reaching depths of about 55-75 km were determined for both classes of data throughout the

analyzed sections, whereas dipping Wadati-Benioff zones could only be determined with foci located with the

temporary network. Beneath the broad volcanic region in northern Ecuador, Gutscher et al. (2000) inferred a flat

Wadati-Benioff zone (WBZ) at 80-100 km depth and related it to the subduction of the Carnegie ridge; Guillier

et al. (2001) show that the continuity of the WBZ –reaching a depth of at least 150 km in northern Ecuador and

200 km at 1-2 °S latitude– shows up only when using the locations of the temporary network.

Northern Central Andes

Seismogenic thickness has been calculated beneath different areas between latitudes 7 °S and 17 °S of the

Peruvian Andes with different crustal thickness (Muñoz, 2005). The brittle crustal seismogenic layer ranges

from about 20 km to 32 km; for some models with thick crust (60-65 km) the brittle regime reinitiates in the

lower crust after the transition into the ductile domain. The uppermost mantle, excepting zones with thin crust,

for most models is in the brittle regime from the CMB to about 70-110 km depth in the continental lithosphere.

Temperatures at the CMB range from about 300 °C to 620 °C, where lower values generally correspond to areas

with 30-50 km of crustal thickness. In northern central Peru, for explaining time residuals observed in

seismological stations, a velocity model for a flat subducting slab at 80-100 km depth was obtained using a

‘festooning’ ray with several reflections inside the slab (Norabuena et al., 1994); the model resulted in velocities

in excess for the flat structure that could not be explained by any effects, and that were found to be inconsistent

with studies of the thermal structure in subduction zones and with the mineralogy of slabs.

Southern Central Andes

In western central Argentina (Precordillera and Sierras Pampeanas) at about 32 ºS rheological modelling

indicates that the crust is in the brittle regime down to 33-45 km depth, in accordance with seismological

observations by Smalley & Isacks (1990) and Smalley et al. (1993), and in contrast with maximum focal depths

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 361-364

363

of 25 km encountered by Alvarado et al. (2005) in the western Sierras Pampeanas. Rheological results show that

maximum depths of seismogenic zones in the mantle could lie between 70 and 100 km (Muñoz 2005);

nonetheless, seismologists argue that foci at ~100 km depth represent a flattening of the subducting Nazca plate

linked to the subduction of the Juan Fernandez ridge (e.g., Smalley et al., 1993; Wagner et al., 2005; Anderson

et al., 2007). A comparison between ISC earthquake locations (with reporting stations located between about

30 ºS and 34ºS in Chile and western Argentina) and locations obtained by using the grid-search multiple-event

(GMEL) relocation algorithm for data of a more appropriate seismic network deployed in the area is presented

by Anderson et al. (2007). Whilst the ISC locations show two main branches of seismicity, one as a nearly flat

structure at ~100 km depth, and a second one as a WBZ dipping at an angle of ~30º, the GMEL locations show

only the branch at ~100 km depth for the same events (use Acrobat Reader to see the ISC locations at ~100 km

depth covered by the stars of the GMEL locations in Fig. 2B of Anderson et al., 2007).

In the main Cordillera of central Chile (33 ºS-35 ºS), the rheological zonation describes a brittle domain down to

about 17-22 km, in accordance with crustal seismic activity (Campos et al., 2002). In southern Chile, in the

south volcanic zone, only shallow tectonic seismic activity could be generated; most of the crust is ductile, and

the temperature at the CMB is of about 880-1100 ºC.

Discussion and conclusions

A large correspondence has been found between the rheological zonation of the crust and maximum depths of

seismic events there produced. The more brittle behaviour of the crust in western central Argentina at about 32ºS

is due to a colder crust impoverished of radiogenic elements, and this cannot be precluded by arguing that crustal

seismicity extends into the active arc south of this area where heat flow could be as high as 100 mWm-2

(Alvarado et al., 2007) –the thickness of the brittle layer at 33-35 ºS is half the thickness at 32 ºS, and when

geochemical estimates of heat flow in geothermal areas are not taken into account there is not such a high heat

flow (~100 mWm-2) in the southern central Andes. For regions of low heat flow in Ecuador, Peru and Argentina

the rheological zonation describes a brittle continental mantle down to 70-100 km depth. Nevertheless, the

activity at these depths has been taken as representing flat segments of the Wadati-Benioff zone, and the

flattening has been related to the cessation of volcanism in the advanced stages of flat-slab subduction. Thermal

modelling of subduction zones (English et al., 2003) shows that for flat-slab subduction magma generation at

distances > 600 km inboard from the trench is not predicted for warm subduction zones and is difficult at best

for cold subduction zones, and that most of the slab dehydration will take place within about 200 km of the

trench when the flat-slab segment occurs at a depth of 90 km or at larger depths. It is noted that late Miocene

Pocho lavas in western Argentina –where a flat-slab segment is believed to be of about 300 km length at 100 km

depth– erupted at a distance of about 750 km from the trench, in conflict with the thermal modelling. Also, in

these regions there is no obvious correlation between the subduction of oceanic ridges –a main subject for the

flat-slab hypothesis– and deformation processes and structural styles of the orogenic phases (Michaud et al.,

2006; Aleman, 2006; Creixell et al., 2006). Finally, the differences between ISC locations of events and GMEL

locations (Anderson et al., 2007) are outstanding at ~32 ºS: a structure that could be a dipping WBZ in the ISC

location disappears in the GMEL location, and less differences between both classes of locations are found just

below the easternmost area –where ISC locations should be more inaccurate– and used to argue about the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 361-364

364

bending and plunging of the ‘flat-slab’. Also, it is not clear why the GMEL locations do accord with classical

ISC locations determining the dipping WBZ at 34-36 ºS where the ISC locations should be at least as inaccurate

as to the north.

It is concluded that the depth of the brittle/ductile transition in the crust along the Andes region is highly

variable and depends largely on regional heat flow, and that brittle behaviour can be expected at ~70-100 km

depth in the continental lithosphere, above dipping Wadati-Benioff zones that could show in some segments an

annihilation of activity.

References Aleman, A.M., 2006. The Peruvian flat-slab. Backbone of the Americas—Patagonia to Alaska, 3–7 April 2006, Mendoza. Alvarado., P., Beck, S. & Zandt, G., 2007. Crustal structure of the south-central Andes Cordillera and backarc region from

regional waveform modelling. Geophys. J. Int. 170: 858-875. Anderson, M., Alvarado, P., Zandt, G. & Beck, S., 2007. Geometry and brittle deformation of the subducting Nazca Plate,

central Chile and Argentina. Geophys. J. Int. 171: 419-434. Audemard, F.A. & Singer, A., 1996. Active fault recognition in northwestern Venezuela and its seismogenic

characterization: Neotectonic and paleoseismic approach. Geof. Int. Méx. 35 (3) (Suppl.): 245–255. Campos, J., Hatzfeld, D., Madariaga, R., López, G., Kausel, E., Zollo, A., Iannacone, G., Fromm, R., Barrientos, S. & Lyon-

Caen, H., 2002. A seismological study of the 1835 seismic gap in south-central Chile. Phys. Earth Planet. Inter. 132: 177-195.

Creixell, C.E., Arriagada, C., Morata, D., Parada, M.A., 2006 — The role of the Juan Fernandez ridge in the tectonic evolution of the central Chilean Andes. Backbone of the Americas—Patagonia to Alaska, 3–7 April 2006, Mendoza.

Dewey, J.W., 1972. Seismicity and tectonics of western Venezuela. Bull. Seism. Soc. Am. 62: 1711-1751. Engdahl, E.R., van der Hilst, R. & Buland, R., 1998. Global teleseismic earthquake relocation with improved travel times and

procedures for depth determination. Bull. Seism. Soc. Am. 88: 722-743. English, J.M., Johnston, S.T. & Wang, K., 2003. Thermal modelling of the Laramide orogeny: testing the flat-slab

subduction hypothesis. Earth Planet. Sci. Lett. 214: 619-632. Guillier, B., Chatelain, J.-L., Jaillard, E., Yepes, H., Poupinet, G. & Fels, J.-F., 2001. Seismological evidence on the

geometry of the orogenic system in central-northern Ecuador (South America).Geophys. Res. Lett. 28: 3749-3752. Gutscher, M.-A., Spakman, W., Bijwaard, H. & Engdahl, E.R., 2000. Geodynamics of flat subduction: seismicity and

tomographic constraints from the Andean margin. Tectonics 19: 814-833. Hamza, V.M. & Muñoz, M. 1996. Heat flow map of South America. Geothermics 25: 599-646. Kafka, A.L. & Weidner, D.J., 1981. Earthquake focal mechanisms and tectonic processes along the southern boundary of the

Caribbean plate. J. Geophys. Res. 86: 2877-2888. Kawakatsu, H. & Proaño-Cadena, G., 1991. Focal mechanisms of the march 6, 1987 Ecuador earthquakes – CMT inversion

with a first motion constraint. J. Phys. Earth 39: 589-597. Michaud, F., Witt, C., Bourgois, J., Bustillos, J., Peñafiel, L., 2006 — Influence of the subduction of the Carnegie ridge on

Ecuadorian geology: reality or fiction? Backbone of the Americas—Patagonia to Alaska, 3–7 April 2006, Mendoza. Muñoz, M., 2005. No flat Wadati-Benioff zone in the central and southern central Andes. Tectonophysics 395: 41-65. Norabuena, E.O., Snoke, J.A. & James, D.E., 1994. Structure of the subducting Nazca plate beneath Peru. J. Geophys. Res. 99: 9215-9226.

Pennington, W. D., 1981. Subduction of the Eastern Panama basin and seismotectonics of Northwestern South America. J. Geophys. Res. 86: 10753-10770.

Ranalli, G., 1991 — "Regional variations in lithosphere rheology from heat flow observations". In Cermak, V., Rybach, L. (eds.): Terrestrial Heat Flow and the Lithosphere Structure, Berlin, Springer: 1-22.

Smalley, R., Jr. & Isacks, B.L., 1990. Seismotectonics of thin- and thick-skinned deformation in the Andean foreland from local network data: Evidence for a seismogenic lower crust. J. Geophys. Res. 95: 12487-12498.

Smalley, R., Jr., Pujol, R., Regnier, M., Chiu, J.-M., Chatelain, J.-L., Isacks, B.L., Araujo, M. & Puebla, N., 1993. Basement seismicity beneath the Andean Precordillera thin-skinned thrust belt and implications for crustal and lithospheric behavior. Tectonics 12: 63-76.

Suárez, G., Molnar, P. & Burchfiel, B.C., 1983. Seismicity, fault plane solutions, depth of faulting, and active tectonics of the Andes of Peru, Ecuador and southern Colombia. J. Geophys. Res. 88: 10403-10428.

Wagner, L.S., Beck, S. & Zandt, G., 2005. Upper mantle structure in the south central Chilean subduction zone (30º to 36ºS), J. Geophys. Res. 110: B01308, doi:10.1029/2004JB003238, 2005.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 365-368

365

Nature of a topographic height in the Tarapacá pediplain, Northern Chile

Violeta Muñoz1, Gérard Hérail

2, & Marcelo Farías

3

1 SERNAGEOMIN, Santa María 0104, Providencia, Santiago de Chile, Chile (vimunoz@sernageomin.cl)

2 LMTG, IRD, CNRS, Université de Toulouse, 14 Avenue Edouard Belin, Toulouse 31400, France

3 Dpto. de Geología, Universidad de Chile, Plaza Ercilla 805, Santiago, Chile

KEYWORDS : Choja peneplain, Tarapacá pediplain, Incaic relief, North Chile forearc, Neogene deformation

Introduction

In the Tarapacá Region (N-

Chile), the Altiplano Western

Flank consists in 4 morpho-

structural units parallel to the

orogen (Fig. 1a). The Precor-

dillera, which concentrated most

of the activity in the Oligo-

Neogene Andean deformation, is

characterized by its flat shape

(Tarapacá Pediplain). This

morphology was formed since the

Oligocene as a consequence of

strong aggradation of sediments

and ignimbrites coeval to the

activity of the West Vergent Thrust

System (WTS; Muñoz and

Charrier, 1996). These deposits

cover a regional erosional surface

(Choja Peneplain) related to an

intense pre-Oligocene exhumation

and denudation episode (Galli, 1967).

The study region is situated in the western limit of the Precordillera at 19º50’S-20º30’S, where an isolated and

NS trending topographic high (~350 m above the regional elevation) stands out in the Tarapacá Peneplain (Cerro

Violeta Range, CVR; Fig. 1b). In this contribution, we analyze this topographic anomaly based on the

morphology, structure, and stratigraphy of the deposits and landforms, proposing that the CVR would be a result

of a particular morphotectonic evolution, in which the pre-Neogene evolution controlled the locus of lithological

units that (1) resisted more strongly the erosion respect the surrounding units, as occurred south of 21ºS

(Domeyko Cordillera) and (2) concentrated the Oligo-Neogene deformation in the western domain of the WTS,

continuing the tectonic framework recorded to the north, but displacing the thrusting front ~20 km westward.

Figure 1. Morpho-structural framework of Tarapacá Region and location of the studied area. A. Longitudinal units are indicated in black, while red structures configure the WTS. Black structures record poorly to none Neogene reactivation. Holocene volcanoes are indicated in green triangles B. ASTER image in 3D view of the studied area, showing the CVR, an isolated and NS trending topographic high surrounded by the Miocene Tarapacá Pediplain

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 365-368

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Geological particularities of pre-Oligocene substratum

In the study region, the pre-Oligocene substratum consists of Paleozoic and Mesozoic stratified sequence

intruded by late Cretaceous-Eocene plutons submitted to an intense and polyphasic deformation, accommodating

20-30 km of shortening before the Oligocene [Harambour, 1990]. Two major features stand out in the study

region respect to the regional pre-Oligocene geological framework: (1) an unusual concentration of intrusive

outcrops and (2) the lithology of Paleozoic and Mesozoic pelitic units (the Permian Juan de Morales Fm. and the

Jurassic El Tranque Fm.).

The Mesozoic-Late Cenozoic unconformity: The Choja peneplain

The Mesozoic-late Cenozoic contact is a regional erosive flat unconformity (Choja Peneplain; Galli, 1967).

The origin of this pediplain correlates to the latest Eocene-earliest Oligocene exhumation recorded by supergene

enrichment at the Cerro Colorado porphyry copper deposit [Bouzari and Clark, 2002] and fission-track data in

the Antofagasta region [Maksaev and Zentilli, 1999]; exhumation has been interpreted as a result of shortening

and uplift related to the Incaic phase [Harambour, 1990; Haschke and Günther; 2003; Tomlinson et al., 1999].

However, in the study region, the pediplain is disrupted by isolated hills and belts. Moreover, east of the CVR,

the presence of paleovalleys and paleochannels (as deep as 300 m) characterizes the Mesozoic-Cenozoic

unconformity. There is no data about the age of this relief growth in this region. However, the age of the main

supergene enrichment at the Cerro Colorado porphyry copper deposits (lowest Oligocene) probably mark the

beginning of local incision, which is older than the Oligo-Neogene deformation. Therefore, it is likely that this

incision was related to the erosion post-Incaic deformation (similar to pedimentation), which in this region was

not capable to produce pedimentation as occurs to the north and to the south.

The topographic anomalies are related to intrusive bodies, mostly in their western flank (Fig. 2). Therefore, this

suggests that during the Choja pedimentation, this lithology inhibited and delayed upward propagation of

Figure 2. Geological map of the studied region. Red lines indicate position of sections in fig. 4

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 365-368

367

erosion. In fact, intrusive bodies can resist >1 order of

magnitude the erosion than other rocks (e.g., Stock and

Montgomery, 1999). Hence, previous geologic evolution

determined the locus of inselbergs and anomalous heights in

the Choja Peneplain as also occurs in the Domeyko

Cordillera (Fig. 3).

The Oligo-Miocene synorogenic sequence

The Cenozoic cover consists of clastic deposits interbedded

with 22.71-19.25 Ma ignimbrites (Altos de Pica Fm) overlaid

by middle to upper Miocene conglomerates (El Diablo Fm.)

(Fig. 2). East of CVR, its thickness is <500 m, much minor

than that observed to the north and south (>1 km; Pinto et al.,

2004; Victor et al., 2004). In addition, intraformational unconformities within this sequence (paleovalleys, that

in some cases contact the lower 22.71 Ma ignimbrite with El Diablo Formation, Fig. 4B) evidence minor

aggradation and more erosion in this region, thus suggesting a major relative uplift in the study region.

The sequence presents syntectonic deposition before 13.7 Ma immediately west of the CVR (Fig. 4); It is

related to a W-vergent thrust with <500 m throw and 30-8ºE tilting of the hanging wall. To the east, there is no

significant deformation affecting these deposits, thus this fault accommodated a similar relative uplift that north

and south of this region (~2 km; Victor et al., 2004; Farías et al., 2005; however, there was accommodated by

several structures). In turn, deformation in the study region differs from that observed along the Precordillera

because the main fault is located ~20 km westward.

Figure 4. Synsedymentary deformation in the Oligo-Miocene sequence. A.

Progressive unconformities in the Pachica and Tarapacá localities (see

position in fig.2). B. View of Altos de Pica Fm. tilted to the east forming

growing strata (30-8ºE).

A.

B.

Figure 3. Late

Cretaceous to

Eocene outcrop of north Chile. Note

the absent of plutonic outcrop in

most part of the

Tarapacá Region

and there

concentration in

the studied area

(square), as well as

in the Domeyko

Cordillera

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 365-368

368

Discussion

Structural and morphological data show that the study region presents a late Cenozoic uplift similar to that

observed along the Precordillera between 18-21ºS, even though the CVR would be in part a remain of the Incaic

orogen.

The fact that deformation is concentrated in only one fault located 20 km westward the main faults presented

elsewhere suggests that not only the substratum resisted more the erosion than in others place, but also there are

particularities on the faults prolongation at depths. We suspect that the pelitic constitution in the substratum

would produce detachment layers capable of accommodate major shortening (and thus uplift) during and before

the Oligocene-Miocene tectonic event: exhumation of intrusive bodies during the Incaic phase and concentration

of deformation in only one fault.

Alternatively, there are two more features in this zone within the regional context (Fig 1): (1) a change in the

general trend in WTS structures (NNW to the north and NS to the south) and (2) its position in the north extreme

of Volcanic Pica Gap (Wörner et al., 2000). This suggests that this zone also represents a major scale tectonic

transition zone. Therefore, the study region would correspond to a local tectonic, lithological, and morphological

anomaly within the Northern Chile forearc evolution since the Eocene, which would be ultimately determined by

a long-term regional tectonic transition zone.

References Bouzari, F. and Clark, A. 2002. Anatomy, evolution, and metallogenic significance of the supergene orebody of the Cerro

Colorado porphyry copper deposit, I Región, Northern Chile. Economic Geology 97: 1701-1740.

Farías, M., R. Charrier, D. Comte, J. Martinod y G. Hérail. 2005. Late Cenozoic deformation and uplift of the western flank

of the Altiplano: Evidence from the depositional, tectonic, and geomorphologic evolution and shallow seismic activity

(northern Chile at 19º30’S). Tectonics 24 (4): TC4001.

Galli, C. 1967. Pediplain in northern Chile and the Andean uplift. Science 158: 653 – 655.

Harambour, S. 1990. Geología pre-cenozoica de la Cordillera de los Andes entre las quebradas Aroma y Juan de Morales, I

Región. Memoria de Título. Departamento de Geología, Universidad de Chile, Santiago. 228 p.

Haschke, M. and Günther. 2003. Balancing Crustal Thickening in arcs by tectonic vs magmatic means. Geology 31: 933-936

Maksaev, V. and M. Zentilli. 1999. Fission track thermochronology of the Domeyko Cordillera, northern Chile: Implications

for Andean tectonics and porphyry copper metallogenesis. Explor. Min. Geol. 8: 65 – 89.

Muñoz, N. and R. Charrier. 1996. Uplift of the western border of the Altiplano on a west-vergent thrust system, northern

Chile. J. S. Am. Earth Sci. 9: 171 – 181.

Pinto, L., G. Hérail, R. Charrier. 2004. Sedimentación sintectónica asociada a las estructuras neógenas en el borde occidental

del plateau andino en la zona de Moquella (19º15’S, Norte de Chile). Revista Geológica de Chile.31 (1), 19-44.

Riquelme, R., G. Hérail, J. Martinod, R. Charrier and J. Darrozes (2007). "Late Cenozoic geomorphologic signal of Andean

forearc and tilting associated with the uplift and climate changes of the Southern Atacama Desert." Geomorphology 86(3-

4): 283-306.

Stock, J. D. and D. R. Montgomery (1999). "Geologic constraints on bedrock river incision using the stream power law."

Journal of Geophysical Research-Solid Earth 104(B3): 4983-4993

Tomlinson, A., Cornejo, P. and Mpodozis, C. 1999. Hoja Potrerillos, Región de Atacama. Servicio Nacional de Geología y

Minería. Mapas geológicos 14. 1 Map, Santiago

Victor, P., Oncken, O. y Glodny, J. 2004. Uplift of the western Altiplano plateau: Evidence from the Precordillera between

20º and 21ºS (northern Chile). Tectonics 23: TC4004

Wörner, G., Hammerschmidt, K., Henjes-Kunst, F., Lezaun, J. y Wilke, H. 2000. Geochronology (40 Ar/ 39AR, K-Ar and

He-exposure ages) of Cenozoic magmatic rocks from Northern Chile (18°-22°S): implications for magmatism and tectonic

evolution of central Andes. Revista Geológica de Chile, 27, 2: 205-240.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 369-372

369

Stratigraphy of the synorogenic Cenozoic volcanic rocks of Cajamarca and Santiago de Chuco, northern Peru

Pedro Navarro, Cristina Cereceda, & Marco Rivera

Instituto Geológico Minero Metalúrgico (INGEMMET), Av. Canada 1470, Lima 41, Peru

(pnavarro@ingemmet.gob.pe, ccereceda@ingemmet.gob.pe, mrivera@ingemmet.gob.pe)

KEYWORDS: volcanism, stratigraphy, volcanic zones, migration, Cenozoic

Introduction

Cenozoic volcanic deposits that cap the Western Cordillera of Northern Peru originally were mapped as one

unit called “Volcánicos Calipuy” by Cossío (1964). Later studies (Reyes, 1980; Cobbing, 1981; Wilson, 1984)

used the name Calipuy Group including within it several volcanic sequences with local informal names as

“Volcánicos Llama, Porculla, Huambos, Chilete and Tembladera”. Nevertheless, Hollister & Sirvas (1978) were

the first in establish an association between those volcanic deposits and processes like volcanoes growth and

explosive activity.

Facies interpretation, updated mapping and volcanic stratigraphy allow the recognition of three segments

(Figure 1) constituted by volcanic rocks of Cenozoic age in northern Peru: Huancabamba (4º - 6º S), Cajamarca

(6º - 7º 30’ S) and Santiago de Chuco (7º 30’ – 9º S), all them emplaced between Eocene and upper Miocene

from 54.8 ± 1.8 to 8.2 ± 0.2 Ma (Noble et al., 1990; Turner, 1997; Davies, 2002; Noble & Loayza, 2004; Noble

et al., 2004; Longo, 2005; Rivera et al., 2005). So that, these volcanic rocks overlies unconformably carbonated

and clastic sequences from Mesozoic time.

The volcanic activity was characterized by intense and continuous explosive and effusive phases that emplaced

thick pyroclastic and lava flow deposits corresponding to five volcanic stages from early Eocene, upper Eocene,

Oligocene, early Miocene to upper Miocene, suggesting a continuity in the emplacement and eastward migration

of the magmatic arc during approximately 46 Ma.

Mainly the volcanic centres show a trend NW-SE, similar with regional faults and trending fractures.

This paper shows geological cartography, regional structural setting and stratigraphy results carried out only in

Santiago de Chuco and Cajamarca segments.

Regional geology

Early Paleozoic polydeformed rocks consist in shales, schists and volcanic deposits; and crop out east of the

Cenozoic volcanic rocks. They are overlain a volcanic and sedimentary sequence of Late Paleozoic age

represented by lava, conglomerate and sediments, related to a Permian-Triassic rift. A subsidence during Late

Triassic-Jurassic, with the development of a carbonate basin is documented by the deposition of limestones,

marls and shales; interbedded some with volcanic sills. Early Cretaceous time is marked by a fill of the basin

with sandstone, shales and minor limestones. Late Cretaceous is characterized by a new subsidence and

establishment of a carbonated shelf. The Cenozoic volcanic rocks lies upon a Paleozoic-Mesozoic basement,

finally structured by 40 Ma.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 369-372

370

The region was affected by the Peruvian and Incaic tectonic events that generated folds, faults and uplift.

Erosion of the created relief and further deposition interbedded of conglomerate and typical red mudstones

during Late Cretaceous and Early Paleogene time. The Eocene to Miocene volcanic rocks were emplaced during

the tectorogenesis that developed the Western Cordillera of Northern Peru.

All sequences were affected by distinct tectonic events, called Incaic and Quechua.

Figure 1. Location map showing Volcanic zones and study area.

Volcanic zones

Cajamarca zone

Volcanic deposits located between 6º - 7º 30’ S were generated from a continuous magmatic arc activity

developed since Early Eocene to Upper Miocene (54.8 ± 1.8 - 8.2 ± 0.2 Ma) showing five eruptive periods

(Figure 2): Early Eocene (55 – 43 Ma), Upper Eocene (43 – 33 Ma), Oligocene (33 – 24 Ma), Early Miocene (24

– 14 Ma) and Upper Miocene (14 – 8 Ma). They were separated by volcanic gaps, development of small

synorogenic basins, and erosional and angular unconformities

Geological mapping and stratigraphical studies suggest an intense and intermittent explosive and effusive

volcanic activity that built at least thirteen volcanic centres (e.g. stratovolcanoes and volcanic complexes):

Yatahual, San Lorenzo, La Colmena, Niepos, Santa Cruz, Anchipan-Mutis, Chuño-Chinchín, Huayquisongo,

Chicche-Hueco Grande, Rumiorcco, San Pedro, Tantahuatay and Yanacocha. Also, there are older volcanic

sequences: Chancay, Chilete-Ayambla and Tantachual, whose eruptive centres are overlain by younger volcanic

deposits. In addition, rhyolitic to dacitic welded crystal-rich ash-flow tuffs probably were emitted from the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 369-372

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recently described Catan Caldera (Navarro et al, 2007).

Stratigraphy puts in evidence the eastward episodic migration of magmatic arc during Cenozoic time in

northern Peru.

Figure 2. Spatial and temporal evolution of the Cenozoic volcanism in Cajamarca zone.

Santiago de Chuco zone

Between 7º 30’ - 8º 30’ S, a thick volcanic sequence is recognized overlying Cretaceous and Paleocene

sedimentary rocks. These rocks are thought to be erupted during four volcanic periods. The first period began

probably in the Eocene, because volcanic deposits were intruded by subvolcanic bodies that yielded an age of

35 Ma. Next event ocurred in Oligocene time, ages from 34 to 24 Ma and angular unconformity point out this

eruptive activity. The volcanism continued until the Miocene, lava and pyroclastic flows aged from 18 to 16 Ma

characterize stages three and four (i.e. Upper Oligocene-Early Miocene and Early Miocene). Each event is

separated by erosional and angular unconformities and volcanic gaps. Stratigraphy and chronology put in

evidence the migration of continuous magmatic arc during Cenozoic time from West to East (Figure 3).

These periods generated at least thirteen volcanic centres (stratovolcanoes): Ultocruz-Ticas, Macón, Matala,

Alto Dorado, San Pedro, Cururupa, Quiruvilca, Paccha-Uromalqui, Totora, Quesquenda, Payhual, Urpillao-

Rushos and Piedra Grande. Besides, an older volcanic sequence called Tablachaca whose source vent is

unrecognized, crops out along the northern flank of the Tablachaca river. Also there are andesitic to rhyolitic

thick ash-flow tuffs erupted from two calderas: Carabamba and Calamarca.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 369-372

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Figure 3. Spatial and temporal evolution of the Cenozoic volcanism in Santiago de Chuco zone.

References Cossío, A. 1964. Geología de los Cuadrángulos de Santiago de Chuco y Santa Rosa. Lima, INGEMMET, 8 (A), 69 p. Cobbing, E. 1981. Estudio geológico de la Cordillera Occidental del norte del Perú. Lima, INGEMMET, 10 (D), 260 p. Davies, C. 2002. Tectonic, magmatic and metallogenic evolution of the Cajamarca mining district, Northern Peru. Ph. D.

Thesis, James Cook University, Australia, 323 p. Hollister, V., Sirvas, E. 1978. The Calipuy Formation of northern Peru and its relation to volcanism in the northern Andes.

Journal of Volcanology and Geothermal Research 4: 89-98. Longo, A. 2005. Evolution of volcanism and hydrothermal activity in the Yanacocha mining district, Northern Perú. Ph.D.

Thesis, Oregon State University, U.S.A., 469 p. Navarro, P., Monge, R., Flores, A. 2007. “Informe Geocientífico: Avances del Año 2006 - Proyecto de Investigación GR4:

Volcanismo Cenozoico (Grupo Calipuy) y su asociación con los yacimientos epitermales, Norte del Perú”. INGEMMET, Reporte interno, 50 p.

Noble, D., Loayza, C. 2004. “Field trip: Volcanic rocks and paleogene geological history in the vicinity of Chilete. Guía de campo”. In: XII Congreso Peruano de Geologia, Lima, 12 p.

Noble, D., McKee, E., Mourier, T., & Mégard, F. 1990. Cenozoic stratigraphy, magmatic activity compressive deformation, and uplift in Northern Peru. Geological Society of America Bulletin 102: 1105-1113.

Noble, D., Vidal, C., Perelló, J., & Rodríguez, O. 2004. Space-time relationships of some porphyry Cu-Au, epithermal Au, and other magmatic-related mineral deposits in northern Perú. Society of Economic Geologists Special Publication 11: 313-318.

Reyes, L. 1980. Geología de los cuadrángulos de Cajamarca, San Marcos y Cajabamba. Lima, INGEMMET, 31 (A), 67 p. Rivera, M., Monge, R., & Navarro, P. 2005. Nuevos datos sobre el Volcanismo Cenozoico (Grupo Calipuy) en el Norte del

Perú: Departamentos de La Libertad y Ancash. Boletín Sociedad Geologica del Perú 99: 7-21. Turner, S. 1997. The Yanacocha epithermal gold deposits, northern Peru: high sulfidation mineralization in a flow–dome

setting. Ph.D. thesis, Colorado School of Mines, U.S.A., 341 p. Wilson, J. 1984. Geología de los cuadrángulos de Jayanca, Incahuasi, Cutervo, Chongoyape, Chota, Celendín, Pacasmayo y

Chepén. Lima, INGEMMET, 38 (A), 104 p.

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Characterization of the Sierras de Córdoba eastern boundary from gravimetry, magnetotelluric and DEM (Argentina)

Luz A. Orozco1, Eduardo A. Rossello

2, Cristina Pomposiello

1, Alicia Favetto

1, & Cristóbal P.

Bordarampé

1 Instituto de Geocronología y Geología Isotópica, Pabellón INGEIS, Universidad de Buenos Aires, Ciudad

Universitaria, C1428EHA, Buenos Aires, Argentina 2

Departamento de Ciencias Geológicas, Universidad de Buenos Aires, CONICET. Ciudad Universitaria,

C1428EHA, Buenos Aires, Argentina

KEYWORDS : Sierras Pampeanas, gravimetry, magnetotelluric, morphostructural

Introduction

The Sierras de Cordoba constitutes a part of the faulted, rotated, tilted and peneplained blocks of the Sierras

Pampeanas located between 29°S and 33° 30´S and between 64°W and 66°W, and represents the easternmost

expresion of the Andean deformation on its foreland. The Andean deformation inverted by compresion the pre-

existing tectonic discontinuities particularly the listric growth extentional faulting bounding the Cretaceous rifts

towards the west (Cobbold et al., 1996).

The Sierras de Córdoba is the easternmost orographic feature of the Sierras Pampeanas. It is constituted by

several NS trending belts extending being the most important ones the Sierra Norte, the Sierra Chica and the

Sierra Grande. The studied area includes part of the Sierra Chica and the westernmost portion of the Llanura

Chacopampeana (Figure 1a).

In the Sierra Chica, the exposed basement is composed by a metamorphic-migmatitic complex, where the

prevailing rocks are tonalitic – biotitic gneisses, locally alternating with micaceous schists and migmatites. The

pleneplained top of the range plunges towards the east underneath the adjacent plain (Llanura Chacopampeana),

which was reached at the Santiago Temple well (Figure 1b). Here an olivinic metagabbro was found at a depth

of 997 meters deep, with a reported K/Ar data of 787+/-150 Ma (Russo et al., 1987).

The Chacopampean plain (covering more than 1,000,000 km2 of central Argentina) lacking any surficial

feature representing the tectonic activity in such huge extension, has on the other side a rich story of

underground tectonic events. These events, even when they took place at different geological times and are

partially recognized by only a few oil wells and seismic records in diverse places having similar geometries,

corresponding to other major geological structures of the Argentine geology (Chebli et al., 1999).

In this paper we present morphostructural results obtained from both gravimetric and magnetotelluric surveys

associated to digital elevation model (DEM) allowing to establish relationships between the Sierras de Córdoba

and its eastern sedimentary cover.

Gravimetric results

The Bouguer gravimetric map (Figure 1a and 1b) including six surveyed lines shows the top of the crystalline

basement and its onlaping sedimentary cover dipping gentle to the east. In the northern part, the isolines show a

regular interface surface that dips to the east and it corresponds to the discordance between the basement and the

sedimentary cover. Near Córdoba city the gravimetric contours suffer a deflection indicating a change in the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 373-376

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basement top geometry. To the south, this change is more significant showing the presence of the gravimetric

high anomaly that corresponds to an uplifted basement.

Magnetotelluric results

In Figure 1c the MT section across the eastern boundary of the Sierras de Córdoba and the Chacoparanense

sedimentary basin model is presented. It shows a huge area with resistivities greater than 200 ohm m, which is

associated to the crystalline basement that outcrops at the Sierras de Córdoba and plunges under the Llanura

Chacopampeana to the east. The sedimentary sequences show layers with resistivities between 1 and 100 ohm

m, where are recognized different lithologic units by comparison to the information provided from wells drilled

by oil companies (Favetto et al., 2004).

From 755 sp up to 780 sp the 2D MT model exhibits a thin tertiary sedimentary records (Sierra Chica Basin,

from Ramos, 1999) on the uplifted and peneplained surfuce worked on the top of the basement. The Sierra Chica

basin is bounded by two vergence towards west reverse faults which control the foothill morphology (among

other the La Calera fault), which extends approximately N-S along all the studied area.

The MT model (with a 4 times vertically increased scale), shows the top of the crystalline basement with a

major plunging gentle to the east. Nevertheless, between the 755 sp and 790 sp this regional trend is modified by

an uplifted basement block. Also, the thickness of the sedimets overlying this block is around 500 meters, while

a few kilometers to the east, the basement was found at a depth of 997 meters deep (Santiago Temple well). It

indicates that the basement – sedimentary sequence discontinuity has been tilted increasing its angle due to the

rotation of the block (Figure 1c) .

DEM results

From Córdoba city to the south we observed that the alluvial quaternary deposit levels are more deformed. The

Córdoba province geological map (Figure 2a), shows two NS trending structures that converge toward the south.

Between these structures it is possible to observe a great doming of the quaternary deposits expresed by surficial

irregular texture, the antecedent dranaige pattern of the Primero and Segundo rivers, and the confining of the

modern sediments west of the present depositation line (Figure 2b).

Vertically exaggerated topographic sections across these deformed deposits show a major doming limited by

faults with decreasing of its displacements toward the north (Figure 2c).

Discussion

Geophysical results together with morphotectonic observations from DEM show a unconformity surface which

overlaps independently reactivated basement blocks. Early Paleozoic basement blocks that were peneplained

during the late Paleozoic were uplifted and differentially rotated to the east during the Neogene through the

reactivation of still active Cretaceous extensional faults.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 373-376

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Figure 1. (a) Bouguer gravimetric anomalies Map, MT sites ( ), gravimetric lines (+), Santiago Temple well (STW), Sierra Grande (SG), Sierra Chica (SC); (b). Gravimetric model corresponding to Line 5 and (c) 2D MT model.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 373-376

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Figure 2. (a) Geological map of Córdoba Province; (b) Digital elevation model, Antecedent Drainage (AD), Deformed Quaternary (DQ) and (c) Quaternary topographic sections (vertically exaggerated). Acknowledgements This project is supported by the UBACYT (Projet N° Ex-272), the ANPCYT (PICT 2005 Nº 38253) and the NSF (Grants EAR0310113 and EAR0739116). MT equipment is from the EMSOC Facility supported by NSF Grant EAR026538. This project also received support from References Chebli, G.A., Mozetic, M.E., Rossello, E.A., & Buhler, M., 1999. Cuencas sedimentarias de la llanura Chacopampeana. In:

Caminos, R. (Ed): Geología Argentina. Instituto de Geología y Recursos Minerales, Buenos Aires, Anales 29 (20): 627-644.

Cobbold, P.R., Szatmari, P., Lima, C., & Rossello, E.A., 1996. Cenozoic Deformation Across South America: Continent-wide data and Analogue Models. III° International Symposium on Andean Geodynamics, Orstom-Géosciences Rennes (Saint Maló, Francia), 21-24.

Favetto A., Pomposiello, M.C., Benedit, T., & Booker, J., 2004 -“Magnetotelluric model of the Chacoparanense sedimentary basin at 31.5 S Argentina”. Proceedings of the 17th Workshop Electromagetic Induction in the Earth, Available at (http://www.emindia2004.org).

Michaut, L., & Gamkosián, A. 1995, Mapa Geológico de la Provincia de Córdoba. Escala 1:500.000. Servicio Geológico Argentina.

Ramos, V.A., 1999. Rasgos estructurales del territorio argentino. 1. Evolución tectónica de la Argentina. In: Caminos, R. (Ed): Geología Argentina. Instituto de Geología y Recursos Minerales, Buenos Aires, Anales 29 (24): 715-784.

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Crustal seismicity and 3D seismic wave velocity models in the Andes cordillera of Central Chile (33°-34.5°S) from local earthquakes

M. Pardo1, E. Vera

1, T. Monfret

2 , & G. Yañez

3

1 Departamento de Geofísica, Universidad de Chile, Santiago, Chile

2 Géosciences Azur, Université de Nice, IRD, Sophia-Antipolis, Valbonne, France

3 Codelco, Teatinos 258, piso 8, Santiago, Chile

KEYWORDS : crustal seismicity, tomography, seismic hazard

Introduction

Central Chile is located in a transition zone where Nazca plate changes from a flat slab subduction north of

33°S to a “normal” subduction south of this latitude, with dip angle of about 30°E. The seismogenic zones

related with these two segments are well known along the downgoing slab: great and large shallow thrust

interplate events along the coast (0-50 km depth), and large deeper (60-200 km) tensional as well compressive

events in the subducted Nazca plate. However, the crustal shallow seismicity within the continental plate is

poorly understood, both in genesis and rates of activity.

The crustal seismicity in central Chile occurs mainly at the Andes cordillera and foothills, and it is related to

the actual Andes deformation and uplift, generated by the Nazca and Southamerican plate interaction. The

different subduction styles control the crustal seismicity; this seismicity is low in the fore-arc and high at the

back-arc at the flat slab zone, while abundant seismicity is observed at the fore-arc and low at the back-arc at the

southern steep subduction zone.

In our studied Andean region (33°-34.5°S), the crustal seismicity is concentrated mainly along the western

foothills and in the Chilean side of the Principal Cordillera. The largest earthquakes reported in the region took

place in September 4, 1958 (M=6.9, Lomnitz, 1961), and September 13, 1987 (M=5.9, Barrientos and Eisenberg,

1988). The 1958 event damaged structures in Santiago, capital city of Chile, and a maximum Mercalli Intensity

of X was reported at Las Melosas. The focal mechanisms of these crustal earthquakes show northwest-southeast

to east-west maximum compressive stress (Barrientos et al., 2004). At least two major faults have been

described in the region: the San Ramon fault (Rauld, 2002) and the Chacayes-Yesillo or El Fierro fault (Charrier

et al., 2002 and 2004). The former is probably one fault branch of a major system separating the Principal

Cordillera from the Central Depression, and associated to Andes uplift. The last one represents the eastern

contact of the Cenozoic deposits of the Abanico formation with the uplifted Mesozoic units on the east-side

block of the fault.

This Andean region is close to Santiago city which concentrates more than 30% of the Chilean population. It

includes two giant porphyry copper deposits (El Teniente and Río Blanco-Los Bronces), several medium and

small size hydroelectric power plants, and a gas pipeline coming from Argentina. It also represents the water

supply region for Santiago. In the region reappears the active volcanism that continues to the south, which is

absent in the flat slab zone to the north since 9-10 Ma (Jordan et al., 1983; Kay et al., 1988). The Principal

Cordillera is narrow, with average elevation of 4000 m and peaks over 5000 m shifted about 30 km to the East

relative to the main Andean summits north and south of this segment.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 377-380

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Crustal seismicity

From November, 2005 until March, 2006, we deployed a temporary broadband and short period seismological

network, which was complemented by the permanent stations of the University of Chile in the region. All the

stations recorded in continuous mode, and were able to detect earthquakes with magnitude over than 1.5 (Fig. 1).

Figure 1. Local seismicity recorded by the temporary seismological stations (grey diamonds), and located with a 1D velocity model. Solid line indicates the Chile-Argentina border line, related with the highest peaks of the Andes. Crustal events (green) and intraslab events (red) are also shown in E-W (bottom left) and N-S projection (top right). Number of located events v/s depth is shown at the bottom right.

The crustal seismicity, in the depth range 0-30 km, is well correlated with known geological faults and gives

new data to improve the assessment of the local seismic hazard. It also clusters around and beneath the giant

porphyry copper deposits in the region (Fig. 1), down to 25 km depth. Discounting the mine-blasts and possible

induced events by the mining activity, this clusters show considerable more earthquakes than the surrounding

area, suggesting that the deposits are emplaced in weaker and more fractured zones of the crust. As in previous

seismological deployments in the zone (Pardo et al., 2006), shallow seismicity is observed beneath Santiago city

that can be correlated with an almost vertical hidden fault capable to generate an M~5 earthquake.

The average stress tensor, derived from focal mechanisms of the crustal events, indicates that the Andean zone

is under compression in the plate convergence direction.

P and S waves velocity models

The three-dimensional body wave velocity models were determined using the TLR3 algorithm (Latorre et al.,

2004), with an initial 1D velocity model obtained using mine-blasts in the zone (Vera et al., 2006). Travel times

and hypocenters from the temporary network, improved with selected data of the best recorded an accurate

located events of the University of Chile network, were the database to perform the inversion (14901 P and

14596 S travel times from 1190 events).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 377-380

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The velocity tomography shows channels of high Vp/Vs (>1.8) connecting the subducted slab with the surface

(Fig. 2). Considering that high Vp/Vs ratios may indicate presence of fluids, this observation suggests upward

migration of hydrous or melted rocks. At depths of about 20 to 25 km, a layer of high Vp/Vs is observed

beneath the Andes cordillera that could be associated to changes in the rheological properties between the upper

and lower crust, or to accumulation of magma.

The zones of the seismic clusters related to the porphyry copper mines exhibits high Vp/Vs, which may

indicate fluid phases located in the weakest and more fractured zone of the crust.

Hypocenter locations are improved using the obtained 3D velocity models. The maximum depth of the crustal

seismicity reaches 25 km.

Figure 2. Cross-sections of Vp/Vs velocity ratio with depth. (Left) E-W projection at latitudes along the El Teniente (top) and Río Blanco-Los Bronces (bot.) copper mines. (Right) N-S projection along the El Fierro fault (top), and along the mentioned copper mines (bot.). Seismicity (black dots) was located using the 3D velocity models obtained from tomography.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 377-380

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Conclusions

Continuous mode recording of the temporary network permits to locate low magnitude crustal earthquakes that

in general are well correlated with known faults, but also with hidden faults of presumably high damage

potential. Future seismic hazard evaluations in this highly populated region must then consider low magnitude

seismicity.

The giant copper mines in the zone are located above areas of high Vp/Vs ratios, which are consistent with

partially saturated and fractured crustal zones that exhibits high seismicity.

Seismic velocity tomography improve our models of the geodynamic processes that are taking place in the

region, from the subducted slab to the surface.

Acknowledgements We thanks to the Seismological Service of the University of Chile for providing instruments and their database. This work was partially funded by projects FONDECYT 1050758, ACT-18, and IRD-France.

References

Barrientos, S., Eisenberg, A., 1998. Secuencia sísmica en la zona cordillerana al interior de Rancagua. V Congreso Geológico Chileno, Santiago, F121-132.

Barrientos, S., Vera, E., Alvarado, P., Monfret, T., 2004. Crustal seismicity in Central Chile. J. South Am. Earth Sci., 16, 759–768.

Charrier, R., Baeza, O., Elgueta, S., Flynn, J., Gans, O., Kay, S., Muñoz, N., Wyss, A., Zurita, E., 2002. Evidence of Cenozoic extensional basin development and tectonic inversion south of the flat-slab segment, southern Central Andes, Chile (33°-36°S). J. South Am. Earth Sci., 15, 117-139.

Charrier, R., Bustamante, M., Comte, D., Elgueta, S., Flynn, N., Iturra, N., Muñoz, N., Pardo, M., Thiele, R., Wys, A., 2004. The Abanico extensional basin: Regional extension, chronology of tectonic inversion and relation to shallow seismic activity and Andean uplift. N. Jb. Geol. Paläont. Abh., 236, 43-77.

Jordan, T., Isacks, B., Allmendinger, R., Brewer, J., Ramos, V., Ando, C., 1983. Andean tectonics related to geometry of subducted Nazca plate. Geol. Soc. Am. Bull., 94, 341-361.

Kay, S.M., Maksaev, V., Moscoso, R., Mpodozis, C., Nasi, C., Gordillo, C., 1988. Tertiary magmatism in Chile and Argentina between 28 and 33: correlation of magmatic chemistry with changing Benioff zone. J. South Am. Earth Sci., 1, 21-38.

Lomnitz, C., 1961. Los terremotos del 4 de Septiembre de 1958 en el cajón del Maipo. Anales de la Facultad de Ciencias Físicas y Matemáticas, 18, 279-306.

Pardo, M., Vera, E., Monfret, T., Yañez, G., Eisenberg, A., 2006. Sismicidad cortical superficial bajo Santiago: Implicaciones en la tectónica Andina y evaluación del peligro sísmico. XI Congreso Geológico Chileno 2006, Antofagasta-Chile, 7-11 Agosto, 2006. Actas, Vol.1, Geodinámica Andina, 443-446.

Latorre, D., Virieux, J., Monfret, T., Monteillet, V., Vanorio, T., Got, J.-L., and Lyon-Caen, H., 2004. A new seismic tomography of Aigion area (Gulf of Corinth, Greece) from the 1991 data set, Geophys. J. Int., 159: 1013-1031.

Rauld, R.A., 2002. Análisis morfoestructural del frente cordillerano de Santiago oriente, entre el río Mapocho y la quebrada de Macul. Thesis, Departamento de Geología, Universidad de Chile.

Vera, E., Pardo, M., Monfret, T., Eisenberg, A., Yañez, G., Triep, E., 2006. Eventos Sísmicos Corticales en los Andes Centrales de Chile y Argentina. XI Congreso Geológico Chileno 2006, Antofagasta-Chile, 7-11 Agosto, 2006. Actas, Vol.1, Geodinámica Andina, 469-472.

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Why is the passive margin of Argentinean Patagonia uplifting?: An insight by marine terrace and tidal notches sequences

K. Pedoja1, V. Regard

2, L. Husson

3, J. Martinod

2, & M. Iglesias

4

1 UMR M2C, Université de Caen, 2-4 Rue de Tilleuls 14000 Caen, France, kevin.pedoja@unicaen.fr

2 LMTG - UMR 5563 UR 154 CNRS Université Paul-Sabatier IRD Observatoire Midi-Pyrénées Observatoire Midi-

Pyrénées - 14, avenue Edouard Belin – 31400 regard@lmtg.obs-mip.fr, martinod@lmtg.obs.mip.fr 3 Géosciences Rennes (UMR CNRS 6118), Université de Rennes1, Bâtiment 15, Campus de Beaulieu, CS

74205, F-35042 Rennes Cedex, France laurent.husson@univ-rennes1.fr 4 Departamento de Geologia, Universidad Nacional de la Patagonia San Juan Bosco Ciudad Universitaria -

Comodoro Rivadavia - Km 4 - Chubut - CP (9005), Argentina, maximiliano_iglesias@hotmail.com

KEYWORDS : marine terraces, marines notches, uplift, passive margin, mantle convection

The Quaternary coastal deposits and morphologies found along the Patagonian coast have been known since

the middle of the last century. The most complete descriptions of the chronostratigraphy, lithology and

paleontology of the Quaternary marine terraces of Patagonia derive from the work of Ferruglio (1933, 1950). In

terms of age control for the Patagonian raised terraces, the first radiocarbon ages of mollusc shells were reported

by Codignotto (1983). Rutter et al. (1989, 1990) and Schellmann and Radtke (2000) provided aminostratigraphy

and ESR dating from various locations. Their ESR results yielded Last Interglacial ages (MISS 5e) for some

locations whereas ages obtained on the basis of aminostratigraphy suggest penultimate interglacial (MIS 7) or

older deposits. ESR and Th/U dating by Radtke (1989) showed that Holocene beaches are at higher elevation in

the south of Patagonia than in the north. Theses dating works concentrated mostly on Holocene (data on zone 2,

3, 4, 5, 6, 8) and MISS 5e (last interglacial, data on zone 2,3,4,5,6, on Figure 1)… Older terraces have been dated

and tentatively correlated to MIS 7, 9 & 11 in zone 3,5,6.

Previous works concentrate on dating (see above) and palaeontology (for example Aguirre et al., 2003) but,

with one noticeable exception (Rostami et al., 2000), no effort was made to use theses features as tectonics tools.

In particular it appears that no extensive and precise mapping and altimetry has been achieved and therefore no

morpho-stratigrpahic correlation or comparison was not possible from one site to another.

Therefore our work concentrate on 3D repartition of the marine terraces sequences and accurate altitudes of

their shoreline angles. We divided the area in 8 zones (see Figure 1). Our mapping was done combining field

observation with interpretation on Landsat image (google earth) and on Shuttle Radar DEM (Geomapapp).

Altitude were taken using precise altimeter and telemeter. We took about 100 altitude of shoreline angle and

therefore we have been able to calculate mean uplift rates since ~ 440 ka (MISS 11) with the method proposed

by Pedoja et al., (2006a, b, c ; in press). Then we focused on Last Interglacial Maximum terrace (MISS 5e), the

better constrained marine terraces (both in term of age and altitude) and we used it as a tectonical benchmark to

reveal positive vertical deformation (i.e. uplift) on the Argentinean passive margin. This approach allow us to

scrutinize the relationship between glacio-isostatic rebond versus “long” term tectonic for the Holocene sequence

and long term (ie Quaternary) tectonic. For the latter one, we discuss its geodynamical origin. More particularly

we reject the possibility of uplift due to the vicinity of the Chilean subduction zone (as proposed by rostami et

al., 2000) and we propose that mantle convection anomaly are responsible for the deformation.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 381-383

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Figure 1 : Main localisation of Quaternary coastal deposits and morphologies along the Argentinean Patagonia coast

References Aguirre, M.L., 2003. Late Pleistocene and Holocene palaeoenvironments in Golfo San Jorge, Patagonia: molluscan evidence.

Marine Geology 194, 3 –30. Codignotto, J.O., 1983. Depo sitos elevados y/o de Acrecion Pleistoceno– Holoceno en la costa Fueguino–Patagonica.

Simposio Oscilaciones del nivel del mar durante el ultimo hemiciclo deglacial en la Argentina. (IGCP200). Universidad Nacional de Mar del Plata Actas, 12– 26.

Feruglio, E., 1933. I terrazi marini della Patagonia. Giornale di Geologia. Annali Reale Museo geologico di Bologna, 1– 288. Feruglio, E., 1950. Descripcion geologica de la Patagonia. Direccion General de Y.P.F., T 3, 431 pp. Buenos Aires. Pedoja, K., Ortlieb, L., Dumont, J-F., Lamothe, J-F., Ghaleb, B., Auclair, M., Labrousse, B. 2006 Quaternary coastal uplift

along the Talara Arc (Ecuador, Northern Peru) from new marine terrace data. Marine Geology, 228 : 73-91. Pedoja, K., Dumont, J-F., Lamothe, M., Ortlieb, L., Collot, J-Y., Ghaleb, B., Auclair, M., Alvarez, V., Labrousse, B., 2006.

Quaternary uplift of the Manta Peninsula and La Plata Island and the subduction of the Carnegie Ridge, central coast of Ecuador. South American Journal of Earth Sciences, 22: 1-21.

Pedoja, K., Bourgeois, J., Pinegina, T. and Higman, B., 2006. Does Kamchatka belong to North America? An extruding Okhotsk block suggested by coastal neotectonics of the Ozernoi Peninsula, Kamchatka, Russia. Geology, 34, (5) : 353-356.

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Pedoja, K., Kershaw, S., Shen, J-W., Tang, C. Coastal Quaternary morphologies on the northern coast of the South China Sea, China, and their implications for current tectonic models: A review and preliminary study, Marine Geology in press (2008)

Radtke, U., 1989. Marine Terrassen und Korallenriffe— Das Problem der quartaren Meeresspiegelschwankungen erlautert an Fallstudien aus Chile, Argentinien und Barbados. Duseldorfer Geographische Schriften, vol. 27. Geograph. Inst. D. Heinrich Heine Universitat p. 246.

Rostami, K., Peltier, W.R., Mangini, A., 2000. Quaternary marine terraces, sea-level changes and uplift history of Patagonia, Argentina: comparisons with predictions of the ICE-4G (VM2) model of the global process of glacial isostatic adjustment. Quaternary Science Reviews 19, 1495–1525

Rutter, N., Schnack, E., Del Rio, L., Fasano, J., Isla, F., Radtke, U., 1989. Correlation and dating of Quaternary littoral zones along the Patagonian coast, Argentina. Quaternary Science Reviews 8, 213–234.

Rutter, N., Radtke, U., Schnack, E., 1990. Comparison of ESR and amino acid data in correlating and dating Quaternary shorelines along the Patagonian coast, Argentina. Journal of Coastal Research 6 (2), 391–411.

Schellmann, G., Radtke, U., 2000. ESR dating of stratigraphically well-constrained marine terraces along the Patagonian Atlantic coast (Argentina). Quaternary International 68–71, 261– 273.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 384-386

384

Neotectonics and mass wasting phenomena in the eastern slope of the southern Central Andes (37º-37º30’S)

Ivanna M. Penna1, Reginald L. Hermanns

2, & Andrés Folguera

1

1 Laboratorio de Tectónica Andina y Consejo Nacional de Investigaciones Científicas y Técnicas, Departamento

de Geología, Facultad de Ciencias Exactas y Naturales (Pabellón II), Ciudad Universitaria, Buenos Aires,

Argentina (penna@gl.fcen.uba.ar) 2 Norges Geologiske Undersøkelse, Norway (Reginald.Hermanns@NGU.NO)

KEYWORDS : neotectonic activity, rock avalanches, Southern Central Andes

Introduction

The transitional area between the Central and Patagonian Andes (37º-39ºS) was recognized as an area of

neotectonic activity both in the forearc and western retroarc (Melnick et al., 2006; Folguera et al., 2004). In the

study area the neotectonic activity manifests itself as the N-S striking Antiñir-Copahue fault zone, which is

20-40 km wide. Crustal seismicity is mainly constrained in the forearc area, but some crustal events are

recognized in the retroarc zone where neotectonics evidences were identified. This area of young deformation is

associated with more than 40 rock avalanche deposits (González Díaz et al., 2006; Hermanns et al., 2006).

Further to the north, huge mass wasting phenomena are related to neotectonics features and hence with the

Pampean flat slab zone (Costa et al., 2005). Here, seismicity is associated with high-angle basement reverse

faults, located in a broken Laramide-like foreland area, that produce slope instability.

Ta árová (2004), by gravimetrics models based on seismic data, proposed that this part of the Andes

experiments a decrease of about 10º in the angle of Wadatti-Benioff zone, in contrast with the adjacent segments

where the subduction angle is near 30º.

The aim of this paper is to present new geological evidence of neotectonic activity and deformation mechanism

and their relationship with the spatial distribution of rock avalanches in the Southern Central Andes.

Neotectonics and rock avalanches

The morphologic expression of neotectonic activity corresponds to a series of rectilineous scarps and drainage

anomalies in the Quaternary coverage.

In the fault-valley intersection, it is possible to recognize a direct association between N-S trending scarps and

faults with transpresive dextral mechanism. Those faults are parallel to the strike of Oligocene-Miocene strata of

Cura Mallín basin, which suggests a flexural deformation mechanism (Figure 1).

The Chacayco and Cerro Guañacos faults (Figure 1) are N-S continuous scarp with 200 and 30 meters of

vertical displacement respectively. The intersection between the first one and Reñileuvú valley is linked to a

huge rock avalanche named Chacayco (0.81 km3, Figure 1). Another important feature is related to the Cerro

Guañacos fault. Here, radiometric determinations in highly deformed glacifluvial sequences have allowed us to

constraint neotectonic activity along the Reñileuvú valley up to at least 26,540+510/-480years BP; what

evidence an ongoing tectonic activity in the foothills of the southernmost Central Andes, despite the absence of

several events of shallow seismicity (Bohm et al., 2002). Additionally, morphological evidence show drainage

disturbances in relation to this structure such as abandoned courses, deflected and beheaded courses, young

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 384-386

385

drainage systems developed in the hangingwalls of growing structures, and springs associated with fault scarps

that flooded the footwall faults.

In the Reñileuvu, Ñireco and Guañacos valleys the main evidence of the linkage between neotectonic activity

and slope instability is the occurrence of several avalanches among which six are the most important ones. The

source area of these deposits is located along the strike of the thrusts, in the intersection with deeply incised

glacial valleys (with slopes around 40-30º).

Three of these landslides have yielded around ~3,000 Ka (Hermanns et al., 2006), which allows us to ascribe

these deposits to postglacial times. One would be synglacial based on the finding of a fluvio/glacial coverage

meanwhile another, based on morphologic criteria such as erosional degrees of break away zones and

hummocky topography, and connectional degree of drainage networks over the landslide surface, allow us to

assign its to postglacial times.

Figure 1. Main neotectonics faults and evidences of the young deformation. Block diagram reflect the flexural deformation mechanism, proposed for the area. Note that the Chacayco rock avalanche is located in the fault-valley intersection.

Conclusions

Tectonic activity could be preparatory mechanisms of slopes collapses in two different ways: a) “internal

causes” referring to the spatial coincidence of rock avalanches with the fronts of important thrust faults and b) as

trigger mechanisms by seismic activity.

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The spatial relationship between the rock avalanches and the neotectonic faults that cut orthogonally the

eastern Reñileuvú creek, suggests that the last one plays an important role as a slope instability trigger factor.

This strong linkage between young deformation and landsliding, points to seismic shaking as a trigger factor for

slope failure. Rock avalanche dimensions could contribute to reinforce this hypothesis.

Clustering ages of around 3,000 yr indicates multiple landslide occurrences at the same time. Those are likely

produced by earthquakes which would have also been responsible for local deformation. The necessary

earthquake magnitude to generate a slope collapse of a rock avalanche type would be in the order of 6.0M

according to the statistical approach of Keefer (1984).

The instrumental seismicity recorded in the western retroarc zone at these latitudes is poor (Bohm et al., 2002),

and reflects crustal and interplate events of around 4-5M.

During the last times no significant landslides were produced in this area in spite of this is located only 300 km

to the trench, where megathrust earthquakes such as the Chile 1960 M9.5 event took place. This indicates that

the energy source has to be related to local events.

References Bohm, M., Lüth, S., Helmut E., Asch, G., Bataille, K., Bruhn, C., Rietbrock, A. & Wigger, P., 2002 —The Southern Andes

between 36º and 40ºS latitude: seismicity and average seismic velocities. Tectonophysics, 356(4): 275-289. Costa C. H., 2005 — Large Holocene earthquakes in the Sierras Pampeanas and sorrounding plains: more likely than once

thought. ICSU-IGCP 480, Holocene environmental catastrophes in South America: From the lowlands to the Andes. Laguna Mar Chiquita, Córdoba.

Folguera A, Ramos V, Hermanns R.L, Naranjo J., 2004 — Neotectonics in the foothills of the Southernmost Central Andes (37º-38ºS). Evidence of the strike-slip displacement along the Antiñir-Copahue fault zone. Tectonics. Vol 23, TC5008.

González Díaz, E. F., Folguera, A., Costa, C., Wright, E. & Elisondo, M., 2006 — Los grandes deslizamientos de la región septentrional neuquina entre los 36º y los 38ºS: una propuesta de su inducción por un mecanismo sísmico. Revista de la Asociación Geológica Argentina. 61 (2): 197-217.

Hermanns, R. L., Folguera, A., Penna, I., Naumann, R. & Niedermann, S., 2006 — Morphologic characterization of giant flood deposits downriver landslide dams in the northern Patagonian Andes. Geophysical Research Abstracts. Vol 8.

Keefer, D.K., 1984 — Rock-avalanches caused by earthquakes: source characteristics. Science, Vol 223: 1288- 1290. Melnick. D., Charlet F., Echtler H. P., & De Batist. M., 2006 — Incipient axial collapse of the Main Cordillera and strain

partitioning gradient between the central and Patagonian Andes, Lago Laja, Chile. Tectonics. v. 25, TC5004. Tasarova, Z., 2004 — Gravity data analysis and interdisciplinary 3D modelling of a convergent plate margin (Chile, 36–

42°S). PhD thesis, Freie Universität Berlin, http://www.diss.fu-berlin.de/2005/ 19/indexe.html

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 387-390

387

Current erosion rates in the Northern and Central Andes: Evaluation of tectonic and climatic controls

Emilie Pépin1, Sébastien Carretier

1, Jean-Loup Guyot

2, Elisa Armijos

3, Hector Bazan

4,

Pascal Fraizy5, Luis Noriega

6, Julio Ordóñez

7, Rodrigo Pombosa

3, & Philippe Vauchel

5

1 LMTG-Univ. de Toulouse-CNRS-IRD-OMP, 14 Av. E. Belin, 31400 Toulouse, France (pepin@lmtg.obs-mip.fr)

2 CP 7091, Lago Sul, 71619-970 Brasilia (DF), Brazil

3 INAMHI, 700 Iñaquito y Correa, Quito, Ecuador

4 UNALM – FIA, Avenida La Molina s/n, Lima 12, Peru

5 IRD,Casilla 18-1209, Lima 18, Peru

6 INAMHI, CP 9214, 00095 La Paz, Bolivia

7 SENAMHI – DGH, Casilla 11-1308, Lima 11, Peru

KEYWORDS : present erosion rates, climate, tectonics, total volume eroded, response time

Introduction

The average erosion rate ( [L/T]) of an active uplifting mountain results from a long-term tendency, related to

history of the incision on geological times and with fluctuations linked to climatic variations and production of

sediments. The relative amplitudes of these components for current flows are quite unknown especially in the

Andes. In this work, we have studied and proposed a new method in order to understand what can control the

current erosion rate at catchments scale.

Method and study area

Figure1: Localization of the nine studied basins, they are mostly mountain basin. Gauging stations are represented by red squares. For each one, local slope, area and total eroded volume are calculated from SRTM, average rainfall, rainfall variability, sediment flows and yield are calculated with HYBAM (www.mpl.ird.fr/hybam/) series of data, unless the Colombian basin which data come from Respeto J., Kjerfve, B., 2000.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 387-390

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We propose a simple method allowing the evaluation of a possible first order control between the transient

answer of erosion to uplift and the mean erosion rates of adjacent elevated basins. If it is confirmed, it means that

the eroded total volume standardized by the surface of the basins (V/A) is positively correlated with the erosion

rates (see figure 2).

Moreover, we establish analytical relations between V/A and (figure 3). These relations allow us to calculate

a response time to the uplift or its age, making an assumption on the form of (t) and A(t) during geological

times.

Figure 3: A. Relation between the data and the models of the developed equations. The models predict ages of uplift or time required to incise the plate between 6 and 10 My. B. developed equations. Equation (1) is valid when all the basins have the same co but a different uplift time. Equation (2) is valid for different co but a same uplifting initial time.

Figure 2: Schematic explanation of the correlation method between the eroded volume standardized by the drained surface (V/A) and the erosion rate ( ). Graph on the top shows the temporal evolution of V/A and for catchments having different times of connectivity

co. As long as time is lower at least co (ie. that the basin which develops most quickly did not reach its maximum surface), V/A are correlated positively with . B- Strong variations are added to the response to tectonics. In this case, it is very unlikely that a correlation would be obtained between V/A and because the erosion rate of each basin responses with different amplitudes and dephasings when the basins are subjected to climatic or internal variations (Tucker and Slingerland, 1997).

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389

This approach is tested on 9 basins Western slopes of the central and north Andes, whose current average rates

of erosion (0.2-1.2 mm/a) are calculated starting from data series of discharge and suspended sediments long

from 2 to 37 years.

Figure 4: Average yield and rainfall calculated on the area. A. Current rates of erosion calculated from interannual sediments flux averages observed at the gauging station of each basin. B. Average rainfall and rainfall annual mode for each catchments area.

Results

The current erosion rates are correlated with the variability of precipitations and anti-correlated with average

precipitations on the unit data. (figure 5).

Figure 5: Correlation between average erosion rate and rainfall: is correlated with rainfall variability and anti-correlated with average rainfall for scale-catchments area.

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390

After developing two kinds of initial surface (envelop surface and initial plate) to calculate the total eroded

volume, we did not find a clear relationship between the current erosion rates and the total eroded volumes of the

nine basins. So, it seems that for our studied area, there is no first order control of the transient answer of erosion

to uplift on the mean erosion rates. Considering only the four bolivian basins, a linear relationship between and

V/A can be found for any initial surface choosen. This result permits us to think about the possible control of the

mean erosion rate, only on the bolivian basins, by the transient answer of erosion to uplift controls. This thought

confirms Safran et al. 2005 and Aalto et al. 2006 results.

Discussion and conclusion

The interpretation of the relative role of average rainfall and rainfall variability depends on the real role of the

long-scale component; first possibility, the response time of erosion to uplift controls effectively the rates of

erosion, in which case it is the average rainfall which plays a main role by controlling these response times.

Second possibility, on the contrary it is the variability of the climate which explains the space variations, the

erosion rates being stronger as variability is high.

The respective roles of the variability and the average of the climate on the current average rates of erosion

cannot be evaluated without taking in account of the state of the erosive answer on very a long-term scale. There

is not clear first order control of the transient answer of erosion to uplift on the mean erosion rates on this area

unless on the bolivian basins. Moreover, it could be useful to investigate the way to calculate more precisely the

total eroded volume.

References Aalto, R., Dunne, T., Guyot, J., 2006 — Geomorphic Controls on Andean Denudation Rates. The Journal of Geology 114,

85-99. Restrepo, J., Kjerfve, B., 2000 — Magdalena River: interannual variability (1975-1995) and revised water discharge and

sediment load estimates. Journal of Hydrology 235, 137-149. Safran, E., Bierman, P., Aalto, R., Dunne, T., Whipple, K., Caffee, M., 2005. Erosion rates driven by channel network

incision in the bolivian andes. Earth Surf. Proc. Land. 30, 1007–1024. Tucker, G. E., Slingerland, R., 1997 — Drainage basin responses to climate change. Water Resources Research. 33, 2031-

2047.

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391

The volcanic rocks of the Mondaca river, Cordillera Principal (31°45'S), San Juan province, Argentina

Daniel J. Pérez & Juan Manuel Sánchez-Magariños

Laboratorio de Tectónica Andina, Departamento de Ciencias Geológicas, Universidad de Buenos Aires, Ciudad

Universitaria 1428, Buenos Aires, Argentina (daniel@gl.fcen.uba.ar)

KEYWORDS : Pelambres, Abanico, Miocene, Principal Cordillera, Mondaca River, San Juan, Argentina

Introduction

The objective of the present contribution is to analyze the volcanic rocks and their relationship within

Mesozoic sedimentary deposits in the Mondaca river region, to the west of Cerro Mercedario. New field

structure data from these region, indicate that these volcanic deposits are Oligocene and Miocene in age. The

study region is located in the Frontal and Principal Cordillera at 31º45'S and 70°15'W, to the west of the

Mercedario Mountain and east of the Paso del Mondaca, in San Juan province, Argentina. This region is in the

southern part of the modern non-volcanic “flat-slab” (Cahill e Isacks, 1992) region betwen 28° and 33°S, under

which the Nazca plate forms a broad sub-horizontal bench between about 100 and 150 km. The first studies in

the region were done by Groeber (1951), Polanski (1964), Olivares Morales (1985), Rivano and Sepulveda

(1991); and more recent studies by Alvarez (1996), Pérez (2001), Ramos et al. (1998).

Stratigraphy and structure

The stratigraphic sequences of the area begin with Permo-Triassic rhyolitic and rhyodacitic rocks of the

Choiyoi Group. The continue with Permo-Triassic and Triassic rocks of the Rancho de Lata Formation and

Jurassic sequences of the Los Patillos, La Manga and Tordillo Formations. Without stratigraphic relationships

continued the Auquilco formation whit gypsum and diapirs.

It follows a sequence of volcanic rocks defined in the Chile region as Los Pelambres Formation by Rivano and

Sepúlveda (1991) and that to the Alitre and Mondaca Pass enter to Argentina region. These same rocks in the

The La Ramada located immediately to the south of the study area, they were assigned to the Juncal Formation

(Ramos et al. 1990), by the way, these same volcanic rocks immediately to the north of the study region, they

would have thrown upper Oligocene lowe Mioceno in ages, being assigned to the Abanico Formation (Bertens,

2006). Unconformable above these rocks continued the volcanic rocks of Farellones Formations with Miocene

age. Above all these mentioned units, they are quaternary deposits broadly distributed inside which deposits

were recognized glaciers, alluvial fun, etc.

The study region presents two structural styles, one of thin skinned and another of thick skinned; which

affected to different rocks and in different periods of times. The first style can see in the Mondaca and

Carnicerias river, by Los Pelambres thrust of low angle, which would be uplifting the volcanic sequences of

upper Oligocene and lower Mioceno in age, over Los Patillos Formation of Jurassic age. Toward the west and in

Chilean territory another landslide of low angle would be the responsible of uplifting the Cretaceous deposits of

the homonymous formation on the tertiary volcanic rocks. These are attributed to the different deformation

phases of out of sequences thrust and in Miocene times. Similar structure they would already have been

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 391-392

392

describes to the south of the study region and affecting to cretaceous volcanic rocks of Los Pelambres formation

(Cristallini et al., 1996).

Discussion

With relation to the volcanic rocks located between the Alitre and Mondaca pass, and these continue for the

Mondaca river, and then toward the north for the Carnicerias river, they would correspond to the Farellones

formation according to Olivares Morales (1985). On the other hands the volcanic rocks to the north of the

Mondaca pass would be Cretaceous in age and corresponding to the Río Totoral or Los Pelambres formations

(Rivano and Sepúlveda, 1991). In this work and over above discusses, we are assigned to the volcanic rocks

located on, Alitre and Mondaca pass, high Mondaca river, continued along the chilen boundary toward the north

of Pachón river, continues for the Carniceria river, reaching the Yeso and Pantanosa rivers, a upper Oligocene

lower Miocene age; and decides denominate to these volcanic rocks as Mondaca Formation.

The volcanism rocks of Mondaca river, represent the initial volcanic sequences in this region, for oligocene

miocene times, before Farellones formations, and represents the initial of volcanism for tertiary times. These

data require reconsideration of paleogeogarphic reconstructions for this time.

References Alvarez, P.P., 1996. Los depósitos triásicos y jurásicos de la Alta cordillera de San Juan. En V.A. Ramos (ed). Geología de la

región del Aconcagua, provincias de San Juan y Mendoza. Subsecretaría de Minería de la Nación. Dirección Nacional del Servicio Geológico. Anales 24 (5): 59-137, Buenos Aires.

Bertens, A., Clark, A.H., Barra, F. y Deckart, K., 2006. Evolution of the Los Pelambres-El Pachón porphyry copper-molybdenum district, Chile-Argentina. XI Congreso Geológico Chileno, Actas, Vol.2, Geología Económica, Antofagasta.

Cahill, T. y Isacks, B.L., 1992. Seismicity and shape of the subducted Nazca plate. Journal of Geophysical Research . Nº97, p. 17503-17529.

Cristallini, E.O. y Ramos, V.A., 1996. Los depósitos continentales cretásicos y volcanitas asociadas. En V.A. Ramos (ed). Geología de la región del Aconcagua, provincias de San Juan y Mendoza. Subsecretaría de Minería de la nación, Dirección Nacional del Servicio Geológico. Anales 24 (8): 231-273, Buenos Aires.

Groeber, P., 1951. La Alta Cordillera entre las latitudes 34° y 29°30'. Instituto Investigaciones de las Ciencias Naturales. Museo Argentino de Ciencias Naturales B. Rivadavia, Revista (Ciencias Geológicas) I(5): 1-352, láminas I-XXI, Buenos Aires.

Olivares Morales América Patricia, 1985. Geología de la Alta Cordillera de Illapel entre los 31°30 y 32° Latitud Sur. Tesis de Grado, Universidad de Chile, Facultad de Ciencias Físicas y Matematicas Departamento de Geología y Geofísica.

Pérez, D.J., 2001. El volcanismo neógeno de la cordillera de las Yaretas, Cordillera Frontal (34°S) Mendoza. Revista de la Asociación Geológica Argentina, 56 (2):221-23, Buenos Aires.

Polanski, J., 1964. Descripción geológica de la hoja 25a Volcán San José, provincia de Mendoza, Dirección Nacional de Geología y Minería, Boletín 98: 1- 94, Buenos Aires.

Ramos, V.A., Rivano, S., Aguirre-Urreta M.B., Godoy, E. y Lo Forte, G.L., 1990. El Mesozoico del Corcón del Límite entre Portezuelo Navarro y Monos de Agua (Chile-Argentina). XI Congreso Geológico Argentino, Actas II: 43-46, San Juan.

Rivano, G. y Sepulveda, H.,1991. Hoja Illapel Región de Coquimbo, Servicio Nacional de Geología y Minería, Carta Geológica de Chile. Nº69. 132pp..

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 393-396

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Geophysical modeling of intrusive bodies: A case study in the Fuegian Batholith. Argentina

J. I. Peroni1, A. Tassone

1, M. Cerredo

1, H. Lippai

1, M. Menichetti

2, E. Lodolo

3, & J. F. Vilas

1

1

CONICET-INGEODAV, Dpto. de Ciencias Geológicas, UBA, Pabellón 2, Ciudad Universitaria, Buenos Aires,

Argentina (peroni@gl.fcen.uba.ar) 2

Istituto di Scienze della Terra, Universita’ di Urbino, Campus Universitario, Urbino 61029, Italy 3

Istituto Nazionale di Oceanografia e Geofisica Sperimentale, Borgo Grotta Gigante, 42/c. 34010 Sgonico,

Trieste, Italy

KEYWORDS : geophysic modelling, magnetic anomalies, tectonics, intrusive bodies, Tierra del Fuego

The Ushuaia Pluton (UP) is one of the five isolated intrusive bodies of the Fuegian Batholith (FB) in the

Argentine sector of Tierra del Fuego north of the Beagle Channel (Figure 1). These intrusive bodies are located

in the southernmost tip of Andean Cordillera where there is an abrupt change of about 90° in its strike around

53°S from a roughly north-south direction in continental Patagonia into an east-west orogen in the island of

Tierra del Fuego. This sharp orogene curvature is presently contentious, as it has been considered either as a post

Early Cretaceous feature likely related to the opening of the Drake Passage (Dalziel et al., 1974) or that it is due

to major strike-slip offsets (Cunningham, 1993).

Figure 1: A) Sketch showing the location of Northern Patagonian Batholith (NPB), Southern Patagonian Batholith (SPB) and Fuegian Batholith (FB). B) Major fault systems and structural domains in southern Tierra de Fuego. SCB Beagle Channel Fault System, FC: Cadic Fault, FA: Andorra Fault, FBE: East Beagle Fault, MFS: Magallanes-Fagnano Fault System (after Menichetti et al. 2007). Main outcrops of Fuegian Batholith in Argentinean Tierra del Fuego are also indicated.

The UP is a poorly exposed epizonal intrusive body cropping out on the northern margin of Beagle Channel; it

is hosted in the Early Cretaceous Yahgan Formation which was strongly deformed by the Late Cretaceous

Andean compression phase. An extended metamorphic contact aureole ( > 1km) is recognized within the

turbidites of Yahgán Fm. reaching up to the garnet zone (Peroni, 2006). Both, pluton and host are affected by a

set of normal and strike-slip faults associated with the main Beagle Channel Fault System (Figure 2a).

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The UP is compositional varied and includes several facies: a basic to ultrabasic facies, a heterogenous facies

where magma mixing processes are evident and marginal roof facies of volcanic/hypabyssal character (Cerredo

et al., 2007). Along the northern shoreline of Ushuaia bay two main facies were recognized (Figure 2b): the

easternmost areas are dominated by ultrabasic-basic rocks (hornblendite facies), whereas the western sector is

more heterogeneous with widespread syenite-monzonite rocks and lesser amounts of ultrabasic-basic rocks

(syenite-hornblendite facies).

Figure 2: A) Geological sketch depicting the main units and structures in northern Beagle Channel (after Menichetti et al, 2007). B) Detail of petrographic facies in the UP along Ushuaia bay shoreline (after Peroni, 2006).

The 5569-II aeromagnetic chart (not reduced to the pole, SEGEMAR, 1998) was employed to model the UP

(Fig. 3a); this 1:250.000 chart was produced from high level survey (120 meters), with N-S flight lines each

500 m and E-W control flights each 5000 m. This high level, tightly constrained aeromagnetic grid is appropriate

for mapping regional, subsurface geology

Within the Ushuaia bay area the aeromagnetic map shows maximum with bell form (Fig. 3A) to the N and an

elliptic minimum to the south. The former displays asymmetric distribution of contour lines, with the steepest

gradients to the S and associated highest values of 350 nT. The minimum in turn, with an E-W oriented axis

attains the lowest values of around -320 nT.

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The 3D modeling of the UP was performed with the Encom ModelVision Pro 7.0 software (Encom

Technology, 2002), which employs the Won and Bevis (1987) algorithm to calculate the magnetic anomalies by

means of a polygon of n sides within a bidimensional space. Software computing requires entering the

International Geomagnetic Reference Field (IGRF) of the studied area, as well as the magnetic susceptibility

and the remanent magnetization of intervening rock units.

The obtained 3D model of UP displays a laccolithic profile with a central thickest area (around 4000 m) and

thinner borders (around 500 meters). Modeled pluton is slightly oval on plant view with N-S major axis (around

12 km), E-W minor axis (10 km) and total volume of 140 km3.

Figure 3: A) Map of aeromagnetic anomalies -not reduced to the pole- 5569-II (SEGEMAR, 1998) (total magnetic field). Contour lines each 5 nT. B) Magnetic profile 1-2, location indicated in A; C) Interpreted schematic cross- N-S section built from the numeric modeling combined with available geological data (after Menichetti et al, 2004).

The modeled UP was integrated with a regional cross-section (Fig.3c), which was built on combined surface

geology and seismic reflection seismic profiles (Menichetti et al, 2004; Menichetti et al., 2008). Topography

was obtained from a digital elevation model.

The modeled body overprints the thrust complex, which includes two south-dipping duplexes, and is affected

by extensional features. The normal faults are splay of the Beagle channel strike-slip fault that shows a structure

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 393-396

396

with double vergence in both side of the Beagle Channel. The faults offset are mainly extensional of few km

respect to the strike-slip. In the northern part of the section the positive structure of the Valle Carbajal is shown.

Since the Late Cretaceous, strike-slip tectonics, would have dominated the Tierra del Fuego area both on-

shore and the Atlantic off-shore (Cunningham, 1993; Menichetti et al., 2008; Tassone et al., 2008). The above

presented structural framework (Figs. 1, 2), combined with regional geology (Fig.3) and magnetic modeling of

UP support a transtensive regime as a likely scenario for pluton emplacement.

References Cerredo, M. E., Remesal, M.B., Tassone, A., Menichetti, M., Peroni, J. I. 2007, 2007 Ushuaia Pluton: petrographic facies and

geochemical signature. Tierra del Fuego Andes. Geosur 2007, pp 31, Santiago de ChileCunningham, W. D. 1993 Strike-slip faults in the southernmost Andes and the development the Patagonian orocline. Tectonics, v. 12 (1) : 169-186.

Dalziel, I.W.D., de Wit, M.J., Palmer, F.K., 1974. Fossil marginal basin in the southern Andes. Nature 250, 291–294. Encom Technology, 2002. ModelVision Pro v.7.0. Encom Technology, Sydney, Australia. Menichetti M., Acevedo R. D., Bujalesky G. G., Cenni M., Cerredo M. E., Coronato A., Hormachea J. L., Lippai H., Lodolo

E., Olivero E. B., Rabassa J. & Tassone A.2004. Field Trip guide of the Tierra del Fuego. Geosur meeting Buenos Aires 2004, 39 p.

Menichetti, M., Tassone, A., Peroni, J. I., Gonzàlez Guillot, M., Cerredo M.E. 2007 Assetto strutturale, petrografia e geofisica della Bahía Ushuaia – Argentina. Rend. Soc. Geol. It., 4 (2007), Nuova Serie, 259-262, 3 ff.

Menichetti M., Lodolo, E., Tassone A., 2008. Structural geology of the Fuegian Andes and Magallanes fold-and-thrust belt – Tierra del Fuego Island. Geologica Acta, 6 (1) : 19-42.

Peroni J.I. 2006. Anomalía magnética en Bahía Ushuaia (Tierra del Fuego). Estudio Geofísico de la continuidad de las unidades geológicas en subsuelo. Trabajo final de Licenciatura. Dpto. de Geología. Facultad de Ciencias Exactas y Naturales. Universidad de Buenos Aires, 90 p.

Servicio Geológico Minero Argentino (SEGEMAR). 1988. Levantamiento geofísico aéreo magnetometría aérea de Tierra del Fuego. Proyecto PASMA. Hoja Ushuaia 5569 II Escala 1:250.000

Tassone, A., Lodolo, E., Menichetti, E., Yagupsky, D., Caffau, M. and Vilas. J. F. 2008. Seismostratigraphic and structural setting of the Malvinas Basin and its southern margin (Tierra del Fuego Atlantic offshore). Geologica Acta, 6 (1) : 1-13

Won, I. J. y Bevis, M. G. 1987. Computing the Gravitational and Magnetic Anomalies due to a Polygon: Algorithms and Fortran subroutines. Geophysics, 52 (2) : 232-238.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 397-400

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Influence of tectonic and magmatic parameters in the deformation of the Andean subduction margin in Central Chile based on analogue models

L. Pinto, F. Albert, & R. Charrier

Departamento de Geología, Universidad de Chile, Plaza Ercilla 803, Casilla 13518, Correo 21, Santiago, Chile

(lpinto@ing.uchile.cl, francisca.albert@gmail.com, rcharrie@cec.uchile.cl)

KEYWORDS : analogue modelling, tectonic parameters, flat-slab, magmatic chambers, Central Andes

Introduction

Tectonic segmentation in the Central Andes is a primary feature related to the geometry of the Nazca Plate

(Baranzangi and Isacks, 1976; Cahill and Isacks, 1992; Jordan et al., 1983a,b). Zones of low dipping (<10º)

subduction planes (flat-slab subduction zones or segments), like the region located between ca. 28° and 33°S,

generally correspond to zones of strong coupling (Gutscher, 2002; Scheuber et al., 1994). North and south, of

this zone the dip of the subduction plane increases to ~25º-30º and coupling decreases (Scheuber et al., 1994).

Moreover, in Central Chile, between the flat-slab zone and the normal subduction zone located to the south

(33°S-33°45'S) there is a transitional zone with particular morphostructural characteristics (Rivera and

Cembrano, 2000). It has been postulated that the flat-slab zone evolved over time (eg. Yañez et al., 2001; Kay

and Mpodozis, 2002). Throughout this evolution eastward shift of magmatism has been a direct function of the

decrease of the subduction angle, such that today in the flat-slab zone along the prolongation of the present-day

volcanic axis, north and south of this zone, there is a complete absence of volcanic activity since 5 myr ago

(Jordan et al., 1983b; Kay et al., 1999; Ramos et al., 2002). It has been postulated (eg. Hervé, 1994; Cembrano

et al., 1996) that in the southern Central Andes (south of ~39ºS) the presence of the magmatic arc has been an

important factor that influenced the partition of deformation. There are several studies that attempted to explain

the factors that have influenced the geometry of deformation along the different subduction segments (eg. Yañez

et al., 2001). However, no modelling studies have been performed to evaluate the influence and interaction of

tectonic and magmatic parameters in the morphostructural configuration of the Central Andes. This prompted a

series of analogue modelling experiments with which we analyzed the influence on deformation of: pre-existent

morphostructural features, the presence of magmatic zones, the degree of coupling, and the angle of convergence

for a region with a tectonic regime like the one predominating in Central Chile in Cenozoic times. In this

contribution we analyse the effects on deformation caused by the zone of magmatic chambers (MCZ) and the

degree of coupling along the continental margin. Analyses on the effects caused by the angle of convergence

have been presented by Albert et al. (2008).

Experimental setup

From Late Eocene to Miocene times the magmatic arc developed in a fault bounded extensional basin

(Abanico Basin) located between the present-day Coastal Cordillera and the easternmost Principal Cordillera in

Central Chile. Based on the aforementioned tectonic setting, the experimental models developed took primarily

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 397-400

398

into account the presence of (Fig. 1): a) A rigid wedge at the W border of the experiment (CC), b) an area

consisting of an upper brittle layer and a lower ductile layer between the wedge and the mobile piston (CD and

PC), and c) a silicone trapeze to simulate the ductile area corresponding to the MCZ. All elements are N-S

oriented and parallel to each other. The brittle behavior was represented by quartz sand (diameter <500 μm,

density 1,400 kg/m3), ductile behavior by silicone (density 1,400 kg/m3, viscosity 4x104 Pa/s), and the rigid

wedge by wood (Fig. 1). The rate used was 3 cm/hr for 2 hr. The dimensions of the various morphostructural and

magmatic units are shown in Fig. 1b. To analyze the effects of the degree of coupling two different models were

performed, the first, with a 0.5 mm thick galvanized sheet fixed to the mobile piston (E side) and situated below

the silicone (Fig. 2c). This sheet distributes the movement at the experimental base simulating a high degree of

coupling. The second model was prepared without sheet simulating a weak coupling (Fig. 2a).

a) .W mobile piston .E b)

Figure 1. a) A W-E profile of the experimental device. The dotted lines in P and Q show the areas where faulting associated with the MCZ. The morphostructural zones are given only for reference. b) Dimensions of the device. The middle column contains the values observed in nature and the column to the right contains the values at the model.

Results

The analogue models showed that the MCZ ductile zone (arc) controls significantly the development of crustal

structures. The models also showed that the angle of obliquity is a parameter of second order that only modifies

the resulting geometry and magnitude of displacement of the structures formed in the MCZ region (Albert et al.,

2008). We observed that a low angle of obliquity together with a high degree of coupling causes greater uplift in

the MCZ than in experiments with higher angle of obliquity (Albert et al., 2008).

The degree of coupling resulted to be a significant parameter in the configuration of the structural system. In

experiments with weak coupling (Fig. 2a, b) a slight east-vergent thrusting occurred in the MCZ (Q in Fig. 1a,

F2 in Fig. 2b). Intense deformation was concentrated in the mobile piston area, where two thrust-faults with

opposite vergencies were formed; the western fault presented a greater displacement than the other one and

developed a fold in its front (F1, Fig. 2b). In the case of high coupling (Fig. 2c,d), the deformation was

concentrated on the MCZ with the development of a pop-up structure (F1 and F2, Fig. 2d), where the edges of

the MCZ define the position of main structural systems (P and Q in Fig. 1a). This concentration of deformation

led to a greater uplift of the MCZ and also a greater displacement of the structures than in the experiments

developed with weak coupling (Fig. 2d).

[km] [cm]

a 20 5,9

b 20 5,9

c 150 44,1

d 40 11,8

e 4,3 1,3

f 6,6 2

g 33 10

h 66 20

i 9,2 2,7

P Q

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 397-400

399

N N

Therefore, a preliminary analysis of the results shows that tectonic factors, the presence of morphostructural

units, and the existence of a ductile zone (MCZ) have a strong influence on the style of deformation. The degree

of coupling defines the area where the greater deformation will occur; for example, in a strong coupling system,

the MCZ will concentrate the deformation and in this case the angle of obliquity will define the geometry and

amount of displacement along the structures next to it (see Albert et al., 2008).

a) CA = 6 cm; %A = 100% c) CA = 6 cm; %A = 100% A ------------------------------------------------ A’ B --------------------------------------------------- B’

compression

compression

A F2 F1 A’ B F1 F2 B’

b) d) Figure 2. Experiments with different coupling degrees: a) Weak coupling: View from above of the last state of the experiment. b) A-A’ section showed in a. c) Strong coupling: View from above of the last state of the experiment. A galvanized sheet was fixed to the mobile piston (right side), below the silicone. d) B-B’ section showed in c. In this case, the model developed two thrust faults with opposite vergencies (F1 and F2) defining a pop-up. The scale showed is 5 cm. Colors of fault traces represent the time at which they appeared: the yellow fault trace appeared before the red one. CA represents the value of shortening in centimeters. %A represents percentage of shortening respect to the initial length.

Application to regional problems

In other regions of Chile, south of the transitional zone, it has been shown that the presence of a magmatic arc

has a direct influence on the distribution of deformation, which, according to Cembrano et al. (1996) is

concentrated along the transpressive Liquiñe-Ofqui Fault Zone (40º-46ºS). This situation results in the almost

complete lack of deformation in the foreland area (eg. Hervé, 1994). In the flat-slab zone, Maksaev et al. (1984)

described a system of faults that forms a large pop-up structure bounded by north-south oriented thrust-faults

(the Vicuña-San Félix Fault, to the west, and the Baños del Toro Fault, to the east). The main stage of faulting in

this case would have occurred at some moment in the upper Tertiary (Mpodozis and Davidson, 1980), causing

inversion of the northern prolongation of the Abanico Basin that hosted the Late Eocene to Late Oligocene/early

Miocene magmatic arc/intra-arc (Charrier et al., 2005). However, in this region the deposits accumulated in the

basin are scarcely exposed. Further south, in the transition zone, the basin deposits are well exposed. Here, a

system of thrust-faults with opposite vergencies (the San Ramón-Pocuro Fault, to the west, and the El Diablo-El

Fierro Fault, to the east) caused uplift of these deposits (the Abanico Formation) coevally with volcanic activity

associated with the Farellones Formation (Charrier et al., 2002). In this case, both faults were most probably

located over the MCZ, similarly as shown in Fig. 1a, and are responsible for the inversion of the Abanico Basin

(Charrier et al., 2005; Fock et al., 2006).

In the examples given for the flat-slab and transitional zones, deformation was probably controlled, apart from

the existence of a ductile zone corresponding to the arc/intra-arc, by pre-existent weakness zones defined by the

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400

normal faults that participated in the extensional stage of the basin. In relation with the example given for further

south, deformation was controlled by the Liquiñe-Ofqui Fault Zone and concentrated along the fault-zone and

magmatic arc. Considering the geometry of the main thrust-faults with opposite vergencies developed in the flat-

slab and transition zones, we propose that a high degree of coupling was fundamental for the development of the

thrust-fault systems.

Acknowledgements We acknowledge funding by the Departamento de Investigación y Desarrollo, Universidad de Chile (Project DI 2004, I2 04/02-2) and Proyecto Anillo ACT Nº 18.

References Albert, F., Pinto, L., Charrier, R. 2008. Influencia del ángulo de oblicuidad en la deformación del margen de subducción

andino en Chile Central basada en modelos análogos. Submitted to the 17th. Congreso Geológico Argentino, San Salvador de Jujuy, October 7-10, 2008.

Barazangi, M., Isacks, B. 1976. Spatial distribution of earthquakes and subduction of the Nazca plate beneath South America. Geology 4: 686-692.

Cahill, T., Isacks, B.L. 1992. Seismicity and shape of the subducted Nazca plate. J. Geophys. Res. 97: 17503-17529. Cembrano, J., Hervé, F., Lavenu, a. 1996. The Liquiñe–Ofqui fault zone: a long-lived intra-arc fault system in southern

Chile. Tectonophysics 259: 55– 66. Charrier, R., Baeza, O., Elgueta, S., Flynn, J.J., Gans, P., Kay, S.M., Muñoz, N., Wyss, A.R., Zurita, E. 2002. Evidence for

Cenozoic extensional basin development and tectonic inversion south of the flat-slab segment, southern Central Andes, Chile (33°-36°S.L.). J. S. Am. Earth Sci. 15: 117-139.

Charrier, R., Bustamante, M., Comte, D., Elgueta, S., Flynn, J.J., Iturra, N., Muñoz, N., Pardo, M., Thiele, R., Wyss, A.R. 2005. The Abanico Extensional Basin: Regional extension, chronology of tectonic inversion, and relation to shallow seismic activity and Andean uplift. Neues Jahrbuch für Geologie und Paläontologie Abh. 236 (1-2): 43-77.

Fock, A., Charrier, R., Farías, M., Muñoz, M. 2006. Fallas de vergencia Oeste en la Cordillera Principal de Chile Central: Inversión de la cuenca de Abanico (33º-34ºS). Rev. Asoc. Geol. Argent., Serie D, Publicación Especial No. 10, 48-55.

Gutscher, M.-A. 2002. Andean subduction styles and their effect on thermal structure and interplate coupling. J. S. Am. Earth Sci. 15: 3-10.

Hervé, F. 1994. The southern Andes between 39º and 44ºS latitude: the geological signature of a transpressive tectonic regime related to a magmatic arc. In: Reutter, K.-J., Scheuber, E., Wigger, P.J. (eds). Tectonics of the Southern Central Andes, Springer, Berlin, pp. 243– 248.

Jordan, T.E., Isacks, B., Ramos, V.A., Allmendinger, R. 1983a. Mountain building in the Central Andes. Episodes 3: 20-26. Jordan, T.E., Isacks, B., Allmendinger, R., Brewer, J., Ramos, V., Ando, C., 1983b. Andean tectonics related to geometry of

subducted Nazca Plate. Geol. Soc. Am. Bull. 94 (3): 341-361. Kay, S.M., Mpodozis, C., Coira, B. 1999. Neogene Mamatism, tectonism, and Mineral Deposits of the Central Andes 22° to

33°S latitude. In: Skinner, B.J. (ed.). Geology and Ore Deposits of the Central Andes. Society of Economic Geology, Special Publication 7: 27-59.

Kay, S.M., Mpodozis, C. 2002. Magmatism as a probe to the Neogene shallowing of the Nazca plate beneath the modern Chilean flat-slab. J. S. Am. Earth Sci. 15: 39–57.

Maksaev, J., Moscoso, M., Mpodozis, C., Nasi, C. 1984. Las unidades volcánicas y plutónicas del Cenozoico Superior en la alta cordillera del norte chico (29°-31°S): geología, alteración hidrotermal y mineralización. Rev. Geol. Chile 21: 11-51.

Mpodozis, C., Davidson, J. 1980. Estructuras gravitacionales en los Andes del Norte Chico de Chile. Rev. Geol. Chile 10: 17-31.

Ramos, V.A., Cristallini, E., Pérez, D.J. 2002. The Pampean flat-slab of the Central Andes. J. S. Am. Earth Sci. 15: 59-78. Rivera, O., Cembrano, J. 2000. Modelo de formación de cuencas volcano-tectónicas en zonas de transferencia oblicuas a la

cadena andina: el caso de las cuencas oligo-miocénicas de Chile Central y su relación con estructuras NWW-NW (33°00’–34°30’S). In: Proceedings of 9th Congreso Geológico Chileno, Puerto Varas, vol. 2, pp. 631–636.

Scheuber, E., Bogdanic, T., Jensen, A., Reutter, K-J., 1994. Tectonic development of the North Chilean Andes in relation to plate convergence and magmatism since the Jurassic. In: Reutter, K-J., et al. (eds). Tectonics of Southern Central Andes, Springer, Berlin, pp. 121-137.

Yañez, G.A., Ranero, C.R., von Huene, R., Díaz, J. 2001. Magnetic anomaly interpretation across a segment of the Southern Central Andes (32-34°S): implications on the role of the Juan Fernández Ridge in the tectonic evolucion of the margin during the upper Tertiary. J. Geophys. Res. 106 (B4): 6325-6345.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 401-404

401

Structural styles in the Eastern Cordillera, Subandean Ranges - Santa Barbara System transition, and Lomas de Olmedo Trough (northern Argentine Andes)

Josep Poblet1,2

, Mayte Bulnes1,2

, Raul E. Seggiaro3,4

, Néstor G. Aguilera3, L. Roberto

Rodríguez-Fernández2,5

, Nemesio Heredia2,6

, & Juan Luis Alonso1,2

1 Departamento de Geología, Universidad de Oviedo, C/Jesús Arias de Velasco s/n, 33005 Oviedo, Spain

(jpoblet@geol.uniovi.es, maite@geol.uniovi.es, jlalonso@geol.uniovi.es) 2 Consolider Team “Topo-Iberia”

3 Universidad Nacional de Salta, Buenos Aires 177, 4400 Salta, Argentina (raul@unsa.edu.ar)

4 Servicio Geológico y Minero Argentino, Avda. de Bolivia 4750, 4400 Salta, Argentina (nestor@unsa.edu.ar)

5 Instituto Geológico y Minero de España, C/Ríos Rosas 23, 28003 Madrid, Spain (lr.rodriguez@igme.es)

6 Instituto Geológico y Minero de España, C/Matemático Pedrayes 25, 33005 Oviedo, Spain (n.heredia@igme.es)

KEYWORDS : inversion tectonics, structural style, Eastern Cordillera, Subandean Ranges, Santa Barbara System

Introduction

The Andean Cordillera, that resulted from convergence between the Nazca subducted plate and the South

American plate, underwent a complex tectonic history and exhibits a notable along-strike segmentation in terms

of structural style, lithospheric thickness and geometry of the subducted plate. This caused the definition of

different geological segments. A critical region to understand the transition between geological provinces with

different features was studied. This area is located between parallels 23°S and 24°S in the northwest corner of

Argentina, close to the Bolivian border, in the Jujuy province (Fig. 1). Several interesting geological features

occur in this region: transitions from thick to thin lithosphere and from thin- to thick-skinned belts, termination

of a Cretaceous rift, a large Cretaceous thermal dome, the nature of the main Andean thrust, and important

economic geological resources.

Near the study area, two large-scale sections across the Eastern Cordillera and Subandean Ranges-Santa

Barbara System (Mon & Salfity, 1995; Drozdzewski & Mon, 1999), a section across the Santa Barbara System

(Cahill et al., 1992), a section across the Subandean Ranges-Santa Barbara System transition (Mingramm et al.,

1979) and a section across the Eastern Cordillera (Rodríguez-Fernández et al., 1999) are available, however,

many unknowns still remain. This work seeks to gain insight into the structural styles of the Eastern Cordillera

and of the transition between the thin-skinned Subandean Ranges and the thick-skinned Santa Barbara System

interfered by the Lomas de Olmedo rift. To achieve these goals we constructed a geological map and three

sections across the study area, merged into a single transect (Fig. 2), using both geological interpretation of

satellite images and field mapping.

Comparison of structural styles

Several evidences support the hypothesis that the Eastern Cordillera remained uplifted in relation to the

Subandean Ranges-Santa Barbara System during large periods of time. The topographic relief is generally higher

in the Eastern Cordillera than in the Subandean Ranges-Santa Barbara System. The Eastern Cordillera moved up

during Andean times along the main Andean thrust that separates it from the Subandean Ranges-Santa Barbara

System located in the dowthrown fault block. The Precambrian top is shallower in the Eastern Cordillera than in

the Subandean Ranges-Santa Barbara System, so that large outcrops of Precambrian and Cambrian rocks

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 401-404

402

predominate in the Eastern Cordillera, whereas Paleozoic, Mesozoic and Cenozoic rocks crop out in the

Subandean Ranges-Santa Barbara System (occasionally, a large mass of Ordovician rocks, cored by Cambrian

and Precambrian rocks, crops out in the Subandean Ranges-Santa Barbara System -between the Valle Grande

and Cianzo synclines- interpreted as the crest of a Cretaceous rollover anticline linked to the Lomas de Olmedo

rift). The upper part of the Paleozoic succession is missing in the Eastern Cordillera due to non deposition,

denudation or both, so that Mesozoic rocks lay on top of a sequence made up of Precambrian, Cambrian and

Ordovician rocks, whereas Mesozoic rocks overlay an almost complete Paleozoic sequence from Precambrian to

Carboniferous-Permian in the Subandean Ranges-Santa Barbara System. This resulted from the uplifted position

of the Eastern Cordillera in the footwall of a normal fault linked to the Cretaceous Lomas de Olmedo rift.

Figure 1. Structural sketch of the north Argentina Andes (after Uliana et al., 1995 modified) with location of the study area and the geological transect. The fault including double triangles corresponds to the main Andean thrust which is the boundary between the Subandean Ranges-Santa Barbara System to the east and the Eastern Cordillera to the west.

The Caimancito anticline and the adjacent inverted structure (Callilegua anticline) both located in the

Subandean Ranges-Santa Barbara System (Fig. 2) suggest the occurrence of a sole thrust. Assuming that the

thrust responsible for these structures is parallel to the west limb (backlimb) of the Callilegua anticline, a

minimum depth to detachment of about 20 km depth is achieved for this portion of the cross section. The large

wavelength folds in the Subandean Ranges-Santa Barbara System is in accordance with the deep detachment

obtained, however, the smaller dimension of the structures in the Eastern Cordillera suggests that another

shallower detachment level may occur in this portion of the Andes. The Cianzo syncline and the adjacent

anticline permits estimating the depth to detachment for this part of the section. Assuming that the main Andean

thrust responsible for these structures is parallel to the west limb (backlimb) of the anticline, a minimum depth to

detachment between 10 and 15 km depth is obtained. The detachment proposed in most published geological

cross sections close to the study area is too shallow to be compatible with the depth obtained here or the

Subandean Ranges-Santa Barbara (except for the Cahill et al., 1992 section based on earthquake data) but agrees

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 401-404

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with the depth proposed for the Eastern Cordillera. Thus, in three cross sections south of our transect, the

detachment is located at a depth: 1) around 25 km in the Santa Barbara System (Cahill et al., 1992), 2) around

10 km in the Santa Barbara System and below 10 km in the Eastern Cordillera (Mon & Salfity, 1995), and 3)

around 15 km in the Eastern Cordillera (Rodríguez-Fernández et al., 1999). In two cross sections north of our

transect, the detachment is located at a depth: 1) around 5 km in the external part of the Subandean Ranges-Santa

Barbara System and around 15 km in the internal part (Mingramm et al., 1979), and 2) slightly more than 5 km

in the Subandean Ranges and slightly less than 10 km in the Eastern Cordillera (Drozdzewski & Mon, 1999).

According to a geological interpretation of 3D seismic survey of an oil field slightly north of our transect,

Masaferro et al. (2003) suggested that faults must detach at a minimum depth of 10 km in the Subandean

Ranges-Santa Barbara System transition. Alternatively, Cahill et al. (1992) propose a solution in which no

detachment occurs beneath the Santa Barbara System. The depths at which earthquakes took place agree in

outline with the detachment depth estimated here for the Subandean Ranges-Santa Barbara System. Thus, some

earthquakes recorded in the studied portion of the Subandean Ranges-Santa Barbara System were estimated to

occur at depths of 20-25 km indicating a nodal plane with a shallow dip to the west that could correspond to a

detachment (e.g., Cahill et al., 1992). According to the USGS earthquake database, a number of earthquakes

occurred along the geological transect took place at a depth of around 30 km both beneath the Subandean

Ranges-Santa Barbara System and the Eastern Cordillera.

Figure 2. Geological transect across the Eastern Cordillera, Subandean Ranges-Santa Barbara System and Lomas de Olmedo Through obtained from merging three cross sections. See Fig. 1 for location.

In the literature, the Subandean Ranges are supposed to have the classical features of a thin-skinned belt,

whereas the Santa Barbara System and the Eastern Cordillera are supposed to be typical thick-skinned belts (e.g.,

Mon, 1976; Mingramm et al., 1979; Allmendinger et al., 1983; Kley et al., 1999). The structural style of these

tectonic units in the study area is more complex than simple thin- and thick-skinned belts because they resulted

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404

from various tectonic events. Thus, some Andean reverse faults correspond to reactivated previous extensional

structures of both Cretaceous and/or Ordovician age, and some Cretaceous features resulted from reactivation of

previous Ordovician faults. The influence exerted by inherited structures in the development of the Subandean

Ranges-Santa Barbara System and of the Eastern Cordillera is crucial because basin inversion varies from

inexistent in some places to strong so that old basins were totally inverted and synrift beds were put into

contraction. The main features of the tectonic units studied here are shown in Table 1.

FOLDS Frequency Wavelength Interlimb

angle

Asymmetry Type

Subandean - Santa Barbara

high

> 10 km

gentle-open

symmetrical east-vergent

fault-bend, fault-propagation folds and others?

Eastern

high

< 5 km

gentle

symmetrical east-vergent west-vergent

fault-bend, fault-propagation folds and others?

FAULTS Frequency Dip Type Age Reactivation Non

reactivated Maximum displacement

Subandean - Santa Barbara

medium

sub-horizontal to steep

reverse normal

Andean Cretaceous

strong (Andean)

no

4 km

Eastern

high

sub-horizontal to steep

reverse normal

Andean Cretaceous Ordovician

mild-strong (Andean Cretaceous)

yes (some Ordovician)

10 km

Table 1. Summary of main structural features of the Subandean Ranges-Santa Barbara System transition and of the Eastern Cordillera. Acknowledgements We acknowledge financial support by projects CGL2006-12415-C03-02/BTE, CSD2006-0041 and CGL2005-02233/BTE funded by the Spanish Ministry for Education and Science.

References Allmendinger, R. W., Ramos, V. A.. Jordan, T. E.. Palma, M. & Isacks, B. L. 1983. Paleogeography and Andean structural

geometry, northwest Argentina. Tectonics 2: 1-16. Cahill, T., Isacks, B. L., Whitman, D., Chatelain, J.-L., Pérez, A. & Chiu, J. M. 1992. Seismicity and tectonics in Jujuy

province, northwestern Argentina. Tectonics 11: 944-959. Drozdzewski, G. & Mon, R. 1999. Oppositely-verging thrusting structures in the North Argentine Andes compared with the

German Variscides. Acta Geológica Hispánica 34: 185-196. Kley, J., Monaldi, C. R. & Salfity, J. A. 1999. Along-strike segmentation of the Andean foreland: causes and consequences.

Tectonophysics 301: 75-94. Masaferro, J.L., Bulnes, M., Poblet, J. & Casson, N. 2003. Kinematic evolution and fracture prediction of the Valle Morado

anticline inferred from 3-D seismic data, Salta province, NW Argentina. AAPG Bulletin 87: 1083-1104. Mingramm, A., Russo, A., Pozzo, A. & Cazau, L. 1979. Sierras Subandinas. In: II Simposio de Geología Regional

Argentina, Córdoba (Argentina), 1979, 1: 95-138. Mon, R. 1976. The structure of the eastern border of the Andes in northwestern Argentin. Geologische Rundschau 75: 211-

222. Mon, R. & Salfity, J. 1995. Tectonic evolution of the Andes of Northern Argentina. In Tankard, A. J., Suárez Soruco, R.,

Welsink, H. J. (ed): Petroleum basins of South America. AAPG Memoir 62: 269-283. Rodríguez Fernández, L., Heredia, N., Seggiaro, R. E. & González, M. A. 1999. Estructura andina de la cordillera oriental en

el área de la Quebrada de Humahuaca, provincia de Jujuy, de Argentina: Trabajos de Geología 21: 321-332. Uliana, M. A., Arteaga, M. E., Legarreta, L., Cerdán, J. J. & Peroni, G. O. 1995. Inversion structures and hydrocarbon

occurrence in Argentina. In Buchanan, J. G., Buchanan, P. G. (ed): Basin inversion. Geological Society Special Publication 88: 211-233.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 405-408

405

Paleomagnetic results from the Antarctic Peninsula and its relation with the Patagonian Andes

Fernando Poblete & César Arriagada

Department of Geology, University of Chile, Santiago, Chile

KEYWORDS : Gondwana, Antarctic Peninsula, South Shetland Island, Patagonia, Permian-Triassic to Cretaceous

Introduction

In the Paleozoic, South America, South Africa and Antarctica were part of Gondwana. By the mid-Cretaceous

South Atlantic opening was under way, and East Antarctica and the Antarctic Peninsula acted as a single plate

(König et al., 2006). The Weddell Sea began to form at about 146 Ma, after rifting between the Antarctic

Peninsula and southernmost South America (Ghidella et al., 2002). Much uncertainty still exists about the

geometrical fit and subsequent drift history between Patagonia and Antarctica. Geophysical and geological data

which describe the tectonic history are sparsely distributed and often of poor quality. During the last two years

we have collected more than 1000 samples (70 sites) for paleomagnetic studies. Here we present the preliminary

results obtained in seven localities (King George Island, Robert Island, Yankee Bay, Half Moon Island, Byers

Peninsula and Snow Island) from the South Shetland Islands and Anderson Island in the northern tip of Antarctic

Peninsula (Fig. 1&2). Our main objective is to provide first-order constraints on latitudinal displacements and

the amount of tectonic rotations as an essential test of published tectonic models.

Fig. 1: Main structural features of the region.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 405-408

406

Paleomagnetic sampling

Within Anderson Island four sites were drilled in Neogene volcanic rocks (Fig. 2). In King George Island

samples are from andesitic dikes and andesitic-basaltic lava flows from the Paleocene-Eocene Fildes Formation

(Smellie et al., 1984) and Upper Cretaceous-Lower Paleocene Martel Inlet Group (Birkenmajer, 2001). Within

Robert Island we sampled several andesitic-basaltic lava flows from Late Cretaceous Coppermine Formation

(Smellie et al., 1984). In Yankee Bay and Snow Island Localities samples were drilled from Late Cretaceous-

Early Tertiary basaltic to andesitic rocks. In Half Moon Island Locality samples were drilled from the Jurassic

Antarctic Peninsula Volcanic Group and from a Late Cretaceous gabbro while in Byers Peninsula samples are

from Anchorage Formation (163 ±16 Ma) and Cerro Negro formation (~119 Ma) (Smellie et al., 1984; Hathway

1997; Hathway and Lomas, 1998). Usually, magnetite is the main magnetic carrier of the magnetization. During

thermal demagnetization, most samples showed univectorial magnetizations going through the origin with

characteristic vectors defined in the range of unblocking temperatures between 310–610º C. A large dispersion

of the paleomagnetic directions is observed in the volcanic rocks of the King George Island locality, however,

both polarities are observed. In all cases, volcanic rocks from the Robert Island locality have a well-defined

normal polarity magnetization. In the Anderson Island locality we observed a well-defined reverse polarity

magnetization. Samples collected in Byers Peninsula shows univectorial magnetizations with inclination smaller

than the observed in the other localities (67°). In Half Moon Island almost samples are well-grouped in in situ

coordinates. In Yankee Bay and Snow Island most samples showed univectorial magnetizations very similar to

that observed in Robert Island locality.

Fig. 2: Paleomagnetic sampling in the South Sheetland Islands.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 405-408

407

Results

The results obtained in Anderson Island suggest no rotation or latitudinal displacement in the last 5 Ma. On the

other hand although the dispersion observed in King George Island could be related with secular variation,

however, some sites seems to be antipodal and we cannot reject the occurrence of a large clockwise rotation (up

to 80°) in this area (Fig. 2&3). For Robert Island we interpret the dominant occurrence of the normal polarity

directions as evidence for acquisition of the magnetization during the Cretaceous Normal Polarity Superchron.

The presence of this polarity agrees with published K-Ar dates which indicate a “mid” Cretaceous age (Smellie

et al., 1984). Here, our results suggest that no latitudinal displacement and rotation occurred since the “mid”

Cretaceous in the middle portion of the Shetland Island.

Fig 3: From left to right: Anderson I. Locality (5 Ma), King George I. Locality (55Ma) and Robert I. Locality (80Ma).

Preliminary results from Yankee Bay and Snow Island localities (Fig. 2) are in good agreement with the

expected direction, however several samples from Yankee Bay show a direction that differ with the expected

one. In in-situ coordinates the volcanic rocks of the Half Moon Island have a similar direction than those

obtained in the gabbroic unit suggesting a Late Cretaceous remagnetization of the volcanic rocks. Observed

inclinations are nearly similar to the expected inclination but up to 10º of discrepancy occur in the declination-

component suggesting a slight counterclockwise rotation. The oldest sampled sequences (Byers Peninsula) show

considerable differences between observed and expected paleomagnetic directions.

Discussion and conclusion

Although preliminary, the differences between observed and expected paleomagnetic direction in the Byers

Peninsula Locality suggest a southward displacement of the Antarctic Peninsula during the Late Jurassic.

However, no major latitudinal displacement occurred since the “mid” Cretaceous times (Robert, Snow,

Anderson) supporting the idea that the Antarctic Peninsula and the Eastern Antarctic were a single plate since

the mid Cretaceous (Barker et al., 2001). In this context if the Antarctic Peninsula was part of the Patagonian

Andes, the break-up should have occurred before ~150Ma. Results obtained in Late Cretaceous units of the Half

Moon Locality suggest a slight counterclockwise rotation while the Tertiary sequences of the northern tip of the

Shetland Island may have been affected by significant clockwise tectonic rotations.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 405-408

408

References Barker, P., 2001. Scotia Sea regional tectonic evolution: implications for mantle flow and palaeocirculation. Earth-Science

Reviews 55 1-39. Birkenmajer, K., 2001. Mesozoic and Cenozoic stratigraphic units in parts of the South Shetland Islands and Northern

Antarctic Peninsula. Studia Geologica Polonica, Vol. 118. Ghidella ME, Yáñez G, LaBrecque JL (2002) Revised tectonic implications for the magnetic anomalies of the Western

Weddell sea, Tectonophysics 347: 65-86. Hathway, B., 1997. Nonmarine sedimentation in an Early Cretaceous extensional continental-margin arc, Byers Peninsula,

Linvingston Island, South Shetland Islands. Journal of Sedimentary Research vol. 67, 686-697 Hathway, B. and Lomas, SA. (1998), The Upper Jurassic-Lower Cretaceous Byers Group, South Shetland Islands,

Antarctica: revised stratigraphy and regional correlations. Cretaceous Research 19, 43-67. König, M., and W. Jokat (2006), The Mesozoic breakup of the Weddell Sea, J. Geophys. Res., 111, B12102,

doi:10.1029/2005JB004035. Smellie J.L., Pankhurst R. J., Thomson M. R. A. and Davies R. E. S., (1984), The Geology of the South Shetland Islands:

VI. Stratigraphy, Geochemistry and Evolution. Britsh Antarctic Survey Scientific Report vol. 87.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 409-412

409

Altiplano-Puna elevation budget and thermal isostasy

Claudia Prezzi1, Hans-Jürgen Götze

2, & Sabine Schmidt

2

1 CONICET - Universidad de Buenos Aires, INGEODAV, Dpto. de Ciencias Geológicas, FCEyN, Universidad de

Buenos Aires, Pabellón 2, Ciudad Universitaria, Buenos Aires 1428, Argentina (prezzi@gl.fcen.uba.ar) 2 Institut für Geowissenschaften, Christian Albrechts Universität zu Kiel, Otto-Hahn Platz 1, 24118, Kiel, Germany

KEYWORDS : Central Andes, Altiplano-Puna, elevation, crustal structure, thermal isostasy

Introduction

The most remarkable feature of the Central Andes is the Altiplano-Puna plateau (Fig. 1). This plateau is

300 km wide, 2000 km long, has an average elevation of 3.5 km and a crustal thickness of approximately 70 km.

Very high heat flow values characterize this portion of the Andean chain (e.g. Springer and Foster 1998).

Furthermore, below the Altiplano-Puna the existence of a partial melting zone at mid-crustal depth (Altiplano-

Puna Magma Body) has been established by a number of independent observations (e.g. Yuan et al. 2000). This

interpretation is strongly supported by the presence of a huge concentration of Neogene ignimbrites: the

Altiplano-Puna Volcanic Complex (De Silva 1989). These features suggest that thermal isostasy could play a

role in the compensation of the Altiplano-Puna. Thermal isostasy is the geodynamic process whereby regional

variations in the lithospheric thermal regime cause changes in elevation. Elevation changes result from variations

in rock density in response to thermal expansion (Hasterok 2005). However the thermal input to continental

elevation is difficult to asses, because variations in crustal density and thickness can mask it. The objective of

this study is to reveal the thermal and/or geodynamic contributions to the elevation of the Altiplano-Puna.

Methodology

The effects on elevation of compositional variations (involving both crustal thickness and density) are removed

through an isostatic adjustment. This adjustment normalizes any crustal column to a crustal standard (Hasterok

2005). We considered a standard crustal thickness of 40 km, a standard crustal density of 2.8 Mg/m3, and a

mantle density of 3.3 Mg/m3. Three parameters must be estimated for each studied point to carry out the

normalization: actual topography, actual crustal thickness and actual crustal density. Actual topography is

obtained from the digital elevation model GTOPO30. Actual crustal densities and actual crustal thicknesses are

derived from the 3D crustal density structure developed for the Central Andes between 19°S and 30°S from 3D

forward gravity modeling (Prezzi et al. 2005). To construct the 3D model we considered 6500 gravity

measurements and used the 3D modelling software IGMAS (Götze et al. 1990). The geometry of our gravity

model is very well constrained by a large amount of geophysical and geological data: seismic reflection and

refraction profiles, receiver function analysis, hypocenter locations, magnetotelluric data, different tomographic

studies, thermal models and numerous structural balanced cross sections (e.g. Yuan et al. 2000). The density

values assigned to the different bodies forming the model were computed based on documented chemical and/or

mineralogical composition (e.g. Lucassen et al. 1999) and information and assumptions about pressure-

temperature conditions expected for each body. We included a partial melting zone at midcrustal depths under

the Altiplano-Puna and took into account the presence of a rheologically strong block beneath the Salar de

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 409-412

410

Atacama basin. We considered the existence of upper, middle and lower crust. In our model the crust is divided

from west to east in different units which represent the Coastal Cordillera, Precordillera, Western Cordillera,

Altiplano-Puna, Eastern Cordillera, Subandean Ranges and Chaco (Prezzi et al. 2005).

Once the elevation was adjusted for compositional buoyancy, a thermal isostatic relationship was estimated in

order to predict the elevation changes expected for different lithospheric thermal states (Hasterok 2005). We

constructed a family of geotherms assuming 1D steady-state temperature conditions, considering exponential

decrease of heat production with depth and constant thermal conductivity. Using as reference the geotherm that

corresponds to a surface heat flow of 40 mW/m2 and assigning a lithosphere having this thermal state an

elevation of 0 km, we can predict the thermal contribution to the actual elevation for different surface heat flow

data. We compiled surface heat flow data for the Altiplano-Puna, including new values recently published (e.g.

Springer and Foster 1998, Hamza et al. 2005) (Fig. 2).

Figure 1. Location map and morphotectonic units. Figure 2. Heat flow data (mW/m2) and depth to the top of the asthenosphere (km) superimposed on shaded relief map.

Results and conclusions

No correlation exists between heat flow values and the corresponding actual topography (Fig. 3). In contrast,

the compositionally adjusted elevation shows direct correlation with heat flow (Fig. 3), and shows a very good

fit with the predicted thermal elevation. While the RMS misfit between compositionally normalized elevation

and predicted thermal elevation is of 0.76 km, the RMS misfit between actual topography and predicted thermal

elevation is of 3.58 km. Our results suggest that while the thermal contribution to the actual topography of the

Altiplano (north of 21.3°S) and the southern extreme of the Puna (27.3-28.7°S) would be of ~ 0.5 km, the

thermal contribution to the actual topography of the southern Puna (24-27°S) would be of ~ 1.3 km. Previous

works highlighted the fact that the Puna has higher elevation than the Altiplano in spite of showing lower

amount of structural shortening and thinner crust (e.g. Allmendinger et al. 1997, Gerbault et al. 2005).

Shortening values are sufficient to account for crustal cross sectional area in the Altiplano north of 22°S, but are

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 409-412

411

less than that needed in the Puna south of 22°S (McQuarrie 2006). Our estimates of the thermal contribution to

the Altiplano and Puna elevation could explain these features. Moreover, our 3D gravity model shows the

presence of deeper asthenosphere below the Altiplano than below the Puna (Fig. 2) (Prezzi et al. 2005).

Particularly, the shallowest asthenosphere is found below the southern Puna (24-27°S), suggesting a possible

relationship between the depth to the top of the asthenosphere and the higher heat flow and the greater thermal

contribution to the elevation. The existence of thinner lithosphere below the southern Puna than below the

northern Puna was previously suggested by other authors (e.g. Whitman et al. 1996). Kay et al. (1994)

documented the eruption of Plio-Quaternary intraplate mafic lavas concentrated around 26°S over thin

continental lithosphere above the central part of a seismic gap in the modern seismic zone. They proposed that

mechanical delamination of a block (or blocks) of continental lithosphere took place during the late Pliocene

below this part of the sourthern Puna. Kay et al. (1994) pointed out that the loss of such lithosphere resulted in

an influx of asthenosphere. These facts coincide with and support our results for the southern Puna.

Unfortunately, there are no surface heat flow data available for the northern Puna (22-24°S), preventing the

evaluation of possible correlations between thermal elevation, asthenospheric depth, the existence of the

Altiplano-Puna Magma Body and of the Altiplano-Puna Volcanic Complex.

Figure 3. Actual topography and compositionally adjusted elevation vs. surface heat flow for the southern Puna (24-27°S), the southern extreme of the Puna (27.3-28.7°S) and the Altiplano (north of 21.3°S). The thermal contribution to elevation predicted for each heat flow value by the thermal isostatic relationship is also shown (Thermal elevation predicted).

With the aim of further validating our results, we compared the estimated thermal contribution to elevation

(normalized elevation) with the residual topography. To calculate the residual topography we assumed that the

Altiplano-Puna is under local isostatic compensation. We used the moho geometry and the densities predicted by

our 3D gravity model (Prezzi et al. 2005) to compute the expected topography considering Airy isostasy (Airy

topography). Then, we obtained the residual elevation by subtracting Airy topography from actual topography.

When we compared the residual topography with the thermal component of the elevation (thermal elevation) a

very good fit is observed (correlation coefficient of 0.98) (Fig. 4) supporting our results. However, the linear

regression parameters (particularly the slope value of 0.80) showed that a portion of the actual topography

(~20%) cannot be explained considering only compositional and thermal effects, suggesting additional

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 409-412

412

geodynamic and/or flexural support. Other authors (e.g. Gerbault et al. 2005) pointed out that horizontal ductile

flow in the lower crust could have occurred along the Central Andes, explaining the observed variations in

crustal shortening between the Altiplano and the Puna. Such a mechanism of channel flow may provide

geodynamic support to the Altiplano-Puna plateau. Regarding flexural support, it has been broadly accepted that

the Altiplano-Puna is locally compensated. Moreover, recent estimates of the Altiplano-Puna elastic thicknesses

(e.g. Tassara et al. 2006) range approximately between 0 and 30 km, indicating the existence of weak lithosphere

and local compensation. The obtained results suggest that the thermal state of the lithosphere would play a

significant role in the elevation of the Central Andes, and may be responsible of some of the geological

differences displayed by the Altiplano and the Puna.

Figure 4. Residual topography vs. thermal elevation (normalized elevation) for the Altiplano-Puna. The regression line is presented in black (correlation coefficient 0.98, slope 0.8).

References Allmendinger, R., Jordan, T., Kay, S., & Isacks, B. 1997. The evolution of the Altiplano-Puna plateau of the Central Andes.

Annual Review of Earth and Planetary Science 25:139-174. De Silva, S. 1989. Altiplano-Puna volcanic complex of the central Andes. Geology 17: 1102–1106. Gerbault, M., Martinod, J., & Hérail, G. 2005. Possible orogeny-parallel lower crustal flow and thickening in the Central

Andes. Tectonophysics 399: 59-72. Götze, H.-J., Lahmeyer, B., Schmidt, S., Strunk, S., & Araneda, M. 1990. Central Andes Gravity Data Base. Eos 71(16):

401-407. Hamza, V., Silva Dias, F., Gomes, A., & Delgadilho Terceros, Z. 2005. Numerical and functional representations of regional

heat flow in South America. Physics of the Earth and Planetary Interiors, 152: 223-256. Hasterok, D. 2005. Thermal isostasy on continents: applications to north America. Thesis Master of Science in Geophysics,

University of Utah, U.S.A., 129 p. Kay, S., Coira, B., & Viramonte, J. 1994. Young mafic back arc volcanic rocks as indicators of continental lithospheric

delamination beneath the Argentine Puna plateau, central Andes. Journal of Geophysical Research 99(B12): 24323-24339. Lucassen, F., Lewerenz, S., Franz, G., Viramonte, J., & Mezger, K. 1999. Metamorphism, isotopic ages and composition of

lower crustal granulite xenoliths from the Cretaceous Salta Rift, Argentina. Contributions to Mineralogy and Petrology 134: 325–341.

McQuarrie, N. 2006. “Revisiting shortening estimates along the Bolivian orocline: implications of thermal heating, erosion and crustal flow on the development of a high elevation plateau”. In Backbone of the Americas Patagonia to Alaska, GSA Specialty Meetings, Abstracts with Programs 2: 86, Mendoza, Argentina, 2006.

Prezzi, C., Götze, H.-J., & Schmidt, S. 2005. “Density structure of the Central Andes from 3D integrated gravity modelling”. In 6th International Symposium on Andean Geodynamics,Extended Abstracts: 574-577, Barcelona, España, 2005.

Springer, M., & Förster, A. 1998. Heat-flow density across the central Andean subduction zone. Tectonophysics 291: 123-139.

Tassara, A., Swain, C., Hackney, R., & Kirby, J. 2006. Elastic thickness structure of South America estimated using wavelets and satellite-derived gravity data. Earth and Planetary Science Letters 253: 17-36.

Withman, D., Isacks, B., & Kay, S. 1996. Lithospheric structure and along-strike segmentation of the Central Andean Plateau: seismic Q, magmatism, flexure, topography and tectonics. Tectonophysics 259 : 29-40.

Yuan, X., Sobolev, S., Kind, R., Oncken, O., & Andes Working Group, 2000. Subduction and collision processes in the Central Andes constrained by converted seismic phases. Nature 408: 958-961.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 413-416

413

Subduction partitioning evidenced by crustal earthquakes along the Chilean Andes

Jorge Quezada & Klaus Bataille

Universidad de Concepción, Casilla 160-C, Concepción, Chile (jquezad@udec.cl, bataille@udec.cl)

KEYWORDS : oblique, subduction, partitioning, crustal, earthquakes

Oblique subduction could produce partitioning or not. In the last case, the deformation is distributed decreasing

from trench to arc (Bevis & Martel, 2001). In the first case there are crustal (upper plate) strike slip faults arc

parallel located in the arc zone that generated a sliver in the forearc (Bevis & Martel, 2001). A good example of

this situation is the Sumatran fault in Indonesia (Fitch, 1972, McCaffrey et al., 2000). In the chilean segment of

central and southern Andes (18,5-46ºS) occurs the subduction of Nazca Plate beneath South American one at

average direction of N77ºE with a convergence velocity between 6,1-7,9 cm/y (DeMets et al., 1994; Tamaki,

1999; Bevis et al., 2001) increasing the magnitude of this velocity southward. The arc has a trend of ~N10ºW

north 22ºS and ~N10ºE south 23ºS (except 31-33ºS, N10ºW), so the subduction is oblique to margin. Along the

chilean Andes, there are two main fault zones that evidences subduction partitioning, the Precordilleran Fault

Zone, PFZ (~20-25ºS) and the Liquiñe-Ofqui Fault Zone, LOFZ (~39-47ºS). Both fault zones had dextral strike

slip activity during the Neogene and Quaternary (Cembrano et al., 2002, Victor et al., 2004, Hoffmann-Rothe et

al., 2006). So a forearc sliver was moved northward. Recent GPS studies (Wang et al., 2007) evidence a

northward sliver movement west of Liquiñe-Ofqui Fault Zone between 42-44ºS.

AFZ

PFZ

LO

FZ

N

Chile Argentina

Bolivia

Peru

0 500 km

V~7 cm/y

2001

1995

1960

2001

2002

1987

2004

2006

1989

2007

1965

2007

Chusmiza

Aroma

Cipreses

Curicó

Ralco

Lonquimay

Ays én

Hudson

Pacific

Oce

an

1985

JFR

1958

Figure 1. Focal mechanisms of crustal earthquakes considered in this study (Hudson are approximate, taken from Lavenu & Cembrano, 1999 and Hoffmann-Rothe et al., 2006). The main faults are indicated. Ellipses shows rupture areas of subduction earthquakes.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 413-416

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The occurrence of subduction partitioning along chilean Andes is evidenced also by crustal earthquakes in the

arc zone with strike-slip focal mechanism. Table 1 summarizes the main crustal earthquakes of strike slip focal

mechanism that occurred in the arc zone in the last half century. Fig. 1 shows the focal mechanisms and location

of these events. All earthquakes are shallow and have focal mechanism indicating high angle ~N/S faults, right

lateral movements. Another earthquake occurred in 1958 (Cajón del Maipo / Las Melosas Earthquake) at

33,83ºS/70,16ºW with a magnitude close to 7. Lomnitz (1958) established for this earthquake right lateral

movement for a N10ºW fault using teleseismic analysis but Pardo & Acevedo (1984) established left lateral

movement for a ~NS fault. Because this contradiction, both papers exhibits strong arguments for their results,

and the fact that this earthquake occurred in the area of the Juan Fernandez Ridge subduction (JFR, Fig. 1) that

can produced local stress changes, so the west-east fault could be the real, we eliminated this earthquake in our

analysis. The Aroma, Chusmiza, Curicó, Ralco and Aysén earthquakes occurred recently. In the case of Aroma

and Chusmiza earthquakes, the aftershocks are aligned N-S (Comte et al., 2003) indicating that this is the

orientation of the fault. Bulletins from Servicio Sismológico of the Departamento de Geofísica of the

Universidad de Chile (www.dgf.uchile.cl) indicate that the aftershocks of Curicó and Aysén earthquakes are

aligned N-S, so this is the orientation of the fault that generated these earthquakes. Hudson, Aysén, Lonquimay

and Ralco earthquakes could be associated to Liquiñe-Ofqui Fault Zone branches. Aroma and Chusmiza

earthquakes are located north of the northern end of the Precordilleran Fault Zone. Farias et al. (2005) related

both earthquakes with strike changes of flexures (associated to west vergent reverse faults) in the Precordillera

(forearc) area. Curicó and Cipreses earthquakes are located north of the Liquiñe-Ofqui Fault Zone. The northern

end of the Liquiñe-Ofqui fault Zone is diffuse but there are some N-S lineaments in the chilean Andes (more

clear in Digital Elevation Models rather Satellite Images) between 35-38ºS that suggest a northward

prolongation of this fault zone or similar structures. Liquiñe-Ofqui Fault Zone evidences subduction partitioning

between 38 to 47ºS; the Ralco, Lonquimay, Aysén and Hudson earthquakes and the GPS studies of Wang et al.

(2007) indicates at least recent movement in segments of Liquiñe-Ofqui-Fault Zone. The Curicó and Cipreses

earthquakes located north of this fault zone, suggests that the subduction partitioning continues northward

Liquiñe-Ofqui Fault Zone at least by 300 km considering also that there are not many geometric, tectonic and

geologic differences along chilean Andes between 34-47ºS. Similar case occurs with Aroma and Chusmiza

earthquakes that could be indicators that subduction partitioning continues north of Precordilleran Fault Zone. A

great number of faults and folds along chilean Andes (forearc and arc zone) suggests the presence of high angle

west vergent reverse faults (Muñoz & Charrier, 1996; Cembrano et al., 2002; Victor et al., 2004; Farias et al.,

2005). These are consequence of normal arc component of the subduction. In the back arc zone in bolivian and

argentinian Andes, the main faults are thrust (low angle east vergent) also due to normal component of the

subduction. The arc parallel component of the subduction is lesser than the normal component. Lavenu &

Cembrano (1999) estimated 2,8 cm/y of arc parallel component of the 7,9 cm/y convergence velocity in the

Liquiñe-ofqui Fault Zone. The small along strike component is enough to generate a forearc sliver. The

Precordilleran and Liquiñe-Ofqui fault zones, the majority of the crustal earthquakes shown in Table 1 and the

GPS studies (Wang et al., 2007) indicates the presence of a sliver, but the distributed deformation (Bevis &

Martel, 2001) in the forearc can not be neglected due to the lack of a dense GPS network that could constrain

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 413-416

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better the deformation and the relatively few amount of crustal earthquakes with focal mechanism known in the

arc and forearc zone of chilean Andes. Both cases of subduction partitioning (Bevis & Martel, 2001) are shown

in Hoffmann-Rothe et al. (2006) and occurs along the arc zone of chilean Andes. Rosenau et al. (2006)

concluded that there is complete partitioning in the south end of Liquiñe-Ofqui Fault, decreasing northward. It

can be considered that there is a continue subduction partitioning along all chilean Andes between 18,5-47ºS but

the magnitude of the earthquakes considered indicates ruptures of segments between 5-20 km, very small in the

context of 3200 km length of chilean Andes considered in this study and there are no clear faults as

Precordilleran and Liquiñe Ofqui Fault Zone in the arc zone of chilean Andes between both. The along strike

heterogeneities, like reological propierties, seismic interplate coupling, age of plates, Juan Fernandez Ridge/flat

subduction (29-34ºS), local buttressing effects, changes in the orientation of the arc (north and south 23ºS, and

Wadatti-Benioff zone, ie. Ghepart symmetry zone ~21ºS) and stress changes due to the occurrence of subduction

interplate thrust earthquakes may influence the degree of partitioning. So this partitioning must be non uniform

along all chilean Andes.

Table 1. Main earthquakes of NS strike, right lateral movement faults along chilean Andes. 1 Chinn & Isacks, 1983.

2

Barrientos & Eisenberg, 1988. 3 Harvard University Centroid Moment Tensors.

The P axis orientation of the earthquakes considered in this study (Table 1) is similar to the convergence

velocity direction between Nazca and South American plates. P axis is the principal shortening axis in a fault so

these similar values indicates that the earthquakes considered in Table 1, are consequence of the subduction

process indicating partitioning.

If we consider the context of these earthquakes in the subduction seismic cycle, one possibility to classify these

events is during the interseismic stage because they are interplate and both plates remains coupled during the

rupture. Coseismic reactivation of strike slip faults during a subduction earthquake must be left lateral like

Atacama Fault Zone (AFZ) and faults located in the Mejillones Peninsula (23,2ºS) during the 1995 Antofagasta

Earthquake (Delouis, et al., 1998) because the extension of the South American Plate during the decoupling

process is in the same direction but opposite sense of convergence (towards SW). Other possibility is that some

of these events are posteismic. In the case of the southern Chile 1960 Mw 9,5 megathrust earthquake, most slip

occurred in the central part of the rupture than in the borders (Barrientos & Ward, 1990), so the Hudson 1965

earthquake located in the southern border of the 1960 may be influenced by the 1960 rupture or change in stress

regime. Similar explanation was done for the Aroma 2001 earthquake (Comte et al., 2003) that followed by few

months the 2001 southern Peru Mw=8,4 subduction earthquake (Fig. 1). Also the Cipreses earthquake is located

Earthquake Date Epicenter

(Lat/Long)

Depth

(km)

Magn.

NP 1

Stk./Dip/Slip

NP 2

Stk./Dip/Slip

P axis

Azm/Plg

T Axis

Azm/Plg

Hudson1 28/11/65 -45,77/-72,9 33 6 Ms

Cipreses2,3

13/9/87 -34,2/-70,15 6,7 5,7 Ms 27/58/176 110/87/32 249/20 348,24

Lonquimay3 24/2/89 -39,2 / -71,83 15 5,3 Mw 9/70/150 110/62/23 61/5 327/35

Aroma3 24/7/01 -19,44/ -69,18 15 6,3 Mw 14/46/-169 276/82/-44 225/36 333/23

Chusmiza3 14/1/02 -19,22/-68,6 38,4 5,6 Mw 13/53/-157 275/80/-37 228/33 329/17

Curicó3 28/8/04 -35,21/-70,36 16 6,5 Mw 21/61/-178 290/88/-29 241/21 339/19

Ralco3 31/12/06 -38,04/-71,4 17,7 5,6 Mw 31/86/178 121/88/4 256/1 346/4

Aysén3 23/1/07 -45,46/-73,07 12,8 5,4 Mw 354/89/-179 264/89/-1 219/2 129/0

Aysén3 3/2/07 -45,51/-73,03 12 5,4 Mw 182/84/-174 91/84/-6 47/8 316/0

Aysén3 23/02/07 -45,51/-73,08 16,6 5,7 Mw 181/79/-160 87/70/-12 46/22 313/6

Aysén3 21/04/07 -45,48/-72,95 12 6,3 Mw 354/88/176 84/86/2 39/1 309/5

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 413-416

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in the southern border of the 1985 Mw 8 central Chile subduction earthquake and occurred two years after this

event in a similar situation of the former cases.

In summary, the arc parallel component of subduction are lesser than normal one in central and southern

Andes, the energy cumulated in this process is enough to generates earthquakes of dextral strike slip focal

mechanism of Mw=6,5 (like Curicó earthquake) along the arc zone of chilean Andes but the existence of

Liquiñe-Ofqui Fault Zone or the Precordilleran Fault Zone indicates the possibility of bigger ruptures and may

be considered in seismic risk analysis.

References Barrientos, S., Eisenberg, A., 1988. Secuencia sismica en la zona cordillerana al interior de Rancagua. In: V Congreso

Geologico Chileno, Santiago, Vol. II: F121-F132. Barrientos, S., Ward, S. 1990. The 1960 Chile earthquake: Inversion for slip distribution from surface deformation.

Geophysical Journal International 103: 589-598. Bevis, M., Martel, S. 2001. Oblique plate convergence and interseismic strain accumulation. Geochem Geophys Geosyst 2:

doi 2000GC000125 Cembrano, J., Lavenu, A., Reynolds, P., Arancibia, G., Lopez, G., Sanhueza, A. 2002. Late Cenozoic transpressional ductile

deformation north of the Nazca–South America– Antarctica triple junction, Tectonophysics 354: 289–314. Chinn, D, Isacks, B. 1983. Accurate source depths and focal mechanisms of shallow earthquakes in western South America

and in the New Hebrides island arc. Tectonics 2(6):529–563. Comte, D., Dorbath, C., Dorbath, L., Farías, M., David, C., Haessler, H., Glass, B., Correa, E., Balmaceda, I., Cruz, A., Ruz,

L. 2003. Distribución temporal y en profundidad de las réplicas del sismo superficial de Aroma, norte de Chile del 24 de Julio de (2001). In: X Congreso Geológico Chileno 2003, Universidad de Concepción, Chile: CD.

Delouis, B., Philip, H., Dorbath, L., Cisternas, A. 1998. Recent crustal deformation in the Antofagasta region (northern Chile) and the subduction process. Geophysical Journal Internacional 132: 302-338.

Demets, C., Gordon, R., Argus, D., Stein, S., 1994. Effect of recent revisions to the geomagnetic reversal time scale on estimates of current plate motions. Geophysical Research Letters 21: 2191-2194.

Farias, M., Charrier, R., Comte, D., Martinod, J., Hérail, G. 2005. Late Cenozoic deformation and uplift of the western flank of the Altiplano: Evidence from the depositional, tectonic, and geomorphologic evolution and shallow seismic activity (northern Chile at 19º30’S). Tectonics 24: TC4001, doi:10.1029/2004TC001667, 2005.

Fitch, T 1972. Plate convergence, transcurrent faults and internal deformation adjacent to southeast Asia and the western Pacific. J. Geophys. Res. 77: 4432–4460.

Hoffmann-Rothe, A., N. Kukowski, N. Dresen, G. Echtler, H., Oncken, O., Klotz, J., Scheuber, E., Kellner, A. 2006. Oblique convergence along the Chilean margin: Partitioning, margin-parallel faulting and force interaction at the plate interface. In Oncken, O. (Springer) Eds: The Andes: Active Subduction Orogeny: 125-146.

Lavenu, A., Cembrano, J., 1999. Compressional and transpressional stress pattern for the Pliocene and Quarternary (Andes of central and southern Chile). Journal of Structural Geology 21:1669– 1691.

Lomnitz, C. 1958. Actividad Sísmica en el Cajón del Maipo. Anuario 1959. Boletín Sismológico para 1958. Universidad de Chile: 31-32.

McCaffrey, R., P. Zwick, Y. Bock, L. Prawirodirdjo, J. Genrich, C. Stevens, S. Puntodewo, C. Subarya. 2000. Strain partitioning during oblique plate convergence in northern Sumatra: Geodetic observations and numerical modelling. J. Geophys. Res. 105: 28363–28375.

Muñoz, N., Charrier, R.,1996. A west vergent fault system at the westem border of the Altiplano in Northem Chile: implications for the uplift of the Altiplano-Puna plateau. Jour. of South American Earth Sciences 9: 171-181.

Pardo, M., Acevedo, A. 1984. Mecanismos de foco en la zona de Chile Central. Tralka 2 (3): 279-293. Rosenau, M., D. Melnick, H. Echtler. 2006. Kinematic constraints on intra-arc shear and strain partitioning in the southern

Andes between 38°S and 42°S latitude. Tectonics 25: TC4013, doi:10.1029/2005TC001943. Tamaki, K., 1999. Nuvel-1A calculation results. Ocean Research Institute, University of Tokyo. http://manbow.ori.u-

tokyo.ac.jp/tamaki-bin/post-nuvella. Victor, P., Oncken, O., Glodny, J. 2004., Uplift of the western Altiplano plateau: Evidence from the Precordillera between

20° and 21°S (northern Chile). Tectonics, 23: TC4004, doi:10.1029/2003TC001519. Wang, K., Hu, Y., Bevis, M., Kendrick, E., Smalley, R, Barriga, R, Lauria, E. 2007. Crustal motion in the zone of the 1960

Chile earthquake: Detangling earthquake-cycle deformation and forearc-sliver translation, Geochem. Geophys. Geosyst. 8: Q10010, doi:10.1029/2007GC001721.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 417-420

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Constraints on delamination from numerical models

Javier Quinteros1, Víctor A. Ramos

1, & Pablo M. Jacovkis

2

1 Lab. de Tectónica Andina – Univ. de Buenos Aires – Pabellón 2, Ciudad Universitaria, 1428 Buenos Aires,

Argentina (jquinte@dc.uba.ar, andes@gl.fcen.uba.ar) 2 Instituto de Cálculo and Depto. de Computación – Univ. de Buenos Aires – Pabellón 2, Ciudad Universitaria,

1428 Buenos Aires, Argentina (jacovkis@dc.uba.ar)

KEYWORDS : Puna, delamination, numerical models, orogenic collapse, isostatic rebound

The main objective for this experiment was to model the effects of eclogitization on the base of an orogen

whose vertical section is similar to the ones near central Andes (southern Puna/northern Argentina), (Oncken et

al., 2006). To study the phenomenon in a proper way, it was isolated from other types of forces that modify the

dynamics of lithosphere, namely compression, extension or thermal anomalies.

Delamination is one of the explanations for the absence of mantle beneath high plateaus and a particular type

of magmatism in the last stage of the orogenic process (Kay and Kay, 1993). Details about delamination

(England et al., 1988) are known from a conceptual point of view, but some of its aspects are quite difficult to

quantify.

The domain studied consists of the lithosphere and asthenosphere up to 150 km depth and 300 km width. The

orogen is considered to be located in the middle of the domain. Crust is 36 km deep far from the orogen and

60 km deep in its axis, similar to the present Puna (Beck et al., 1996). The crust is divided into upper and lower

crust. The underlying mantle is divided into lithosphere and asthenosphere, bounded by the 1250ºC isotherm.

The boundary conditions for the thermal model are: 20ºC over the surface and 1350ºC over the bottom

boundary. On the lateral boundaries free-slip vertical conditions are imposed and horizontal displacements are

not allowed. Vertical displacements over the bottom boundary are also forbidden.

The orogen is about 3 km height after the stabilization time steps, at the moment when the eclogitized root

appears, as it was proposed in the Puna (Oncken et al., 2006).

All the material that suffers the pressure of more than 55 km of crust is considered to transform into an eclogite

during the first 3.5 My, due to the presence of fluids in the system.

The transformation of the crustal root to eclogite causes orogen collapse due to the increment of weight. In the

base of the crust, due to the density difference between the eclogite and the asthenospheric mantle, the former

tries to go down but this is difficult during the first My because it is stuck to the rigid lower crust. It should be

pointed out that the crustal roots are usually distributed in a horizontal direction that can sometimes prevent the

vertical column from having the necessary mass difference to start the detachment process.

One can see in figure 1 how the crustal root evolves during the simulation. This is pushed from the sides by the

acting convective cells, mainly composed of hot asthenosphere.

Laboratory experiments performed by Leech (2001) show that eclogite is much more ductile than the original

rock and thus would suffer a greater deformation as soon as the system turns unstable and the stress increases.

This ductility will be one of the weak points of the domain from a mechanical point of view. The other will be

the contact between eclogite and the lower crust. The downward pressure that the eclogite exerts and the crustal

rigidity turns the contact between them into a low pressure region, and thus unstable. One can see the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 417-420

418

distribution of density and viscosity for the crustal root of the orogen in figure 2. The contact will receive from

the sides a hot incoming mantle flow, due to its low viscosity.

Figure 1. Evolution of the crustal root during the transformation into eclogite and later detachment. The images belong to the evolution at 0.5 My, 3.4 My, 5 My, 6.2 My, 6.7 My and 8.2 My.

The lateral force exerted by the incoming mantle is not only deforming the eclogite, but also introduced into

the contact between the latter and the crust. The exposure of the contact to higher temperatures and the intense

deformation in the zone that increases the strain rate, results in a decreasing viscosity.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 417-420

419

Figure 2. Detail of the contact between the eclogite and lower crust.

After 3.5 My, the income of fluids ends and the transformation from lower crust to eclogite stops. However,

the delamination process evolves until 7/8 My, when the eclogite detaches from the lower crust and sinks into

the asthenosphere.

Figure 3. Evolution of the maximum orogen elevation during the entire process (0-9 My).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 417-420

420

Other process that is tightly associated with delamination is the isostatic rebound. This effect happens after the

detachment of the eclogite from the lower crust. Then, the crust should elevate in order to compensate the losing

of the heavy crustal root. The evolution of topography associated with the different stages of the delamination

process can be seen in figure 3.

It is shown that the maximum orogen elevation reduces continuously since the first eclogite appears. As the

orogen collapses to compensate the extra weight, lower crust is converted to eclogite, due to the tectonic

stacking and the presence of fluids. Once the system is completely out of fluids, it starts to stabilize slowly.

However, the system gets unstable again as soon as the root detaches from the lower crust. At that moment, the

descendant force that was applied at the base of the crust diminishes and the orogen elevates until a new

stabilization equilibrium is reached.

The evolution of the delamination process in this work is in very good agreement with the results from

tomographic inversion performed by Schurr et al. (2006) further to the north. The presence of a high Qp body at

the base of the easternmost Puna crust is interpreted as a detached part of the roots completely delaminated and

resting on top of the Nazca slab.

Also, the possibility to quantify the orogenic collapse and the isostatic rebound for the andean orogen by

means of numerical models can be a powerful tool in order to establish the time span expected at each stage of

the process.

References Beck, S. L., Zandt, G., Myers, S. C., Wallace, T. C., Silver, P. G., & Drake, L. 1996. Crustal-thickness variations in the

central Andes. Geology 24(5): 407-410. England, P. C., Houseman, G. A., Osmaston, M. F., & Ghosh, S. 1988. The mechanics of tibetan plateau. Philosophical

Transactions of the Royal Society of London 326(1589): 301-320. Kay, R. W., & Kay., S. M. 1993. Delamination and delamination magmatism. Tectonophysics 219: 177-189. Leech, M. L. 2001. Arrested orogenic development: eclogitization, delamination, and tectonic collapse. Earth and Planetary

Science Letters 185: 149-159. Oncken, O., Hindle, D., Kley, J., Elger, K., Victor, P., & Schemmann, K. 2006. “Deformation of the central andean plate

system - Facts, fiction, and constraints for plateau models”. In Oncken, O., Chong, G., Franz, G., Giese, P., Götze, H. J., Ramos, V. A., Strecker, M. R., & Wigger, P. (eds.): The Andes - Active subduction orogeny, Berlin-Heidelberg, Springer: 3-27.

Schurr, B., Rietbrock, A., Asch, G., Kind, R., & Oncken, O. 2006. Evidence for lithospheric detachment in the central Andes from local earthquake tomography. Tectonophysics 415: 203-223.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 421-422

421

Magmatic history of the Fitz Roy Plutonic Complex, Southern Patagonia (Argentina)

C. Ramírez de Arellano, B. Putlitz, & O. Müntener

Institut de Minéralogie et Géochimie, Université de Lausanne, Batiment Anthropole CH-1015 Lausanne,

Switzerland (cristobal.ramirez@unil.ch)

KEYWORDS : Patagonia, Fitz Roy, Miocene, magmatism, deformation

The Fitz Roy Plutonic Complex (FRPC) belongs to a chain of isolated Miocene plutons in Southern Patagonia,

which are located in an exotic position between the volcanic arc and the Patagonian plateau basalts It has been

suggested that the intrusion of these plutons is related to the subduction of the Chile ridge (e.g. Michael. 1991).

However, the FRPC is not well studied (Kosmal 1997). Here, we present first results based on field observations

and petrography and speculate about the magmatic evolution of the FRCP.

The FRPC is formed by at least four different magmatic units (phase I to IV): a main central granitoid body

(granitic to tonalitic), which is partially surrounded by a syn-magmatically deformed tonalite. We further

distinguish a mafic series with variably deformed diorites, gabbros and gabbro breccias, and an ultramafic series

with pyroxenites and olivine gabbros. The host rocks are composed of Paleozoic – Mesozoic sedimentary and

volcano-sedimentary sequences.

Contact-metamorphism is characterized by the formation of calc-silicates and by the development of cordierite

and andalusite in pelitic host rock composition. The contact between diorites/gabbros and meta-volcanic host

rocks (rhyolitic to dacitic compositions) is typically formed by garnet-bearing mylonites. Preliminary field

observations suggest that these garnets most probably formed due to partial melting of semi-pelitic host rocks.

The central biotite-hornblende granite displays abundant schlieren structures and is locally rich in miarolitic

cavities. Its contacts with the host rock and other plutonic bodies are steep, sharp and everywhere intrusive. The

tonalite shows beautiful synm-agmatic deformation features and variable microstructures (grain size, flow

textures) related to the distance to the gabbro unit. Gabbro xenoliths within the tonalite demonstrate that the

gabbro complex is relatively older with respect to the tonalite. The ultramafic unit shows a wide variety of

mafic-ultramafic rocks and magmatic textures from more continuous domains (100m size) of layered gabbros

and massive pyroxenite to domains rich in magmatic breccias. At the contact to the diorite the ultamafic unit is

characterized by a wide brecciated zone, with meter-sized angular hornblende gabbros, hornblendite and

gabbroic blocks in a tonalitic matrix. Ductile deformation at the macro and microscopic scale are widespread in

the gabbro unit. Locally, a penetrative foliation with a metamorphic mineralogy developed along shear bands

(amph, bt, fsp, qz) was found, indicating deformation under upper greenschist-lower amphibolite facies

conditions.

Based on our field observations and petrographic criteria we propose that the FRPC was formed by several

magmatic cycles. The oldest magmatic phase is probably represented by the mafic to ultramafic intrusions,

which are brecciated along their margins by intruding gabbroic rocks, which forms the second magmatic pulse

(phase II) dominated by olivine-bearing gabbros. These rocks display little evidence of synmagmatic

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deformation. This third intrusive cycle (phase III) is characterized by gabbros (ol-poor or absent) and diorites,

which are heavily deformed under upper greenschist – lower amphibolite facies conditions. The youngest

granitoid-tonalite phase (phase IV) is largely unaffected by post-intrusive deformation. The granitoid textures

and its intrusive contacts with host tonalite and gabbros demonstrate that the granites postdate the older

ultramafic-mafic tonalite suite. The brecciated nature of the tonalite-gabbro contact suggests a brittle - ductile

environment that is probably defined by the different solidus temperatures of gabbroic and tonalitic rocks. We

speculate that there is a time gap between the ultamafic-mafic and the granitoid-tonalite magmatic cycles.

However, at present the age of the FRPC is determined only by a single K-Ar whole rock (granite) age of 18 ± 3

Ma (Nullo et al., 1978). New age determinations are in progress to test whether our field observations are

resolvable on absolute time scales. Preliminary Ar-Ar- ages on amphibole and biotite separates suggest that the

gabbros and diorites are older than the granites.

References Kosmal A., 1997 — Nuevos aportes a la geología de la zona del Cerro Fitz Roy Departamento Lago Argentino, Provincia de

Santa Cruz. Trabajo final de licenciatura. Departamento de Geología de la Universidad de Buenos Aires, Argentina, 111 p. Michael P.J., 1991 — Intrusion of basaltic magma into a crystallizing granitic magma chamber: The Cordillera Paine pluton

in southern Chile, by in situ fractional crystallization. Contributions to Mineralogy and Petrology, 108: 396-418. Nullo F., Proserpio C., Ramos V. & Rabassa J., 1978 — “Estratigrafía y Tectónica de la vertiente este del hielo Continental

Patagónico, Argentina-Chile” Actas del VII Congreso Geológico Argentino, Neuquén, 1978. Tomo I: 455-470.

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Late Cretaceous synorogenic deposits of the Neuquén Basin (36-39°S): Age constraints from U-Pb dating in detrital zircons

Victor A. Ramos1, Marcio Pimentel

2, & Maisa Tunik

3

1 CONICET & Lab. de Tectónica Andina, Universidad de Buenos Aires, Argentina (andes@gl.fcen.uba.ar)

2 Geochronology Laboratory, Universidade de Brasilia, Brazil (marcio@unb.br)

3 CONICET & Universidad Nacional del Comahue, Neuquén, Argentina (mtunik@jetband.com.ar)

KEYWORDS : Central Andes, uplift, detrital zircons, synorogenic deposits, Cretaceous

Introduction

The inception of orogenic uplift in the Neuquén Basin was subject of intense debate in the last decades. Some

authors favour the initiation of the orogenic shortening and development of the foreland basin stage in the latest

Cretaceous associated with the deposition of the Malargüe Group during Maastrichtian-Paleocene times (Uliana

& Dellapé 1981, Legarreta &Uliana 1991), with the major orogenic uplift during the Miocene time. However,

since the early proposals of Keildel (1921) of the Patagónides orogeny, which was associated by Groeber (1951)

with the intersenonian movements, some authors favour the Late Cretaceous as the main orogenic episode that

developed the Agrio fold-and-thrust belt in the Neuquén Basin (Fig. 1) (Ramos & Folguera 2005). Although a

Late Cretaceous age was generally accepted for the Neuquén Group continental deposits (see Fig. 2), some

authors support depositional systems derived from the eastern foreland. In order to elucidate the provenance of

the Neuquén Group and add some constraints to its age a detrital zircon analysis has been made.

Regional geology and tectonic setting

The Neuquén Basin is a typical multistage retroarc basin developed in the eastern slope of the Andes during

Jurassic to Cenozoic times (Ramos & Folguera 2005). The early sedimentation was controlled by the

paleogeography of the rift systems developed during the Triassic and Early Jurassic (Franzese & Spalleti 2001,

Vergani et al. 1995). It was followed by thermal subsidence associated with the initiation of the subduction

along the Pacific margin, and influenced by changes in the direction and intensity of the convergence vector

(Mosquera & Ramos 2006). As a consequence of these variations several cycles of sedimentation and no-

deposition associated with sea-level changes were recorded in a thermal subsiding retroarc basin, with minor and

local interruptions. A drastic change occurred in the Late Cretaceous when the first continental molasses were

deposited in the basin represented by the Neuquén Group deposits. The ages of these sequences were poorly

constrained by abundant dinosaurs, charophytes, ostracods, and plant remains (Legarreta & Uliana 1999). Figure

2 shows the main units and their accepted ages.

Methodology

Sampling was conducted beneath and above the angular unconformity that separates the Rayoso Formation

from the Neuquén Group. All the samples in the Neuquén Group were taken from the basal units corresponding

to the Candeleros Formation. Sampled stratigraphic horizons are indicated in figure 2.

Zircon separation was carried out at the Geochronology Laboratory of the Universidade de Brasilia. Heavy

mineral concentrates were obtained using conventional gravimetric and magnetic techniques. Final purification

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 423-426

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was achieved by hand picking using a binocular microscope and selected zircons grains were mounted on an

epoxy mount. The mount was polished to obtain an even surface exposing the interior of zircons grains. Prior to

LA-ICPMS analysis, the mounts were cleaned by carefully rinsing with dilute (2%) HNO3. Once fully dry, the

samples were mounted in a laser cell especially adapted for thick sections, and loaded into an UP 213 Nd:YAG

laser ( = 213 mm), linked to a multi-collector, high-resolution Neptune ICPMS.

Helium was used as the carrier gas and mixed with Argon before entering the ICPMS. Normalization was

carried out using GJ-1 standard zircon (608.5 ± 1.5 My; Jackson et al. 2004) and age calculation were performed

using an in-house developed Excel worksheet, based on ISOPLOT V3 formulas. Correction for common Pb

was carried out in samples with 206Pb/204Pb lower the 1000, using Stacey and Kramers model for the age of

crystallization. U-Pb data were plotted using ISOPLOT V3. Errors for isotopic ratios are presented at the 2

level. More than 50 zircon crystals were analyzed from each sample.

Results

Figure 3 shows the probability density of detrital zircons from representative samples indicated in figure 1. As

a general characteristic, it can be seen that there is a striking difference between detrital zircon distributions

above and beneath the unconformity. The zircons from Rayoso Formation (Fig. 3a) show a dominant provenance

of different basement provinces, and are not derived from the magmatic arc. The youngest zircon in this sample

is 188 million years old, and the main peaks are 272 Ma (Choiyoi province); 482 Ma (Famatinian arc), 523-560

Ma (Pampean arc), and 1070 Ma (Grenville ages), and some older ones. On the other hand, samples derived

from the base of Neuquén Group are dominated by a magmatic arc provenance (Figs. 3b, 3c, 3d).

Figure 1. General location of the Neuquén Basin in the provinces of Mendoza, Neuquén and Río Negro, with indication of the thrust front of the Late Cretaceous deformation. There are indicated the sample localities used in this study.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 423-426

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a) b)

c) d)

Figure 3. Relative probability ages from detrital zircons from representative samples of the Rayoso Formation and the Neuquén Group of the Neuquén Basin. See location in the figure 1.

It is worth to note that in all the zircons of the Neuquén Group the main peak is representing ages that vary

from 98.6 to 130 Ma (Fig. 4), which clearly demonstrate that they are derived from the magmatic arc. There are

no Early Cretaceous rocks exposed east of the orogenic front. This striking change in the provenance rules out

the possibility that the sediments of the base of the Neuquén Group could be derived from the foreland. Besides,

in these samples only very few recycled zircons are from the Choiyoi province or the Famatinian belt.

Figure 2. Stratigraphic location of the sampled units above and beneath of the Late Cretaceous unconformity.

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Figure 3. Relative probability ages from detrital zircons from the Neuquén Group showing the Cretaceous and Jurassic peaks, both derived from the magmatic arc. Plot of all GN-samples shown in the figure 1.

Concluding remarks

The age patterns of the detrital zircons analyzed by ICPMS show contrasting features indicating different

sources in their provenances. One population is very similar to the standard pattern of the Early Cretaceous

deposits, where the detrital zircons are mainly derived from the foreland region and therefore matching the age

distribution of the basement provinces. In these sequences only some isolated and rare tuffs are derived from the

magmatic arc (see Aguirre-Urreta et al. 2008). On the other hand, the detrital zircons from the synorogenic

deposits have a conspicuous pattern of Early Cretaceous ages associated with a minor frequency of Jurassic age.

Both Early Cretaceous and Jurassic ages indicate that the main arc was uplifted and under erosion at the time of

deposition of the Neuquén Group. The youngest ages of the Neuquén Group between 98 and 100 Ma constrain

the maximum age of deposition to the base of the Cenomanian as it has been postulated by its fossil content.

Once more the detrital zircons from sedimentary rocks are contributing to date fossils of poor stratigraphic value.

References Aguirre-Urreta, M.B., Pazos, P.J., Lazo, D.G., Fanning C.M., Litvak, V.D., 2008. First U-Pb SHRIMP age of the Hauterivian

stage, Neuquén Basin, Argentina. Journal of South American Earth Sciences (in press, on line). Franzese, J.R., Spalletti, L.A., 2001. Late Triassic continental extension in southwestern Gondwana: tectonic segmentation

and pre-break-up rifting. Journal of South American Earth Sciences, 14: 257-270. Groeber, P., 1951. La Alta Cordillera entre las latitudes 34° y 29°30'. Instituto Investigaciones de las Ciencias Naturales.

Revista Museo Argentino de Ciencias Naturales Bernardino Rivadavia, (Ciencias Geológicas) 1(5): 1-352. Jackson, S.E., Pearson, N.J., Griffin, W.L., Belousova, E.A., 2004. The application of laser ablation inductively coupled

plasma mass spectrometry to in situ U-Pb zircon geochronology. Chemical Geology 211: 47-69. Keidel, J., 1921. Sobre la distribución de los depósitos glaciares del Pérmico conocidos en la Argentina y su significa¬ción

para la estratigrafía de la serie del Gondwana y la paleogeografía del Hemisferio Austral. Academia Nacional de Ciencias, Boletín 25: 239 368.

Legarreta L., Uliana, M.A., 1991. Jurassic-Cretaceous marine oscillations and geometry of back-arc basin fill, central Argentine Andes. International Association of Sedimentology, Special Publication 12: 429-450.

Legarreta, L., Uliana, M.A., 1999. El Jurásico y Cretácico de la Cordillera Principal y la Cuenca Neuquina. In R. Caminos (ed.) Geología Argentina, Instituto de Geología y Recursos Minerales, Anales 29(3): 399-416.

Mosquera, A., Ramos, V.A., 2006. Intraplate deformation in the Neuquén Basin. In Kay, S.M., Ramos, V.A. (eds.) Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S latitude). Geological Society of America, Special Paper 407: 97-124.

Ramos V.A., Folguera, A. 2005. Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc and foreland deformation In G. Veiga et al. (eds.) The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics. The Geological Society, Special Publication 252: 15-35.

Uliana, M.A, Dellapé, D.A., 1981. Estratigrafía y evolución paleoambiental de la sucesión Maastrichtiana-eoterciaria del engolfamiento Neuquino (Patagonia Septentrional). 8° Congreso Geológico Argentino (San Luis), Actas 3: 673-711.

Vergani, G., Tankard, A.J. Belotti H.J., Welsnik, H.J. 1995. Tectonic Evolution and Paleogeography of the Neuquén basin. In Tankard, A.J., Suárez Sorucco, R., Welsnik, H.J. (eds.) Petroleum Basins of South America. AAPG Memoir 62: 383-402.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 427-430

427

Revisiting accretionary history and magma sources in the Southern Andes: Time variation of “typical Andean granites”

C.W. Rapela1, R.J. Pankhurst

2, J.A. Dahlquist

3, E.G. Baldo

3, C. Casquet

4, & C. Galindo

4

1

CIG-CONICET-UNLP, Calle 1 Nº 644, 1900, La Plata, Argentina (crapela@cig.museo.unlp.edu.ar) 2

British Geological Survey, Keyworth, Nottingham NG12 5GG, UK (rjpankhurst@gmail.com) 3 CICTERRA-CONICET-UNC, Av. Vélez Sarsfield 1611, X5016CGA-Córdoba, Argentina

(jdahlquist@efn.uncor.edu). 4

Dep. Petrol. y Geoquím., Universidad Complutense, 28040 Madrid, Spain (casquet@geo.ucm.es)

KEYWORDS : Southern Andes, Andean granites, subduction, continental accretion

Southern Andes: accretionary history of the basement blocks

The composition and distribution of Andean magmas are strongly influenced by the age and extent of the

different basement blocks beneath the modern Andes. In particular, radical revision of southwestern Gonwana

assembly models depicted in recent studies has to be taken into account when considering the variation with time

of the early pre-Andean and Andean subduction-related granitic magmas (Fig. 1): (i) Palaeomagnetic studies

indicate that one of the most important assembly episodes occurred during the Pampean-Araguaia collisional

orogeny (540-520 Ma), between an Amazonia craton group and the West Africa, Congo-São Francisco, Paraná

and Río de la Plata cratons (Trindade et al., 2006 and references therein); (ii) The Amazonia craton group

included the Arequipa-Antofalla and Western Sierras Pampeanas basement blocks (Rapela et al., 2007), for

which a common metamorphic and magmatic history has been established (Casquet et al., 2006; 2008); (iii)

Further evidence shows that the large Neoproterozoic turbiditic sequence of the Eastern Sierras Pampeanas

(Pampean belt), now bounded to the east by the Palaeoproterozoic Rio de la Plata craton, is a transcurrent terrane

resulting from right-lateral movements along the SW Gondwana margin (Rapela et al., 2007). This dextral

displacement was associated with the oblique collision of the Western Sierras Pampeanas during the Pampean–

Araguaia orogeny, following closure of the intervening Clymene ocean (Fig. 2); (iv) South

Figure 1. Group of cratons and minor blocks amalgamated in the Pampean-Araguaia orogeny (540-515 Ma) (modified from Trindade et al., 2006 and Rapela et al., 2007). TB = Transbrasiliano shear zone.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 427-430

428

of 36ºS, the last recognised accretion to the already assembled Gondwana took place during mid-Carboniferous

time, when Early Carboniferous I-type granites representing a subduction-related magmatic arc was followed by

a collision between continental areas typified by the Deseado and North Patagonian massifs (Pankhurst et al.,

2006).

Figure 2. Precambrian and Early Palaeozoic terrains in the Southern Andes, disclosed by back-thrusting above Miocene “flat-slab” subduction (28º - 33ºS).

Subduction associated with pre-Andean episodes was related to Wilson cycles of ocean opening and closing,

where the final event is either a continent-continent collision or large-scale back-arc closure. At 30º-34ºS, three

main episodes of pre-Andean plate convergence are well established: (1) Pampean: 540-528 Ma subduction,

followed by oblique continent-continent collision at 528-515 Ma. The supercontinent grew westwards by lateral

accretion of the Western Sierras Pampeanas Grenvillian block (Rapela et al., 2007), including the Precordillera

(Fig.2). (2) Famatinian: 484-463 Ma convergent episode associated with the opening and closing of a large back-

arc basin in Early to Mid Ordovician times (Pankhurst et al., 2000; Dahlquist et al., 2008) (Fig.2). (3)

Gondwanan: 320-190 Ma. After the intrusion of Devonian and Early Carboniferous (c. 380 and 340 Ma) intra-

plate A-type granites in the Sierras Pampeanas, a new subduction regime started along the palaeo-Pacific margin

in Late Carboniferous times (c. 320 Ma), which included younger pulses (Parada et al., 1999). At 33ºS the Late

Palaeozoic batholiths occur both in the coast range of Chile and in the Frontal Cordillera, suggesting that no

major continental accretion took place after the collision of the Western Sierras Pampeanas and associated

Grenvillian blocks.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 427-430

429

Isotopic and chemical variations of the subduction-related granites

The isotopic and chemical characteristics of the granites emplaced in the different pre-Andean episodes

described above are compared with the typical Andean I-type granites emplaced near the continental margin

after the break-up of Gondwana. The latter are mostly Cretaceous in age at 30º-34ºS, but a complete record from

185 Ma to Tertiary is exposed in the Patagonian batholith and subcordilleran belts (Rapela et al., 2005, Hervé et

al., 2007). The Pampean and Famatinian rocks have chemical and isotopic characteristics that contrast with the

younger Andean bodies. The older pre-Andean rocks show a wide silica range and, although metaluminous I-

type varieties from gabbro to granodiorites are abundant, cordierite-bearing S-type granites are also conspicuous.

S-type granites are rare in the Carboniferous and Andean granites, indicating that melting of sedimentary

material was not common in these episodes. The Ndt values decreases with time, suggesting derivation from

progressively more primitive and depleted sources. Only the younger Gondwanan and the majority of the

Andean granites plot in the “mantle array” of the (87Sr/86Sr)0 – Ndt isotopic diagram, in contrast to the

Palaeozoic granites, most of which lie outside the mantle field, with Ndt < -2 (Fig. 3). This is a remarkably

consistent feature of the Pampean and Famatinian events, as they include abundant amphibole-bearing and

noritic gabbros with less than 50% SiO2 that share the same crustal signature as the intermediate rocks. As there

is no evidence for massive in situ contamination during emplacement in the upper crust, this signature must

reflect the composition of the middle or lower crust (Pankhurst et al. 1998). Depleted mantle model ages (TDM)

for most of the Cambrian and Ordovician rocks, both I- and S-types, are in the interval 1400–1700 Ma indicating

involvement of Palaeo- to Mesoproterozoic sources. Altogether the chemical and isotopic evidence suggests that

the Pampean and Famatinian episodes did not involve significant recycling of young underplated material.

Rather, it indicates melting of an old crustal section, including the underlying subcontinental mantle, to produce

the basic rocks with enriched isotopic signatures. Although isotopically less evolved than the Cambrian–

Ordovician granites, the Carboniferous coastal batholiths of Chile also plot off the “mantle array”, but with

younger (mostly Neoproterozoic) model ages. Recycling of the immature 1000–1200 Ma juvenile Grenvillian

lithosphere in which they are emplaced seems to fit the source isotopic constraints. Only the Andean and

younger Gondwanan granites show depleted signatures (Parada et al. 1999): this is not only a characteristic of

central Chile but also in the Patagonian Andes (Pankhurst et al. 1999, Rapela et al. 2005, Hervé et al., 2007).

Figure 3. Variation of Ndt versus initial 87Sr/86Sr for the granitic rocks emplaced during the main convergence episodes in the Andean sector at 28º- 33ºS (a) and Patagonia (b). Data sources are Pankhurst et al., 1999; 2000, Parada et al., 1999, Rapela et al., 2005, Hervé et al., 2007 and references therein.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 427-430

430

Remarkably, the Andean granites also show a change in isotopic composition with time. For example the Sm–

Nd relationships of granitoids from the Patagonian batholith at 44º–46ºS indicate source compositions that

change from slightly LIL-enriched for the Jurassic and Early Cretaceous rocks, to significantly depleted for the

Late Cretaceous to Early Miocene plutons ( Ndt values between +4 and +6), the latter in turn very similar to

those of the Tertiary to Recent mafic strato-volcanoes of the Southern Volcanic Zones (Pankhurst et al. 1999)

(Fig. 3). This cannot be explained by upper or lower crustal contamination and it has been suggested that melting

occurs in progressively more LIL-depleted mantle sources underlying the Patagonian batholith (Rapela et al.,

2005). The obvious conclusion is that under the label of “typical I-type Andean granite” there is a wide range of

isotopic compositions that shows a general variation towards more depleted mantle sources with time. Isotopic

equivalents of modern Andean granites are uncommon or absent in the extensive Early Paleozoic metaluminous

suites.

References Casquet, C., Pankhurst, R.J., Fanning, C.M., Baldo, E., Galindo, C., Rapela, C.W., González-Casado, J.M. & Dahlquist, J.A.,

2006. U–Pb SHRIMP zircon dating of Grenvillian metamorphism in Western Sierras Pampeanas (Argentina): correlation with the Arequipa Antofalla craton and constraints on the extent of the Precordillera Terrane. Gondwana Research 9: 524-529.

Casquet, C., Pankhurst, R.J., Rapela, C.W., Galindo, C., Fanning, C.M. , Chiaradia, M., Baldo, E., González-Casado, J.M. & Dahlquist, J. A., 2008. The Mesoproterozoic Maz terrane in the Western Sierras Pampeanas, Argentina, equivalent to the Arequipa–Antofalla block of southern Peru? Implications for West Gondwana margin evolution. Gondwana Research 13: 163-175.

Cordani, U.G., D’Agrella-Filho, M.S., Brito-Neves, B.B. & Trindade, R.I.F., 2003. Tearing up Rodinia: the Neoproterozoic palaeogeography of South American cratonic fragments. Terra Nova 15: 350-359.

Dahlquist, J. A., Pankhurst, R. J. , Rapela, C. W., Galindo, C., Alasino, P., Fanning, C. M., Saavedra, J. & Baldo, E. , 2008. New SHRIMP U-Pb data from the Famatina Complex: constraining Early–mid Ordovician famatinian magmatism in the Sierras Pampeanas, Argentina. Geologica Acta (in press).

Hervé, F., Pankhurst, R.J., Fanning, C.M., Calderón, M. & Yaxley, G.M. 2007. The South Patagonian batholith: 150 my of granite magmatism on a static plate margin. Lithos 97: 373-394.

Pankhurst, R.J., Rapela, C.W.& Fanning, C.M., 2000. Age and origin of coeval TTG, I- and S-type granites in the Famatinian belt of NW Argentina. Transactions of the Royal Society of Edinburgh: Earth Sciences 91: 151-168.

Pankhurst, R. J., Rapela, C. W., Saavedra, J., Baldo, E., Dahlquist, J., Pascua, I. & Fanning, C. M. 1998. “The Famatinian magmatic arc in the central Sierras Pampeanas”. In Pankhurst, R. J. & Rapela, C. W. (eds.): The Proto-Andean margin of South America. Geological Society (London) Special Publication 142: 343-367.

Pankhurst, R.J., Weaver, S.D., Hervé, F. & Larrondo, P., 1999. Mesozoic–Cenozoic evolution of the North Patagonian Batholith in Aysén, southern Chile. Journal of the Geological Society, London 156: 673-694.

Pankhurst, R.J., Rapela, C.W., Fanning, C.M. & Márquez, M., 2006. Gondwanide continental collision and the origin of Patagonia. Earth Science Reviews 76: 235-257.

Parada, M.A., Nyström, J.O. & Levi, B., 1999. Multiple sources for the Coastal Batholith of central Chile (31-34º): geochemical and Sr-Nd isotopic evidence and tectonic implications. Lithos 46: 505-521.

Rapela, C.W., 2000. “Accretionary history and magma sources in the Southern Andes”. 31st. t International Geological Congress, Rio 2000, Special Simposia F-2 "Andean Tectonics and Magmatism" -(LP). Abstract Volume (CD-ROM), 4p. 2figs.

Rapela, C.W., Pankhurst, R.J., Fanning, C.M. & Hervé, F., 2005. “Pacific subduction coeval with the Karoo mantle plume: the Early Jurassic Subcordilleran Belt of northwestern Patagonia”. In Vaughan, A. P. M., Leat, P. T. & Pankhurst, R. J. (eds.): Terrane Accretion Processes at the Pacific Margin of Gondwana. Geological Society (London) Special Publication 246: 217-239.

Rapela, C.W. , Pankhurst, R.J., Casquet, C., Fanning, C.M., Baldo, E.G., González-Casado, J.M., Galindo, C. & Dahlquist, J., 2007. The Río de la Plata craton and the assembly of SW Gondwana. Earth Science Reviews 83: 49-82.

Trindade, R.I.F., D´Agrella-Filho, M.S., Epof, I. & Brito Neves, B.B., 2006. Paleomagnetism of Early Cambrian Itabaiana mafia dikes (NE Brazil) and the final assembly of Gondwana. Earth and Planetary Science Letters 244: 361-377.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 431-434

431

Recent debris-flows and megaturbidite in a confined basin of the North Ecuador subduction trench

G. Ratzov, J.-Y. Collot, M. Sosson, & S. Migeon

Geosciences Azur: UNSA – CNRS – IRD – UPMC. BP 48 port de la Darse, 06230 Villefranche-sur-Mer, France

(ratzov@geoazur.obs-vlfr.fr, collot@geoazur.obs-vlfr.fr, sosson@geoazur.unice.fr, migeon@geoazur.obs-vlfr.fr)

KEYWORDS : trench, debris flow, megaturbdite, Ecuador

Introduction

Mass wasting plays a major role on sediment transport and distribution, and is largely responsible for shaping

the seafloor of both deep sea and coastal environments. Multiple processes have been proposed for slope failures

and associated deposits, ranging from rock falls to turbidity currents [Mulder and Cochonat, 1996]. Factors

promoting submarine landslides include rapid sediment accumulation, slope increase, excess pore pressure,

eustatic sea-level variations, and tectonics and earthquakes [Hampton, et al., 1996].

Tectonics and earthquakes are of particular importance along active margins where slope failures deposits can

be used as a marker for tectonic activity [Goldfinger et al., 2003]. At subduction zones, the nature and physical

properties of sediment dragged into the subduction play a major role on inter plate frictional conditions

[Calahorrano, et al., 2008] and thus on upper plate erosion, accretion, and earthquake rupture propagation. The

objectives of this study are 1) to discriminate mass wasting and turbidite deposits into the trench, 2) constrain

their extent, 3) estimate their time recurrence and 4) determine their origin, in order to better constrain the nature

of sediments entering subduction.

To achieve these objective, the Amadeus cruise conducted onboard of the R/V L’Atalante in 2005 acquired

new multibeam bathymetry data (150m resolution), 3-5kHz Chirp high resolution lines, 6 channels multichannel

seismic data, as well as gravity core. Only bathymetric and Chirp data will be presented here.

Geological setting

The north Ecuador / south Colombia active margin is located along Northwestern South America, where the

Nazca plate underthrusts eastward the South America plate (Fig.1) with a 58mm/year convergence rate

[Trenkamp, et al., 2002]. Between the Carnegie Ridge at latitude 0° and the Galera Seamounts at latitude

1°30’N, the trench is poorly sedimented and shows numerous topographic asperities such as fault scarps and

seamounts. In this region, the margin undergoes tectonic erosion [Collot, et al., 2002] and shows a small frontal

wedge. The area is seismically active and has been submitted to major subduction earthquakes in 1906

(Mw=8.8) and 1942 (Mw=7.7)[Beck and Ruff, 1984]. Onshore, the Coastal cordillera undergoes active uplift

probably since 1.1 Ma [Pedoja, et al., 2006]. Coastal cordillera’s uplift has caused migration of the Andean

drainage system northward to Esmeraldas, and southward to the Gulf of Guayaquil so that, in the study area

sediment supply to the trench is limited.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 431-434

432

Results and interpretations

Trench Morphology

Two series of structural lineations control the trench in the study area (Fig 2). NNE – SSW-trending, linear

scarps parallel to the trench are interpreted as normal faults related to bending of the plate, and WSW- ENE

trending lineaments reflect the structural grain of the oceanic plate. The later lineaments form a series of small

relief, structural ridges (SR), which delineate sub-basins in the trench. The most spectacular is the rectangular

Galera Basin, bounded northward by the 1200-m-high Galera Seamount, and by SR1, southward. Numerous

arcuate scarps affect both the toe of the continental slope and the basement highs in the trench. These scarps may

be scars related to slope instabilities.

Sedimentary units

The trench sedimentary fill reveals six main seismic units in the Galera basin (Fig.3)

- H1 and H2 have a well-layered continuous and horizontal medium to high amplitude reflection facies.

- U1df and U2df are both characterised by a semi transparent low amplitude chaotic facies with a hummocky

top surface. U2df exhibits irregular upper and lower surfaces, the lower being clearly erosive on unit H2.

- U2mt exhibits transparent facies with regular and sharp upper and lower boundaries. The base of U2mt is

outlined by a high amplitude reflection.

- U3 shows transparent facies and seems concordant with both underlying and overlaying stratas.

Core data collected in the southern part of the study area (KAMA03) reveal that seismic facies of unit H1 and

H2 are associated with hemipelagic stratified deposits with few turbidites. No sedimentological data is available

on units U1df to U3. Units U1df and U2df are interpreted as debris flows according to their chaotic semi

transparent facies, irregular boundaries, and erosional base similar to examples described offshore the Iberian

Peninsula by [Lastras, et al., 2004]. U2mt slightly differ from Units U1df and U2df, but has the same

transparent facies, regular boundaries, and high amplitude basal reflection as megaturbidites identified in the

eastern Mediterranean sea [Rebesco, et al., 2000].

Spatial organisation

The debris flows and the megaturbidite are trapped in trench structural basins controlled by the ENE-trending

ridges and NNE-trending faults (Fig4): U1df deposited exclusively within Galera Basin, with thickness ranging

from 2.5 meters in the center of the basin to 7.5 meters along its southern boundary. SR1 acts as a boundary for

U2df and U2mt: U2df is only 7 meters thick southward of SR1, compared to its maximum 45-m thickness in the

Galera basin. U2mt is only present southward from SR1, and the thickness of U2mt decreases abruptly across

each SR (15 m along SR1, 8m and 2m respectively northward and southward of SR2). U2mt is also bounded

westward by the NNE-trending normal faults.

Age of units

In absence of absolute dating we estimated the ages of the mass wasting events based on the geometry of the

trench fill and convergence rate [Mountney and Westbrook, 1997]. Unit H1 mean sedimentation rate was

estimated to be 4 mm/yr. U1df outcrops at the seafloor, and is consequently considered the most recent event.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 431-434

433

No hemipelagic deposits were identified between U2df and U2mt suggesting that U2 megaturbidite is part of the

same event as the U2 debris flow. A 3m-thick H1 layer overlies U2mt. According to the above sedimentation

rate, U2df and U2mt would have deposited ~750 years ago, and U3 ~7000 years ago.

Discussion and conclusions

The structural segmentation of the trench basins and the thickness distribution of the debris flows allow to

discriminating the geographic origin of the flows. U1df’s greatest thickness in the Galera basin is along the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 431-434

434

northern flank of the SR1 suggesting U1df derived from SR1. A small package of U1df deposited on the NW

flank of the Galera basin suggests a contribution of the Galera seamount sedimentary cover to U1df. The

geometry of scours at the base of U2df supports a westward transport, indicating that U2df, and likely U2mt,

derived from the margin toe. However, differences in seismic facies indicate that the U2 debris flow deposited

dominantly in the Galera basin, whilst the more volatile component of the flow have overflown SR1 and

deposited further south in the form of the U2 megaturbidite (Fig4). Alternatively multiple sources located along

the toe of the margin may also account for the observations: the debris flow would have originated from a

northern source, whereas a southern source would be responsible for the contemporaneous megaturbidite. The

origin of U3 is not constrained.

A trigger factor cannot be doubtlessly established. Sediment overloading is unlike to be of importance as

sediment supply is limited. Slope oversteepening and fracturation of underlying rocks may be considered as

facilitating factors because the coast undergoes active uplift and seamount subduction promotes margin basal

erosion. Earthquake is certainly the main trigger factor along such a seismically active margin. The hypothesis of

multiple contemporary sources along the margin for units U2df and U2mt acts in favour of an earthquake

trigger. The time recurrence of the described events is not regular (750 and 6250 years time intervals). If they are

seismically triggered, the slope failures may be related to exceptional earthquakes.

The total thickness of mass wasting deposits identified in our chirp profiles represents ~40% of the total

deposit in the trench and ~65% in the Galera basin indicating a relatively large component of mass wasting

deposits entering subduction. Physical properties of mass wasting deposits (porosity, fluid overpressure…) are

considered different from those of the hemipelagic / turbiditic trench fill. Such differences may be of particular

importance for tectonic accretion, basal erosion, and variations of the interplate coupling.

References Beck, S. L., & L. J. Ruff (1984), The rupture process of the great 1979 Colombia earthquake: evidence for the asperity

model, Journal of Geophysical Research, 89, 9281-9291. Calahorrano, A. B., et al. (2008), Nonlinear variations of the physical properties along the southern Ecuador subduction

channel: Results from depth-migrated seismic data, Earth Planet Sc Lett, 267, 453-467. Collot, J.-Y., et al. (2002), Exploring the Ecuador-Colombia active margin and interplate seismogenic zone, EOS

Transactions, American Geophysical Union, 83, 189-190. Collot, J.-Y., et al. (2006), Mapas del margen continental del Norte de Ecuador y del Suroeste de Colombia : Batimetría,

Releive, Reflectividad Acústica e Interpretación Geológica, publicación IOA-CVM-03-Post. Goldfinger, C., et al. (2003), Holocene earthquake records from the Cascadia subduction zone and northern San Andreas

Fault based on precise dating of offshore turbidites, Annu Rev Earth Pl Sc, 31, 555-577. Hampton, M. A., et al. (1996), Submarine Landslides, Review of Geophysics, 34, 33-59. Lastras, G., et al. (2004), Characterisation of the recent BIG'95 debris flow deposit on the Ebro margin, Western

Mediterranean Sea, after a variety of seismic reflection data, Marine Geology, 213, 235-255. Mountney, N. P., and G. K. Westbrook (1997), Quantitative analysis of Miocene to Recent forearc basin evolution along the

Colombian convergent margin, Basin Research, 9, 177-196. Mulder, T., & P. Cochonat (1996), Classification of offshore mass movements, Journal of Sedimentary Research, 66, 43-57. Pedoja, K., et al. (2006), Plio-Quaternary uplift of the Manta Peninsula and La Plata Island and the subduction of the

Carnegie Ridge, central coast of Ecuador, J S Am Earth Sci, 22, 1-21. Rebesco, M., et al. (2000), Acoustic facies of Holocene megaturbidites in the Eastern Mediterranean, Sediment Geol, 135,

65-74. Trenkamp, R., et al. (2002), Wide plate margin deformation, southern Central America and northwestern South America,

CASA GPS observations, J S Am Earth Sci, 15, 157-171.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 435-438

435

Geomorphology of the Fitzcarrald Arch, Peru, and its relationships with the Nazca plate subduction

V. Regard1, R. Lagnous

1, N. Espurt

1,*, J. Darrozes

1, P. Baby

1,2, M. Roddaz

1, Y. Calderon

2, &

W. Hermoza2,$

1 LMTG, Université de Toulouse-CNRS-IRD-OMP, 14 av. E. Belin 31400 Toulouse, France (regard@lmtg.obs-

mip.fr, renaud.lagnous@wanadoo.fr, nicolas.espurt@ifp.fr, darrozes@lmtg.obs-mip.fr, Patrice.Baby@ird.fr,

mroddaz@lmtg.obs-mip.fr) 2 PerùPetro SA, Lima, Peru (ycalderon@perupetro.com.pe, whermoza@perupetro.com.pe)

* Now at IFP, Rueil-Malmaison, France $ Now at Repsol, Madrid, Spain

KEYWORDS : geomorphology, hypsometry, ridge subduction, Nazca Ridge, Peru

Figure 1. Geodynamic setting of the Peruvian Andes and its associated Amazonian foreland basin (taken from Espurt, Baby et al. 2007). The flat slab segment is illustrated by isodepth contours of Wadati-Benioff zone (Gutscher, Olivet et al. 1999), and plate convergence vector is from NUVEL1A plate kinematics model (DeMets, Gordon et al. 1990). The western part of the Amazon basin consists of two main subsiding basins —the northern Amazonian foreland basin (NAFB or Marañon-Ucayali basin) and the southern Amazonian foreland basin (SAFB or Beni-Mamore basin) — separated by the antiformal Fitzcarrald Arch. This arch is superimposed on the present-day reconstruction of the subducted part of the Nazca Ridge (black line, Hampel 2002). The ridge reconstruction at 11.2 Ma is shown (white line, Hampel 2002). The easternmost edge of the Nazca Ridge is not involved in the flat slab; it is brought by the sinking slab: its projection at surface may differ from the reconstruction represented by the dotted line. The empty rectangle indicates the study area covered by next figures. The forebulge is located after the works of Aalto et al. (2003) in the SAFB and Roddaz et al. (2005) as the Iquitos Arch in the NAFB.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 435-438

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Introduction

In Peru, the Fitzcarrald Arch constitutes a major geomorphic feature spreading more than 400 000 km2 in the

Amazon basin, extending from southern Peru to western Brazil (fig. 1) (Espurt, Baby et al. 2007). It lies in front

of the subducting Nazca ridge, which is supposed to have important effects on geodynamics and superficial

tectonics (Hampel 2002; Saillard, Audin et al. 2008). This Arch is characterized by a radial, but dissymmetrical,

drainage network, originating at a point close to main Andean river basin outlets, which argues for simple mega-

fan geometry. Other hypotheses were proposed recently for the Arch formation. First, Jacques (2003) related it

to a ENE-trending transfer zone, probably inherited from old structures affecting the overall lithosphere, and

crossing the entire continent. Second, the buoyant Nazca Ridge is thought to have underplated the South-

American lithosphere resulting in a regional uplift (fig. 1). Hampel (2002) reconstructed a likely shape and

position of Nazca Ridge subducted part which matches pretty well the actual arch shape. These results lead

Dumont (1996) and Espurt et al. (2007) to hypothesize a link between flat subduction and the Fitzcarrald Arch.

To decipher between these hypotheses about the Fitzcarrald Arch our study aims at exploring “classical”

quantitative geomorphology, by using indicators of basin maturity, such as hypsometry or basin shape. Multiple

indicators of basin maturity are used to ensure reliable conclusions.

The results presented here are currently in review for publication in Geomorphology.

Figure 2. Fitzcarrald Arch 7th-order basins’ hypsometric integrals (in percent), with contours for I=35% and I=50%.

Data and processing

The area was studied using Shuttle Radar Topography Mission (SRTM) Digital Elevation Model (DEM).

River networks were extracted and classified according to Strahler’s method (1952); 2207 5th-order and 90

7th-order basins were extracted for a “local” and regional signal, respectively (at the Amazonian scale, they are

respectively 134 and 3660 km2 in average). For these basins the following values are calculated.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 435-438

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- Hypsometry and hypsometric normalized integral (H), given in percent (Strahler 1952);

- Elongation LSE /)/(2= where S is the basin area and L the basin length (Schumm 1956). For

circular, mature basins, E~1, for highly elongated, immature, basins E<<1.

- Local average azimuth (Al) is the average 5th-order basin azimuth over 80 km x 80 km square cells.

Figure 3. Fitzcarrald Arch 5th-order basins’ locally averaged azimuths. Big star represents the approximate location for Fitzcarrald Arch divergent drainage centre. Small stars represent local centres for diverging drainages that disturb the Fitzcarrald drainage shape. The forebulge axis is shown; to the Southeast, after Aalto et al. (2003), to the Northeast after Roddaz et al. (2005).

Results

The main part of the Fitzcarrald Arch is characterized by relatively intermediate to high hypsometric integrals

(between 40 and 50%), and higher values at its north-eastern and eastern boundaries (more than 50%, up to 65%,

fig. 2). Low hypsometric interval values (values between 10 and 25%) are found to the northwest, around the

Moa Divisor range and low to intermediate values (15%-35%) are also present to the southwest, at the boundary

between the Fitzcarrald Arch and the Subandean zone, where some basins cross the Subandean thrust front.

7th-order basins elongation E-values range from 0.33 to 0.88. High E-values are found in the north-western part

(0.42 to 0.78) and in the south-western part of the Arch, near the subandean zone (0.50 to 0.85). Low E-values

are found to the south-east (from 0.33 to 0.66). Intermediate E-values are near the Arch centre and to the north-

east (0.42-0.88; very high E-values being in low-elevation areas).

Local average azimuths for 5th-order basins (fig. 3) show that Arch can be described by one major and a couple

of minor centers for radial drainage systems. The major centre is situated at 10.5°S and 72.5°W (star in fig. 3)

and may explain the first-order drainage pattern of the Arch. It corresponds to the centre of the radial drainage

organization. Superimposed to this large scheme whose wavelength is about 500 km, three second-order

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 435-438

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drainage centers (wavelength ~100 km) cause local divergences (see fig. 3)

Discussion/Conclusion

Both elongations and hypsometric integrals for 5th order and 7th-order basins show a relief maturity decrease

from north-west at Moa Divisor to south-east near the Madre de Dios basin and to the north-east (Altos de Acre);

the most mature basins are found to the south-east. A small part, to the south-west near by the Subandean zone,

displays mature basins. Drainage azimuth helps understanding this scheme. It indicates that there is a first-order

relief, which can be called the Fitzcarrald Arch sensu stricto, whose centre is situated ~100 km north-east from

the Subandean thrust front at 10.5°S and 72.5°W, and which covers the entire study area at the exception of the

Moa Divisor and the Subandean zone. According to the maturity gradient, from mature to the North-West to

immature to the South-East, this relief seems to have formed recently with a north-west to south-east

progression. The Subandean zone is separated from the Amazonian basin by the Subandean thrust front which is

the main tectonic feature observed in the area. Contrary to what expected, it does not affect significantly the

drainage organization, indicating this structure has no or little activity in recent times since it does not disturb

significantly the Fitzcarrald Ach drainage organization.

In sum, there is a clear progress from old basins to young basins from north-west to south-east, fully

compatible with Hampel (2002)’s reconstruction which imply a sliding from NW to SE (cf. fig. 1). Conversely,

our observations do support neither the hypothesis of alluvial fan, since its centre is not located near major basin

outlet, or the Pisco–Abancay–Fitzcarrald lineament which would imply a geanticline structure, different from

our radial structure.

References Aalto, R., L. Maurice-Bourgoin, et al. (2003). "Episodic sediment accumulation on Amazonian flood plains influenced by El

Nino/Southern Oscillation." Nature 425(6957): 493-497. DeMets, C., R. G. Gordon, et al. (1990). "Current Plate Motions." Geoph. J. Int. 101: 425-478. Dumont, J. F. (1996). "Neotectonics of the Subandes-Brazilian craton boundary using geomorphological data: the Maranon

and Beni basins." Tectonophysics 259(1-3): 137. Espurt, N., P. Baby, et al. (2007). "How does the Nazca Ridge subduction influence the modern Amazonian foreland basin?"

Geology 35(6): 515-518. Gutscher, M. A., J. L. Olivet, et al. (1999). "The "lost Inca Plateau": cause of flat subduction beneath Peru?" Earth And

Planetary Science Letters 171(3): 335-341. Hampel, A. (2002). "The migration history of the Nazca Ridge along the Peruvian active margin: a re-evaluation." Earth And

Planetary Science Letters 203(2): 665-679. Jacques, J. (2003). "A tectonostratigraphic synthesis of the Sub-Andean basins: implications for the geotectonic segmentation

of the Andean Belt." Journal of the Geological Society 160: 687. Roddaz, M., P. Baby, et al. (2005). "Forebulge dynamics and environmental control in Western Amazonia: The case study of

the Arch of Iquitos (Peru)." Tectonophysics 399(1-4): 87. Saillard, M., L. Audin, et al. (2008). Pleistocene uplift rates variability along the Andean active margin inferred from marine

terraces. 7th International Symposium on Andean Geodynamics (ISAG), Nice. Schumm, S. A. (1956). "Evolution of drainage systems and slopes in badlands at Perth Amboy, New Jersey." Bull Geol. Soc.

Am. 67: 597-646. Strahler, A. N. (1952). "Dynamic basis of geomorphology." Geological Society of America Bulletin 63: 923-938.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 439-441

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Orientation of current crustal stresses in the South America plate between 30° and 55° S

Claus-Dieter Reuther & Elmar Moser

Department of Geosciences, University of Hamburg, Bundesstr. 55, 20146 Hamburg, Germany

(reuther@geowiss.uni-hamburg.de)

KEYWORDS : active stress directions, southern South America plate

Introduction

Recent crustal stresses control the current structural deformation style of a region. Information about principal

stress directions (stress tensors) provide hints about the acting forces and allow conclusions regarding their

tectonic and / or gravitational origin. The orientation of the principal stresses governs fault and fold development

and determines potential movements along pre-existing faults.

Crustal stresses are generally due to plate tectonic forces and / or gravitational forces. Current stress

accumulation and stress release induce a form of permanent “crackling” within the earth crust. On a molecular

scale this mechanical disturbance is reflected by the breaking of atomic bonds leading to the creation of electric

dipoles. The resulting electromagnetic waves propagate normal to the dipole and therefore parallel to the crack.

The orientation of the opening microcracks depends on the state of stress, the principal stress directions and the

mode of deformation. Measuring pulsed electromagnetic geogenic signals in a specific frequency range leads to

the determination of the direction of the emitted electromagnetic waves. Based on certain physical and rock-

mechanical requirements these measurements allow statements about the maximum stress orientation in the

uppermost crust.

With in-situ measurements of electromagnetic emissions from rocks we determined the maximum horizontal

stress orientation in the southern part of the South America Plate between 30° S and 55° S. From the onshore

active margin along the Chilean Pacific coast across the Andes into the Argentinean foreland and in Patagonia as

far as to the passive Atlantic margin we took more then 500 readings and identified directions of the current

maximum horizontal stress (fig. 1).

Active horizontal tectonic stresses affecting the earth crust are acting inside a structural unit from the surface

into depth with the same magnitude. Thus in the uppermost crust at plate margins and in intraplate settings the

maximum prevailing stress direction is primarily horizontal and exceeds the gravitational stresses. Long before

failure and faulting of the rock the initial and prevailing structures are microcracks reflecting micro longitudinal

splitting. The maximum of the emitted electromagnetic waves corresponds to the maximum active stress

direction. During increasing confining pressure with depth hybrid cracks will form. The opening of tensional

cracks before fracturing still parallels the maximum stress direction and is again the source of directed

electromagnetic emissions. If the confining pressure increases more and one horizontal stress is still the

maximum stress, secondary order structures, pre-running rock failure will develop. The emitted electromagnetic

magnetic waves range within a dispersion cone still allowing the identification of a maximum stress direction

(Reuther & Moser 2007).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 439-441

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Fig.1: Active stress directions in the southern South America Plate

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 439-441

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Active maximum horizontal stress directions

Between South of the Talinay Peninsula and Valparaiso we observed a consistent W-E direction of the stresses

from the Chilean Pacific coast to the Western border of the Andes. From Valparaiso to Concepción, the coastal

area is characterized by mainly WNW-ESE directed stresses. Inland between Talca and Temuco, this area

corresponds to the Longitudinal Valley, the stress directions switch into a NNW-SSE orientation. On the Talinay

and Arauco peninsulas the maximum horizontal stress directions differ remarkably from the adjacent areas and

trend more or less parallel to the coast (NNE-SSW respectively N-S).

Along the Argentinean margin of the Andes and in the Andean foreland between San Juan down to the latitude

of Neuquén (linear distance about 800 km) the stress direction is quite uniform and trends WNW-ESE. Chiloé

Island (Chile) is characterized by a uniform stress-field with NE-SW orientations of SHmax differing from

mainland Chile. A significant stress anomaly occurs in the area of Lago Buenos Aires / Lago General Carrera.

West of the lake the stresses are varying between 80° and 100°. North and South of the lake that exhibits a

length of ca. 160 km and a width to about 25 km, the active stresses are trending between 170° and 10°. Towards

the region adjoining the lake to the east, SHmax switches in a 100° direction.

Across the Patagonian plains the stress directions change from WNW-ESE in the North into a NW-SE

direction towards the South. From Rio Gallegos and along the Magellanes the stresses vary between NNE and

NW. In the Puerto Natales – Torres del Paine region SHmax turns into a WNW-ESE direction.

Stress directions obtained from analysis of neotectonic / subrecent tectonic structures observed at several

measurement-locations along the Pacific coast and along the Southern and Patagonian Andes correspond to the

stress directions deduced from electromagnetic measurements. Results from paleostress analysis in the Lago

General Carreras / Buenos Aires area carried out by Lagabrielle et al. (2004) show stress data comparable to the

obtained stress directions of our study.

This study allows the identification of different active stress regions on the southern South America Plate and

supports the modelling of tectonic processes along the active plate margin from the onshore forearc area across

the magmatic arc into the backarc region and in Patagonia until the passive continental margin. In areas with

relative poor outcrop conditions, the determination of electromagnetic emissions is an useful tool to identify

current stress fields.

Acknowledgements We thank Dr. H. Obermeyer, GE&O, Karlsruhe (Germany) who provided us with two CERESKOP-Instruments for detecting natural electromagnetic emissions. The field trip to Patagonia was financed within the TIPTEQ Project by the German Bundesministerium für Bildung und Forschung (BMBF). We thank our Chilean friends and colleagues Prof. Dr. Arturo Quinzio, Prof. Ramiro Bonilla and Gian Carlo D’Ottone (Universidad de Concepción, Chile) for the logistic support of the field-campaigns.

References Reuther, C.-D., Moser, E. 2007 - Orientation and nature of active crustal stresses determined by electromagnetic

measurements in the Patagonian segment of the South America Plate. Int. J. Earth. Sci. (Geol. Rundsch.) (preprint online version Dec. 2007, DOI 10.1007/s00531-007-0273-0)

Lagabrielle, Y., Suarez, M., Rosello, E.A., Hérail, G., Martinod, J., Regnier, M., de la Cruz, R. 2004 - Neogene to Quaternary tectonic evoluion of the Patagonian Andes at the latitude of the Chile Triple Junction. Tectonophysics, 385: 211-241.

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New field studies in the Gonzanamá, Catamayo and Malacatos-Vilcabamba basins, Ecuador: Preliminary results

Pedro Reyes1,2

, François Michaud1,2

, Pierre Carbonel3, & Michel Fornari

2

1 Escuela Politécnica Nacional, Departamento de Geología, Andalucía n/s, C.P. 17-01-2755, Quito, Ecuador

(preyes@epn.edu.ec) 2 Geosciences-Azur (IRD-UPMC) BP 48, 06235, Villefranche/Mer, France (micho@geoazur.obs-vlfr.fr)

3 EPOC CNRS, Université Bordeaux 1, Bordeaux, France (p.carbonel@epoc.u-bordeaux1.fr)

The North Andean Block northward drifting has been related to lateral (opening of the Guayaquil Gulf) and

vertical motions (tectonic inversion of the Loja, Catamayo, Gonzanamá and Malacatos - Vilcabamba basins,

Fig.1). Hungerbühler et al. (2002) propose that the sedimentary infill evolution of these basins took place during

two stages. The first “Pacific coastal” stage took place between 15-10 Ma, with deltaic to brackish marine

environment deposits. At 10-9 Ma occurred the deformation of the sedimentary infill of the basins. The second

stage took place between 9 and 5 Ma with continental intermountain series. Two aspects of this model have

important implications: 1) development of marine embayments (15-10 Ma) throughout the basins (Steimann et

al., 1999; Hungerbühler et al., 2002) gives a reference point for the quantification of lateral as well as vertical

motions (Fig. 1) and 2) deformation period, between 10-9 Ma, involves surface uplift and thrusts.

We have realized the geological map (1:50 000 scale) between the towns of Catamayo, Gonzanamá and

Malacatos (approximately 700 km2, Fig 1.and 2A). We propose a new geological model for the Gonzanamá,

Catamayo and Malacatos - Vilcabamba basins evolution.

Ostracods fauna revisited

All Ostracod sample locations of Hungerbühler et al. (2002) were revisited. The samples yielded a rich fauna

of ostracods. The ostracofauna is present in most of the samples, sometimes, very abundant (more than 10000

individuals). It appears two characteristics:

- the diversity is poor (less than 5 species. That indicates instability of the physico-chemical conditions at the

water-substratum interface.

- all the genera of these ostracodes live at the present day in fresh or brackwaters. The dominance of Cyprideis

confirms the indication of instability of the waters. This instability is due to 2 main factors: the climate

(evaporation, seasonality) and/or the hydrothermalism marked also by intensive dissolution in several samples

Evidence for a regional intrusive episode affecting the Gonzanamá Basin

The sedimentary sequence of the Gonzanamá basin (MiGz, Fig. 2B) extends from the north of Nambacola to

the south of Gonzanamá (1500 m.s.n.m.) and reaches to the west the town of Purunuma (2400 m.) (Fig 2A). The

sediments have a regional eastward inclination between Nambacola and Gonzanamá and westward near

Purunuma. Nevertheless, in several places there are local changes of the inclination of the sediments which are

strongly deformed (kink fold, cf. see Figure 5.10.F in Hungerbühler, 1997). Towards the south of Nambacola

town (Fig. 2B) the sediments are intruded by andesitic intrusions (Ingaurcu and Yaramina hills, MiPor, Fig

2B). No metamorphism evidences near the contact of the sediments with the intrusion are present. Nevertheless,

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the strongest deformation observed in the sediments is located in the vicinity of these intrusions. Towards the top

of the sedimentary sequence the inter-bedded volcanic layers are more numerous. In some places sediment clasts

(which size ranges from few centimeters to several meters) are observed mixed with intrusive rocks. All these

observations allow concluding the existence of a magmatic intrusion and peperite-like phreatic eruption

(Skillling et al., 2002). The local deformation that was interpreted as slump (Kennerley, 1973, 1975; Jaillard et

al., 1996) or related to regional thrusting (Hungerbühler et al., 2002), results of the intrusion of magmatic bodies

within the soft sediments containing an important amount of water. Three zircon fission-track datings

(Hungerbühler et al., 2002) were made. Taking into account our observations and the sample locations of

Hungerbühler et al., (2002) (Fig. 2B), it seems that two of them are in the intrusive rocks (15.7 +/- 2.0 Ma and

14.4 +/- 1.8) whereas the last one is in volcanic levels to the top with sediments of Gonzanamá (14.0 +/-3 Ma)

(Fig 2.B). It is hoped to confirm these hypotheses with the results of analysis (petrologic, geochemistry, Ar/Ar

ages at the moment, in process) of the intrusive complex samples. Towards the east, between Gonzanamá and

Purunuma towns, several consecutive inselberg bodies (Colambo hill, Fig. 2A) ) can be observed (similar to the

well-known Cariamanga inselberg bodie located about 20 km southwest of Gonzanamá town), which display a

breccia crown that is indicative of a phreatic environment. Between Las Lagunas and Sasaco (NE of Purunuma

town, Fig 2A) the sediments are also affected by intrusions (hill of Colca, Fig 2A).

Evidence for a regional volcanic episode between the deposits in the

Catamayo Basin and those in the Malacatos - Vilcabamba Basin

The sediments of the Catamayo basin extend from the Catamayo town towards the south to the Santa Rita town

(Fig 2.A and 2C). The series includes conglomerates at the base followed by coarse yellow sands, fine

multicolored sandstone interbedded with limolites; finally to the top there is discontinuous limestone levels

interbedded with brown shales (MiCm, Fig.2C) Regionally, the layers plunge towards the east. At the top of the

sediments (between Catamayo and Boqueron) white and gray pyroclastic levels can be observed that vary from

tuffs to breccias with a total thickness between 20 to 80 meters (MiQuTa, Fig. 2C). These series outcrop in

discontinuous form, drawing channels within the top of the sediments. Above thick levels (up to 50 m) of grey

Figure 1. Left: Location of the marine embayments of Steinmann et al. (1999) and Hungerbühler et al. (2002), which 15-10 Ma ago extended throughout the Nambacola-Gonzanamá and Catamayo-Malacatos basins and was connected to the Progresso basin (yellow = basin locations). This would imply >100 km of lateral motion since 15 Ma. Red rectangle: mapped area. Above: Miocene marine basin (yellow) and its position today (after Steinmann et al., 1999), implying an uplift of >6000 m since 10 Ma.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 442-445

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volcanic breccias (MiQuB, Fig. 2C) underlay volcanic agglomerate levels interbedded with very well welded

green breccias (MiQu, Fig. 2C).

This volcanism extends towards the south until the La Era town (Fig 2A). Near Tambo town, in the sector of

San Antonio and along the highway Catamayo - Loja, this volcanism is in contact by fault with Paleozoic

metamorphic rocks. Between the La Era and La Merced towns (Fig 2A), the Malacatos - Vilcabamba basin

Figure 2. (A) Area mapped at scale 1:50000. (B) Map near Nambacola; the Gonzanamá basin (MiGz-MiGzc) are intruded by sub-volcanic rocks (hills Ingaurcu and Yaramina, MiPor). (C) Map between Catamayito and Matalá; strata of the Catamayo basin (MiCm-MiCmc) are conformably underlain by the Quinara Formation (white tuffs at the base (MiQuTa), followed by breccias and megabreccias levels (MiQuB), and finally sequences of agglomerates (MiQu)). To the south the Suche fault has possibly controlled uplift of the Loma Blanca formation (OlLB-OlLBP). The volcanic levels (MiGzCG-MiGzV) represent events that occurred before and after filling of the Gonzanamá basin. PcSa = Sacapalca Formation.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 442-445

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sediments (Marocco et al., 1995), rests directly on the volcanic agglomerate levels and breccias, previously

mentioned. This volcanism is lithologically different as well as structurally from the Loma Blanca Formation

described by Kennerley et al. (1973). No evidence of thrusting (as it was proposed by Hungerbühler et al., 2002)

of these volcanic series on sediments of the Catamayo basin has been found in the area, according to Kennerley

et al. (1976) and Jaillard et al. (1996). This volcanism which is intercalated between the Catamayo basin

sediments and Malacatos-Vilcabamba basin sediments is lithologically equivalent to the volcanic Quinara

Formation described by Hungerbühler et al. (2002) south of Vilcabamba town (dated 15 Ma). This age is

coherently older than the ages from 12 to 13,5 Ma, proposed by Hungerbühler et al. (2002) for the sediments of

the Malacatos-Vilcabamba basin.

Along the Catamayo-Nambacola highway outcrop pyroclastics levels whose age is 29 Ma (Hungerbühler, et

al., 2002). Faults as the Suche fault (Fig. 2C) possibly control the vertical position of this volcanic set which

could explain the relative altitude between Catamayo basin (1200 m) and Gonzanamá basin (1600-2400 m).

Conclusions

Based on the geological mapping (1:50000 scale) among the Catamayo - Gonzanamá - Malacatos towns, we

propose a new interpretation of this area. The presence of peperites related to magmatic intrusion and associated

phreatic eruption into the sediments of the Gonzanamá basin is reported for the first time. The local deformation

in sediments is related to this magmatic episode and not to thrusting (10-9 Ma). Moreover the ostracofauna do

not agree with to the marine embayments model proposed (Steinmann et al., 1999; Hungerbühler et al. (2002).

The Gonzanamá basin sediments are affected by an intrusive event before deposition of the Catamayo basin

sediments. The Catamayo basin deposits are separated from the Malacatos - Vilcabamba basin deposits by an

interbedded volcanic event.

Acknowledgments. This work was supported by IRD (France) and by the Departamento de Geologia de EPN (Ecuador).

References Hungerbühler, D., 1997. Tertiary basins in the Andes of southern Ecuador (3º00’ – 4º20’): sedimentary evolution,

deformation and regional tectonic implications. PhD Thesis, Institute of Geology ETH Zurich, Switzerland, 182 pp. Hungerbühler D., Steinmann, M., Winkler W., Seward D., Egüez A., Peterson D.E., Helg U., and Hammer C., 2002,

Neogene stratigraphy and Andean geodynamics of southern Ecuador, Earth Science Reviews, 57, p. 75-124. Jaillard, E., Odóñez M., Berrones G., Bengtson P., Bonhomme M., Jimenez N., y Zambrano I., 1996, Sedimentary and

tectonic evolution of the arc zone of Southwestern Ecuador during the Late Cretaceous and Early Tertiary times, Journal of South American Earth Sciences, 12, p. 51-68.

Kennerley, J.B., 1973. Geology of Loja Province, southern Ecuador. Institute of Geological Sciences Overseas Division, London. Unpublished Report 23, 34 pp.

Kennerley, J.B., Almeida, L., 1975. Mapa geológico del Ecuador, hoja de Cariamanga (39), escala 1:100.000. Instituto Geográfico Militar IGM, Ministerio de Recursos Naturales y Energéticos MRNE, Dirección General de Geología y Minas, DGGM, Institute of Geological Sciences London IGS .

Kennerley, J.B., Almeida, L., 1975. Mapa geológico del Ecuador, hoja de Loja (57), escala 1:100.000. Instituto Geográfico Militar IGM, Ministerio de Recursos Naturales y Energéticos MRNE, Dirección General de Geología y Minas, DGGM, Institute of Geological Sciences London IGS.

Marocco, R., Lavenu, A., Baudino, R., 1995. Intermontane Late Paleogene–Neogene basins of the Andes of Ecuador and Peru: sedimentologic and tectonic characteristics. In: Tankard, A.J., . Suarez, R., Welsink, H.J. Eds., Petroleum Basins of South America. American Association of Petroleum Geologists Memoir, vol. 62, pp. 597–613.

Skilling I.P., White J.D.L., y McPhie J., 2002, Peperite : a review of magma-sediment mingling, Journal of volcanology and geothermal research, 114, p.1-17.

Steinmann, M., Hungerbuhler, D., Seward, D., Winkler, W., 1999, Neogene tectonic evolution and exhumation of the southern Ecuadorian Andes: a combined stratigraphy and fission-track approach. Tectonophysics 307, 25.

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71º 30’ 71º 00’ 70º 30’ W

16º00'S

16º30’

Petrology of the 2006-2007 tephras from Ubinas volcano, southern Peru

Marco Rivera1,2

, Marie-Chistine Gerbe3, Alain Gourgaud

1, Jean-Claude Thouret

1, Hervé

Martin1, Jean-Luc Le Pennec

1, & Jersy Mariño

2

1 IRD and Laboratoire Magmas et Volcans, Université Blaise-Pascal, 5 rue Kessler, 63038 Clermont-Ferrand,

France (mrivera@opgc.univ-bpclermont.fr) 2 INGEMMET, Dirección de Geología Ambiental. Av. Canadá 1470, San Borja, Lima, Peru

3 Université Jean Monnet, UMR 6524 CNRS, 22 rue Paul Michelon, 42023 Saint Etienne, France

KEYWORDS : Ubinas, vulcanian eruption, lava, tephra, petrography, mineralogy, geochemistry

Introduction

The Ubinas volcano (16º 22' S, 70º 54' W) is located in the Quaternary volcanic range in southern Peru,

~60 km east of Arequipa city (Fig. 1). Ubinas is historically the most active volcano in southern Peru with 24

volcanic events (VEI 1-3) recorded since 1550 AD (Hantke and Parodi, 1966; Simkin and Siebert 1994; Rivera

et al. 1998). These events are largely intense degassing episodes, with some ashfall and ballistic blocks

(<10.106m3) produced by vulcanian and phreatomagmatic explosive activity (Thouret et al. 2005; Rivera et al.

1998). The events caused damage to crops and cattle and affected approximately 3,500 people living in six

villages within 12 km from the volcano (Fig. 1).

The most recent explosive activity began on 27 March 2006 and lasted two years with intermittent eruptive

events, while degassing is still ongoing at present. Based on the characteristics of activity and the erupted

products the eruptive episode has progressed in four stages: 1) initial phreatic and phreatomagmatic activity

(27 March to ~19 April 2006), including high eruption columns that dispersed ashfall as far as 7 km from the

summit; 2) vulcanian explosions (~20 April to 11 June 2006) formed 3 to 4 km-high columns that ejected blocks

Fig. 2 Fig. 1

Fig. 1. Location map of the Ubinas volcano. Inset shows its location with respect to the volcanic range in southern Peru.

Fig. 2. Ballistic blocks were ejected from the volcano on 24 May 2006. This photograph shows an impact crater created where a 2-m-diameter bomb struck the caldera floor ~200 m from the crater.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 446-449

447

up to 40 cm in diameter to distances 2 km from the vent (Fig. 2). Fresh lava reached the vent bottom on 20 April;

3) strong degassing interspersed with at least 12 events that produced 2 to 3 km-high columns between mid June

2006 and April 2007, dispersing ash as far as 40 km from the vent; 4) mild degassing produces a permanent

200 to 800 m-high plume and occasional light ashfall around the summit (May 2007 until the present).

Short-lived lasting and slug columns, cannon-like explosions, small amounts of juvenile material, and the

andesitic composition of bread-crust bombs indicate a vulcanian style of behaviour at Ubinas. The behaviour is

similar to the first phase of the Nevado Sabancaya eruption in 1990-1998 (Gerbe and Thouret, 2004) or to the

behaviour of Sakurajima, Japan since 1955 (Morrisey and Mastin, 2000), and to Ngauruhoe, New Zealand in

1974-1975 (Hobden et al. 2002). Petrographical and geochemical characteristics of juvenile blocks and scoriae

erupted during the 2006-2007 explosive activity allow for the description of newly erupted magma and therefore

leads to a better understanding of the origin of the eruption.

Petrography and mineralogy of the 2006 tephra

The juvenile dense and poorly vesicular blocks erupted on 27 April (Ub-04), 7 May (Ub-13), 24 May (Ub-14)

and 28 October 2006 (Ub-18) are porphyritic (Fig. 3a, 3b) and contain phenocrysts (250μm-1.6mm in size,

2-5% vol.) and microphenocrysts (80-250μm in size and 30 to 40 vol.%) of subhedral to euhedral plagioclase

(An41-68) and a small amount of amphiboles and clinopyroxenes. The plagioclase phenocrysts are variably zoned.

Some display reverse zonations, low-An cores (An33-56) surrounded by relatively high-An “dusty”rims (An47-68)

containing abundant small (1-20μm) melt inclusions (Fig. 3b). Other plagioclase phenocrysts lack the

“dusty”rims, being normally zoned with high-An (An47-66) cores and An41-59 rims.

Fig. 3. Photomicrographs of thin sections in scoriae and dense blocks: a) “dusty-rimmed” plagioclase phenocryst and clinopyroxene phenocryst. b) Amphibole phenocryst showing reaction rims.

In addition, some tephra samples (Ub-04, 13, 14) contain scattered phenocrysts of subhedral and anhedral

amphibole, namely pargasite with Mg# 66-70, and 200 to 300μm in size. They show reaction rims (20 to 150μm

in width) suggesting resorption or dissolution surfaces, probably due to decompression effect during magma

ascent. Clinopyroxene, specifically augite (En39-48 Wo38-49 Fs14-19), are either phenocrysts up to 800 μm in size or

microphenocrysts; some showing reverse zonation (Mg# 68-74). Phenocrysts of orthopyroxene, i.e enstatite

(En65-71 Wo2-7 Fs23-32) are up to 600μm in size and sometimes show slight reverse zonation (Mg# 71-73).

Numerous glomerophenocrysts of plagioclase, clinopyroxene, orthopyroxene, and Fe-Ti oxydes are in reaction.

In all blocks and scoriaes, Fe-Ti oxydes (<200 μm and <2-4 vol. %) are euhedral and dispersed in the

cpx

plg

amp

0 150 m 0 150 m

a) b)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 446-449

448

0.0

0.5

1.0

1.5

2.0

2.5

3.0

3.5

4.0

4.5

50 55 60 65 70 75

SiO2 wt%

K2O

wt%

April

07-May

24-May

Historical times

Basaltic

Andesite

Andesite Dacite Rhyolite

High-K

Calco-Alkaline

groundmass, and also appear as inclusions in the phenocrysts (orthopyroxene, clinopyroxene, amphibole, and

olivine).

Blocks that were sampled in April and May 2006 include scattered, subhedral and anhedral olivine

phenocrysts (<200 μm and 1-2 vol. %), which display normal zonation (Fo72-76 cores and Fo63-71 rims). They are

usually surrounded by a fine microlite aggregate, while resorbed shapes suggest that they may be xenocrysts.

The groundmass (<80 μm) consists of plagioclase, ortho- and clinopyroxene, and dacitic glass (67-68 wt%

SiO2). Both olivine and amphibole are missing in lavas erupted in October 2006, and contrast to lavas erupted

earlier in April and May 2006. However, the mineral assemblage does not display any significant variation in

mineral composition throughout the eruptive episode. On the other hand, pre-eruption temperatures have been

estimated using different calibrations of the two pyroxenes (Wood and Banno 1973; Wells 1977). Pre-eruption

temperature is estimated to range between 1000 and 1090 ºC.

Geochemistry of juvenile 2006 lavas

The juvenile magma, represented by lava blocks and scoriae, comprise high-K calc-alkaline andesite showing

a restricted range of composition (56.7-57.6% SiO2; 2.0-2.3% K2O: Fig. 4) compared to historical lavas. In

addition, trace element compositions are characterised by a high LILE (K, Rb, Ba, Th) and LREE contents with

respect to HREE (Fig. 5). The trace element composition of the juvenile tephra is similar to the average

composition of the erupted andesites over the last 1500 years (Thouret et al. 2005; Fig. 5). Depleted Y and

HREE is attributed to mixing and assimilation processes of magmas near the base of the >60-km-thick

continental crust (Thouret et al. 2005).

Discussion and conclusion

Over the past 1500 years Ubinas has erupted magma ranging from basaltic andesites to dacites, with andesites

being most common. Chemical characteristics of the magmas mainly result from fractional crystallisation

processes in a shallow magma chamber and assimilation at various crustal levels (Thouret et al. 2005). In

0.1

1

10

100

1000

RbBa Th U Nb K La CePb Pr Sr P Nd ZrSmEu Ti Dy Y Yb

Ro

ck/N

MO

RB

Sun/McDon. 1989-NMorb

Scoria of 1662

Lava flows of stratoconebetween 370,000 - 142,000 yr B.P.

Bombs of may and october 2006

Fig. 5

Fig. 4

Fig. 4. Alkali-silica diagram showing that the composition of erupted lavas in 2006 at Ubinas.

Fig. 5. Spiderdiagram of the 2006 - 2007 tephras for the purpose of comparison with Ubinas lavas prior to 2006 (Thouret et al. 2005). Both figures show that the composition of the 2006 erupted lavas is similar to the average composition of historically erupted lavas

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 446-449

449

addition few erupted pyroclastic products show evidence of magma mixing as mineralogical disequilibrium, in

combination with shallow aquifers of the hydrothermal system, which may have contributed to the triggering of

eruptions (e.g., the AD 1677 scoria and ash flow deposits of basaltic andesite composition. On the other hand,

the 980 yr BP-old plinian event produced a voluminous dacitic pumice fall deposit that does not contain

evidence for magma mixing. Thus, recent Ubinas magma has displayed a large range in compositions and the

magma chamber may have been periodically and partially emptied during the historical eruptions.

The geochemical composition of the juvenile magma erupted at Ubinas between April and October 2006 is

similar to that erupted during the last 1500 years, suggesting that all magmas have the same mantle source.

However, some petrographic textures and chemical zoning pattern of phenocrysts suggest that part of the mineral

assemblage was not in equilibrium with the melt prior to, or during the eruptive activity. The juvenile tephra

erupted at Ubinas between April and October 2006 exhibit plagioclase phenocrysts (some “dusty-rimmed”, with

reverse zonation), orthopyroxene, clinopyroxene with reverse zonation, amphibole with disequilibrium features

as reaction rims, and olivine xenocrysts. Based on the types of textures and mineral geochemistry two

hypotheses can be considered for the triggering mechanism of the most recent eruptive activity: 1) re-supply of

mafic magma into the conduit or a shallow magma chamber containing cooler andesitic magma, which has

triggered the eruption through the addition of heat and/or volatiles to the resident magma; or 2) repeated and

continuous ascent of small batches of new magma, which incorporated xenocrysts of magma erupted previously.

Both processes may have led to over pressurisation of the magma chamber and probably triggered the mild

eruptive episode.

References Gerbe M.-C., Thouret J.-C., 2004. Role of magma mixing in the petrogenesis of lavas erupted through the 1990-98 explosive

activity of Nevado Sabancaya in south Peru. Bulletin of Volcanology, 66, 541-561. Hantke G., Parodi I., 1966. The active volcanoes of Peru. Catalogue of the active volcanoes of the world including sofatara

fields, part. XIX, Colombia, Ecuador and Peru, International Association of Volcanology, Roma;65-73. Hobden B.J., Houghton B.F., Nairn I.A., 2002. Growth of a young, frequently active composite cone: Ngauruhoe volcano,

New Zealand. Bulletin of Volcanology, 64, 392-409. Morrisey M.M., Mastin L.G., 2000. Vulcanian eruptions. In; Sigurdsson H (ed) Encyclopedia of volcanoes. Academic Press,

San Diego, p 463-475. Simkim T., Siebert L., 1994. Volcanoes of the World - A Regional Directory, Gazeteer and chronology of volcanism during

the last 10,000 years. Global Volcanism Program, Smithsonian Institution, Washington DC. pp. 348. Rivera M., Thouret J.C., Gourgaud A. 1998. Ubinas, el volcán más activo del sur del Perú desde 1550: Geología y

Evaluación de las amenazas volcánicas. Bol. Soc. Geol. Perú 88 ; 53-71. Thouret J.C., Rivera M., Worner G., Gerbe M.C., Finizola A., Fornari M., Gonzales K., 2005. Ubinas: the evolution of the

historically most active volcano in southern Peru. Bulletin of Volcanology; 67: 557 - 589. Wells P.R.A. 1977. Pyroxene thermometry in simple and complex systems. Cont. Mineral. Petrol. 62; 129-140. Wood B.J., Banno S. 1973. Garnet-orthopyroxene and orthopyroxene-clinopyroxene relationship in simple and complex

systems. Cont. Mineral. Petrol. 42; 109-124.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 450-453

450

Comparative methodological considerations for estimating fracture parameters

Wilmer Robles, Andreas Kammer, Mauricio Marentes, & Wilmer Espitia

Universidad Nacional de Colombia, Cra 30 Calle 45, Edif. Manuel Ancízar of. 310, Bogotá, Colombia

(wilmerangel@gmail.com)

KEYWORDS : Methodology, fracture parameters, synthetic and natural pavements, Arcabuco anticline

Introduction

Studies of natural fractures bear a great importance on the evaluation of geological structures, transport

mechanisms of fluids, the stability of rock units and paleostresses and urge a neat methodological approach for

characterizing them. Quantitative outcrop studies are based on the evaluation of different parameters that include

intensity, density and medium lengths. Intensity is defined as averaged length of fracture traces per unit area,

density as numbers of center points per unit area and mean length as an averaged length of a fracture population

(Rohbaugh, 2002). In general terms, methods applying for the quantification of these fracture parameters include

the sampling of fractures along scan lines or in areas. Problems related to these sampling methods principally

pertain to their variable orientations and censoring bias. Further problems refer to the heterogeneity of fracture

distributions, i.e. to spatial changes in the fracture pattern and to the observational resolution. Thus, long fracture

traces are more easily sampled than shorter ones. In order to obviate these problems different methods have been

proposed (Mauldon et al., 2001). In this paper we implement various recently published methods designed to

estimate fracture parameters (Mauldon et al., 1998; Mauldon, 2001; Nieto et al., 2003; Wu & Pollard, 1995 ) and

confront their results, in order to identify possible limitations in their calculations and to compare their results in

an objective way. We perform this evaluation utilizing synthetic models and natural fracture populations of the

Arcabuco Anticline, located near the town of Villa de Leyva, Department of Boyacá, Colombia.

Methodological implementation

Figure 1. Quantities involved in the calculation of fracture parameters. Circles designate the n intersections between fractures and scan lines (both circular and rectangular), while triangles indicate the m terminations of fracture traces within a window.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 450-453

451

A first step in obtaining the fracture parameters consists in digitizing observed or synthetic fracture patterns. In

a subsequent step for which we designed different programs we superposed scan lines and windows (circular and

rectangular), in order to obtain some numerical quantities that allow for an estimation of the fracture parameters

(Figure 1). These fracture parameters were assessed using both single fracture sets (like the ones differentiated

by means of various colors in Figure 1) as well as indiscriminate fracture populations.

Circular windows

In this method scan lines represent the circumferences and sample areas the inside of a circular window.

Parameters can be assessed by counting the number of fracture traces (n), which intercepts the circumferences,

and the number of fractures the terminations (m) which are completely contained within a circular window

(Mauldon et al. 2001). In this procedure fracture parameters are defined as follows:

Density

22

ˆr

mcircir = (1)

Intensity

r

nI cir

cir4

ˆ = (2)

Mean trace lengh

=cir

circir

m

nr

2μ̂ (3)

where r represents the radius of the circular window.

Rectangular windows

In addition to the aforementioned parameters m and n the methodological approach for rectangular windows

requires the determination of an orientation angle between fracture plane and the long side of the rectangle. In

order to establish a direct comparison between rectangular and circular windows we applied squares of length r

that circumscribe previously defined circumferences (Figure 2). Considering these special conditions fracture

parameters were defined in the following way (Mauldon, 1998):

Density

24

ˆr

mrectrect = (4)

Intensity

]cossin[

4ˆ

+=

E

nrI rect

rect (5)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 450-453

452

Mean trace lengh

rect

rectrect

m

n

E

rI

]cossin[

ˆ

+= (6)

Figure 2. Sampling of the fracture parameters utilizing multiple scan lines and windows (both circular and retangular)

Direct estimation of intensity

Taking the definition of Mauldon (2001), it is possible to estimate the fracture intensity parameter summing all

traces enclosed completely within a window and dividing it by the area of the window, as is indicated by Nieto et

al. (2003).

Intensity

=A

lI i

dir

) (7)

Estimation of a fracture spacing

Wu & Pollard (1994) determined a fracture spacing parameter, which is interpreted as being the reverse of the

spacing, obtaining a similar results that the direct calculation.

=

+

=n

i

io ll

AS

1

) (8)

where lo is the length of a square and A = lo x lo represents a window’s area.

Using the appropriate expressions for each case and utilizing especially designed computing programs we

automated the estimation of the fracture parameters. These routines allow considering multiple windows and

varying their size, introducing, therefore, much versatility in the evaluation of the fracture parameters. In Figure

2 fracture parameters are assigned to specific sizes of windows which have been spawn in a random manner over

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 450-453

453

a pavement. In Figure 3 an arrow marks the results obtained for the pavements of Figures 1 and 2, utilizing an

input of 1000 windows and scan lines per size. Windows encompass a range starting from 1 cm and covering

intervals of 1cm until reaching the maximum size of a completely inscribed circle in the reference area.

This procedure was tested by means of a synthetic fracture pattern that mimic simple situation with known

parameters. Additionally, we tested these methods on natural pavements in different structural positions of the

Arcabuco Anticline (Villa de Leyva, Colombia) (Robles et al, 2007)

Figure 3. Results obtained for pavements shown in figures 1 and 2 deploying 1000 scan lines y windows of variable sizes. Red lines depict results obtained employing circular windows, blue lines refer to rectangular windows. For fracture intensity we used direct estimates based on circular (cyan lines) and rectangular (black lines) windows. Green lines refer to the method of Wu y Pollard (1995).

Results

Our synthetic model runs confirm the efficiency of our methodology, though we found situations where the

methodology fails to reproduce the known parameters.

Applying different methods for the evaluation of fracture parameters we found a good coincidence, though

values calculated for fracture intensities are slightly higher, utilizing the method of Wu & Pollard (1994). (See

Figure 3)

Evaluating fracture patterns along different structural positions of the Arcabuco Anticline of the town of Villa

e Leyva (Boyacá, Colombia) we find that fracture parameters positively correlate with structural positions.

References Mauldon M. 1998. Estimating mean fracture trace length and density from observations in convex windows. Rock Mechanics

and Rock Engineering 31: 201-216. Mauldon M.,Dunne W. &Rohrbaugh M. Jr. 2001. Circular scanlines and circular Windows: new tools for characterizing the

geometry of fracture traces. Journal of Structural Geology 23: 247-258. Nieto A., Alaniz S., Tolson G., Xu S.& Perez J. 2003. Estimación de densidades, distribuciones de longitud y longitud total

de fracturas; un caso de estudio en la falla de Los Planes, La Paz, B.C.S. Boletín de la Sociedad Geológica Mexicana tomo LVI, Nº 1:. 1-9.

Rohbaugh M. Jr., Dunne W., & Mauldon M. 2002. Estimating fracture trace intensity, density, and mean length using circular scan lines and windows. AAPG Bulletin, V.86, Nº 12: 2089-2104.

Robles W., Buitrago J., Kammer A., Marentes M.& Caro M. 2007. “Esrimación deparámetros estadísticos para sistemas de fracturas, caso de studio en el anticlinal de Arcabuco sector de Villa de Leyva”. In: XI Congreso colombiano de Geología. Bucaramenga, 2007

Wu H. & Pollard D. 1995, An experimental study of the relationship between joint spacing and layer thickness. Journal of Structural Geology vol.17: 887-905.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 454-457

454

Subduction control on chemical composition of Oligocene to Quaternary sediments of the Ecuadorian Amazonian foreland basin from major and trace elements and Nd-Sr isotopes

Martin Roddaz*, Frédéric Christophoul, Jean-Claude Soula, José David Burgos-Zambrano,

Patrice Baby, & Joachim Déramond

LMTG, Université de Toulouse, CNRS, IRD, OMP, 14 avenue Edouard Belin, 31400 Toulouse, France

* corresponding author (mroddaz@lmtg.obs-mip.fr)

KEYWORDS : foreland basin, major and trace elements, Nd-Sr isotopes, provenance, Andes, alluvial fan sediments, Ecuador

Introduction

The Ecuadorian Andes are characterized by frequent earthquakes with magnitudes greater than M=5 (Legrand

et al., 2005) and numerous active volcanoes with compositions ranging from calk-alkaline to shoshonitic which

can be related to a subduction context perturbed by the subduction of the Carnegie ridge (Bourdon et al., 2003).

The Ecuadorian Andes have been shown to have risen during the Neogene (Spikings et al., 2000), with

elevations and uplift rates similar to the nearby Peruvian and Bolivian Andes to the south and Colombian Andes

to the north. In contrast with other Andean foreland basins where progressive shortening lead to forward

(cratonward) migration of the thrust wedge and adjacent foredeep (DeCelles and Horton, 2003; Hermoza et al.,

2005), the Ecuadorian Amazonian foreland basin (Oriente foreland basin) experienced little shortening and

depocentre migration since the Oligocene which make its foredeep the widest of the Amazonian foreland basin

system. The most spectacular characteristic of the Ecuadorian foreland is the presence in front of the thrust and

fold belt, of a large-scale (60,000 km2) humid tropical alluvial fan, the Pastaza megafan (Räsänen et al., 1992).

This modern megafan partly overlaps the Miocene and Pliocene–Pleistocene fans (Christophoul et al., 2002).

Figure 1: a) Location map and geologic map of the studied area. 1A: Geologic map of the studied area; Ab.G: Abitagua Granitoid, Nap.D: Napo Dome, Cu.D: Cutucu Dome, ST: Subandean Thrust, SF: Subandean Frontal thrust. Dur to industrial confidentiality, location of oil wells is approximate. 1B: Simplified structural map of Ecuador. Co.B: Coastal Block, W.C.:

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 454-457

455

Western Cordillera; I.A.D: Interandean Depression; E.C.: Eastern Cordillera; S.Z.: Subandean Zone. White stars represent sampled outcrops (modified from Burgos Z et al., submitted)

Geochemical studies (major and trace elements and Nd-Sr isotopic compositions) of Neogene Amazonian

foreland basin sediments have been used successfully to investigate the control of the foreland basin dynamics

on the geochemistry (Roddaz et al., 2006) and unravel the drainage evolution of the Amazonian foreland basin

(Roddaz et al., 2005).

Results and conclusions

46 sediment samples were collected along river and road cuts in Ecuador. All samples were analyzed by ICP-

MS for trace elements at the LMTG (University of Toulouse, France). Forty one samples were selected for

determination of Sc concentrations (Chemex Labs, Canada). Thirty two samples were selected for major element

analysis by ICP-AES at the CRPG (Nancy, France). Sr and Nd isotopic compositions of ten representative

samples were measured by thermal ionization mass spectrometer at the University of Toulouse (France).

Based on major and trace elements and Nd-Sr isotopic compositions of Oligocene to Quaternary sediments of

the Ecuadorian foreland basin, we suggest that:

1) Sedimentary sorting played a minor role in the chemical differentiation of Oriente foreland basin

sediments;

2) Most of the analyzed sediments have CIA values close to that of the PAAS indicating that they are

moderately weathered. In addition, they have lower CIA values than those of the other Neogene Amazonian

foreland deposits (Roddaz et al., 2006) indicating that weathering was less intense in the Ecuadorian Amazonian

foreland basin;

3) The Ecuadorian foreland basin has been continuously fed by poorly to moderately weathered sediments

issued from andesitic protoliths since the Oligocene in non steady-state weathering conditions as indicated in the

A-CN-K diagram (Fig. 2a);

4) When compared with the other Amazonian foreland sediments (Roddaz et al., 2005; Roddaz et al.,

2006), the analyzed sediments have contributions of volcanic arc rock sources as indicated by their high Cr/Th

ratios and Nd(0) values and low Th/Sc ratios and Eu anomalies. The Quaternary sediments are derived from

more “basic” sources (Fig. 2b). As the Ecuadorian foreland basin is continuously alimented by volcanic arc

detritus since the Oligocene, we suggest that the chemical change observed in the Quaternary sediments is due to

a change in the nature of the volcanism in the Quaternary.

5) The peculiar chemical characteristics of the Ecuadorian foreland basin sediments can be best explained

in regard to the particular characteristics of the foreland basin geodynamics. The exhumation of the Eastern

Cordillera occurring in the Pleistocene causes the Western Cordillera to have been the main source of the

foreland basin deposits from Oligocene till Quaternary allowing export of andesitic detritus to the Amazonian

lowland. The Late Pliocene-Pleistocene subduction of the Carnegie ridge triggered main and back arc volcanism

inducing export of more fresh and basic sediments into the Amazonian foreland basin.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 454-457

456

Figure 2 : a) Ternary A-CN-K diagrams (Nesbitt and Young, 1984; Nesbitt and Young, 1989) for analyzed sediments. Pl:

Plagioclase; Ks: K-feldspar; Ad: Andesite (from (Fedo et al., 1995); b) 87Sr/86Sr- Nd(0) diagram for Western Amazonian sediments. Data sources: A: Quaternary Ecuadorian lavas (Barragan et al., 1998); B: Type I sand field (Basu et al., 1990); C: Type II sand field (Basu et al., 1990). Mesozoic and Neogene volcanic rocks from (Rogers and Hawkesworth, 1989) and from (Kay et al., 1994). Data for Central Depression, Altiplano, Oriental Cordillera, Subandean Zone fields are available in (Pinto, 2003)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 454-457

457

References Barragan, R., Geist, D., Hall, M., Larson, P. and Kurz, M., 1998. Subduction controls on the compositions of lavas from the

Ecuadorian Andes. Earth Planet. Sci. Lett., 154: 153-166. Basu, A.R., Sharma, M. and DeCelles, P.G., 1990. Nd, Sr-isotopic provenance and trace element geochemistry of Amazonian

foreland basin fluvial sands, Bolivia and Peru: implications for ensialic Andean orogeny. Earth Planet. Sc. Lett., 100: 1-17. Bourdon, E., Eissen, J.-P., Gutscher, M.-A., Monzier, M., Hall, M.L. and Cotten, J., 2003. Magmatic response to early

aseismic ridge subduction: the Ecuadorian margin case (South America). Earth and Planetary Science Letters, 205(3-4): 123-138.

Christophoul, F., Baby, P., Soula, J.-C., Rosero, M. and Burgos, J., 2002. Les ensembles fluviatiles neogenes du bassin subandin d'Equateur et implications dynamiques: The Neogene fluvial systems of the Ecuadorian foreland basin and dynamic inferences. Comptes Rendus Geosciences, 334(14): 1029-1037.

DeCelles, P. and Horton, B.K., 2003. Early to middle Tertiary basin development and the history of Andean crustal shortening in Bolivia. Geological Society of America Bulletin, 115: 58-77.

Fedo, C.M., Nesbitt, H.W. and Young, G.M., 1995. Unraveling the effects of potassium metasomatism in sedimentary rocks and paleosols, with implications for paleoweathering conditions and provenance. Geology, 23(10): 921-924.

Hermoza, W., Brusset, S., Baby, P., Gil, W., Roddaz, M., Guerrero, N. and Bolanos, R., 2005. The Huallaga foreland basin evolution: Thrust propagation in a deltaic environment, northern Peruvian Andes. Journal of South American Earth Sciences, 19(1): 21-34.

Kay, S., Coira, B. and Viramonte, J., 1994. Young mafic back arc volcanic rocks as indicators of continental lithospheric delamination beneath the Argentina Puna plateau, central Andes. J. Geophys. Res., 99: 24323-24339.

Legrand, D., Baby, P., Bondoux, F., Dorbath, C., Bes de Berc, S. and Rivadeneira, M., 2005. The 1999-2000 seismic experiment of Macas swarm (Ecuador) in relation with rift inversion in Subandean foothills. Tectonophysics, 395(1-2): 67-80.

Nesbitt, H.W. and Young, G.M., 1984. Prediction of some weathering trends of plutonic and volcanic rocks based on thermodynamic and kinetic considerations. Geochim.Cosmochim.Acta, 48: 1523-1534.

Nesbitt, H.W. and Young, G.M., 1989. Formation and diagenesis of weathering profiles. J.Geol., 97: 129-147. Pinto, L., 2003. Traçage de l'érosion Cénozoïque des Andes Centrales à l'aide dela minéralogie et de la géochmie des

sédiements (Nord du Chili et Nord-Ouest de la Bolivie). PhD Thesis, University Paul Sabatier, Toulouse, 196 pp. Räsänen, M.E., Neller, R., Salo, J. and Jungner, H., 1992. Recent and ancient fluvial deposition systems in the Amazonian

foreland basin, Peru. Geol. Mag., 129: 293-306. Roddaz, M., Viers, J., Brusset, S., Baby, P., Boucayrand, C. and Herail, G., 2006. Controls on weathering and provenance in

the Amazonian foreland basin: Insights from major and trace element geochemistry of Neogene Amazonian sediments. Chemical Geology, 226(1-2): 31-65.

Roddaz, M., Viers, J., Brusset, S., Baby, P. and Herail, G., 2005. Sediment provenances and drainage evolution of the Neogene Amazonian foreland basin. Earth Planet. Sc. Lett., 239(1-2): 57-78.

Rogers, G. and Hawkesworth, C.J., 1989. A geochemical traverse across the North Chilean Andes: evidence for crust generation from the mantle wedge. Earth Planet. Sci. Lett., 91: 271-275.

Spikings, R.A., Seward, D., Winkler, W. and Ruiz, G.M., 2000. Low-temperature thermochronology of the northern Cordillera Real, Ecuador: Tectonic insights from zircon and apatite fission track analysis. Tectonics, 19(4): 649-668.

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Neogene erosion and relief evolution in the Central Chile forearc (33°-34ºS) as determined by detrital heavy mineral analysis

María P. Rodríguez1, Luisa Pinto

1, & Gérard Hérail

2

1 Departamento de Geología Universidad de Chile, Plaza Ercilla 803, 13518 Santiago, Chile

(mariarod@ing.uchile.cl, lpinto@ing.uchile.cl) 2 IRD, LMTG, 14 Avenue Edouard Belin, 31400, Toulouse, France (gerard.herail@ird.fr)

KEYWORDS : Central Chile, morphostructural unit, peneplain, heavy mineral analysis, Leterrier diagrams

Introduction

At the forearc of Central Chile (33-34ºS) three morphostructural units are recognizable from west to east: the

Costal Cordillera (CC), the Central Depression (CD) and the Principal Cordillera (PC) (Fig 1). Flat-shaped

erosional surfaces forming the highest summits of the eastern CC (ECC) and similar surfaces at the western PC

(WPC) have been interpreted as relicts of ancient peneplains and indicated that the CC and the PC once formed

part of a single relief (Brüggen, 1950; Borde, 1966; Farías et al., 2006).

Figure 1 Geological map of Central Chile forearc (33-34ºS)

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As peneplains formed near to absolute base level (Phillips, 2002) and at present the flat-shaped surfaces are

located at high elevations, they represent uplift markers and indicate > 2 km of regional uplift of the entire

forearc during Late Neogene (Farías et al., 2006, 2008). The outcrops of Late Neogene sediments at Central

Chile that could represent the erosion material related to the regional uplift event are the Navidad, Lincancheo,

Rapel y La Cueva marine formations (Encinas, 2006) (Fig 1). Standard detrital heavy mineral (HM) analysis and

single grain microprobe analysis of detrital heavy minerals were carried out to recognize the lithological units

subjected to erosion during the deposition of each marine formation, allowing us to reach a better understanding

of the paleogeographic evolution of Central Chile forearc during Late Neogene times.

Results and conclusions

The principal differences in the sources areas of Late Neogene marine formations are indicated by the detrital

volcanic and metamorphic HM suite and are reflected in the composition of pyroxene-amphibole and garnet

respectively. For sandstones deposited between 12.8 and 4.6 Ma (deposition of lower levels of Navidad

Formation), the geochemical characteristics of pyroxenes (Fig 2) indicate a source formed mainly by tholeittic

and calcoalkaline volcanic rocks, like rocks from Abanico West Formation at eastern CD (ECD) south of 35ºS

(López-Escobar y Vergara, 1997), while the composition of garnet is typical of garnet of metamorphic rocks

from the west part of the CC (WCC). These features show that during the indicated 12.8 - 4.6 Ma period two

main reliefs formed the Central Chile forearc: one located at the actual WCC and the other at the actual ECD.

Figure 2. Leterrier et al. (1982) diagrams for clinopyroxene from a) lower levels (orange symbols) and upper levels (red symbols) of Navidad Formation and b) Lincancheo (yellow and dark green symbols), Rapel (pale green symbols) and La Cueva (purple, cyan and blue symbols) formations.

By 4.6 Ma (deposition of upper levels of Navidad Formation), the provenance of sandstones is determined by

the alkalis and AlIV contents of detritical volcanic amphibole (Fig 3), which mimics the contents of these

elements in amphibole from the hypabyssal volcanic basement rocks of the ECD (Manquehue type rocks). The

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absence of the Abanico West Formation source and the erosion of the ECD basement indicates that by this time

the Central Depression was being formed.

Figure 3 Alkalis and AlIV contents on amphibole from upper levels of Navidad Formation (red symbols), Lincancheo Formation (yellow symbols), Rapel Formation (pale green symbols) and La Cueva Formation (purple and blue symbols). The pale grey field represent the composition of amphibole from Manquehue type hypabissal rocks (unpublished data from Daniel Sellés) and the dark grey field represent the composition of amphibole from Farellones Formation (Fuentes, 2004).

The detrital volcanic HM associations of sandstones deposited between 4.6 and 2.7 Ma (deposition of

Lincancheo, Rapel and La Cueva Navidad formations) are characterized by the geochemical characteristics of

both pyroxene and amphibole. A main source represented by calcoalkaline intermediate to basic rocks, like

outcrops of Farellones Formation from western CC (WCC) is indicated by Leterrier diagrams in pyroxene (Fig

2) and alkalis and AlIV contents of amphibole (Fig 3). A secondary source represented by volcanic and contact

metamorphic rocks from the eastern CC (ECC) is also recognized by Leterrier diagrams in pyroxene and the

presence of garnet from the andradite-grosular series respectively. The rock sources identified show that between

4.6 and 2.7 Ma the reliefs subjected to erosion at the forearc of Central Chile were the WCP and the ECC.

Acknowledgments We acknowledge funding by the FCFM (Proyecto de inserción), Departamento de Investigación y Desarrollo (DI), Universidad de Chile (Project DI 2004, I2 04/02-2), IRD, Bourse Amerique Latine (UPS) and Proyecto Anillo ACT Nº 18.

References Borde, J., 1966. Les Andes de Santiago et leur avant-pays: étude de géomorphologie. Bordeaux, France. Union Française

d’Impression, 559 p. Brüggen, H., 1950. Fundamentos de la Geología de Chile. Santiago, Chile. Instituto Geográfico Militar, 510 p. Encinas, A., 2006. Estratigrafi a y sedimentologi a de los depo sitos marinos miopliocenos del área de Navidad (33ºS-

34º30´S), Chile Central: implicaciones con respecto a la tecto nica del antearco. Tesis, Doctor en Ciencias, mencio n Geologi a. Universidad de Chile.

Farías, M., Charrier, R., Carretier, S., Martinod, J., Comte, D., 2006. “Erosión versus tectónica en el origen de la Depresión Central de Chile Central” In Actas XI Congreso Geológico Chileno. 7-11 Agosto, Antofagasta, Chile.

Farías, M., Charrier, R., Carretier, S., Martinod, J., Fock, A., Campbell, D., Cáceres, J., Comte, D., 2008. Late Miocene high and rapad surface uplift and its erosional response in the Andes of Central Cehile (33-35ºS), Tectonics, 27, TC 1005, doi: 10.1029/2006TC002046

Fuentes, F., 2004. Petrologi a y metamorfismo de muy bajo grado de unidades volca nicas oligoceno-miocenas en la ladera occidental de Los Andes de Chile Central (33ºS). Tesis, Doctor en Ciencias, mencio n Geologi a. Universidad de Chile.

Leterrier J., Maury, R., Thonon, P., Girard, D., Marchal, M., 1982. Clinopyroxene composition as a method of identification of the magmatic affinities of paleo-volcanic series. Earth and Planetary Sciences Letters, 59, 139-154.

López-Escobar, L., Vergara, M., 1997. Eocene-Miocene Longitudinal Depression and Quaternary volcanism in the Southern Andes, Chile (33-42.5°S): a geochemical comparision. Revista Geológica de Chile, 24 (2) 227-244.

Phillips, J.D., 2002. Erosion, isostatic response, and the missing peneplains, Geomorphology, 45 (3-4), 225-241.

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The Loncopué Trough: A major orogenic collapse in the western Agrio fold-and-thrust belt (Andes of Neuquén, 36º40´-38º40´S)

Emilio Rojas-Vera1, Andrés Folguera

1, Gonzalo Zamora-Valcarce

2, & Victor A. Ramos

1

1 Laboratorio de tectónica andina, CONICET-UBA, 2º Pabellón, Ciudad Universitaria, Buenos Aires, Argentina

(emilio@gl.fcen.uba.ar) 2 Repsol-YPF, Exploración Neuquén, Argentina

KEYWORDS : neotectonics, extensional retroarc basin, strike-slip faults, orogenic collapse, basament-controlled structure

Introduction

The Pliocene to Quaternary Loncopué trough is located in the Andean retroarc zone between 38º 40’ and 36º

40’S approximately, east of the main Andes and the Present volcanic arc west of and the Agrio fold and thrust

belt (Figure 1). Neotectonic activity has been proposed in the area by several authors (Ramos and Folguera 2005,

García Morabito et al. 2005, Foguera et al. 2006) who based on morphotectonic features have interpreted as

controlled by normal faults located between the Agrio fold and thrust belt and the Loncopué trough (Figure 1).

However, scarce to none evidence of faulted Quaternary materials has been described. More recently, Yuan et al.

(2006) determined a crustal attenuated area beneath the Loncopué trough using receiver function techniques.

Zapata et al. (1999), Zamora Valcarce et al. (2006) and Jordan et al. (2001) based on limited seismic reflection

and field data, proposed a basement west-dipping normal fault controlling a Late Oligocene depocenter to the

west beneath the Loncopué trough, in coincidence with that area of reported young tectonic activity. However,

Cobbold and Rossello (2003) interpreted this limit as produced by a major backthrust that overrides the

Mesozoic sequences over Tertiary strata. In this stuy we present field evidence on the transtensional nature of the

limit between the Agrio fold and thrust belt and the Loncopué trough, as well as the complex structure of the

axial part of the trough that was formed by a series of ten-of- kilometers wide pull-apart depocenters that

evolved during Quaternary times.

Agrio fold and thrust belt

This deformational belt has been divided in two sectors with contrasting geology and structural styles. The

western sector near the Loncopué trough (Figure 1) is characterized by the strong inversion of Late Triassic to

Early Jurassic normal faults uplifting synrift sequences of the Neuquén basin. Compressional deformation begun

in the Late Cretaceous times (Zapata et al. 1999, Zamora Valcarce et al., 2006). Thin-skinned deformation

becomes more important at the eastern sector where post Late Jurassic sequences are detached from

Kimmeridgian evaporites which form the main decollment in the area. Late Miocene contraction has modified

Late Cretaceous uplifts as revealed by the development of synorogenic depocenters (Zapata et al 2002, Zamora

Valcarce et al. 2006).

Loncopué trough

Extensional tectonics associated with the Loncopué trough is superimposed to the western Agrio Fold and

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thrust belt to the east (eastern Loncopué fault system), where mainly Jurassic sequences and Quaternary volcanic

products are extensionally deformed and is expanded up to the eastern Present volcanic arc to the west (western

Loncopué fault system). To the north, this trough is limited by the NE Mandolegüe arch, a basement cored uplift

composed of Eocene rocks, partially covered by Quaternary retroarc volcanic products (Figure 1). Kinematic

indicators at the eastern Loncopué fault system show left lateral displacements associated with extensional

components at NS fault planes. Transtensional deformation has affected the axial part of a series of basement

cored uplifts such as Agua Fría and Moncol anticlines, producing asymmetric grabens (Figure 2) that controlled

the emplacement of monogenetic field basalts. These rocks were lately affected by transtensional deformation

that produced 5-10 meter scarps superimposed to faults affecting Mesozoic strata with normal displacements.

Alluvial deposits are locally faulted indicating ongoing orogenic collapse at the western Agrio fold and thrust

belt.

Figure 1. Argentine Andes between 37º45´and 38º30´S, where the main Andes are separated from the Agrio fold and thrust belt by the Loncopué trough. The eastern sector is modified from Ramos (1998).

Western Loncopué fault system is formed by a rectilineous NS east-dipping set of normal faults that deform

Pliocene to Quaternary volcanic sequences. These faults are associated with ignimbritic eruptions, dome and

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caldera complexes located immediately to the east of the Present volcanic arc (Figure 1). Geometry of pull-apart

basins associated with those structures (Codihue depocenter) indicates left lateral components along the edges of

the trough.

The central part of the trough is formed by a series of rhomboedric depocenters that are filled by basaltic

eruptions interfingered with lacustrine deposits. These depocenters are limited by two fault systems, one NE-

trending another perpendicular to it. Geometry and the association to long NNW fault systems that diverge from

the western Loncopué trough, suggest that they could be a response to right-lateral displacements, which

segmented the axial trough (Figure 1). The northernmost of these features is the Huecú depocenter (Figure 1)

that has controlled the eruption of postglacial products all along the retroarc basin (Rojas Vera, 2007). There are

associated with liquefaction phenomena, in youngest lacustrine deposits interfingered with the youngest lavas in

the area, which indicates strong paleoseismic events in the region.

Structural cross-section at 38º10´S

The structural cross-section has been built with limited seismic information mainly restricted to the Agrio fold-

and-thrust belt and to the axial central part of the Loncopué trough. Field data and gravimetric and

magnetometric information is available for the entire trough. Gravimetric analyses show that the western

Loncopué trough controls a nearly 8 km deep depocenter, which implies that at least 3 km of the sedimentary

column can not be explained by Tertiary to Quaternary strata. Therefore a Mesozoic depocenter is expected

beneath the thickest part of the Loncopué basin (western Loncopué fault system), next to the main Andes (Figure

2), this depocenter has no relation to the morphologically more prominent East Loncopué fault system, where

most previous studies considered the thickest section of the basin.

Figure 2. Cross section at 38º10´S, based at the Loncopué trough on surface data, gravimetric and magnetometric surveys and limited coverage of seismic lines. The structural style of the Cerro Mocho area is controlled by seismic data (Zamora Valcarce et al. 2006).

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Contrastingly, field data, and gravity and limited seismic information show that the axial part of the trough

would controls a smaller depocenter associated with both east and west dipping extensional fault systems. To the

eastern extensional structure is localized by the westernmost basement uplifts of the Agrio fold and thrust belt

and do not expand over the Cerro Mocho basement uplift (Figure 2). This could indicate that the structure of the

Agua Fría and Moncol anticlines has locally relaxed in order to produce extensional morphologies at surface.

Concluding remarks

The Loncopué structure is intimately linked to the Late Triassic rift architecture slightly modified during Late

Cretaceous compressional stage. Most of the rhomboedric features which are aligned at the axial part of the

trough coincide with gravimetric lows determined between basement highes which are exposed immediately to

the west in the Agrio fold and thrust belt. However, Mesozoic depocenters do not necessarily coincide with

Tertiary to Quaternary areas of maximum thickness. Silicic to mafic volcanism in the area is controlled by

transtensional structures both in right and left lateral pull apart basins. Last deformation can be considered as

Holocene based on faulted alluvial fans at the eastern sector of the trough. On the other hand western

deformation seems to be fossilized by Pleistocene caldera complexes. Further work is necessary in order to

refine kinematics of the main fault systems.

References Cobbold, P.R., Rossello, E.A., 2003. Aptian to Recent compressional deformation, foothills of the Neuquén Basin,

Argentina. Marine and Petroleum Geology, v. 20, no 5, p. 429-443. Folguera, A., Zapata, T., Ramos, V.A., 2006. Late cenozoic extension and the evolution of the Neuquén Andes, in Kay, S.M.,

and Ramos, V.A., (eds): Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35º-39º lat): Geological Society of America Special Paper 407:267-285.

García-Morabito, E., Folguera, A. 2005. El alto de Copahue – Pino Hachado y la fosa de Loncopué: un comportamiento tectónico episódico, andes neuquinos (37°-39°S). Revista de la Asociación Geológica Argentina 60 (4): 742-761, Buenos Aires.

Jordan, T., Burns, W., Veiga, R., Pangaro, F., Copeland, P., Kelley, S., Mpodozis, C., 2001. Extention and basin formation in the southern Andes caused by increased convergence rate: A mid-Cenozoic trigger for the Andes. Tectonics 20 (3): 308-324.

Ramos, V.A..1998. Estructura del sector occidental de la faja plegada y corrida del Agrio, cuenca Neuquina, in X Congreso Latinomericano de Geología (Buenos Aires): Actas, v. II, p. 105-110.

Ramos, V., Folguera, A. 2005. Tectonic evolution of the Andes of Neuquén: constraints derived from the magmatic arc and foreland deformation, en Spalletti, L., Veiga,G., Schwarz, E. y Howell, A. (eds). The Neuquén Basin: A case study in sequence stratugraphy and basin dynamics: Geological Society of London Special Publications, 252: 15-35.

Rojas-Vera, E. 2007. Estudio tectónico del sistema de fallas de Mandolegüe: La cuenca cuaternaria del Huecú (37°43’S, 70°41’O) Provincia de Neuquén. Tesis de licenciatura, Universidad de Buenos Aires, 89 p.

Yuan, X., Asch, G., Bataille K., Bock, G., Bohm, M., Echtler, H., Kind, R., Oncken, O., Wólbern, I., 2006. Deep seimic images of the Southern Andes, in Kay, S.M., & Ramos, V.A. (eds): Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35º-39ºS): Geological Society of America Special Paper 407: 61-72.

Zamora-Valcarce, G., Zapata, T., Del Pino, D., Ansa, A., 2006. Structural evolution and magmatic characteristics of the Agrio Fol.-and-thrust belt in Kay, S.M., and Ramos, V.A. (eds): Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35º-39º lat): Geological Society of America Special Paper 407:125-145.

Zapata, T., Brissón, I., Dzelalija, F., 1999. The role of basement in the Andean fold and thrust belt of the Neuquén Basin, en K. McClay, (ed.): Thrust Tectonics, 122 – 124, Chapman and Hall, New York.

Zapata, T. R., Corsico, S., Dzelalija, F., Zamora-Valcarce, G., 2002. La faja plegada y corrida del Agrio: Análisis estructural y su relación con los estratos Terciarios de la Cuenca Nequina, in 5° Congreso de exploración y desarrollo de Hidrocarburos. Actas electronicas. Mar del Plata.

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The Cordillera Blanca fault system as structural control of the Jurassic-Cretaceous basin in central-northern Peru

Darwin Romero

INGEMMET, Av. Canadá 1470, San Borja, Lima, Peru (romerofdarwin@gmail.com)

KEYWORDS : Cordillera Blanca fault system, Jurassic-Cretaceous basin, central-northern Peru

Introduction

Many works related to the Cordillera Blanca fault system exist (e.g. Bonnot, 1984; McNulty and Farber,

2002). These studies were mainly in the central part of the Cordillera Blanca fault, between northern Yungay and

southern Recuay, without taking account the northern outcrops in the Pallasca zone and to the south the

Cajatambo zone. The present study presents a new interpretation of the Cordillera Blanca fault system, based on

stratigraphic and structural observations of the Jurassic-Cretaceous Chicama-Goyllarisquizga basin in central-

northern Peru between Pallasca-Huaraz-Cajatambo, which form part the Cordillera Blanca fault system (CBFS).

Structure and stratigraphy of the Chicama-Goyllarisquizga basin

The termed deposits Chicama Group (Middle-Upper Jurassic) and Goyllarisquizga Group (Berriasian-Aptian),

define the termed Jurassic-Cretaceous Chicama-Goyllarisquizga basin in central and southern Peru. This basin is

part of the Western Cordillera and to be more precise corresponds to the Cordillera Negra and Blanca. In the

north the basin surrounds the Pallasca, Corongo and Huaylas areas, central the Huaraz, Recuay and Aija areas,

and south the Cajatambo, Oyon and Churin areas. The basin basement has not been possible determinate.

However, in the Aija and Churín area, along the anticline core has been observed ignimbrites intercalated with

volcanic breccias, probably corresponding to the Oyotun Formation of Lower Jurassic age.

In the central-northern Peru (8° 30’ a 10° 30’), we divided the zone in three stratigraphic basins (Figure 1):

The Cretaceous volcano-sedimentary Casma basin (KVSCB), the Jurassic-Cretaceous Chicama-Goyllarisquizga

basin, and the Permian-Triassic Mitu-Pucara basin.

The Jurassic-Cretaceous Chicama-Goyllarisquizga basin to the west is limited with the volcano-sedimentay

basin by the Tapacocha fault system and to the east is limited with the Permian-Triassic Mitu-Pucara basin by

the Chonta fault systems.

The Jurassic-Cretaceous Chicama-Goyllarisquizga basin is characterized by ignimbrites and volcanic breccias

of the Oyotun Formation (Lower Jurassic), sandstone sequences intercalated with mudstone to the top and

bottom of the Chicama Formation (Middle Upper Jurassic). Those follow by the deposits of the Goyllarisquizga

Group (Berriasian-Aptian) characterized by sandstones intercalated with mudstones and limestones, changing to

quartz rich sandstones of the Chimú Formation, limestones with mudstones of the Santa Formation, developing

to quartz rich sandstones, grauwacas intercalated with gray-red-green mudstones of the Carhuaz Formation,

ending in white quarzt rich sanstones of the Farrat Formation. Finally, we observe the carbonate sequence

(Albian-Campanian) characterize by the Parihuanca, Chúlec, Pariatambo, Jumasha and Celendín units.

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The Cretaceous volcano-sedimentary Casma basin basin is localized to the west of the study area and the

boundary with the Chicama-Goyllarisquizga basin correspond to the Tapacocha fault. This basin is characterized

by mudstones intercalated with chert, ignimbrites and limestones of the Cochapunta Formation (Albian-

Cenomanian).

The Permian-Triassic Mitu-Pucara basin is limited to the west with the Chicama-Goyllarisquizga basin by the

Chonta fault. In this basin the Paleozoic-Precámbriam basement overlap with unconformity the sandstones and

conglomerates, red mudstones intercalation of Mitu Group (Upper Permic-Middle Triassic), limestones of

Pucara Group (Upper Triassic-Lower Jurassic) on the top of this deposits and with erosional angular

unconformity are the red mudstones and sanstones intercalations that developed from white quartz rich

sandstones to conglomerates of the Goyllarisquizga Formation (Berrisian-Aptian?) and to the end we observe

limestones sequences of the Chulec-Pariatambo Formation (Albian).

Figure 1. Structural section D-D’, showing the three Stratigraphic basins.

Structural controls of the Jurassic-Cretaceous Chicama-Goyllarisquizga basin.

This basin is limited to the west by three reverse fault systems with vergence to west: 1) Tapacocha, 2)

Huacllan-Churín, and 3) Huaraz-Recuay faults systems; and to east is limited by two reverse fault systems with

vergence to east: 1) Cordillera Blanca and Chonta fault systems. We will now describe the structural sections

that cross the Cordillera Blanca fault system.

The section A-A’ is located to the northern, has E to W direction, between Cabana and Pallasca towns. We

observe the Huaraz-Recuay reverse fault systems with west dip, outcropping the Goyllarisquizga Group rocks

and overlying the Tablacacha sequence (Upper Cretaceous-Paleocene, Navarro et al., in preparation). The

sediment deformation of the Tablachaca sequence corresponds to folds with west dip. In this area the fold has

NE-SW direction and the Huaraz Recuay fault system has NNE-SSO direction, thus indicate a reverse sinistral

motion for the Huaraz Recuay fault system.

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Figure 2.

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The structural section B-B’ is located in the central southern part of Cordillera Blanca fault system near

Recuay town, from NE to SO direction, here we observe the Huaraz Recuay fault system as reverse with west

dip and the Cordillera Blanca fault system as reverse with east dip, which due tectonic inversion actually show

sinistral motion with normal component and generating a Plio-Qauternary basin with horst and graben showing

NW to SE direction.

The section C-C’ is located to the south, has NE to SW direction, near to Cajatambo town. Here we can

observe the three western fault systems of the JKCHGB, corresponding reverse faults with west dip, affecting

Jurassic rocks and the volcano-sedimentary sequence of Upper Cretaceous. This section is very important

because we can observe in the HCH fault system a positive flower structure by tectonic inversion.

The section D-D’, located in the central-southern part of Jurassic-Cretaceous Chicama-Goyllarisquizga basin,

has a NE to SO direction, between Chiquian and Recuay towns. This section is more regional, cross the three

stratigraphic basins and all fault systems that controlled the JKChGB. To the southwest we observe the

Cochapunta Formation (Albian-Cenomanian) of the KVSCB basin and limited by the TFS. In the central part we

observe the Jurassic-Cretaceous sediments controlled to the west by the Tapacocha, Huacllan-Churin, Huaraz

Recuay fault systems, and to the east controlled by the Chonta fault system. These fault systems show clear

distensive tectonic inversion to compressive. However, between the Huaraz-Recuay and Cordillera Blanca fault

systems we observe sinistral motion with normal component that affect Plio-Quaternary deposits. Toward NE

the Goyllarisquizga Formation is overlying with angular unconformity the Mitu-Pucara Group and the

Paleozoic-Precambrian basement in the Mitu-Pucara basin. Therefore, this basin corresponds to a horst during

the Cretaceous.

Conclusions

From the stratigraphic and structural analyses, we interpret that the Jurassic-Cretaceous Chicama-

Goyllarisquizga basin was originated and controlled by the Tapacocha, Huacllan-Churín, and Huaraz-Recuay

fault systems in the western boundary and by the Cordillera Blanca and Chonta fault systems in the eastern

boundary of the basin. These faults at the beginning have presented normal motion, later due to compressive

tectonic inversion change to reverse fault with west and east dip. Along the Chicama-Goyllarisquizga basin axes,

limited by Cordillera Blanca and Huaraz-Recuay faults systems, we observe sinistral slip with normal

component affecting Plio-Quaternary deposits. This last tectonic style indicates sinistral transtensive motion for

the Cordillera Blanca zone.

References Bonnot D., 1984 — Neotectonique et Tectonique active de la Cordillere Blanche et du callejon de Huaylas, Andes nord-

péruviennes. (Ph.D.), Orsay, Universite de Paris 115p. McNulty & Farber, 2002 — Active detachment faulting above the peruvian flat slab. Geology, v.30, p.567-570. Navarro P., Rivera M., Monge R., (in preparation) — Estratigrafía del Volcanismo Cenozoico (Grupo Calipuy), Cordillera

Occidental del norte del Perú, departamentos La Libertad y Ancash 7º 30’ – 9º 00’ latitud sur. Boletín INGEMMET.

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Block rotations within the northern Peruvian Altiplano

Pierrick Roperch1, Víctor Carlotto

2, & Annick Chauvin

3

1 IRD, LMTG & Géosciences Rennes, campus de Beaulieu, 35042 Rennes, France (pierrick.roperch@ird.fr)

2 INGEMMET, Av. Canada 1470, San Borja, Lima 41 & UNSAAC, Cusco, Peru (vcarlotto@ingemmet.gob.pe)

3 Géosciences Rennes, campus de Beaulieu, 35042 Rennes, France (annick.chauvin@univ-rennes1.fr)

KEYWORDS : central andes, tectonics, paleomagnetism, rotation

Introduction

Counterclockwise tectonic rotations in the northern central Andes and clockwise rotations in the southern

Central Andes have been systematically reported (see Roperch et al. 2006 and Arriagada et al., 2006 for a recent

summary) and interpreted to be mostly driven by oroclinal bending associated with shortening in the Eastern

Cordillera and in the subandean belt. Large counterclockwise rotations have been found in the Eastern Cordillera

of Southern Peru (Gilder et al. 2003). While these rotations were initially attributed to a Cretaceous event of

deformation, Gilder et al. (2003) interpreted these rotations to be coeval with the rotations found along the

forearc (Roperch et al., 2006). Rotations along the forearc from Arequipa to Caravelli are larger than 40° and

occurred mainly during the late Eocene - Oligocene. However, the lack of data within the Peruvian Altiplano

precludes a good description of the spatial and temporal evolution as well as a clear understanding of the

different tectonic processes leading to rotations. Here we report results from a paleomagnetic study from Nazca

to Cusco (Figure 1). This transect corresponds to the location of the northern end of the Altiplano and a

transition with the central Peruvian Andes. Near Cusco, the Eastern Cordillera is also strongly deflected toward

the east with a complex deformation as shown by the curved fold and thrust system associated with the Manu

Indenter.

Paleomagnetic results

Paleomagnetic sampling

Fifty one paleomagnetic sites have been sampled from Nazca to Cusco. Six sites were drilled in lower Miocene

Nazca ignimbrites that cover the forearc between Nazca and Puquio. Near Puquio, 5 sites were drilled in undated

tertiary volcanics that are overlain by the Lower Miocene Nazca tuffs. Four sites were also sampled in upper

Miocene ignimbrites and two sites corresponds to a Semca Ignimbrite within the Altiplano. Six sites were drilled

in Paleocene to Eocene red bed sediments intruded by dykes north of Chalhuanta. Near Cusco, 9 sites were

sampled in Oligocene red beds from the Sonco Formation of the San Jeronimo group. While a late Cretaceous

age was initially attributed to the thick red beds sequences that outcrop in the southern Peruvian Altiplano, a mid

Oligocene 40Ar-39Ar age at the top of the sequence provides an upper age limit for the red bed sequence

(Carlotto, 1998). Several sites were also drilled in red beds either contemporenaeous or slightly older than the

San Jeronimo group (3 sites west of Sicuani) or in the Anta Formation.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 469-472

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Figure 1. Simplified geological map from Peru (modified from the 1/1000000 Peruvian map from Ingemmet ; see also Roperch et al. 2006) The paleomagnetic sampling was carried from Nazca to Cusco. a) sites drilled in the Sonco Formation (results reported in Figure 2); b) sites drilled in Paleocene- Eocene red beds and dyke near Chalhuanca (Figure 4); c) sites drilled in other Paleogene red beds near Cusco (Figure 3) ; d) sites drilled in Oligocene – Miocene volcanics near Puquio (Figure 5) ; c) sites (results not discussed). The dashed line is the northern boundary of the Arequipa domain identified by geochemical data (Mamani et al., 2008). Angles between the arrows and the north direction correspond to tectonic rotations. New results shown with white confidence interval. The colour (yellow, orange, magenta, cyan) correspond to age of rock unit (Miocene, Oligocene, Paleogene, Mesozoic).

Characteristic directions

All the samples in red beds were progressively thermally demagnetized. Except at a few sites where it was

possible to isolate a secondary magnetization, pre-tectonic components of magnetization were determined at

most sites. Near Cusco, all the samples drilled in 9 sites from the Sonco Formation have a reverse polarity

(Figure 2). The magnetic characteristic upon demagnetization and the better grouping after tilt correction

demonstrates that the magnetization is a primary component. Although the sampling is not stratigraphically

continuous enough for magnetostratigraphic dating, the lack of samples with normal polarity suggests high

deposition rate during a time interval dominated by reverse polarity, possibly during the late Eocene – early

Oligocene (35-31Ma).

Figure 2. Paleomagnetic results in red beds in the type section of the San Jeronimo group located to the south of Cusco. Equal areal projection of site-mean directions with the 95 angle of confidence. Filled circles are projection in the lower hemisphere. Filled red square is the expected late Eocene-early Oligocene direction for stable south America. The filled red circle and associated angle of confidence correspond to the mean direction.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 469-472

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Normal and reverse polarity magnetizations were found in other red bed units at different localities (Figure 3).

After bedding correction the observed inclinations is in good agreement with the expected inclination but large

counterclockwise deviations of the declinations are observed.

The dykes sampled north of Chalhuanca do not record a primary component of magnetization but a well-

defined primary magnetization was observed in the sediments baked contacts. Away from the dyke contacts,

characteristic directions were also determined in the red bed sediments (Figure 4).

Sites in volcanic rocks near Puquio have a well-defined primary characteristic magnetization of normal or

reverse polarity (Figure 5). Sampling in mid Miocene to Pliocene rocks is not sufficient to enable an accurate

averaging of secular variation. The average direction calculated for sites in the lower Miocene ignimbrites and

Figure 3. Paleomagnetic results in red beds at other sites near Cusco. Equal areal projection of site-mean directions with the 95 angle of confidence. Filled (open) symbols are projection in the lower (upper) hemisphere. The squares correspond to the late Eocene-early Oligocene expected directions.

Figure 4. Paleomagnetic results in red beds and dykes from the Paleocene-Eocene basin located to the north of Chalhuanca. Equal areal projection of site-mean directions with the 95 angle of confidence. The grey circle correspond to a remagnetization of unknown age recorded by pyrrothite in a site of Jurassic limestone south of Chalhuanca.

Figure 5. Paleomagnetic results in the lower Miocene Nazca ignimbrites and the underlying volcanics sampled near of Puquio. The red circle with angle of confidence is the mean direction calculated for these sites. Results highlighted in orange colour corresponds to younger Miocene or Pliocene volcanics and are not included in the mean calculation. The red square is the expected direction. for the late Oligocene Early Miocene.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 469-472

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underlying volcanics has an inclination similar to the expected direction.

Tectonic rotations

Near Cusco, large counterclockwise rotations are observed at some Paleogene sites while only 10° rotation is

determined in the Sonco Formation. Large rotations were also reported for the Permian to lower Jurassic

sediments (Gilder et al., 2002) in the Eastern Cordillera at a distance of about 15km to the east of the sites

sampled in the Sonco section. Taking into account that the upper part of the Sonco section is of mid-Oligocene

age, the result in the Sonco formation may provide an upper bound in the age of the rotation. Further work is

however needed to understand the cause of the large variation in the magnitude of the rotations.

North of Chalhuanca, the preliminary results from the late Paleocene- Eocene red bed basin indicate that this

area is affected by large counterclockwise rotation (~60°). In the Chalhuanca region to the west of the Eocene

basin, fold axes and faults in Mesozoic units have an anomalous WSW-ENE orientation. These structures can be

traced for more than 50km. Despite the fact that the Eocene basin presents little folding, the sediments record a

low magnitude anisotropy of magnetic susceptibility whose lineations are also oriented WSW-ENE. We

speculate that the ~60° counterclockwise rotation recorded in the Eocene basin corresponds to the block rotation

of a larger block postdating or associated with the main event of folding in the region. The Nazca ignimbrites

and underlying volcanics record a <5° counterclockwise rotation in agreement with the result previously

published by Roperch et al. (2006) indicating almost no block rotation within the forearc during the Neogene

even in front of the Nazca ridge.

The new available results confirm the widespread occurence of large (>40°) counterclockwise tectonic

rotations in southern Peru during a major phase of deformation in the time interval late Eocene- early Oligocene.

Geochemical data (Carlier et al., 2005; Mamani et al., 2008) show significant crustal and lithosperic

heterogeneities in the Central Andes and especially in Southern Peru. Mamani et al. (2008) argue that the

different rheologies are an important factor in controlling the deformation pattern of the central Andes and the

localization of the Andean plateau. They suggest that most of the rotations are the results of intense shearing

along the border of the Arequipa block. While the geochemical data suggest a sharp boundary between the

Arequipa block and the Central Peruvian Andes, large rotations are however observed in a much wider area

north and south of the geochemical boundary (Figure 1).

References Arriagada, C., Roperch, P., Mpodozis C., & Fernández, R. 2006. Paleomagnetism and tectonics of the southern Atacama

Desert region (25-28ºS) Northern Chile. Tectonics, 25: TC4001, doi:10.1029/2005TC001923. Carlier, G., Lorand, J.P., Liegeois, J.P., Fornari, M., Soler, P., Carlotto, V. & Cardenas, J. 2005. Potassic-ultrapotassic mafic

rocks delineate two lithospheric mantle blocks beneath the southern Peruvian Altiplano, Geology, 33: 601–604 doi: 10.1130/G21643.1

Carlotto, V. 1998. Evolution andine et raccourcissement au niveau de Cusco (13°-16°S, Pérou), Thèse Doctorat, Univ. Grenoble, 159p.

Mamani, M., Tassara, A. & G. Woerner, 2008. Composition and structural control of crustal domains in the central Andes, Geochem. Geophys. Geosys., doi:10.1029/2007GC001925.

Gilder, S., Rousse, S., Farber, D., Sempere, T., Torres, V., & O. Palacios 2003. Post-Middle Oligocene origin of paleomagnetic rotations in Upper Permian to Lower Jurassic rocks from northern and southern Peru, Earth and Planetary Science Letters, 210: 233-248.

Roperch, P., Sempere, T., Macedo, O., Arriagada, C., Fornari, M., Tapia, C., García, M. & C. Laj, 2006. Counterclockwise rotation of late Eocene – Oligocene fore-arc deposits in southern Peru and its significance for oroclinal bending in the central Andes, Tectonics, 25: TC3010, doi:10.1029/2005TC001882.

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From steady state to climate-driven denudation across the Central Andes in SE Peru

Geoffrey M.H. Ruiz1,2

& Víctor Carlotto3

1 University of Neuchatel, Switzerland (Geoffrey.ruiz@unine.ch)

2 ETH Zurich, Switzerland

3 INGEMMET, Lima, Peru

To better constrain the orogenic growth of the Andean chain, we investigated the time-Temperature paths of

bedrocks from the two morpho-structural highs of the Central Andes that are separated by the vanishing

Altiplano, i.e. the Eastern and Western Cordilleras of SE Peru.

The Western Cordillera is a volcanic to volcano-detrital chain that developed ~40-35 Ma ago and is

characterized by a 4000m high mean altitude whose origin is poorly constrained. Fission-Track data on apatite

and zircon crystals extracted from an Eocene pluton yield ages comprised between 24 and 14 Ma, and 38 and 30

Ma respectively. One of the noteworthy aspects of the data is that analyses reveal a steady-state phase of

exhumation from the late Eocene to at least the middle Miocene (38-14 Ma) with no disruption of the

exhumation path since 38 Ma either by sedimentary burial and/or rapid exhumation. The uplift of the Western

Cordillera was thus probably steady since, avoiding the deposition of foreland basin sequences as in the

Altiplano region. Further east, Apatite Fission-Track ages are much younger and range between 7.6 and 2.5 Ma

for the Eastern Cordillera and between 11.2 and 1.5 Ma for the Sub-Andean Zone. Age-altitude relationships

suggest that denudation increased from a more quiescent Late Miocene period to a high rate of 0.9 km/my for the

Pliocene. Such abrupt change is supported by a net in sediment accumulation rates in the Andean Amazon Basin

but as far as offshore the Amazon fan. A global climate change is usually invoked for high Pliocene rates;

however it post-dated a documented period of surface uplift in the Eastern Andes.

Denudation patterns are thus much contrasted across the Andes of SE Peru. The western Cordillera, despite

significant topography and deep river valleys in the studied area, still yield information that suggest a steady and

slow uplift from the late Eocene until at least the middle Miocene. We thus propose a coupled scenario: first the

Andean orographic barrier developed from the Eocene by tectonism as more recently in the eastern Cordillera

(Late Miocene), later focusing the Amazon moisture (5m/y of annual precipitation today) and as a result

denudation since the Pliocene along the eastern flank of the Andes. The localization of erosion modified in turn

the structure of the belt, limiting the deformation in the narrow Sub-Andean Thrust Belt.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 474-476

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Pleistocene uplift rates variability along the Andean active margin inferred from marine terraces

M. Saillard1, L. Audin

1, G. Hérail

1, S. Hall

2, D. Farber

2,3, J. Martinod

1 , & V. Regard

1

1 LMTG, Université de Toulouse, CNRS, IRD, OMP, 14 Av. E. Belin, F-31400 Toulouse, France

2 University of California, Santa Cruz, Dept. of Earth Sciences, Santa Cruz, CA 95060, USA

3 Lawrence Livermore National Laboratory, LLNL, Livermore, CA 94550, USA

KEYWORDS : Beryllium-10, marine terrace, Pleistocene, uplift rate, Central Andes

Introduction

The southern Pacific coast morphology and especially the presence of marine terraces give information on the

dynamics of Andean forearc evolution from the Pleistocene. Marine terraces preserve a record of eustatic sea

level changes together with the uplift history of the coastal area in the Andean forearc. Along most of the

Southern Peru and Northern Chilean coasts, discontinuous uplifts are recorded by wave-built terraces and wave-

cut platforms. To investigate these processes, we studied sequences of marine terraces in various coastal sections

either in southern Peru or in Chile: at Chala (15°50'S) and Ilo (17°32'S-17°48'S), situated above a steep

subduction segment and at San Juan de Marcona (15°20'S), situated above the southern part of the Nazca ridge,

and along the coastal part of the Altos de Talinay area, from Tongoy (30°15'S) to ~31°20'S, situated above a flat

subduction segment (Figures 1, 2 and 3).

We chose various sites in order to sample possibly different response of the continental plate to the subduction

process. Various studies were already undertaken on such problems either in Peru or Chile but mainly leaded to

the datation of the 5th isotopic stage. In this study, differential GPS and cosmogenic datations are pursued in

order to propose robust and absolute ages on these sites and subtract the effects of eustatic sea-level changes

from local curves, identifying tectonic uplifts. We dated ~15 levels of marine terraces either in Peru or Chile.

Results

In Chile, we show that, since 700 ka, Pleistocene uplift rates have been highly variable along the Andean

margin near 31°S (Saillard et al., submitted). The uplift of the Chilean forearc has been recorded by a sequence

of five wave-cut platforms that have been dated using in situ produced 10Be. These platforms formed during

interglacial periods corresponding to marine isotopic stages (MIS) or substages (MISS) 1, 5e, 7e, 9c and 17. Our

mapping in conjunction with the new chronology we present shows that the surface uplift rate varied from

103 ± 69 mm/ka between 122 and 6 ka, to 1158 ± 416 mm/ka between 321 and 232 ka. The absence of preserved

marine terraces related to the MIS 11, 13 and 15 highstands likely reflects slow uplift rates during these times.

Consequently, we propose that this area essentially uplifted during 2 short periods following MIS 17 and

MISS 9c with a contemporaneous superficial normal faulting. This episodic uplift of the Chilean coast in the

Pleistocene may result from subduction related processes, such as pulses of tectonic accretion at the base of the

forearc wedge. To our knowledge, this is the first time that 10Be exposure ages of a succession of wave-cut

platforms, has revealed non-steady long-term uplift rates on the Andean margin.

In Peru, the same type of studies is in progress. New ages on marine terraces are now available and provide

different uplift rates along the coast and higher uplift rates in the San Juan de Marcona area, above the Nazca

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 474-476

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Ridge. In Peru, the chosen areas belong to the transition zone where the megathrust dip angle changes from flat

(~15°) to normal (~30°) and extend further south towards Arica region, where the distance from the coast to the

trench increases. The coastal morphologies could be related to those major geodynamic changes that affect the

Peruvian coast from Paracas to Arica.

Figure 1: DEM SRTM 90 m mosaic of Central Andes. Location of study areas (orange rectangles): the southern coast of Peru, between San Juan de Marcona and Ilo, and the area between Tongoy and Los Vilos, in the Norte Chico of Chile. Ridge outlines correspond to the -3700 m bathymetric curve, obtained from GMT. FS: Flat subduction.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 474-476

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Figure 2: Type example of the morphology of wave-cut platforms from the Altos de Talinay area in Chile. 3D view of the wave-cut platform sequence along the Altos de Talinay and their respective age (Error 1 ). TI, TII, TIII, TIV, TV: Talinay I, Talinay II, Talinay III, Talinay IV, Talinay V. Color lines delimit the shoreline angle of each marine terrace. Images are extracted from Google Earth (http://earth.google.fr). Copyright: Terrametrics, DigitalGlobe, Europa Technologies image NASA, 2007.

Figure 3: Panoramic view of marine terraces sequences in the San Juan de Marcona area, along the Cerro El Huevo (A) and the Cerro Tres Hermanas (B) (Photographs M. Saillard).

References Saillard, M., Hall, S.R., Audin, L., Farber, D.L., Hérail, G., Martinod, J., Regard, V., Finkel, R.C. & Bondoux, F. Submitted.

Non-steady long-term uplift rates and Pleistocene marine terrace development along the Andean margin of Chile (31°S).

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477

3D structure of the Tres Cruces synclinorium from seismic data and serial balanced cross-sections, Eastern Cordillera, Argentina

Luiraima Salazar1, Jonas Kley

1, Eduardo Rossello

2, Ruben Monaldi

3, & Miriam Wiegand

1

1 Institut für Geowissenschaften, Friedrich-Schiller Universität Jena, Burgweg 11, D- 07749 Jena, Germany

2 Universidad de Buenos Aires-CONICET –Depto. de Ciencias Geológicas, 1428 Buenos Aires, Argentina

3 Universidad de Salta-CONICET-Depto. de Ciencias Geológicas, La Rioja 698, 4400 Salta, Argentina

KEYWORDS : fault-propagation fold, Cretaceous rift, tectonic inversion, shortening, Northwestern Argentina

Introduction

The geologic structures in

orogenic belts often show

complex kinematics on different

scales. These changes of local

movement direction have

frequently been interpreted as a

direct effect of plate kinematics.

However, local processes could

obviously play an important role

here. Only an active orogen with

a relatively simple geologic

history offers good possibilities

to examine these relationships. The Central Andes developed during the last 30 Ma in a constant tectonic setting

and are the product of crustal thickening and magmatism, in response to subduction of the oceanic Nazca plate

beneath continental South America, with a convergence vector of almost constant orientation. The structural

grain of the Central Andes reflects imbricated fold and thrust systems that in general agree well with continuous,

roughly E-directed shortening throughout the history of the active plate margin. However, there are geologic

structures indicating a perpendicular, approximately N-S shortening direction. The magnitude of N-S shortening

is relatively small, but the structures occur over a large area. The Tres Cruces synclinorium, located in Northwest

Argentina (Fig. 1) offers ideal conditions to study this type of perpendicular structures and their kinematic and

genetic relationship with the regional setting. The aim of this study is to understand and explain the tectonic

mechanisms that have created these complex and local structures inside a regional tectonic environment through

a three-dimensional model of this region. We have analysed field and subsurface data in an area of c. 180 Km2,

integrating seismic profiles, four exploratory wells with appearances of hydrocarbons and gas, satellite images

and previous geologic maps, in an attempt to clarify the spatial and temporal relationships between N-S- and E-

W-trending structures.

Figure 1. Simplified structural map (modified from Kley et al., 2005) of the Tres Cruces basin with location of Cerro Colorado. Right: Satellite image of the Cerro Colorado structure.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 477-480

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The Tres Cruces Synclinorium

Geologic Setting

The Tres Cruces synclinorium is an internally

deformed intermontane basin (Coutand et al. 2001)

located on the western border of the Eastern Cordillera

(Turner, 1972), adjacent to the Puna plateau. It exhibits

a typical thrust belt structural style, with subparallel

folds and imbricated thrust sheets trending

preferentially WNW-ESE and separated by wide

synclinal depressions.

The Tres Cruces synclinorium (Fig. 2) contains

sediments of Paleozoic to Cenozoic age. The oldest

rocks exposed in the area are slightly metamorphosed

marine shales and sandstones of Late Proterozoic to

Early Cambrian age. They are overlain in regional

angular unconformity by Late Cambrian to Ordovician

shallow marine sandstones and shales. The

disconformably overlying, Cretaceous-early Tertiary

Salta Group includes a lower unit of red conglomerates and sandstones with rare intercalations of basalts,

followed by light-coloured sandstones, limestones and varicoloured shales (Salfity, 1982). These strata represent

the synrift and postrift (sag) successions of a continental rift underlying much of northwest Argentina (Mon et al.

2005). They were deposited in alluvial fan, fluvial, eolian and lacustrine environments (Monaldi et al, 2008;

Coira et al., 1982). A thick synorogenic clastic wedge of continental foreland basin strata, with coarsening

upward sequence of mudstones, sandstones and conglomerates and attaining more than 2500 m thickness rests

disconformably on the postrift strata (Jordan and Alonso, 1987, Boll and Hernández, 1986).

Structures

The Tres Cruces synclinorium developed in the footwall of a major eastward verging thrust, which runs along

the western border of the intramontane basin and has uplifted the basin-bounding Sierra de Aguilar. However,

the structures and the amount of horizontal shortening change along strike (Coutand et al., 2001). Only the

eastern part of the basin is well exposed whereas the western part is largely covered by Neogene and Quaternary

sediments. In recent contributions the Tres Cruces synclinorium was interpreted as an inverted segment of the

Tres Cruces subbasin of the Salta rift (Boll and Hernández, 1986; Gangui, 1998; Coutand et al., 2001; Kley et

al., 2005; Monaldi et al., 2008). Cenozoic folding and thrusting superimposed on the earlier rift structures has

produced a complex structural style (Fig. 3), of which the Cerro Colorado structure, located in the center of the

Tres Cruces synclinorium is an excellent example (Boll y Hernández, 1986; Monaldi et al., 2008).

Figure 2. Stratigraphic column of the Tres Cruces area

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 477-480

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The Cerro Colorado anticline

The Cerro Colorado structure is an

approximately 15 km long and up to 6 km

wide asymmetrical anticline with a sigmoidal

main fold axis that trends NNE-SSW. The

anticline involves more than 5 km of

sedimentary strata including the Paleozoic

basement, Salta Group and Cenozoic foreland

basin sequences.

The integration of field data, seismic

interpretation, satellite-foto analysis,

geological profile construction and cross

section balancing with quantification of

shortening across and near the Cerro Colorado

indicate that it is a fault-propagation fold

above an eastward verging thrust (Vizcarra

Fault) with a listric and asymptotic geometry

merging into a detachment surface within

Ordovician strata (Fig. 4). In the central part it

is cut by a transverse fault which is expressed

as a topographic depression. Its origin is

probably associated with Cretaceous rifting

and it was reactivated as transverse reverse

fault during the Neogene orogeny. This fault

subdivides the anticline into two different

structural domains: in the northern part the

structure is a relatively simple east-directed

thrust sheet with a frontal anticline (Kley et al.

2005, Monaldi et al., 2008). The irregular

geometry of the Cerro Colorado anticline

reflects the close control on structure and topography by the amount of internal deformation, resulting from the

combined effects of two perpendicular shortening phases in N-S (10% shortening) and W-E (30-40% shortening)

direction.

Figure 3. Structural map of the central part of the Tres Cruces synclinorium.

Figure 4. Structural cross-section along the seismic line 90-47. It shows the internal structure of the Cerro Colorado anticline and other structures nearby.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 477-480

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The Sierra de Cajas anticline

The Sierra de Cajas anticline (Fig. 5) is located

south of Cerro Colorado and was uplifted on the

Cajas thrust, which has emplaced Ordovician

rocks over Tertiary sequences. The anticline

trends north, is doubly-plunging and asymmetric,

with the steep to overturned western flank

dissected by the steeply dipping, westward

verging thrust (Monaldi et al., 2008). The

horizontal shortening associated with these

structures is ~46 %.

Conclusions

The structures of the Cerro Colorado and Sierra de Cajas fault propagation anticlines demonstrate the effects

of Neogene contraction with and without added complexity by preexisting transverse faults. The transverse

contraction structures appear to be genetically associated with the reactivation of transverse faults that originated

during the rift period, since all documented examples of N-S shortening occur within the Salta rift. Overfolding

relationships indicate that N-S shortening mostly postdates the main E-W shortening phase. However, the

strongly curved fault trace of the N-directed thrust in the central part of Cerro Colorado may indicate folding of

the fault and a temporal overlap of both shortening events. The reason for along-strike contraction in an

essentially straight segment of the Andean thrust belt is still not clear. Kinematically, it could be explained as an

effect of strain partitioning; dynamically it would require the E-W trending faults to be extremely weak.

References Boll, A., and Hernández, R., 1986 - Interpretación estructural del área Tres Cruces. Boletin de Informaciones Petroleras,

Tercera Epoca III ( 7): 2-14. Coira, B., Davidson, J., Mopodozis, C, and Ramos, V., 1982 - Tectonic and magmatic evolution of the Andes of northern

Argentina and Chile. Earth Science Reviews, 18: 303-332. Coutand, I., Gautier, P., Cobbold, P., De Urreiztieta, M., Chauvin, A., Gapais, D., Rossello, E., Lopéz Gamundi, O., 2001-

Style and history of Andean deformation, Puna Plateau, northwestern Argentina. Tectonics, 20: 210-234. Gangui, A. H., 1998 - A combined structural interpretation based on seismic data and 3D gravity modelling in the northern

Puna/Eastern Cordillera, Argentina. Berl. Geowiss. Abh. Reihe C, 27: 176 pp. Jordan, T., and Alonso, R., 1987 - Cenozoic stratigraphy and basin tectonics of the Andes Montain, 20°-28° south latitude.

AAPG Bulletin, 71: 49-64. Kley, J., Rossello, E., Monaldi, C., and Habighorst, B., 2005 - Seismic and field evidence for selective inversion of

Cretaceous normal faults, Salta rift, Northwest Argentina. Tectonophysics, 399: 155-172. Mon, R., Monaldi, C y Salfity, J.A., 2005 - Curved structures and interference fold patterns associated with lateral ramps in

the Eastern Cordillera, Central Andes Argentina. Tectonophysics: 173-179. Monaldi, C.R., Salfity, J.A., y Kley, J., 2008 - Preserved extensional structures in an inverted Cretaceous rift basin,

northwestern Argentina: Outcrop examples and implications for fault reactivation. Tectonics 27: TC1011, 21p. Salfity, J.A., 1982 - Evolución paleogeográfica del grupo Salta (Cretácico-Eogénico), Argentina. Quinto Congreso

Latinoamericano de Geología, Buenos Aires, Argentina: 11-26. Suppe, J and Medwedeff, D., 1990 - Geometry and kinematics of fault-propagation folding, Eclogae Geologicae Helvetiae,

83: 409-453. Turner, J. 1972 - Cordillera Oriental. In: Leanza, A. F. (Ed), Geología regional de Argentina, vol. 1. Academia Nacional

de Ciencias: 117-142.

Figure 5. Structural cross-section along the seismic line 90-49. It shows the westward verging Cajas thrust and the associated fold with an overturned front limb.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 481-484

481

Analysis of microseismicity in the Precordilleran Fault System at 21°S in Northern Chile

Pablo Salazar1,2

, Jörn Kummerow1, Günter Asch

1,3, Dorothee Moser

1, & Peter Wigger

1

1 Freie Universität Berlin, FR Geophysik, Malteserstrasse 74, D12249 Berlin, Germany

(wigger@geophysik.fu-berlin.de- 2 Universidad Catolica del Norte, Av. Angamos 0610, Antofagasta, Chile

3 GeoForschungsZentrum Potsdam, Telegrafenberg, D14473 Potsdam, Germany

KEYWORDS : microseismicity, earthquake location, earthquake cluster, Precordilleran Fault system, Northern Chile

Introduction

The Precordilleran Fault system (PFS) is one of the most important features of the Andean forearc in Northern

Chile. The PFS has a length of more than 1000 km [Camus, 2003], with a predominant N-S orientation and

associated strike-slip movements. The fault system has influenced the emplacement and mineralization of a

number of the largest porphyry-copper-related intrusion around the world. These faults are composed of various

regional segments, each having undergone a distinct series of deformation events [Lindsay et al., 1995]. In the

northern part the regional branch of the PFS is known as West Fissure Fault Zone (WFFZ). The WFFZ is partly

well exposed at the surface. In spite of its geological, tectonic and economic importance very little is known

about its seismic activity and the continuation to greater depth. Recent studies of the area focus on the seismicity

associated with the subduction of the Nazca Plate beneath the South American Plate e.g., [Haberland and

Rietbrock, 2001], [Heit, 2005], the Tarapacá earthquake on 13 June, 2005 [Peyrat et al., 2006], the Aroma

earthquake on 24 July, 2001 [Legrand et al., 2007], intracontinental seismicity associated to Central Andes

Orocline [David, 2007] and tectonics evidences of the western Altiplano plateau e.g., [Janssen, 2002], [Victor,

2004], [Farías, 2005]. The present work focuses on the following aspects: Does the WFFZ have related

seismicity? Down to which depth crustal seismicity can be observed? Does seismicity give a hint to the width of

the WFFZ? How are the focal mechanisms and what do they tell us about the stress field on WFFZ? Our

analysis is placed on the area at ~21°S due to a series of geophysical observations at this latitude [ANCORP

Working Group, 1999, 2003; Wigger et al., 1994, Yuan et al., 2000].

Tectonic setting

The WFFZ has been active since the Late Eocene/Early Oligocene [Janssen, 2002], [Reutter et al., 1996]. Most

authors assume that an older dextral strike-slip motion caused by subduction-related magmatic arc tectonics of

the Incaic tectonic phase was followed by sinistral shear corresponding to a time of reduced convergence rate

[Reutter et al., 1996]. The youngest event is the reactivation of dextral slip under the same kinematic conditions

as described for the older phase. Deformation structures, which include folds, foliations, brittle faults and thrusts,

trend obliquely to the main fault trace [Carrasco et al., 1999]. The tectonic shows that west side the WFFZ is

subject to a compressional behaviour where thrust and reverse faults are observed. On the other hand, at the

eastern side a tensional behaviour with normal faults and basin development is described [Victor et al., 2004].

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 481-484

482

Experiment set-up

In order to monitor the seismicity at the WFFS, we installed a temporary seismic network in November 2005

(Fig. 1). The net has been recording continuously since that time and the operation will be maintained at least

throughout the year 2008. The seismic monitoring system covers an area of approx. 50 times 50 km at elevations

between 3000 m and 5000 m a.s.l. and consists of 14 surface stations with 3-component short period

seismometers (Mark L4-1Hz). The data recording is continuous at a sample rate of 200Hz. The seismic stations

are stand-alone and each seismic station is serviced every 3-6 months. So far, the observations until February

2007 have been processed and the first results are presented here.

Data Analysis and Results

Seismicity

For the recording period November 2005 to February 2007 about 700 local seismic events with magnitudes -

0.5 < ML<4.2 and focal depths between 2km and 50km have been located (earthquakes from the subduction

zone are not incorporated). The seismic event locations were determined using a grid-search algorithm in a

regional 2-D Vp model [Lueth 2000] and automatically corrected arrival times. We observe upper crustal

microseismicity (Fig. 1), which is only partly associated with the known branches of the WFFS. Focal depths

north of 21°S are smaller than 20 km whereas south of 21°S depths down to 50 km can be observed. In the W-E

section a distinct lower boundary of the seismicity is obvious, dipping to the West.

Swarm Events

Among the seismic activity, two seismic clusters were detected, one at 35-40 km depth in the SW of the

monitoring area (~110 events, between September and November 2006), and one in the central part at 9-10 km

depth beneath sea level (~120 events between March 31, and April 28, 2006, and ~40 events between January 8-

10, 2007).

The event cluster at 9-10 km depth exhibit characteristics of an earthquake swarm (Wigger et al. 2007). The

majority of events have very high cross correlation coefficients (greater than 0.8), indicating that the events

locate in a small volume (the maximum inter event distance is in the order of 1 km). To resolve the finer scale

structure, we applied a master event location approach. The results show that the earthquakes cluster on a ~600

m times ~500 m patch along a near-vertical west-east orientated fracture plane (perpendicular to the main

orientation of the WFFS).

Focal Mechanism stress tensor inversion

Focal mechanisms for ~200 events were determined based on the polarity of P wave arrivals and amplitude

ratios (SV/P, SH/P and SH/SV) [Snoke, 2003]. From the focal mechanisms we inverted the stress field in the

study area. We applied two approaches in the stress tensor calculations. First way we considered a single stress

tensor solution that explains all focal mechanisms in the area. In a second step we divided the study area

considering the tectonics frame and calculated the tensors respectively. For both calculations we have selected

events with more than 6 P polarity measurements. The results show that on the west side of WFFZ the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 481-484

483

predominant regime is transpressional whereas on the east side of WFFZ is transtensional. For the single

solution we have calculated a transtensional regime for the area. These results are consistent with WFFZ fault

motion derived from geological studies.

Fig 1: Local seismicity around the West Fissure Fault Zone (WFFZ) at about 21°S detected during the first 15 months of monitoring (11/2005-2/2007). Seismicity at the Wadati-Benioff zone is not shown. Short period 3-D station locations are plotted by red squares, WFFZ (black lines) after Reutter et al. (1994) Camus (2003), SERNAGEOMIN (2003) and Victor et al. (2004). The black W-E crooked line indicates the ANCORP profile with shot points (white stars). Earthquake locations (blue triangles) are plotted in map view (top left) and two depth sections (N-S: right and W-E: bottom), depth=0km corresponds to sea level.

References ANCORP Working Group (1999), Seismic reflection image revealing offset of Andean subduction-zone earthquake locations

into oceanic mantle, Nature, 397, 341 - 344. ANCORP Working Group (2003) Seismic imaging of an active continental margin and plateau in the central Andes (Andean

Continental Research Project 1996 (ANCORP '96)), J. Geophys. Res., 108 (B7), doi:10.1029/2002JB001771.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 481-484

484

Camus 2003. Geología de los sistemas porfídicos de los Andes de Chile. Servicio Nacional de Geología y Minería, 267pp. Santiago, Chile.

Carrasco, P., H. Wilke, H. Schneider (1999). Post-Eocene deformational events in the North segment of the Precordilleran Fault system, Copaquiri (21°S). Fourth ISAG, Göttingen (Germany), 04-06/10/99.

David, C. (2007). Comportamiento actual del ante-arco y del arco del codo de Arica en la orogénesis de los Andes centrales. PhD. Thesis, Universidad de Chile, Chile, 290 pp.

Farías, M., R. Charrier, D. Comte, J. Martinod, and G. Hérail (2005), Late Cenozoic deformation and uplift of the western flank of the Altiplano: Evidence from the depositional, tectonic, and geomorpholic evolution and shallow seismic activity (northern Chile at 19°30'S), Tectonics, 24, TC4001, doi: 10.1029/2004TC001667.

Haberland C, Rietbrock A (2001) Attenuation Tomography in the Western Central Andes: A detailed insight into the structure of a magmatic arc. JGR 106 (B6): 11151-11167.

Heit, B. (2005) Teleseismic Tomographic Images of the Central Andes at 21°S and 25.5°S: An inside look at the Altiplano and Puna plateaus, Scientific Technical Report, STR06/05GeoForschungsZentrum Potsdam, 139pp.

Janssen, C., A. Hoffmann-Rothe, S. Tauber, H. Wilke (2002). Internal strucutre of the Precordillera fault system (Chile) – insights from structural and geophysical observations. Journal of Structural Geology, 24, pp. 123 – 143.

Legrand, D, B. Delouis, L. Dorbath, C. David, J. Campos, L. Marquéz, J. Thompson, D. Comte (2007). Source parameters of the Mw=6.3 Aroma Crustal earthquake of July 24, 2001 (northern Chile), and its aftershock sequence. Journal of South American Earth Sciences, 24, pp. 58 – 68.

Lindsay, D. D., M. Zentilli, A. J. Rivera (1995). Evolution of an active ductile to brittle shear system controlling mineralization at Chuquicamata porphyry copper deposit, Northern Chile. International Geology Review, 37, pp. 945 – 958.

Lüth, S. (2000), Results of wide-angle investigations and crustal structure along a traverse across the central Andes at 21 degrees south. PhD thesis (in german), Institute of Geology, Geophysics and Geoinformatics, Free University of Berlin, Berliner Geowissenschaftliche Abhandlungen, Band 37, Reihe B.

Peyrat S., J. Campos, J. B. Chabalier, A. Perez, S. Bonvalot, M.-P. Bouin, D. Legrand, A. Nercessian, O. Charade, G. Patau, E. Clévédé, E. Kausel, P. Bernard and J.-P. Vilotte (2006). Tarapacá intermediate-depth earthquake (Mw 7.7, 2005, northern Chile): A slab-pull event with horizontal fault plane constrained from seismologic and geodetic observations. Geophysical Research Letters, 33, L22308.

Reutter, K.-J., E. Scheuber, G. Chong (1996). The precordilleran fault system of Chuquicamata, Northern Chile: evidence for reversals along arc-parallel strike-slip faults. Tectonophysics, 259, pp. 213 – 228.

Snoke, J. A. (2003). FOCMEC: FOCal MEChanism determinations, International Handbook of Earthquake and Engineering Seismology (W. H. K. Lee, H. Kanamori, P. C. Jennings, and C. Kisslinger, Eds.), Academic Press, San Diego, Chapter 85.12.

Victor, P., O. Oncken, and J. Glodny (2004), Uplift of the western Altiplano plateau: Evidence from the Precordillera between 20° and 21° (northern Chile), Tectonics, 23, TC4004, doi:10.1029/2003TC001519.

Wigger, P., Schmitz, M., Araneda, M., Asch, G., Baldzuhn, S., Giese, P., Heinsohn, W.-D., Martìnez, E., Ricaldi, E., Röwer, P., and Viramonte, J. ((1994), Variation of the crustal structure of the Southern Central Andes deduced from seismic refraction investigations. In: Tectonics of the Southern Central Andes, Reutter, Scheuber & Wigger (eds.), Springer Verlag, Berlin Heidelberg, p 23-48.

Wigger, P., Kummerow, J., Salazar, P., Asch, G., Moser, D. (2007), Microseismicity in the West Fissure fault system, Northern Chile. Geophysical Research Abstracts, Vol. 9, 07136.

Yuan, X., S. V. Sobolev, R. Kind, O. Oncken, and Andes Seismology Group (2000), New constraints on subduction and collision processes in the central Andes from P-to-S converted seismic phases, Nature, 408, 958 – 961.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 485-488

485

Relations between plutonism in the back-arc region in southern Patagonia and Chile Rise subduction: A geochronological review

Alejandro Sánchez1, Francisco Hervé

1, & Michel de Saint-Blanquat

2

1

Departamento de Geología, Universidad de Chile, casilla 13518 correo 21, Santiago, Chile

(alsanche@cec.uchile.cl) 2

CNRS-LMTG/Observatoire Midi-Pyrénées, Université de Toulouse, 14 av. Edouard-Belin, 31400 Toulouse,

France (michel@lmtg.obs-mip.fr)

KEYWORDS : plutonism, geochronology, Patagonia

Introduction

Southern Patagonia is constituted by 3 main tectono-stratigraphic units, from west to east: the Permo-Triassic

deformed basement, the Late Jurassic – Early Cretaceous extensional basins, and the Late Cretaceous to

Cenozoic compressional Magallanes foreland basin (MFB). There are also two intrusive and two extrusive

magmatic belts: the Southern Patagonian Batholith (SPB) (Mesozoic-Cenozoic), which intrudes the basement,

and the back-arc region plutons (Neogene) which intrude the MFB and the extensional basins, specially in the

northern area. The extrusive units are the Miocene-Pleistocene basaltic plateau lavas, mainly in Argentina, and

lastly the active volcanic arc located at the western

margin of South America.

The back-arc region plutons (figure 1), are mainly

miocene isolated granites to diorites bodies. Michel

(1983) had been the first in considerate it as a north-

south magmatic lineament in the back-arc region, and

furthermore he linked this magmatic lineament with

the Chile Rise (CR) subduction, which started 15-14

Ma at 55º lat.S and since that time the triple joint has

migrated to the north until their actual position

(~46º30' lat. S) (Cande and Leslie, 1986).

Nevertheless, there is not a good correlation between

Triple Joint migration and magmatism occurrence.

The latter is widespread in Neogene times (figure 1),

including several 15-25 Ma plutons in the SPB

(Hervé et al., 2007); the basaltic plateau lavas, which

are Pliocene to mid-Miocene (~14 Ma) (Guivel et al.,

2006); also there is several adakite type intrusives

miocene in age in the back-arc region (Ramos et al.,

2004).

The aim of this contribution, is to help to solve the following questions: Is this north to south miocene

lineament a real lineament? And: are this igneous bodies related to CR subduction?

Figure 1: Magmatic map of southern Patagonia. After Michel,

1983.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 485-488

486

In this way we present a geochronologic compilation of cenozoic intrusive bodies in the back-arc region of

Patagonia together with preliminary new age data of plutons that have been include in this “lineament” (Sánchez

et al., in prep.). The main goal of this contribution is that age data of the back-arc plutons reveals that most of

them are older than the CR subduction and a direct relation between plutonism and slab windows is still unclear.

Geochronology

A compilation of available geochronologic data of intrusives rocks in southern Patagonia is presented in Table

1 and illustrated in Fig. 2. Its include the different type of rocks mentioned above, together with the age and

location of Chile rise segments collision (accord to Cande & Leslie, 1986). This compilation is only of

radiometric dating methods. Even K/Ar method may reflect younger ages than plutons crystallization ages, is

include because many plutons are only dated by this methodology. Even the database includes diverse

radiometric methods, some general observations can be made:

In all the region, the age data for the back-arc plutons

is mainly concentrated between 9 and 18 Ma

(exceptions are Las Nieves granite and San Lorenzo

granite). In most cases these age are older than Chile

Rise collision age. Only few temporal coincidence

between magmatism and Chile rise subduction exist.

The most remarkable are Cerro Pampa adakite, San

Lorenzo granite and Torres del Paine granite.

In the arc region (SPB), there are plutons which range

in age between 15 and 25 Ma (Hervé et al., 2007).

There is only one Pliocene pluton of 4 Ma and this is

the only dated pluton of the SPB in this area, younger

than CR collision.

In Fig. 2, also can be noted that the original plutons

in the lineament (enclosed by circles), do not show a

pattern of been younger northward as the CR collision

age.

Discussion

The geochronology used should be more representative of crystallization ages for allow good comparison of

emplacement times of the plutons. So U/Pb zircon dating is being carried out by the authors to avoid bad

interpretations, as can occur here comparing pluton ages obtained by differents radiometric methods.

Nevertheless, it seems to be clear that most plutonism in both, arc and back-arc regions, is previous to CR

collision. This underestimate astenospheric windows source for these plutons. And it can be related to previous

buoyancy of the subducted slab related magmatism as is suggest by Espinoza (2007).

Among the back-arc plutons, there are several complex which includes Miocene, Oligocene and/or Cretaceous

plutons (e.g. San Lorenzo and Puesto Nuevo) the coincidence in the back-arc region location of these, may

Figure 2. Southern Patagonian back-arc plutons ages

diagram. Crosses represents ages of plutons of SPB (Hervé

et al., 2007). horizontal bars represent location and age of

collision of CR segments (from Cande & Leslie).LL: Las

Llaves, LN: Las Nieves, CN: Cerro Negro del Ghío, CI:

Cerro Indio, SL: San Lorenzo, CP: Cerro Pampa, PN: Puesto

Nuevo, FR: Fitz-Roy, Ch: Chalten, CM: Cerro Moyano, TP:

Torres del Paine, CD: Cerro Donoso, CB: Cerro Balmaceda

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 485-488

487

reflect ancient discontinuities in the patagonian continental crust This mechanism should allow magmas to reach

upper levels of the crust in high magmatic periods, as may be miocene times.

Table 1. Localities and dating methods of the plutons ages plotted in Fig.2. 1: Pankhurst et al. (1999); 2: Petford & Turner

(1996); 3: Suarez & de la Cruz (2001); 4: Morata et al. (2002); 5: Ramos (2002); 6: Welkner (1999); 7: Welkner (2000); 8:

Fanning, pers.com.; 9: Pino (1976); 10 Ramos & Palma (1981); 11: Pankhurst, com.pers to Giacosa Franchi; 12: Ramos et al. (1991); 13: Ramos et al. (2004); 14: Motoki et al. (2003); 15: Nullo et al. (1978); 16: Linares & Gonzales (1990); 17:

Halpern, 1973; 18: Sánchez et al. (2006); 19: Altenberger et al. (2003); 20: Sánchez et al., in prep.; 21: in Skarmeta &

Castelli (1997)

Pluton Lat.S Lon.W Dating method Age (Ma) Error (Ma) Material

Paso de las llaves1

46°40' 72°15' Rb/Sr 10.3 0.4 WR-KFd Bt (isochron)

Paso de las llaves2

46°40' 72°15' Ar-Ar isochron 9.6 0.5 Biotite

Paso de las llaves2

46°40' 72°15' Ar-Ar 9.6 0.4 Biotite

Paso de las llaves3

46°40' 72°15' K/Ar 10 1.1 Biotite

Aviles3

46°45' 72°15' K/Ar 9.6 0.6 Biotite

Rio de las Nieves4

46°41' 72°06' K/Ar 3.2 0.4 Biotite

Cerro Indio5

47º6 71º53' K/Ar 13.2 0.9 WR

Cerro Negro del Ghío5

47º7' 71º52' K/Ar 18.1 1.2 WR

Cerro Negro del Ghío5

47º7' 71º52' K/Ar 15.8 0.7 WR

Cerro Negro del Ghío5

47º7' 71º52' K/Ar 15.8 0.6 Hornblende

Co Sn Lorenzo3

47°35’ 72°20’ K/Ar 6.6 0.5 Biotite

Co Sn Lorenzo6

47°35’ 72°20’ K/Ar 6.4 0.4 Biotite

Co Sn Lorenzo7

47°35’ 72°20’ WMPA Ar/Ar 5.76 0.18 K-feld

Co Sn Lorenzo7

47°35’ 72°20’ WMPA Ar/Ar 6.2 0.12 Biotite

Co Sn Lorenzo8

47°35’ 72°20’ SHRIMP U/Pb 6.44 0.28 circon

Co Sn Lorenzo9

47°35’ 72°20’ K/Ar 8.8 6.1 Not published

Co Sn Lorenzo10

47°35’ 72°20’ K/Ar 8 1 Not published

Co Sn Lorenzo11

~47º40' ~72º15 Rb/Sr 9.2 1.6 Not published

Cerro Pampa12

47º54' 71º20' K/Ar 12.1 0.7 Not published

Cerro Pampa12

47º54' 71º20' K/Ar 12 0.7 Not published

Cerro Pampa13

47º54' 71º20' WMPA Ar/Ar 11.39 0.61 Hornblende

Cerro Pampa13

47º54' 71º20' WMPA Ar/Ar 12.87 0.24 WR

Puesto Nuevo13

48º56' 72º12,5' WMPA Ar/Ar 13.12 0.55 Hb

Puesto Nuevo13

48º56' 72º12,5' WMPA Ar/Ar 13.29 3.97 WR

Puesto Nuevo14

48º56' 72º12,5' U/Pb 14.1 3.6

Fitz Roy15

49°15' 73° K/Ar 18 3 WR

Chalten13

49°25,5' 72º59,5' WMPA Ar/Ar 14.5 0.29 Amphibole

Cerro Moyano16

50º27,2 72º23,7' K/Ar 16 1 WR

Paine17

51° 73° Rb/Sr 12 2 Bt

Paine17

51° 73° K/Ar 13 1 Bt

Paine18

51° 73° SHRIMP U/Pb 12.65 0.13 circon

Paine (external gabro)19

51° 73° K/Ar 29.4 0.8 Biotite

Co Donoso20

51º13,3 73º9.5' SHRIMP U/Pb ~26 Circon

Co Balmaceda21

51°25' 73°11' K/Ar ? 28 Not published

Co Balmaceda20

51°25' 73°11' SHRIMP U/Pb ~15 Circon

References Altenberger, U., Oberhaensli, R., Putlitz, B. & Wemmer, K. 2003. Tectonic controls of the Cenozoic magmatism at the

Torres del Paine, southern Andes (Chile, 51°10'S). Revista Geologica de Chile 30: 65-81 Cande, S.C., & Leslie, R.B., 1986. Late Cenozoic tectonics of the Southern Chile Trench. J. Geophys. Res., 3: 471-496. Espinoza, F. 2007. Evolución Magmática de la región de trasarco de la patagonia central (47ºS) durante el Mio-Plioceno.

Tesis U.Chile (inedito).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 485-488

488

Guivel, C., Morata, D., Pelleter, E., Espinoza, F., Maury, R., Lagabrielle, Y., Polve, M., Bellon, H., Cotten, J., Benoit, M., Suárez, M. & de la Cruz, R. 2006. Miocene to Late Quaternary Patagonian basalts (46–478S): Geochronometric and geochemical evidence for slab tearing due to active spreading ridge subduction. Journal of Volcanology and Geothermal Research 149: 346-370

Halpern, M., 1973. Regional Geochronology of Chile south of 50° S latitude. Geol. Soc. Am Bull 84: 2407-2422. Hervé, F., Pankhurst, R.J., Fanning, C.M., Calderón, M. & Yaxley, G.M. 2007. The South Patagonian batholith: 150 my of

granite magmatism on a plate margin. Lithos 97: 373-394 Linares, E. & González, R.R. 1990. Catálogo de edades radimétricas de la República Argentina 1957-1987. Asociación

Geológica Argentina, Publicaciones Especiales Serie B, Didáctica y Complementaria 19: 1-628, Buenos Aires. Michael, P.J. 1983. Emplacement and differentiation of Miocene plutons in the foothills of the southernmost Andes. Ph.D.

Thesis (Unpublished), Columbia University, 367 p. Morata et al. 2002. Early Pliocene magmatism and high exhumation rates in the patagonian cordillera (46°40'S): K-Ar, and

fission track data. V ISAG Motoki, A., Orihashi, Y., Cario, F.D., Hirata, D., Haller, M.J., Ramos, V., Kawano, H., Watanabe, Y., Schilling, M., Iwano,

H. & Anma, R. 2003. U-Pb dating for single grain zircon using Laser Ablation ICP Mass Spectrometer and fission track ages for zircon grains of back-arc adakitic bodies, Argentine Patagonia. IV International Symposium of Isotope Geology Abstracts: 219-220, Salvador.

Nullo, F.E., Proserpio, C. & Ramos, V.A. 1978. Estratigrafía y tectónica de la vertiente este del hielo continental patagónico, Argentina-Chile, VII Congreso Geológico Argentino Actas I: 455-470.

Pankhurst, R., Weaver, S., Hervé, F. & Larrondo, P. 1999. Mesozoic–Cenozoic evolution of the North Patagonian Batholith in Aysén, southern Chile. Journal of the Geological Society 156: 673-694

Petford, N. & Turner, P. 1996. Reconnaissance 40Ar-39Ar age age and palaeomagnetic study of igneous rocks around Coyhaique, S. Chile (45º30'-47ºS). III ISAG 17: 625-628

Pino, M. 1976. Reconocimiento geolo gico de los departamentos de Cochrane y Baker, Unde cima Regio n, Ayse n. Tesis Universidad de Chile, 155 p, (inedito)

Ramos, V., Kay, S. & Singer, B. 2004. Las adakitas de la Cordillera Patagónica: Nuevas evidencias geoquímicas y geocronológicas. RAGA 59: 693-706

Ramos,V. & Palma, 1981. El batolito granitico del monte san Lorenzo, cordillera Patagonica, provincia de Santa Cruz. VIII Congreso Geologico Argentino Actas 3: 257-280

Ramos, V.A., Kay, S.M. y Márquez, M. 1991. La Dacita Cerro Pampa (Mioceno - provincia de Santa Cruz): evidencias de la colision de una dorsal oceánica. VI Congreso Geológico Chileno Actas I: 747-751.

Ramos, V.A. 2002. El magmatismo neógeno de la Cordillera Patagónica. In M.J. Haller (ed.) Geología y recursos naturales de Santa Cruz. XV Congreso Geológico Argentino (El Calafate) Relatorio I(13): 187-200, Buenos Aires.

Sánchez, A., de Saint Banquat, M., Hervé, F., Pankhurst, R.J. & Fanning, C.M. 2006. A SHRIMP U-Pb zircon late Miocene crystallization age for the Torres del Paine pluton, Chile. V SSAGI : 196-199

Sánchez, A., de Saint Banquat, M., Hervé, F., & Fanning, C.M. In prep. SHRIMP U-Pb geochronology and geochemical signatures of cenozoic back-arc region plutons in southern Patagonia:insights for genesis and emplacement of magmas.

Skarmeta, J. & Castelli, J.C. 1997. Intrusión sintectónica del Granito de Las Torres del Paine, Andes Patagónicos de Chile. Revista Geológica de Chile 24:.55-74.

Suarez, M. & de la Cruz, R. 2001. Jurassic to Miocene K-Ar dates from eastern central Patagonian Cordillera plutons, Chile (45°-48° S). Geol Mag, 138: 53-66.

Welkner, D., 1999. Geología del área del cerro San Lorenzo: Cordillera Patagónica Oriental, XI Región de Aysén, Chile (47°25'-47°50'S). Tésis Universidad de Chile, 151 p.( Inédito)

Welkner, D. 2000. Geocronología de los plutones del área del cerro San Lorenzo, XI Región Aysén. IX Congreso Geológico Chileno Actas v.2: 269-273.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 489-492

489

Gravity field analysis and preliminary 3D density modeling of the lithosphere at the Caribbean-South American plate boundary

Javier Sánchez1, Hans-Jürgen Götze

1, Michael Schmitz

2, & Carlos Izarra

3

1

Christian-Albrechts-Universität zu Kiel, Kiel 24118, Germany (jsanchez@geophysik.uni-kiel.de) 2 FUNVISIS, Caracas, Venezuela (mschmitz@funvisis.gob.ve)

3 Universidad Simón Bolívar, Dpto. Ciencias de la Tierra, Caracas, Venezuela (cizarra@usb.ve)

KEYWORDS : gravity, 3D modeling, Euler deconvolution, curvature, Caribbean, subduction zone

Introduction

The present Caribbean-South American configuration of subduction and geodynamics results from a

transpressive evolution that started in the Tertiary and continued in the Quaternary (Pindell, 1994; Meschede and

Frisch, 1998). In fact, western Venezuela shows a very complex geodynamic setting where the South America,

Nazca and Caribbean plates and several smaller crustal blocks are interacting (Audemard, 1993; 1998). The

Caribbean plate moves eastward with about 2 cm/yr (Mann, et al., 1990, Pérez et al., 2001) relatively to the

South American continent in compressive, extensional and strike-slip tectonic regimes. Those regimes are

associated with significant E-W-trending topographic reliefs (Mérida Andes, the Coast and Interior ranges) and

are still active. Atlantic/South America oceanic lithosphere subducts obliquely under the Caribbean plate with

about 4 cm/yr (DeMets et al., 1994) in the last 5 Myr (Audemard, 2000).

In recent times different projects have been conducted to collect geophysical, geological and geodetic data

which help us to understand this plate boundary. Some projects are e.g. BOLIVAR (Broadband Ocean-Land

Investigations of Venezuela and the Antilles arc Region) and GEODINOS (Recent Geodynamics of the Northern

Limit of the South American Plate) (Levander et al., 2006).

Gravity database and data analysis

In total more than 100,000 stations have been compiled and homogenized. The data stem from archives of the

Simon Bolivar University (Izarra, 2005), the National Geophysical Data Center (NGDC) and the Venezuelan

Foundation for Seismological Research (FUNVISIS) and comprises about 80,000 observations onshore and

more than 20,000 stations offshore. Taking into account all different sources of errors in the databases the

accuracy of the computed Bouguer anomaly map is in the range of ± 5–10 10-5m/s2. The calculated anomaly

map consists of Bouguer anomalies onshore (correction density of 2.67 t/m3) and Free Air anomaly offshore

(Figure 1). The magnitude of the anomaly map ranges from –250 to 250 x 10 m/s2 (mGal) with a prominent

anomaly low observed in eastern Venezuela and two gravity lows in both flanks of Merida Andes. Offshore

gravity anomalies are observed along the east-west trending subduction zone in northern Venezuela, and in the

oblique subduction in eastern Venezuela. The anomaly map does not really show a significant trend or pattern

which is related to diverse sources along the deformation zone on the Caribbean-South-American plate

boundary.

Curvature techniques were applied to analyze directional pattern of the observed anomaly and highlight the

main feature within study area. Various curvature attributes have been calculated and compared to detect faults

and other features in the potential field data that were not visible in the Bouguer gravity anomaly field directly.

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In general, curvature attributes shows main features observable in the analyzed data. Best results were taken

from dip curvature, minimum curvature and most positive curvature. In dip curvature maps (e.g. Figure 2) thrust

systems in central and eastern Venezuela were highlighted by minimum values. High values characterize

lineaments which trends from north Trinidad until Eastern Venezuela basin.

Toward this end the most important implications are related with the detached oceanic slab beneath continental

crust (VanDecar et al , 2003; Russo et al, 1996); the lineaments analyzed in the gravity map seems to be related

more with shallow structures and the attached material in the Easter Cordillera and not from the slab itself. This

behavior it is also observed in depth of anomalies sources from Euler deconvolution method which shows a

gather of solutions situated within the upper crust and mantle below eastern basin which can be a result of a

decrease of density contrast (Figure 3) associated with sinkinning of the basin. Many other source solutions seen

to be relate with Moho. In figure 4 many depth solutions located beneath Maracaibo block shows a clear

correlation with de Caribbean slab underneath South America.

Figure 1. Gravity anomaly map of Northern Venezuela and the Southern Caribbean Sea. EVB: Eastern Venezuela Basin, MA: Merida Andes, FB: Falcon Basin, BAB: Barinas Apure Basin, MB: Maracaibo Block, CCTB: Coastal Cordillera Thrust Belt, SCTB: Serranía del Interior Thrust Belt, GS: Guayana Shield.

3D density model

The presented model was calculated by the inhouse IGMAS software (Götze and Lahmeyer, 1988 Schmidt and

Götze, 2002) and represent an interpretation of many possibilities which in constraining by published and

unpublished studies. Here two parallel sections are presented to illustrate the main characteristics of the 3D

model. This preliminary 3D model is constrained by wide-angle seismic refraction sections (Schmitz et al.,

2005), Moho depth estimations from receiver functions (Niu et al., 2007), and hypocenters. Additionally we

used depths estimations from Euler deconvolution, a Venig-Meinesz isostasy map, and mapped surface geology

and other geodynamic information for model constrains. Model densities have been taken from previous models

of the region and were calculated by the method of Sobolev and Babeyko (1994). Their vp-density conversion

takes into count in situ temperature and pressure conditions of the lithosphere in the investigated region.

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The density model consists of ~30 bodies which represent the main geological units,: oceanic crust and

lithospheric mantle of the subducting slab, continental crust, the asthenospheric mantle and the oceanic water

cover. Up to now it is built by 20 vertical planes. In Figures 3 and 4 we present the planes located at the same

location of the velocity models of Schmitz et al. (2005) at 70 W and 65W respectively. The cross section along

the 70°W meridian (Fig. 3) shows the sinking slab of the Caribbean plate under the Maracaibo block and

significant crustal thinning in the area of the Falcón basin where Moho depths are reduced to approximately

10 km (Bezada, 2007). The cross section located at 65°W (Fig. 4) shows the sedimentary thickness of the

Eastern Venezuela Basin where the crustal thickness reaches up to 50 km (Schmitz et al, 2005); it varies

significantly between the area of the Guayana shield and the Venezuela Basin.

Figure 2. Map of dip curvature superimposed by the coastline (light grey lines) and fault Zones (black teeth lines). Box A shows lineaments in the South Caribbean deformation belt; box B: contains of lineaments of the Serranía del Interior Thrust Belt; Box C the lineaments of the Merida Andes are visualized.

Figure 3. The cross-section of the density model along 70° W. It represents the segments which were identified in the analysis and modeling of the gravity field and in geologic mapping. The upper part shows three lines : the measured gravity curve in red, the gravity of the 3D density model (black dotted line) and the modeled gravity of a 2D approach (black dashed line). The lower part represents the modeled structures of the lithospheric densities. Also included are the locations of the Euler source points (black dots) and local seismicity (circles).

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Figure 4. For comparison this cross-section of the density model along 64° W is shown. It represents the segments which were identified in the analysis and modeling of the gravity field and in geologic mapping. The upper part shows three lines: the measured gravity curve in red, the gravity of the 3D density model (black dotted line) and the modeled gravity of a 2D approach (black dashed line). The lower part represents the modeled structures of the lithospheric densities. Also included are the locations of the Euler source points (black dots) and local seismicity (circles). Blue lines indicate the findings of the wide-angle seismic experiment.

References Audemard, F.A., 1993. Néotectonique, sismotectonique et aléa sismique du nord-ouest du Vénézuela (système de failles

d’Oca-Ancón). Thesis doctoral, University of Montpellier II. 369 p. Audemard, F.A., Machette, M., Cox, J., Dart, R., Haller, K., 2000, Map and Database of Quaternary Faults and Folds in

Venezuela and its Offshore Regions: USGS Open-File report 00-0018. Audemard, F.A., 1998. Evolution géodynamique de la facade nord Sud-américaine: nouveaux apports de l'histoire

geéologique du Bassin de Falcón, Vénézuéla. Proceedings XIV Caribbean Geological Conference Trinidad: 327-340. Bezada et al. 2007. Crustal structure in the Falc´on Basin area, northwestern Venezuela, from seismic and gravimetric

evidence. Journal of Geodinamics. Accepted 16 November 2007 DeMets, C., Gordon, R.G., Argus, D.F. and Stein, S., 1994, Effects of recent revisions to the geomagnetic time scale on

estimates of current plate motions. Geophysical Research Letters 21: 2191-2194. Götze, H.-J. and Lahmeyer, B., 1988. Application of three-dimensional interactive modeling in gravity and magnetics.

Geophysics Vol. 53, No. 8: 1096-1108. Levander, A., M. Schmitz, H.G. Avé Lallemant, C.A. Zelt, D.S. Sawyer, M.B. Magnani, P. Mann, G. Christeson, J. Wright,

D. Pavlis y J. Pindell, 2006. Evolution of the Southern Caribbean Plate Boundary. EOS: 87, nr. 9: 97-100. Meschede, M. and Frisch, W., 1998, A plate-tectonic model for the Mesozoic and Early Cenozoic history of the Caribbean

plate, Tectonophysics 296: 269-291. Mann, P., Schubert, C. and Burke, K., 1990, Review of Caribbean neotectonics, in Dengo G., Case J.E. (eds.): The Caribbean

Region. Geological Society of America, Boulder, Colorado, v. H: 307-338. Niu, F., Baldwin,T., Pavlis, G., Vernon, F.,Rendón, H. y Levander, A., 2007. Receiver function study of the crustal structure

in the south eastern Caribbean plate boundary and Venezuela. Earth and Planetary Science Letters, submitted. Perez, O.J., Bilham, R., Bendick, R., Velandia, J.R., Hernandez, N., Moncayo, C., Hoyer, M. and Kozuch, M., 2001,

Velocity field across the Southern Caribbean plate boundary and estimates of Caribbean-South-American plate motion using GPS geodesy 1994-2000, GRL 28: 2987-2990.

Pindell, J., 1994, Evolution of the Gulf of Mexico and the Caribbean, in Donovan, S.k., Jackson, T.A., (eds): Caribbean Geology, and Introduction: 13-39.

Russo, R.M., Silver, P.G., Franke, M., Ambeh, W.B., James, D.E., 1996. Shear-wave splitting in northeast Venezuela, Trinidad, and the eastern Caribbean. Phys. Earth Planet. Inter. 95, 251– 275.

Schmitz, M., Martins, A., Izarra, C., Jácome, M.I., Sánchez, J. y Rocabado, V., 2005. The major features of the crustal structure in north-eastern Venezuela from deep wide-angle seismic observations and gravity modelling. Tectonophysics, doi:10.1016/j.tecto.2004.12.018.

Schmidt, S., and Götze, H.-J., 1998: Interactive visualization and modification of 3D-models using GIS-functions. Physics and Chemistry of the Earth 23-3: 289-295.

VanDecar, J.C., Russo, R.M., James, D.E., Ambeh, W.B., Franke, M., 2003. Aseismic continuation of the Lesser Antilles slab beneath northeastern Venezuela. J. Geophys. Res. 108, doi:10.1029/2001JB000884.

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Upper lithospheric structure of the subduction zone in south-central Chile: Comparison for differently aged incoming plate

Martin Scherwath, Eduardo Contreras-Reyes, Ernst R. Flueh, & Ingo Grevemeyer

Leibniz Institute of Marine Sciences, IFM-GEOMAR, Wischhofstr. 1-3, 24116 Kiel, Germany

(mscherwath@ifm-geomar.de)

KEYWORDS : subduction zone processes, forearc deformation, lithospheric structures, ray-tracing

Crustal and upper mantle structures of the subduction zone in south central Chile, between 42°S and 46°S, are

determined from seismic wide-angle reflection and refraction data as part of the TIPTEQ (from The Incoming

Plate to mega-Thrust EarthQuake processes) project (Scherwath et al., 2006). Three profiles along differently

aged segments of the subducting Nazca plate are analysed here in order to study dependencies of subduction

zone structures on age, i.e. thermal state, of the incoming plate. The age of the oceanic crust at the trench ranges

from 3 Ma on the southernmost profile, immediately north of the Chile triple junction, to 6.5 Ma old about

100 km to the north, and to 14.5 Ma old another 200 km further north, off the Island of Chiloe (Figure 1).

Figure 1. Basemap of TIPTEQ working area: (a) Relative to South America. (b) All five TIPTEQ profiles and the 2001 SPOC profile (Krawcyzk et al., 2003). (c) Central TIPTEQ profiles; high resolution bathymetry from R/V Sonne cruise 181 (Flueh and Grevemeyer, 2005) and from Bourgois et al. (2000).

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Figure 2. Velocity models of TIPTEQ Corridors 2-4, obtained by a combination of ray-tracing (Zelt and Smith, 1992) and tomorgraphy (Korenaga et al., 2000).

Remarkable similarities appear on the structures of both the incoming as well as the overriding plate (Figure

2). The oceanic Nazca plate is around 5 km thick, with a slightly increasing thickness northward, possibly due to

temperature changes along the actively spreading Chile Ridge. The trench basin is about 2 km thick except in the

south where the Chile Ridge is close to the deformation front and only a small, 800 m thick trench could

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develop. Roughly half the trench fill subducts and half of it accretes above the decollement (Bangs and Cande,

1997). Similarities in the overriding plate are the width of the active accretionary prism, 35-50 km, and a strong

lateral crustal velocity gradient zone about 75-80 km off the deformation front, where upper crustal velocities

decrease seaward from over 5.0-5.4 km/s to around 4.5 km/s within about 10 km, which possibly represents a

paleo-backstop. This zone is also accompanied by strong intraplate seismicity.

Differences in the subduction zone structures exist in the outer rise region, where the northern profile exhibits a

clear bulge of uplifted oceanic lithosphere prior to being subducted whereas the younger structures have a less

developed outer rise. This plate bending is accompanied by strongly reduced rock velocities on the northern

profile due to fracturing and possible hydration of the crust and upper mantle (Contreras-Reyes et al., 2007,

2008). The southern profiles do not exhibit such a strong alteration of the lithosphere, although this effect may

be counteracted by plate cooling effects, which are reflected in increasing rock velocities away from the

spreading centre. Overall there appears little influence of incoming plate age on the subduction zone structure

which may explain why the Mw=9.5 great Chile earthquake from 1960 ruptured through all these differing age

segments (Cifuentes et al., 1989).

References Bangs, N.L., and Cande, S.C. 1997. Episodic development of a convergent margin inferred from structures and processes

along the southern Chile margin. Tectonics, 16(3): 489-503. Bourgois, J., Guivel, C., Lagabrielle, Y., Calmus, T., Boulegue, J., and Daux, V. 2000. Glacial-interglacial trench supply

variation, spreading-ridge subduction, and feedback controls on the Andean margin development at the Chile triple junction area (45-48°S). J. Geophys. Res., 105(B4): 8355-8386.

Cifuentes, I.L. 1989. The 1960 Chilean Earthquake. J. Geophys. Res., 94(B1): 665-680. Contreras-Reyes, E., Grevemeyer, I., Flueh, E.R., Scherwath, M. and Heesemann, M. 2007. Alteration of the subducting

oceanic lithosphere at the southern central Chile trench-outer rise. Geochem. Geophys. Geosyst. 8(7), Q07003, doi:10.1029/2007GC001632.

Contreras-Reyes, E., Grevemeyer, I., Flueh, E.R. and Scherwath, M. 2008. Seismic structure of the incoming Nazca plate at the Southern Central Chile outer-rise. Geophys. J. Int., 173(1): 142-156, doi:10.1111/j.1365-246X.2008.03716.

Flueh, E.R. and Grevemeyer, I. 2005. FS Sonne Fahrtbericht / Cruise Report SO181 TIPTEQ. IFM-GEOMAR Report 2, pp. 539.

Korenaga, J., Holbrook, W.S., Kent, G.M., Kelemen, P.B., Detrick, R.S., Larsen, H.-C., Hopper, J.R., Dahl-Jensen, T. 2000. Crustal structure of the southeast Greenland margin from joint refraction and reflection seismic tomography. J. Geophys. Res., 105, 21,591-21,614.

Krawcyzk, C., and the SPOC Team. 2003. Amphibious seismic survey images plate interface at 1960 Chile Earthquake. Eos Trans. AGU, 84(32).

Scherwath, M., Flueh, E.R., Grevemeyer, I., Tilmann, F., Contreras-Reyes, E. and Weinrebe, W. 2006. Investigating Subduction Zone Processes in Chile. Eos Trans. AGU, 87(27): 265-269.

Zelt, C.A., and Smith, R.B. 1992. Seismic traveltime inversion for 2-D crustal velocity structure. Geophys. J. Int., 108: 16-34.

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Are the Falkland Plateau and the Deseado Massif part of the same Mesoproterozoic lithospheric block?

Manuel Schilling1,2

& Andrés Tassara3,*

1 Departamento de Geología Regional, Servicio Nacional de Geología y Minería, Av. Santa María 0104,

Providencia, Santiago, Chile (mschilling@sernageomin.cll) 2 Departamento de Geología, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile. Plaza Ercilla

803, Santiago, Chile. (mschilli@cec.uchile.cl) 3 Departamento de Geofísica, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile. Blanco

Encalada 2002, Santiago, Chile (andres@dgf.uchile.cl)

* now at Departamento de Ciencias de la Tierra, Facultad de Química, Universidad de Concepción, Casilla 160-

C, Concepción, Chile

KEYWORDS : southern South America, Falkland/Malvinas Islands, Deseado Massif, lithospheric block

Introduction

The reconstruction of Rodinia and Gondwana supercontinents has been the target of several studies during the

last decades. The Falkland Islands represent a key element for this purpose (e.g. Marshall, 1994). This is in part

because its ancient geological history started in Mesoproterozoic times, as is evidenced by rocks from the Cape

Meredith Complex with ages from 1140 to 1000 Ma, that are related to the Namaqua-Natal Mesoproterozoic

orogenic belt (e.g. Thomas et al., 1994). During the Neoproterozoic – early Paleozoic occurred the collision

between East Africa and East Antarctica that produced the final amalgamation of Gondwana. This phenomenon

generates the lateral escape of several microplates, including the Falkland, the Ellsworth-Haag and the Filcher

blocks, in the southern part of the orogen (Jacobs and Thomas, 2004). Approximately during mid Carboniferous

to early Permian is recognized a Gondwanide Fold Belt at Sierra de la Ventana (Argentina), Cape Fold Belt

(Africa), Ellsworth Mountains (Antarctica) and the Falkland Islands (e.g. Pankhurst et al., 2006). Finally, during

the break-up of Gondwana from Jurassic to the present, the continental plate reconstructions are characterized by

major changes, including the rotation and translation of fragments such as the Falkland Plateau (Taylor and

Shaw, 1989) and southern Patagonia (Vizán et al., 2005).

The geological evolution of southern South America is characterized by the accretion of several continental

terranes to the southwestern proto-margin of Gondwana since late Proterozoic times (Ramos, 1984, 1988).

Recently, Pankhurst et al. (2006) presented a review of the post-Cambrian igneous, structural and metamorphic

history of Patagonia. In their new collision model, they propose that the southern continental block represented

by the Deseado Massif together with the southern part of the continent, was separated from SW Gondwana from

Cambrian until Carboniferous times. The final amalgamation was consequence of a northeasterly (present

coordinates) subduction beneath the north Patagonian Massif, and produced intense metamorphism and granite

emplacement in the upper plate that continued until the Early Permian. Predrift restoration shows that Patagonia

was positioned closer to both Africa and Antarctica (Marshall, 1994) and more recent models, e.g., Ghindella et

al. (2002), concur with respect to placing the Deseado Massif much closer to the southern tip of South Africa

and the northern tip of the Antarctic Peninsula outboard of southernmost Patagonia.

In this contribution, we combine novel geochronological and geophysical evidences to suggest that the

Falkland Plateau and the Deseado Massif could be part of the same ancient continental block, which is now

partially hidden under the Atlantic Ocean and the sedimentary cover. If this is the case, the geological evolution

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of southern Patagonia should be considerably older then currently thought, and the models for the tectonic

evolution of this region should be significantly modified.

Geochronological evidences

The Mesoproterozoic history of the Falkland Islands is well known since the first geochronological works

carried on its basement metamorphic rocks from the Cabo Belgrano (Cingolani and Varela, 1976; Rex and

Tanner, 1982). Contrary, most published geochronological data of the few basement rocks of the Deseado

Massif indicate a history of Neoproterozoic sedimentation and metamorphism followed by Silurian and

Devonian granite magmatism (Pankhurst et al., 2003). Nevertheless, the weathered and altered granitoids and

their metasedimentary host rocks analyzed by Pankhurst et al. (2003) using the U-Pb zircon method by

SHRIMP, evidenced prominent components at 1000-1100 Ma. Similarly old zircons were found at the low grade

metamorphic complexes of the Patagonian Andes, located west and southwest of the Deseado Massif, at the

western edge of South America plate (Hervé et al., 2003). Also, neodymium model ages of 1200 Ma were

obtained for Jurassic volcanic rocks from the Deseado Massif by Pankhurst et al. (1994).

Recently, Schilling et al. (2008) presented a Re-Os isotopic study for widely dispersed mantle xenoliths carried

to the surface of southern South America (36º - 52ºS) by Eocene to recent alkaline magmatism. This isotopic

system gives unique chronological information about the time of mantle depletion that is associated with

lithosphere formation. The presented results indicate Mesoproterozoic ages for lithospheric mantle formations on

the Cuyania terrane, which is accepted to be a continental block formed during the Mesoproterozoic, and the

Deseado Massif. Contrary, the lithospheric mantle of the rest of the South America continent is similar to the

present suboceanic mantle, suggesting a relatively recent lithospheric mantle formation from the convecting

mantle. Based on these results and considering the geochronological, geographical and geomorphological

characteristics of the Falkland Islands and the adjoining areas (Fig. 1), Schilling et al. (2008) propose that the

Deseado Massif and the Falkland Plateau can be derived from the same tectonic microplate.

Geophysical evidences

The subsurface extension of the Deseado Massif to the southeast is suggested by geophysical data showing the

presence of an offshore basement high, the Rio Chico-Dungeness Arch (Biddle et al., 1986). More recently,

Tassara et al. (2007) estimated the elastic thickness (Te) over South America and its surrounding plates using a

wavelet formulation of the classical spectral isostatic analysis inverting satellite-derived gravity and

topography/bathymetry. Te is a proxy for lithospheric thickness and lateral variations of this parameter over

continents have been interpreted as reflecting spatial changes in the age-dependent thermal structure of the

continental lithosphere (e.g. Tassara et al., 2007; Pérez-Gussinyé and Watts, 2005). In Figure 1 we present the Te

estimates of southern South America and the surrounding plates, together with a map of the bathymetry and

topography of the area modified from the work of Tassara et al. (2007). The maximum Te values for this region

reach 30 to 40 km just between the Deseado Massif and the Falkland Islands with an apparent extension to the

southwest. These high Te values apparently reflect the presence of a thicker, colder, more rigid and presumably

older lithospheric block compared to surrounding regions. The presence of this block is probably related to the

Proterozoic and relatively depleted lithospheric mantle of the Deseado Massif identified by Schilling et al.

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(2008). Consequently, the Río de la Plata craton also exhibits high Te values (Fig. 1). These results are coherent

with the idea that the Falkland Island and Plateau, and the Deseado Massif belong to the same continental block.

Other striking feature of the Te map is a E-W low Te zone, which coincides relatively well with the NW-SE

collision zone inferred by Pankhurst et al. (2006) between the Deseado Massif and the North Patagonian Massif

(Fig. 1). This low Te region is possibly a consequence of the thermomechanical weakening imposed by the

tectonomagmatic activity occurred in this zone during Carboniferous times, when the intervening oceanic crust

was subducted to the northeast under the North Patagonian Massif, and after the collision that produced

significant crustal anatexis until Early Permian times (Pankhurst et al., 2006).

Our interpretation of Te variations over the southern tip of South America are further supported by global- and

continental-scale seismic tomography results (Ritsema et al., 2004; Vdovin et al., 1999) that show a coherent

high-velocity anomaly at upper mantle depth coinciding with our proposed Falkland-Deseado lithospheric block.

Figure 1. Topography and bathymetry of southern South America (left) and the computed elastic thickness map (right). The main tectonic elements and continental terranes are shown. Abbreviations: ChT: Chilenia Terrane; CT: Cuyania Terrane; PT: Pampia Terrane; RPC: Río de la Plata craton; NPM: North Patagonian Massif; ICZ: Inferred collision zone of Pankhurst et al. (2006); DM: Deseado Massif; FI: Falkland Islands; FDLB: Falkland-Deseado lithospheric block.

Geological implications

If the hypothesis that the Deseado Massif and the Falkland Plateau belong to the same ancient and rigid

continental block is correct, it is possible that they have been together since relatively long time

(Mesoproterozoic?), and suffered considerably less stretching then the surrounding areas. Considering the

generally accepted model where the Falkland Islands need to be rotated some 180º counterclockwise to fit with

eastern front of the Cape Fold Belt, at the southeast cost of Africa, the Deseado Massif should be displaced

considerably southeastward in reconstructions to times before the Atlantic opening. It is also possible that the

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relative motion northward of this rigid microplate against the southern tip of Gondwana, correspondent to the

present southern border of the North Patagonian Massif and southern Africa, was the responsible for the genesis

of the Cape Fold Belt orogen during Carboniferous times. Apparently, the Deseado-Falkland lithospheric block

was amalgamated to southern Gondwana margin, and during its fragmentation and the Atlanctic Ocean opening,

the Deseado-Falkland continental block kept together with the South America continent. This resolves the

problem of the driving force by which the Falkland Islands achieve its position. It seems that a continental region

of this block that suffered considerably more stretching is the Maurice Ewing Massif to the east.

If the Deseado Massif was located considerably eastward previous to the Gondwana fragmentation, it is

possible to think that the Neoproterozoic-early Paleozoic granites outcropped at the NE of the massif are related

to the East Africa-Antarctica orogen that occurred during that time (Jacobs and Thomas, 2004).

Finally, the significant differences of the Deseado Massif and the North Patagonian Massif continental

lithospheres, seems to have important economic implications. In the Deseado Massif, the Jurassic bimodal

volcanism related to the break-up of Gondwana is related with gold mineralization, contrary to the situation

observed in the North Patagonian Massif, where the same volcanism is not associated with gold deposits.

References Biddle, K.T., Uliana, M.A., Mitchum Jr., R.M., Fitzgerald, M.G., Wright, R.C., 1986. “The stratigraphical and structural

evolution of the central and eastern Magallanes Basin, southern South America”. In: Allen, P.A., Homewood, P. (eds), Foreland Basins. International Association of Sedimentology, Special Publications, vol. 8, pp. 41–61.

Cingolani, C.A., Varela, R., 1976. Investigaciones geológicas y geocronológicas en el extremo sur de la Isla Gran Malvina, sector cabo Belgrano (Cabo Meredith), Islas Malvinas. 6° Cong. Geol. Argentino, Actas 1, 457–474.

Ghidella, M.E., Yañez, G., LeBreque, J.L., 2002. Revised tectonic implications for the magnetic anomalies of the western Weddell Sea. Tectonophysics 347, 65–86.

Hervé, F., Fanning, C.M., Pankhurst, R.J., 2003. Detrital zircon age patterns and provenance of the metamorphic complexes of southern Chile. J. South Am. Earth Sci. 16, 107–123.

Jacobs, J., Thomas, R.J., 2004. Himalayan-type indenter-escape tectonics model for the southern part of the late Neoproterozoic–early Paleozoic East African–Antarctic region. Geology 32, 721– 724.

Marshall, J.E.A., 1994. The Falkland Islands: A key element in Gondwana paleography. Tectonics 13 (2), 499–514. Pankhurst, R.J., Hervé, F., Rapela, C.W., 1994. Sm–Nd evidence for the Grenvillian provenance of the metasedimentary

basement of Southern Chile and West Antartica. 7° Cong. Geol. Chileno, Concepción, Actas 2, 1414–1418. Pankhurst, R.J., Rapela, C.W., Loske,W.P., Márquez, M., Fanning, C.M., 2003. Chronological study of the pre-Permian

basement rocks of southern Patagonia. J. South Am. Earth Sci. 16, 27–44. Pankhurst, R.J., Rapela, C.W., Fanning, C.M., Márquez, M., 2006. Gondwanide continental collision and the origin of

Patagonia. Earth-Sci. Reviews 76, 235–257. Pérez-Gussinye, M. and Watts, A.B. 2005. The long-term strength of Europe and its implications for plate-forming processes,

Nature 436 (2005), pp. 381–384. Ramos, V.A., 1984. Patagonia, Un continente a la deriva? 10° Cong. Geol. Argentino, Actas 2, 311–325. Ramos, V.A., 1988. Tectonics of the late Proterozoic–early Paleozoic: a collisional history of Southern South America.

Episodes 11, 168–174. Rex, D.C., Tanner, P.W.G., 1982. “Precambrian age for the gneisses at Cape Meredith in the Malvinas/Falkland islands”. In:

Antartic Geoscience, Campbell Craddock (eds), Symposium on Antarctic geology and geophysics. The University of Wisconsin press, pp. 107–108.

Ritsema, J. and Hendrik, J.V.H. 2004. Global transition zone tomography. J. Geophy. Res., 109, B02302. Schilling, M.E., Carlson, R.W., Conceição, R.V., Dantas, C., Bertotto, G.W. and Koester, E. 2008. Re–Os isotope constraints

on subcontinental lithospheric mantle evolution of southern South America, Earth Planet. Sci. Lett. 268, 89-101. Taylor, G.K., Shaw, J., 1989. The Falkland Islands: new palaeomagnetic data and their origin as a displaced terrane from

southern Africa. American Geophysical Union, Geophysical Monograph 50, 59–72. Thomas, R.J., Cornell, D.H., Moore, J.M., Jacobs, J.J., 1994. Crustal evolution of the Namaqua-Natal metamorphic province,

southern Africa. South African Journal of Geology 97, 8–14. Vizán, H., Somoza, R., Taylor, G., 2005. Paleomagnetic testing the behaviour of Patagonia during Gondwana break-up. In:

Pankhurst, R.J., Veiga, G.D. (Eds.), Gondwana 12: Geological and Biological Heritage of Gondwana, Abstracts. Academia Nacional de Ciencias, Córdoba, Argentina, p. 368.

Oleg, V., Rial, J.A., Levshin, A.L. and Ritzwoller, M.H. 1994. Group velocity tomography of South America and the surrounding oceans. Geophys. J. International 136, 324-340.

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Principal results of the Caracas, Venezuela, Seismic Microzoning Project

Michael Schmitz1, Julio J. Hernández

2, Cecilio Morales

1, Danna Molina

1, Maxlimer Valleé

1,

Jean Domínguez1, Elise Delavaud

3, André Singer

4, Moralis González

1, Victoria Leal

1, & the

Caracas Seismic Microzoning Project Working Group

1

FUNVISIS, Final calle Mara, Urb. El Llanito, Caracas, Venezuela (mschmitz@funvisis.gob.ve) 2

Consultant in Earthquake Engineering, Caracas, Venezuela (julher@cantv.net) 3

IPGP Paris; now at Univ. Potsdam, Germany 4

Consultant to FUNVISIS, Caracas, Venezuela

KEYWORDS : microzoning, seismic hazard, site effects, response spectra, landslide hazard

Abstract We present the principal results of the Caracas Seismic Microzoning Project realized in the years 2005 to 2007

in the Venezuelan capital. Its location close to the plate boundary between the South America and the Caribbean plates, the emplacement within a sediment filled half graben and the extensions on steep hills are responsible for the moderate to high seismic hazard of Caracas. During the execution of the project, extensive geological and geophysical investigations were done in order to determine the distribution of the different units within the valley. For the hillside areas, the landslide hazard was estimated based on available geotechnical information. A detailed analysis of the seismic hazard at the outcropping rock was derived, deconvolved to bedrock and used as input for the determination of response spectra at different subsoil conditions within the sedimentary valley, and later calibrated with actual earthquake spectra.

Introduction

During its history, Caracas has suffered several destructive earthquakes. The most recent one, the July 1967

Caracas earthquake, a magnitude 6.6 earthquake which occurred some 25 km northwest of Caracas as a multi-

event earthquake (Suárez & Náb lek, 1990), caused damage to numerous buildings, and the collapse of 4

multistory buildings, with more than 300 people killed. Damage investigations of buildings were performed in

detail, including soil and building dynamical characteristics, and the earthquake engineering characteristics of

the deposits, seen as the fundamental factor for earthquake damage (FUNVISIS, 1978; Seed et al., 1970). The

particular behavior of the thick soil deposits in the east of Caracas valley had attracted attention during the past

decades, leading to detailed studies of seismic response and ground shaking characteristics (e.g. Papageorgiou

and Kim, 1991; Abeki et al., 1998; Semblat et al., 2002; Rocabado et al., 2006). In the years 2003 to 2005, the

Japan International Cooperation Agency JICA developed a “Basic Study on Disaster Prevention”, studying

various scenarios for earthquake disaster (Yamazaki et al., 2005). Nevertheless, basin effects as observed during

the Caracas 1967 earthquake could not be modeled thoroughly. A fast growing part of the city with 3.5 million

inhabitants comprises informal housing at steep hillsides surrounding the valley. Earthquake or rainfall triggered

landslide is the principal hazard in these areas. Despite all the previous studies, local conditions have not been

taken into account so far for building regulations in Caracas, which is located in the seismic zone 5 with

horizontal accelerations at rock sites of 0.3 g for 475 years mean return period (COVENIN, 2001). In order to

prepare a concise base for local urban planning and regulations, a seismic microzonation study was developed

from 2005 to 2007. Zones of distinctive seismic response within the sedimentary valley, as well as areas of

seismic landslide hazard, have been determined by the project.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 500-503

501

Methodology

The estimation of the soil movement in different areas of a city is an efficient tool for mitigation of seismic

risk, being site amplifications crucial for local behavior (Bard, 1999). A resume of the methodology applied in

this study is presented in Hernández et al. (2006). The principal elements used are: 1) Probabilistic assessment of

seismic hazard at rock sites, 2) Identification of soil and basin site effects, 3) Definition of microzones of similar

seismic response. For the latter one, information from various sources was use, as geomorphologic, geological

and geophysical analysis of the sediment in the valley, geotechnical analysis of the rocks exposed at the hillsides

and analysis of the detailed damage information from the 1967 Caracas earthquake. We developed generic

models of dynamic response using equivalent linear analysis (Schnabel et al., 1972), considering variations in

sediment thickness (between 10 and 350 m) and shear wave velocity of the upper 30 m (Vs30 between 150 and

650 m/s), which were grouped into 12 classes (Table 1) according to their typical behavior. As an important

feature, the dynamic results are calibrated and corrected by comparing them with actual earthquake spectra

(PEER, 2005).

Table 1. Groups of generic soil profiles used for dynamic response of sedimentary sites.

Outside the sedimentary areas, the earthquake triggered landslide hazard is evaluated using information

regarding geology, geomorphology, slope, weathering and anthropic modifications. Thus, priority areas for

intervention may be identified. Additionally, topographic effects are taken into account for seismic response at

hillside areas. Part of the study comprises the evaluation of buildings regarding their typified structural behavior,

which will point out the priorities for retrofitting of existing buildings regarding their location within the

different microzones. All the information generated within the project is introduced in a Geographic Information

System (GIS), which will enable the interaction with local institutions and urban planners for fast

implementation of the recommendations. Interaction with local communities is organized by the “Aula Sísmica

Madeleilis Guzmán” at FUNVISIS, a unit that works in disaster prevention education.

Principal results

The seismic hazard in the area of Caracas, which is determined to 0.3 g following the seismic building code

(COVENIN, 2001) was detailed with values ranging from 0.3 g in the north and 0.21 g in the south (Figure 1).

The microzones of similar seismic behavior within the sedimentary valley were assigned following the above

described parameters. The resulting spectra, which will include basin effects derived from 3-D modeling of the

seismic response (Delavaud, 2007), are compared to the building codes COVENIN (2001), NEHRP (BSSC,

2003) and EUROCODE 8 (CEN, 2003); they are closer to the last one. As an example, the spectra for group

GP03 and weathered bedrock and for group GP12 are displayed (Figure 2). The distribution of microzones was

VS,30 (m/s) H, deposit (m)

185 185 to 325 > 325

< 60 GP-01 GP-02 GP-03

60 to 120 GP-04 GP-05 GP-06

120 to 220 GP-07 GP-08 GP-09

> 220 GP-10 GP-11 GP-12

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 500-503

502

calibrated by different means, as there are “realistic” soil profiles from deep boreholes, predominant periods

from H/V, and by experimental transfer functions. Data from three deep (110 to 280 m depth) accelerographic

stations will help to constrain the results in the future.

Figure 1. Zoning of seismic hazard at outcropping rock and microzones of similar seismic behavior for the Caracas valley.

Figure 2. Adjusted spectra for group GP-03 and weathered bedrock (left) and sedimentary thickness of more than 220 m (right) are displayed; both examples for Vs30 > 325 m/s.

Conclusions

During the 1967 Caracas earthquake, the damage distribution evidenced strong site effects within the

sedimentary valley. Nevertheless, the principal parameters which control the seismic response, as sediment

thickness of more than 50 m and basin geometry, are not considered in the Venezuelan building code

(COVENIN, 2001). The results of the project presented here allow assign modified response spectra for the

different parts of Caracas. An evaluation of earthquake triggered landslide is also included in the project. An

efficient elaboration of recommendations and local building codes will be crucial for the implementation of the

project results.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 500-503

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Acknowledgments Funding was provided by the “Proyecto de Microzonificación Sísmica en las Ciudades Caracas y Barquisimeto” (FONACIT–BID II 2004000738). Further members of the Caracas Seismic Microzoning Project Working Group are: A. Aguilar, I. Aguilar, L. Alvarado, E. Amarís, M. Andrade, F. Anzola, J. Araque, F. Audemard, J. Azuaje, P.Y. Bard, H. Cadet, V. Cano, E. Caraballo, A. Castillo, C. Cornou, J. Delgado, P. Feliziani, Y. Flores, K. García, J. Guzmán, A. Hernández, A. Justiniano, R. López, W. Marín, G. Malavé, J. Moncada, R. Ollarves, J. Oropeza, M. Palma, A. Petitjean, B. Quintero, H.Rendón, V. Rocabado, J. Rodríguez, L. Rodríguez, G. Romero, S. Safina, J. Sánchez, M. Tagliaferro, F. Urbani, R. Vásquez, M. Villar, J.P. Vilotte, A. Zambrano, H. Zambrano, J. Zamora.

References

Abeki N., Seo K., Matsuda I., Enomoto T., Watanabe D., Schmitz M., Rendón H., Sánchez A., 1998. “Microtremor observations in Caracas city, Venezuela.” In: Irikura et al., (ed.). The Effects of Surface Geology on Seismic Motion, Rotterdam, AA Balkema, 619-624.

Bard, P.Y., 1999. “Microtremor measurements: a tool for site effect estimation?” In: Irikura, K., Kudo, K., Okada, H. & Sasatani, T. (eds.), The Effects of Surface Geology on Seismic Motion - Recent progress and new Horizon on ESG Study, vol. 3, Balkema, Rotterdam, 1251-1279.

BSSC, 2003. NEHRP recommended provisions for seismic regulations for new buildings and other structures (FEMA 450). Building Seismic Safety Council (BSSC), NIBS, Washington.

CEN, 2003. Eurocode 8: Design of structures for earthquake resistance. European Standard, English version, Comité Européen de Normalisation (CEN), Brussels.

COVENIN, 2001. Edificaciones sismorresistentes, COVENIN 1756:2001. Comisión Venezolana de Normas Industriales (COVENIN), FONDONORMA, MCT, MINFRA, FUNVISIS, Caracas.

Delavaud, E., 2007. Simulation numérique de la propagation d'ondes en milieux géologiques complexes : application à l'évaluation de la réponse sismique du bassin de Caracas. PhD thesis, IPGP, France, pp. 155.

FUNVISIS, 1978. Segunda Fase del Estudio del Sismo ocurrido en Caracas el 29 de Julio de 1967. Ministerio de Obras Públicas, Comisión Presidencial para el Estudio del Sismo, FUNVISIS, Caracas, Venezuela, Vol. A, pp. 517.

Hernández, J.J., Schmitz, M., Audemard, F., Malavé, G., 2006. “Marco conceptual del proyecto de microzonificación de Caracas y Barquisimeto”. VIII Congreso Venezolano de Sismología e Ingeniería Sísmica, Valencia, Venezuela, 2006, Memorias en CD.

Papageorgiou A.S., Kim J., 1991. Study of the propagation and amplification of seismic waves in Caracas valley with reference to the 29 July 1967 earthquake: SH waves. Bull. Seis. Soc. Am., 81, 2214-2233.

PEER, 2005. PEER NGA Database. Pacific Earthquake Engineering Research Center (PEER), California. Rocabado, V., Schmitz, M., Rendón, H., Vilotte, J.-P., Audemard, F., Sobiesiak, M., Ampuero, J.-P., Alvarado, L., 2006.

Modelado numérico de la respuesta sísmica 2D del valle de Caracas. Revista de la Facultad de Ingeniería de la U.C.V., vol. 21 (4), 81-93.

Schnabel, P., Lysmer, J., Seed, H., 1972. SHAKE - A computer program for earthquake response analysis of horizontally layered sites. Earthquake Engineering Research Center, Report No. UCB/EERC-72/12. University of California, Berkeley.

Seed HB, Idriss IM, Dezfulian H., 1970. Relationships between soil conditions and building damage in the Caracas earthquake of July 29, 1967. EERC-Report 70-2, Berkeley, California, 40 pp.

Semblat, J.F., Duval, A.M., Dangla, P., 2002. Seismic site effects in a deep alluvial basin: numerical analysis by the boundary element method. Computers and Geotechnics, 29, 573-585.

Suárez G, Náb lek J., 1990. The 1967 Caracas earthquake: Fault geometry, direction of rupture propagation, and seismotectonic implications. J. Geophys. Res., 95, 17459-17474.

Yamazaki, Y., Audemard, F., Altez, R., Hernández, J., Orihuela, N., Safina, S., Schmitz, M., Tanaka, I., Kagawa H., and JICA Study Team-Earthquake Disaster Group, 2005. Estimation of the seismic intensity in Caracas during the 1812 earthquake using seismic microzonation methodology. Revista Geográfica Venezolana, Número Especial 2005, 199-216.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 504-507

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Anatomy of the Central Andes: Distinguishing between western, magmatic Andes and eastern, tectonic Andes

Thierry Sempere1 & Javier Jacay

2

1 LMTG, Université de Toulouse, CNRS, IRD, OMP, 14 avenue Edouard Belin, F-31400 Toulouse, France

(sempere@lmtg.obs-mip.fr) 2 EAP Ingeniería Geológica, Universidad Nacional Mayor de San Marcos, Lima, Peru (j_jacay@yahoo.com)

KEYWORDS : Andes, crustal thickening, magmatism, tectonic shortening, southern Peru

Introduction: The Andes under the weight of a paradigm

Scientific activity and production take place under the light of paradigms (Kuhn, 1962) and the geosciences

make no exception. Plate tectonics is the paradigm that currently governs our large-scale understanding of the

Earth. Paradigms orient research at all scales, a phenomenon which geoscientists are rarely aware of. In the case

of the Central Andes, most studies conducted since the late 1980s have admitted, explicitly or not, that crustal

thickening has been primarily achieved through tectonic shortening of the South American margin, and thus that

magmatic additions to the crust represented only a minor contribution to crustal thickening. Because this idea

was particulary well developed in Isacks’s (1988) landmark paper, this paradigm may be referred to as “the

Isacksian paradigm”. Since then, many researchers in the Central Andes have concentrated on tectonic

shortening; however, crustal thickness cannot be accounted for by the available shortening estimates, especially

in the arc and forearc (e.g., Schmitz, 1994; Sempere et al., 2008, and references therein).

Yet, aside from the tectonic, i.e. mechanical, interaction of the converging plates, the other first-order

characteristic feature of subduction zones is the production of arc magmatism. Simple logics implies that

tectonic and magmatic processes should therefore be viewed as two related aspects of one same system. The idea

that arc orogens are formed through magmatic accretion forced by subduction is widely admitted in island arc

contexts (e.g., Tatsumi and Stern, 2006), but has only received minor attention in the case of the Central Andes

— albeit with noteworthy exceptions (e.g., James, 1971b; Thorpe et al., 1981; Kono et al., 1989; James & Sacks,

1999; Haschke and Günther, 2003) —, a situation largely due to the dominance of the Isacksian paradigm.

The belief that the Central Andes originated by shortening has commonly biased cartography, for instance by

forcing high-angle or poorly-exposed faults to be mapped as reverse faults and thrusts. Some areas were mapped

in dramatically different ways by geologists who favored distinct models (e.g., Sempere, 2000; Wörner &

Seyfried, 2001), and extensional structures were often overlooked. Moreover, observations and models from a

variety of undoubtedly extensional settings in Europe and Africa now teach that some structural geometries

previously thought to be typical of contractional processes, as in the Central Andes, in fact also occur in

extensional contexts, in particular where normal faults were initiated as flexure-forming blind faults (e.g., Finch

et al., 2004). We have thus undertaken the revision of traditional mapping in southern and central Peru in order

to better understand the detailed anatomy of this portion of the Central Andes.

“Western Andes” and “eastern Andes” in southern Peru

Southern Peru provides a convenient observatory for a detailed anatomy of the Central Andean Orocline

(CAO). Identification and correction of mapping biases result in major revisions: the forearc, arc, and SW

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 504-507

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Altiplano (henceforth “western Andes”) appear to have been dominated by transcurrence (including

transpressional deformation) and extension since ~30 Ma (Sempere & Jacay, 2006, 2007), in contrast with the

NE Altiplano, Eastern Cordillera (EC), and sub-Andean belt (henceforth “eastern Andes”), where shortening has

been indeed significant. Separating these two contrasting orogenic domains, seismic tomography detected a

subvertical lithospheric-scale boundary in the northern Altiplano of Bolivia (Dorbath et al., 1993) and in its

prolongation, i.e. along the SW edge of the EC of southern Peru, the distribution of magmatic rocks (Sempere et

al., 2004) and the isotopic geochemistry of mantle-derived rocks (Carlier et al., 2005) also mapped a subvertical

lithospheric boundary, which coincides at the surface with the SFUACC major fault system (Fig. 1).

Figure 1. Approximate partition between the western, magmatic Andes and the eastern, tectonic Andes in the Central Andean Orocline. The magmatic Andes (forearc, Western Cordillera [WC], SW Altiplano) are characterized by little or no shortening and a crust thickest across the arc, whereas the tectonic Andes (NE Altiplano, Eastern Cordillera [EC], sub-Andean belt) display evident, substantial shortening. This implies that crustal thickening was achieved by magmatic accretion in the former, and by tectonic shortening in the latter. In southern Peru, the boundary between the magmatic and tectonic Andes is marked by the lithospheric-scale Urcos-Ayaviri-Copacabana-Coniri fault system (SFUACC in Spanish abbreviation), but may be transitional elsewhere. Hatched pattern: areas affected by Cenozoic shortening older than ~25 Ma.

In the western Andes, i.e. SW of the SFUACC, synorogenic basins formed in extension and along transcurrent

faults. At least one low-angle extensional detachment, placing near-vertical Miocene conglomerates over a

Cretaceous unit, occurs just west of Lake Titicaca (Sempere & Jacay, 2006). Significant transcurrent faulting,

including transpressional deformation, developed along specific structures over southern Peru, including the

western Andes. However, transpressional structures in the forearc and arc, such as the Cordillera de Domeyko in

Northern Chile, can only account for relatively minor shortening and crustal thickening. The Pacific Andean

escarpment is the locus of oceanward reverse faulting, suggesting incipient gravitational collapse of the WC

(Wörner & Seyfried, 2001; Wörner et al., 2002; Sempere & Jacay, 2006, 2007) instead of tectonic shortening.

WC EC

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 504-507

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A magmatic origin for the western Andes

Although the lack of surface evidence for significant shortening in the western Andes was accomodated in

some graphic constructions by supposing blind crustal duplexes or insertion, at the base of the crust, of crustal

slices tectonically displaced from the margin, no evidence has been obtained yet for any of such hypotheses,

which appear to be largely paradigm-driven. The combined facts that in the western Andes the orogeny was

accompanied by extensional and transcurrent tectonics, and that transpressional deformation cannot account for

significant crustal thickening (as it produces only localized shortening), imply that the Isacksian paradigm, i.e.

the assumption that the Central Andean orogenic build-up was mostly driven by tectonic shortening, has to be

questioned in the western Andes.

A key insight into this issue is provided by the fact that in southern Peru the crust is thickest across the arc

(Fig. 2), as demonstrated by seismic and gravity studies (James, 1971a; Kono et al., 1989). Association of

maximum crustal thickness with the arc region simply points to magmatism as the cause of crustal thickening in

the western Andes, reinforcing previous similar interpretations (e.g., James, 1971b; Thorpe et al., 1981; Kono et

al., 1989; Schmitz, 1994; James & Sacks, 1999), which unfortunately have been largely disregarded.

The idea that the arc crust was primarily thickened by magmatic mass transfer from the mantle is supported by

the fact that the isotopic characteristics of most Andean magmas indicate that they largely consist of material

extracted from the mantle (e.g., Pitcher et al.,

1985). Furthermore, I-type magmatism, a

typical feature of Andean arc batholiths

(Pitcher et al., 1985), is now understood to

result from the reworking of crustal materials

by mantle-derived magmas, and is even

viewed to drive the coupled growth and

differentiation of continental crust (Kemp et

al., 2007). Crustal growth rates at arcs are now

known to be at least 40-95 km3/km.Myr (e.g.,

Tatsumi & Stern, 2006), i.e. at least twice the

rates estimated by Reymer and Schubert

(1984), who nevertheless mentioned a few

cases with arc crustal growth rates as high as

~300 km3/km.Myr. Volumes of volcanic rocks

erupted at the surface were invoked to discard

magmatic addition as a significant cause of

crustal thickening, but updated estimates are

much higher (de Silva and Gosnold, 2007);

besides, no secure constraints are available on

the ratio of volcanic volumes to total

magmatic volumes, and this ratio might well

be anomalously low in the case of thick crusts.

magm

ati

c c

ord

ill.

= a

rc =

WC

SW NE

A

B

C

tecto

nic

cord

ille

ra =

EC

Figure 2. Topography (A), Bouguer anomaly (B), and distribution of crustal densities and thicknesses (C) in southern Peru, after Kono et al. (1989). Crustal thickness is clearly maximum across the magmatic cordillera (the volcanic arc, or Western Cordillera [WC]) and decreases toward the Eastern Cordillera (EC) across the Altiplano. The NE edge of the Eastern Cordillera, which is undisputably of tectonic origin, corresponds to a marked subvertical stair-step in the crustal structure (red arrows).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 504-507

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Conclusion

Updated geological mapping in southern Peru is confirming that tectonic shortening has been insignificant

southwest of the SFUACC fault system, and certainly cannot explain the outstanding crustal thickening in the

western Andes. The western Andes, which include the arc region, must therefore have formed by magmatic

accretion, as already suggested by the abundant geochemical database. In contrast with the eastern Andes, which

are indeed of tectonic origin, the western Andes have been built by magmatic processes, confirming previous but

disregarded conclusions (e.g., James, 1971a,b; Kono et al., 1989; Schmitz, 1994; James and Sacks, 1999). After

all, crustal growth and orogenic build-up by subduction-related magmatism are known elsewhere to be common

processes in island arcs as well as continental arcs (e.g., Tatsumi and Stern, 2006; Lee et al., 2007).

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rocks delineate two lithospheric mantle blocks beneath the southern Peruvian Altiplano. Geology 33, 601-604. de Silva, S.L., Gosnold, W.D., 2007. Episodic construction of batholiths: Insights from the spatiotemporal development of an

ignimbrite flare-up. Journal of Volcanology and Geothermal Research 167, 320–335. Dorbath, C., Granet, M., Poupinet, G., Martinez, C., 1993. A teleseismic study of the Altiplano and the Eastern Cordillera in

northern Bolivia: New constraints on a lithospheric model. Journal of Geophysical Research 98: 9825–9844. Finch, E., Hardy, S., Gawthorpe, R., 2004. Discrete-element modelling of extensional fault propagation folding above rigid

basement fault blocks. Basin Research 16, 489–506. Haschke, M., Günther, A., 2003. Balancing crustal thickening in arcs by tectonic vs. magmatic means. Geology 31, 933-936. Isacks, B.L., 1988. Uplift of the central Andean plateau and bending of the Bolivian orocline. Journal of Geophysical

Research 93, 3211–3231. James, D. E. 1971a. Andean crust and upper mantle structure. Journal of Geophysical Research 76: 3246-3271. James, D. E. 1971b. Plate tectonic model for the evolution of the Central Andes. Geological Society of America Bulletin 82:

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From island arcs to accretion and continental arc magmatism. Earth and Planetary Science Letters 263, 370–387. Pitcher, W.S., Atherton, M.P., Cobbing, E.J., Beckinsale, R.D. (Eds), 1985. Magmatism at a Plate Edge: The Peruvian

Andes. Glasgow: Blackie / New York: Halsted Press, 323 p. Reymer, A., Schubert, G., 1984. Phanerozoic addition rates to the continental crust and crustal growth. Tectonics 3, 63–77. Schmitz, M., 1994. A balanced model of the southern Central Andes. Tectonics 13, 484–492. Sempere, T., 2000. Discussion of “Sediment accumulation on top of the Andean orogenic wedge: Oligocene to late Miocene

basins of the eastern Cordillera, southern Bolivia” (Horton, 1998). Geol. Society of America Bulletin 112, 1752–1755. Sempere, T., & Jacay, J. 2006. Estructura tectónica del sur del Perú (antearco, arco, y Altiplano suroccidental). Extended

abstract, XIII Congreso Peruano de Geología, Lima, 324-327. Sempere, T., & Jacay, J. 2007. Synorogenic extensional tectonics in the forearc, arc and southwest Altiplano of southern

Peru, Eos Trans. AGU 88(23), Joint Assembly Suppl., Abstract U51B-04. Sempere, T., Jacay, J., Carlotto, V., Martínez, W., Bedoya, C., Fornari, M., Roperch, P., Acosta, H., Acosta, J., Cerpa, L.,

Flores, A., Ibarra, I., Latorre, O., Mamani, M., Meza, P., Odonne, F., Orós, Y., Pino, A., Rodríguez R., 2004. Sistemas transcurrentes de escala litosférica en el sur del Perú. In: J. Jacay, T. Sempere (Eds.), Nuevas contribuciones del IRD y sus contrapartes al conocimiento geológico del sur del Perú. Sociedad Geológica del Perú, Publicación Especial 5, 105-110.

Sempere, T., Folguera, A., & Gerbault, M. (eds.). 2008. New insights into the Andean evolution: An introduction to contributions from the 6th ISAG symposium (Barcelona, 2005). Tectonophysics, in the press.

Tatsumi, Y., Stern, R.J., 2006. Manufacturing continental crust in the subduction factory. Oceanography 19, 104–112. Thorpe, R. S., Francis, P. W., & Harmon, R. S. 1981. Andean andesites and crustal growth. Philosophical Transactions of the

Royal Society of London A 301: 305–320. Wörner, G., Seyfried, H., 2001. Reply to the comment by M. García and G. Hérail on “Geochronology (Ar–Ar, K–Ar and

He-exposure ages) of Cenozoic magmatic rocks from northern Chile (18–22°S): Implications for magmatism and tectonic evolution of the central Andes” by Wörner et al. (2000). Revista Geológica de Chile 28, 131–137.

Wörner, G., Uhlig, D., Kohler, I., Seyfried, H., 2002. Evolution of the west Andean scarpment at 18°S (N. Chile) during the last 25 Ma: Uplift, erosion and collapse through time. Tectonophysics 345, 183–198.

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Direct versus indirect thermochronology: What do we truly trace? An example from SE Peru and its implication for the geodynamic development of the Andes

Diane Seward1, Geoffrey M.H. Ruiz

2,1, & Julien Babault

3

1 ETH Zurich, Switzerland

2 University of Neuchatel, Switzerland (Geoffrey.ruiz@unine.ch)

3 Universitat Autónoma de Barcelona, Spain

To quantify long-term denudation rates, research groups commonly applied low-temperature

thermochronometric methods to rock now exposed at the surface. This approach on bedrocks from the hinterland

is sometimes limited since erosion has often removed the record of earlier stages of orogenic growth. To

overcome this shortcoming, researchers have increasingly studied since 20 years orogenic sedimentary records

combining detrital thermochronological analyses with sedimentary petrography but also modelled detrital age

populations from true bedrock catchments.

We propose here to study the denudation history of a region located in the Eastern Cordillera of SE Peru. Our

approach consists on analysing present-day erosional products along five different river catchments for the

Apatite Fission-Track (AFT) thermochronometer. Up to four age populations were extracted from the analyses

of 100 grains per sample. Age populations range between 80 and 0.5 Ma with a majority of age populations

younger than 10 Ma. These AFT analyses from the ‘true’ present-day erosion product of the chain are compared

with ones from an 'artificial' one we generated and this to investigate the recent evolution of the eastern Andes.

The artificial detrital record was engendered by the combination of 197 individual grain ages we recently

produced from a bedrock profile in the region (Ruiz and Andriessen, in press.). Interestingly, the 'artificial' sand

express a clear homogeneous AFT signal with a single and pooled AFT age of 4.1 ± 0.1 Ma. This age is identical

to the youngest age population (P1) we extracted from the 'true' sand within the same catchment (4.4 ± 0.4 Ma)

and suggest that the ‘true’ dated grains of the P1 population were derived from, if not this one, a region with

similar thermal record. Our results are of main importance because they indicate for the first time that a detrital

age population, once statistically individualized and limitations of the method perfectly excluded, most likely

reflects the erosion in a single part of a catchment. In the eastern Andes of Peru, the older age populations we

extracted are probably derived from upper levels within the catchment that reflect by their presence, but not

directly quantify, former denudation. Reversely, the youngest age populations for all present-day river sands are

younger than 6.8 Ma. These data point towards lower levels of the eastern Andes that undergo rapid denudation

and this since recent time (Ruiz and Andriessen, in press.) because of the preservation of older thermal record.

The approach we developed is innovative and aims to reduce the amount of necessary analysis to constrain

long-term denudation rates in different orogenic settings. It also hosts a methodological aspect by comparing

results from direct (bedrock) and indirect (present-day river sands) thermochronological analyses within the

same catchment.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 509-512

509

Major mid-Cretaceous plate reorganization as the trigger of the Andean orogeny

Rubén Somoza

CONICET - Departamento de Ciencias Geológicas, FCEyN, Universidad de Buenos Aires. Pabellón 2, Ciudad

Universitaria, C1428EHA Buenos Aires, Argentina (somoza@gl.fcen.uba.ar)

KEYWORDS : mid-Cretaceous, plate tectonics, Andes, Alps, Himalayas

Introduction

Global plate reorganizations are inescapable events in plate tectonics and must episodically occur. A major

mid-Cretaceous plate reorganization marked the final dismembering of Gondwana leading to consolidation of

the major present-day continents and oceanic basins. In those times occurred the physical disconnection between

South America and Africa as well as the separation of Australia from Antarctica and India from Madagascar.

Precise dating of this widespread event is precluded by the lack of seafloor magnetic anomalies from earliest

Aptian to the end of Santonian. However extrapolated ages of 95 ± 5 Ma have been assigned to changes in

relative plate motion in the South Atlantic Ocean, southwest and southeast Indian Ocean, and Weddell Sea. It is

worth noting that the pole of opening for the Central Atlantic does not change for the Aptian-Santonian time

interval, suggesting that the mid-Cretaceous plate reorganization mainly affected the former Gondwana region.

The hotspot (HS) fixity axiom was intensively used for many years for tectonics and geodynamics analyses, in

particular to determine plate motions with respect to the mantle. However recent findings point to failure of the

fixed-HS hypothesis, indicating that the emerging, more realistic scenario where sub-lithospheric melting

anomalies move and deform in concert with flow in the surrounding mantle need to be allowed for assaying

tectonic and geodynamic models. In this report, a moving-HS model (O´Neill et al., 2005) and paleomagnetism

are applied in the analysis of the mid-Cretaceous plate reorganization and its implications for the development of

major present-day orogenic systems in general, and the Andean Cordillera in particular.

Cretaceous to Recent evolution of the Andean margin

The Andean magmatic arc that parallels the western margin of South America was almost permanently active

since at least the Early Jurassic, pointing out a long-lived subduction history. The coeval evolution of the

continental margin may be divided into two periods. During Jurassic to Early Cretaceous times most of the

margin was very close to sea level, with backarc shallow seas and extensional basins. In contrast, the Late

Cretaceous to Recent interval is characterized by rising of arc massifs and increasing predominance of horizontal

shortening, leading to progressive crustal thickening, uplift, development of thrust belts and associated foreland

basins.

Figure 1 shows the 120 Ma reconstruction to the moving-HS framework of O´Neill et al. (2005). The synthetic

flowlines describing the motion of Africa with respect to the moving-HS suggest that slab pull force in the

eastern Tethys subduction zone was an important factor in controlling the motion of that continent during the

120-100 Ma time interval. By those times South America was physically connected to northwest Africa

throughout an incipient extensional region in the present day equatorial Atlantic. This way, South America must

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 509-512

510

have felt both the slab pull force at the Tethys trench and the competing force derived from suction at the

Andean subduction zone (Fig. 1). The moving-HS model predicts that about 75% of the 120-100 Ma full

spreading in the South Atlantic Ocean is associated with African “absolute” motion, implying eastward motion

of the young mid-ocean ridge and, by inference, little (~1.5 cm/yr average in the model) westward motion of

South America. This scenario, with South America experiencing little motion with respect to the mantle, allows

considering episodes in which oceanward motion of the Andean trench due to slab rollback was faster than

westward continental motion, yielding a mechanism to account for the extensional conditions in the western

continental margin during the considered time interval. Although no moving-HS reconstructions older than

120 Ma are available, paleomagnetism indicates that South America experienced counterclockwise rotation

about a northern pole between 135-125 Ma, suggesting that the continent moved away from its western

subduction zone in those times, also consistent with development of extensional conditions at the Andean

margin. Hence, paleomagnetic and moving-HS kinematics allow interpreting the development of extensional

tectonics in the early Andean margin as the product of episodic divergence between the trench and the

continental interior.

Extensional conditions dominated in Peru and central-northern Chile until the Cenomanian (Cobbing et al.,

1981; Atherton and Webb, 1989; Mpodozis and Allmendinger, 1993). On the other hand, the first widespread

contractional events in the Andean Cycle seem to have occurred in Santonian-Campanian times (Mégard, 1987;

Ladino et al., 1999; Tomlinson et al., 2001), suggesting that they began a little later than the final disconnection

between Africa and South America in the present day Equatorial Atlantic.

Rifting in the sheared Equatorial Atlantic margins (Fig. 1) started in Aptian times and complete continental

disconnection occurred some time during the Cenomanian-Turonian (Basile et al., 1998), although it seems that

deepwater connection between central and south Atlantic was not established until Turonian-Coniacian or even

Santonian times. The moving-HS model predicts that the westward motion of South America substantially

increased (Central Andean average ~4.5 cm/yr between 90-60 Ma) after the final continental disconnection in

the Equatorial Atlantic region (inset “a” in Figure 1). This faster westward drift likely led to episodes in which

the continent effectively overrode the Andean trench, in agreement with the development of compressive events

at its leading edge. This tectonic behavior dominated the Late Cretaceous to Recent evolution of the margin,

resulting in an important (predominant?) factor for mountain building in the Andean region.

The moving-HS framework further implies that increasing westward motion of South America was associated

with a desacceleration of African drift at about 90 Ma (inset “b” in Figure 1). This African motion slowdown

and the continued expansion in the South Atlantic imply that the spreading ridge must have began to move

westward with respect to the mantle, substantially increasing the westward drift of South America. In particular,

a velocity increment of ~200 % is predicted between 90 and 80 Ma (inset “a” in Figure 1), a time interval that

includes the beginning of contractional deformation in the Andes.

The Alps and the Himalayas

The moving-HS model allows envisaging a kinematic scenario where, prior to 90 Ma, northeastern Africa and

northern India (or Greater India) constituted the leading edges of these independently drifting landmasses

towards the eastern Tethys trench (Fig. 1). The ~90-88 Ma separation of India from Madagascar (Storey et al.,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 509-512

511

1995; Torsvik et al., 2000; Raval and Veeraswamy, 2003) and the associated development of the Central Indian

oceanic ridge (inset “b” in Figure 1) led to Africa to be almost surrounded by plate-border-parallel spreading

ridges. The latter configuration greatly inhibited African motion excepting towards the Mediterranean region, the

only remaining “free face” of Africa in the Late Cretaceous. Thus, African motion slowdown at 90 Ma may be

related to the establishment of an almost complete girdle of spreading systems around this continent.

Tethys

120

908060

a

b

120

90

60

90

AfricaSAm.

South Atlantic fullspreading 120-100 Ma

W E

SB

MB

Figure 1. Earliest Aptian reconstruction to the moving-HS framework. Oceanic spreading systems (mainly based on identifications of the M0 magnetic anomaly) are shown in red. SB and MB depict Somali and Mozambique basins, respectively. Northern (grey) star is the 120-100 Ma stage pole for Africa-South America relative motion. Southern (black) star is the pole describing the motion of Africa with respect to the moving-HS between 120 and 100 Ma (dashed small circle sectors being the associated synthetic flowlines). Box in the lower part of the draw show the Africa-South America divergence (at the “X” site) decomposed into motion of each one of these continents relative to the moving-HS framework. Inset “a” depicts the westward motion of the Andean region between 120 and 60 Ma. Note acceleration between 90 and 80 Ma, coincident with the beginning of compressive tectonics in the Andes. Inset “b” shows the motion of Africa with respect to the moving-HS between 120 and 60 Ma, note the motion slowdown after 90 Ma. Dashed lines represent the 90-60 Ma synthetic flowlines of the motion of Africa and India with respect to the moving-HS. India and Madagascar are reconstructed at 90 Ma in order to show the paleogeography at the beginning of spreading in the Central Indian Ocean.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 509-512

512

In this context, the beginning of Africa-Europe convergence at ~90 Ma triggered the Alpine orogeny, leading

to the development of magmatic arcs and the build-up of regional compressional stresses and associated

metamorphic events (Ziegler, 1988; Dewey et al., 1989; Okay et al., 2001; Carrapa and Wijbrans, 2003; Ziegler,

2005; Stampfli and Kozur, 2006). On the other hand, the coeval, almost E-W standstill of Africa together with

the continued accretion of oceanic lithosphere at its eastern and western margins resulted in fast motion of both

India and South America because the South Atlantic and Central Indian spreading ridges also moved apart from

the then leisurely drifting Africa. In particular, this kinematics led to fast northward drift of India, accounting for

almost the whole oceanic expansion in the early Central Indian ridge, which culminated with its collision with

Asia and the associated formation of the Himalayas.

Thus, it is suggested that the above described mid-Cretaceous plate reorganization triggered the Andean and

Alpine orogenies as well as the beginning of the plate tectonic conditions that led to the formation of the

Himalayas.

References Atherton, M.P. & Webb, S., 1989. Volcanic facies, structure and geochemistry of the marginal basin rocks of central Perú,

Journal of South American Earth Sciences 2: 241-261. Basile, C., Mascle, J., Benkhelil, J., & Boullin, J.P., 1998. Geodynamic evolution of the Côte d´Ivore-Ghana transform

margin: an overview of Leg 159 results. In Mascle, J., Lohmann, G.P., Moullade, M. (ed.): Proceeding Ocean Drilling Program Scientific Results 159: 101-110.

Carrapa, B. & Wijbrans, J., 2003. Cretaceous 40Ar/39Ar detrital mica ages in Tertiary sediments shed a new light on the Eo-Alpine evolution. In Forster, M., Wijbrams, J. (ed.): Geochronology and Structural Geology, Journal of Virtual Explorer 13: paper 2

Cobbing, E.J., Pitcher, W.S., Wilson, J.J., Baldock, J.W., Taylor, W.P., McCourt, W. & Snelling, N.J., 1981. The geology of the western cordillera of northern Perú, London Institute of Geological Sciences, Overseas Memoir 5: 1-143

Dewey, J.F., Helman, M.L., Knott, S.D., Turco, E. & Hutton, D.H.W., 1989. Kinematics of the western Mediterranean. In Coward, M.P., Dietrich, D., Park, R.G. (ed.) Alpine tectonics, Geological Society Special Publication 45: 265-283

Ladino, M., Tomlinson, A.J. & Blanco, N., 1999. New constraints for the age of Cretaceous compressional deformation in the Andes of northern Chile (Sierra de Moreno, 21°-22° 10´S). Fourth International Symposium on Andean Geodynamics, IRD: 407-410

Mégard, F., 1987. Cordilleran Andes and marginal Andes: a review of Andean geology of the Arica elbow (18oS), In Monger J.W.H., Francheteau, J. (ed.) Circum-Pacific orogenic belts and evolution of the Pacific basin, American Geophysical Union Geodynamic Series 18: 71-95

Mpodozis, C. & Allmendinger, R.W., 1993. Extensional tectonics, Cretaceous Andes, northern Chile (27oS), Geological Society of America Bulletin 105: 1462-1477

Okay, A.I., Tansel, I. & Tüysüz, O., 2001. Obduction, subduction and collision as reflected in the Upper Cretaceous-Lower Eocene sedimentary record of western Turkey, Geological Magazine 138: 117-142

O´Neill, C., Müller, R.D. & Steinberger, B., 2005. On the uncertainties in hot spot reconstructions and the significance of moving hot spot reference frames, Geochemistry Geophysics Geosystems 6 (4): Q04003, doi: 10.1029/2004GC000784.

Raval, U. & Veeraswamy, K., 2003. India-Madagascar separation: breakup along a pre-existing mobile belt and chipping of the craton, Gondwana Research 3: 467-485.

Stampfli, G.M. & Kozur, H.W., 2006. Europe from Variscan to the Alpine cycles, In Gee. D.G., Stephenson, R.A., (ed.) European Lithosphere Dynamics, Geological Society of London Memoir 32: 57-82

Tomlinson, A.J.; Martin, M.W.; Blanco, N. & Pérez de Arce, C. 2001. U-Pb and K-Ar geochronology from the Cerro Empexa Formation, 1st and 2nd Regions, Precordillera, Northern Chile. In Proceedings of South American Symposium on Isotope Geology, 3: 632-635

Storey, M., Mahoney, J.J, Saunders, A.D., Duncan, R.A., Kelley, S.P. & Coffin, M.F., 1995. Timing of hot spot-related volcanism and the breakup of Madagascar and India, Science 267: 852-855.

Torsvik, T.H., Tucker, R.D., Ashwal, L.D., Jamtveit, B., Vidyadharan, K.T. & Venkataramana, P., 2000. Late Cretaceous India-Madagascar fit and timing of break-up related magmatism, Terra Nova 12: 220-224.

Ziegler, P.A., 1988. Evolution of the Artic–North Atlantic and the Western Tethys, American Association Petroleum Geologist Memoir 43: 1-198

Ziegler, P.A., 2005. Europe/Permian to Recent evolution, In Selley. R.C., Cocks, L.R., Plimer, I.R. (ed.), Encyclopedia of Geology 5th. 2: 102-124

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 513-516

513

Linkage between Neogene arc expansion and contractional reactivation of a Cretaceous fold-and-thrust belt (southern Central Andes, 36º-37ºS)

Mauro G. Spagnuolo, Andrés Folguera, & Victor A. Ramos

Laboratorio de Tectónica Andina, Universidad de Buenos Aires and CONICET, Buenos Aires, Argentina

(mauro@gl.fcen.uba.ar, andresfolguera2@yahoo.com.ar, andes@gl.fcen.uba.ar)

KEYWORDS : magmatism, arc expansion, orogenic front

Introduction

The Andean margin between 34º and 37º south latitude is considered a key area in order to constrain plate

motion dynamics and especially the geometrical relation between the overriding and the subducting plate

through time. Variations in the Wadatti-Benioff zone during the last 20 My strongly affected and controlled

foreland deformation and emplacement of anomalously thick accumulations of arc-related rocks far away from

the trench. Here, we analyze Miocene arc expansion, which may be related to shallowing of the subduction zone

and related to a maximun in uplift and deformation during that time.

The region was affected by intense compression since Late Cretaceous times. Magmatic activity is generally

explained by a progressively intense subduction coupling at the western plate margin and alternations between

normal and flat subduction stages.

Variations in the Wadatti-Benioff zone during the last 20 My at these latitudes strongly affected and controlled

foreland deformation and emplacement of anomalously thick accumulations of arc-related rocks far away from

the trench (Kay et al., 2006). The Late Miocene orogenic front at 36ºS extended as much as 430 km from the

Pacific trench and was associated with arc-related products with mean ages around 11 Ma (Ramos and Folguera,

2005). All this magmatic activity was superimposed to the Cretaceous fault-and-thrust belt that partially

controlled the emplacement of the volcanic products during the Miocene contractional phase. Slab flattening

have produced migration of volcanic arc and the reactivation of the basement structures that deformed the

Miocene deposits.

Pre-Cenozoic history

Since Early-Jurassic to Early Cretaceous, deep marine sediments interfingered with basaltic rocks accumulated

along extensional depocenters associated with negative trench roll-back velocities of the overridden plate

(Ramos, 1999). Eventually, important changes in magmatism and deformation occurred near the end of the

Early Cretaceous due to increase in absolute westward motion of South American plate that started a positive

roll back regime. There are several evidences of contractional deformation from Late Cretaceous to Paleogene

times, such as the distribution of Late Cretaceous synorogenic deposits and magmatic cross cutting relationships

with previous deformed rocks all along the Northern Patagonia and Neuquén Andes (Llambías et al. 1979, Kay

2001, Burns 2002; Orts and Ramos 2006; ZamoraValcarce et al., 2006).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 513-516

514

Cenozoic history

Contractional setting lasted until Oligocene times, when the arc remained nearly stationary north of Cortaderas

Lineament. Later on, during Oligocene to Early Miocene times, the Neuquén Andes were characterized by

generalized extension caused by a negative trench roll back regime (Somoza 1998; Folguera et al. 2003).

Following this period or regional extension dramatic changes occurred during Middle to Late Miocene with the

development of Malargüe fold-and-thrust belt. This time was characterized by contractional deformation and an

eastward expansion of the arc magmatism. This eastward expansion of magmatism occurred at the time of

important crustal shortening and uplift in the Malargüe fold-and-trhust belt between 15 and 8 Ma with a peak

deformation between 10.5 and 8 Ma (Giambiagi et al. 2007).

The Cerro Domuyo and Cerro Mayán case study

In order to study Middle to Late Miocene arc migration and contractional deformation, two areas were studied:

Cerro Domuyo and Cordillera de Mayán. Both are prominent structures cored by Miocene granitoids that were

emplaced in Mesozoic sequences unconformably underlying volcanic products.

Cerro Domuyo

Cerro Domuyo is a high structural dome that exhumes Choiyoi Group around its core. It has been proposed

that it was uplifted in two different phases (Lisjak 2007). The first phase occurred during Cretaceous times by

analogy with Cordillera del Viento, located immediately to the south, which has been uplifted between 80 and

70 Ma. That first uplift was followed by an intense deformation during Late Miocene-Pliocene times. This fact

was confirmed by the tilting between the Charilehue volcanic sequences and the Mesozoic deposits. Palinspastic

restoration of the two sequences shows that major contraction took place in the Miocene times.

It was also found nearby Domuyo center, several structures with high obliquity respect to the active margin.

The Cerro Domuyo shows a quadrangular shape due to basement control and a series of oblique anticlines and

synclines that affected the Charilehue sequences west of the Barrancas river (Fig. 1b).

Cerro Mayán

Cerro Mayán is also a volcanic center found northeast of Cerro Domuyo (Fig. 1b). Its core shows Late

Miocene granites and Mesozoic deposits indicating a high degree of exhumation, similarly to the Domuyo case.

In Cerro Mayán, we have measured 30º dips for Miocene volcanic sequences and 50º dips in the Cretaceous

deposits. This fact indicates that the main phase of deformation was again produced during Miocene times. The

main structure is basically a wide anticline with an east-vergence and it is surrounded by oblique structures

(Fig. 1).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 513-516

515

Figure 1. a) Relation between Oligocene and the Miocene arcs. b) Distribution of the Miocene arc rocks and main oblique basement structures.

Conclusions

These facts imply that major deformation occurred during Miocene times in the foreland, similarly to the

deformation observed in the Main Cordillera at these latitudes by Giambiagi et al. (2007). Moreover, highly

oblique structures, with east-northeast to west-northwest trends, have been observed and interpreted as a series

of major uplift structures affecting the Miocene volcanic rocks. That oblique deformation must be controlled by

basement structures, which are also linked to tectonic inversion of normal faults in the area. The arc migration

may explain many geological features such as the irregular broken foreland basement uplifts and the consequent

orogenic collapse as a function of the arc-retreat to the trench.

The arc may have migrated in a series of discrete branches of the order of a few tens of kilometres that

controlled the emplacement of basement-cored structures such as the Cerro Domuyo and Cerro Mayán in the

foreland area, as a result of development of fragil-ductil discontinuities. Those transtensionally-reactivated

structures were the conduits by which arc melts reached the surface. Subsequent intraplate volcanism with OIB

characteristics were extensionaly extruded, process that continued until historical times.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 513-516

516

References

Burns, W. M., 2002. —Tectonic and depositional evolution of the Tertiary Cura Mallín Basin in the southern Andes ( 36.5º to 38ºS lat.) Ph. D Thesis. Cornell University, Ithaca, New York, 218 p.

Folguera, A., Ramos, V. A., Melnick, D., 2003. Recurrencia en el desarrollo de cuencas de intraarco. Colapso de estructuras orogénicas. Cordillera Neuquina (37º30’). Revista de la Asociación Geológica Argentina, 58: 3-19.

Giambiagi L., Bechis, F., García, V., Clark, A., 2007 – “Temporal and spatial relationships of thick- and thin-skinned deformation: a case study from the Malargüe fold and thrust belt, Southern Central Andes”. In Sempere, T., Folguera, A., Gerbault, M. (ed.): Tectonophysics Special Issue-New insights into Andean evolution ISAG 2005. In Press.

Kay, S.M. 2001. —Tertiary to Recent magmatism and tectonics of the Neuquén Basin between 36º05’ and 38ºS latitude. Buenos Aires, Internal report to repsol YPF, 125p.

Kay, S.M., Burns, M., Copeland, P., 2006 – “Upper Cretaceous to Holocene Magmatism over the Neuquén basin: Evidence for transient shallowing of the subduction zone under the Neuquén Andes (36°S to 38°S latitude)”. In Kay, S.M. and Ramos, V.A. (eds.). Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén basin (35º-39ºS)-Geological Society of America, Special Paper, 407: 19-60.

Lisjak, M., 2007— Geología, estratigrafía y estructura de las nacientes del arroyo Manchana Cobunco. Area del cerro Domuyo, Neuquén. Trabajo final de licenciatura . Buenos Aires.100 p.

Llambias, E., Danderfer, J., Palacios, M., Broggioni, N., 1979. Las rocas ígneas Cenozoicas del volcán Domuyo y áreas adyacentes: 7th Congreso Geológico Argentino (Neuquén, 1978) Actas 2:569-584.

Orts and Ramos, V. A. 2006. Evidence of Middle to Late Cretaceous compressive deformation in the high Andes of Mendoza. Backbone of the Americas abstract. Argentina 5-35.

Ramos, V. A., 1999. Plate tectonic setting of the Andean Cordillera: Episodes 22: 183-190 Ramos, V., Folguera, A., 2005 – “Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc

and foreland deformation”. In Veiga, G., Spalletti, L., Howell J. and Schwarz E. (eds.). The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics-Geological Society of London, Special Publication, 252: 15-35.

Somoza, R. 1998. Updated Nazca (Farallón) –South America relative motion during the last 49 m.y.; implications for mountain building in the Central Andean region. Journal of South American Earth Sciences, 11: 211-215.

Zamora-Valcarce, G., Zapata, T., Del Pino, D., Ansa, A., 2006 – “Structural evolution of the Agrio fold and thrust belt“. In Kay, S.M., Ramos, V.A. (eds.). Evolution of an Andean margin: a tectonic and magmatic view from the Andes to the Neuquén basin (35°- 39°s lat.). Geological Society of America, Special Paper, 407: 125-145.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 517-520

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Tectonic response of the central Chilean margin (35°-38°S) to the collision and subduction of heterogeneous oceanic crust: A thermochronological study

R. Spikings1, M. Dungan

1, J. Foeken

2, A. Carter

3, & L. Page

4

1 Department of Mineralogy, University of Geneva, 13 Rue des Maraîchers, Geneva 1205, Switzerland

(spikings@terre.unige.ch, dungan@terre.unige.ch) 2 Scottish Universities Environmental Research Centre, Rankine Avenue, Scottish Enterprise Technology Park,

East Kilbride, G75 0QF, Scotland (jf99w@udcf.gla.ac.uk) 3 School of Earth Sciences, Birkbeck College, Malet Street, London WC1E 7HX, England (a.carter@ucl.ac.uk)

4 Dept of Geology, GeoBiosphere Science Centre, Lund University, Sölvegatan 12, 223 62 Lund, Sweden

(Laurence.Page@geol.lu.se)

KEYWORDS : thermochronology, tectonics, exhumation, Chile, Miocene

Hotspot activity and ocean plate rearrangements since at least 25 Ma have formed structural, thickness and

density heterogeneities in the approaching and subducting oceanic crust offshore central Chile. Numerous

studies relate spatial and temporal variations in slab-dip, structure and thickness of the upper plate, seismicity,

and arc geochemistry to the location of heterogeneities in the subducting oceanic crust, and the distribution of

relict continental rift structures. However, there is a paucity of studies which attempt to quantify the timing,

spatial extent and amount of Neogene vertical displacement experienced by the crust in the flat-slab region of

central Chile, which currently hosts the Juan Fernandez Ridge at ~33°S.

We present the results of 40Ar-39Ar (hornblende; biotite), fission-track (FT; apatite), and (U-Th)/He (zircon;

apatite) analyses of Miocene granitoids, which crop-out along a north-south oriented traverse along the western

slope of the Principal Andean Cordillera of Chile, between 35-38°S. Each of the minerals studied provides

thermal history information over a specific temperature range, and when the results are integrated they yield

temperature-time paths. These thermal histories permit an assessment of the timing, magnitude and duration of

thermal events, which are subsequently used to quantify the Neogene exhumation history of the Principal

Andean Cordillera in central Chile.

Driving forces for cooling and exhumation

18-15 Ma

High cooling rates during 18-15 Ma (Figure 1) within the Principal Cordillera between 35–38°S were partly

driven by exhumation, although we can not resolve the amount of cooling caused by sub-solidus, thermal

relaxation. As along-strike variations in exhumation depth during 18–15 Ma can not be approximated accurately,

the potential for identification of responsible driving forces is limited. The along-strike extent of exhumation

during 18–15 was at least 400km, and may even extend continuously to the Puna and Altiplano plateaux where

crustal shortening during the Early and Middle Miocene is evident (Victor et al. 2004), giving rise to an along-

strike distance of ~1500km.

Kay et al. (2005) report a geochemical change in central Chile from low-K tholeiites to calc-alkaline dacites

with more evolved isotopic compositions. This is consistent with crustal thickening during inversion tectonics

(Charrier et al. 2002) in the late Early Miocene. Therefore, exhumation along the central Chilean margin during

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 517-520

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18-15 Ma was synchronous with crustal thickening, suggesting they are the result of a common driving force.

The thickness of the crust increases abruptly northwards at 36°S, and the thickened crust persists towards the

Puna and Altiplano plateaux (Tassara et al. 2006). Therefore, assuming that i) exhumation was driven by periods

of crustal thickening, and ii) large-scale thickening of the crust over an along-strike distance of >1500 km was

occurring in the Early and Middle Miocene, we suggest that rock uplift of the Principal Cordillera in central

Chile during 18-15 Ma was driven by a large-scale process, and not by the subduction of local heterogeneities,

such as the Juan Fernandez Ridge, which was subducting beneath the South American Plate at ~20°S, 18 million

years ago (Yáñez et al., 2002).

10-0 Ma

The onset of rapid exhumation during the Late Miocene follows a younging trend from ~10 Ma in the Sierras

Pampeanas (~30°S; Coughlin et al. 1998), ~8.4 Ma at 34.5°S in the Principal Cordillera (Kurtz et al. 1997), to

~7.5 Ma south of 35°S. Furthermore, the depths of exhumation since ~7.5 Ma increase from 1km at 38°S to

~5km at 35°S, with the greatest increase in rates occurring to the north of ~36°S, assuming a constant

geothermal gradient of 40°C/km. Periods of elevated exhumation between 38°S and 36.2°S terminated at ~5 Ma,

at which time the current surface was at temperatures 50°C, implying depths of <1km. However, the Rio Teno

valley (35°S) continued to exhume rapidly until 1–0 Ma.

The spatial coincidence of the younging trend of the onset of Late Miocene exhumation with the Juan-

Fernandez Ridge and flat-slab (Yáñez et al. 2002) strongly suggests a cause and effect relationship (Figure 1).

The flat-slab is located ~200 km north of the study region, and hosts the Juan Fernandez Ridge. Yáñez et al.

(2002) predict that the zone of subduction of the ridge migrated rapidly southwards from ~20°S to ~30°S until

~10 Ma, since when it has migrated more slowly to its current location at 33°S. We suggest that elevated

exhumation rates at 10 Ma in the Sierras Pampeanas, and after 7.5 Ma within the study region, were driven by i)

progressive southward flattening of the slab due to ridge subduction, resulting in increased coupling between the

subducting and upper plates, ii) increased compressive stress caused by collision of the topographically

prominent, and thick volcanic ridge with the upper plate. The study region does not lie directly above the region

of the ridge and the flattened slab, although it is feasible to suggest that the upper plate response to increased

coupling and horizontal stress attenuates beyond the region of flattening, as shown by the latitudinal dispersion

of recorded seismicity (Barientos et al. 2004), and numerical modelling (Yáñez et al. 2002).

Kay et al. (2006) present a model for transient shallowing and steepening of the Nazca Plate between 36-38°S

during 8–5 Ma. This period was synchronous with elevated exhumation rates north of 38°S, in the Principal

Cordillera of Chile. Therefore, it is sensible to suggest that elevated, Late Miocene exhumation rates in the

Principal Cordillera may also be a consequence of Late Miocene flattening of the slab by a mechanism that is

unrelated to the subduction of the Juan Fernandez Ridge, although this does not solely account for the along-

strike exhumation trend in the cordillera between 35-38°S.

Volcanic ridge subduction, rock-uplift, surface uplift and exhumation

Spikings et al. (2001) utilised the FT method to show that the Carnegie Ridge collided with the northern

Andean margin at 15 Ma, resulting in a sudden increase in exhumation rates in the upper plate during

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 517-520

519

~15-14 Ma, and the erosion of ~6 km of crust. Gullier et al. (2001) showed that the Carnegie Ridge is currently

not associated with a flat-slab, which contrasts with the flattened slab that hosts the Juan Fernandez ridge.

Therefore, the tectonic response of the upper plate is not solely a simple function of slab-dip.

Wipf (2006) obtained apatite FT data from Cretaceous and older granitoids along coastal Peru, and were

unable to detect distinct periods of elevated exhumation rates since the Late Miocene, which may have been

driven by subduction of the Nazca Ridge, as it migrated southwards from ~11.5°S to its current location at

~15°S, since its collision at ~11.2 Ma (Hampel et al. 2002). However, Pleistocene marine terraces crop-out at a

maximum elevation of ~800m along coastal Arequipa (Hsu, 1992). Current annual rainfall along the region of

coastal Peru is less than 5cm/yr, and arid conditions persisted along the coast throughout the Cenozoic (Dunai et

al. 2005). We suggest that low exhumation depths along coastal Peru, compared with the upper plate above the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 517-520

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Juan-Fernandez and Carnegie ridges, corroborates the combined effects of the lower erosive power of the

climate in coastal Peru, and the short life-span of the dynamically supported land-surface (~3.5 my; Hsu 1992)

during the SE directed displacement of the Nazca Ridge, relative to the South American Plate.

A comparison of thermochronological and geophysical data from rocks that have been subducted by volcanic

ridges along the western margin of the South American Plate suggests that i) the collision and subduction of

topographically prominent, and thick volcanic ridges with continental plates drives rock uplift in the continental

plate, ii) the amount and spatial extent of exhumation that occurs is strongly dependant on climate induced

erosion, and the life-span of the dynamically supported, topographically prominent crust, and iii) slab-flattening

and plate coupling may play a less important role than horizontal compressive stress originating at the trench

during ridge collision.

References Barrientos, S., Vera, E., Alvarado, P., & Monfret, T. 2004. Crustal seismicity in central Chile. Journal of South American

Earth Sciences 16: 759-768. Charrier, R., Baeza, O., Elgueta, S., Flynn, J.J., Gans, P., Kay, S.M., Muñoz, N., Wyss, A.R., & Zurita, E. 2002. Evidence for

Cenozoic extensional basin development and tectonic inversion south of the flat-slab segment, southern Central Andes, Chile (33°-36° S.L.). Journal of South American Earth Sciences 15: 117-139.

Coughlin, T.J., O’Sullivan, P.B., Kohn, B.P., & Holcombe, R.J. 1998. Apatite fission-track thermochronology of the Sierras Pampeanas, central western Argentina: Implications for the mechanism of plateau uplift in the Andes. Geology 26: 999-1002.

Dunai, J.T., Gabriel, A.G.L., & Juez-Larre, J. 2005. Oligocene-Miocene age of aridity in the Atacama Desert revealed by exposure dating of erosion-sensitive landforms. Geology 33: 321-324.

Gullier, B., Chatelain, J.L., Jaillard, E., Yepes, H., Poupinet, & G., Fels, J.-F. 2001. Seismological evidence on the geometry of the orogenic system in central-northern Ecuador (South America). Geophysical Research Letters 28: 3749–3752.

Hampel, A. 2002. The migration history of the Nazca Ridge along the Peruvian active margin: a re-evaluation. Earth and Planetary Science Letters 203: 665-679.

Hsu, J.T. 1992. Quaternary uplift of the Peruvian coast related to the subduction of the Nazca Ridge: 13.5 to 15.6 degrees south latitude. Quaternary International 15/16: 87-97.

Kay, S.M., & Kurtz, A. 1995. Magmatic and tectonic characterization of the El Teniente region. CODELCO (unpublished report), 180 pp.

Kay, S.M., Godoy, E., & Kurtz, A. 2005. Episodic arc migration, crustal thickening, subduction erosion, and magmatism in the south-central Andes. Geological Society of America Bulletin 117: 67-88.

Kay, S.M., Burns, W.M., Copeland, P., & Mancilla, O. 2006. Upper Cretaceous to Holocene magmatism and evidence for transient Miocene shallowing of the Andean subduction zone under the northern Neuquén Basin. In: Kay, S.M. & Ramos, V.A. (eds) Evolution of an Andean Margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°-39°S lat). Geological Society of America Special Paper 407: 19-59.

Kurtz, A., Kay, S.M., Charrier, R., Farrar, E. 1997. Geochronology of Miocene plutons and exhumation history of the El Teniente region, Central Chile (34°-35°S). Revista Geológica de Chile 24: 75-90.

Spikings, R.A., Winkler, W., Seward, D., & Handler, R. 2001. Along-strike variations in the thermal and tectonic response of the continental Ecuadorian Andes to the collision with heterogeneous oceanic crust. Earth and Planetary Science Letters 186: 57-73.

Somoza, R. 1998. Updated Nazca (Farallón)-South America relative motions during the last 40 my: Implications for mountain building in the central Andean region. Journal of South American Earth Sciences, 8: 17-31.

Tassara, A., Götze, H-J., Schmidt, S., & Hackney, R. 2006. Three-dimensional density model of the Nazca plate and the Andean continental margin. Journal of Geophysical Research 111: doi:10.1029/2005JB003976.

Victor, P., Oncken, O., & Glodny, J. 2004. Uplift of the western Altiplano: Evidence from the Precordillera between 20° and 21°S (northern Chile). Tectonics 23: TC4004, doi:10.1029/2003TC001519.

Wipf, M. 2006. Evolution of the Western Cordillera and Coastal margin of Peru: Evidence from low-temperature Thermochronology and Geomorphology. PhD Thesis, ETH Zürich, Switzerland.

Yáñez, G., Cembrano, J., Pardo, M., Ranero, C., & Selles, D. 2002. The Challenger-Juan Fernández-Maipo major tectonic transition of the Nazca-Andean subduction system at 33-34°S: geodynamic evidence and implications. Journal of South American Earth Sciences 15: 23-38.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 521-523

521

Fluvial responses to regional tectonic and local tectonic evolution of Oxaya anticline in hyper-arid area, Arica (North Chile)

M. Strub1, J. Darrozes

1,*, L. Audin

1,3, E. Maire

1, G. Hérail

1,2, J.-C. Soula

1 , & J. Deramond

1

1 UMR 5563 UPS, Toulouse, France (* corresponding author : darrozes@lmtg.ups-tlse.fr)

2 IRD UR 154, Santiago, Chili

3 IRD UR 154, Lima, Peru

KEYWORDS : tectonic, fluvial network, erosion rate, Chilean forearc

Introduction

The purpose of this study is to analyse the relief and to point out the development of the fluvial network as

result of the combined effects of tectonics, orographic rainfall variation and erosion.

The studied area (yellow dot on world map, Fig. 1) is located in a hyper arid region of the northern Chile, the

Northermost Atacama desert. The Atacama Desert is situated in the South-Western part of the Bolivian Orocline,

which is characterized by a thrust-slated deformation in the Fore Arc region (Garcia et al. 2002).

Figure 1: Structural map of Arica area and location of Oxaya Anticline and AB topographic profile.

Results and discussion

The morphological evolution of fluvial network is analyzed with respect to the development of the tectonic

structures. The study is focused on the growth of the Oxaya fault-propagation anticline in the Arica area. The

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 521-523

522

relief initiates in the forelimb of the fold and progresses backward. As a result, the transverse trunk river was

gradually diverted and then captured (see AA’, BB’ and CC’ on fig.2) by a fold-parallel river formed in the

backlimb.

Figure 2: Fluvial network evolution correlated to Oxaya anticline growing, note the southward capture of upper streams due to the backward development of the fold.

Figure 3: Map of Incised Volume performs by the BTH method

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 521-523

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In order to correlate incision rates with regional tectonic, the incised volumes (Fig. 3) in the deep valley was

computed by subtraction of a valley filled DEM and the initial DEM (spatial resolution 50 m, vertical

resolution 15 m) using the Black Top Hat (BTH) method developed by Rodriguez et al. (2002).

Although somewhat higher the long-term maximum average incision rates are in good agreement with those

estimated in similar arid conditions (Riquelme, 2003; Alpers and Brimhall, 1988 and Riquelme et al. 2008). In

addition to the structural growth, the profile of the trunk rivers emphasizes the role of orographic rainfall effect

in their upstream reach. The higher incision rate values we obtained may be explained by the greater discharge of

strongly uplifted area of the Bolivian Orocline. These main summits, which exceeds the 4000 m asl, collect

humid tropical fluxes and store important water reserve in the form of snow or of ice for highest summits. These

reserves make it possible to have important water discharge, throughout the year, for the rivers which take their

source in the main summit reserves.

References Alpers, C.N.; Brimhall, G.H. 1988 - Middle Miocene climatic change in the Atacama Desert, northern Chile: Evidence from

supergene mineralization at La Escondida. - G.S.A.B., 100: 1640-1656. Garcia M.; Hérail G. (2002) -Evolution oligo-néogène de l’altiplano occidental- ; Thèse 3ème cycle, Université J. Fourier,

Grenoble. Riquelme R. (2003), - Evolution géomorphologique néogène des Andes Centrales -; Thèse 3ème cycle, Université Paul

Sabatier, Toulouse. Riquelme R., Darrozes J., Maire E., Hérail G., J. C. Soula, (2008), - Long-term denudation rates from the Central Andes

(Chile) estimated from a Digital Elevation Model using the Black Top Hat function and Inverse Distance Weighting: implications for the Neogene climate of the Atacama Desert -, Rev. Geol. . Chil., 35 (1): 105-121.

Rodriguez F., Maire E., Courjault-Rade P. (2002) - The Black Top Hat function applied to a DEM: a tool to estimate recent incision in a mountainous watershed -; G.R.L., 29, 0, 10.1029.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 524-525

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Tithonian to Aptian volcanism in the central Patagonian Cordillera, Aysén, Chile (45°-46°30’S): U-Pb shrimp new data

Manuel Suárez1, Rita De La Cruz

1, & Marc Fanning

2

1 Servicio Nacional de Geología y Minería, Av. Sta. María 0104, Santiago, Chile (msuarez@sernageomin.cl,

rcruz@sernageomin.cl) 2

RISE, ANU, Canberra, ACT 0200, Australia (mark.fanning@anu.edu.au)

Ten new U-Pb SHRIMP dates, ranging between 150 and 118 Ma, obtained from volcanic rocks from the Aysén

Region of Chile, in central Patagonian Cordillera (45-46ºS) are reported herein. Five are from the subduction-

related Ibáñez Formation, 3 from overlying dinosaur-bearing beds interpreted as a deltaic association of the

Toqui Formation and two from the Divisadero Formation (Time Scale of Gradstein et al., 2004). The Ibáñez

Formation is the youngest association that has been included in the felsic large igneous province known as the

Chon Aike Province emplaced during rifting that preceded the dismembering of Gondwana. It is mainly formed

by rhyolitic and dacitic ignimbrites, with subordinate andesitic and basaltic lava flows and sedimentary rocks,

including volcaniclastic sandstones, black shales, matrix-supported volcanic breccias and calcareous silicified

laminites. Marine beds of the overlying Toqui Formation, exposed north of the studied area, include Tithonian

ammonites and, in turn, marine intercalations in the Ibáñez Formation with Berriasian ammonites implied the

continuation of volcanism during the earliest Cretaceous (Covacevich et al., 1994; De La Cruz et al., 1996;

Suárez et al., 2007). Seven earlier zircon U-Pb SHRIMP analyses from the Ibáñez Formation rendered ages of

ca. 150 and 136 Ma indicative of a Tithonian and early Hauterivian ages (Pankhurst et al., 2000, 2003; Suárez et

al., in press), supporting earlier paleontologic and K-Ar information. Previous Early Berriasian ammonite

bearing beds are overlain by a 137 Ma ignimbrite (late Valanginian), which suggests the presence of a hiatus (or

condensed section).

Three new concordant SHRIMP ages of ca. 147 Ma were obtained from volcanic rocks intercalated in the

dinosaur-bearing beds of the Toqui Formation (De La Cruz et al., in press). Our new data corroborates that

volcanism was active between the Tithonian and early Hauterivian coeval with marine sedimentation of the

Aysén Basin. Volcanism reappeared in the region at approximately 120 Ma, in the Aptian, with the eruption of

surtseyan tuff cones of the Baño Nuevo Volcanic Complex, during the waning stages of the Aysén Basin

(Demant et al., 2007; Suárez et al., 2007, in prep.).

Two new Aptian SHRIMP ages from ignimbrites of the Divisadero Formation, one of 118,5±0,8 Ma from an

exposure approximately 25 km S of Chile Chico, and the other of 116,7±0,7 Ma (Aptian), from an area near the

city of Coyhaique, north of the studied area, are reported herein. They are concordant with SHRIMP ages of

ca. 118 and 116 Ma from ignimbrites of the same formation (Pankhurst et al., 2003).

Twenty two biotite K-Ar dates from the Ibáñez Formation, some of which have been published in abstracts of

geologic congresses or in geologic sheets, range between 159 and 132 Ma (Suárez and De La Cruz, 1997a,b;

Suárez et al., 1997; De La Cruz et al., 2003, 2004; De La Cruz and Suárez, 2006). Where SHRIMP and K-Ar

ages were obtained from the same bed, in general a discrepancy exists, indicative of loss or excess Ar. Recently,

De La Cruz and Suárez (in prep.) obtained latest Jurassic and earliest Cretaceous K-Ar and Ar/Ar alteration ages

in rocks of the Ibáñez Formation exposed in a mining district west of Chile Chico.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 524-525

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Acknowledgements We thank Leonardo Zuñiga for collaboration in the field. The work was financed by project FONDECYT 1030160 and the Servicio Nacional de Geología y Minería. References Covacevich, V.; De La Cruz, R.; Suárez, M. 1994. Primer Hallazgo de fauna del Berriasiano Inferior Neocomiano en la Formación Ibáñez,

XI Región, Aisén. In Congreso Geológico Chileno, No. 7, 1: 425-429. Concepción. De La Cruz, R.; Suárez, M.; Covacevich, V.; Quiroz, D. 1996. Estratigrafía de la zona de Palena y Futaleufú (43º15'-43º45' Latitude S), X

Región, Chile. In Congreso Geológico Argentino, No. 13 y Congreso de Exploración de Hidrocarburos, No. 3, Actas 1: 417-424. De La Cruz, R.; Suárez, M.; Belmar, M.; Quiroz, D.; Bell, M. 2003. Geología del área Coihaique-Balmaceda, Región de Aisén del General

Carlos Ibáñez del Campo. Servicio Nacional de Geología y Minería, Serie Geología Básica, Carta Geológica de Chile, No. 80, escala 1:100.000.

De La Cruz, R.; Welkner, D.; Suárez, M.; Quiroz, D. 2004. Geología del área oriental de las hojas Cochrane y Villa O’Higgins, Región Aisén del General Carlos Ibáñez del Campo. Servicio Nacional de Geología y Minería. Carta Geológica de Chile, Serie Geología Básica, No. 85, escala 1:250.000.

De La Cruz, R.; Suárez, M. 2006. Geología del área Puerto Guadal-Puerto Sánchez, Región Aisén del General Carlos Ibáñez del Campo. Servicio Nacional de Geología y Minería. Carta Geológica de Chile, Serie Geología Básica, No. 95, escala 1:100.000.

De La Cruz, R.; Salgado, L.; Suárez, M.; Fernández, M.; Gasparini, Z.; Palma-Heldt, S. In press. First Late Jurassic Dinosaur Bones from Chile. Journal of vertebrate Paleontology.

Demant, A.; Suárez, M.; De La Cruz, R. 2007 Lower Cretaceous surtseyan volcanoes in the eastern central Patagonian Cordillera (45°15’-45º40’S): the Baño Nuevo volcanic complex. Geosur 2007: 51 p.

Gradstein, F.M.; Ogg, J.G.; Smith, A.G.; Cooper, R.A.; Sadler, P.M.; Hinnov, L.A.; Villeneuve, M; McArthur, Howarth, R.J.; Agterberg, F.P.; Robb, L.J.; Knoll, A.H.; Plumb, K.A.; Shields, G.A.; Strauss, H.; Veizer, J.; Bleeker, W; Shergold, J.H.; Melchin, M.J.; House, M.R.; Davydov, V.; Wardlaw, B.R.; Luterbacher, H.P.; Brinkhuis, H.; Hooker, J.J.; Monechi, S.; Powell, J.; Röhl, U.; Sanfilippo, A.; Schmitz, B.; Lourens, L.; Hilgen, F.; Shackleton, N.J.; Laskar, J.; Wilson, D.; Gibbard, P.; van Kolfschoten, T. 2004. A geologic time scale 2004. Cambridge University Press:500 p.

Pankhurst, R.J.; Riley, T.R.; Fanning, C.M.; Kelley, S.P. 2000. Episodic silicic volcanism in Patagonia and the Antarctic Peninsula: chronology of magmatism associated with the break-up of Gondwana. Journal of Petrology, 41: 605-625.

Pankhurst, R.J.; Hervé, F.; Fanning, M.; Suárez, M. 2003. Coeval plutonic and volcanic activity in the Patagonian Andes: the Patagonian Batholith and the Ibáñez and Divisadero Formations, Aisén, southern Chile. In Congreso Geológico Chileno No. 10. Concepción.

Suárez, M.; De La Cruz, R. 1997a. Edades del Grupo Ibáñez en la parte oriental del Lago General Carrera (46º-47º LS), Aysén, Chile. In Congreso Geológico Chileno No. 8, Actas 2: 1548-1551. Antofagasta, Chile.

Suárez, M; De La Cruz, R. 1997b. Cronología magmática de Aysén Sur, Chile (45º-48º30’ Latitud Sur) In Congreso Geológico Chileno No. 8, Actas 2: 1543-1547. Antofagasta, Chile.

Suárez, M.; De La Cruz, R.; Bell, M. 2007. Geología del Área Ñireguao-Baño Nuevo. Servicio Nacional de Geología y Minería. Carta geológica de Chile, Serie Geología Básica, No. 108, escala 1:100.000. Santiago, Chile.

Suárez, M.; De La Cruz, R. 1997a. Cronología magmática de Aysén Sur, Chile (45º-48º30' Latitud Sur). In Congreso Geológico Chileno, No. 8, Actas 2:1543-1547.

Suárez, M.; De La Cruz, R. 1997b. Edades K-Ar del Grupo Ibáñez en la parte oriental del Lake General Carrera (46º-47º LS). Aysén, Chile. In Congreso Geológico Chileno No. 8, Actas 2: 1548-1551.

Suárez, M.; Márquez, M.; De La Cruz, R. 1997. Nuevas edades K-Ar del Complejo El Quemado a los 47º13'-47º22' Latitud Sur. In Congreso Geológico Chileno No. 8, Actas 2: 1552-1555.

Suárez, M.; De La Cruz, R.; Aguirre-Urreta, B.; Fanning, C.M. In press. Relationship between volcanism and marine sedimentation in northern Austral (Aysén) Basin, central Patagonia: Stratigraphic, U-Pb SHRIMP and paleontologic evidence. Journal of South American Earth Sciences.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 526-529

526

Anatomy of the Andean forearc controlling short-term interplate seismogenesis and long-term Cordilleran orogenesis

Andrés Tassara1, Ron Hackney

2, & Denis Legrand

3

1 Departamento de Ciencias de la Tierra, Universidad de Concepción, Casilla 160-C, Concepción, Chile

(andres@udec.cl) 2 Petroleum and Marine Division, Geosciences Australia, Canberra ACT 2601, Australia

3 Departamento de Geofísica, Universidad de Chile, Blanco Encalada 2002, Santiago, Chile

KEYWORDS : Andean margin, forearc structure, seismogenesis, orogenesis

Introduction

The convergent Andean margin of western South America is characterized by two levels of along-strike

segmentation (fig. 1). These levels are manifested at different time-scales and related to apparently different

geodynamic processes: a short-term (102-103 yr), regional-scale (102-103 km) segmentation of the seismically

coupled interplate contact, which is defined by segments rupturing the megathrust fault during “characteristic”

earthquakes of large magnitude (M>7.5); and a long-term (106-107 yr), continental-scale (103-104 km)

morphostructural segmentation of the entire margin, which is expressed in along-strike variations of the

morphology and topographic volume of the Andes. Some authors have recognized that the processes resulting in

these two levels of segmentation should be related to a common phenomenon; the mechanical coupling between

the subducted slab and the overriding forearc (e.g. Lamb and Davis, 2003; Iaffaldano and Bunge, 2008). Plate

coupling makes convergence to be translated into a main component of elastic strain that is stored in the strong

forearc and suddenly released during megathrust earthquakes, and another comparatively minor component of

permanent plastic strain that is slowly accumulated along the weak orogenic region to form the Cordillera. Here

we suggest that a key control on plate coupling, and hence on short- and long-term processes acting along the

margin, is the 3D anatomy of the upper plate inherited from the geological configuration of the margin and

producing significant along-strike variations of stress and strength along the plate interface.

Our approach considers the computation of two parameters characterizing the mechanical coupling along the

interplate contact. By one side, we use the digital results of a 3D density model of the Andean margin (Tassara et

al., 2006) to calculate the vertical stress produced by the weight of the forearc column on the subducting slab.

This vertical stress is the main component of the normal stress on the slab, which along with friction, controls

the shear stress to be surpassed in order to generate an earthquake. By the other side, we applied a wavelet

formulation (Kirby and Swain, 2008) of the classical spectral isostatic analysis to invert topography/bathymetry

and satellite gravity data into flexural rigidity. This parameter is a measure of the integrated mechanical strength

of the lithosphere, which in the context of the slab-forearc system likely depends on the strength along the plate

interface and hence on frictional properties of the megathrust (Hackney and Tassara, 2008).

Vertical stress anomaly on the megathrust

The 3D density model of Tassara et al. (2006) integrates several sources of geophysical data to produce a

continental-scale and digital representation of the 3D geometry for several density discontinuities below the

Andean margin between 5º and 45ºS.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 526-529

527

Figure 1. (a) Depth below see level to the Intra-Crustal density Discontinuity (ICD) that separates light upper crust from dense lower crust in the 3D model of Tassara et al. (2006); cold/hot colors indicate a mafic-/felsic- dominated crustal composition. (b) Vertical Stress Anomaly ( v) on the plate interface computed considering the crustal density structure given by (a); blue/red colors indicate deficit/excess stress on the interplate contact compared to a forearc of averaged crustal density structure. (c) Flexural Rigidity (D) inverted from topography and gravity data using a wavelet formulation of spectral isostatic analysis; cold/hot colors indicate high/low strength of the plate interface (i.e. high/low plate coupling). The right panel shows segments for: short-term, regional-scale seismogenic segmentation of the forearc as defined by the rupture length of the last large subduction earthquake and commonly coinciding with rupture of historical events (tj Trujillo 1996, lm Lima 1940, pc Pisco 2007, nz Nazca 1942, aq Arequipa 2001, iq Iquique 1877, at Antofagasta 1995, cp Copiapo 1922, ls La Serena 1943, vp Valparaiso 1985, cn Constitución 1928, va Valdivia 1960); and long-term, continental-scale orogenic segmentation defined by major morphostructural units along the orogen (PeC Peruvian Cordillera, AP Altiplano-Puna Plateau, FC Frontal Cordillera, PpC Principal Cordillera, PgC Patagonian Cordillera). Black points are earthquakes from NEIC catalog, 3<M<7, years 1973-2007, depth < 65 km and occurring 5 km around the slab model of Tassara et al. (2006). Yellow and red dots are earthquakes 7<M<8 and M>8, respectively, from the “Significant Historical Worldwide Seismicity” catalog of NOAA. White start and red contours are the epicenter and slip distribution of the giant M9.5 Valdivia 1960 earthquake.

For instance, it contains the geometry of the top slab surface, the continental Moho and an Intra-Crustal

Discontinuity (ICD) separating light upper crust (density = 2.7 g/cm3) from dense lower crust ( = 3.1 g/cm3).

The depth to the ICD (fig. 1a) is a proxy to lateral density variations that, for the thermodynamic conditions of

the forearc, are mostly due to spatial changes on the lithological configuration of the crust; shallow/deep ICD

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 526-529

528

represents a mafic-/felsic- dominated crystalline crust. Using the slab, Moho and ICD geometries of the 3D

model and appropriate densities, we computed the vertical stress on the interplate megathrust as the weight of the

forearc column on top of the subducted slab between the trench and the 60 km slab-depth contour (~ downdip

limit of the seismogenic zone). The vertical stress is dominated by the thickening of the forearc column from the

trench landward. In order to isolate the effect of crustal density variations, we calculated a Vertical Stress

Anomaly ( v) by subtracting the vertical stress from the one produced by a forearc of homogeneous density

averaging 2.9 g/cm3. The v map is shown in figure 1b.

Flexural rigidity of the slab-forearc system

The method of Kirby and Swain (2008) allows the calculation of continuous grids of flexural rigidity (D) by

considering the statistical coherence in the spectral domain between wavelets transforms of topography and

Bouguer anomaly at different wavelengths around each grid node. At plate interiors, low D implies local, Airy-

type compensation of loads and hence a very weak lithosphere, while high D means a regional compensation by

the deflection of a rigid, strong and thick lithosphere. Thus, D is a measure of the lithospheric strength that has

been shown to depend mostly on the compositional and thermal structure of tectonic plates. Along ocean-

continent convergent margins, Hackney and Tassara (2008) have proposed that the main factor controlling

spatial D variations is the strength of the plate interface that regulates the mechanical coupling between both

plates and the transmission of elastic stress one to each other. The strength of the interface should depend on the

frictional properties of the subduction channel that are likely dominated by the kind and amount of subducted

material, the amount of water and the physical properties of the upper and lower plates. In figure 1c we show the

flexural rigidity map of our study area, a result which is part of the work of Hackney and Tassara (2008).

Results and Discussion

Lateral variations of stress (fig. 1b) and strength (fig. 1c) along-strike the seismically coupled Nazca-South

America interplate contact are not only strongly correlated one to each other (low/high v regions coincide with

low/high D regions), but also to both levels of segmentation. At regional-scale, the limit of seismic segments

seems to coincide with remarkable changes in v, and less pronounced changes in D. It is very significant that

earthquakes of all magnitude have a strong tendency to occur in regions where v > 0 MPa and D > 1021 Nm, a

result that could have important consequences for understanding earthquake-generating processes along

subduction zones. Of particular interest is the situation of the giant (Mw9.5) Valdivia 1960 earthquake, which

nucleated in a high- v and high-D region and then propagated into a large zone of low- v and low-D. This

suggest that once the megathrust fault breaks after cumulated sufficient elastic strain in a region of high shear

stress controlled by high vertical stress and high interplate friction, the subsequent earthquake can grow to giant

dimensions if the rupture front eventually propagates into a sufficiently large region of the plate interface

characterized by low stress and strength. Flexural rigidity maps computed along other subduction zones

worldwide are being interpreted in the context of seismogenesis and would help to further refine this hypothesis

(Hackney and Tassara, 2008).

At continental-scale we also observe a remarkable correlation between the along-strike segmentation of the

Andean orogen and our proxy for stress and strength of the plate interface. The forearc region in front of the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 526-529

529

Altiplano-Puna Plateau (4000 m high, 400 km width) is characterized by very high v and D values from the

trench to the downdip limit of the seismogenic zone. The high shear stress of the interplate contact and high

integrated strength of the forearc suggested by these results are in agreement with the observations of other

authors (Lamb and Davis, 2003; Yáñez and Cembrano, 2004; Tassara, 2005; Lamb, 2006; Iaffaldano and Bunge,

2008) and reinforce the idea that high stress/strength along the interplate contact is a necessary condition for the

dynamic support of the large buoyancy stresses associated with the huge topographic mass and crustal thickness

of the Altiplano-Puna. North and southward of the Plateau, the Peruvian and Frontal Cordilleras have similar

elevations than the plateau, but they form narrower orogenic chains. This is notable correlated with a change in

the stress and strength distribution along the forearc, whose outer part north of 15ºS and south of 29ºS is

characterized by v < 0 and relatively lower D. This could imply a less pronounced capability of the slab-

forearc system to support orogenic topography compared to the Altiplano-Puna Segment. The forearc region in

front of the low-elevation (<2000 m) Patagonian Cordillera south of 38ºS has the lowest v and D values of the

entire Andean margin, suggesting that a quite weak and low-shear stress plate interface is incapable of

supporting any significant orogenic topography there.

Conclusions

Along-strike variations on plate coupling exert a primary control on short-term seismogenic processes causing

a characteristic segmentation of the megathrust fault and on the larger scale and long-term segmentation of the

Cordillera. As shown by other authors (Song and Simons, 2003; Iaffaldano and Bunge, 2008), plate coupling

seems to be manifested in the forearc gravity field, which is the primary observable behind the calculations done

here for v and D. What is significant of our work is that the forward modelling of the Bouguer anomaly

performed by Tassara et al. (2006) explains the variations of the gravity field along the forearc as a consequence

of lateral density variations (fig 1a) that are mostly due to changes of the bulk crustal composition and are

strongly correlated with spatial changes of the old geological configuration observed at the surface. Our main

conclusion is, therefore, that it is the geologically-derived anatomy of the forearc what finally controls the shear

stress level and the mechanical strength of the interplate contact, producing significant along-strike variations of

plate coupling and influencing short-term seismogenesis and long-term orogenesis of the Andean margin.

References Iaffaldano, G. & Bunge, H-P. 2008. Strong plate coupling along the Nazca/South America convergent margin. In press in

Geology. Hackney, R, & Tassara, A. 2008. Subduction zone flexural rigidity and giant earthquake rupture. Submitted to Nature. Kirby, J. & Swain, C. 2008. An accuracy assessment of the fan wavelet coherence method for elastic thickness estimation.

Geochem. Geophys. Geosyst., 9, Q03022, doi:10.1029/2007GC001773. Lamb, S. (2006), Shear stresses on megathrusts: Implications for mountain building behind subduction zones, J. Geophys.

Res., 111, B07401, doi:10.1029/2005JB003916 Lamb, S. & Davis, P. 2003. Cenozoic climate change as a possible cause for the rise of the Andes, Nature, 425, 792– 797. Song, T. & Simons, M. 2003. Large trench-parallel gravity variations predict seismogenic behavior in subduction zones.

Sciences, 301, 630-633. Tassara, A. (2005), Interaction between the Nazca and South American plates and formation of the Altiplano-Puna plateau:

Review of a flexural analysis along the Andean margin (15º–34ºS), Tectonophysics, 399, 39–57. Tassara, A., Gotze, H.-J., Schmidt, S. & Hackney, R. 2006. Three-dimensional density model of the Nazca plate and the

Andean continental margin, J. Geophys. Res., 111, B09404, doi:10.1029/2005JB003976. Yáñez, G., & J. Cembrano. 2004. Role of viscous plate coupling in the late Tertiary Andean tectonics, J. Geophys. Res., 109,

B02407, doi:10.1029/2003JB002494.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 530-533

530

A geochemical survey of geothermal resources in the Tarapacá and Antofagasta regions (northern Chile)

F. Tassi1, F. Aguilera

2, O. Vaselli

1,3, & E. Medina

2

1 Department of Earth Sciences, University of Florence, Via La Pira 4, 50121, Florence, Italy (franco.tassi@unifi.it)

2 Universidad Católica del Norte, Av. Angamos 0610, Antofagasta, Chile (faguilera@ucn.cl, emedina@ucn.cl)

3 CNR-IGG Institute of Geosciences and Earth Resources, Via La Pira 4, 50121, Florence, Italy

(orlando.vaselli@unifi.it)

KEYWORDS : geothermal field, Northern Chile, hydrothermal system, geothermometry, fluid chemistry

Introduction

The Tarapacá and Antofagasta regions (northern Chile) host part of the Central Volcanic Zone (CVZ) of the

Andes Range that is characterized by an intense volcanism likely triggered by subduction thrusting the oceanic

Nazca Plate beneath the South America Plate (Stern, 2004). The main tectonic features of the CVZ are a series of

NNW–SSE oriented grabens (Francis and Rundle, 1976; Lahsen, 1976), where several geothermal fields, not

necessarily associated with active volcanism, are located. During the 70’s, preliminary geochemical

investigations on thermal fluids from several hydrothermal systems of this extended area, i.e. Surire, Puchuldiza-

Tuja and El Tatio, were performed (e.g. Cusicanqui et al., 1975; Lahsen, 1975; 1976; Lahsen and Trujillo, 1976;

Urzua et al., 2002), in response to the increasing Chilean demand for energy from alternative sources. At El

Tatio, in the period 1968-1974, 13 wells were drilled to (600 to 1,820 m in depth), where up to 256 °C was

measured. The available capacity from three successful production wells was estimated around 15 MW.

However, in spite of the promising geothermal potential, the chemical and isotopic features of the majority of the

hydrothermal discharges of the northern Chilean regions are still almost unknown. In this work we present the

results of a geochemical survey, carried out from October 2002 to May 2007, on the thermal springs from: i)

Surire (250 km NE the city of Iquique), Puchuldiza-Tuja (200 km NE the city of Iquique), Pampa Lirima (136

km NE the city of Iquique), Pampa Apacheta (120 km NE the city of Calama), El Tatio (100 km NE the city of

Calama) and La Torta de Tocorpuri (90 km NE the city of Calama); ii) the pre-Andean basins (Pozo 3, Peine),

iii) the Precordillera chain (Chusmiza), and iv) the Andean basins (Jurase, Las Cuevas, Putre, Pumire,

Chinchillani, Enquelga, Cancosa, Puritama, Puripicar, Aguas Calientes Norte, Ojos de Hécar) (Fig. 1). On the

whole, 72 waters and 57 gases were collected and analysed. The main aims of this study are to: i) recognize the

different sources of the thermal fluids, and ii) provide information, i.e. the physical–chemical conditions

regulating fluid chemistry, that can be helpful to evaluate the geothermal potential of the various hydrothermal

systems.

Compositional features and origin of thermal fluids

Water chemistry

The thermal discharges have outlet temperature varying from 16 to 88 °C and pH values comprised between

1.7 and 7.9. Most of the collected thermal waters have a Na-Cl composition, typical of geothermal fluids, with

the only exception of water samples from La Torta de Tocorpuri and several Andean thermal springs (Jurase,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 530-533

531

Chinchillani, Putre, Enquelga and Cancosa), which show a Na-SO4 composition likely related to the interaction

of meteoric-originated water with dacitic-riolitic rocks, typically found in the Andean environment, in presence

of a CO2(H2S)-rich gas phase (Ellis and Mahon, 1967).

Arica

Iquique

Calama

Antofagasta

CHILE

BOLIVIA

ARGENTINA

Legend

Volcano

Active volcano

Puchuldiza-Tuja

Geothermal field

0 100 200 Km

Surire

Torta de Tocorpuri

El Tatio

Pampa

Apacheta

Pampa Lirima

Thermal spring

Putre

Jurase

Las Cuevas

Chusmiza

Pumire Chinchillani

Cancosa

Pozo 3

Puritama

Aguas Calientes

NorteOjos de Hecar

Enquelga

Peine

Puripicar

City

Figure 1. Map of northern Chile with the location of the geothermal fields and the collected thermal springs.

The concentrations of boron, one of the most useful tracers for geothermal fluids (e.g. Giggenbach, 1991), in

the Na-Cl waters are relatively high (up to 1,020 mg/L), whereas those of the water samples from the La Torta

de Tocorpuri area do not exceed 0.6 mg/L. The high boron concentrations (up to 1,350 mg/L) measured in some

Ca-SO4 thermal springs (Jurase, Putre and Cancosa) are to be related to the presence of borate mineralization that

characterizes the Salar deposits in northern Chile (Chong et al., 2000). The 18O and D values are comprised in

a wide range (from -13.9 to -2.4 ‰ and from -116.6 to -37.5 ‰ V-SMOW, respectively) and indicate that all the

studied hydrothermal systems are recharged by meteoric water, which, during their underground circulation, is

affected by water-rock interactions able to provoke an O-shift. On the basis, contributions from magmatic

sources can be regarded as negligible.

Gas chemistry

The gas/vapour ratios of the thermal discharges is mainly regulated by condensation processes at very shallow

depth, being the fumaroles, present only at Pampa Apacheta and El Tatio hydrothermal areas, and the bubbling

pools generally characterized by dominating H2O and CO2 concentrations, respectively. The composition of the

dry gas phase, besides of CO2, shows significant concentrations of N2 (up to 31,000 μmol/mol) and highly

variable amounts of H2 (from 0.2 to 2,000 μmol/mol) and CH4 (from 0.05 to 2,100 μmol/mol). Hydrogen

sulphide, which is absent in the thermal discharges of Pampa Lirima and La Torta de Tocorpuri, varies from 25

to 1,000 μmol/mol. Highly acidic gas compounds, i.e. HF, HCl and SO2, are virtually absent (<0.1 μmol/mol),

with the only exception of the Pampa Apacheta thermal discharges, where significant concentrations of SO2 and

HCl were detected (up to 1.53 and 6.34 μmol/mol, respectively). CO contents are below the detection limit (0.01

μmol/mol), likely due to its complete formiatation into shallow aquifers. Concerning the organic gas fraction,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 530-533

532

light hydrocarbons are marked by a high speciation, a feature that has been commonly observed in fluids of

worldwide geothermal areas (e.g. Capaccioni et al., 2004; Tassi et al., 2005). Gas species pertaining to the

alkanes group are generally the most abundant ones, although at Pampa Apacheta the light alkenes contents

prevail over those of their homologue alkanes.

In the N2-CH4-Ne ternary diagram (Fig. 2) the gas discharges plot within the area delimited by the following

end-members: i) air and air saturated water (ASW), ii) hydrothermal, and iii) andesitic. Therefore, the origin of

the thermal fluids from all the investigated systems is mainly related to the contribution of at least three different

sources: 1) low-temperature atmospheric–rich fluids, dominating at Surire and La Torta de Tocorpuri, 2)

medium–temperature hydrothermal fluids, particularly at Pampa Lirima and Puchildiza-Tuja, and 3) a high–

temperature magmatic–related component, strongly affecting Pampa Apacheta and El Tatio systems.

mix

ing

mix

ing

andesitic

hydrothermal

Ne*1000 N2/50

CH4*5

aswair

Andinean spring

Torta Tocorpuri

Pampa Lirima

Surire

El Tatio

Apacheta

Puchuldiza-Tuja

Figure 2. N2-CH4-Ne ternary diagram for the gas discharges from Tarapacà and Antofagasta regions (northern Chile).

Geothermometry

The CO2-H2-Ar system (Giggenbach, 1991) was investigated in order to evaluate the reservoir temperatures of

the main hydrothermal areas. As shown in Fig. 3a, gases from Puchuldiza-Tuja, El Tatio and Surire seem to be

produced by a liquid-dominated system at equilibrium temperature comprised between 250 and 300 °C.

Differently, the chemistry of Pampa Apacheta gas samples can be referred to the presence of a vapour phase

equilibrated at temperatures ranging between 300 and 350 °C, likely affected by significant contribution of

magmatic-related fluids, as also indicated by the presence HCl and SO2. On the contrary, gas samples from

Pampa Lirima, La Torta de Tocorpuri and the Andean springs are likely derived from shallow, low temperature

environments where most of H2 is lost and/or consumed by oxidation processes. Equilibrium temperatures

calculated on the basis on the Na+-K+-Ca2+-Mg2+ system (Giggenbach, 1988) (Fig. 3b) are in agreement with

those estimated by gas-geothermometry.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 530-533

533

0 1 2 3 4 5 6

-3

-2

-1

0

1

2

3

4

250 °CAndinean spring

Torta Tocorpuri

Pampa Lirima

Surire

El Tatio

Apacheta

Hconsum

ptio

nand/or

loss

liqu

id

350 °C

300 °C

200 °C

150°C

vapor

log(H

2/A

r)

log(CO2/Ar)

CO dissolution

Puchuldiza-Tuja

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0

0.0

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0.9

1.0

Granite

300°C

250°C

50°C

100°C

150°C

200°C

10M

g2+

/(10M

g2+

+C

a2+

)

10K+/(10K

++Na

+)

CrustBasalt

Fig. 3a-b. Geothermometric plot for gas and water thermal discharges from northern Chile; a) log(H2/Ar) vs. log(CO2/Ar) binary diagram (Giggenbach, 1991); b) 10Mg2+/(10Mg2++Ca2+) vs. 10K+/(10K++Na+) binary diagram (Giggenbach, 1988).

Concluding remarks

The results of the present geochemical survey have provided useful constraints on the reservoir conditions of

the main hydrothermal systems in northern Chile. Most of these thermal areas can be regarded as important

geothermal resources, although further detailed geochemical and geophysical investigations are necessary for an

exhaustive estimation of the geothermal potential. Accordingly, in the forthcoming years the research activity of

the Chilean–Italian group will mainly be aimed to generate a regional framework for the geothermal fields of the

Andean Central Volcanic Zone.

References Capaccioni, B., Taran, Y., Tassi, F., Vaselli, O., Mangani, F., & Macias, J.-L., 2004. Source conditions and degradation

processes of light hydrocarbons in volcanic gases: an example from the Chichon Volcano (Chiapas State, Mexico). Chem. Geol. 206: 81-96.

Chong, G., Pueyo, J., Demergasso, C., 2000. Los yacimientos de boratos de Chile. Rev. Geol. Chile 27: 99-119. Cusicanqui, H., Mahon, W.-A., & Ellis, A.-J., 1975. “The geochemistry of the El Tatio geothermal field, Northern Chile.“ 2nd

UN Symposium Development and Utilization of Geothermal Resources, San Francisco, 703-711. Ellis, A., & Mahon, W., 1967. Natural hydrothermal systems and experimental hot water/rock interactions (Part II). Geochim.

Cosmochim. Acta 31: 519-538. Francis, P., & Rundle, C., 1976. Rates of production of the main magma types in the Central Andes. Geol. Soc. Am. Bull. 87:

474–480. Giggenbach, W., 1988. Geothermal solute equilibria, derivation of Na-K-Mg-Ca geoindicators. Geochim. Cosmochim. Acta

52: 2749-2765. Giggenbach, W., 1991. “Chemical techniques in geothermal exploration.” In D’Amore, F. (ed): Application of geochemistry

in geothermal reservoir development, UNITAR: 253–273. Lahsen, A., 1975. Evaluación preliminar del sistema geotérmico Puchuldiza. Unpubl. Report CORFO, 23 p. Lahsen, A., 1976. “La actividad geotermal y sus relaciones con la tectónica y el volcanismo en el Norte de Chile.“ 1st Chilean

Cong. Geol., B105–B127. Lahsen, A., & Trujillo, P., 1976. El campo geotermico El Tatio, Chile. Unpubl. report CORFO, 21 p. Stern, C., 2004. Active Andean volcanism: its geologic and tectonic setting. Rev. Geol. Chile 31: 161-206. Tassi, F., Martínez, C., Vaselli, O., Capaccioni, B., & Viramonte, J., 2005. The light lydrocarbons as a new geoindicator for

temperature and redox conditions of geothermal fields: Evidence from the El Tatio (Northern Chile). App. Geochem. 20: 2049-2062.

Urzua, L., Powell, T., Cumming, W., & Dobson, P., 2002. Apacheta, a New Geothermal Prospect in Northern Chile. Geoth. Res. Counc. Tran. 26: 65-6.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 534-537

534

The June 23, 2001, southern Peru earthquake

Hernando Tavera & Isabel Bernal

Instituto Geofísico del Perú, Calle Badajos 169 Urba. Mayorazgo IV Etapa , Ate, Lima, Peru

(hjtavera@geo.igp.gob.pe)

Introduction

The western border of South America is one of the most important sismogenic regions in the world. In this

region did occur the most damaging earthquakes known and reported in the news. One of these earthquakes

occurred in June 23, 2001 (Mw=8.1-8.2) and produced death and damages in the whole southern region of Peru.

This earthquake was originated by a friction process between Nazca and Sudamericana Plates and affected an

area of about 300km x120km defined by the distribution of more than 220 aftershocks recorded by a local

seismic network that operated 20 days. The epicenter of the main shock was localized in the northwestern

extremity of the aftershock area and this suggests that the rupture propagated towards the SE direction. The P-

wave modelization for teleseismic distances permited to define a focal mechanism of reverse type with nodal

planes oriented NW-SE and a possible fault plane dipping gently toward the NE. The STF suggest a complex

process of rupture during 85 seconds with 2 succesive sources, the second one of greater size, and located

approximately 100-120 km toward the SE direction. It was estimated a rupture velocity of about 2 km/seg on a

28°-dipping plane to the SE (1xx°). A second event happened 45 seconds after the first one with an epicenter

130km farther to the SE and a complex STF. This event and the second source of the main shock gave origin to a

tsunami with waves from 7 to 8 meters that propagated almost orthogonally to the coast line affecting mainly the

Camana area.

From all the aftershocks, three presented magnitudes greater or equal to Mw=6.6, two of them occurred in

front of Ilo and Mollendo (June 26 and July 7) with focal mechanisms similar to the main seismic event.

Aftershock of July 5 corresponded to a normal mechanism at a focal depth of 75km, with a probable origin

inside the Nazca plate under the friction zone. The aftershocks of June 26 (Mw=6.6) and July 5 (Mw=6.6) show

plain STF with short duration. The aftershock of July 7 (Mw=7.5) with duration of 27 seconds suggests a

complex process of energy release with the possible occurrence of a secondary shock with lower focal depth and

focal mechanism of inverse type with a great lateral component. Plain focal mechanisms and composite ones

were calculated for the aftershocks, and all of them show characteristics similar to the main one.

The June 23 earthquake induced major damages in the whole southern Peru. The damage estimation in towns

of Arequipa, Moquegua allow to consider maximum intensities from 6 to 7 (MSK79). In Alto de la Alianza and

Ciudad Nueva zones from Tacna, the maximum intensity was of 7- (MSK79).

Discussion and conclusion

The southern Peru region was affected after 133 years by an Mw=8.2 earthquake that occurred in June 23,

2001. Preliminary studies allowed to consider this earthquake as a recurrence of the one of August 1868, that

was assigned Mw=9.0, a rupture length of about 500km and intensities of about X-XI (MM) (Dorbath et al,

1990; Comte and Pardo, 1991). Later, Tavera et al. (2001), Kikuchi and Yamanaka (2001), Tavera et al. (2002),

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 534-537

535

Giovanni et al. (2002), Bilek and Ruff (2002), and Dewey et al. (2003) demonstrated that June 2001 earthquake

presented a lower magnitude (Mw=8.2), rupture length of 300km and maximum intensity of VII-VIII (MM),

therefore it is not a recurrent earthquake.

The focal depth was estimated to 23km and from the waveform inversion to 29 km (Table 3), approaching

those values of USGS (33 km); Kikuchi and Yamanaka (2001), 32 km and Giovanni et al (2002), 20 km. The

earthquake magnitude was estimated by the IGP to ML (d)=6.9 and from waveform inversion to Mw=8.1 similar

to those reported by other international agencies (mb=6.7, GS; Ms=8.2, GS; Mw=8.4, HRV; Mw=8.2, Kikuchi

and Yamanaka (2001); Mw=8.2, Giovanni et al (2002); Mw=8.4, Bilek and Ruff (2002)).

The relocation of 220 aftershocks with magnitudes ML between 2.4 and 4.8 allowed to define a rupture area

of about 300km x 120 km with the epicenter of the main shock located in the NW extremity of this area, that

suggest an unilateral propagation of the rupture toward the SE, as suggested by Giovanni et al, (2002) and Bilek

and Ruff (2002). According to the aftershock distribution, the rupture stopped abruptly in front of Ilo town,

producing two aftershocks with magnitude Mw=6.6 y 7.5, and thus delimiting the initiation of a new area of

energy accumulation. The aftershocks form three clusters, the first one concentric fully around the main shock,

the second one near the trench and the third, spread in the SE end. In between those clusters can be observed the

presence of another area that would not have experienced rupture and on the contrary the displacement would

have taken place in aseismic way.

Spatial distribution of June 23, 2001 earthquake and aftershock series. The focal mechanisms were obtained from the P-wave model and polarity. The crosshatched area corresponds to aftershocks of the 1996 November aftershock (Tavera et al, 1998) and the shaded one to the asperity of the aftershock area of the June 23 earthquake. The discontinuous line indicates the various aftershock swarms. In the lower part is presented a vertical cross section with the aftershocks, indicated as A-B. Triangles are indicating the seismic station disposed during this study.

The focal mechanisms

obtained for the main shock and

major aftershocks correspond to

reverse type with NW-SE nodal

planes, being the possible fault

plane the one dipping gently to

the NE. The focal mechanisms

corresponding to lower

magnitude already events, whither composed or simple are similar to the main shock one, even if cluster GRUP1

presents some lateral component. The 5 of July aftershock, with epicenter inside the continent located NE to the

main shock, presents a normal focal mechanism with NW-SW planes and a possible fault plane nearly vertical.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 534-537

536

This earthquake focal depth has been estimated to 75km, therefore it should be associated to the Nazca plate

internal deformation, below the friction level, and suggests a depth zone in which compression stresses change

to extension. In northern Chile, this zone this located to 60 km of depth approximately (Comte and Suarez,

1995).

Synthetic (sint) and observed (obs) waveform corresponding to the main shock obtained with inversion after the Nabelek method (Nabelek, 1984). The record amplitude has been normalized to a gain of 5000 and a distance of 40°. The inversion window is indicated with vertical lines on the record. The station identification is found at the extreme left side of the record and below the epicentral distance and azimuth in degrees. The focal sphere corresponds to its projection in the lower hemisphere, after P and T axis represented by black and white circles. The SFT is presented below the focal mechanism, just as the record scale. In the upper part is presented the focal mechanisms corresponding to both seismic events (E1, E2) and the solution after Tavera et al (2002) thanks to the P-wave polarity (TB). The upper left side figure shows the epicenter localization of 2 seismic events associated to the rupture process of the June 23, 2001 earthquake.

The STF characteristics suggest that the

main shock presented a very complex process

of rupture during 85 seconds. During this time

period occurred two main ruptures, the first

one at the onset of the earthquake that lasted 25 seconds before to slowly propagate until to produce a major

rupture of 45 seconds approximately 100 km in the SE. As suggested by Giovanni et al (2002) and Bilek and

Ruff (2002), in between both sources is encountered an asperity, but in this case of lower size. A second event,

complex too, occurred 40 seconds after the first one, with its epicenter localized 120-130km toward the SE

respecting to the beginning of the rupture. This second event and the second source of the main event produced

the greatest energy release in front of Camana town, just as suggested by Kikuchi and Yamanaka (2001),

Giovanni et al (2002) y Bilek and Ruff (2002). This whole rupture process developed on a surface dipping about

28° with a velocity of 2 km/seg. The rupture velocity explains the lasting time of the shock and perhaps the

damage extent and induced effects in surface that were not so big compared to the earthquake size. The June 26

and July 5 aftershocks, presents one SFT very simple and of short duration. For July 7 aftershock, the SFT lasted

27 seconds and suggested an occurrence due to a complex rupture process. A second event, with epicenter

located to the west of first one, produced 7 seconds after with duration of 23 seconds. The first event presented

an 18km-depth and the second one a 12km-depth consistent with its epicenter localization and that could suggest

the propagation of the rupture to the west. The second event focal mechanism is of reverse type with a big lateral

component that suggests the occurrence of a more complex rupture processes that could implicate the two

internal plates. In Table 3 is presented a summary of the source parameters obtained in this study for the main

shocks and its great aftershocks.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 534-537

537

The results obtained in this study demonstrate that the June 2001 earthquake has produced one of the most

complicated rupture processes known for Peruvian earthquakes. The local tsunami characteristics are complex

too and gave birth to the waves that affected to Camana town. Data collected from 15 people, confirmed the

formation of marine currents that circulated parallel to the coastline, from Chala to Atico toward the SE, and

from Mollendo and Matarani toward the NW, apparently with major velocity and that could have run into each

other near from Camana and facilitate the formation of waves that propagated oblique to the coastline. This

could explain the 32km-inland flooding de 32 km of beaches south of Camana. Obviously the tsunami

characteristics could be explained by the complex pattern of the main shock rupture.

Considering the rupture propagation to the SE, to the margin of the engineering parameters, major damages

and effects should be produced in towns and villages of the southern Peru. In terms of acceleration, the

accelerometer located in Huancavelica city (420 km to the NE of the epicenter of the earthquake) has registered

an acceleration of 11.6 cm/seg2; whereas, the accelerometers of the Arica-Chile city (455 km to of epicenter of

the earthquake) registered a acceleration of 284 cm/seg2, which suggests it energy propagated in direction SE,

coherent with the damage assessed in the southern region. The intensity estimations show that in this area the

maximum intensity was about 6+ to 7- MSK79, excepted in the districts of Alto de la Alianza and Ciudad Nueva

where the major percentage of damages occurred for the houses because of the low quality level of the build

work. Similar damages occurred in Moquegua, but in this case the houses were mainly old ones built with mud

and without any building techniques. Lower damages were reported in Arequipa and near town areas. The

presence of geologically inadequate grounds to build houses and public edifices played a major role in the

increase of damages induced by the earthquake (sewers, water pipes, public phones and electrical maintenance).

The June 23 earthquake is one of the most important shocks that occurred in this region, as well as one of the

most complex one in terms of rupture process. This magnitude Mw=8.1 shock showing a rupture length of 300

km, cannot be seen as a repetition of the August 1868 earthquake. This earthquake presented magnitude of

Mw=9.0 and one length of rupture of 500 km; that is to say, 200 km but that the earthquake of the 2001. This

new energy accumulation zone should give birth to a new high magnitude earthquake in the future.

References Bilek, S. and Ruff, L., 2002, Analysis of the 23 June 2001 Mw=8.4 Peru under thrusting earthquake and its aftershock.

Geophys. Res. Lett., 29, 21:1-21:4. Comte, D. and Pardo, M., 1991, Reappraisal of great historical earthquakes in the northern Chile and southern Peru seismic

gaps. Natural Hazards, 4, 23-44 Dorbath, L., Cisternas, A. and Dorbath, C., 1990, Assessment of the size of large and great historical earthquake in Peru.

Bull. Seism. Soc. Am., 80, 551-576. Giovanni, M., Beck, S. and Wagner, L., 2002, The June 23, 2001 Peru earthquake and southern Peru subduction zone.

Geophys. Res. Lett., 29, 14:1-14:4. Kikuchi, M., and Yamanaka, Y., 2001, Near coast of Peru earthquake (Mw=8.2) on June 23, 2001 (revised). EIC

Seismological Note: N°105, posted on the website of the University of Tokio Earthquake Information Center. Nabelek, J., 1984, Determination of earthquake source parameters from inversion of body waves. PhD Thesis, MIT,

Cambridge, MA. Nishenko, S., 1985, Seismic potential for large and great interplate earthquakes along the Chilean and southern Peruviann

margins of South America: a quantitative reappraisal. J. Geophys. Res., 90, 3589-3615. Ocola, L., 1979, Intensidades sísmicas del sismo. XII Congreso de Ingeniería, Universidad de Ingeniería, Lima-Perú. Tavera, H., Buforn, E., Bernal, I. and Antayhua, Y., 2002, The Arequipa (Peru) earthquake of June 23, 2001. Journal of

Seismology, 6, 279-283. Tavera, H. and Buforn, E., 2001, Source mechanisms of earthquakes in Peru. Journal of Seismology, 5, 519-539.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 538-541

538

Preliminary petrological, geochemical and stratigraphical characterization of the Sotará volcano, SW Colombia

L. Téllez1, M.I. Marín-Cerón

1, G. Toro

1, & B. Pulgarín

2

1 EAFIT University, Department of Geology, Medellin, Colombia (ltelleza@eafit.edu.co, mmarince@eafit.edu.com,

gtoro@eafit.edu.co) 2 INGEOMINAS, Observatorio Vulcanológico de Popayán, Popayán, Colombia (bpulgarin@ingeominas.gov.co)

KEYWORDS : NVZ, Sotará volcano, andesites, ignimbrites deposits

Introduction

The Sotará Volcanic Complex (SVC) is located in the SW Colombian volcanic arc between Doña Juana and

Puracé-Coconucos volcanic complexes, and is part of the Andean Central Cordillera. It is formed by Sotará,

Cerro Gordo, Cerro Negro and Azafatudo volcanoes. The complex has been constructed during Plio-Pleistocene

time. It is composed by lavas ranging from basic andesites to andesites and dacites at the lava domes, andesitic

ignimbrites, pyroclastic flows and air fall deposits.

The Sotará volcano is an active strato-volcano which combines effusive and explosive type eruptions. At least

two stages have been identified: external pre-caldera and external post-caldera. The external pre-caldera lavas

are basic andesites associated to lava flows with less viscosity; they are located 3 to 4 km to the North of the

actual Sotará volcano. After that effusive activity an explosive period started associated to ignimbrite deposits

and the external caldera formation (Acevedo and Cepeda, 1982). During the Pleistocene glaciation the oldest

volcanic edifice may be eroded and a relatively gap-activity permitted the drainage system incision forming the

typical deep valleys of that volcanic area. At the beginning of the Quaternary, the external post-caldera stage

concentrated to the SW flank of the actual volcanic center has begun with eruptions of ignimbrites, lavas and

pyroclastic flows.

Taking into account the explosive-effusive behaviour of this volcanic complex, a detailed study is much

needed to understand the spatial distribution and temporal variations of this volcano. Here, we present a new and

detailed stratigraphic column of the Sotará volcano, together with major and trace element and Sr-Nd isotopic

systematics for representative samples. The aim of this study is to contribute in the understanding of the Sotará

volcanic complex activity and to compare these results with the recently obtained data from other volcanoes in

the SW of Colombia.

Tectonic setting and regional geology

The active volcanism within the Andes is divided into three zones (e.g. Thorpe et al., 1982): North Volcanic

Zone (NVZ), Central Volcanic Zone (CVZ) and South Volcanic Zone (SVZ), related to the interaction of Nazca

plate beneath South American plate.

The Colombian volcanic arc consists of some forty Cenozoic-Quaternary volcanoes, of which twenty are

considerate as active (Hanke & Parodi, 1966; Mendez, 1989). The volcanoes at this arc are strongly controlled

by tectonics and especially by two main fault zones of Cauca-Patía and Romeral-Dolores, together with

associated secondary faults.

Sotará volcanic complex is located in the Cauca department (Fig 1), at approximately 180 km to the depth of

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 538-541

539

the Waddati-Beniof zone. The basement under Sotará volcano is composed by Precambrian-Paleozoic

metamorphic constituted by green and carbonaceous schists and quarcites separated of the diabasic rocks of the

Diabasic Group by Romeral fault system (Paris & Cepeda, 1978) (Fig 1). Small porphyritic bodies have intruded

the area during Tertiary time. Several authors have proposed that the lower crust in this region may be related to

a highly Pb-radiogenic crust (e.g. Weber et al., 2002; Marín-Cerón, 2007) which could be associated with a

crustal-make up event during the accretion of the Caribbean-Colombian Oceanic plateau.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 538-541

540

Generalized stratigraphic column

The geological map and the generalized stratigraphic column of the study area are summarized in the figure

1a-b. The general feature of the SVC behaviour since the late Tertiary is the series of effusive and explosive

periods (Acevedo & Cepeda, 1982). It is clear from this stratigraphic sequence that the Late Tertiary activity is

mainly exposed along the Guachicono River which ended with a large ignimbrite event. After the last glaciation,

the new Sotará volcanic complex started to build up.

Several ignimbrites deposits have been recognized during the Quaternary (see Fig. 1b): Auca, La Virgen, El

Barrial, La Quinquina y las Cabras ignimbrites. The above deposits range in thickness from 20 m to 200 m,

filing the valleys of Rio Negro and Rio Blanco rivers. In general small ignimbrites deposits (20-30 m) are

followed by andesites lavas flows with 80 to 100 m in thickness. These ignimbrites events could be related to the

formation of the caldera. There after, a period of relatively calm has been related with small pyroclastic flows,

air fall deposits, base surge deposits and the formation of the dacitic lava-dome complex inside of the caldera.

At the moment, the stratigraphic sequence presented here has been determined by field work correlations. A

detailed age dating determination is carrying out using fission track method in zircon grains for the most

representative volcanic events.

Petrography

Sotará volcanic rocks consist mainly of andesites with some variations to basic andesites and dacites. Detailed

Petrographic analyses of volcanic rocks reveal a complex phenocryst assemblage:

Pl+Opx±Cpx±Ol±Qz+Mt+Ilm. The general texture is seriate-porphyritic texture (Fig. 2), with several

disequilibrium features (e.g. sieve texture, reactions rims, crystal aggregates, complex plagioclase and pyroxenes

zoning and silica-rich melt inclusions in plagioclase following the zoning pattern). In the case of dacitic rocks

they appear highly vesiculated and are mainly restricted to the actual lava domes complex.

The modal distribution on those rocks is variable in terms of amphibole and pyroxene content. Plagioclase

modal composition is always the highest (15-30%) follow by orthopyroxene (5-10%). The amount of amphibole

can reach up to 10% in the youngest andesites lavas and ignimbrites compared with the amount of clinopyroxene

(>5%) at the early stages. Andesites with olivine appear as a small window in the northern flank near the

Cuchilla la Ensillada.

Geochemistry and isotope analyses

Whole-rock geochemical (major and trace elements) and isotopic (Sr- Nd) data of representative samples of

the SVC were obtained at the laboratories of Shimane University (Japan). Representative whole-rock, major and

trace element composition of the SVC are shown in the figure 3a-c. The range of SiO2 content of the volcanic

lavas is between 53.4 and 68.4 wt % which are classified as basaltic andesite to dacites using TAS classification

diagram. Rocks plot just on the border with alkali- subalkaline field (Fig. 3a). On the base of potassium content

all those rocks fall in the transitional line of medium-k to high-k series calc-alkaline series (Figure 3b).

The general pattern of the samples studied in PM-normalized diagrams is showing the typical subducted

signatures such as Nb and Ta negative anomalies and enrichment of fluid-mobile elements compared with the

other volcanoes in the SW Colombian volcanic arc (Marín-Cerón et al., 2008). In general those rocks have high

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 538-541

541

values of Ba (746-1643 ppm), Sr (376-591 ppm), Zr (99-149 ppm) and Nb (7-12 ppm), comparing with the other

volcanoes located at its south such as Azufral, Galeras and Doña Juana volcanoes and less concentrations in the

same elements compared with the Purace-Coconucos volcanic complex (Marín-Cerón et al., 2008).

Isotopically, both andesitic and dacitic volcanic rocks of the SVC have a little variation in 87Sr/86Sr (0.704061-

0.704343) during the Plio-Pleistocene time (Fig. 3c) and fall within the mantle-array field. Those data are

consistent with the data reported by Wilson (1997) for the NVZ and the reported data by Marín-Cerón (2007) for

the volcanoes in the SW of Colombia. The variation in 143Nd/144Nd ratio (0.512709, 0.512805) is also very

narrow and fall within the values reported in the NVZ with values between (0.5127 - 0.5130) (e.g. Wilson, 1997;

Calvache et al. 1997; Marín-Cerón, 2004; Marín-Cerón, 2007).

Conclusions

To the light of our preliminary results, we can suggest that the conditions at the magma chamber may

drastically changed, increasing the pressure gradient and driving magma out to the reservoir in a short time

intervals combining lava extrusions and ignimbrite deposits. On the basis of stratigraphy, petrography and

geochemistry (major, trace elements and Sr-Nd isotopic systematics) we can conclude the magma source of the

volcanic products of the Sotará volcanic complex indicates the dominance of mantle-derived subducted-related

magmas. However, the petrographical and geochemical variations shown by this volcanic complex cannot be

explained just by simple fractional crystallization processes and therefore much more complicated processes

such as assimilation and fractional crystallization (AFC) may be invoked. The formation of those silicic-rich

rocks could be explained by the model proposed by Marín-Cerón et al. (2008) for the andesites generation at the

SW Colombian volcanic arc. So that, the variations in the modal composition of the volcanic products in terms

of pyroxenes and amphibole contents could be related to the amount of water and volatiles in the mantle-derived

primary magmas during their time of residence at the lower crustal region. Finally, the variation noticed on Sr

and Nd isotopic data with silica content can be related with more crustal assimilation.

References Acevedo, A.P. & Cepeda, H., 1982. El volcaán Sotará : Geologia y Geoquimica de elementos mayores. Publicaciones

Geológicas Especiales del Ingeominas V.10, pp.19-30. Calvache, M.L. & Williams S., 1997. Geochemistry and Petrology of the Galeras volcanic complex, Colombia. Journal of

Volcanology and Geothermal Research 7721-38. Hanke G. & Parodi I., 1966. Catalogue of the active volcanoes of the worl including solfatara fields. Part XIX. International

Association of Volcanology, 73 pp. LeMaitre, R., Bateman, P., Dudek, A., Keller, J., Lameyre-Lebas, M., Sabine, P. Schmid, R., Soresen, R., Streckeisen, A.,

Woolley, A. & Zanetting, B., 1989. Classification of Igneous Rocks and Glossary of Terms Oxford : Blackwell. Marín-Cerón, M.I, 2004 Geochemical variation of late Cenozoic volcanic rocks in time and space, SW Colombia.

Unpublished, MSc. Thesis Shimane University Japan. 105p. Marín-Cerón, M.I, 2007. Major, trace element and multi-isotopic systematics of SW Colombian volcanic arc, northern

Andes: Implication for the stability of carbonate-rich sediment at subduction zone and the genesis of andesite magma. Unpublished, PhD. Thesis Okayama University, Japan. 131 p.

Marín-Cerón, M.I, Moriguti, T. & Nakamura, E., 2008. Andesite magma generation at the Plio-Quaternary SW Colombian volcanic arc. Inthis symposium: 7th International Symposium on Andean Geodinamics.

Méndez, R.A., 1989. Catálogo de los volcanes activos de Colombia. Bol. Geol. V30, no. 3, Ingeominas, Bogotá, 75 p. Thorpe, R. S., Francis, P. W, Hammill M. & Baker M.C.W., 1982. The Andes. Andesites. Ed Thorpe, R.S., pp. 187-205. Paris, G. & Cepeda, H., 1978. Algunos complejos ultramaficos en los departamentos del Cauca y Narino, Colombia. Informe

IRP-011. Ingeominas, Popayan. 22 p. Weber, M.B.I., Tarney, J., Kempton, P.D. & Kent, R. W., 2002. Crustal make-up of the northern Andes: evidence based on

deep crustal xenolith suites, Mercaderes, SW Colombia. Tectonophysics 345, p. 49–82. Wilson, M., 1989. Igneous Petrogenesis: a Global Tectonic Approach. London: Unwin Hyman.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 542-544

542

Quaternary soft-linked fault systems highlighted through drainage anomalies in the northwestern Precordillera Sur (32ºS), Central Andes of Argentina

C. M. Terrizzano & J. M. Cortés

Laboratorio de Neotectónica (LANEO), FCEyN, Universidad de Buenos Aires, Pabellón 2, Ciudad Universitaria,

C1428EHA Buenos Aires, Argentina (cterrizzano@gl.fcen.uba.ar, cortes@gl.fcen.uba.ar)

KEYWORDS : soft linkage, drainage anomalies, Neotectonics, Precordillera Sur

Introduction

At the Central Andes, between 31° and 34° SL, the paleotectonic features represented by sutures zones

between allochtonous Paleozoic terranes and branches of Mesozoic extensional basins concentrate much of the

Quaternary and active deformation. At 32º SL, the Cuyana basin northern segment, which is Triassic in age,

controls the location of seismically active NW trending morphotectonic units. At this location, the northwestern

piedmont of the Precordillera Sur (Cortes et al. 2005) has been uplifted during the Quaternary (Cortes y Costa

1993, Cortés y Cegarra 2003, Terrizzano 2006 a and b) through folds, blind faults, shear faults and faulted

blocks that form discrete areas of brittle and brittle-ductile shear zones at different scales (Terrizzano et al.

2007). Quaternary deformation is made evident through the presence of piedmont fault scarps, Quaternary

pressure ridges, Quaternary tilted deposits and different kinds of drainage anomalies.

This contribution analyzes the geomorphic and structural features at the piedmont of the El Naranjo and

Ansilta ranges, northwestern Precordillera Sur (Figures 1a and 1b). As a result of this analysis it was possible to

develop different criteria for characterizing such Pleistocene to Recent piedmontal tectonic geophorms.

Drainage anomalies areas as indicators of soft-linkage in Quaternary fault

systems

Contractional active deformation of Quaternary alluvial deposits in piedmont environments becomes evident

from an accentuated relief modification to subtle perturbations of fluvial geomorphologic elements.

The cumulative deformation associated to growth structures develops tectonically controlled topographic highs

(TH) characterized by definite and clear borders, lineal, trapezoidal or irregular design at surface (Figure 1c).

This phenomenon takes place when the tectonic uplift rate is greater than the erosion and sedimentation rate.

These features are mainly related to faults, folds, faulted blocks, push-up structures and imbricate thrusts.

On the other hand, the subtle tectonic perturbation of the fluvial landforms is made evident by different kinds

of drainage anomalies (Howard 1967, DeBlieux 1949) such as deflected streams, asymmetric basins, zonal

variations in drainage density, anomalous stream sinuosity and incision anomalies (Figure 2a). The association

of these geomorphologic features defines tectonically controlled drainage anomalies areas (DAA), which are

characterized by indefinite or diffuse borders and lineal, trapezoidal or irregular design at surface (Figure 2b).

These elements are associated with subtle uplift, rotation, propagation or migration of structures.

An analysis of the distribution and spatial relationship between the tectonically controlled topographic highs

and the tectonically controlled drainage anomalies areas at piedmont plains of northwestern Precordillera Sur has

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 542-544

543

made possible to elucidate the role of the drainage anomalies areas as a superficial expression of soft-linkage

areas in fault systems (Figure 2b).

Figure 1. a) Study area location, at San Juan and Mendoza provinces, in the Central Andes of Argentina. b) Main ranges at the northwestern edge of the Precordillera Sur. Notice some uplifted features at the piedmont area. c) Neotectonic faults, Quaternary fault scarps, Quaternary anticlines and tectonically controlled topographic highs which are located at the same region.

Conclusions

A useful tool at the moment of determining slight tilting and ductile strain is given by a detailed study of

drainage areas of anomalous behavior. As an example in fault bridges, which work as soft linkage zones between

major structures.

The tectonically controlled topographic highs would be interpreted as isolated features without the

consideration of subtle drainage anomalies between them. However, the set of tectonically controlled

topographic highs linked by tectonically controlled drainage anomalies areas makes possible to clarify major

regional structures like Quaternary deformation belts on their initial development stages, which exhibit

mechanical interconnection.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 542-544

544

Figure 2. a) Neotectonic faults, Quaternary fault scarps, Quaternary anticlines and drainage anomalies at the study area. Green squares: sinuosity anomalies; orange circles: deflected streams; pink triangles: asymmetrical basins; purple crosses: incision anomalies. b) Tectonically controlled topographic highs (TH) in light orange and controlled drainage anomalies areas (DAA) in light yellow. Notice the spatial distribution between them which highlights the TH mechanical interconnection through soft-linkage zones (DAA). In black dot lines: Yalguaraz belt edges.

At 32º SL, a wide neotectonic soft-linkage deformation belt, the Yalguaraz belt (Figure 2b) could be

highlighted. This belt links the northwestern edge of the Precordillera Sur (Barreal block, El Naranjo range,

Ansilta range, El Abra high) with minor tectonic blocks (Cucaracha range) of the Cordillera Frontal (Figures 1b

and 2b).

The presence of an anisotropic basement with Paleozoic and Mesozoic previous structures at the northwestern

side of the Precordillera Sur seems to favour the ductile linkage of Quaternary faults along wide deformation

belts.

References Cortés, J., Yamin, M. & Pasini, M. 2005. La Precordillera Sur, Provincias de Mendoza y San Juan. Actas 16° Congreso

Geologico Argentino, Tomo 1: 395-402, La Plata, Argentina. Cortés, J. M. & Cegarra, M. 2004. Plegamiento Cuaternario transpresivo en el piedemonte suroccidental de la Precordillera

sanjuanina. Revista de la Asoción Geológica Argentina, Serie D, Publicación Especial 7: 68-75; Buenos Aires, Argentina. Cortés, J.M. & Costa, C.H. 1993. La deformación cuaternaria pedemontana al norte de la pampa Yalguaraz, margen

occidental de la Precordillera de San Juan y Mendoza. Actas del 12º Congreso Geológico Argerntino y 2º Congreso de Exploración de Hidrocarburos, 3: 241-245, Buenos Aires, Argentina.

DeBlieux, C.W. 1949. Photogeology in Gulf Coast exploration. American Association of Petroleum Geologists Bulletin, v. 33: 1251-1259.

Howard A.D. 1967. Drainage analysis in geologic interpretation: a summation. American Association of Petroleum Geologists Bulletin, v. 51: 2246-2259.

Terrizzano, C. M. 2006a Deformación transpresiva pleistocena en el piedemonte de la depresión de Barreal – Uspallata, Precordillera Sur, Argentina. Actas 11º Congreso Geológico Chileno: 465-468, Antofagasta, 2º Región, Chile.

Terrizzano, C.M. 2006b. Deformación cuaternaria en las Lomitas Negras, cinturón Barreal_Las Peñas, provincia de San Juan. Resúmenes de la 13º Reunión de Tectónica: 57, San Luis, Argentina.

Terrizzano C. M., Fazzito S. Y.,Cortés J. M. & Rapalini A. E. 2007. Quaternary transpressive zones in the Barreal – Las Peñas belt, Precordillera Sur, Argentina: an structural and geophysical approach. Abstracts 20th Colloquium on Latin American Earth Sciences: 60-61, Kiel, Germany.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 545-548

545

Neogene ignimbrites and volcanic edifices in southern Peru: Stratigraphy and time-volume-composition relationships

J.-C. Thouret1, M. Mamani

2, G. Wörner

2, P. Paquereau-Lebti

1, M.-C. Gerbe

3, A. Delacour

3,

M. Rivera1,4

, L. Cacya4, J. Mariño

4, & B. Singer

5

1 Laboratoire Magmas et Volcans UMR 6524 CNRS, OPGC et IRD, université Blaise Pascal, 5 rue Kessler,

63038 Clermont-Ferrand cedex, France (thouret@opgc.univ-bpclermont.fr) 2 Abt. Geochemie, GZG, Universität Göttingen, Goldschmidtstrasse 1, 37077 Göttingen, Germany

3 Laboratoire Magmas et Volcans UMR 6524 CNRS et Université J. Monnet, Saint-Etienne, France

4 Ingemmet, Instituto Geológico, Minero y Metalúrgico, Av. Canadá 1470, San Borja, Lima, Peru

5 Departement of Geology and Geophysics, University of Wisconsin, Madison, WI 53706, USA

KEYWORDS : volcanic arc, Neogene, Peru, ignimbrites, volcanic edifices, chronostratigraphy

Introduction

In the Central Andes of Peru, four volcanic arcs, termed Tacaza, Lower and Upper Barroso, and Frontal arc,

have been active over the past 30 Ma (Fig. 1). They form five units between Moquegua and Nazca (14°30–

17°15’°S and 70–74°W). The ‘Neogene ignimbrites’ (<25 Ma) comprise six generations of widespread sheets

(>500 km2 and >20 km3 each), representing a major crustal melting event, triggered by thickening and advective

heat input from the mantle wedge. Also, four generations of edifices (i.e shields, composite cones, and dome

clusters) and monogenetic fields mostly overly the ignimbrites based on ages, stratigraphy and mapping.

Figure 1. Extent of five volcanic units over the past 30 Ma in southern Peru (Mamani et al., 2008a).

Pre and post-valley incision ignimbrite sheets and western CVZ evolution

Our new stratigraphy (Fig. 2) records changing magma composition, uplift and valley incision of the Central

Andes, and the rate of growth and degradation of the Early Miocene to Holocene volcanoes.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 545-548

546

Figure 2. Stratigraphy and chronology of ignimbrites and lava edifices in southern Peru. The evolution of the western Andean range in Peru is also indicated. ‘Dad’ stands for debris avalanche deposit.

The older ignimbrite sheets pre-date valley incision and are intercalated with voluminous conglomerates that

reflect major phases of surface uplift as a response to tectonic phases in a crust weakened by massive crustal

melting. 1) The 24.6-21.8 Ma-old, welded Nazca ignimbrite caps extensive plateaus to the N and W of the area

as well as further S near Moquegua. 2) The welded, 19.4-18 Ma-old Alpabamba ignimbrite and 3) the

14.3-12.7 Ma-old Huaylillas ignimbrite form extensive plateaus between 4000 and 4500 m S of Coropuna and N

of Cotahuasi. They blanket the polygenetic ‘Puna’ peneplain formed between >40 and 14 Ma (Gunnell et al.,

2008). The ignimbrites erupted from calderas (e.g. N of Alca, NW of Oyolo) during growth of the Western

Cordillera between 24 and 13 Ma. Distal tuffs of these ignimbrites are interbedded in forearc deposits towards

the top of the Moquegua Formation conglomerates (Roperch et al., 2006) in the Majes, Sihuas and Vitor valleys.

The younger, less widespread ignimbrites that filled tectonic basins or were channelled in deep valleys,

postdate valley incision. 1) The 9.4-8.8 Ma-old Caraveli ignimbrites fill an irregular topography cut in the

Huaylillas ignimbrites and crown small and low plateaus at 3000 m asl. to the W of the area. They were

emplaced in shallow wide valleys cut in the peneplain, thus indicating that the fluvial incision had already begun

by 9 Ma. 2) The 4.9-3.6 Ma-old non-welded lower Sencca ignimbrites (Lower Barroso equivalents) crop out in

conglomeratic piedmonts or are preserved on deep valley flanks. The 4.86 Ma-old La Joya ignimbrite

(Paquereau-Lebti et al., 2006) fills the Arequipa depression. 3) The non-welded 2.3-1.4 Ma-old upper Sencca

ignimbrites (Upper Barroso equivalents) crop out in similar stratigraphic positions and comprise the non-welded

Arequipa Airport ignimbrite (1.63 Ma: Paquereau-Lebti et al., 2007) filling the Arequipa depression. Calderas

are not clearly identified but magnetic fabric and AMS measurements (Paquereau-Lebti et al., 2007) indicate that

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 545-548

547

Sencca ignimbrites are probably sourced at calderas or crater clusters that are buried beneath younger volcanic

complexes such as Chachani, Coropuna, and Ampato. A second phase of valley incision took place after 2.2 Ma

(Río Colca valley) or 1.4 Ma, the age of non-welded pumice-flow deposits, which had largely filled the canyon

of Río Cotahuasi. Younger ignimbrites exist but none exceeds 200 km2 and 10 km3. One such example is the

Yura Tuff N and W of Chachani, that may be contemporaneous with the Capillune Formation.

Four generations of edifices and time-volume-composition relationships

Dated lava flows and pyroclastic deposits indicate that four generations of composite and monogenetic

edifices have crowned the Western Cordillera and mostly overlie the ignimbrites (Fig. 2). 1) Although the

Tacaza arc is deeply eroded, roots of hydrothermally-altered edifices remain in the Caylloma area 60 km N of

the Frontal arc. 2) The 9 to 4 Ma-old Lower Barroso edifices are moderately eroded, subdued shields with a core

of 6-4 Ma-old basaltic andesite and andesite lava flows (e.g. near Cora Cora and Laguna Salinas); 3) The Upper

Barroso 3-1 Ma-old stratovolcanoes and dome complexes, with a wider range of composition from mafic

andesites to rhyolites, have been carved by glacial erosion and abundant scars of flank failures (e.g. Pichu Pichu

and Chachani); 4) The Pleistocene – Holocene volcanoes are composite cones such as El Misti, Ubinas, and

domes on caldera edges such as Ticsani. Most of these composite cones are younger than 0.8-0.6 Ma (Thouret et

al., 2001, 2005). The frontal arc includes coeval monogenetic fields like the Andahua-Orcopampa-Huambo field,

where strombolian cones and lava flows formed between 0.5 Ma and historic times (Delacour et al., 2007).

The 40Ar/39Ar chronology combined with volumes of composite cones allow eruption rate estimation, which

are minimums because of glacier erosion and explosive destructions. Eruption rate is apparently lower during the

first phase of the growth of stratovolcanoes over a long period (400 – 800 ka) and apparently accelerates during

maturation and growth of the summit cones: 0.6 km3/ky at Misti over 110 ka and 0.22 km3/ky at Ubinas over

250 ka. Eruption rates fluctuate between 0.1 and 1 km3/ky according on the time span considered and with

respect to magma composition and eruptive style (mafic effusive vs. evolved and pyroclastic). Composite cones

have changed between Plinian eruptions that form summit calderas (Misti 13-11 ka; Ubinas 25-9 ka). Large

debris avalanches occurred at composite cones and dome complexes during the last 0.5 Ma. The largest collapse

at Ticsani produced a 20 km3 debris-avalanche deposit but smaller, recurring debris avalanches as young as

middle-late Holocene have also occurred at the Ubinas cone and Tutupaca.

The 40Ar/39Ar chronology and petrology of lava flows and pyroclastic deposits allow us to estimate the

magmatic evolution through time (Fig. 3). Andesite and mafic andesite magmatism forms the base of

Figure 3. Nd and 87Sr/86Sr Plot of ignimbrites of the Cotahuasi and Arequipa areas. Isotope values of igneous rocks support the concept that Andean magmas are controlled by the composition and age of the Andean crust. The Arequipa and the Cotahuasi ignimbrites define a domain that overlaps the average CVZ magma composition domain. The Arequipa ignimbrites eNd is lower than the Cotahuasi ignimbrites eNd. These differences may be the result of the assimilation of crustal materials with different isotopic signatures during magma genesis. Recent geochemical and geophysical data pointed out two distinct crustal domains, termed Cordillera and Arequipa, in southern Peru (Mamani et al., 2008a,b).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 545-548

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stratovolcanoes beneath summit cones and are present in monogenetic compound lava flow fields throughout the

region, mostly along deep-seated, normal N80-trending faults (e.g. Ichupampa Fault). The monogenetic field of

Andahua-Orcopampa-Huambo has 25-50 km3 of lava, with an eruption rate of 0.09-0.18 km3/ky. The ascent of

magma producing coeval compound lava fields has bypassed the reservoirs of composite cones in the upper

crust: the magma genesis is attributed to partial melts of the lowermost part of the thick Andean continental crust

added to mantle-derived arc magmas in a high pressure MASH zone (Delacour et al., 2007).

Conclusion: implications on eruption frequency and hazards

From the chronostratigraphy, large-scale ignimbrite sheets (>20 km3) have erupted on intervals of 5 Ma but

many individual smaller ignimbrites have also occurred. Each of the four generations of composite and shield

volcanoes has lasted between 1 and 4 My but this belies the rapid growth of short-lived (<0.8 My) composite

cones, which have erupted at a rate of 0.2–1 km3/ky on average over the past 250 ka.

A 50 ka-long record of identified and dated tephra and lava flows is linked to 10 volcanic edifices and

monogenetic fields between Nevado Sara Sara and Yucumane. The record of the Frontal arc displays at least 50

events, i.e one eruption every kyr over the past 50 ka, including 12 large Plinian-type eruptions with >1 km3 of

tephra. If the more complete 15 kyr-old tephra record is taken at face value, the eruption frequency increases to 3

events per kyr, comprising two moderate-sized ashfalls every kyr and one voluminous Plinian pumice fall on a

2400–3600 yr interval. Very large eruptions such as the Huaynaputina AD1600 event potentially would have a

large effect on southern Peru, western Bolivia, and northernmost Chile. Such eruption could occur in the area

comprised between Huaynaputina, Ticsani and Tutupaca (Fig. 1): in this region, a long volcanic history and

recent eruptions of silicic magma suggest that similar vigorous eruptions may occur in the near geological future.

In addition, debris avalanches and landslides from from ignimbrites cliffs and from hydrothermally-altered

composite cones, even without any eruption, and subsequent debris flows pose serious threats to the population.

References Delacour A., Gerbe M.-C., Thouret J.-C., Wörner G., Paquereau-Lebti P., 2007 Magma evolution of Quaternary minor

volcanic centres in Southern Peru, Central Andes. Bull. Volc, 69, 6, 581-606. Gunnell Y., Thouret J.-C., Brichau S., Carter A., 2008 A low-temperature thermochronology of denudation, crustal uplift

and canyon incision in the Western Cordillera of southern Peru. Geoph. Res. Abs., vol. 10, EGU2008, Vienna. Mamani M., Tassara A., Wörner G., 2008a Composition and structural control of crustal domains in the central Andes. G3,

Geoch., Geoph., Geos., 9 (3) 10.1029/2007GC001925. Mamani M., Wörner G., Thouret J.-C., 2008b “Tracing a major crustal domain boundary based on the geochemistry of

minor volcanic centres in southern Peru”. Extended Abstract, 7th ISAG, Nice, September 2008, this volume. Paquereau-Lebti P., Thouret J.-C., Wörner G., Fornari M., Macedo O., 2006 Neogene and Quaternary ignimbrites in the

area of Arequipa, southern Peru: stratigraphical and petrological correlations. J. Volc. Geoth. Res., 154 : 251-275. Paquereau-Lebti P., Fornari M., Roperch P., Thouret J.-C., Macedo O., 2007 Paleomagnetic, magnetic fabric properties,

and 40Ar/39Ar dating, of Neogene - Quaternary ignimbrites in the Arequipa area, Southern Peru. Flow directions and implications for the emplacement mechanisms. Bull. Volcanol., DOI 10.1007/s00445-007-0181-y.

Roperch P., Sempere T., Macedo O., Arriagada C., M., Tapia C., Laj C., 2006 Counterclockwise rotation of late Eocene-Oligocene fore-arc deposits in southern Peru and its significance for oroclinal bending in the central Andes. Tectonics 25, TC3010.

Thouret J.-C., Suni J., Finizola A., Fornari M., Legeley-Padovani A., Frechen M., 2001 Geology of El Misti volcano near the city of Arequipa, Peru. Geol. Soc. Amer. Bull., 113 : 1593-1610.

Thouret J.-C., Rivera M., Wörner G., Gerbe M.-C., Finizola A., Fornari M., Gonzales K., 2005 Ubinas: evolution of the historically most active volcano in Southern Peru. Bull. Volc., 67 : 557-589.

Thouret J.C., Wörner G., Gunnell Y., Singer B., Zhang X., Souriot T., 2007 Geochronologic and stratigraphic constraints on canyon incision and Miocene uplift of the Central Andes in Peru. Earth Plan. Sci. Letters, 263 : 151-166.

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The Proterozoic basement of the Arequipa massif, southern Peru: Lithologic domains and tectonics

Paul Torres1,2

, Aldo Alván1, & Harmuth Acosta

1

1 INGEMMET, Av. Canadá 1470, San Borja, Lima, Peru (ptorres@ingemmet.gob.pe)

2 Universidad Nacional de Ingeniería (UNI), Av. Tupac Amaru 210, Lima, Peru

KEYWORDS : Arequipa massif, lithologic domains, Proterozoic, tectonic evolution, Grenville

Introduction

The Proterozoic rocks outcrops along the southern coast of Peru, on the margin of the Pacific Ocean and

extends in the Western Cordillera of the Andes, forming a major exhibition of pre-Andean rocks where

tectonothermal activity has been recognized preliminarily throughout Proterozoic history; being Grenville ages

(~1000-1200 Ma) well documented. There are many gaps in knowledge of the Arequipa massif, for example, in

lithological and tectonic part, there are so many authors taking the massif as an undifferentiated complex. We

show in this paper a new Proterozoic basement mapping supported by field works. In addition we raised a

discussion of the tectonic evolution of this massif.

Regional geological setting

The Arequipa Massif is the main basement of central Andes (Wasteneys et al., 1995), displays a magmatic and

metamorphic evolution complex. This massif is mainly made of migmatitic gneiss rocks, thus between Camana

and Mollendo, the occurrence of the mineralogical joint: orthopyroxene-sillimanite-quartz is common in gneiss,

migmatite and granulite rocks, so that they are described as ultra-high-temperature rocks (Martignole et al.,

2003), that apparently is extending through all Proterozoic basement of central Andes between the southern Peru

and northern Chile. The oldest rocks displays ages between 2000-1900 Ma (Wasteneys el al., 1995; Dalmayrac

et al., 1977; Ries, 1976). Preliminary Rb-Sr and U-Pb geochronology implies granulites-amphibolites

metamorphic facies between ~1900 to 1800 Ma (Cobbing et al., 1977; Shacklenton et al., 1979), but recently

U-Pb geochronology analysis in zircons from gneisses near the Mollendo, Quilca and Camaná areas indicates a

high-degree metamorphism between 1200 and 970 Ma, whose ages put in evidence the orogenic-metamorphic

event named Grenville (Wasteneys et al., 1995), alternatively Dalmayrac et al., (1977), proposes that those rocks

underwent two metamorphic events, during the Paleoproterozoic (1950 Ma) and Neoproterozoic (600 Ma) times.

In the Southernmost part of Arequipa massif in Ilo, to the north of the environs of the Arica elbow, there are

outcrops one of the few Proterozoic anorthosite rocks occurrences (it is showing an Sm-Nd model age at

1150 Ma) [(Martignole et al., 2005)] documented in the Andes basement, also demonstrating the orogenic-

metamorphic Grenville belt indicated above. Whereas in the north part of the Arequipa massif, in the locality of

San Juan, there outcrops diamictites, interpreted as glaciers deposits (tillites) of Chiquerio Formation, probably

Neoproterozoico age (Caldas, 1979). Recent analyses of zircons and 13C isotopes from Chiquerío and San Juan

formations that showing an age of deposition for these glaciers deposits (unique in the proto-Andes belt)

between 635-750 Ma (Chew et al., 2007b). At the moment it is thought that Arequipa massif accreted to the

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 549-552

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Figure 1. Map of lithologic domains of the Proterozoic of the Arequipa massif, supported by unpublished field works and completed with INGEMMET maps. Numbers are range of ages at Ma taken from authors mencioned in text.

main Amazonia craton during Sunsas orogeny (~1000 Ma, equivalent in South America to Grenville event)

[Loewy et al., (2004), on the basis of U-Pb geochronology in zircons].

Lithologic domains

The mapping of the metamorphic rocks outcrops was based on domains, and was supported on field works,

where it has been possible to differentiate ten litologic domains showed in the margin from the South coast of

Peru (I) and in the western margin of the Western Cordillera (II) (Figure 1). The predominance of the migmatitic

gneisses and granulites of greater antiquity are

observed, as well as metamorphic rocks from

sedimentary protolite, evidenced by first time

in the western margin of the Western

Cordillera (e.g. in Condesuyos and

Pampacolca, see Figure 1). Also structural

alignments were identified, which they are

removing the main basement from the central

Andes.

Discussion

Ries et al., (1976) was the first that indicates

metamorphism in gneiss rocks of Mollendo in

granulites facies with an age of 1960 ± 33 Ma

then Dalmayrac et al., (1977), indicates two

orogenic-metamorphic events: (1) the first

prograde event, produces gneisses,

characterized by biotite-estaurolite, garnet-

kyanite-sillimanite-potassium feldspar from

relicts associations of a type of average

pressure with cordierite in catazonal

paragenesis on granulites facies, dated at 1950

Ma to age same that Ries et al.; but Dalmayrac

et al. mentioned a metamorphism second more

(2) pressure low, characterized by chlorite-

muscovite-epidote-cordierite, that corresponds

to a epizonal retrograde metamorphism dated

at 600 Ma, whereas Cobbing et al., (1977),

indicates three metamorphic events: (1) the

first event in granulites facies, produced an

extensive area of undifferentiated gneiss rocks

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 549-552

551

dated at 1811 Ma, (2) the second event of sedimentary deposition and subsequent metamorphism produced schist

and gneiss rocks in amphibolites facies dated at 1340 Ma, (3) and the third, a migmatization event that probably

affected to gneiss and schist rocks, and that could be contemporary with the mentioned metamorphic event in

amphibolites facies or it could taken place in later Precambrian or Cambrian times?. The same way Shacklenton

et al., (1979), mentions three metamorphic events: (1) the first denominated Mollendo event in the sillimanite-

gneiss rocks of Mollendo in granulites facies, dated at 1918 Ma, also producing, probably a estaurolite-

andalusite schist rock, (2) followed by a metamorphism denominated Atico event where a series of basic and

acid igneous rocks were intruded and deformed in amphibolites facies, having begun in 679 ± 12 Ma (Stewart et

al., 1974, in the Charcani gneiss) and ending in 440 Ma, (3) and a third event denominated Marcona, it happened

previous erosion of the Atico complex, depositing discordantly to the sediments of the Marcona Formation, with

slight deformation, associated to a metamorphism in green schists facies, dated at 392 Ma.

Figure 2. Summary of the tectonic of the Arequipa massif. Data taken from authors mentioned in text.

But Wasteneys et al., (1995), discusse the age of metamorphism at granulites facies, and it indicated a

younger age, and this is prevailed to the previous metamorphism mentioned before, thus, in Quilca gave an age

of 1198 +6/-4 Ma and 970 ± 23 Ma in the Mollendo area [these ages are according with preliminary ages of

James & Brooks, (1976), in Dalmayrac et al., (1977) that were the firsts that mencioned a Grenville age indicate

a metamorphism in Charcani gneiss at 1012 ± 52 Ma] concluding that the isochronal ones published of 1900 Ma

of Rb-Sr for the gneiss rocks of the Arequipa massif register ages of metagranitoid protolite and were not

affected by the high metamorphism degree, relating them to orogeny of Grenville, furthermore Martignole et al.,

(2003), indicates recently an age of metamorphism at 998 ± 11 Ma for a migmatitic gneiss rock from Camana

that reinforce the presence of Grenville ages on the south coastal of Peru. Protolites of the gneissitic basement in

San Juan and Mollendo crystallized between 1851-1819 Ma and the age of crystallization of the granite of San

Juan was dated at 1793 Ma by U-Pb in zircons (Loewy et al., 2004). This would preliminarily indicate us that a

Paleoproterozoic metamorphic event existed (Figure 2) associated to a magmatism, that affected deepest meta-

sedimentary sequences of the Arequipa massif, producing gneisses in granulites facies, thus it is demonstrated

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 549-552

552

that intrusive ones which cut to the Precambrian gneiss rocks of the San Juan from Marcona. Loewy et al.,

(2004), indicate that the Arequipa massif underwent three different pulses from metamorphism and deformation:

(1) 1820-1800 Ma, (2) 1200-940 Ma and (3) 440 Ma.

Conclusions

Analysis of the evolution of the northern part of Central Andes, in southern Peru and northern Chile shows a

magmatic and metamorphic polycyclic evolution in Proterozoic time, with a magmatism-metamorphism event at

~1000 Ma (Mesoproterozoic later) associate to metamorphism in granulites facies demonstrated in San Juan and

Mollendo, and an acid magmatism in San Juan and metamorphism in Camana-Mollendo and San Juan in the

Mesoproterozoic time, these evidences would be according to the accretion of the massif to the Amazonia

craton. The radiometric analyses carried out in Arequipa massif indicate three groups of well differentiated ages

in all the Proterozoic (Figure 2). The lithologic domains (Figure 1) show the first mapping and lithologic

division of the Arequipa massif and variety of the rocks from this massif as well as its relation with the

metamorphic degrees.

References Caldas J. (1979). “Evidencias de una glaciación Precámbrica en la costa sur del Perú”. Segundo Congreso Geológico

Chileno, Arica. p. J-29 a J-38. Chew, D., Kirkland, C., Schaltgger, U., Goodhue, R. (2007b). “Neoproterozoic glaciation in the Proto-Andes: Tectonic

implications and global correlation”. Geology, v.35, n.12, p.1095-1098. Cobbing, E., Ozard, J., Snelling, N. (1977). “Reconnaissance geochronology of the basement rocks of the Coastal Cordillera

of southern Perú”. Geological Society of American Bulletin, v.88, p. 241-246. Dalmayrac, B., Lancelot, J., Leyreloup, A. (1977). “Two-Billion-Year Granulites in the late Precambrian metamorphic

basement along the southern peruvian coast”. Science, v. 198, p. 49-51. Loewy, S., Connelly, J., Dalziel, I. (2004) “An orphaned basement block: The Arequipa-Antofalla Basement of the central

Andean margin of South America”. GSA Bulletin, v. 116, p. 171-187. Martignole, J., and Martelat, J. (2003). “Regional-scale Grenvillian-age UHT metamorphism in the Mollendo-Camana block

(basament of the Peruvian Andes)”, J. Metamorphic Geol. v. 21, p. 99-120. Martignole, J. Stevenson, R., & Martelat, J. (2005). “A Grenvillian anorthosite-mangerite-charnockite-granite suite in the

basement of the Andes: The Ilo AMCG suite (southern Perú)”, 6th International Symposium on Andean Geodinamics ISAG, Barcelona, Extend Abstacts: p. 481-484.

Ries, A. (1976). “Rb/Sr ages from the Arequipa Massif, southern Peru”. Instituto de estudios africanos, Boletín de la Universidad de Leeds.

Shackleton, R., Ries, A., Coward, M., Cobbold, P. (1979). “Structure, metamorphism and geochronology of the Arequipa massif of coastal Peru”. J. Geol. Soc. London 136:195—214

Stewart, W., Everden, J., Snelling, N. (1974). “Age Determinations from Andean Peru: A Reconnaissance Survey“, Geological Society of American Bulletin, v. 85, p. 1107-1116.

Wasteneys, H., Clarck, A., Farrar, E. & Langridge R. (1995). “Grenvillian granulite-facies metamorphism in the Arequipa Massif, Peru: a Laurentiia-Gondwana link”. Earth and Planetary Science Letters 132, p. 63-73.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 553-554

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Trachydacitic domes in the caldera of Pino Hachado, province of Neuquén, Argentina

Cynthia Tunstall1, Jorge E. Clavero

2 , & Victor A. Ramos

1

1 Laboratorio de Tectónica Andina, FCEyN, Universidad de Buenos Aires, Consejo Nacional de Investigaciones

Científicas y Técnicas, Argentina (c_tunstall@gl.fcen.uba.ar; andes@gl.fcen.uba.ar) 2 Sevicio Nacional de Geología y Minería, Avda. Santa María 0104, Santiago, Chile (Jclavero@sernageomin.cl)

KEYWORDS : domes, Pino Hachado, Neuquén Andes

Introduction

The existence of the caldera in the Pino Hachado region has been known in the literature since the work of

Muñoz and Stern (1985) & Mazzoni & Iñiguez Rodríguez (1986).

The Pino Hachado caldera is located between the 38o and 39o SL in the Main Cordillera of Neuquén on the

border of Argentina and Chile (Neuquén and Cautín provinces) in a retroarc geotectonic setting, which occurs

approximately 60 kms to the east of the present volcanic arc (figure 1).

Figure 1: Location map and satelite image of the Pino Hachado caldera.

Andesitic eruptions began in the Paleogene in this sector of the Andes but their major development took place

during the Neogene. The Pino Hachado volcanic centre was generated during the last retroarc volcanic episode

and is composed of the Pino Hachado caldera, in which ignimbrites and breccias dacitic and rhyolitic

composition were generated in an initial stage. Subsequent trachyandesitic lava flows were generated in a later

stage. Trachyandesitic domes are only found inside of the caldera (Muñoz & Stern, 1985; Muñoz & Stern 1988;

Muñoz, 1988b; Muñoz et al., 1989).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 553-554

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Previous works have allowed the definition of the principal units of the Pino Hachado postcaldera. Five formal

units have been proposed for Plio-Quaternary volcanism: Pailaleo Andesite, Litrán Basalt, Suarzo Basalt, Las

Tres Hermanas Andesites and El Volcán Basalt (Tunstall 2005).

The presence of numerous trachydacitic domes in the intracalderic region are the predominant volcanic

feature, among with the Cerros Las Tres Hermanas are the most outstanding examples where well-developed

conic form occurs without evidence of glacial erosion. They show columnar disjunction, a succession of flows

and maffic inclusions. Lithologically, they have porphyritic textures with groundmasses of hyalopylitic textures.

Phenocrysts (40 %) are composed of plagioclase, biotite, pyroxene and less amount of opaque minerals that

occasionally shows hexagonal sections.

According to the geochemical data these rocks have trachydacitic compositions in accordance to the diagrams

of Le Maitre (2002) and Winchester & Floyd (1977). The SiO2 and Na2O+K2O compositions vary between 65-

75% and 9-10% respectively, and classify in the class of subalkaline rocks (Irvine & Baragar 1971). The

calcalkaline signature come from the amount of alkalis, FeO* and MgO that evidence the whole sequence.

Conclusions

Those three eruptive centres are recognised independently as a postcaldera volcanic activity and their relative

ages has been established. Their similar petrography and geochemical signatures suggest that they should have

the same magmatic origin. The following contribution described new alkaline units of trachydacitic composition

inside the Las Tres Hermanas Andesites formation. These volcanic events contributes to a better understanding

of the magmatic evolution of the Pino Hachado caldera.

References Irvine, T. N. y Baragar, W. R. A., 1971. A guide to the chemical classification of the common volcanic rocks. Can. J. Earth

Sci., 8 (5): 523-548. Le Maitre, R. 2002. Igneous Rocks. A classification and glossary of terms. Recommendations of the Intarnational Union of

Geological Sciences, Subcomisión on the Systematics of Igneous Rocks. 2nd edition. Cambridge University Pres. 236 p. Mazzoni, M. e Iñiguez Rodríguez, A., 1986. Depósitos piroclásticos neógenos y cuaternarios en el área de Pino Hachado,

Prov. del Neuquén. 1º Reunión Argentina de Sedimentología (La Plata) Resúmenes 97-100, La Plata. Muñoz, J., 1988b. Evolution of Plioceno and Quaternary volcanism in the segment of the southern Andes between 38º and

39º S. University of Colorado (Unpublished Thesis), 160. Muñoz, J. y Stern, C., 1985. El complejo volcánico Pino Hachado en el sector nor-occidental de la Patagonia (38º-39ºS):

volcanismo plio-cuaternario de trasarco en Sudamérica. IVº Congreso geológico Chileno (Antofagasta), Actas 3: 381-412, Antofagasta.

Muñoz, J. y Stern, C., 1988. The Quaternary volcanic belt of the southern continental margin of South America: Transverse structural and petrochemical variations across the segment between 38° and 39°S. Journal of South American Earth Science 1 ( 2): 147-161

Muñoz Bravo, J., Stern, C., Bermúdez, A., Delpino, D., Dobbs, M.F., y Frey, F. A., 1989. El volcanismo plio-cuaternario a través de los 38º y 39ºS de los Andes. Revista de la Asociación geológica Argentina, 44: 270-286. Buenos Aires.

Tunstall, C., 2005. Geología de la caldera de Pino Hachado. Trabajo Final de Licenciatura. Universidad de Buenos Aires (inédito). Buenos Aires.

Winchester, J. A. y Floyd, P. A. (1977). Geochemical discrimination of different magma series and their differentiation products using immobile elements. Chemical Geology, 20:325-343.

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Controls on erosion and clastic sediment flux in the Central Andes during the Late Cenozoic

Cornelius E. Uba, Gerold Zeilinger, Manfred Strecker Institut für Geowissenschaften, Universität Potsdam, Karl-Liebknechtstr. 24, 14476 Potsdam, Germany

(uba@geo.uni-potsdam.de)

Introduction

The Central Andes of south-central Bolivia is an integral part of the Andean orogenic system that is related to

the basement-involved shortening, uplift, thrust loading, and the subsequent eastward propagation of the Andean

deformation. The study area we examine here lies in the Chaco foreland basin that consists of the Subandean

Zone and the Chaco plain (Fig. 1; Uba et al., 2005). The basin development was as a result of the interaction of

the Nazca and the South American plates and its related simultaneous under-thrusting of the Brazilian Shield.

This activity led to widespread and pronounced shortening in the Eastern Cordillera in the Oligocene, which

produced folding and eastward migration of thrusting in the Interandean and Subandean Zones (e.g., Gubbels et

al., 1993; Uba et al., 2006). The study area lies within three major river catchments (rios Grande, Parapeti, and

Pilcomayo).

To unravel the controls on erosion and sediment flux in the Andes, we use isopach maps (well logs, seismic

lines, and measured sections) and recently published zircon U-Pb age data from Mio-Pliocene sedimentary strata

(Uba et al., 2007) to produce a 2D mass flux budget for the central Andes.

64°W 63°W

Quaternary

Rio

Pilcom

ayo

20°S

21°S

0 50

km

25

Ordovician

Silurian

Measured sections

Devonian

Carboniferous

Cretaceous

Tertiary

Neogene tuffs

and lava flows

RioPara

pe

ti

Villamontes

28

24

20

16

12

74 70 66

Salta

La Paz

Chaco

Plain

EasternC

ord

illera

Altip

lan

o

Puna

Su

ba

nd

ea

nZ

on

e

Western Cordillera

Figure 1. Geological map of the study area in the southern Bolivia showing the measured sections.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 555-557

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Geological setting

Many authors have previously postulated that the deformation front arrived in the Chaco foreland basin at

10 Ma (e.g., Gubbels et al., 1993). Recently, however, Uba et al. (2007) present ~12.4 Ma for the arrival of the

deformation front into the Subandean Zone. The Chaco foreland basin is characterized mostly by in-sequence,

thin-skinned thrusting, which include ramp anticlines and passive roof duplexes [Baby et al., 1992] separated by

thrust faults and synclines. Blind thrusting is another documented structural style in the basin (Uba et al., 2006).

The Chaco basin consists of up to 7.5 km thick late Cenozoic strata. The base of the foreland stratigraphy is

defined by the 250-m-thick, 26 - 12.4 Ma Petaca Formation. This formation consists of paleosol, conglomerate,

sandstone, and mudstone, which accumulated in a fluvial environment (Uba et al., 2005). Overlying the Petaca

Formation is the ~400-m-thick, 12.5-8 Ma Yecua Formation (Uba et al., 2007). This lacustrine-fluvial-shallow-

marine strata compose mostly of mudstone and sandstone deposited in a distal foredeep (Uba et al., 2005; 2006).

The Yecua Formation is overlain by the up to 4000-km-thick, 8 to 6 Ma, fluvial-megafan-deposited sandstone

and mudstone Tariquia Formation, which represents medial foredeep deposit. The up to 1600-m-thick,

5.94-2.1 Ma Guandacay Formation overlies transitionally the Tariquia Formation. The Guandacay Formation

consists of sandstone, conglomerate, and subordinate mudstone deposited in a medial fluvial megafan

environment (Uba et al., 2005; 2007). The late Cenozoic succession is capped by more than 1500-m-thick

2.1 Ma to present conglomerate and sandstone dominated Emborozú Formation that represents a proximal fluvial

megafan (Uba et al., 2005; 2007).

Results and conclusions

Our mass accumulation budget result shows that a total of 262,637 km3 of clastic sediment was deposited in

the Chaco Basin between ~26 Ma and the present (Fig. 2). Of this volume, 4200 km3 of sediment, representing

309 km3/Ma sediment supply rate, was deposited during the Petaca time. During 12.4-8 Ma (Yecua Fm) and

8-6 Ma (Tariquia Fm) 37,011 (sediment supply rate: 10,280 km3/Ma) and 103,322 km (51,661 km /Ma) of

sediments were deposited. Finally, the 6-2.1 Ma (Guandacay Fm) and 2.1 to present (Emborozú Fm) time slices

recorded 81,363 km (20,862 km /Ma) and 36,741 km (17,495 km /Ma) of sediment volume respectively.

Fig. 2. 3D model showing the late Cenozoic Chaco foreland basin geometry and configuration and the summary of sediment budget estimates for the five different late Cenozoic stratigraphic units.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 555-557

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These results show a continuous increase in sediment volume from the late Oligocene till its maximum during

the late Miocene, when a three-fold (103,322 km ) increase in the sediment supply and corresponding five-fold

(51,661 km /ma) increase in sediment supply rate is observed. Interestingly, the three-fold increase in mass flux

is consistent with rapid increase in sedimentation rate during the same period documented by Uba et al. (2007).

In addition, Uba et al. (2006) suggest that the basin witnessed accommodation space creation between 12-8 Ma

as a result of tectonic loading and coeval eastward advance of thrusting. The created space was then filled with

clastic sediments as a result of enhanced erosion during 8-6 Ma (Tariquia Fm) interval.

Using sediment flux as a proxy for continental erosion, our result shows that the eastern part of the central

Andes witnessed enhanced erosion and deposition during 8-6 Ma. This enhanced erosion from the late Miocene

to Pliocene is supported by apatite fission track thermochronometry (Ege, 2004; Barnes et al., 2006).

Furthermore, the increase in sediment discharge and enhanced erosion coincides with the deposition of

progradational sequences, changing from lacustrine-shallow marine mudstone to fluvial sandstone.

In addition, the rapid increased erosion and deposition and high erosion rate observed during the 8-6 Ma

interval after a low at 12.5-8 Ma reflect probably the reorganization of the paleo-drainage systems from small

rivers to the proto-Rios Grande, Parapeti, and Pilcomayo fluvial megafans. The change in river capture was

probably driven by stronger, wetter South American monsoon, which brought moisture to this previously semi-

arid part of the Andes (Strecker et al., 2007).

Although we recognize the potential importance of tectonic in influencing high sediment discharge and

erosion in the central Andes, however, apatite fission track thermochronologic and paleosurface studies in the

Eastern Cordillera and the Interandean zone show that the structures there were inactive during this time (Ege,

2004; Barke and Lamb, 2006). The interpretation of inactive structures is further supported by the enhanced

erosion during this time, which might have led to the retardation of deformation in the Subandean zone.

References Barke, R. & Lamb, S. 2006. Late Cenozoic uplift of the Eastern Cordillera, Bolivian Andes. Earth and Planetary Science

Letters 249: 350-367. Baby, P., Hérail, G., Salinas, R. & Sempere, T. 1992. Geometry and kinematic evolution of passive roof duplexes deduced

from cross-section balancing: Example from the foreland thrust system of the southern Bolivian Subandean Zone. Tectonics 11: 523-536.

Ege, H. 2004. Exhumations- und Hebungsgeschichte der zentralen Anden in Südbolivien (21°S) durch Spaltspur-Thermochronologie an Apatit. Ph. D. Thesis, Freie Universität Berlin, Berlin, 159 p.

Gubbels, T.L., Isacks, B.L. & Farrar, E. 1993. High-level surface, plateau uplift, and foreland development, Bolivian central Andes. Geology 21: 695-698.

Strecker, M.R., Alonso, R.N., Bookhagen, B., Carrapa, B., Hilley, G.E., Sobel, E.R., & Trauth, M.H. 2007. Tectonics and climate of the southern central Andes. Annual Review of Earth and Planetary Sciences 35: 747-787.

Uba, C.E., Heubeck, C. & Hulka, C. 2005. Facies analysis and basin architecture of the Neogene Subandean synorogenic wedge, southern Bolivia. Sedimentary Geology 180: 91-123.

Uba, C.E., Heubeck, C. & Hulka, C. 2006. Evolution of the late Cenozioc Chaco foreland basin, Southern Bolivia. Basin Reserach 18: 145-170.

Uba, C.E., Strecker, M.R., & Schmitt, A.K. 2007. Increased sediment accumulation rates and climatic forcing in the central Andes during the late Miocene. Geology 35: 979-982.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 558-561

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Diente Verde and Mario, Cañada Honda, San Luis, Argentina: Porphyry-type deposits in the South Pampean flat-slab region of the Central Andes

Nilda E. Urbina1 & P. Sruoga

1,2

1 Universidad Nacional de San Luis, Ej. de los Andes 950, 5700 San Luis, Argentina (urbina@unsl.edu.ar)

2 CONICET, SEGEMAR, Av. J.A. Roca 651, 1322 Buenos Aires, Argentina (patysruoga@yahoo.com.ar)

KEYWORDS : San Luis, porphyries, gold-copper, flat-slab, Mio-Pliocene

Introduction

The San Luis Tertiary Metallogenic Belt (SLMB) located in the Sierras Pampeanas of San Luis is related with

the subduction zone shallowing between 27º and 33º S (Fig. 1 A). Mineralization and volcanic rocks occur

within a west-northwest-trending magmatic belt from La Carolina in the west to El Morro in the east (Fig. 1 B).

Mesosilicic magmas belong to normal to high-K calc-alkaline and shoshonitic types. In a close spatial and

temporal linkage several mineralizations of epithermal and porphyry types are associated. Volcanic activity began at

about 12-13 Ma in the west and ended at 1.9 Ma in the east (Ramos et al. 1991, Urbina, 2005, Urbina y Sruoga,

In press). Volcanics and associated mineralization formed 600-700 kilometers east from the trench for about

10 m.y. and over a west-east distance of 80 kilometers.

The Diente Verde and Mario deposits are copper-gold porphyry mineralizations genetically related to the Late

Miocene-Late Pliocene volcanic activity. Both deposits are located at Cañada Honda district and are part of an

arc-transverse magmatic lineament at 33º S coincidently with the change of the subduction angle.

Diente Verde deposit

An intrusion centered district was suggested by Urbina et al. (1997) for Cañada Honda, based on the spatial

distribution of several low-sulfidation epithermal veins with regard to the high-level stock at Diente Verde.

Diente Verde is a gold-copper porphyry deposit consisting of stockwork sulfide veining associated with a small

intrusion centered within an andesitic stratovolcano (Fig. 1 C). Hydrothermal alteration and mineralization have

a symmetrical distribution surrounding the porphyritic subvolcanic intrusion. The sulfide ore-mineral

assemblage occurs either disseminated or in a stockwork. The alteration affects the rocks of the core, and spreads

outside from the volcanic edifice in an extensive hydrothermal halo (Fig. 1 C). K silicate alteration present in the

central zone is characterized by the presence of quartz veinlets that may occur as multidirectional stockwork or

subparallel arrays of closely spaced WNW-ESE, NE-SW, NW-SE-striking veinlets suggesting structural control

on emplacement. Veinlets from a few millimeters to 5 cm in width, are planar to slightly sinuous, and have

alteration halos composed mainly of biotite replacements accompanied by hydrothermal K feldspar. Quartz

veinlets have centrally located sulfides grains. Chalcopyrite and magnetite are the principal hypogene minerals

in K silicate alteration with traces of electrum, digenite, bornite, tennantite, covellite, enargite, pyrite and

pyrrhotite, which are A-type veinlets (Gustafson and Hunt, 1975) (Fig. 2). Veinlets with central suture of

chalcopyrite belonging to B-type (Gustafson and Hunt, 1975) are also present although in lesser amounts

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 558-561

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(Fig. 2) (Suarez Funes, 2007). Outward from the central core phyllic and argillic alteration mineral assemblages

characterized by sericite, illite, quartz, kaolinite, smectite, albite, chlorite, ilmenite, pyrite, rutile and specularite

affect mainly volcanic breccias and andesite flows and overprint the early K silicate alteration. Magnetite from K

Figure 1. Simplified maps showing: A) location of the San Luis Metallogenic Belt in the Central Andean flat-slab region, B) regional distribution of Tertiary volcanic rocks in the Sierra de San Luis, and C) location of Diente Verde and Mario deposits at Cañada Honda district. Note the symmetrically centered distribution of hydrothermal alteration surrounding the volcanic edifice.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 558-561

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silicate alteration is variably martitized. In the outer alteration zone, abundant disseminated pyrite essentially

without chalcopyrite defines a peripheral, broad pyritic halo. Propylitic alteration is present as a discontinuous

outer aureole. Supergene alteration is represented by goethite, hematite, malachite and azurite. The gold-copper

ore zone coincides with the most densely stockworked porphyry averaging approximately 0.93 ppm Au and

7200 ppm Cu (Suarez Funes, 2007). Considering that the average gold content is 0.4 ppm, the Diente Verde

porphyry copper deposit may be defined as gold-rich (Sillitoe, 2000).

Mario deposit

Mario deposit is located 1200 m east from Diente Verde gold-copper porphyry deposit and was emplaced

approximately at the same topographic level. The outcrop of Mario is restricted to a 180-m long bulldozer

trench, with most of the inferred extent of the deposit concealed beneath the present surface. The mineralization

hosted by hornblende-bearing andesitic rock occurs as a multidirectional stockwork and disseminations (Arce et

al., 2005). Chalcopyrite with traces of bornite are coincident with the stockwork (Fig. 3). Abundant pyrite is

present in dominantly disseminated form. Hematite and magnetite occur abundantly disseminated in altered

rock, and alone as M-type veinlets (Clark and Arancibia, 1995) or with quartz. Much of the hematite is

developed by hypogene martitization of hydrothermal and accessory magmatic magnetite, although minor

specular hematite is also present. The associated hydrothermal alteration is intense and comprises K silicate

alteration (biotite accompanied by K feldspar) and peripheral propylitic (chlorite, calcite, epidote) assemblages.

Intermediate argillic (sericite, chlorite, calcite, smectite) and sericitic (sericite, quartz, pyrite) alterations are

believed to have overprinted and partially destroyed earlier K silicate alteration. Much of the hydrothermal

quartz is present as stockwork veinlets that range from 0.2mm to 3cm in width. They were formed along with

biotite-rich K silicate alteration. Chalcopyrite + pyrite ± magnetite are also thought to have been introduced with

early K silicate alteration and constitute, together with quartz, the A-type veinlets (Gustafson and Hunt, 1975)

(Fig. 3). Supergenic products are present as iron oxides (goethite and hematite) and copper carbonates (malachite

and azurite), in the form of coatings and as complete replacements of sulfides from quartz-veinlets stockwork in

the upper part of Mario. The gold contents averaging 0.3ppm Au, correlate well with the intensity of A-type

quartz veinlets, whereas the copper contents range from 100 to 200ppm (Vazquez, 2007).

Figure 2. Left: Quartz A-type veinlet with K silicate alteration halo. Rigth: B-type veinlet with chalcopyrite suture. bt: biotite, kfs: k feldspar, mt: magnetite, qtz: quartz, cp: chalcopyrite.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 558-561

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Figure 3. Left: Quartz veinlet stockwork. The veinlets are highlighted by goethite and hematite of supergene origin. Rigth: A-type veinlet of quartz (1mm width) with magnetite and chalcopyrite. Note the biotite in the K silicate halo.

Conclusions

The superficial gold-copper contents of Diente Verde and Mario porphyry mineralizations make the deposits

an interesting target for exploration. Besides, the remnants of K silicate alteration and the strongly intermediate

argillic/sericitic overprinted alteration at Diente Verde and Mario suggest that the present level of erosion is

showing the shallow parts of the copper-gold deposits and can be considered as evidence to expect a well

developed K silicate alteration at deeper levels. On the other hand, mineralogical affinities, and spatial and

temporal linkages suggest that Diente Verde and Mario deposits are genetically associated. Therefore, this aspect

might lead to assume the existence of cluster-type porphyry manifestations at Cañada Honda district.

The Diente Verde and Mario porphyry deposits are in SLMB, which reflects the extraordinary broadening of

the magmatic arc in the flat-slab segment at 33ºS. The unusual setting of SLMB at 600-700km from the trench

and the high cortical levels of emplacement, suggest a structural control in the magma ascent coincidently with

the change of the subduction angle, a situation that seems to be alike to the Farallón Negro district at 27ºS.

References

Arce, M., Urbina, N., Sruoga, P. 2005. “A new porphyry-type mineralization in Cañada Honda district, San Luis, Argentina”. In: 19th Colloquium on Latin American Geosciences. Potsdam, Germany. Terra Nostra 05/1: 13.

Clark, A., and Arancibia, O. 1995. “Occurrence, paragenesis and implications ofmagnetite=rich alteration=mineralization in calc=alkaline porphyry copper deposits”. In Clark, A. (ed) Giant ore deposits-II. Kingston, Ontario: 511-581.

Gustafson, L., and Hunt, J. 1975. The porphyry copper deposit at El Salvador, Chile. Economic Geology, 70: 857-912. Isacks, B. 1988. Uplift of the central Andean plateau and bending of the Bolivian Orocline. Journal of Geophysical Research

93: 3211-3231. Ramos, V., Munizaga, F. y Kay, S. 1991. “El magmatismo Cenozoico a los 33ºS de Latitud: Geocronología y Relaciones

Tectónicas”. In: 6º Congreso Geológico Chileno. Chile. Actas 1: 892-896. Sillitoe, R. H. 2000. “Gold-rich Porphyry Deposits: Descriptive and Genetic Models and Their Role in Exploration and

Discovery”. In Hagemann, S. and Brown, P. (eds) Gold in 2000. Society of Economic Geologists, Reviews in Economic Geology 13: 315-345.

Suarez Funes, L. 2007. Geología y metalogénesis del sector suroeste del Cerro Diente Verde, San Luis, Argentina. Tesis de Licenciatura, Universidad Nacional de San Luis (inédito), 80 p.

Urbina, N. 2005. “New insights into the timing of gold systems in the Tertiary metallogenic belt of San Luis, Argentina”. In: 6th International Symposium on Andean Geodynamics. Barcelona. Actas: 752-755.

Urbina, N.E. and Sruoga, P. In press. “K-Ar mineral age constraints on the Diente Verde porphyry deposit formation, San Luis, Argentina”. In: VI South American Symposium on Isotope Geology. 2008, San Carlos de Bariloche. Argentina. 4 p.

Urbina N. E., Sruoga P. and Malvicini L. 1997. Late Tertiary Gold-Bearing Volcanic Belt in the Sierras Pampeanas of San Luis, Argentina. International Geology Review. 39 (4): 287-306.

Vazquez, S. 2007. Geología y metalogénesis del sector NE del Cerro Diente Verde y mineralizaciones asociadas, Cañada Honda, San Luis, Argentina. Tesis de Licenciatura, Universidad Nacional de San Luis (inédito), 93 p.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 562-565

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Relationship between topography and seismicity in the Peruvian Andes: Influence of topography on stress field

V. Manuel Uribe1, Laurence Audin

2, Hugo Perfettini

2, & Hernando Tavera

3

1 Universidad Nacional Mayor de San Marcos, Av. Venezuela 3400, Lima 1, Peru (vm_uribe@yahoo.es)

2 Institut de Recherche pour le Développement, Teruel 357, Lima 18, Peru

3 Instituto Geofísico del Perú, 4 etapa de Mayorazgo, Ate Vitarte, Lima, Peru

KEYWORDS : seismicity, topography, Andes, Peru, tectonic yield stress

Abstract Seismicity in Peru and the Andean Orogeny are two direct effects of the subduction process that occur for My

between the Nazca Plate and the South American Plate, eventhough both of them differ in time scale. During the inter-seismic period the background seismicity in Peru show a complex spatial distribution. We observed that the seismic activity anti correlates with the highest topography. As demonstrated by Bollinger et al., 2004 in Nepal, but in this case for a subduction zone, the stress field varies with topographic loading of the upper plate on the lower plate. Effective diminution until extinction of seismicity below the higher Andes (>2000m) is associated with a notable change in the state of stress, from compression to extension and also correlates with the fault kinematics on the upper plate.

Introduction

The seismic activity in Peru is heterogeneous with wide-ranging spatial and time distributions. We used the

seismic catalog provided by the Peruvian Institute of Geophysics (IGP: 1982-2005 and its 2007 update) (+ F.

Grange data set) and/or Tele-seismic data provided by the Harvard CMT catalog. We added the corrections

made by Engdahl & Villaseñor (2002). The IGP- Engdahl data set has 34088 seismic events with magnitude

range between 1-7.7 ML. The second data set was taken from F. Grange (1984), who implemented a dense

temporal net (~43 local stations) for the period 1980 – 1981. We used 888 events in total with superficial and

intermediate depths (<300 Km) between 13°30’S and 17°30’S, to take in consideration the changing slab

geometry from plane to normal in this special area. We used two data sets to analyze the focal mechanisms.

Tavera & Bufom (2001) create a data set used to study and analyze earthquakes in Peru during 1990-1996. They

took as reference 19 characteristic earthquakes in zones of high seismicity and these correspond to average

hypocenter parameters in the main seismogenic zones. The second data set is the Harvard Centroid-Moment-

Tensor (CMT). We used the global seafloor topography provided by Sandwell & Smith (1997) for topography

modeling.

The western edge of the South American Plate (from the trench to the western Cordillera along the Peruvian

fore-arc) shows compressive focal mecanisms which are typical from active subduction zones. In the sub-

Andean zone, the seismicity shows compressional pattern in the superficial part and extensional patterns below

(60-300 Km). Between these two areas, the high Cordillera of the Central Andes is present and the upper plate

shows extensional active faults whereas a net decrease in seismic activity can be noticed below the main relief.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 562-565

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Discussion

Bollinger et al. (2004), showed the influence of the topography on seismic activity during the inter-seismic

period in the Himalayas. He analyzed seismic data from the regional catalog and modeled the Coulomb stress.

Our study tries to demonstrate whether these parameters have a similar relationship in the Cordillera of the

Andes during the inter-seismic period and if the type of Coulomb modelling could be done in the Andes. The

Magnitude of completeness is Mc= 3.9 ML for the IGP data set. We did a decluster process to identify the

seismic crisis and aftershocks and substract it of the studied catalog.

Figure 1. Comparison between the Himalayan Model (after Bollinger et al., 2004) and the Andean Model. a) Microseismicity recorded between April 1995 and April 2000 by Nepal Seismological Center, Department of Mines of Geology. Ml 3.0. The grey band and red lines present the assumed location of the locked portion of the fault and the location of the 3500 m contour line (DEM-Gtopo30/USGS), respectively. B) Recorded seismicity between January 1982 and December 2005 (Peruvian Institute of Geophysics – IGP). The black dots represent superior events to 3.9 Ml. Focal mechanisms after Tavera & Buforn (2001). The grey section on the geological profile corresponds to the locked part between the two plates and red line present the location of the 2000 m contour line (Sandwell & Smith, 1997).

The seismic activity was separated in three zones:

1) Fore-arc zone, where the focal mechanisms showed a compressive tendency (feature commonly found in

active subduction zones) and this compressive regime dominates the state of stress. It generates that the first

stress component ( 1) is oriented ENE-WSW to E-W along the convergence boundary.

2) Sub-Andean zone (<60 Km) and the Amazone flat terrain. Here, focal mechanisms show reverse/thrust with

orientation E-W and ENE-WSW parallel to the Andes where the first stress component ( 1) has an orientation E-

W similar to the fore-arc zone. This compressive process here is mainly associated to the Brazilian convergence

shield beneath the Eastern Cordillera (Suárez et al., 1983).

a)

b)

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3) Intermediate depths (60-300 Km) beneath the sub-Andean zone where extensional processes are deduced

from the analyse of focal mechanisms. This is due to the “detachment” or separation of the oceanic slab by the

gravity effect. Their axis are oriented NW-SE.

The stress produced by the plate convergence and the presence of the Andean mountain range play a role in

seismicity generation near the trench. In the fore-arc and back-arc zones, some continuity exists in the seismicity

with a quite dense occurrence of events. However, beneath the high Andes we notice the extinction of the

seismic activity, either in cross section as presented here or in map view, following the irregular geometry of the

mountain chain piedmont which roughly correspond to the 2000 meters isocontour (Figure 1).

Because the vertical stress increase with the lithosphere thickness and weight, therefore at greater depths the

vertical stress is larger (Turcotte & Schubert, 2002). So, when it increases below the high Andes, it may reach

and equals the compressive horizontal stress, resulting from the collision between the two plates.

To show in a better way how the topography model induces the seismicity decrease, we have quantified the

vertical stress deviation produced over the 2000m level in ~53 MPa. This result is similar to the value

determined by Bollinger et al. (2004), in the Himalayas (~35 MPa, Figure 2).

Figure 2. Model of the topographic cross section, geologic profile (scale 1:3) and seismic profile (scale 1:1) showing the lithosphere weight effect of the higher Andes (>2000 m). The focal mechanisms are issued from Harvard CMT catalog.

In the fore-arc and back-arc zones (>2000m), the deformation process has a main compressive component with

lithospheric volumes much lesser than the high Andes (>2000m), where the main stress tensor component is

vertical ( 3). Therefore when the vertical stress ( 3) increases in the high Andes and overcomes the horizontal

stress ( 1), the tectonic process becomes extensional generating a zone where they are similar in magnitude. The

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lithosphere weight generates a tectonic compensation which in turn produces the observed extinction in the

seismicity.

References Bollinger L., Avouac J.P., Cattin R., Pandey M.R., 2004. Stress buildup in the Himalaya. Journal of Geophysical Research,

109-B11405 Engdahl, E.R., Villaseñor A. 2002. Global Seismicity: 1900–1999, in W.H.K. Lee, H. Kanamori, P.C. Jennings, and C.

Kisslinger (editors), International Handbook of Earthquake and Engineering Seismology, Part A, Chapter 41, 665–690. Grange, F. 1984. Etude sismotectonique detaille de la subduction lithospherique au Sud-Pérou, Ph.D. Thesis, IRIGM,

Grenoble, France. Manrique M.O., 2003. Estimación del espesor de la corteza continental en el centro y sur del Perú a partir de fases PmP.

Compendio de trabajos de investigación CNDGIGP, Lima, 9p. Smith W. H., Sandwell D.T. 1997. Global sea floor topography from satellite altimetry and ship depth soundings. Science,

277: 1956-1962 Suárez O., Molnar P., Burchfiel C., 1983. Seismicity, fault plane solutions, depth of faulting and active tectonics of the

Andes of Peru, Ecuador and Southern Colombia. Journal of Geophysical Research, 88: 10403-10428. Tavera H. Buforn E., 2001. Source Mechanism of Earthquakes in Peru. Journal of Seismology, 5: 519-539. Turcotte D. L. & Schubert G., 2002. Geodynamics. Cambridge University Press, 2002, 406 p.

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The Peruvian Pataz, Parcoy and Huachón districts: Evidence for a coherent, 400 km-long, Carboniferous orogenic gold belt along the Eastern Andean Cordillera?

Edina Vágó & Robert Moritz

Section des Sciences de la Terre, Université de Genève, Rue des Maraîchers 13, 1205 Geneva, Switzerland

(Edina.Vago@terre.unige.ch, Robert.Moritz@terre.unige.ch)

KEYWORDS : gold, orogenic, fluid, geochemistry, isotopes

Abstract The Eastern Peruvian Cordillera is the host of a major Carbonifeours belt of shear-zone hosted, auriferous

quartz veins. The major mining districts are Pataz and Parcoy in the north, and the Huachón district located 400 km to the south is interpreted as the southern extension of this belt. Previous investigations have interpreted these auriferous vein systems as orogenic gold deposits, however there is still an open debate about any possible magmatic link.

The gold-bearing veins are emplaced along NNW-oriented and NE-E dipping brittle-ductile shear zones, within or along the western margin of a granodioritic batholith. Our preliminary observations indicate that the orebodies share similar structural, paragenetic and hydrothermal alteration characteristics in all three districts. Typically, an early milky quartz-pyrite-arsenopyrite stage is followed by blue-grey quartz-galena-sphalerite-native gold, and a final barren calcite-quartz stage, accompanied by sericite, chlorite and carbonate alteration of the host rocks. The similarities support the existence of a ~400 km long gold belt along the Eastern Cordillera.

Detailed structural, fluid inclusion, isotopic and radiometric studies will compare ore forming events along this major gold belt, and address the issue about the controversial relationship of the gold deposits with any contemporaneous magmatic event, which still remains to be identified.

Figure 1. Situation of Pataz, Parcoy and Huachón districts along the Eastern Cordillera in the peruvian Andes (Haeberlin et al., 2002/b).

Figure 2. Schematic geological map of the Pataz gold province with the location of the main deposits. (Haeberlin et al., 2002/b).

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 566-568

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Geological and structural setting

In relation with the Early Carboniferous calc-alkaline magmatism in the Eastern Cordillera of northern Peru

granodioritic, dioritic and monzogranitic bodies have been emplaced along a NW-SE oriented structural line

(Fig.s 1 & 2). Shear stress related tectonism dominated the Eastern Cordillera while gold-bearing quartz-sulfide

veins were emplaced within the batholith. In the north, the mineralized veins are located near the western margin

within the batholith body but towards the south they appear in central or eastern marginal position. The massive

quartz veins of Pataz deposit contain gold bearing pyrite, galena and sphalerite. Two main sets of mineralized

veins have economic importance in Pataz district: a NW-striking E-dipping (~45° to E) and an E-W striking flat

(~10° to E) extensional vein system. Mineralization in both types of vein has similar mineralogical composition

and structural control but gold concentrations are slightly higher in the flat veins.

Huachón district (10°40.6’S, 75°53.5’W), which could presumably belong to the southern extension of the

Maranón Valley gold belt, has ~48-60° dipping quartz veins to the southwest, within the west marginal part of

the granodiorite body of Paucartambo batholith with similar Fe and base-metal sulfide mineralization. Wall-rock

alteration around (~1-2m) the veins contains mainly sericite and chlorite, and is similar to Pataz district’s

alteration characteristics.

Ore and gangue mineralogy

In all districts pyrite is the ore related mineral and has two main types of appearance. Fine grained pyrite

which is mostly the gold bearing phase and coarse grained barren, euhedral pyrite which occurs in massive

stocks or disseminated related to the first and second stages of the mineralization (Haeberlin, 2004). The gold

concentrations are divers, generally 10-20 g/t and irregularly can attain higher values, up to ~3-4 oz/t. The most

common gangue mineral, quartz appears at least in three different generations, an early milky quartz, a second

stage blue-grey microgranular quartz related to the gold precipitation phase and a late stage white quartz

(Haeberlin, 2004). Beside the quartz, carbonate minerals appear in the late stage of the paragenesis. Alteration

minerals, principally chlorite and muscovite, do not show important variations related to the change in the

composition of the host rock.

Fluid geochemistry

Previous microthermometric measurements on the mineralization of Pataz speak about three fluid inclusion

populations which can be well related to the three stages of mineralization. The results reflect a mixing event

between an early hot brine fluid with a late phase low-temperature saline water (Haeberlin, 2002/a).

From Huachón district fluid inclusion study is currently undertaken. Our aim is to compare the different types

of fluids involved in the ore formation at Pataz, Parcoy and Huachón area. Therefore a comparative O isotopic

study will be done as well on quartz from divers elevations of the different ore mineralizations to constrain

sources of the ore forming fluids.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 566-568

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Summary

Current study consists of structural, petrological, dating and fluid-geochemical investigations of host rock and

mineralized bodies at Huachón district for tracing the possible evidences of a ~400 km long, coherent gold belt

in the central Eastern Cordillera of Peru.

References Haeberlin, Y. (2002/a) Geological and structural setting, age and geochemistry of the orogenic gold deposits at Pataz

Province, Eastern Andean Cordillera, Peru. Doctoral Thesis, University of Geneva, Terre et Environnement 36: 182 p.

Haeberlin, Y.,, Moritz, R., Fontboté, L., (2002/b) Paleozoic orogenic gold deposits in the eastern Central Andes and its

foreland, South America, Ore Geology Reviews 22, p. 41-59

Haeberlin, Y., Moritz, R., Fontboté, L., Cosca M., (2004) Carboniferous orogenic gold deposits at Pataz, Eastern Andean

Cordillera, Peru: Geological and structural framework, paragenesis, alteration, and 40Ar/39Ar geochronology, Economic

Geology, Vol. 99 p. 73-112.

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Chemical and mineralogical characterization of the River Huasco (Norte Chico, Chile)

Ana Valdés1, Mireille Polvé

1, & Diego Morata

2

1 Laboratoire des Mécanismes et Transferts en Géologie, UMR 5563 / UR 154 CNRS–Université Paul Sabatier–

IRD–Observatoire Midi-Pyrénées, 14 avenue Edouard Belin, 31400 Toulouse, France (valdes@lmtg.obs-mip.fr) 2

Departamento de Geología, Facultad de Ciencias Físicas y Matemáticas, Universidad de Chile, Plaza Ercilla,

Santiago, Chile

KEYWORDS : waters, sediments, chemical composition, hydrothermalism, heavy metals

Geographical context

The area of study is located south of the third region of Atacama between 28°S and 29°S, corresponding to the

Huasco River Valley. The origin of this river is located at the confluence of the rivers “del Transito” and “del

Carmen”, which flows into the northern part of the city with the same name. The river Huasco extends for 88

kilometers and presents a basin of 9850 km2, with a general orientation east west.

Geographically the total surface of the basin is equivalent to a 13% of the regional surface area (III region of

Atacama). Through the Huasco River valley different urban centers can be found, for example, Puerto Huasco,

Freirina, Vallenar y Alto del Carmen. According to the 2002 census, there exist 15 sites, from which two are

cities and the rest rural areas. The cities found on the basin are Vallenar with 48040 inhabitants, and Huasco with

7945.

Objectives

Considering geologic, geographic and meteorological characteristics of the study area, the goal of the present

study is to characterize chemically and mineralogically the river Huasco through the analysis of water and

sediment samples, collected between the cities of Alto del Carmen and Huasco. This will allow us to know the

origin of the elements analysed, the determination of the transport mode of a metal by a surface drainage, and to

know some of the factors that control the dynamics of the river Huasco.

Methodology

Two field work sessions were organised within the framework of this project. The first field work was done on

April 2007 and the second was performed on Januray 2008. Each corresponding to 10 days of work.

The data obtained in the first field work period allows us to have a general knowledge of the study area. From

these results the general distribution of the elements allowed a more detailed sampling exercise during the

second field work campaign.

To accomplish the detailed sampling, two sections of the Huasco valley were defined. The first between Alto

del Carmen and Vallenar, and the second between Vallenar and Huasco. A total of 11 locations were sampled,

with one river water sample and one sediment sample taken at each location). Both sample groups were analysed

by ICP-MS (Inductively Coupled Plasma – Mass Spectrometry) at the LMTG, Toulouse, France.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 569-570

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Results

For each station, water and sediment are compared, and sediment mineralogy taken into account. Evolution of

water and sediment chemical compositions from upstream to downstream traces the local influence of each

formation, as the two sections sampled cross cut all the N-S geological formations. Water analysis from the two

field sessions show inter seasonal variations as they have been sampled during two successive summers. These

observations can be enlarged using the DGA (Dirección General de Aguas) data base.

The water analysis results obtained in both field sessions are compared with previous data from waters in this

area, in particular data from hydrothermal waters in order to quantify the dilution factors compared to these

hydrothermal waters. River sediments are compared with the mean value of the Andean continental crust.

All data are interpreted in terms of regional signature, local hydrothermal influence and mining activities.

Acknowledgements This work was possible by the support of the ALßAN EU grant to A.V. and the IRD-Chile. The first field work session was realized with the logistic support of professors Rodrigo Riquelme and Arturo Jensen of the Universidad Católica del Norte, Chile. In addition, the authors thank Maria Eliana Lorca from the Universidad de Chile for her assistance during field work. Finally A.V. thanks Raul Martinez for his help with English corrections.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 571-572

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Climatic impact on the erosive dynamics of the Pacific Central Andes revealed by cosmogenic and hydrological records of river sediments

Riccardo Vassallo1, Emilie Pépin

1, Vincent Regard

1, Jean-Loup Guyot

1, Sébastien Carretier

1,

Eric Gayer2, Laurence Audin

1, Frédéric Christophoul

1, Rodrigo Riquelme

3, Juan Julio

Ordóñez4, Fernando Escóbar-Cáceres

5

1 Laboratoire des Mécanismes et Transferts en Géologie, Toulouse, France

2 Institut de Physique du Globe, Paris, France

3 Universidad de Antofagasta, Chle

4 SENHAMI, Lima, Peru

5 DGA, Santiago, Chile

KEYWORDS : Pacific Andes, climate, erosion, cosmonucleids, suspended load

The Pacific side of the Central Andes, characterized by a similar tectonic pattern and by a strong North-South

climatic gradient, offers the opportunity to estimate the impact of climate variability on catchments denudation

rates. To understand the mechanisms and the rates of the geomorphic evolution along this mountain range, it is

important to quantify landscape processes over different time scales. For this reason, in our approach we coupled

historical records of hydrological events (suspended load data of the Chilean DGA and ongoing measurements in

Peruvian rivers), used to constrain erosion and transport rates over the last decades, with cosmogenic terrestrial

nuclides analysis in sediments, to extend the study of these processes over millennial timescales. This study is

supported by the project ANR "ANDES" jeunes chercheurs ANR-06-JCJC-0100.

We combined records of river suspended load fluxes over the last 40 years with measurements of cosmogenic

nuclide (10Be, 26Al, 21Ne) concentrations in alluviums and colluviums in 20 main catchments of Peru and Chili

(Figure 1). Short-term erosion rates are very low from South Peru up to the Central Chili, then increase with a

strong gradient around the region of Santiago, well correlating with the water discharge pattern over the range.

However correlations disappear in the zone of Chile characterized by large glaciers, which probably disturb the

erosion signal (Figure 2).

Lima

Santiago

Figure 1. Location of the samples of the cosmogenic nuclides analysis within the main Pacific catchments of South Peru and North Chili, and pictures showing examples of rivers sampling.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 571-572

572

Figure 2. Annual mean water discharges and catchments erosion rates from data of the DGA (Chile).

To determine long-term erosion rates, we sampled river sands for cosmogenic analysis in similar

morphostructural areas for each catchment. Moreover, for some of them we extended the sampling to several

points of the river profile and to hillslopes, and collected different sizes of cobbles (Figure 3).

This approach will allow us to detect local variations of the erosion rates within the single catchments and to

better constrain the dynamics and transport of sediments in rivers. Preliminary long-term results of this study

will be presented for the first time at this conference.

Figure 3. Picture of the Ocona river (Peru) showing a valley morphology characterized by a wide active riverbed transporting alluviums (from sand to cobble size), and sapping hillslopes covered by regolith.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 573-576

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Thermotectonic history of the Northern Andes

D. Villagómez1, R. Spikings

1, D. Seward

2, T. Magna

3, W. Winkler

2, & A. Kammer

4

1 Département de Minéralogie, Université de Genève, 13 rue des Maraîchers, 1205 Genève, Switzerland

(Diego.Villagomez@terre.unige.ch, Richard.Spikings@terre.unige.ch) 2 Geologisches Institut, ETH-Zurich, 8092 Zürich, Switzerland (diane.seward@erdw.ethz.ch,

wilfried.winkler@erdw.ethz.ch) 3 Institut de Minéralogie et Géochimie, Université de Lausanne, L’Anthropole, 1015 Lausanne, Switzerland

(Tomas.Magna@unil.ch) 4 Departamento de Geociencias, Universidad Nacional de Colombia, A.A. 14490 Bogotá, Colombia

(akammer@unal.edu.co)

KEYWORDS : Northern Andes, Tahami terrane, accretion, Caribbean plateau

Introduction

The Northern Andean Zone stretches from Ecuador, through Colombia to Venezuela and is characterized by

three sublinear topographic ridges, referred to as the Western, Central and Eastern Cordillera. The basement of

the Western Cordillera of Colombia and Ecuador is formed by multiple oceanic terranes mainly accreted during

the Mesozoic; the terranes are juxtaposed against the paleo-continental margin (Central Cordillera) across the

Romeral Fault Zone. The aim of this contribution is to provide a temporal framework for the pre-, syn- and post-

accretionary tectonic framework of Western Colombia. Our study is based on multiphase thermochronological

methods (40Ar/39Ar, apatite and zircon fission-track), geochronological (zircon U/Pb LA-ICPMS) and

geochemical (ICPMS and XRF) analyses of crystalline, sedimentary and mafic volcanic rocks from several

traverses across the Colombian Andes.

Paleocontinental margin

In Colombia the continental province crops out in the Central Cordillera. West of the Otú-Pericos Fault (Figure

1), this province (the so-called Tahami terrane, (Touissant and Restrepo, 1994)) is partly composed of medium-

to low-pressure Paleozoic metamorphic rocks (detrital U/Pb zircon ages ranging from 270-380 Ma), which are

intruded by syn- and post-collisional granitic stocks and granitic gneisses with U/Pb zircon ages between

200-300 Ma (Vinasco et al., 2006; this work). These granitic sequences are related to the agglutination of the

Pangea supercontinent during Permian and late break-up during Triassic times.

All these pre-Triassic rocks are intruded by Jurassic calc-alkaline granites of the Ibague Batholith (160 ± 3 Ma,

U/Pb) and Late Cretaceous intrusive rocks of the Antioquia Batholith (U/Pb ages between 83.75 ± 0 and

94.5 ± 1.7 Ma, (Ibañez-Mejía et al, 2007; this work)).

The Paleozoic metamorphic belt is bounded to the west by a narrow volcanic sedimentary body of the

Quebradagrande Complex (Figure 1), which geochemically has been interpreted as having formed in a zone of

back-arc spreading during the Albian-Aptian (Nivia et al., 2006). To the west, these calc-alkaline rocks are in

tectonic contact with isolated tectonic slices of amphibolites, eclogites and metamorphic basic intrusive rocks

exposed on the western flank of the Central Cordillera (Arquía Complex). Recently obtained U/Pb data of an

igneous body (Amagá granite) which intrudes these HP rocks yielded 227.6±4.5 Ma (Vinasco et al, 2006),

implying that the Arquía Complex is pre-Triassic and hence supports the ensialic marginal basin origin for the

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Quebradagrande Complex. Geochemical and geochronological characterization of these HP rocks of the Arquía

and Quebradagrande complexes is in progress.

Figure 1. Geochronology and Thermochronology of the Central and Western Cordillera of Colombia.

Allochthonous terranes

The accreted Cretaceous terranes occur to the west of the regional Romeral Fault zone (more specifically the

Cauca-Almaguer Fault) in the western flank of the Central Cordillera (Figure 1). They seem to be of oceanic

plateau affinity. They consist of basalts, gabbros and ultramafic cumulates (Amaime Fm., Bolívar Complex,

Volcanic and Barroso Fms.), which are characterized by flat mantle-normalized REE patterns (Kerr et al., 1997,

this project) and may represent a portion of the large late-Cretaceous Colombian-Caribbean oceanic plateau

(CCOP, Kerr et al., 1997).

We have proceeded to date two mafic hornblende-bearing pegmatites from the Bolivar Ultramafic Complex of

Western Cordillera, which has been described as an internal part of the CCOP (Kerr et al., 2004), yielding ages

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 573-576

575

of 94.0±2.4 and 95.5±1 Ma (U/Pb). New (U/Pb) ages on granitic rocks from the Buga Batholith are 90.6±1.2 and

91.5±1.3 Ma and may perhaps represent the initial stages of east-facing island arc activity that formed at the

juvenile active margin of the eastward migrating CCOP (and may be genetically related to the coeval Aruba

Batholith in the Southern Caribbean). We have also dated the arc-related Cordoba Batholith, which intrudes and

hence post-dates the mafic terranes. This rock yields a U/Pb age of 79.3±1.5 Ma, which constrains the minimum

age of the accreted terranes. Therefore, the age of the accreted terranes is likely to be between 95 and 80 Ma.

Time of accretion of the allochthonous terranes

In Colombia, several preliminary fission track (FT) ages have been obtained from the rocks of the Tahami

Terrane in the Central Cordillera (Figure 1). In the central section the apatite fission track (AFT) ages range

between ~77 and ~36 Ma. An older group of nearly indistinguishable AFT ages (ages within 1 error interval) of

77±6, 69±4 (sample 05DV82) and 60±8, 59±10Ma are representative of the region and were obtained from

samples located relatively far (>8 km) from local fault traces (e.g. Ibagué Fault). Samples close to sheared rocks

within the Ibagué Fault yielded younger AFT ages of 31±3 and 36±3 Ma (sample 05DV06). Four zircon fission

track (ZFT) ages on those same samples are indistinguishable within error (78±5, 81±5, 85±9, 88±6 Ma; Fig. 1).

We performed inverse modeling on some of these samples (05DV06 and 05DV82), using the annealing model

of Carlson et al. (1999), and the Monte-Carlo inverse modeling procedure of Ketcham et al. (1999) (Figure 2), to

constrain their potential thermal histories using the ZFT age as time constraint at ~250±50°C. Sample 05DV82

(mean track length of 14.54±0.76μm), located 20 km from the Ibagué Fault (Figure 1), cooled rapidly through

250˚C to ~ 60˚C at ~ 80-70 Ma. Sample 05DV06, located closer to the Ibagué Fault (Figure 1) hosts partially

annealed FT lengths in apatite (mean track length of 13.60±1.62 μm), suggesting they may have resided for a

significant amount of time within the apatite partial annealing zone (APAZ). The best-fit model indicates a rapid

cooling through ~ 250˚C to ~ 90˚C between 80-70 Ma, followed by slower cooling from 70-10 Ma and finally

renewed rapid cooling from ~ 60˚C to ~25˚C between 10 Ma to present.

Figure 2. Modeled T-t paths and track length distribution for two samples which crop out in the Central Cordillera. Monte-Carlo inverse modeling following procedure of Ketcham et al. (1999) and multi-kinetic approach (based on Dpar) according Carlson et al. (1999). APAZ = apatite partial annealing zone

Our study indicates that Jurassic granitoids (Ibagué Batholith) emplaced along the eastern border of the Central

Cordillera in Colombia cooled rapidly through ~ 250˚C to ~ 60˚C between 80-70 Ma. We conclude that an

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 573-576

576

important tectonic event produced exhumation in the paleocontinental margin during late Cretaceous and ascribe

this event to the accretion of the CCOP.

Preliminary conclusions

Initial stages of interaction between the CCOP and the Ecuadorian margin took place in Late Campanian –

Mastrichtian (75 - 65 Ma, Vallejo et al., 2006). However, our study shows that this event occurred slightly

earlier in Colombia. Hence it is plausible to suggest, pending additional data, that CCOP accretion took place in

southward direction due to the northward drifting of South America during the Late Cretaceous. This Campanian

accretion is coeval with the cessation of the late Cretaceous arc in Northern Colombia (Antioquia Batholith) due

to the clogging of the subduction zone caused by the collision between South America and the CCOP. As

already mentioned, this was synchronous with accelerated surface uplift and exhumation within the buttressing

continental rocks and is temporally corroborated by the onset of clastic sedimentation derived from the Central

Cordillera into the Upper-Middle Magdalena Valley (UMV-MMV, Figure 1) located immediate to the east.

Sediment progradation to the east (El Cobre sandstone) and the initiation of eastward shifting of the axis of

deposition of the MMV began during Campanian time (Villamil, 1999).

This work is supported by the Swiss National Science Foundation (DV & RS)

References Carlson, W.D., Donelick, R.A., Ketcham, R.A., 1999. Variability of apatite fission-track annealing kinetics: I Experimental

results. American Mineralogist. 84, 121 –1223. Ibañez-Mejia M., Tassinari C.C.G.,; Jaramillo J.M. 2007. U-Pb zircon ages of the “Antioquian Batholith”: geochronological constraints of late Cretaceous magmatism in the central Andes of Colombia. In Proceedings of XI

Congreso Colombiano de Geologia. Bucaramanga. Colombia. Kerr, A.C., Marriner, G.F., Tarney, J., Nivia, A., Saunders, A.D., Thirlwall, M.F., Sinton, C.W. 1997. Cretaceous Basaltic

Terranes in Western Colombia: Elemental, Chronological and Sr-Nd Isotopic Constraints on Petrogenesis. Journal of. Petrology. 38, 677–702.

Kerr, A.C., Tarney, J., Kempton, P.D., Pringle, M., Nivia, A. 2004. Mafic pegmatites intruding oceanic plateau gabbros and ultramafic cumulates from Bolivar, Colombia; evidence for a "wet" mantle plume? Journal of Petrology. 45, 1877-1906.

Ketcham, R.A., Donelick, R.A., Carlson, W.D. 1999. Variability of apatite fission-track annealing kinetics: III. Extrapolation to geological time scale. American Mineralogist, 84, 1235-1255.

Nivia A., Marriner G., Kerr A., Tarney J. 2006. The Quebradagrande Complex: A lower cretaceous ensialic marginal basin in the Central Cordillera of the Colombian Andes. Journal of South American Earth Sciences. 21. 423-436.

Toussaint, J.F., Restrepo, J.J., 1994. “The Colombian Andes during Cretaceous times”. In (J.A. Salfity), Cretaceous tectonics of the Andes, Verlag Braunschweig, Wiesbaden, Germany, 61–100.

Vallejo, C., Spikings, R., Luzieux, L., Winkler, W., Chew, D., Page, L., 2006. The early interaction between the Caribbean Plateau and the NW South American Plate. Terra Nova 18, 264–269.

Villamil T. 1999. Campanian-Miocene tectonostratigraphy, depocenter evolution and basin development of Colombia and western Venezuela. Palaeogeography, Palaeoclimatology, Palaeoecology. 153. 239-275.

Vinasco C.J., Cordani U.G., Gonzalez H., Weber M., Pelaez C. 2006. Geochronological, isotopic, and geochemical data from Permo-Triassic granitic gneisses and granitoids of the Colombian Central Andes. Journal of South American Earth Sciences 21. 355-371.

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Cenozoic high-strontium andesites in the Eastern Cordillera of Northwestern Argentina, Central Andes

José Maria Viramonte1, Néstor Suzaño

1, Carolina Prescott

2,3, Raul Becchio

1, José Germán

Viramonte1, Marcelo Arnosio

3, & Marcio M. Pimentel

2

1

University of Salta and CONICET, Geonorte Institute, Av. Bolivia 5051, 4400 Salta, Argentina

(jviramonte@hotmail.com) 2 University of Brasilia, Geoscience Institute, Geochronology Laboratory, 70910-900, Brasilia, Brazil

3 University of Salta, Geonorte Institute, Av. Bolivia 5051, 4400 Salta, Argentina

KEYWORDS : Cenozoic, andesites, high-strontium, Eastern Cordillera, Central Andes

Cenozoic magmatism in the southern Central Andes occurs generally in the N–S trending volcanic arc and in

NW–SE trending transverse volcanic belts (Viramonte et al., 1984).

In the Eastern Cordillera of northwestern

Argentina, near the Huachichocana town (24º-65º,

Jujuy province; Figure. 1) crop out Cenozoic

andesites with high Sr contents. These rocks are

associated with the Lipez transverse belt (Figure 1)

and were firstly described by Ramos et al., (1967) as

“Huachichocana Andesite”. They are located ~590

Km from the trench in a back arc position and

represent one of the easternmost magmatism

between 23º and 25º LS with the Diego de Almagro

Complex (Hauser, 2005) and Alemania-Pampa

Grande Andesites (Figure. 1). The Huachichocana

rocks occur as a 200m-thick and 700m-large sill

along the “Tilcarica” unconformity between the

Puncoviscana Formation (Upper Precambrian-Lower

Cambrian) and the Meson Group (Mid-Upper

Cambrian). This body present columnar jointing and

have a porphyritic texture comprising by phenocrysts

of hornblende, biotite, clinopyroxene, orthopyroxene

and zoned plagioclase included in a cryptocrystalline

matrix. Titanite occurs as an accessory mineral.

Major and trace elements composition are

presented in Table 1. They have 61-63 % wt. SiO2

and according to the SiO2 vs K2O (Figure 2) diagram

Huachichocana rocks plot in the high-K andesite field. Figure 3 show that these rocks present the highest Sr

contents (950-1200 ppm) comparing with other andesites of the southern Central Andes. On chondrite

normalized diagram (Figure 4) these rocks display coherent REE patterns characterized by enrichment in LREE

Figure 1. Huachichocana location regarding the magmatic arc.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 577-579

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relative to HREE as indicated by La/SmN= 3.99 to 4.65 and Gd/YbN= 2.04 to 2.31. Also, the samples show very

minor Eu anomalies (Eu/Eu*=0.85-0.91).

Table 1. Major and trace elements of Huachichocana rocks

Sample Huachichocana HO-4 HO-5 HO-6 HO-7 SiO2 61.87 61.17 62.81 61.22 61.56

TiO2 0.57 0.60 0.496 0.56 0.56

Al2O3 16.40 17.16 17.18 17.45 17.18

Fe2O3 5.53 5.48 4.92 5.25 5.27

MnO 0.12 0.140 0.129 0.13 0.14

MgO 2.44 2.04 2.01 2.02 2.07

CaO 5.61 6.00 5.194 6.02 5.76

Na2O 3.18 3.20 3.194 3.58 3.51

K2O 3.05 2.69 3.08 2.48 2.54

P2O5 0.38 0.37 0.337 0.36 0.36

PPC 0.6 1.78 0.96 0.54 0.89

total 99.785 100.65 100.31 99.64 99.88

Ba 762 764 840 874 806

Rb 86 81 91 75 79

Sr 954 1061 1088 1202 1164

Zr 201 225 213 230 224

Y 25 26 25 25 25

Nb 17 20 19 20 21

Ni 10 11 9 9 8

Cr 6 4 3 3 2

La 35 46.5 41.1 45.5 46.6

Ce 65.7 85.4 76.4 83.9 85.2

Pr 7.79 9.89 8.9 9.67 9.93

Nd 30.3 37.5 33.7 37.1 37.7

Sm 5.65 6.61 6.09 6.71 6.46

Eu 1.64 1.82 1.62 1.89 1.82

Gd 5.31 6.02 5.53 6.03 5.96

Tb 0.75 0.8 0.75 0.8 0.81

Dy 3.95 4.2 3.88 4.19 4.15

Ho 0.76 0.82 0.75 0.8 0.8

Er 2.24 2.4 2.23 2.36 2.37

Tm 0.31 0.34 0.32 0.33 0.34

Yb 2.15 2.15 2.12 2.22 2.21

Lu 0.33 0.34 0.33 0.33 0.34

Sm-Nd and Sr isotopic data for selected samples are presented in table 2. These rocks show low 87Sr/86Sr ratios

(0.70591-0.70665), negative Nd values ranging from -2.2 to -5.4 and TDM model ages in the interval between

0.81 to 0.99 Ga.

Figure 3. SiO2 vs Sr diagram (modified from Ramos et al., 2004).

Figure 2. SiO2 vs K2O diagram (Peccerilo and Taylor, 1976).

Figure 4. Chondrite normalized REE patterns of Huachichocana rocks (normalizing values of Sun and Mc Donough, 1989)

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 577-579

579

Table 2. Nd and Sr isotopic composition of Huachichocana rocks

Samples Sm

(ppm)

Nd

(ppm)

147Sm/

144Nd

143Nd/

144Nd

± 2SE

(o)

TDM

(Ga)

87Sr/

86Sr

± 2SE

Huachich 5,867 31,748 0,1117 0,512459+/-20 -3,49 0,87 0,70665+/-1

h-01 6,733 31,488 0,1292 0,512525+/-18 -2,21 0,94 0,70591+/-3

h-04 6,596 36,881 0,1081 0,512358+/-19 -5,45 0,99 0,70636+/-2

h-05 6,512 37,788 0,1042 0,512388+/-18 -4,88 0,91 0,70655+/-2

h-06 6,299 34,911 0,1091 0,512456+/-15 -3,55 0,86 0,70631+/-5

h-07 6,648 37,444 0,1073 0,512481+/-11 -3,07 0,81 0,70635+/-2

Following the ideas of Kay et al., (1999) the geochemical and isotopic data, we suggest a possible origin of

the Huachichocana rocks through partial melting of lower mafic crust (Sunsas?) and subsequent contamination

with upper crustal rocks during its ascent to shallow crust levels. The high Sr values as well as the Eu/Eu* values

allow a little, if any, plagioclase in the restite.

These rocks present some differences when comparing with typical andesites of the southern Central Andes

indicating a different origin and evolution. So, further studies are carried out in order to better constrain the

source and the evolution of the Huachichocana rocks and similar rocks outcropping southward in the Eastern

Cordillera (Diego de Almagro Complex and Alemania-Pampa Grande Andesites).

Acknowledgements We would like to thank to R. Pereyra and A. Nieva (Universidad Nacional de Salta) for help with the sample preparation and chemical analyses. Financial support for field and laboratory works was provided by a CIUNSA project Nº 1350/3 and SECyT, PICT 2005 Nº 07 - 38131.

References Hauser, N. 2005. Estudio petrográfico y geoquímico de las volcanitas aflorantes al sur de la localidad estacion Diego de

Almagro, departamento Rosario de Lerma, provincia de Salta. Tesis profesional. Universidad Nacional de Salta. Inédito. Kay, S.M.; Mpodozis, C., Coira, A.B., 1999. Neogene magmatism, tectonism, and mineral deposits of the central Andes. In:

Skinner, B.J. (Ed.), Geology and Ore Deposits of the Central Andes. Society of Economic Geology, Special Publication, vol. 7, pp. 27– 59.

Peccerillo, A. and Taylor, S.R. 1976. Geochemistry of Eocene Calkalcaline volcanic rocks from the Kastamanou area, northern Turkey. Contributions to Mineralogy and Petrology 58: 61-63.

Ramos, V.; Turic, M.A.; Zuzek, A. B. 1967: Geología de las quebradas de Huichaira-Pecoya, Purmamarca y Tumbaya Grande en la margen derecha de la quebrada de Humahuaca. Revista de la Asociación Geológica Argentina, 22 (3): 209-221.

Ramos, V.; Kay, S.M.; Singer, B.S. 2004. Las adakitas de la Cordillera Patagónica: Nuevas evidencias geoquímicas y geocrnológicas. Revista de la Asociación Geológica Argentina. 59 (4): 693-706.

Sun, S.S. and Mc Donough, W.F. 1989. Chemical and isotopic systematic of ocean basalts: implication for mantle composition and processes. In: Saunders, A. D. Norry, M..J. (Eds.), Magmatism in ocean basins. Geol.. Soc. London Spec. Pub. 42: 313-345.

Viramonte, J.G.; Galliski, M. A.; Araña Saavedra,V.; Aparicio, A.; García Escacho L. y Martín Escorza, C. M. 1984: El finivolcanismo básico de la depresión de Arizaro, provincia de Salta. IX Congreso Geológico Argentino, Bariloche. Actas III: 234-251.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 580-582

580

Heterogeneous thermal overprint of a Late Palaeozoic fore-arc system in north-central Chile (32°–31°S) discernible by small scale equilibration and age domains (Ar-Ar; fission track)

Arne P. Willner1,2

, Hans-Joachim Massonne2, Masafumi Sudo

3, & Stuart Thomson

4

1 Institut für Geologie, Mineralogie & Geophysik, Ruhr-Universität, D-44780 Bochum, Germany

(arne.willner@rub.de) 2 Institut für Mineralogie und Kristallchemie, Stuttgart University, Azenbergstr. 18, D-70174 Stuttgart, Germany

3 Institut für Geowissenschaften, Potsdam University, Karl-Liebknechtstr. 24, D-14674 Potsdam, Germany

4 Department of Geology & Geophysics, Yale University, P.O. Box 208109, New Haven, CT 06520-8109, USA

KEYWORDS : Ar-Ar dating, fission track dating, accretionary system, thermal overprint, age resetting

Geological setting

The metamorphic basement in north-central Chile at lat. 31°-32°S shows various levels of a coastal

accretionary system which are telescoped to a short distance in outcrop by Mesozoic tectonic processes at the

southernmost end of the Atacama strike-slip system. The Choapa Metamorphic Complex (CMC) comprises low

grade rocks (metagreywacke, greenschist) and has the same structural inventory as the Western Series south of

34°S that originated by basal accretion (Willner et al. 2005). The Arrayán Formation (AF), dominated by very

low to low grade metagreywacke, shows similar structures as the Eastern Series south of 34°S which was formed

by frontal accretion (Richter et al. 2007). The Huentelauquén Formation (HF), unconformably overlying the

Arrayán Formation, is a heterogeneous sedimentary sequence of shelf deposits involving platform limestone,

conglomerate and neritic clastic sediments with an Upper Carboniferous to Permian biostratigraphic age

(Rebelledo and Charrier 1994). U/Pb-dating yielded a detrital magmatic zircon population in the metamorphic

basement which demonstrates a maximum age of deposition at 303 Ma for the HF. This deposition was

concomitant with that of the CMC (maximum deposition age 308 Ma) and , thus, occurred in a retrowedge basin.

On the other hand, deposition of the AF was considerably older (maximum deposition age 343 Ma; Willner et al.

2008) confirming the general relationship, as observed south of lat. 34°S, that basal accretion follows frontal

accretion in time in Chile (Richter et al. 2007). The oldest intrusion ages (Rb-Sr isochron) obtained by Irvine et

al. (1988) at 31°S are 220±20 Ma for a gabbro and 200±10 Ma for a monzogranite documenting the end of the

accretion process at the same time as in north-central Chile (Willner et al. 2005) and south-central Chile (Glodny

et al. 2005). This magmatic event is concomitant to an extensional event with ubiquitous basin opening in central

Chile. It was followed by bimodal dyke intrusions of a Jurassic fore-arc according to K/Ar ages of hornblende in

mafic dykes 133-213 Ma (Irvine et al. 1988). This corresponds to a similar K/Ar age spectrum of hornblende in

metabasite within the CMC at 154-220 Ma.

Pressure-temperature constraints

In the CMC metabasite with the assemblage Ca-amphibole - white mica - chlorite - quartz - plagioclase ±

epidote contains two generations of amphibole, white mica and plagioclase at thin section scale: actinolite, albite

and phengite (Si 3.25-3.33 pfu) define a relic metamorphic stage I at 6.0-7.5 kbar, 305-360°C, whereas

magnesiohornblende or tschermakitic hornblende, oligoclase and muscovite (Si 3.15-3.20 pfu) equilibrated at

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 580-582

581

stage II at 480-540°C, 4.5-6.0 kbar. Similar pressure conditions are corroborated by metapsammopelitic rocks,

where two white mica generations occur as well. A local rock in the CMC is a garnet mica-schist with the

assemblage garnet-phengite-chlorite-epidote-quartz-rutile. Garnet is replaced by chlorite and phengite which are

not in equilibrium with the garnet rim. Multivariant reactions based on chlorite-phengite pairs show maximum

PT data of 400-430°C and 12-14 kbar. These highest PT-conditions in the basement are presumably associated

with an anticlockwise PT-path and the highest ages. Hence this rock shows the same characteristics as

equivalent rare rocks described at 33°S and 41°S (Willner et al. 2005; Willner et al. 2004) indicative of

conditions at the onset of the accretion process.

On the other hand, in a rhyolite pebble within a

conglomerate and in K-feldspar bearing metagreywackes of

the HF pressures did not exceed 3 kbar (muscovite Si

3.2 pfu). Detrital phengite grains with Si-contents up to

3.3 pfu were eroded from the contemporaneous basal

accretionary prism (CMC). This confirms the deposition of

the HF in a retrowedge basin. Local presence of biotite in HF

rocks indicates metamorphic temperatures as high as 350°C.

In most parts of the AF metagreywacke micas are only

muscovite (Si 3.2 pfu) occurring as metamorphic and

detrital mineral. Minimum pressures were in the same range

as in the HF.

Ar-Ar and fission track dating

In order to to obtain ages at thin section scale white mica in

eleven samples from the three units CMC, AF and HF was

systematically studied by an in situ Ar-Ar UV laser ablation

method (50 μm spot size). In addition, detrital zircon from

seven samples was investigated by fission track dating.

Only two rocks were recognized showing mainly relic ages

which correspond to the peak of HP/LT metamorphism and

hence to the the accretion process: (1) the above mentioned

garnet mica-schist yielded an age range for single white mica

between 276±3 and 310±3 Ma. The oldest age gives a

constraint to the maximum age of accretion whereas younger

ages correspond to white mica formation by mineral reactions

during cooling. This effect was shown to be independent of

internal rock deformation by Willner et al. (2005) due to

static growth after the accretion process. Minor ages at 247-

200 Ma are interpreted to be due to incipient Mesozoic

resetting. (2) The age cluster of a surrounding metagreywacke

Fig. 1. The metamorphic basement in north-central Chile.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 580-582

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shows a range of single white mica ages between 244±12 and 281±3 Ma. Again following our results of studies

south of lat. 34°S (Willner et al. 2005) we interpret the oldest age as the age of peak HP/LT metamorphism and

the younger age as that of white mica formation during cooling under 300°C. In both rocks about 35-40 Myrs of

retrograde mineral formation are documented. A zircon FT-age of 248±27 Ma is equivalent to the youngest

white mica age and represents cooling under ~280°C.

All further studied rock samples from the CMC, AF and HF show ages younger than 248 Ma with age clusters

at 190-219 Ma, 154-173 Ma and 131-137 Ma. These clusters match the time of first Mesozoic intrusions into the

accretionary wedge and Jurassic bimodal vein intrusions into the fore-arc setting and, thus, do not correspond to

HP/LT metamorphism and accretion. Most FT-ages are older than the Ar-Ar white mica ages. For instance, in

two samples from the CMC zircon FT-ages are at 274±18 and 272±40 approaching the maximum age of

metamorphism, whereas Ar-Ar white mica ages typically show wide age ranges at 199-248 Ma and 129-143 Ma.

We interpret the Upper Triassic to Jurassic Ar-Ar ages as resetting ages. Resetting affects detrital as well as

metamorphic white mica in a similar way. No primary detrital white mica ages are preserved. In some cases

several broad age clusters occur, whereas in some cases age peaks are narrow and can be interpreted as complete

resetting. Resetting is assumed to be due to influx of fluids ascending in hydrothermal systems within

extensional environments that are known from Upper Triassic to Jurassic times.

Furthermore, FT dating of three local zircon samples show astonishing young ages of 96±5 Ma and 104±5 Ma

and also a mixture of single values ranging from 99 to 281 Ma. The event at 100 Ma corresponds to a regional

short-time compressional episode with slight crustal thickening. Locally, a deeper part of the accretionary system

was probably exhumed along a fault and cooled below 280°C in Cretaceous times.

Summarising the in insitu Ar-Ar age study reveals a rather complex overprint history in a fore-arc setting of a

long-lived convergent margin that cannot be resolved into great detail by integrating methods.

References Glodny, J., Lohrmann, J., Echtler, H., Gräfe, K., Seifert, W., Collao, S. & Figueroa, O., 2005 - Internal dynamics of a

paleoaccretionary wedge: insights from combined isotope tectonochronology and sandbox modelling of the south-central Chilean forearc. Earth and Planetary Science Letters 231: 23-39.

Irvine, J.J., García, C., Hervé, F. & Brook, M., 1988 - Geology of part of a long-lived dynamic plate margin: the coastal cordillera of north-central Chile, latitude 30°51´-31°S. - Canadian.Journal of Earth Sciences 25: 603-624.

Rebelledo, S.& Charrier, R., 1997 - Evolución del basamento paleozoico en el área de Punta Claditas, Región de Coquimbo, Chile (31-32°S). Revista Geológica de Chile 21: 55-69.

Richter, P., Ring, U., Willner, A.P. & Leiss, B., 2007 - Structural contacts in subduction complexes and their tectonic significance: The Late Paleozoic coastal accretionary wedge of central Chile. - Journal of the Geological Society of London 164, 203-214.

Willner, A.P., Glodny, J., Gerya, T.V., Godoy, E. & Massonne, H.-J., 2004 - A counterclockwise PTt-path in high pressure-low temperature rocks from the Coastal Cordillera accretionary complex of South Central Chile: constraints for the earliest stage of subduction mass flow. – Lithos 75: 283-310.

Willner, A.P., 2005 - Pressure-temperature evolution of an Upper Paleozoic paired metamorphic belt in Central Chile (34°-35°30´S). - Journal of Petrology 46: 1805-1833.

Willner, A.P., Thomson, S.N., Kröner, A., Wartho, J.A., Wijbrans, J & Hervé, F., 2005 - Time markers for the evolution and

exhumation history of a late Palaeozoic paired metamorphic belt in central Chile (34°-35°30´S). - Journal of Petrology 46: 1835-1858.

Willner, A.P., Gerdes, A. & Massonne, H.-J., 2008a - History of crustal growth and recycling at the Pacific convergent margin of South America at latitudes 29°-36°S revealed by a U-Pb and Lu-Hf isotope study of detrital zircon from late Paleozoic accretionary systems. Chemical Geology, in press.

Willner, A.P., Richter, P., Ring, U., 2008b - Structural modification of a late Paleozoic accretionary system in north-central Chile (34°-35°S) during postaccretional shortening episodes at a long-lived active margin. Revista Geologica de Chile, in press.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 583-586

583

Upper Pleistocene deglaciation as a conditioning factor for catastrophic mass redistribution in Las Cuevas basin, Mendoza, Argentina

C. G. J. Wilson1, R. Hermanns

2, L. Fauqué

1, M. Rosas

1, V. Baumann

1, & K. Hewitt

3

1 Servicio Geológico Minero Argentino. Av. Julio A. Roca 651, piso 10 (1322) Ciudad de Buenos Aires, Argentina

(carloswilson71@gmail.com) 2 Norges geologiske undersøkelse, Leiv Eirikssons vei 39, N-7491 Trondheim, Norway

(Reginald.Hermanns@NGU.NO) 3 Cold Regions Research Centre, Wilfrid Laurier University, Waterloo, Ontario N2L 3C5, Canada (khewitt@wlu.ca)

We analysed large rock slope failures in the Las Cuevas basin, Mendoza valley (Figure 1 ) using air photos,

satellite images, digital elevation models as well as intensive field work including absolute dating of deposits by

means of 14C dating of organic matter and surface exposure dating using the cosmogenic nuclide 36Cl. A total

of 8 large rock slope failures were identified. The large landslides recognized in Las Cuevas basin are (Figure 2):

Rock Avalanche of Penitentes (Fauqué, 2008) (Figure 3); Megalandslide in the Southern wall of the Aconcagua

(Fauque et al 2008b) (Figure 4), Rock Avalanche of Tolosa valley (Figure 5) and Rock Avalanche of Las

Cuevas (Figure 6) (Rosas et al 2008). The corresponding ages are shown in Table 1.

Figure 1: Location map.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 583-586

584

Figure 2:View of the Rock Avalanche of Las Cuevas from the west

Figure 3: Orthogonal view of Rock Avalanche of Penitentes in a North direction

All results of our dating indicate

postglacial events spanning from about 1500

years after the Last Glacial Maximum to the

early Holocene. Similar to other mountain

regions three of the deposits (Horcones

deposits and Penitentes deposit) have been

earlier misinterpreted as glacial deposits

resulting in an erroneous interpretation of

glacial stratigrapphy (Espizua L.E., 1989) of

the western Central Andes. However, detailed

sedimentologic analyses based on

mineralogic and grain roundness

investigation show that these deposits are the

result of complex interaction of rockslides

resulting from the large mountain failures

with valley glaciars and glacial valley fill

resulting in a mixture of various deposits.

Avalanche of Las Cuevas

Main scarp

i ii

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 583-586

585

These mountain slope failures started as high mobility rock avalanches however due to erosion of glacial ice

and glacial deposits these flow type landslides got more water saturated and continued for 34 km down from the

south wall of Aconcagua. The resulting deposits obstructed Horcones and Las Cuevas valleys, and thus caused

due to catastrophic dam failures several cascading catastrophic events.

Table 1

Process Stratigraphy Material Age Error Dating

Megalandslide in the Southernwall of the

Aconcagua i

- Lacustrine deposits

(pelites) 8.260 - 8.254 Carbon 14

Megalandslide in the Southernwall of the

Aconcagua i -

Lacustrine deposits

(pelites) 14.798 - 13.886 Carbon 14

Megalandslide in the Southernwall of the

Aconcagua ii Aconcagua group Volcanic rocks 11.110 - 8.170

Cosmogenic

nuclide

Rock Avalanche of Las Cuevas event i

Tordillo Formation and

Puente del Inca´s

Traquites

Conglomerates and red

sandstones. Traquites 13.125 875

Cosmogenic

nuclide

Rock Avalanche of Las

Cuevas event ii

Tordillo Formation and

Puente del Inca´s

Traquites

Conglomerates and red

sandstones. Traquites 10.250 750

Cosmogenic

nuclide

Rock Avalanche of Tolosa

Stream Juncal Formation

Volcanic and pyroclastic

rocks 9.307 -

Cosmogenic

nuclide

Rock Avalanche of Penitentes

Choiyoi Group Granite 13.890 920 Cosmogenic

nuclide

Rock Avalanche of Penitentes

Choiyoi Group Dacitic breccia, riolitic

and riodacitic igninbrites 11.820 790

Cosmogenic

nuclide

Rock Avalanche of Penitentes

Choiyoi Group Dacitic breccia, riolitic

and riodacitic igninbrites 10.620 730

Cosmogenic

nuclide

Furthermore rock slope failures along the south wall of Cerro Aconcagua caused glacial capture of glaciers

flowing east previously to rockslope. Today these glaciars are hanging glaciars in the South wall of Cerro

Aconcagua which collapse into the upper Horcones valley. Today's frequent glacial surges in lower Horcones

valley influencing security of mountaineer camps along the main rout towards Cerro Aconcagua are seen as a

long term effect of this post glacial slope collapses.

The absolute ages obtained vary between the Tardiglacial (15-1014 C ka BP) and the Postglacial Periods which

corresponds to the Early Holocene (10-8 14C Ka BP). Most of the dates belong to the Tardiglacial Period that is

characterized by a rapid change towards warmer conditions.

We interpret that in the Las Cuevas basin, the important loss of glacial ice generated destabilization of the

Mountain Geomorphic System due to the glacial “debuttresing” and changes of the elevation of permafrost

causing these multiple rockslides. Fauqué et al (2005) studied an example of this process at Puente del Inca

characterized by deep slope gravity deformation showing that slope processes due to the change of climatic

conditions is ongoing to the present.

The Las Cuevas River Basin analysis shows that the Upper Pleistocene climatic change caused an important

perturbation in high mountain processes. This evidence allows us to speculate the implications that present

global warming might have in this high mountain environment. However, it is likely that the possible changes

will be much smaller than those associated with the Upper Pleistocene Deglaciation (15000-10000 years

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 583-586

586

BP).This study was part of a project “Geoscientific study applied to land use planning in Puente del Inca”

(SEGEMAR, 2007) which was developed within the framework of the Multinational Andean Project,

Geosciences for Andean Communities (MAP-GAC) by the Argentine Geological Service (SEGEMAR) and the

Environmental Planning and Urban Development Direction (DOADU). It was developed with technical and

financial support of the Government of Canada.

The purpose of the study was to provide guidelines to the administrative and territorial organizations with

jurisdiction in Puente del Inca. These guidelines will allow to take decisions for rural planning and further

development of the community.

References Espizua, L.E.1989 Glaciaciones Pleistocenicas en la quebrada de los Horcones y rió de las Cuevas. Mendoza. Republica

Argentina Evans , S.G. and Clague, J.J., 1993. Glacier-related hazards and climate change. In : R. Bras (Editor), The world at Risk :

Natural Hazards and Climate Change. Am. Inst. Phys. Conf. Proc., 277 : 48-60. Fauqué, L., M. Rosas, R. Hermanns, V. Baumann, S. Lagorio, C. Wilson y K. Hewitt, 2008. Origen y edad del depósito

asignado al drift Penitentes. Mendoza, Argentina. XVII Congreso Geológico Argentino. En prensa. Jujuy. Fauqué, L., Hermanns, R., Rosas, M., Wilson, C., Lagorio, S., Baumann, V., Di Tommaso, I., Hewitt, K., Coppolecchia, M.

y González, M., 2007. Geomorfología. En: Estudio geocientífico aplicado al ordenamiento territorial de Puente del Inca. PMA-GCA. IGRM-SEGEMAR. Informe Final, 10-34. Buenos Aires.

Fauqué, L., M. Rosas, M. Coppolecchia, R. Hermanns, M. Etcheverría, A. Tejedo y C. Wilson, 2005. Laderas afectadas por deformaciones gravitacionales profundas en el valle del río Cuevas. Provincia de Mendoza. XVI Congreso Geológico Argentino, Actas 3: 515-520. La Plata.

Rosas, M., C. Wilson., R . Hermanns y L. Fauqué, 2008. Las avalanchas de rocas de Las Cuevas. Mendoza, Argentina. XVII Congreso Geológico Argentino. En prensa. Jujuy.

SEGEMAR, 2007. Estudio geocientífico aplicado al ordenamiento territorial de Puente del Inca. PMA-GCA. IGRM-SEGEMAR. Informe Final, 73 pp., Buenos Aires.

Figure 5: View of the Rock Avalanche of Tolosa Stream from the North.

Avalanche of Tolosa

Figure 4: View of the Megalandslide in the Southern wall of the Aconcagua facing the town of Puente del Inca a south east direction

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 587-591

587

Timing and causes of the growth of the Ecuadorian cordilleras, as inferred from their detrital record

Wilfried Winkler1, Cristian Vallejo

1,3, Léonard Luzieux

1,4, Richard Spikings

2, & Nergui Martin-

Gombojav1

1 Geological Institute, Department of Earth Sciences, ETH Zentrum HAD, 8092 Zurich, Switzerland

(wilfried.winkler@erdw.ethz.ch) 2 Department of Mineralogy, University of Geneva, CH-1205 Geneva, Switzerland (spikings@terre.unige.ch)

3 Present Address: Salazar Resources, 10 de Agosto N37-232 y Villalengua, Quito, Ecuador

(cvallejo@curimining.com) 4 Present Address: Holcim Group Support Ltd. CH-5113 Holderbank, Switzerland (leonard.luzieux@holcim.com)

KEYWORDS : Northern Andes, Carribean plateau, provenance analysis, detrital zircon U/Pb ages, paleotectonics

Introduction

The Andean cordilleras of Ecuador are considered to have formed during multiple, continent-ocean accretion

events since the Early Cretaceous. Thus, these distinct collision events should be documented in the sedimentary

record that evolved in response to the growth of the cordilleras. We review the growth of the Ecuadorian

cordilleras using compositional, geochronological, thermochronological data from the i) Late Cretaceous-Present

retro-arc foreland basin (Oriente and Subandean zone), ii) Late Cretaceous-Paleogene sedimentary basins that

precede, are coeval with and post-date the collision of the Caribbean plateau and arcs with the paleo-margin of

Ecuador (Cordillera Occidental and Costa), and other late post-collisional sedimentary rocks (Neogene) that

crop-out in the flat forearc (Costa).

Provenance has been estimated using standard heavy mineral analyses, which we combine with (1) single

detrital zircon grain U/Pb LA-ICPMS ages (to determine source rock ages and possible multiple recycling), and

(2) detrital zircon fission-track (FT) measurements and calculated lag-times (to determine the exhumation

history of the source regions). Finally, the inferred tectonic history of the Andean chain will be calibrated against

thermochronological results from the cordilleras.

Andean Amazon Basin (retro-arc foreland basin)

The Andean Amazon Basin, located east of the Cordillera Real in Ecuador, has been a depocenter since the

Aptian-Albian. The early heavy mineral assemblage, until the late Campanian-Maastrichtian, is characterized by

a simple stable mineral association (zircon-tourmaline-rutile; ZTR), which implies it was derived from the

erosion of shallow granitic continental crust and/or recycling of older sedimentary rocks. Detrital zircons in the

Hollin and Napo fms. yield a broad U/Pb age distribution, ranging from ~ 2.0-0.5 Ga, with a few ages older than

2.5 Ga, and some minor Paleozoic and Mesozoic ages (Martin-Gombojav and Winkler, 2008). A central

question is whether or not the basin was supplied exclusively from the Amazon craton during the early stage of

evolution, as it is believed according to sedimentological interpretations (e.g. Barragán et al. 2004), or if a

primordial cordillera to the west already existed. A primordial Cordillera Real may have existed because, i)

Proterozoic zircons are also frequently observed in the post-Campanian foreland basin series (Tena, Tiyuyacu,

Chalcana fms.), which were predominantly derived from the Cordillera Real (e.g. Christophoul et al. 2002), ii)

The entire basin fill succession yields a group of ZFT ages that range between 270 and 225 Ma. Most likely,

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 587-591

588

these zircons were derived from the Triassic Tres Lagunas granite (Litherland et al., 1994), which is a crustal

anatectite that formed via melting of the older Paleozoic basement and intruded Paleozoic schists of the

Cordillera Real (Litherland et al., 1994). The presence of blue (rutilated) quartz grains in the Hollin Fm., which

frequently occur within the Tres Lagunas Granite, corroborates this interpretation, and iii) Scattered ZFT lag-

times ranging between 60 and 0 Ma, measured in the Hollin and Napo fms., suggest rocks located west of the

basin were being eroded, possibly inherited from exhumation during the Peltetec event (e.g. Litherland et al.,

1994), and Jurassic Misahualli arc may also have been a source region. Independent evidence comes from a

palynologic analysis of the Napo Fm. (Vallejo et al. 2002), which reveals that the basin had no connection with

the paleo-Pacific.

The Turonian-Recent evolution of the Cordillera Real is recorded in the following manner. Since ~ 80 Ma, the

low zircon FT lag-times, combined with a frequent change of source regions, confirm that the cordillera was

exhuming rapidly (± 1 mm/year). Medium metamorphic grade minerals have been reworked since the

Maastrichtian-Paleocene (Tena Fm.), and high grade (kyanite, sillimanite) minerals have been reworked since

the Eocene. This trend documents the exhumation of progressively deep crustal levels in the Cordillera Real.

The appearance of recycled mafic, volcanic minerals (diopsidic augite, hypersthene, olivine and chromian

spinel) from the Late Oligocene on (~ 25 Ma) indicates that the Cordillera Occidental was exhuming. The

importance of this exhumation event is emphasized by subsequent constant lagtimes (± 35 Ma), and the

appearance of a second population of zircons with low lagtimes. This suggests that an important Oligocene event

has brought a large volume of source rocks in the Cordillera Real close to the partial annealing zone and steady

state exhumation prevailed since then.

Cordillera Occidental and Costa (forearc)

Basaltic lavas and hyaloclastites of the Pallatanga and Piñon units form the mafic basement of the Cordillera

Occidental and the coastal blocks, respectively. Radiometric age data (40Ar/39Ar, U/Pb SHRIMP) and

biostratigraphic correlations of overlying sedimentary rocks show the volcanic rocks erupted between ~ 90 and

87 Ma (Luzieux et al. 2006, Vallejo et al. 2006). Age data and geochemical signatures (e.g. Mamberti et al.

2003) indicate a derivation of these Ecuadorian rocks from the Caribbean oceanic plateau, which were shred off

during the collision and NE drift of the plateau with the Ecuadorian and Colombian margin of South America

(Spikings et al., 2001). In the coastal region, (1) paleomagnetic inclinations prove that the mafic basement

extruded at equatorial, low southern latidudes (Fig. 1), and (2) rapid, biostratographically constrained changes in

paleomagnetic declination reveal 20-50o vertical axis clockwise rotations occurring between 73 and 70 Ma.

Several volcanic arcs have been identified (Rio Cala, San Lorenzo and Las Orquideas), which intruded the

plateau sequence. Chronostratigraphic constraints from the Rio Cala Group in the Cordillera Occidental suggest

that initiation of east-facing subduction under the plateau occurred soon after extrusion of the plateau. Arc-

related turbidites solely contain mafic to intermediate volcanic-type heavy minerals, corroborating geochemical

evidence for an intra-oceanic origin (Vallejo et al. 2006), pre-dating the collision of the plateau with the South

America margin. The turbiditic Yunguilla Fm. was deposited along the South American continental margin prior

to, and during the collision of the plateau. Heavy minerals and detrital zircon U/Pb LA-ICPMS data reveal a

dominant component of Proterozoic ages (Vallejo 2007), similar to coeval sedimentary rocks in the Andean

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 587-591

589

Amazon Basin (Tena Fm.), implying that the Cordillera Real represented the prominent feature before the

plateau/arc collision.

Figure 1. Proposed model for the formation and collision of the Carribbean plateau and parts of the Greater Antilles Arc with the northern South America continent during Late Coniacian-Late Campanian (from Luzieux 2007). The large-scale plate tectonic situation is according to Duncan and Hargraves (1984).

The Paleocene Saguangal and Saquisilí fms., deposited on the newly created forearc, are post-accretion

formations. Their heavy mineral compositions show a mixture of continental crust and mafic volcanic grains.

Detrital zircon U/Pb ages from the Sanguagal Fm. correlate with source regions within the Cordillera Real. The

Paleocene-Eocene Angamarca Group as a whole (including the basal Saquisilí Fm.), was derived from medium-

grade metamorphic rocks in the Cordillera Real.

Cenozoic siliciclastic sediments covering the Piñon and Santa Elena blocks, as well as the Progreso basin fill,

depict a mixed detrital supply from accreted mafic volcanic basement and arcs, and from continental crust,

including medium- to high-grade metamorphic rocks in the Cordillera Real. However, the coeval distal forearc

(San Lorenzo, Pedernales and Esmeraldas blocks) and the sedimentary rocks of the Neogene Borbon and

Manabi basins were nearly exclusively derived from mafic rocks (Luzieux 2007). This suggests that already in

the Paleocene-Eocene an axial, southward directed drainage system, parallel to the evolving cordilleras,

developed in the forearc, as it prevails today.

Thermochronological calibrations

Numerous multi-phase 40Ar/39Ar, zircon and apatite FT and apatite (U-Th)/He data, which constrain the thermal,

exhumation and growth history of the cordilleras of Ecuador have been published (e.g. Spikings et al. 2001,

2005, Spikings and Crawhurst 2004). The Cordillera Real and Subandean zone were exhuming at high rates at

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 587-591

590

73-55 and 43-30 Ma. The source rocks of the Paleocene Saquisilí Fm. presumably situated in the Cordillera

Real, cooled rapidly during 74-65 Ma. The Amotape complex experienced significant cooling during 75-65 and

43-39 Ma (both events possibly associated with clockwise block rotations). Reactivation of fault blocks in the

Cordillera Occidental is inferred for the period 42-32 Ma and later during ca. 13 and 9 Ma. Furthermore, high

exhumation rates (> 1km/my) have been recorded in the northern and central Cordillera Real at 15, 9 and 5-3

Ma. The Cretaceous exhumation events correlate with the collision of the Caribbean plateau with the Ecuadorian

margin during ca. 75-65 Ma as also concluded by Luzieux et al. (2006) and Vallejo et al. (2006), using different

analytical techniques. Eocene exhumation was previously considered to be a response to accretion of the

Macuchi arc in the Eocene. However, Vallejo (2007) has shown that the Macuhi arc is autochthonous, and

enhanced exhumation during the Eocene may be a consequence of a significant change in plate convergence

directions.

The paleotectonic model

In conclusion we present a refined model of the Ecuadrian Andes (Fig. 1), which differs in several points from

earlier ones (e.g. Kerr et al. 2002). Soon after extrusion of the Caribbean plateau westward subduction under its

leading edge gave rise to intra-oceanic arc development (Rio Cala, San Lorenzo, Las Orquideas). The plateau

and overlying arcs drifted eastward and collided with the South America margin during the Campanian. This is

inferred from the termination of arc magmatism in the Early Maastrichtian, and the clockwise rotation of coastal

blocks during ca. 73-70 Ma. New, eastward subduction under the accreted oceanic plateau fragments was

established in the Late Maastrichtian. On the new active margin from the latest Maastrichtian to Eocene the

Silante, and subsequently, the Macuchi arc developed. Volcanism occurred coeval with the Saguangal Fm. and

Angamarca Group forearc basin deposition, which were mainly shed from the emerging Cordillera Real. There

is no positive evidence for an Eocene accretion of the Macuchi block, which would be geometrically

challenging, because the coastal blocks situated to the west (Piñon, San Lorenzo etc.) already collided with the

margin during the Late Cretaceous. Enhanced Eocene-Oligocene uplift in the cordilleras as documented by the

erosion of increasing deeper metamorphic levels in the Cordillera Real may also have involved the Cordillera

Occidental. Since the Late Oligocene, the scree of the Cordillera Occidental also contributes to the detrital flux

into the Andean Amazon Basin.

Acknowledgments We acknowledge the support by various Swiss National Science Foundation grants, in particular grant no. 2-72058-05.

References Barragán, R., Christophoul, F., White, H., Baby, P., Rivadeneira, M., Ramirez, F. and Rodas, J., 2004. Estratigraphia

secuencial del Cretacio de la Cuenca Oriente del Ecuador. In Baby, P., Rivadeneira, M., and Barragán, R. (eds.): La Cuenca Oriente: Geologia y Petroléo. Traveaux de l’Institut Français d’Etudes Andines, 144: 45-68.

Christophoul, F., Baby, P. and Dávila, C., 2002. Stratigraphic response to a major tectonic event in a foreland basin: the Ecuadorian Oriente basin from Eocene to Oligocene times. Tectonophysics, 345: 281-298.

Duncan, R.A., Hargraves, R.B. 1984. Plate tectonic evolution of the Caribbean region in the mantle reference frame. In Bonini, W.E., Hargraves, R.B., Shagam, R. (eds.): The Caribbean – South America Plate Boundary and Regional Tectonics. Geological Society of America Memoire, 162: 81–93.

Kerr, A.C., Aspden, J.A., Tarney, J. Pilatasig, L.F. 2002. The nature and provenance of accreted oceanic Blocks in western Ecuador: geochemical and tectonic constraints. Journal of the Geological Society, 159: 577-594.

Mamberti, M., Lapierre, H., Bosch, D., Ethien, R., Jaillard, E., Hernandez, J. and Polve, M., 2003. Accreted fragments of the

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Late Cretaceous Caribbean-Colombian plateau in Ecuador. Lithos, 66: 173–199. Litherland, M., Aspden, J., Jemielita, R.A. 1994. The metamorphic belts of Ecuador. British Geological Survey. Overseas

Memoir, 11: 147 p. Luzieux, L.D.A., Heller, F., Spikings, F., Vallejo, C.F., Winkler, W. 2006. Origin and Cretaceous tectonic history of the

coastal Ecuadorian forearc between 1°N and 3°S: Paleomagnetic, radiometric and fossil evidence. Earth and Planetary Science Letters, 249: 400-414.

Luzieux, L.D.A. 2007. Origin and Late Cretaceous-Tertiary evolution of the Ecuadorian forearc. PhD Thesis, Institute of Geology ETH Zürich, Switzerland, 197 p.

Martin-Gombojav, N. and Winkler W. 2008. Recycling of Proterozoic crust in the Andean Amazon foreland of Ecuador: implications for orogenic development of the Northern Andes. Terra Nova, 20: 22-31.

Ruiz, G. M. H., Seward, D., Winkler W. 2004. Detrital thermochronology – a new perspective on hinterland tectonics, an example from the Andean Amazon Basin, Ecuador. Basin Research, 16: 413-430.

Spikings, R.A. Winkler, W., Seward, D., Handler, R. 2001. Along-strike variations in the thermal and tectonic response of the continental Ecuadorian Andes to the collision with heterogeneous oceanic crust. Earth and Planetary Science Letters, 186: 57-73.

Spikings, R.A. and Crawhurst, P.V. 2004. (U-Th)/He thermochronometric constraints on the late Miocene-Pliocene tectonic development of the northern Cordillera Real and the Interandean Depression, Ecuador. Journal of South American Earth Sciences, 17: 1-13.

Spikings, R.A., Winkler, W., Hughes, R.A., Handler, R., 2005. Thermochronology of allochthonous blocks in Ecuador: unraveling the accretionary and post-accretionary history of the Northern Andes. Tectonophysics, 399: 195–220.

Vallejo, C., Hochuli, P.A., Winkler, W., von Salis, K. 2002. Palynological and sequence stratigraphic analysis of the Napo Group in the Pungarayacu 30 well, Sub-Andean Zone, Ecuador. Cretaceous Research, 23: 845-859.

Vallejo, C., Spikings, R.A., Winkler, W., Luzieux, L., Chew, D., Page, L. 2006. The early interaction between the Caribbean Plateau and the NW South American Plate. Terra Nova, 18: 264-269.

Vallejo, C. 2007. Evolution of the Western Cordillera in the Andes of Ecuador (Late Cretaceous-Paleogene). PhD Thesis, Institute of Geology ETH Zurich, Switzerland, 208 p.

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 592-593

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Damage zone and the occurrence of world-class porphyry copper deposits in the active margin of Chile: Geophysical signatures and tectonomagmatic inferences

Gonzalo Yáñez1,2

, Orlando Rivera3, Diana Comte

2, Mario Pardo

2, Luis Baeza

3, & Emilio

Vera2

1 CODELCO, Teatinos 258, piso 8, Santiago, Chile

2 Departamento de Geofísica, FCFM, Univ. de Chile, Blanco Encalada 2002, Santiago, Chile

3 Exploraciones Mineras Andinas S.A., Apoquindo 4775, oficina 602, Santiago, Chile

KEYWORDS : tectonomagmatic, porphyry copper, seismicity, gravity, magnetotelluric

World-class porphyry copper deposits represent an extremely anomalous tectono-magmatic process in actives

margins, in both, space and time. The delicate combination of favorable structures, stress regime, magmatic and

fluid source migration, traps, and preservation, might end up in an ore deposit if the appropriate equilibrium is

reached. In Chile, two of the most productive porphyry belts of the world, the Paleocene-Oligocene province of

Cordillera de Domeyko (20-26°S), and the Miocene-Pliocene metalotect of the high Andes (32-35°S), are the

best places to get some insights on the associated tectonomagmatic processes. The petrophysics of these

particulars tectono-magmatic domains of the Andean region, is inferred from a series of geophysical experiments

carry out in this project: natural seismicity, a regional gravimetric survey, and magnetotelluric transects. Those

geophysical experiments were initially carried out by Codelco and EMSA in the Cordillera de Domeyko, and

later on further extended to the principal cordillera of Central Chile in the framework of the ANILLO ACT18

and FONDECYT 1050758 projects.

From a regional scale perspective, these deposits are generally located in the flanks of a high gravimetric

anomaly: in the western flank of the Atacama Block in the northern domain (Cordillera de Domeyko), and the

northern and southern flank of the Mapocho-El Volcán Block (Central Chile). Gravity modeling constrained by

seismic tomography indicates that these dense and rigid/impervious blocks are located at mid-crustal depths

(15-20km), with densities at the order of 3.0 gr/cc. Natural seismicity is distributed in the periphery of these

blocks, further suggesting their rigid and impervious nature. In addition, deep crustal natural seismicity is also

aligned in NW, NE, and NS directions, following old and penetrative structures of lithospheric nature

(translithospheric domains), which are controlling the tectonomagmatic evolution of the margin and the

geometry of the dense and impervious blocks. In fact, world-class porphyry copper deposits are precisely located

in this permeable (and seismically active) periphery, generally intersecting the translithospheric fabric. Working

hypothesis in progress suggest a passive and active role of this dense and impervious blocks: in one side (passive

role) directing the magmas and fluid flow towards the flanks, and on the other hand (active role) providing a

likely complementary magma source of Cu for the mineralized bodies. Such a working hypothesis also suggests

a tectono-magmatic perspective beyond the Andean Cycle, as far as the accretion of Precambrian blocks (with

the attachment of young and buoyant slabs) to the proto South American lithosphere.

At a local scale, world-class porphyry copper deposits shows a gravity low (5-10 mgal) in direct response to a

pervasive damage zone (wide structural network, hydrothermal breccias and alteration products). Seismic

activity shows a concentration at depths in the range of 0-18 km, elongated in the direction of the particular

7th International Symposium on Andean Geodynamics (ISAG 2008, Nice), Extended Abstracts: 592-593

593

structural fabric of each deposit. It is interesting to point out that the depth extent of the natural seismicity is

deeper in the central zone (up to 15-18 km) compared with the northern domain (mostly concentrated below

8 km depth). We interpret this behavior as the result of a major volume of pore-fluid pressure in the southern

segment. Vp/Vs ratio within the deposits of the southern segment is characterized by maximum values (1.8-

1.85), whereas in the northern deposits the Vp/Vs ratio is transitional from high gradients towards low Vp/Vs

ratios (1.6-1.7). In the southern zone, this high Vp/Vs ratio is interpreted as a direct consequence of partially

saturated damage zones. In the northern zone, the rather opposite behaviour is attributed to a low Vp response as

a result of the pervasive damage domain, affecting the medium porosity (and density), under low humidity

conditions, further supported by the differences in the seismic depth extent already discussed. The likely

association of this natural seismic clustering and mine-blasts and/or induced mine-related has been thoroughly

analyzed. In mines where daily blast activity is applied (Chuquicamata, Río Blanco Los Bronces), seismicity is

concentrated in a window time of 2 hours after the blasts. Whereas in others (i.e. El Teniente mine), where blast

activity is minor, seismic activity is distributed randomly during the day. The induced effect in the first case is

evident given the direct cause-effect relationship, however, the space (wide in depth and distance) and time

(narrow concentration) distribution is highly peculiar. Our working hypothesis postulates that this behaviour is

the result of a highly deformed zone with a stick-slip recovery time of less than two hours. This time constant is

in agreement with direct observations of surface deformation in the banks of the mine, with a recovery time of

less than two hours.

Geoelectric imaging (MT studies), show that damage zones associated to world class porphyry Cu deposits are

also linked to low resistivities sub vertical domains (5-20 ohm-m) within scales of 5-10 km. These low resistivity

values are also characteristic of structural systems percolated by hydrous phases which is consistent with the

independent inferences provided by gravity and seismic studies.

The geophysical characterization of world-class-copper-deposits provides a robust framework for the necessary

tectonomagmatic conditions required for their spatial distribution and genesis.

This is a contribution to Project PBCT ACT 18.

594

author index

A ABASCAL, L. 13 ACEÑOLAZA, G. F. 17 ACOSTA, H. 549 ADAMS, C. J. 17 AGUDELO, W. 156 AGUILAR, G. 21 AGUILERA, F. 25,530 AGUILERA, N. G. 401 AGUIRRE, L. 29 AGUIRRE-URRETA, B. 33 AGUSTO, M. 101 ALBERT, F. 397 ALBINO, F. 281 ALEJANDRO, V. 265 ALLMENDINGER, R. 238 ALONSO, J. L. 401 ALVÁN, A. 549 ALVARADO, M. J. 37 ALVARADO, P. 41,75 ALVAREZ, V. 84 ANDERSON, M. 75 ANGLADE, A. 223 ARAUJO, S. 339

ARELLANO, S. 67 ARGOLLO, J. 44 ARMIJO, R. 48,77 ARMIJOS, E. 387 ARNOSIO, M. 577 ARON, F. 52,116,238 ARRIAGADA, C. 56,269,354,405 ASCH, G. 481 ASTINI, R. A. 60 ASTROZA, M. 41 AUDEMARD, F. A. 37 AUDIN, L. 64,253,295,474,521, 562,571 AVOUAC, J.-P. 64

B BABAULT, J. 508 BABY, P. 199,322,435,454 BACHMANN, O. 191 BAEZA, L. 592 BALDO, E.G. 427 BANCHIG, A. L. 344 BARBA, D. 67 BARBARAND, J. 199 BARCKHAUSEN, U. 94 BARRIENTOS, S. 41 BASSO, M. 116 BATAILLE, K. 348,413 BAUMANN, V. 583 BAZAN, H. 387 BEAN, C. J. 334

BEATE, B. 124 BECCHIO, R. 577 BECHIS, F. 71,330 BECK, S. 41,75 BÉJAR-PIZARRO, M. 77 BELMAR, M. 29 BELTRAN, A. 273 BENAVENTE, C. 295 BERMÚDEZ-CELLA, M. 81 BERNAL, C. 140 BERNAL, I. 534 BERNET, M. 81 BÉTHOUX, N. 84,302 BISCAYART, P. P. 88

BONDOUX, F. 64 BONVALOT, S. 77,136 BOOKER, J. R. 90 BORDARAMPÉ, C. P. 373 BOURDON, B. 191 BOUVET de MAISONNEUVE, C. 191 BRÄNDLEIN, D. 188 BRASSE, H. 188 BROWN, L. D. 277 BRUSSET, S. 199,322 BUATOIS, L. 195 BULNES, M. 401 BURD, A. 90 BURGOS, J. D. 140,454 BUSKE, S. 245

C CACYA, L. 545 CALAHORRANO, A. 94 CALDERÓN, S. 29 CALDERON, Y. 435 CAMPOS, J. 48 CAÑOLA, E. 97 CAPECCHIACCI, F. 101 Caracas Seismic M. P. group 500 CARBONEL, P. 442 CÁRDENAS, J. 105 CARDONA, A. 120 CARLIER, G. 105 CARLOTTO, V. 105,469,473 CARRASQUERO, S. I. 109 CARRETIER, S. 206,295,387,571 CARRIZO, D. 77,113,238 CARTER, A. 517 CASALLAS, I. F. 216 CASELLI, A. 101 CASQUET, C. 427 CEMBRANO, J. 52,116,238 CERECEDA, C. 369 CERREDO, M. 393 CHABALIER, J.-B. de 77 CHARADE, O. 77 CHARRIER, R. 160,206,269,357,397 CHARVIS, P. 84,223 CHAUVIN, A. 469 CHAZOT, G. 257 CHEW, D. 120 CHIARADIA, M. 124 CHICANGANA, G. 128, 132 CHLIEH, M. 136 CHRISTOPHOUL, F. 140,454,571 CISTERNA, C. E. 144 CLAVERO, J. E. 553 COBBOLD, P. R. 148 COIRA, B. L. 277 COLLO, G. 60,152 COLLOT, J.-Y. 84,156,292,306,431 COMTE, D. 160,206,277,341,592 CONTRERAS-REYES, E. 164,493 CORREA, M. J. 88 CORTÉS, J. M. 542 CORTÉS, J. 168,238 COTTEN, J. 257 CRISTALLINI, E. O. 71,219

D DAHLQUIST, J.A. 427

595

DAMM, T. 242 DARRAH, T. 25 DARROZES, J. 21,140,172,435,521 DÁVILA, F. M. 152,176 DE LA CRUZ, R. 524 DELACOUR, A. 545 DELAVAUD, E. 500 DELOUIS, B. 136 DELPIT, S. 180 DÉRAMOND, J. 454,521 DESHAYES, P. 184 DESSA, J.-X. 223 DÍAZ, D. 188 DÍAZ, J. 84 DOMÍNGUEZ, J. 500 DUNAI, T. 113,238 DUNGAN, M. A. 191,517

E ECHAVARRÍA, L. E. 88 ECHTLER, H. P. 326,348 ENCINAS, A. 195,206 ESCÓBAR-CÁCERES, F. 571 ESCRIG, S. 191 ESMERALDAS team 302 ESPITIA, W. 450 ESPURT, N. 199,322,435

F FABRE, D. 315 FACCENNA, C. 322 FANNING, M. 357,524 FARBER, D. L. 64,203,253,474 FARÍAS, M. 160,206,269,365 FAUQUÉ, L. 583 FAVETTO, A. 90,373 FERRER, O. 261 FINGER, K. L. 195 FINKEL, R. C. 253 FLUEH, E. R. 164,493 FOEKEN, J. 517 FOLGUERA, Alicia 231 FOLGUERA, Andrés 210,289,384,461,513 FONT, Y. 84,214 FORNARI, M. 442 FRAIZY, P. 387 FUNICIELLO, F. 322

G GABALDA, G. 136,168 GAILLER, A. 84,223 GALÁN, R. A. 216 GALINDO, C. 427 GALLEGO, A. 341

GALVE, A. 223 GARCÍA, V. H. 71,219 GARCÍA-CANO, L. C. 223 GARCÍA-MORABITO, E. 227 GAYER, E. 571 GERBE, M.-C. 257,446,545 GIAMBIAGI, L. 71,231,330 GIBERT, G. 235 GILBERT, H. 75 GIORDANENGO, G. 90 GOLDSTEIN, S. 191 GONZÁLEZ, A. 160 GONZÁLEZ, G. 52,113,116,168,238 GONZÁLEZ, M. 500 GONZÁLEZ-BONORINO, G. 13 GÖTZE, H.-J. 242,409,489 GOURGAUD, A. 446

GREVEMEYER, I. 94,164,493 GROSS, K. 245 GUILLAUME, B. 249 GUNNELL, Y. 281 GUTIÉRREZ, A. A. 261 GUYOT, J.-L. 387,571

H HACKNEY, R. 526 HALL, M. L. 67,351 HALL, S. R. 253,474 HANCOCK, G. S. 203 HASSANI, R. 235 HEIT, B. S. 277 HELLO, Y. 84,223 HÉRAIL, G. 172,365,458,474,521 HEREDIA, N. 401 HERMANNS, R. 384,583 HERMOZA, W. 199,435 HERNÁNDEZ, J. J. 500 HERVÉ, F. 485 HEWITT, K. 583 HIDALGO, S. 180,257 HUSSON, L. 249,381

I IAFFA, D. N. 261 IGLESIAS, M. 381 INGLES, J. 172 IZARRA, C. 489

J JACAY, J. 265,504 JACOVKIS, P. M. 417 JARA, P. 269 JWEDA, J. 191

K KAMMER, A. 132,273,450,573 KAUSEL, E. 48 KAY, S. M. 277 KIND, R. 277 KIRKLAND, C. L. 120 KLEY, J. 477 KLOTZ, J. 348 KO LER, J. 120 KUMMEROW, J. 481

L LA RUPELLE, A. de 281 LACASSIN, R. 48 LAFFAILLE, J. 37 LAGNOUS, R. 435 LANGE, D. 326 LANGMUIR, C. 191 LAPRIDA, C. 33 LARA, L. E. 116,285 LARSEN, J. 90 LAZO, D. G. 33 LE PENNEC, J.-L. 67,180,446 LEAL, V. 500 LEGRAND, D. 526 LIPPAI, H. 393 LITHGOW-BERTELLONI, C. 176 LITVAK, V. D. 289 LODOLO, E. 393 LÓPEZ, E. 292 LÓPEZ, S. M. 97 LOVELESS, J. 238 LUZIEUX, L. 587

596

M MACEDO, O. 334 MACHARÉ, J. 295 MAGNA, T. 120,573 MAIRE, E. 21,521 MAKSAEV, V. 357 MAMANI, M. 298,545 MANCHUEL, K. 84,302 MARCAILLOU, B. 156,306 MARENTES, M. 450 MARÍN-CERÓN, M. I. 310,538 MARIÑO, J. 446,545 MARQUES, F. O. 148 MARTELLI, K. 315 MARTIN, H. 257,446 MARTINA, F. 60 MARTÍNEZ, A. 231 MARTÍNEZ-DOPICO, C. I. 319 MARTIN-GOMBOJAV, N. 587 MARTINOD, J. 168,206,238,249,322, 381,474 MASSONNE, H.-J. 580 MCGLASHAN, N. A. 277 MEDINA, E. 25,116,530 MELNICK, D. 326,348 MENA, R. 144 MENICHETTI, M. 393 MERINO, D. 124 MESCUA, J. F. 71,231,330 MÉTAXIAN, J.-P. 334, 339 MICHAUD, F. 442 MIGEON, S. 431 MILLER, H. 17 MI KOVI , A. 120,337 MOLINA, D. 500 MON, R. 144,261 MONALDI, R. 477 MONFRET, T. 136,184,235,377 MONTEILLER, V. 334, 339 MORA, C. 341 MORALES, C. 500 MORATA, D. 29,569 MOREIRAS, S. M. 344 MORENO, H. 191 MORENO, M. 326,348 MORIGUTI, T. 310 MORITZ, R. 566 MOSER, D. 481 MOSER, E. 439 MOTHES, P. A. 67,351 MPODOZIS, C. 56,354 MUÑOZ, Marcia 357 MUÑOZ, Miguel 361 MUÑOZ, V. 365 MÜNTENER, O. 421

N NAKAMURA, E. 310 NARANJO, J. A. 191 NAVARRO, P. 369 NERCESSIAN, A. 77 NORIEGA, L. 387

O O’BRIEN, G. S. 334 OLIVEROS, V. 29 OLLARVES, R. J. 37 ORDÓÑEZ, J. J. 387,571 OROZCO, L. A. 90,373 ORTEGA, V. 116

OTTONE, E. G. 33

P PAGE, L. 517 PANKHURST, R.J. 427 PAQUEREAU-LEBTI, P. 545 PARDO, M. 136,184,377,592 PAZOS, P. J. 33 PEDOJA, K. 381 PENNA, I. M. 384 PÉPIN, E. 387,571 PÉREZ, D. J. 88,391 PÉREZ, P. 116 PERFETTINI, H. 64,562 PERONI, J. I. 393 PIMENTEL, M. M. 423,577 PINO, A. 265 PINTO, L. 397,458 PIRAQUIVE, A. 273 POBLET, J. 401 POBLETE, F. 405 POLVÉ, M. 569 POMBOSA, R. 387 POMPOSIELLO, M. C. 90,373 PONTOISE, B. 84,223,302 PRESCOTT, C. 577 PREZZI, C. 409 PULGARÍN, B. 97,538 PUTLITZ, B. 421

Q QUEZADA, J. 413 QUINTEROS, J. 417

R RAMÍREZ de ARELLANO, C. 421 RAMÓN, P. 67 RAMOS, V. A. 33,210,227,289,417,423, 461,513,553 RANERO, C. R. 94 RAPELA, C.W. 427 RATZOV, G. 431 RAULD, R. 48 REGARD, V. 295,381,435,474,571 RÉGNIER, M. 84,302 REICHERT, C. 94,164 RÉMY, D. 136,168,238 REUBI, O. 191 REUTHER, C.-D. 439 REYES, P. 442 RIBODETTI, A. 156,306 RIQUELME, R. 21,172,571 RIVERA, M. 369,446,545 RIVERA, O. 592 ROBIN, C. 67,180 ROBINSON, D. 277 ROBLES, W. A. 273,450 RODDAZ, M. 199,435,454 RODRÍGUEZ, L. M. 37 RODRÍGUEZ, M. P. 458 RODRÍGUEZ-FERNÁNDEZ, L. R. 401 ROJAS-VERA, E. 461 ROMERO, D. 465 ROPERCH, P. 56,354,469 ROSAS, M. 583 ROSSELLO, E. A. 148,373,477 RUEGG, J.-C. 77 RUIZ, G. M. H. 67,473,508 RUSSO, R. 341

597

S SÁBAT, F. 261 SAEZ, M. 41 SAILLARD, M. 295,474 SAINT-BLANQUAT, M. de 485 SALAZAR, E. 354 SALAZAR, L. 477 SALAZAR, P. 481 SAMANIEGO, P. 67,180 SÁNCHEZ, A. 485 SÁNCHEZ, J. 489 SÁNCHEZ-MAGARIÑOS, J. M. 391 SANDVOL, E. 277 SCHALTEGGER, U. 120,337 SCHERWATH, M. 164,493 SCHILLING, M. 496 SCHMIDT, S. 242,409 SCHMITZ, M. 489,500 SEGGIARO, R. E. 401 SEGOVIA, M. 84,214 SELLÉS, D. 191 SEMPERE, T. 265,281,504 SEWARD, D. 508,573 SHAPIRO, S. A. 245 SHERIDAN, M. 315 SIELFELD, G. 116 SINGER, A. 500 SINGER, B. 545 SOCQUET, A. 77 SOLÍS, C. 44 SOMOZA, R. 509 SOSSON, M. 292,431 SOULA, J.-C. 140,172,454,521 SOURIOT, T. 281 SPAGNUOLO, M. G. 513 SPENCE, G. 306 SPIKINGS, R. 120,517,573,587 SRUOGA, P. 558 STRECKER, M. R. 326,555 STRUB, M. 521 SUÁREZ, M. 524 SUDO, M. 580 SUZAÑO, N. 577

T TAIPE, E. 334 TASSARA, A. 206,496,526 TASSI, F. 25,101,530 TASSONE, A. 393 TAVERA, H. 64,214,534,562 TÉLLEZ, L. 538 TERRIZZANO, C. M. 542 THIELE, R. 48 THOMSON, S. 580 THOURET, J.-C. 281,298,315,446,545 TIPTEQ group 245 TORO, G. E. 97,538 TORRES, P. 549 TOSELLI, A. J. 17

TRIC, E. 235 TUNIK, M. 423 TUNSTALL, C. 553

U UBA, C. E. 555 URBINA, N. E. 558 URIBE, V. M. 562

V VÁGÓ, E. 566 VALDÉS, A. 569 VALETTE, B. 339 VALLEÉ, M. 500 VALLEJO, C. 587 VALLEJO, S. 351 VAN DER BEEK, P. 81 VAN WESTEN, C. 315 VARGAS, C. A. 128,132 VARGAS, G. 48 VARGAS, R. 315 VASELLI, O. 25,101,530 VASSALLO, R. 571 VAUCHEL, P. 387 VELÁSQUEZ, A. 273 VELOSO, E. 52,116,238 VERA, E. 184,377,592 VERGARA, M. 29 VILAS, J. F. 393 VILLAGÓMEZ, D. 573 VILLALBA, R. 44 VILLASEÑOR, A. 84 VIRAMONTE, J. G. 577 VIRAMONTE, J. M. 577

W WAGNER, L. 75 WANG, K. 306 WHITEHOUSE, M. J. 120 WIEGAND, M. 477 WIGGER, P. 245,481 WILLNER, A. P. 580 WILSON, C. G. J. 583 WINKLER, W. 573,587 WÖRNER, G. 298,545

Y YAGUPSKY, D. 71 YÁÑEZ, G. 377,592 YATES, B. A. 223 YEPES, H. 67 YUAN, X. 277

Z ZAMORA-VALCARCE, G. 461 ZANDT, G. 75 ZEILINGER, G. 555