Propagation of orographic barriers along an active range front: insights from sandstone petrography...

Propagation of orographic barriers alongan activerange front: insights fromsandstone petrographyanddetrital apatite fission-track thermochronology in theintramontane Angastaco basin,NWArgentinaIsabelle Coutand,nw Barbara Carrapa,w Anke Deeken,z Axel K. Schmitt,‰ Edward R. Sobelwand Manfred R. StreckerwnUniversite¤ des Sciences etTechnologies de Lille 1, UMR-CNRS 8110, UFR des Sciences de laTerre (ba“ t.SN5),Villeneuve d’Ascq cedex, FrancewUniversit�t Potsdam, Institut fˇr Geowissenschaften, Potsdam, GermanyzInstitut fˇr Geowissenschaften, Freie Universit�t Berlin, Maltesestrasse, Berlin, Germany‰Department of Earth and Space Sciences, University of California, Los Angeles, CA, USA

ABSTRACT

The arid Puna plateau of the southern Central Andes is characterized by Cenozoic distributedshortening forming intramontane basins that are disconnected from the humid foreland because ofthe defeat of orogen-traversing channels.ThickTertiary and Quaternary sedimentary ¢lls in Punabasins have reduced topographic contrasts between the compressional basins and ranges, leading to atypical low-relief plateau morphology. Structurally identical basins that are still externally drainedstraddle the eastern border of thePuna anddocument the eastward propagation of orographic barriersand ensuing aridi¢cation. One of them, the Angastaco basin, is transitional between the highlycompartmentalized Puna highlands and the undeformed Andean foreland. Sandstone petrography,structural and stratigraphic analysis, combinedwith detrital apatite ¢ssion-track thermochronologyfrom a �6200-m-thickMiocene to Pliocene stratigraphic section in the Angastaco basin, documentthe late Eocene to late Pliocene exhumation history of source regions along the eastern border of thePuna (Eastern Cordillera (EC)) as well as the construction of orographic barriers along thesoutheastern £ank of the Central Andes.Onset of exhumation of a source in the EC in late Eocene time as well as a rapid exhumation of the

Sierra de Luracatao (in the EC) at about 20Ma are recorded in the detrital sediments of the Angastacobasin. Sediment accumulation in the basin began �15Ma, a time atwhich the EC had already builtsu⁄cient topography to prevent Puna sourced detritus from reaching the basin. After �13Ma,shortening shifted eastward, exhuming ranges that preserve an apatite ¢ssion-track partial annealingzone recording cooling during the late Cretaceous rifting event. Facies changes and fossil contentsuggest that after 9Ma, the EC constituted an e¡ective orographic barrier that prevented moisturepenetration into the plateau.Between3.4 and2.4Ma, another orographic barrier was uplifted to the east,leading to further aridi¢cation and pronounced precipitation gradients along the mountain front.Thisstudy emphasizes the important role of tectonics in the evolution of climate in this part of the Andes.

INTRODUCTION

TheAltiplano-Puna plateau stretches for �2000 km fromsouthern Peru to northwestern Argentina and is a ¢rst-or-der topographic feature of the Central Andes.The plateauis bounded on the west by active stratovolcanoes of the

Western Cordillera and on the east by the basement-involved thrust system of the Eastern Cordillera (EC)(Fig. 1a and b). Both Cordilleras reach elevations in excessof 6500m and delimit the plateau region of 3800m averageelevation.The present-day elevation of the plateau is com-monly attributed to crustal thickening in the course of dis-tributed tectonic shortening (e.g. Isacks, 1988; Allmendingeretal., 1997; Lamb et al., 1997; Kley &Monaldi, 1998;McQuar-rie, 2002), and subsequentwholesale uplift following delami-nation of mantle lithosphere in its southern part(ArgentinianPuna) (Kay etal.,1994;Allmendingeretal.,1997).

Correspondence: Isabelle Coutand, Universite¤ des Sciences etTechnologies de Lille 1, UMR-CNRS 8110, UFR des Sciencesde la Terre (ba“ t. SN5), 59655 Villeneuve d’Ascq cedex, France.E-mail: [email protected]

BasinResearch (2006) 18, 1–26, doi: 10.1111/j.1365-2117.2006.00283.x

r 2006 The Authors. Journal compilationr 2006 Blackwell Publishing Ltd 1

PUNA

PUNA

(a)

(b)

Fig.1. (a) Geological map of the southeastern Central Andes (redrawn from the geological maps of Salta, Catamarca and Tucumanprovinces, scale1 : 500 000). (b) DEMof the southern central Andes between 24^271S and 69^641W.Data are from SRTM90-m square-gridDEM.The heavydashed line shows the boundarybetween the internallydrainedPuna plateau and the foreland areas that drain to theeast. (c) Coloured mean annual precipitation distribution along the southeastern Central Andes based on data fromWMO (1975) andBianchi&Yan� ez (1992).Location of (a^c) is given by the black square in inset (a).Abbreviations areAcon,Aconquija range;Ang,Angastacobasin;Ant,Antofalla salar;Ar,Arizaro salar;Bar,SantaBarbara ranges;Ca,Calalaste range;Chan,ChangoReal range;OC,Oire intrusivecomplex; Gal, CerroGala¤ n; Lur, Cumbres de Luracatao;Mar, SantaMaria valley; Quil, Quilmes range; SRT, Santa Rosa deTast|¤ l.

r 2006 The Authors. Journal compilationr 2006 Blackwell Publishing Ltd,Basin Research, 18, 1^262

I. Coutandet al.

The Puna plateau comprises internally drained, par-tially over¢lled contractional basins where up to 6000mof late Eocene to Pliocene continental clastics and evapor-ites have accumulated (Jordan & Alonso, 1987; Kraemeret al., 1999; Adelmann, 2001; Coutand et al., 2001). Thesebasins are thrust-bounded and isolated from each otherby �N^S trending basement ranges, typically 1000^1500m higher than the plateau £oor (Coutand,1999; Cou-tand et al., 2001 and references therein) (Fig.1a and b).Thedeposition of thick evaporite units is related to internaldrainage and sustained aridity, at least during the last 27^14Myr (Alonso et al., 1991; Vandervoort et al., 1995; Adel-mann, 2001; Carrapa et al., 2005).The aridity results fromthe latitudinal position in a region of descending airmasses (e.g. Hartley, 2003; R. Alonso, B. Carrapa, I. Cou-tand, M. Haschke, G.E. Hilley, E.R. Sobel, M.R. Strecker&M.H.Trauth, submitted) and the condensation of Atlan-tic-derived moisture along the eastern (windward) £ank ofthe orogen (e.g. Masek et al., 1994; Haselton et al., 2002)(Fig.1c).

Sobel etal. (2003) andHilley&Strecker (2005) have sug-gested a model for the Puna plateau and adjacent regionsin which the low internal relief, aridity and internal drai-nage are intimately linked and have persisted in this regionbecause tectonic uplift has overwhelmed the ine⁄cient£uvial system and caused orogen-traversing channels tobe defeated, thus disconnecting the interior of the orogenfrom the foreland base level. The combination of an aridclimate and tectonic activity along range fronts impliesthat progressive aridi¢cation of the hinterland is pro-motedwhen laterally continuous orographic barriers alongthe margin of the plateau are perpendicular to moisture-bearing winds. If aridity is coupled with low erodibility,and moderate to high uplift rates of orographic barriers,the channel network draining the hinterland will be frag-mented.Consequently,material removed from the leewardside of the uplifts will ¢ll internally drained basins. The

failure of erosional mass export will increase gravitationalstresses in the hinterland and eventually promote the mi-gration of contractional deformation into the foreland, ad-vancing orographic barriers and precipitation in the samedirection (Sobel & Strecker, 2003; Sobel et al., 2003; Hilley& Strecker, 2004).

To further evaluate this model and to understand theupper crustal and sur¢cial mechanisms that promote theformation, maintenance and lateral growth of a highplateau, it is important to better constrain the space^timedevelopment of orographic barriers. Most estimates forthe timing of shortening and associated crustal thickeningand surface uplift in NW Argentina rely on the synoro-genic record of contractional basins within and adjacentto the orogenic belt (Jordan & Alonso, 1987; Strecker et al.,1989; Allmendinger et al., 1997; Jordan et al., 1997; Kraemeretal., 1999; Coutand etal., 2001).However, across and at theborder of the southern Puna, little is known about (1) theonset of tectonic shortening as a proxy for mountain build-ing and (2) the spatial distribution and propagation of con-tractional deformation through time.Whereas some arguethat the propagation of the contractional front was uni-form in space from west to east with shortening startingwithin the plateau in the earlyMiocene (�15^20Ma) (Jor-dan&Alonso,1987; Allmendinger etal., 1997) and reachingthe plateau margin by the lateMiocene (10^6Ma) (Streck-er et al., 1989; Allmendinger et al., 1997;Marrett & Strecker,2000), other studies suggest that early contraction and up-lift occurredwithin and at the eastern periphery of the pla-teau as early as the late Eocene^Oligocene time (Kraemeret al., 1999; Adelmann, 2001; Coutand et al., 2001; Carrapaet al., 2005), similar to observations made to the north inthe Bolivian Andes (Horton et al., 2002; Horton, 2005).

In this study, we evaluate the tectono-sedimentary his-tory of the southeastern Puna margin by reconstructingthe evolution of the intramontane Angastaco basin, lo-cated between the plateau and eastern foreland ranges.The Angastaco basin is part of the semi-arid Calchaqu|¤Valley receiving �200mm of rainfall annually; it is thustrapped between the arid Puna plateau to the west(�100mmyr�1) and the humid Chaco lowlands of theAndean foreland to the east (1200mmyr�1) (Haseltonet al., 2002) (Fig. 1c). Our new structural, sedimentologic,provenance and detrital apatite ¢ssion-track data fromthe Neogene Angastaco basin help reconcile di¡erentviews on the timing of plateau evolution and constrainthe Cenozoic structural and palaeoecologic evolution ofthe southeastern plateau margin. We demonstrate thatcontractional deformation began in the late Eocene, simi-lar to other areas on the southern Puna plateau. Further-more, we show that the migration of compressionaldeformation and related building of orographic barriersfocused precipitation on the eastern, windward £anks ofranges east of the Puna. This starved the western hinter-land of moisture and thus promoted the maintenance ofinternally drained areas in the plateau and the creation oftransient internally drained basins in the transition be-tween the plateau border and the foreland.

PUNA

(c)

Fig.1. Continued

r 2006 The Authors. Journal compilationr 2006 Blackwell Publishing Ltd,Basin Research, 18, 1^26 3

Propagation of orographic barriers along an active range front

REGIONAL GEOLOGICAL SETTING

The narrow, N^S trending Angastaco basin is located inthe distal part of a once contiguous Palaeogene RetroarcForeland basin (Jordan & Alonso, 1987) (Fig. 1a). The re-verse-fault-bounded basin is located in the semi-aridCal-chaqu|¤ valley between the EC to the west, and the forelandranges of the Santa Barbara structural province to the east,an inverted Cretaceous extensional province (Grier et al.,1991; Mon & Sal¢ty, 1995) (Fig. 1a and b). On the west, thebasin is overthrust by late Proterozoic to early Cambrianmetasediments of the Puncoviscana Formation (Fig. 2aand b), and to the east it is limited by outcrops of Cretac-eous coarse clastic rift-related sediments (Marquillas etal.,2005). The Angastaco basin ¢ll comprises continentalclastic deposits of the latePalaeogene toNeogenePayogas-tilla Group (D|¤ az & Malizzia, 1983), which is tilted east-ward and gently folded (Fig. 2b).

The Puncoviscana Formation contains low-grade me-tamorphosed sandstones and shales (Turner, 1960;Toselli,1990; Durand, 1992) that grade southward into schists,gneisses and migmatites (La Paya Fm. and Tolombo¤ ncomplex).These units are intruded by Cambrian to Ordo-vician granites (Lork et al., 1990; Lork & Bahlburg, 1993;Omarini et al., 1999). Further west, the north^south-or-iented Oire intrusive complex of Ordovician granites andgranodiorites crops out (Me¤ ndez et al., 1973; Omarini etal., 1984; Lork & Bahlburg, 1993) (Figs 1a and 2a). Thesebasement lithologies constitute the source for the Payo-gastilla Group. The basement rocks are unconformablyoverlain by the rift-related Neocomian to early Eocenestrata of the Salta Group (Bianucci et al., 1981; Grier et al.,1991; Sal¢ty &Marquillas, 1994).

The Payogastilla Group (D|¤ az & Malizzia, 1983) restsunconformably on these units and consists of four forma-tions (Figs 2a, b and 3): (1) the basalQuebrada de los Color-ados Formation comprises sequences of coarse-grainedsandstones (sometimes conglomeratic) that grade upwardinto ¢ne- to medium-grained sandstones and siltstoneswith palaeosols (D|¤ az et al., 1987; Grier, 1990), interpretedto be alluvial plain deposits (D|¤ az & Malizzia, 1983). Itsage is poorly constrained as either early Miocene (D|¤ azet al., 1987) or Oligocene (Starck & Vergani, 1996). In theAngastaco basin, the base of the formation is not exposedbecause of faulting along the eastern side of Cerro Negro(Fig. 2a and b), but the unit is 340-m thick in the Pucara¤valley in thewest (Ruiz,1993). (2)The unconformably over-lying 4200-m-thick Angastaco Formation is composed ofgrey to beige conglomerates, sandstones and muddy silt-stones organized in upward- ¢ning £uvial cycles. In thelower and upper parts of the section, sandy and silty tex-tures dominate and the sequence is interpreted as a braidedstream system typical of a piedmont plain (D|¤ az & Maliz-zia, 1983) (Fig. 3). In the middle part of the section, con-glomeratic lenses intercalated with coarse-grainedsandstones and massive conglomeratic layers increase, in-dicating a distal to proximal alluvial fan origin (D|¤ az &Malizzia, 1983; D|¤ az et al., 1987; D|¤ az & Miserendino

Fuentes, 1988). The virtual absence of fossils and the pre-sence of desiccation polygons suggest that these stratamay have been deposited under arid to semi-arid condi-tions similar to coeval units in the Santa Mar|¤ a basin tothe south (D|¤ az & Malizzia, 1983; D|¤ az & MiserendinoFuentes, 1988; Bossi et al., 2001; Starck & Anzo¤ tegui,2001). A tu¡ at the base of the Angastaco Formation datedby the 40Ar/39Ar method on biotite yielded an age of13.4� 0.4Ma (Grier &Dallmeyer, 1990), indicating a mid-dle Miocene age. (3) The Palo Pintado Formation is con-formable with the underlying Angastaco Formation andcan be divided into two parts (Fig. 3). The lower 1500mconsist of ¢ne- to coarse-grained sandstones interlayeredwith siltstones and mudstones, interpreted as meanderingriver deposits locally withwetland facies (D|¤ az &Malizzia,1983;D|¤ az &Miserendino Fuentes,1988).The upper 450mcomprise an upward-coarsening succession of coarser-grained sandstones inter¢ngered with massive conglom-eratic strata and intercalated tu¡ beds, interpreted asbraided stream deposits (D|¤ az & Malizzia, 1983; D|¤ az &Miserendino Fuentes, 1988). Mammal fossils in the lowersection indicate a late Miocene to early Pliocene age andpalaeoenvironments characterized by relative humidity(Marshall et al., 1983; Anzo¤ tegui, 1998; Starck & Anzo¤ tegui,2001). A volcanic ash at the top of this unit (sample AI29,Fig. 3) yielded a U^Pb zircon age of 5.27� 0.28Ma (seethe next section). (4) Finally, the San Felipe Formation cov-ers the Palo Pintado strata and consists of 650-m-thickconglomeratic upward- ¢ning sequences (Grier, 1990), in-terpreted as braided stream deposits transitional with dis-tal alluvial fan deposits (D|¤ az & Malizzia, 1983; D|¤ az &Miserendino Fuentes, 1988). Capped by an angular uncon-formity, these units are overlain by gravels that contain re-cycled pyroclastic material at the base, 2.4Ma old (M. R.Strecker, unpubl. data).

U^PB ZIRCON DATING

To better constrain the stratigraphic age of theNeogene se-diments of theAngastaco basin, an ash located at the top ofthe Palo PintadoFormationwas dated by theU^Pbmethodon zircons (Table 1). Epoxy grain mounts of hand-selectedzircons were sectioned to expose grain interiors andpolished with 1mm Al2O3. After ultrasonic cleaning withsoapywater, dilutedHCl and distilledwater, theAu-coatedmounts were transferred into a high vacuum chamber(410� 8Torr) and kept overnight. Zircon analysis was per-formed using the UCLACameca ims1270 ion microprobewith a mass- ¢ltered, �15nA 16O� beam focused to an�30^35mm diameter spot. The analysis surface was£oodedwith O2 at a pressure of �4� 10� 3Pa to enhancePb1 yields. Secondary ions were extracted at 10 kV with anenergy band pass of 50 eV. Following a 4min pre-sputterperiod to reduce surface contamination intensitiesfor 94Zr2O

1, 204Pb1, 206Pb1, 207Pb1, 208Pb1, 238U1,232Th16O1 and 238U16O1 were then sequentially measuredin 10 cycles at a mass resolution of �5000, which is su⁄ -cient to resolve most molecular interferences.The relative


I. Coutandet al.

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Fig. 3. Stratigraphic column of the lateNeogene sedimentary ¢ll of the Angastaco basin measured in the ¢eld (this study). Symbols areas follows: white stars indicate detrital apatite ¢ssion-track samples; grey squares, sandstone petrography samples; white squares, clastcounts in conglomeratic layers. a, age of the bottom of the Angastaco Fm., calculated from b and c assuming a constant sedimentationrate; b, CorteElCan� on tu¡dated by

39Ar/40Ar on biotite (Grier&Dallmeyer,1990); c,mammal fossils (Marshall etal.,1983); d, 206Pb/238Uzircon age (this study).


I. Coutandet al.

sensitivities forPb andUwere determined on reference zir-con AS-3 (Paces &Miller, 1993) using a calibration techni-que similar to Compston et al. (1984). Th and U contentswere calculated from 232Th16O1/94Zr2O

1 and 238U16O1/94Zr2O

1 with relative sensitivities calibrated on referencezircon 91500 (Wiedenbeck et al., 1995), and are reported inTable1. 206Pb/238Uageswere calculated from theU^Pb iso-topic ratios by linear regression and by computing theweighted average of individual results as outlined inSchmitt et al. (2002). In both calculation schemes, a com-mon 207Pb/206Pb typical for anthropogenic sources wasused (e.g. San� udo-Wilhelmy&Flegal, 1994). Both methodsyield ages that are in close agreement: 6.16� 0.42Ma(mean square of weighted deviates MSWD5 8.4) and5.76� 0.42Ma (MSWD5 5.1; uncertainties at the 95%con¢dence levels with errors scaled by the square root ofMSWD). The consideration that zircons from an earliermagmatic phase could be present in the sample led us toexclude some of the older zircon results (open symbols inFig. 4,Table 1) and recalculate the ages for the younger zir-con population. Excluding three older grains signi¢cantlyreduces theMSWDtovalues close to one, andwe obtainedmodi¢ed age estimates forAI29,whichwere 5.29� 0.27Ma(intercept age) and 5.27� 0.28Ma (weighted average age;Fig. 4). These ages are thought to represent the time ofzircon crystallization in the magma that produced theAI29 ash.

SANDSTONE AND CONGLOMERATEMODAL PROVENANCE ANALYSES

Method

Provenance analyses on both sandstones and conglomer-ates were carried out in the Angastaco and Palo PintadoFormations in order to identify di¡erent source terrainsand to relate possible upsection changes to source area re-organization through time (e.g. Ingersoll et al., 1984).Thir-

teen sandstones indicated in Fig. 3 were analysed using theGazzi^Dickinson method (G^D) (Gazzi,1966; Dickinson,1970). All constituents, and at least 250 framework grains,were counted in thin sections following well-establishedcounting procedures (Valloni, 1985; DiGiulio, 1990). Weused a detailed counting form (Table 2) in order to re-cal-culate and present the data according to both theG^D andthe Folk methods (Folk, 1968; Basu, 1976) following theprocedure proposed by DiGiulio and Valloni (1992).Table2 contains detailed information on the framework cate-gories (quartz, feldspars and rock-lithic fragments) de¢n-

0 500 1000 1500 2000 2500 30000.0

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age

[Ma]

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6.5

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w.m.: 5.27±0.28 Ma (95% confidence, MSWD = 0.8)

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Fig.4. 207Pb/206Pb vs. 238U/206Pb of AI29 zircons (uncorrectedfor common Pb; error bars 1s). Xenocrystic zircon g9 (Table1)has been omitted for clarity. Also note that three apparently olderzircons (open symbols) have been excluded in the regression.Theregression age is obtained from intersection of concordia(modi¢ed for disequilibrium assuming initial activity ratio230Th/238U5 0.15) from linear extrapolation throughuncorrected datawith a ¢xed intercept (207Pb/206Pb5 0.83). Insetshows individual 207Pb-corrected ages and the weighted averageage fromyounger zircon population (n5 7; solid symbols).

Table1. ZirconU^Pb results for ash AI29

Analysis

206Pbn/238U� 10� 3

206Pbn/238U1SE� 10� 3

207Pbn/235U� 10� 3

207Pbn/235U1SE� 10� 3

Correlationof concordiaellipses

207Pb corrected.disequilibriumcorrected 206/238 aget (Ma)

1st (Ma)

Radio-genic(%)

U(p.p.m.)

Th(p.p.m.)

AI29-g1 0.9283 0.043 16.34 1.06 0.26 5.42 0.28 89.6 270 275AI29-g2 0.9161 0.0395 20.42 3.06 0.56 5.11 0.31 85.2 535 360AI29-g3 0.9709 0.0262 8.413 0.442 0.85 6.21 0.21 97.9 2164 827AI29-g4 0.8209 0.0504 8.226 1.1 0.89 5.18 0.33 96.6 376 295AI29-g5 0.8393 0.0571 6.765 0.748 0.85 5.42 0.37 98.4 2113 463AI29-g6 0.8631 0.0499 8.23 0.608 0.50 5.49 0.32 97.1 1113 518AI29-g7 0.7757 0.032 6.207 0.572 0.78 5.01 0.21 98.5 1229 674AI29-g8 0.9032 0.0639 6.428 0.503 0.78 5.87 0.41 99.3 4868 2095AI29-g9 143 2 1404 26.2 0.78 858 28 99.5 1068 120AI29-g10 1.005 0.0156 6.53 0.146 0.83 6.54 0.22 99.9 9436 7407AI29-g11 0.9518 0.0186 6.655 0.37 ^0.10 6.18 0.20 99.4 1581 936



Tab

le2.

Sand

ston

epetrographyresults

Parameters

Form

ation:

samplenu

mber

Ang

astaco

Form

ation

PaloPintadoForm

ation

QFL

1C

QFR

Ans

4Ans

5Ans

6Ans

9Ans

13Ans

14Ans

15Ans

17Ans

19Ans

22Ans

28Ans

30Ans

32

QQ

Quartz(Q

Z)straightextinction

1.91.2

1.6

1.922.5

3.2

15.1

0.4

2.0

2.4

8.115.5

1.4

QQ

QZwavyextinction

9.912.4

15.0

10.5

4.114.9

4.2

3.8

13.1

15.2

1.8

1.3

21.7

QQ

Coarsepo

lycrystalline

QZ

3.5

5.14.1

6.1

4.7

10.6

9.62.5

7.55.9

6.0

3.3

9.6Q

QFine

polycrystalline

QZ

0.1

0.0

0.2

0.2

0.0

0.3

0.0

0.6

0.8

0.4

0.3

0.7

1.4

QR

Chert

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

QR

QZ(4

62mm

)inpluton

icrock

3.6

1.6

1.2

0.5

2.0

5.2

2.1

0.3

0.8

1.4

1.0

1.3

2.0

QR

QZ(4

62mm

)inmetam

orph

icrock

3.8

2.8

2.8

2.4

5.6

4.6

11.7

4.3

5.3

2.4

5.7

8.9

4.9F

RQZ(4

62mm

)involcanicrock

0.1

0.0

0.0

0.0

0.0

0.0

0.0

0.2

0.2

0.0

0.0

0.0

0.0

FF

Kmon

ocrystallin

e(altered)

4.7(1.9)

12.4(2.6)

8.1(1.3)

8.3(2.2)

16.1(3.5)

15.7(3.7)

7.54.0

7.9(1.8)

10.7(1.2)

9.6(1)

11.5(3.3)

6.7(0.9)

FR

K(4

62mm

)inpluton

icrock

(altered)

1.5(0.1)

2.8(1)

2.1

4.1(1.5)

4.1(1.2)

4.3(0.6)

3.6(1.2)

0.4

0.4

2.8(0.6)

2.6

3(0.7)

2.9

FR

K(4

62mm

)inmetam

orph

icrock

(altered)

1.73.2

3.3

4.8

8.5(1.8)

7.411.1(1.8)

3.4

6.3

4.2

9.9(0.3)

15.8(1.3)

11.6

FR

K(4

62mm

)involcanicrock

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

FF

Pmon

ocrystallin

e(altered)

1.0

0.4

1.0

0.5

2.0

0.6

1.8

1.0

1.0

0.2

0.3

2.0

0.0

FR

P(4

62mm

)inpluton

icrock

0.0

0.0

0.0

0.0

0.0

0.0

0.9

0.0

0.0

0.0

0.0

0.3

0.0

FR

P(4

62mm

)inmetam

orph

icrock

0.0

0.2

0.0

0.0

0.3

0.3

1.2

0.0

0.6

0.2

0.0

3.0

0.0

FR

P(4

62mm

)involcanicrock

0.0

0.0

0.2

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Ls

RMetam

orph

iclithic(withconstituentso62

mm)

2.6

4.91.8

4.15.8

4.6

8.4

6.8

7.15.1

20.8

17.8

13.0

Ls

RVolcaniclithic

2.9

2.2

0.2

0.0

1.2

0.0

0.0

0.0

0.6

0.4

0.3

0.0

0.0

Ls

RClasticlithic

0.1

0.0

0.0

0.0

0.0

0.3

0.0

0.0

0.0

0.0

0.0

0.0

0.0

xxx

xxx

Mon

ocrystallin

ebiotites/chlorites

0.6

0.2

2.8

0.0

2.3

0.3

0.6

0.8

0.0

0.6

1.8

1.72.3

xxx

xxx

Mon

ocrystallin

emuscovites

0.4

0.4

1.0

0.5

1.8

0.6

0.9

0.6

0.0

0.2

1.3

1.0

0.9

xxx

RMica(4

62mm

)incrystalline

rock

fragment

0.6

0.0

0.2

0.0

0.9

0.9

2.4

0.6

0.2

0.0

0.8

3.3

0.0

xxx

RMica(4

62mm

)involcanicor

clasticrock

fragment

0.6

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.2

0.0

0.0

0.0

0.0

xxx

xxx

Mon

ocrystallin

eheavyminerals

0.1

0.2

0.3

0.0

0.3

0.3

0.9

0.1

0.0

0.2

0.0

0.0

0.0

xxx

RHeavy

mineral(4

62mm

)incrystalline

rock

0.0

0.0

0.2

0.2

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.3

xxx

RHeavy

mineral(4

62mm

)involcanicor

clasticrock

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Ls

RCalc-schist

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

CR

Sparry

calcite

0.0

0.0

0.7

0.0

0.0

0.0

0.0

0.0

0.2

0.0

0.0

0.0

0.0

CR

Dolom

ite

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

CR

Impu

recalcareous

grain

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

CR

Mud

ston

e^wackstone

0.0

0.0

0.0

0.0

0.0

0.0

0.3

0.0

0.0

0.0

0.0

0.0

0.0

CR

Packston

e^grainstone

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Totalframew

orkgrains

31.5

33.9

34.1

31.0

48.2

51.9

63.0

27.8

46.0

37.8

46.8

54.1

68.7

Glaucon

ite0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Peliticclast

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Bioclast

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Alte

redmaterial

3.5

3.4

8.7

1.72.6

0.3

4.2

1.12.0

4.0

3.11.0

3.2


I. Coutandet al.

Tab

le2.

(Con

tinu

ed)

Parameters

Form

ation:

samplenu

mber

Ang

astaco

Form

ation

PaloPintadoForm

ation

QFL

1C

QFR

Ans

4Ans

5Ans

6Ans

9Ans

13Ans

14Ans

15Ans

17Ans

19Ans

22Ans

28Ans

30Ans

32

Authigenicmineralin

unkn

owngrain

0.0

0.0

0.2

0.7

0.3

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Poreinun

know

ngrain

0.3

0.2

0.0

0.0

0.0

0.0

0.0

0.0

0.0

1.2

0.0

0.0

0.3

Coarsesilt

0.9

0.8

1.8

2.0

1.2

0.0

1.2

0.8

1.6

4.0

3.9

3.6

0.3

Siliciclasticmatrix

3.8

6.9

12.7

4.4

2.6

13.5

1.5

2.2

4.7

9.14.9

3.3

6.4

Carbonatematrix

0.0

0.0

0.7

0.2

1.5

0.3

0.6

0.0

0.0

0.0

0.0

0.0

0.0

Qzcement

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Phyllosilicatecement

10.6

7.722.4

2.6

7.35.2

8.7

4.0

6.9

4.4

0.8

1.3

3.8

Carbonatecement

0.1

0.0

2.1

2.4

1.8

0.3

1.2

0.0

0.0

0.0

0.0

0.0

0.3

Othercements

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Interg

ranu

larpo

rosity

40.8

31.2

1.3

41.8

0.0

6.6

0.0

61.9

30.4

24.8

17.1

0.0

7.0Calciticovergrow

th0.0

0.0

3.5

0.2

0.3

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Oversized

pore

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

0.0

Totalrock

100

100

100

100

100

100

100

100

100

100

100

100

100

QFL

Q61

4759

5050

5455

4455

5435

3755

F24

3935

4141

3934

3230

3534

4228

Ls1

C15

146

99

711

2415

1132

2117

Average

1standarddeviation

Q53

610

45F

356

634

Ls1

C12

68

21QFR

Q40

3849

4340

4036

2643

4624

2445

F44

4744

4851

5453

5141

4345

5638

R17

147

99

711

2415

1131

2017

Average

1standarddeviation

Q39

712

34F

494

745

R12

68

21Finelithics

Lm

Fine

QZ

1.15.3

1.4

xxx

3.2

5.8

xxx

xxx

1.15.1

2.9

4.3

xxx

Lm

Chert

0.0

0.0

6.9

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Lm

Radiolarite

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Lm

Quarzite

1.10.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.7

xxx

Lm

SchistQZ

1biotite

14.9

13.2

9.7xxx

19.4

25.6

xxx

xxx

10.6

7.65.7

17.0

xxx

Lm

SchistQZ

1biotite

1muscovite

1

feldspar

1heavymineral

8.5

52.6

16.7

xxx

14.5

24.4

xxx

xxx

16.2

11.4

27.1

14.2

xxx

Con

tinu

ed



Tab

le2.

(Con

tinu

ed)

Parameters

Form

ation:

samplenu

mber

Ang

astaco

Form

ation

PaloPintadoForm

ation

QFL

1C

QFR

Ans

4Ans

5Ans

6Ans

9Ans

13Ans

14Ans

15Ans

17Ans

19Ans

22Ans

28Ans

30Ans

32

Lm

Phyllite

1.115.8

9.7xxx

6.5

4.7

xxx

xxx

12.3

34.2

17.1

40.4

xxx

Lm

Fine

micaschist(o32

mm)

0.0

0.0

2.8

xxx

1.6

3.5

xxx

xxx

1.10.0

2.9

1.4

xxx

Lm

SchistQZ

1muscovite

16.0

10.5

20.8

xxx

24.2

33.7

xxx

xxx

31.8

15.2

30.0

14.9

xxx

Lm

SchistQZ

1muscovite

1heavymineral

1.12.6

15.3

xxx

29.0

2.3

xxx

xxx

25.1

17.7

14.3

5.0

xxx

Lm

SchistQZ

1epidot

1chlorite

1ph

illosilicate

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Lm

Chlorite

Schist

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Lv

Volcanics(crystalline

matrix)

43.6

0.0

13.9

xxx

0.0

0.0

xxx

xxx

1.78.9

0.0

2.1

xxx

Lv

Glass

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Lv

Acidvolcanics

0.0

0.0

1.4

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Lv

Middle-basicvolcanics

0.0

0.0

1.4

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Lv

Pyroclastics

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Ls

Argillite

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Ls

Siltston

e0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Ls

Calcsiltston

e0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

CMud

ston

e^wackstone

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

CDolom

ite

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

CSp

arry

calcite

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

CPackston

e-grainstone

0.0

0.0

0.0

xxx

0.0

0.0

xxx

xxx

0.0

0.0

0.0

0.0

xxx

Totalaph

aniticgrains

100

100

100

xxx

100

100

xxx

xxx

100

100

100

100

xxx

Lm^L

v^Ls1

CLm

49100

82xxx

98100

xxx

xxx

9891

100

98xxx

Lv

510

18xxx

00

xxx

xxx

29

02

xxx

Ls1

C0

00

xxx

20

xxx

xxx

00

00

xxx

Italicnu

mbersinbracketscorrespo

ndto

alteredgrains

ofthesamemineral.


I. Coutandet al.

ing the parameters for the QFLs1C (G^D method) andthe QFR (Folk method) diagrams, respectively (Fig. 5aand b).Additional counting of the ¢ne lithic fraction (lithicgrainswith constituentso62mm)was also performed (Ta-ble 2). Only nine of the 13 samples provided results andonly two (Ans 20 and 32) had more than 100 ¢ne lithics.However, a qualitative description of the lithics is pro-vided, together with a ternary diagram (Lm^Lv^Ls1C;DiGiulio & Valloni, 1992), based on the available data (Fig.5c). Pebble composition was also analysed in the Payogas-tilla Group by counting at least100 pebbles at 25 countingstations in the ¢eld (Fig. 5d).

Results

The compositions of the Angastaco and Palo Pintado For-mations are feldspathic to feldspatho-lithic, with an in-creasing lithic component up sequence (R, Ls1C; Fig. 5aand b). The Lm^Lv^Ls1C diagram shows an increase ofmetamorphic and a corresponding decrease in volcanicclasts upsection (Fig. 5c). Metamorphic lithics are mainlyquartz-muscovite schist and phyllite, whereas volcaniclithics are altered and di⁄cult to di¡erentiate.

Clast count data re£ect an up-section increase in meta-morphic clasts vs. a decrease of granitic andvolcanic clastsupsection; volcanics are absent from Cc 11 upward (Fig.5d), in agreement with the sandstone petrography results.Finally, from the top of the Palo Pintado Formation andupsection, a minor (not re£ected in the plots), but diagnos-tic contribution of red sandstone and limestone clasts ofthe Cretaceous Salta Group is notable. Limited publishedpalaeocurrent data in the Angastaco and Palo Pintado For-mations indicate a source in the adjacent EC to the westand northwest (Sal¢ty et al., 1984; D|¤ az & MiserendinoFuentes, 1988) (Fig.1a and b).

The evidence of a source of volcanic detritus in thePuna interior is ambiguous because the volcanic lithiccontents of sandstones and conglomerates are very minoror absent. However, the feldspathic to feldspatho-lithiccomposition of the sandstones could represent a contam-ination from ignimbrites that are widespread on the Punaand/or an eolian contribution by airfall tu¡s (Figs1a and 3).Occasional volcanic lithics in the basal Angastaco Forma-tion and their eventual disappearance up-section suggestthat a morphological high acting as a barrier already ex-isted to thewest-northwest of the basin in themiddleMio-cene, preventing volcanic material from the Puna frombeing £uvially transported eastward (Fig.1a).

DETRITAL APATITE FISSION-TRACKTHERMOCHRONOLOGY

Fission-trackmethod

The majority of detrital ¢ssion-track studies to date haveutilized zircon ¢ssion-track analysis because this thermo-chronologic system is less susceptible to post-depositionalthermal overprinting (e.g. Hurford, 1986; Brandon et al.,1998). However, we have used apatite crystals because inthe study area, there has been too little Cenozoic exhuma-tion to expose signi¢cant volumes of young zircons(Andriessen & Reutter, 1994; Coutand et al., 2001; Deekenet al., 2004).

Fission-track thermochronology is based on the spon-taneous ¢ssion decay of 238U in uranium-bearing miner-als, which creates a linear damage zone (a ¢ssion track) inthe crystal lattice. For apatite, tracks begin to be retainedat temperatures below ca. 150^120 1C, depending on thecooling rate and the kinetic characteristics of the crystal(e.g. Green et al., 1985; Ketcham et al., 1999). Determining

F

Q

Ans4Ans14

Ans9

Ans6Ans5

Ans13

Ang19

Ans15Ans28Ang17

Ans30Ans22

Ans32

R Ls+C

Q

F

Ans30

Ans15Ans14

Ans4

Ang17

Ang19Ang22

Ang32

Ans28

Ans13-Ans9

Ans6

Ans5

(a) (b)

Fig. 5. Provenance discrimination diagrams. (a^c) Sandstone ternary diagrams (QFR, QFLs1C and LmLvLs1C) from point-counting analysis; rawdata for the calculation of speci¢c parameters are reported inTable 2. (d)Provenance discrimination diagram frompebble-counting analysis.



So

urc

e d

iscr

imin

atio

n f

rom

co

ng

lom

erat

es

Ang

asta

co F

orm

atio

nP

alo

Pin

tado

F

orm

atio

n

020406080100

San Felipe Formation

Gra

nite

Sye

nite

Met

amor

phic

(Pun

covi

scan

a ty

pe)

Peg

mat

ite

Qua

rtzi

te

Vol

cani

cs (

Sed

imen

tary

*)

Ang

asta

co F

orm

atio

n

Pal

o P

inta

do F

orm

atio

n

Ave

rage

Tren

d up

-sec

tion

Lm

Ls+

CLv

Ans

7

Ang

15

Ans

13

Ans

4

Ang

19Ans

30-A

ng17

Ans

28-

Ans

32

(c)

(d)

Fig.5.

Con

tinu

ed


I. Coutandet al.

the density of spontaneous ¢ssion tracks and the concen-tration of 238U yields a single-grain ¢ssion-track age (e.g.Naeser, 1976). Fission tracks are unstable features thatshorten (anneal) by thermal and time-dependent re-crys-tallization processes. Over geological timescales, signi¢ -cant track shortening occurs in £uorine-rich apatitebetween 60 and 110 1C (e.g. Laslett et al., 1987; Laslett &Galbraith, 1996; Donelick et al., 1999); this temperaturerange is called the partial annealing zone (PAZ) (Gleadow& Duddy, 1981; Green et al., 1986). At su⁄ciently low tem-peratures, both old and new tracks are fully retained in thecrystal, whereas at increasing temperatures within thePAZ, newly formed tracks are long and older tracks areprogressively shortened. Therefore, a track-length distri-bution is a measure of the time spent within and belowthe PAZ.

Interpretation of a ¢ssion-track data set in terms of aT^t path requires an integrated analysis of ¢ssion-trackage, track-length distribution and kinetic characteristics ofthe apatite grains.Thermal modelling programmes (Galla-gher, 1995; Ketcham et al., 2000) incorporate quantitativeannealing models for apatite crystals with di¡erent kineticcharacteristics (Laslett et al., 1987; Ketcham et al., 1999).

Twelve sandstone samples, �6^10 kg in weight, werecollected through the continuous �6200-m-thick An-gastaco section, with an average sample spacing of�600m (Table 3, Fig. 3). In the laboratory, bulk sampleswere crushed, washed, sieved and apatite crystals were se-parated using standard magnetic and heavy liquids meth-ods. Analytical details are presented inTable 3. About 100grains were dated per detrital sample (Tables 3 and 4). Foreach sample, ¢ssion-track grain-age (FTGA)distributionswere decomposed following the binomial peak- ¢t method(Galbraith &Green, 1990) incorporated into BINOMFITprogram (Brandon, 1992, 1996; for details see, Brandon,2002) to identify discrete populations (or components, P1to Pn). The ages within each component are by de¢nitionconcordant. Con¢ned track lengths were only measuredin crystals that were dated; the angles between the con-¢ned track and the C-crystallographic axis (C-axis pro-jected data) and the etch pit diameter (Dpar; Donelicket al., 1999) were also measured.

Results

The FTGA distributions typically show a large single-grain age span, from166 to 2Ma (Fig. 6), with central agesranging between 62.6 � 2.4 and 28.9� 2.5Ma (Table 3). Allsamples clearly fail the w2 test (Galbraith, 1981; Green,1981) with P(w2)5 0 (Table 3), indicating that FTGA dis-tributions are discordant and are a mixture of componentsor peak ages (Brandon, 1992, 1996). All observed FTGAdistributions were decomposed into their main grain-agecomponents or peaks andyielded three to ¢ve componentsper sample (P1, . . . , Pn from the youngest to the oldestpeak,Table 4). Assignment of each FTGA component to adetritalDn path (Ruiz etal., 2004) or trend and consequentinterpretations are presented below.

Con¢ned track-lengths and Dpar measurements wereobtained from the greatest possible number of datedgrains.UsingBINOMFIT, each singular dated grain, tracklength and Dpar measurement was assigned to a speci¢cpeak age. Thus, we obtained representative track-lengthdistributions for 14 peak ages (number of tracks 420 ex-cept in three cases), mainly from P4 and P5 (for details, seeTable 5), as expected from their older apparent ages, andhence higher ¢ssion-track density.The mean track lengthsdid not vary signi¢cantly throughout the sedimentary sec-tion (Table 5, Fig. 7); values ranged between 11.55 � 0.31and12.93 � 0.26mm.There was no clear trend in the meantrack-length vs. stratigraphic position (Fig. 7). All popula-tions yielded similar Dpar values of 1.64^1.94mm (Table 5),suggesting that all populations contained monocomposi-tional apatites with similar low resistance to annealing.

EVIDENCE AGAINST THERMALRESETTING

The interpretation of detrital apatite FTages depends onwhether the ages have been partially reset by post-deposi-tional burial and reheating.We are particularly concernedabout this problem because of the large thickness of thesampled sedimentary sequence (46000m). However, forthe following reasons, our observations argue stronglyagainst signi¢cant post-depositional annealing:

(1) All the FTGA distributions are discordant; evenwith multiple ages, there could still be a possibility thatless retentive grains have been selectively reset and giveyounger peak ages (Garver et al., 1999). However, all popu-lations have apatites with similar kinetic characteristics(Table 5), ruling out this possibility. (2) All resolvable peakages in the distributions are older than the depositionalage (Table 4); furthermore, very young age peak recur-rently appear close to the depositional age but neveryounger than it. (3) Both the central ages and the peak agesincrease systematically down-section, whereas partiallyannealed samples would show the opposite (Tables 3 and4). (4) Finally, progressive burial heating should be ex-pressed as a decrease in the mean track-lengthwith depth;however, such a variation was absent (Table 5 and Fig. 7).These observations indicate that the detrital grain agesre£ect cooling in the source area because partial annealingdue to burial reheating during the sedimentary ¢lling ofthe Angastaco basin did not occur as intensely as expectedfor such a thick preserved sedimentary section.Two possi-ble explanations for this unanticipated result include: (1)the e¡ective geothermal gradient in the basinwas reducedbecause of the e¡ects of rapid sedimentation and topogra-phically driven groundwater recharge (Garven, 1989;Deming et al., 1990; Willett et al., 1997); and (2) complexstructural and depositional geometries lead to depocentremigration, preventing the base of the section from beingburied beneath the full stratigraphic thickness. Below, weexplore both of these possibilities.



Tab

le3.

Apatite¢ssion

trackdatafrom

Ang

astaco

basin

Sample

Location

Stratigraphic

unit

Cum

ulative

thickn

ess(m)

Num

bero

fgrains

(N)

Spon

taneou

sr s�

106cm�2

(Ns)

Indu

ced

r i�

106cm�2

(Ni)

Dosim

eter

r d�

106cm

-2

(Nd)

P(w2)

(%)

Central

age�

1s(M

a)Latitu

deLon

gitude

(1S)

(1W)

AI1

25141.584

66110.0670

Ang

astaco

Fm.

096

1.0114

(5456)

3.0767

(16598)

1.0843

(5261)

062.6�

2.4

AI6

25141.104

66109.1350

Ang

astaco

Fm.

601.7

110

1.1698(7165)

4.1051(25144)

1.0932

(5261)

056.7�

2.7

AI9

25141.409

66108.6860

Ang

astaco

Fm.

1066.4

910.8359

(3636)

2.8289

(12305)

1.1022(5261)

059.6�

2.5

AI11

25141.4400

66108.1780

Ang

astaco

Fm.

1620.3

100

0.7919

(5623)

2.7905

(19813)

1.1111(5261)

053.9�

2.3

AI13

25141.9600

66107.6970

Ang

astaco

Fm.

2235.9

480.5577

(1653)

2.2674

(6720)

1.1179(5261)

047.2�

3.0

AI16

25142.7800

66107.5190

Ang

astaco

Fm.

3280.3

115

0.9432

(8542)

3.1059(28128)

1.1335(5261)

059.4�

2.4

AI18

25141.2820

66106.2610

Ang

astaco

Fm.

3945.5

100

0.6655

(3413)

2.5634

(13146)

1.1447(5261)

052.4�

2.6

AI23

25140.4400

66105.5070

PaloPintadoFm.

4578.5

104

0.9746

(7015)

3.6268

(26104)

1.2604

(9321)

045.9�

3.6

AI25

25140.6030

66104.9790

PaloPintadoFm.

5077.3

970.6729

(3175)

3.708(174

96)

1.2869

(9321)

043.1�

2.4

AI26

25140.3130

66104.6540

PaloPintadoFm.

5508.5

101

0.5617

(3331)

3.5825

(21246)

1.3134

(9321)

038.4�

2.7

AI28

25140.4120

66104.3340

PaloPintadoFm.

6027.1

100

0.4183

(2697)

2.9695

(19146)

1.3399

(9321)

028.9�

2.5

AI35

25140.4320

66102.2140

SanFelip

eFm.

xxx

100

0.6567

(4691)

3.0718

(219

43)

1.3664

(9321)

048.4�

2.8

Notes:C

umulativestratigraphicthickn

essism

easuredfrom

thebo

ttom

oftheA

ngastaco

Fm.tothetopoftheP

aloPintadoFm.;N,num

bero

find

ividualgrainsd

ated

persam

ple;r s,spo

ntaneous

trackdensity

;Ns,nu

mbero

fspon

taneou

strackscoun

ted;

r i,ind

uced

trackdensity

inexternaldetector

(muscovite);N

i,nu

mbero

find

uced

tracks

coun

ted;r

d,indu

cedtrackdensity

inexternaldetector

adjacent

todosimetry

glass;N

d,nu

mbero

ftracks

coun

tedin

determ

iningr d

;P(w

2 ),w

2probability;Trackkeysoftware(D

unkl,2002)was

used

forcalculating

thecentral¢

ssiontrackages�

1s.A

patitealiquotsweremou

nted

inaralditeepoxyon

glassslides,groun

dand

polishedtoexpo

seinternalgrainsurfaces,thenetched

for20sin5.5MHNO3at211Ctorevealspon

taneous¢

ssiontracks.A

llmou

ntsw

ereprepared

usingtheexternal-detectorm

etho

d(H

urford

&Green,1983).Sam

plesand

CN-5

glassstandardsw

ereirradiated

with

thermalneutrons

inthewell-thermalized

OregonStateUniversity

reactor.After

theirradiation,thelow-U

muscovitedetectorsthatcoveredapatite

grainmou

ntsandglassdo

si-

meterwereetched

in40

%HFfor45minat211Ctorevealindu

ced¢ssion

tracks.Sam

pleswereanalysed

with

anOlympu

sBX61microscop

eat�

1250

magni¢cationwithadraw

ingtube

locatedaboveadigitizing

tabletanda

Kinetek

compu

ter-controlledstagedriven

bytheF

TStageprogramme(Dum

itru,1993).Fission

-track

ageswerecalculated

usingaw

eightedmeanzetacalib

ration

factor(H

urford

&Green,1983)basedon

IUGSagestandards

(Durango,F

ishCanyonandMou

ntDromedaryapatites)(M

illeretal.,1985,H

urford,1990).B

ased

on19

analyses,the

zvalue

forI.C

outand

is369.6�

5.1(�

1standard

error).


I. Coutandet al.

In a basin that is in a thermal steady state, the basal tem-perature is given by

T ¼ sumðziðQ=kiÞÞ þ T0 ð1Þ

where Q is the heat £ow (mWm� 2), ki is the thermal con-ductivity of a lithologic unit (Wm�1K�1), zi is the thick-ness of that unit and T0 is the mean surface temperature,taken to be10^15 1C.

Heat £ow data are not available in the study area; how-ever, an extensive study in Bolivia provides predictions fordi¡erent tectonic elements within the Central Andes(Springer & F˛rster, 1998), suggesting that the heat £ow intheEC is between 60 and40mWm� 2.Thermal conductiv-ity data are not available.Therefore, we consider a range ofplausible conductivities based on the literature values (e.g.Blackwell &Steele,1988).We subdivided theAngastaco ba-sin section into three units from bottom to top: sandstone(4200m), shale with subordinate sandstone (1500m) andsandstone (1100m). For these three units, we examinedthermal conductivities of 2.0^3.0, 1.4^2.3 and 2.0^3.0Wm�1K�1, respectively. At steady state, these para-meters implied basal temperatures between 107 and238 1C, far in excess of the maximum temperature re-quired to preserve the observed detrital thermochrono-

metric signal. This suggests that the Angastaco basin didnot reach a thermal steady state.

Transient e¡ects associated with rapid sedimentationcan depress the geothermal gradient in a sedimentary ba-sin (e.g. Deming et al., 1990). In the Angastaco basin, accu-mulation rates were 700 � 100mMa�1, which is rapid.Such transient e¡ects could allow for a slightly thicker ef-fective burial depth. Topographically driven groundwaterrecharge may also a¡ect the thermal structure of a basin,cooling proximal areas and heating more distal regions(Garven, 1989; Willett et al., 1997). This e¡ect could per-haps explain 10^20 1C of cooling; however, hydraulic datarequired to quantitatively test this hypothesis are lacking.

The stratal geometry in the Angastaco basin is poorlyconstrained by ¢eld observations because of lack of 3-Dexposure. However, the overall structural style of an east-ward-migrating deformation front observed in this areafavours syntectonic growth strata. In particular, in the ab-sence of an eastward migration of the depocentre, both thesubsidence and the exhumation rate of the basin would beextremely high. Therefore, it is plausible that depocentremigration prevented the sampled base of the section fromhaving been buried beneath the 6-km-thick section.

Although we cannot de¢nitively prove which of theseexplanations is more important, both arguments are con-

Table4. Best- ¢t age populations for apatite FTGA samples from the Angastaco basin

Sample

Strati-graphicage (Ma) N

FTagerange(Ma) P1 P2 P3 P4 P5 P6

AI 35 [3.4^1.6] 100 3.9^118.7 xxx 12.7 � 2.6/2.1 xxx xxx 50 � 3.3/3.1 82.7 � 7.0/6.420.1% 56.4% 23.5%

AI 28 5.6 100 2.6^110.9 6.6 � 1.4/1.2 xxx 22.5 � 2.9/2.6 xxx 60.9 � 3.8/3.6 xxx39.1% 20.3% 40.6%

AI 26 6.5 101 4.0^139.7 xxx 14.3 � 1.5/1.3 xxx xxx 53.2 � 3.5/3.3 103.9 � 20.6/17.241.3% 51.9% 6.9%

AI25 7.3 97 3.8^100.5 xxx xxx 20.6 � 3.2/2.8 45.6 � 20.5/14.1 75.7 � 20.0/15.9 xxx32.8% 44.0% 23.2%

AI23 8.3 104 3.9^154.3 9.9 � 1.5/1.3 xxx xxx xxx 70.8 � 6.1/5.7 123.2 � 56.1/38.640.5% 52.6% 6.9%

AI18 9.3 100 5.9^166.6 xxx 17.7 � 5.9/4.4 xxx 39.8 � 4.0/3.7 67.3 � 4.7/4.4 158.2 � 48.4/37.27.5% 41.2% 48.9% 2.4%

AI16 10.1 115 3.0^156.5 10.2 � 12.8/5.7 xxx 26.8 � 13.1/8.8 56 � 4.6/4.3 74 � 5.4/5.1 119.7 � 35.1/27.25.8% 6.3% 45.1% 39.8% 3%

AI13 11.4 48 14.3^102.3 xxx xxx 31.3 � 5.3/4.5 51.8 � 7.1/6.3 84.7 � 12.6/11.0 xxx39.9% 46.5% 13.6%

AI11 12.2 100 13.8^133.8 15.6 � 8.0/5.3 xxx 36.6 � 4.2/3.8 60.9 � 3.9/3.7 82.2 � 8.6/7.8 128.5 � 91.3/53.63.6% 33.0% 48.9% 13.5% 0.9%

AI9 12.9 91 21.0^158.8 xxx xxx 30.2 � 6.9/5.6 49.1 � 5.4/4.9 75.5 � 5.6/5.2 xxx10.4% 39.5% 50.2%

AI6 13.5 110 11.4^129.1 14.3 � 2.0/1.8 xxx 36.7 � 4.5/4.0 65.3 � 4.6/4.3 88.4 � 12.2/10.7 xxx12.4% 16.7% 54.0% 16.9%

AI1 14.3 96 15.4^147.0 xxx xxx 40.2 � 4.3/3.9 72.1 � 3.9/3.7 xxx 130 � 34.2/27.130% 68.7% 1.2%

Notes: Binomial peak- ¢t ages (P1^P6) were determinedwithBINOMFIT (Brandon,1996; 2002) and are indicated by their mean age � 2SE (equivalent to95% con¢dence interval; note that these intervals are asymmetric), followed by their relative abundance (%);N, total of grains counted.FTGA, ¢ssion-track grain age.



Fig. 6. Detrital apatite ¢ssion-track results.Radial plots of the single-grain age data and probability density plots of ¢ssion-track grain-age distributions (Brandon, 1996). Dashed lines indicate probability density distributions and solid lines indicate componentdistributions estimated by the binomial peak- ¢t method (Galbraith &Green, 1990).


I. Coutandet al.

sistent with the basic observation that the AFT data werenot signi¢cantly reset by burial during ¢lling of the basin.Hence, our data most likely re£ect the thermal history ofthe source terrains of the clastic material.

DETRITALTHERMOCHRONOLOGYANDHINTERLAND EVOLUTION

The distribution of detrital mineral cooling ages in sedi-ments can provide a record of the erosional exhumationof the hinterland source area. However, converting a cool-ing age extracted from a detrital sample into an exhuma-tion rate for the source area requires several majorassumptions that we brie£y discuss below (for a review,see Garver et al., 1999; Brewer et al., 2003; Ruiz et al., 2004).

The procedure used to derive the average exhumationrate of a hinterland source from a population of thermo-chronologic data collected through a reasonably well-dated sedimentary sequence is discussed by Garver et al.(1999). In summary, the average exhumation rate is givenby

dz=dt ¼ ððTc � TsÞ=ðdT=dzÞÞ=Dt ð2Þ

which is equivalent to

dz=dt ¼ Zc=Dt ð3Þ

whereTc is the closure temperature of the thermochron-ometer,Ts is the surface temperature, dT/dz is the geother-mal gradient,Dt is the lag time andZc is the closure depth.The lag time is de¢ned as the time elapsed between thetime of closure of a crystal for a given geochronologicalsystem in the source region and the time at which it is de-posited in the sedimentary basin. The assumption isusually made that the time taken by newly eroded detritusto be transported and then deposited in the sedimentarybasin is negligible compared with the time taken to reachthe surface (and be eroded) from the closure depth (Cer-veny et al., 1988; Garver et al., 1999; Ruiz et al., 2004).Thus,to be valid, the lag time concept implies that during theirtransit from the source to the basin, the detritus are notstored in transient depocentres during time periods longenough (ca. �1Ma) to bias their stratigraphic age in thebasin signi¢cantly. In addition, subsequent to deposition,recycling of non-reset sediments does not signi¢cantlymodify the initial stratigraphic age (Ruiz et al., 2004). Per-haps the most signi¢cant condition for the Angastaco ba-sin is that the detritus must have cooled from atemperature higher than the apatite ¢ssion-trackTc. If this

Table5. Mean-track length and related etch pit (Dpar) values obtained for di¡erent detrital age components

Sample Component

Peak age( � 95%CI)(Ma)

Numberof tracks

Mean tracklength(mm � 1s)

Mean tracklength STD(mm)

Dpar

(mm)Dpar STD(mm)

AI1 P4 72.1 � 3.9/3.7 78 12.3 � 0.17 1.46 1.71 0.12AI6 P3 36.7 � 4.5/4.0 13 12.2 � 0.32 1.16 1.74 0.07

P4 65.3 � 4.3/4.6 61 12.44 � 0.15 1.25 1.70 0.09P5 88.4 � 12.2/10.7 30 11.73 � 0.31 1.69 1.75 0.17

AI9 P5 75.5 � 5.6/5.2 33 11.72 � 0.27 1.55 1.71 0.12AI11 P4 60.9 � 3.9/3.7 22 11.55 � 0.31 1.45 1.69 0.15

P5 82.2 � 7.8/8.6 21 11.76 � 0.40 1.83 1.71 0.10AI16 P4 56.0 � 4.6/4.3 28 12.53 � 0.19 0.99 1.67 0.08

P5 74.0 � 5.4/5.1 75 11.83 � 0.18 1.60 1.64 0.09AI18 P4 39.8 � 4.0/3.7 20 12.88 � 0.28 1.26 1.94 0.42

P5 67.3 � 4.7/4.4 22 12.07 � 0.32 1.49 1.78 0.19AI23 P5 70.8 � 6.1/5.7 57 12.0 � 0.16 1.20 1.75 0.19AI25 P5 75.7 � 20.0/15.9 19 12.07 � 0.19 0.81 1.85 0.10AI26 P5 53.2 � 3.5/3.3 22 11.71 � 0.40 1.88 1.72 0.15

P6 103.9 � 20.6/17.2 17 12.12 � 0.41 1.68 1.92 0.14AI28 P5 60.9 � 3.8/3.6 43 12.93 � 0.26 1.68 1.75 0.16AI35 P5 50.0 � 3.3/3.1 24 12.62 � 0.36 1.76 1.79 0.19

P6 82.7 � 7.0/6.4 33 11.77 � 0.32 1.82 1.79 0.12

Dep

ositi

onal

age

(M

a)

0

2

4

6

8

10

12

14

16

0 2 4 6 8 10 12 14

Mean track-length (µm)

Fig.7. Mean track-length values plotted against depositional age(see alsoTable 5).



were not the case, detrital thermochronologic data col-lected could not be converted into exhumation rates forthe corresponding source regions.

For hinterland regions, our approach requires other as-sumptions. It is necessary to assume a steady geothermalgradient at the time of closure of the geochronological sys-tem, although the geothermal gradient at that time is nottrivial to derive (Garver etal.,1999). Isotherms are assumedto remain horizontal and particle paths toward the surfacevertical. The exhumation rate remains constant from thetime of closure to the time of deposition of the crystal.Hence, this approach neglects isotherm warping in thesource region caused by topography and advection (e.g.Stˇwe et al., 1994;Mancktelow &Grasemann, 1997; Braun,2002) as well as non-vertical displacements of particlesduring exhumation (Batt & Brandon, 2002). As long asthe spatial distribution of cooling ages within a sourcecatchment during a time increment remains di⁄cult todetermine (Brewer et al., 2003), quantifying exhumationrates from detrital thermochronologic data is fraughtwithuncertainty. Finally, the simplest model supposes that theexhuming topography is eroded principally from higherelevations and strictly in a vertical direction. In reality, thedrainage basin may cover a large range of elevations in thesource topography; thus, di¡erent peak ages identi¢ed inthe detrital sediments will not necessarily be derived fromdistinct source terrains but rather could be derived fromdi¡erent elevations within a single catchment (Garver etal., 1999). These restrictions imply that interpretation ofdetrital peak age provenance aswell as long-term exhuma-tion rates derived from them should always be treatedwithcaution.Nevertheless, detrital thermochronology remainsa valuable tool for constraining the cooling and exhuma-tion history of sources terrains partially or totally eroded.

INTERPRETATION OF THE ANGASTACOBASIN DETRITAL SIGNAL

Plotting the ¢ssion-track ages of a given componentvs. thestratigraphic ages of the host strata yields a graphical as-sessment of the lag time variability through the strati-graphic column and thus, through time (e.g. Ruiz et al.,2004), thereby providing information on the long-term ex-humation of the source region for that particular detritalage population (Fig. 8).The source-area signal of the prox-imal Angastaco foreland basin changes through time, forinstance because of sporadic drainage re-organization,and unlike Ruiz etal. (2004), we do not expect every sampleto contain the same number ofPn age components.Conse-quently, ourDn paths or trends are not de¢ned by automa-tically connecting individual components with the samen of each sample. Instead, we sorted the Pn componentssuch as they de¢ned geologically realistic processessmoothly evolving through time following aDn trend.

Here, we describe this population subdivision and sub-sequently interpret each path. In this dataset, lag time foreach path either remains constant or decreases upsection,

which is consistent with exhumation. Increasing lag timescould represent recycling of sediments. Althoughwe can-not exclude such a process, if it exists, the resulting signalappears to be hidden behind the dominant decreasing lagtime trend.

P1component

The ¢rst population (P1) had extremely short lag times(Fig. 8 and Table 4); ¢ve samples had cooling ages between15.6 and 6.6Ma and lag times between 0 and 4Ma.

A constant lag time approaching 0 re£ects either persis-tent and extremely rapid cooling and exhumation in thehinterland or volcanic contamination of the clasticmateri-al shed into the basin (Ruiz et al., 2004). The publishedapatite ¢ssion-track data set in the Andean hinterland atthese latitudes (e.g. Andriessen & Reutter, 1994; Coutandet al., 2001; Deeken et al., 2004; Carrapa et al., 2005) doesnot reveal any ages young enough to produce the detritalages in population P1, suggesting that rapid exhumationwas not the cause of this young age component. Alterna-tively, the hypothesis of volcanic contamination seemsplausible.The region contains numerous airfall tu¡s withsimilar ages (Strecker et al., 1989; Grier & Dallmeyer, 1990;Muruaga, 2001a, b) (Fig. 3). The P1 population is presentfrom bottom to top of the stratigraphic section but spora-dically, as would be expected from episodic input of volca-nic ash into the basin.

P2 component

The next population, P2, contains only three samples withages between 17.7 and 12.7Ma, and similar lag times of 8^10Myr. It ¢rst appears at 9.3Ma; however, it is possiblethat the next older sample (�10Ma) contains some grainswith similar ages that have not been statistically distin-guished from the P1population.

The relatively constant lag time of 8^10Ma is character-istic for a constant, high erosion rate (Bernet et al., 2001;Ruiz et al., 2004) and likely representative of material thathas rapidly cooled through the entire apatite PAZ. For alagtime Dt5 8^10Myr, Tc5110� 5 1C, Ts515 1C, dT/dz5 25� 5 1C km�1, Eqn (2) yields a range of exhuma-tion rates of 0.3^0.6mmyr�1. This signal ¢rst appears at9^10Ma, suggesting that at this time either the ¢rst apa-tites exhumed from belowTc were exposed at the surfaceor that drainage reorganization, which ¢rst supplied thisdetritus to the basin, had occurred. One method to di¡er-entiate between these possibilities is to ascertain whetherapatites representing the exhumed PAZ are preserved inthe basin; this is discussed below (see P4). Alternatively,we can look for appropriate source areas in the vicinity ofthe basin. Palaeocurrent directions throughout the sectionsuggest sediment transport from the west (e.g. Starck &Anzo¤ tegui, 2001). The Cumbres de Luracatao, belongingto the EC (Figs 1a and 2a), is the most prominent range tothe west of the Angastaco basin and contains granite andgranodiorite of the Oire intrusive complex. Three AFTvertical pro¢les from the northern portion of this range


I. Coutandet al.

reveal reset AFTages between 16 and 20Ma over a com-bined elevation range of 2100m (Deeken et al., 2004). Asingle partially reset sample at high elevation de¢nes akink in the age^elevation curve and suggests that relativelyrapid exhumation commenced by 20Ma at ca. 0.4mmyr�1. At this exhumation rate, the ¢rst reset apatiteswouldhave been exposed at the surface at ca. 10Ma.The similarage range and exhumation rate observed at Cumbres deLuracatao and in the D2 trend suggests a genetic link.However, the youngest ages observed in the D2 path havenot yet been retrieved from bedrock samples.

P3 component

The third component (P3) had cooling ages between 40.2and 20.6Ma; D3 path is a ‘forward-moving peak’ with alag time that regularly decreases upsection from�25Myr at the bottom to �15Myr at the top (Fig. 8).Thiskind of pattern is generally interpreted as re£ecting anoverall increase in cooling rate in the source regionthrough time and is classically interpreted as being asso-ciated with a phase of orogenic growth (Garver et al.,1999). However, this statement is valid only if the ages re-corded by the apatite crystals re£ect cooling through theclosure temperature. For young, non-volcanic apatites,this hypothesis is likely correct; for older apatiteswithverylong lag times, and hence very slow cooling rates, this isprobably incorrect. In the case that the samples were fullyreset during the ultimate exhumation event, an exhuma-tion rate can be calculated using the same approach aswas applied forP2, resulting in 0.1^0.2mmyr�1at the baseand increasing to 0.2^0.3mmyr�1 at 5Ma. The lowestsample in theP3 population re£ects cooling during the in-terval from 40Ma (passing throughTc) to 14.3Ma (deposi-tion in the basin); the youngest sample in P3 re£ectscooling between 21 and 7.3Ma. Potential source areas withsimilar ages (e.g. 38� 3 to 29� 3Ma) have been dated inthe Chango Real Range, located ca. 100 km to the SSW

(Fig. 1a and b) (Coutand et al., 2001) and a single age of30.3 � 3.0 has been reported from Santa Rosa de Tastil(Quebrada del Toro) located ca. 100 km to the NNE (Fig.1a and b) (Andriessen & Reutter, 1994). Both localities arepart of the present-day EC.Although neither of these agesprovides an exact match to theP3 component, they suggestthat appropriate sources existed on the eastern £ank of theAndes. Cooling ages of 29 � 1 to 24 � 3Ma have been ob-tained from Sierra de Calalaste, in the southern Puna pla-teau to the west (Carrapa et al., 2005). A large suite of AFTdatawith ages between62 and30Ma characterizes theCor-dillera Domeyko in the Western Cordillera (Chile) (Mak-saev & Zentilli, 1999). However, the scarcity of volcanicmaterial within the Angastaco basin, which would be ex-pected from £uvial systems draining the Puna, strongly ar-gues against source areas on the far western side of theplateau. Although it has not yet been precisely identi¢edin the modern landscape, the source for theP3 componentmust be locatedwithin the present-day EC

P4 component

The fourth population,P4, contains a much broader rangeof cooling ages and a less well-de¢ned trend on the lagtime plot. Cooling ages range between 65.3 and 39.8Ma;this population disappears from the section above ca.7Ma. D4 is also a ‘forward-moving peak’; however, the lagtime decreases upsection in a complicated pattern, from�57Myr at the bottom to �30Myr at 9Ma (Fig. 8).Thestratigraphically youngest FTGA peak has a longer lagtime of 38Myr and an age of 45 � 21/14Ma (sample AI25,seeTable 4 and Fig. 8).The next older FTGA peak in thissample also has a very large error bar; the errors of thesetwo age clusters overlap. Therefore, we suspect that thepeak ages for this sample are poorly resolved and we donot consider them to be highly signi¢cant. More impor-tantly, no FTGA peaks for P4 are identi¢ed higher in thestratigraphic section (Fig. 8).This suggests two possibili-

Lagtime < 0 My

D1 D2 D3 D4 D5

0 10 20 30 40 50 60 70 80 90 100Detrital apatite fission track age (Ma)

0

5

10

15

Str

atig

raph

ic a

ge (

Ma)

0 My

5 My

20 My

30 My

40 My

50 My

60 My

70 My

80 My

90 My

10 My

Fig. 8. Evolution of the detrital paths (D1^D5) through the stratigraphic record. Error bars are � 2 SE for apatite FTages and � 1Mauncertainty for depositional ages. Dashed grey lines indicate iso-lag time lines. Symbols indicate binomial peak- ¢t ages extracted from themixedgrain age distributions (P1^P6, seeTable4 andFig.6) andgrey shaded boxes underline relateddetrital path tendency upsection (D1^D5).



ties: (1) the component representing this source area wasno longer deposited in the basin after ca. 9Ma, or (2) theD4 path became signi¢cantly younger at this time.The for-mer idea cannot be ruled out: similar ages are present inthe Sierra Famatina ca. 350 km to the south (Coughlinet al., 1998). However, 9Ma corresponds closely to the ¢rstappearance of the P2 component.Therefore, it is possiblethat the D4 path represents the exhumed PAZ that for-merly overlay the P2 source area.We link the latter sourcearea with rapid exhumation of Cumbres de Luracatao,commencing at ca. 20Ma. A single sample from the ex-humed PAZ that formerly overlay the range yielded an ageof 56� 2Ma (Deeken et al., 2004), ¢tting well with the agerange recorded by P4.The large scatter of ages within theD4 path is consistent with the exhumation of a PAZ, be-cause a variety of ages would be exposed within a narrowrange of elevation.The pronounced decrease in age upsec-tion in D4 is also more typical of exhumation through aPAZ.Therefore, we conclude that the proximal Cumbresde Luracatao is a likely source for the P4 component.

P5 component

The ¢fth component (P5) has cooling ages between 88.4and 50.0Ma. P4 and P5 together contribute between 50and 70% of the dated crystals (Table 4). D5 is a ‘forward-moving peak’ with ages and lag times decreasing upwardfrom 88 to 50Ma and 75 to 47Myr, respectively. Indeed,the rigorous de¢nition of a lag time applies to neitherthese samples nor the P4 component; the old ages and ex-tremely long lag times suggest that this material was de-rived from an exhumed PAZ that was slowly eroded ratherthan cooled through a closure temperatureTc. A verticalAFT pro¢le through Cerro Durazno, immediately west(Fig. 2a), yielded cooling ages between 95 and 52Ma (Dee-ken et al., 2004; A. Deeken, E.R. Sobel, I. Coutand, M.Haschke, U. Riller & M.R. Strecker, submitted). Track-length modelling of these samples indicates exhumationcommencing between12 and 7Ma (A.Deeken, E.R. Sobel,I. Coutand, M. Haschke, U. Riller & M.R. Strecker, sub-mitted). Similar ages between 93 � 5 and 42 � 2Ma werereported from Santa Rosa de Tastil at 241S (Coutand,1999). The similarity likely re£ects a common tectono-thermal history; however, the detritus in the basin wasmost likely derived from the adjacent Durazno range.Thesimilar age range of the P5 component and the adjacentCerroDurazno suggests that they represent the same (late)Cretaceous PAZ.

P6 component

The oldest population (P6) ranges between 158.2 and82.7Ma, with lag times between 150 and 80Myr. Most ofthese FTGA peaks represent a minor fraction of the datedcrystals (1^7%) and hence have quite large errors; how-ever, the youngest sample AI35 contains 23% of the detri-tal apatite (Table 4).There is a crude decrease in age and lagtime upsection (not represented in Fig. 8); however, withthe exception of the stratigraphically youngest sample,

this pattern is poorly resolved.There are several potentialsources for this minor component. One possibility is thePuna plateau, which includes old AFT source regions(Deeken et al., 2004; A. Deeken, E.R. Sobel, I. Coutand,M. Haschke, U. Riller & M.R. Strecker, submitted), butour sedimentary petrography data argue against this.Alternatively, the presently outcropping Cretaceous andPalaeogene strata could have provided a source for re-cycled detrital apatite with old ages.

DISCUSSION

Sandstone petrography and clast counts in conglomeratesidentify the EC as the main source of the detritus depos-ited in the intramontane Angastaco basin between 14.5and 5.5Ma. A cooling event starting in the late Eocene^Oligocene is clearly visible in the D3 detrital path. Thissource was exhumed until the Pliocene at a rate increasingfrom 0.1 to 0.3mmyr�1. Although the location of thesource area has not yet been identi¢ed, the existence ofmodern outcrops that yielded late Eocene to Oligocenerapid cooling AFTages (Andriessen & Reutter, 1994; Cou-tand et al., 2001) supports a provenance from the EC.Wepropose that this cooling event re£ects an early compres-sional episode in the eastern part of the Andean orogen.This event created enough topography to produce detritus(Fig. 9a). Detrital path D2 highlights faster exhumation oftheEC that initiated 20^18Myr ago and lasted at least untilthe Pliocene at a nearly constant rate between 0.3 and0.6mmyr�1. Palaeocurrent indicators (Starck & Anzo¤ te-gui, 2001), sandstone petrography and apatite ¢ssion-track data (Deeken et al., 2004; A. Deeken, E.R. Sobel, I.Coutand, M. Haschke, U. Riller & M.R. Strecker, sub-mitted) suggest that the adjacent Cumbres de Luracatao(Fig. 2a) was being exhumed and eroded at that time. Up-lift was accommodated by a north^south striking, west-dipping reverse fault on the eastern side of the range (Fig.2a). The evolution of the compressional Angastaco basinbeganwhen su⁄cient accommodation space had been cre-ated in the footwall of the Luracatao thrust to accumulateand preserve erosional products from this range and othersources. If our interpretation of the detrital pathD3 is cor-rect, and that a topography-creating exhumation event inthe EC commenced in the late Eocene^early Oligocene,then the former foreland basin must have been segmentedas itwas farther west in the Puna in the Sierra deCalalaste-Salar de Antofalla area (Carrapa etal., 2005). Subsidence inthe Angastaco basin accelerated at about 15Ma (Fig. 9b).Our sandstone petrographydata con¢rm that the EC musthave constituted an e⁄cient orographic barrier, because itwas able to prevent transport of detritus derived fromwidely outcropping volcanics of the Puna plateau into theAngastaco basin (Figs 1a, 5 and 9b). During sediment de-position between �15 and 9 � 1Ma, the climatic condi-tions are interpreted to have been arid (Starck &Anzo¤ tegui, 2001) as would be expected at this latitude (R.Alonso, B. Carrapa, I. Coutand, M. Haschke, G.E. Hilley,


I. Coutandet al.

W E

Condensation of Atlantic-derivedmoisture along the windward flankof the eastern Cordillera between ~ 9 −3.5 Ma

Condensation of Atlantic-derivedmoisture along the windward flankof the frontal uplifts from ~ 3.5 Ma

Stage 1 : 40-20 Ma

Stage 2 : 20-13 Ma

Stage 3 : 13-3.5 Ma

Stage 4 : 3.5 Ma to present

Eastern Cordillera

Eastern Cordillera

Runno, Durazno & Negro ranges

Pucara-Angastaco basin

Angastaco basin

Proto Eastern Cordillera(?)

Arid Wet

Arid WetSemi-arid

(a)

(b)

(c)

(d)

Fig.9. Cartoons illustrating the tectonic, topographic and climatic evolution of the Eastern Cordillera and Calchaqu|¤ valley at thelatitude of the Angastaco basin, during the last 40Myr.



E.R.Sobel,M.R. Strecker&M.H.Trauth, submitted).Thisassessment is also compatible with oxygen-isotope datafrom the Corral Quemado basin located further south, inthe Sierras Pampeanas province (Latorre et al., 1997).

During or right after the deposition of the AngastacoFormation, regional shortening shifted eastward, drivingthe exhumation of the Runno,Durazno andNegro Ranges(Figs 2a and 9c); this led to compartmentalization in thePucara and Luracatao valleys to the west, and the modernAngastaco basin to the east.The late Cretaceous AFTagesobtained by Deeken et al. (2004) and A. Deeken, E.R.Sobel, I. Coutand, M. Haschke, U. Riller &M.R. Strecker(submitted) across these ranges indicate that the magni-tude of exhumation is less than in the Luracatao Range.The increasing contribution of the D5 detrital populationto over 50% of the apatite fraction from sample AI16 (topof the Angastaco Formation, seeTable 4 and Fig. 3) and anincreased contribution of low-grade metamorphic detri-tus constituting these ranges (Fig. 5b) support the onset ofexhumation in the Runno, Durazno and Negro Rangesbetween13 and10Ma.

Depositional environments changed drastically betweenthe deposition of the Angastaco and Palo Pintado Forma-tions after about 9Ma (Fig. 3) (Marshall et al., 1983; Starck& Anzo¤ tegui, 2001).The Palo Pintado and basal San FelipeFormations contain a rich fossil mammal fauna, freshwatermollusc fossils, pollen, leaves and tree-trunk remnants (e.g.Anzo¤ tegui, 1998; Starck & Anzo¤ tegui, 2001), from which awarm, humid forested environment, similar to the presentforeland (e.g. Vergani & Starck, 1989; Starck & Vergani,1996), can be inferred for the late Miocene to early Plio-cene. In contrast, intramontane basins within the presentPuna record sustained aridity and evaporite deposition atthat time (Alpers&Brimhall,1988;Alonso etal.,1991;Alon-so,1992;Vandervoort etal.,1995).These observations docu-ment that between about 9 and 5Ma, the EC had reached athreshold elevation that forced precipitation from easterlysources, independent of global changes inferred from pa-laeo-temperatures (Starck & Anzo¤ tegui, 2001; R. Alonso,B. Carrapa, I. Coutand, M. Haschke, G.E. Hilley, E.R. So-bel, M.R. Strecker &M.H.Trauth, submitted) (Fig. 9c). Si-milar results were obtained from stable isotope analysis ofpalaeosols in the adjacent Santa Maria valley (Kleinert &Strecker, 2001) (Fig.1).

In the late Pliocene, during deposition of the upper partof the San Felipe Formation, major changes in provenanceand climatic conditions recurred. Palaeocurrent and pro-venance data indicating westward transport (Starck & An-zo¤ tegui, 2001) of sediments derived from the Sierra de losColorados to the east (Fig. 2a and b) mark the onset of ac-tivity along a N^S striking, east-dipping thrust faultbounding the western £ank of the Sierra.The uplift of thisrange (present-day elevation of 3500m) coincided with achangeover to arid conditions in the Angastaco basin andthe formation of calcretes sometime between 5.4 and2.4Ma (Fig. 9d), a scenario identical to the situation re-corded in the Santa Maria basin to the south (Fig. 1)(Strecker etal.,1989;Kleinert &Strecker, 2001).This envir-

onmental change was accompanied by the deformationand exhumation of the entire Angastaco sedimentary se-quence, the establishment of transient internal drainageconditions and the deposition of massive conglomerateswith inter¢ngered lacustrine sediments. In agreementwith Starck & Anzo¤ tegui (2001), we propose that the latePliocene uplift of the Sierras de los Colorados to the eastcreated an orographic barrier that subsequently led to thearidi¢cation of the Calchaqu|¤ valley, by forcing Atlantic-derived moisture to condense along the windward £ank ofthis then frontal foreland uplift (Fig.9d).

CONCLUSIONS

Sandstone petrography, structural and stratigraphic analy-sis combinedwith detrital thermochronology in the intra-montaneAngastaco basin provide valuable information onthe late Eocene to late Pliocene exhumation history of theeastern border of the Puna plateau and the successive es-tablishment of orographic barriers along the southeastern£ank of the Central Andes. Long-term exhumation ratesbased on detrital thermochronology revealed an increasefrom 0.1 to 0.3mmyr�1 from late Eocene^Oligocenethrough Pliocene. However, the Luracatao Range in theEC was exhumed at a nearly constant rate between 0.3and 0.6mmyr�1 from the early Miocene (20^18Ma) untilthe Pliocene.This created su⁄cient crustal thickening toinitiate the Pucara-Angastaco compressional basin atabout15Ma,when clastic material eroded from theEC be-gan to be preserved. Apparently, by this time, the EC hadundergone su⁄cient surface uplift to disrupt the £uvialnetwork, disconnecting the interior of the Puna plateaufrom the foreland. By ca. 13^10Ma, compressional defor-mation had shifted eastward, dividing the early basin intotwo subordinate depocentres: the Pucara valley to the westand the Angastaco basin to the east.

From about 15 to 9Ma, the Angastaco Formation accu-mulated in theAngastaco basin under arid conditions.Theoverlying Palo Pintado Formation marks a drastic environ-mental change from a dry to a humid climate, similar toconditions in the Chaco lowlands and the Subandeanranges (Vergani & Starck, 1989; Starck & Vergani, 1996). Atthe same time, arid conditions characterized the Puna pla-teau to the west. This suggests that between about 9 and5Ma, the EC had built su⁄cient topography to focusAtlantic-derived precipitation and block moisture trans-port toward the Puna plateau.

In the late Pliocene (3.4^2.4Ma), shortening was trans-ferred to the eastern £ank of the Angastaco basin, buildinganother orographic barrier further east and establishingsemi-arid conditions across the Calchaqu|¤ Valley. Alongthe southeastern £ank of the Central Andes, the migrationof compressional deformation and coeval building of oro-graphic barriers in an environmentwith an easterly moist-ure source thus focused precipitation onto the windward(eastern) £anks of uplifts and starved the western hinter-land of moisture.


I. Coutandet al.

The tectonic and climatic conditions in the southernCentral Andes have led to an interesting relationship be-tween tectonics and climate, where the eastward migrationof deformation has been accompanied by successive aridi-¢cation in the same direction. This evolution caused thetectonic defeat of £uvial systems by tectonically driven sur-face uplift, resulting in ine⁄cient mass removal from theinterior of the orogen (Sobel et al., 2003), and helped topreserve and aerially expand the internal drainage condi-tions of the Puna plateau.

ACKNOWLEDGEMENTS

This project is part of Collaborative Research Center(SFB) 267, ‘Deformationsprozesse in den Anden’, fundedby the German Research Council (DFG) to M. Strecker.I. Coutand and B. Carrapa thank the Alexander vonHum-boldt Foundation for providing funding during their stayinGermany.The ion microprobe facility atUCLA is partlysupported by a grant from the Instrumentation and Facil-ities Programme, Division of Earth Sciences, NationalScience Foundation. Simone Gaab is thanked for assis-tance in the ¢eld, S. Bonnet for providing Fig.1b and G.E.Hilley forFig.1c.Constructive reviews byF.M.Da¤ vila, P.G.DeCelles and P.A. van der Beek helped to improve themanuscript.

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Manuscript received 26 May 2005; Manuscript accepted19 January 2006.


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