Tectonic vs. climate forcing in the Cenozoic sedimentary evolution of a foreland basin (Eastern...

Tectonic vs. climate forcing in the Cenozoicsedimentary evolution of a foreland basin (EasternSouthalpine system, Italy)N. Mancin, A. Di Giulio and M. Cobianchi

Dipartimento di Scienze dellaTerra, Universita' di Pavia, Pavia, Italy

ABSTRACT

This paper discusses the Cenozoic interaction of regional tectonics and climate changes.Theseprocesses were responsible for mass £ux from mountain belts to depositional basins in the easternAlpine retro-foreland basin (Venetian^Friulian Basin). Our discussion is based on the depositionalarchitecture and basin-scale depositional rate curves obtained from the decompacted thicknesses ofstratigraphic units.We compare these data with the timing of tectonic deformation in thesurrounding mountain ranges and the chronology of both long-term trends and short-term high-magnitude (‘aberrant’) episodes of climate change. Our results con¢rm that climate forcing (andespecially aberrant episodes) impacted the depositional evolution of the basin, but that tectonics wasthe main factor driving sediment £ux in the basin up to the LateMiocene.The depositional rateremained below 0.1mmyear�1 on average from the Eocene to theMiocene, peaking at around0.36mmyear�1, during periods of maximum tectonic activity in the eastern Southern Alps.Thisdynamic strongly changed during the Pliocene^Pleistocene, when the basin-scale depositional rateincreased to an average of 0.26mmyear�1 (Pliocene) and 0.73mmyear�1 (Pleistocene).This result ¢tsnicely with the long-term global cooling trend recorded during this time interval. Nevertheless, wenote that the timing of the observed increase may be connectedwith the presumed onset of majorglaciations in the southern £ank of the Alps (0.7^0.9Ma), the acceleration of the global cooling trend(since 3^4Ma) and climate variability (in terms of magnitude and frequency). All these factors suggestthat combined high-frequency and high-magnitude cooling^warming cycles are particularlypowerful in promoting erosion in mid-latitude mountain belts and therefore in increasing thesediment £ux in foreland basins.

INTRODUCTION

The question of the relative roles played by tectonics andclimate changes in guiding erosion, building mountainchains and producing sediments that accumulate in fore-land basins remains a matter of debate (Cederbom et al.,2004; Molnar, 2004; Allen, 2008;Vernon et al., 2008).Tec-tonics, climate and erosion interact as tectonics, via hori-zontal shortening, increases the thickness of thecontinental crust, thereby elevating the Earth’s surfaceand erosion removes rock masses, inducing an isostatic re-sponse (rebound) that elevates peaks above the surround-ing surface and produces new relief (Champagnac et al.,2007 and references therein). The two types of processes,however, operate over a very di¡erent timescales (Allen,2008 and references therein): changes in tectonic condi-tions are very slow, on the order of several millions of years(Willett, 1999); whereas climate, in contrast, is ¢ckle, withperiodicities on the order of 105 years (Zachos et al., 2001).

As a result, it is di⁄cult to distinguish global climate ef-fects in the sedimentary records of tectonically active ba-sins. Consequently, it is challenging to determine howand by how much these signals may have been imprinted(Molnar, 2004). Researchers have also debated what arethe most important climate-related factors that in£uenceerosion^deposition dynamics in source-basin systems. Inparticular, we do not knowwhether long-term global tem-perature trends or short-term cooling^warming cycles arethe main factors. Similarly, it remains unclear whether gla-ciation thresholds must be overcome in order to exert pri-mary control on sediment routing systems (e.g. Peizhen etal., 2001;Molnar, 2004).

TheVenetian^ Friulian foreland basin (VFB) provides agood natural laboratory to study if and how climate signalsare recorded in a tectonically controlled basin. In such asystem, the entireCenozoic history of the basin iswell pre-served in the subsurface (mostly) and locally, also in surfaceexposure (Mancin et al., 2007) and the tectonic history ofboth the basin and surrounding belts is fairly well known(Fantoni et al., 2002 and references therein).

This paper aims to explore the relationships betweenregional tectonics and climatic forcing in the depositional

EAGE

Correspondence: Nicoletta Mancin, Dipartimento di Scienzedella Terra, Universita' di Pavia, via Ferrata 1, 27100 Pavia, Italy.E-mail: [email protected]

BasinResearch (2009) 21, 799–823, doi: 10.1111/j.1365-2117.2009.00402.x

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https://www.researchgate.net/publication/12010365_Trends_Rhythms_and_Aberrations_in_Global_Climate_65_Ma_to_Present?el=1_x_8&enrichId=rgreq-8ae2cda7-7b59-43b9-931d-940688474738&enrichSource=Y292ZXJQYWdlOzIyNzgwMjg0MTtBUzoxMDEwOTY5NTM0ODMyNjRAMTQwMTExNDgwMDI3OA==

https://www.researchgate.net/publication/215614600_Willett_S_D_Orogeny_and_orography_the_effects_of_erosion_on_the_structure_of_mountain_belts_J_Geophys_Res_104_28957-28981?el=1_x_8&enrichId=rgreq-8ae2cda7-7b59-43b9-931d-940688474738&enrichSource=Y292ZXJQYWdlOzIyNzgwMjg0MTtBUzoxMDEwOTY5NTM0ODMyNjRAMTQwMTExNDgwMDI3OA==

evolution of theVFB during the last 60Myr. It documentsthat in a tectonically controlled sedimentary basin, climateforcing can imprint an identi¢able signal, and that climateinstability and ‘aberrant’ short periods (as de¢ned by Za-chos et al., 2001) seem to in£uence depositional patternsmore than do long-term climate trends. Moreover, theavailability of a quite continuous Pliocene^Pleistocene de-positional record gives us the opportunity to discuss thefactors controlling sediment £ux. In this respect, we dis-cuss in particular whether simple cooling^warming cycleshave been the main drivers for erosion^deposition bud-gets, or if glaciation^deglaciation cycles are required to al-low climate cycles to exert a primary forcing on clasticproduction^accumulation systems.

CENOZOIC CLIMATE AND SEDIMENTSUPPLY

The Cenozoic is an Era characterised by signi¢cantchanges in global climate with generalised high-latitudecooling that led to aNeogene increase in ice-sheet volumeand a resulting long-term transition from a non-glaciated‘greenhouse world’ to a polar-glaciated ‘icehouse world’(Miller & Kent, 1987; Zachos et al., 1996, 2001; Liu et al.,2004).The mechanisms that caused this global cooling re-main ambiguous (Raymo et al., 1988; Raymo & Ruddiman,1992; Zachos et al., 1993, 2001; De Conto & Pollard, 2003;Nilsen et al., 2003; Hay et al., 2005; Pollard et al., 2005;Thorn & De Conto, 2006), although recent studies arenow providing ¢ner details about Earth’s climate evolution(Wunsch, 2004; Raymo et al., 2006; Hansen et al., 2007;Lisiecki & Raymo, 2007).

Cenozoic climate changes occur with di¡erent periodi-cities (e.g. Zachos et al., 2001; Lisiecki & Raymo, 2007)namely long-term trends (ca. 106^107 years), rhythmic andperiodic cycles (ca.104^105 years) and rare rapid anomalousshifts and extreme climate transients (ca.103^105 years) de-¢ned as ‘aberrations’ by Zachos et al. (2001).

Cenozoic long-term trends have been related primarilyto the growth and/or decay of the Antarctic and subse-quently Arctic ice-sheets. Since the Eocene, these sheetshave experienced both periods of expansion and stability(global cooling) and periods of contraction and instability(global warming). In detail (Fig. 1), since Palaeocene themost pronounced warming trend of the Cenozoic is re-corded from the Palaeocene until the climate optimum ofthe Early Eocene.This interval is then replaced by a gener-alised cooling trend culminating in the Oi-1 glaciation atthe Eocene/Oligocene boundary. Such e¡ects are probablyrelated to the establishment of a permanent ice-sheet onAntarctica.After this period, a secondweakwarming trendoccurs, culminating with the mid-Miocene climate opti-mum. Ultimately this period is interrupted by the Mi-1glaciation at the Oligocene/Miocene transition. Even-tually, a second gradual cooling trend, related to the estab-lishment of a major ice-sheet on Antarctica, occurs untilthe Pliocene, when the last abrupt cooling begins during

the Late Pliocene, leading to the onset of NorthernHemi-sphere glaciations.

Rhythmic and periodic climate cycles are guided by or-bitally driven oscillations of incoming solar radiation andhappenwith periodicities ranging from 41ka (e.g. Plioceneand Early Pleistocene) to 100 ka (e.g. Late Oligocene, EarlyMiocene and Middle^Late Pleistocene). Lengthening ofperiodic isotope rhythms has also been recorded; thesesignals largely re£ect changes in continental ice-volumeand temperature.

Finally, aberrations are de¢ned (Zachos et al., 2001) asbrief anomalies of the climate system that stand out wellabove the normal background variability in terms of rateand/or amplitude, usually accompanied by a major pertur-bation in the global isotope signal (e.g. Palaeocene/EoceneThermalMaximum, Oi-1andMi-1glaciations).

The global climate interacts with regional factors, suchas the elevations of mountain chains, neighbourhoods ofepi-continental seas, and type and extent of the vegetationcover in creating regional climate conditions. One exam-ple is continental Europe during the middle^late Ceno-zoic (Mosbrugger et al., 2005; Bruch et al., 2007).

According to certain authors (Fauquette et al., 1999,2006; Utescher et al., 2000; Fauquette & Bertini, 2003;Mosbrugger et al., 2005; Bruch et al., 2006, 2007; B˛hme etal., 2007; Erdei et al., 2007; Micheels et al., 2007), the Eur-opean continental climate during the Neogene was char-acterised by warm and humid stable climate conditionsuntil theLateMiocene,when itwas replaced by amore dy-namic, cooler regional climate. It appears to be charac-terised by a gradual decrease in annual temperatureleading to an increase in seasonality, with conditions re-maining generally wet but accompanied by the develop-ment of more seasonal rainfall patterns.Then, during thePliocene, the European climate became cooled further in-terrupted only by a short-term warming episode in theearly Zanclean. During the Piacenzian^Gelasian, a moresigni¢cant decline in temperature occurredwith the onsetof the Northern Hemisphere glaciation dated at 2.7Ma(Lisiecki & Raymo, 2007 and references therein), whichled to the cooler glaciated climate of the Pleistocene.

In theAlps-Po Plain system,where theVFB thatwe stu-died is located (Fig. 2), regional climate conditions duringthe Neogene are generally similar to those described forcontinental Europe (Bruch et al., 2007 and referencestherein). During the Early^Middle Miocene, the circum-alpine climate was characterised by subtropical to warm-temperate and humid conditions. From the Serravallianto theTortonian, there was a generally weak cooling trend,leading to an increase in seasonality. This was especiallytrue in the Alpine region (Bruch et al., 2007), where the ef-fects mayhave been accentuated by the uplift of the moun-tain chain. However, such climate change is not onlyrelated to the elevation gains in the Alpine mountains.The retreat of the Paratethys also reduced of the marinein£uence in the Pannonian domain and probably also inthe eastern portion of the Po-Plain basin, with a conse-quent decrease in temperatures and increase in seasonality

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N. Mancinet al.




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C18O

(a)(b)

(c)

Fig.1. Synoptic sketch of relevant Cenozoic climate events, (a) average oxygen and carbon isotope curves and major climate eventsduring Cenozoic (redrawn after Zachos et al., 2001); (b) benthic d 18O stack redrawn and partly modi¢ed after Lisiecki & Raymo (2005);(c) main Cenozoic European and northernMediterranean climate events synthesised from the literature. For details see the text.


Tectonic vs. climate forcing in the Southalpine foreland basin


in the Late Miocene (Fauquette et al., 2006; Bruch et al.,2007;Martinetto et al., 2007). During the Zanclean, the PoPlain climate was still characterised by stable subtropicalto warm-temperate conditions (Bertini, 2001; Fauquette& Bertini, 2003; Bertini & Martinetto, 2008), replacedsince the Piacenzian by cooler and less stable conditions.Fossil pollen data testify to the development of an altitudi-nal forest (mainlyPicea genus) at the borders of the PoPlainbasin since 2.7Ma.This correlates well with the maximumexpansion of theArctic ice-sheet (Bertini, 2001).ThePlio-Pleistocene cooling trend culminates at ca. 0.87Ma withthe ¢rst major glaciation recorded in the Southern Alps(Muttoni et al., 2003, 2007).

Sediment supply to foreland basins is controlled bya complex interplay of climatic and topographic factors,including: regional climate, seasonality of precipitation,vegetation cover and relief and size of the drainage area(Van der Zwan, 2002 and references therein). In parti-cular, the seasonality of rainfall in a semi-arid climateenhances erosion more than an increase in average riverdischarge (Molnar, 2001), even more so when accompa-nied by a reduction in the protective e¡ect of the vegeta-tion cover (Blaine, 1990). This climate-driven forcingcombines with tectonically driven processes that buildmountain belt relief. The result is that sediment yieldtransfers from mountain ranges to foreland basins by

(a)

(c)

(b)

Fig. 2. Location of the studied Venetian-Friulian basin (a,b) and tectonic sketchmap (c) reporting the studiedwells andoutcrop sections with respect to theNNW^SSE trending AA0 pro¢le and theNE^SW-trending BB0 pro¢le shown inFig. 3.


N. Mancinet al.

https://www.researchgate.net/publication/248616571_Pliocene_climatic_cycles_and_altitudinal_forest_development_from_27_Ma_in_the_Northern_Apennines_Italy_Evidence_from_the_pollen_record_of_the_Stirone_section_51_to_22_Ma?el=1_x_8&enrichId=rgreq-8ae2cda7-7b59-43b9-931d-940688474738&enrichSource=Y292ZXJQYWdlOzIyNzgwMjg0MTtBUzoxMDEwOTY5NTM0ODMyNjRAMTQwMTExNDgwMDI3OA==

https://www.researchgate.net/publication/241241682_Paleoclimate_Controls_on_Stratigraphic_Repetition_of_Chemical_and_Siliciclastic_Rocks?el=1_x_8&enrichId=rgreq-8ae2cda7-7b59-43b9-931d-940688474738&enrichSource=Y292ZXJQYWdlOzIyNzgwMjg0MTtBUzoxMDEwOTY5NTM0ODMyNjRAMTQwMTExNDgwMDI3OA==

https://www.researchgate.net/publication/249521038_Onset_of_major_Pleistocene_glaciations_in_the_Alps?el=1_x_8&enrichId=rgreq-8ae2cda7-7b59-43b9-931d-940688474738&enrichSource=Y292ZXJQYWdlOzIyNzgwMjg0MTtBUzoxMDEwOTY5NTM0ODMyNjRAMTQwMTExNDgwMDI3OA==

https://www.researchgate.net/publication/224962450_Climate_change_flooding_in_arid_environments_and_erosion_rates?el=1_x_8&enrichId=rgreq-8ae2cda7-7b59-43b9-931d-940688474738&enrichSource=Y292ZXJQYWdlOzIyNzgwMjg0MTtBUzoxMDEwOTY5NTM0ODMyNjRAMTQwMTExNDgwMDI3OA==

means of natural erosion processes that shape the Earth’ssurface.

METHODS AND DATASET

The database for our study comes from a complete revi-sion of a large set ofVFB stratigraphic data and subsurfacesamples, most of which were originally collected for thepurposes of hydrocarbon exploration. In particular, inte-grated foraminiferal and nannofossil biostratigraphy wascompletely reviewed by analysing over 400 cuttings from13 wells located along two seismic transects roughly ortho-gonal to theSouthalpine andDinaric (Apennine) thrust beltfronts (Figs 2 and 3). The results of this review are sum-marised in the chronostratigraphic schemes of Figs 4 and 5(for biostratigraphic details seeMancin et al., 2007).

Bio-chronostratigraphic data were used to better con-strain the ages of depositional^erosional events recorded inthe basin, in order to calculate the average depositional ratesof the units and to detail the geometries of the sedimentarybodies consistentwith available seismic data (Fig. 3).

The original depositional thickness of each stratigraphicunit was assessed through decompaction performed withBasinMod 1D software (by Platte River Associates, http://www.platte.com/), which o¡ers three mechanical decom-paction algorithms for three di¡erent types of depositionalunits.These are: the Exponential Mechanical Compaction(EXP), the Reciprocal Mechanical Compaction (RCP) andthe Baldwin and Butler method (BB).

The exponential method (EXP)was empirically derivedby Sclater & Christie (1980) mainly for terrigenous sedi-ments with high sedimentation rates. It implies an expo-nential relationship for porosity reduction with depthaccording to the equation f ¼ f0 exp

�kz where j is thepresent porosity, j0 the initial porosity, k an exponentialfactor related to the di¡erent lithologies andz is the depth.

The reciprocal method (RCP)was proposed byFalvey&Middleton (1981)mainly for carbonate platform sedimentscharacterised by shallow depths and low sedimentationrates. It assumes that porosity changes proportionally to

the change in sediment load according to the equation1/j5 1/j01kz wherej andj0 are the present and initialporosity, respectively, k a reciprocal factor and z is thedepth.

Finally, the BB method was proposed by the authorsmainly for pelitic sediments (Baldwin & Butler, 1985).Themethod is mostly focussed on the lithology according tothe equation z5 6.02 S6.35 where z is the depth (in km)and S is the solidity, that is the inverse of the porosityj.

The choice of the decompactionmethod for the subsur-face depositional units was tentatively made taking inaccount mainly the di¡erent lithologies and the accumu-lation rates. Afterwards, when possible the present poros-ity values measured in some wells have been comparedwith the porosity loss curves calculated by BasinMod1Dsoftware to calibrate the obtained values. The input dataused to perform decompaction and calculate the sedimen-tation rate curves for each of the wells are summarised inTable 1. The resulting curves for sedimentation rate areshown in Figs 7 and 8 whereas the schematic geometriesof the sedimentary bodies are reported in Figs 9 and10.

Consistentwith the decompacted thicknesses of the stra-tigraphic units, the depositional rate was calculated overtime for both single wells and the entire basin. In so doing,an age model was obtained for each well by calculating theaverage depositional rate between each pair of successivechronostratigraphically calibrated biostratigraphic events.The resulting value was applied to each time interval di-vided into steps of 0.5Myr, from 60 to 5Ma, and into stepof 0.1Myr, from 5Ma to the present. Afterwards, the meandepositional rate of the basin was estimated for each 0.5 or0.1Myr time step by averaging the depositional rates as in-dicated by data from the wells over the same time interval.

The results that we calculated for the basin must clearlybe considered estimates, as the chronological resolutionsof the obtained curves strongly depend on the biostrati-graphic resolution of each interval and are not comparableto climatic variability at the orbital scale (10^100 ka).Nevertheless our data can be used to compare mass £uxesfrom the highlands to the basin during climatically di¡er-ent time spans.

Fig. 3. Geologic pro¢les derived from seismic sections oriented orthogonal to the South Alpine (AA0) andDinaric (BB0 ) directions oftectonic transport; the locations of the wells used to calibrate the sections are also reported.Transect AA0 is after Barbieri et al. (2002),and transect BB0 is after Pieri &Groppi (1981).



http://www.platte.com/






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Fig.4.

Chron

ostratigraph

icdiagram

oftheCenozoicsubsurface

succession

oftheFriulian^Venetian

area

alon

gtheAA0 transect(biostratigraph

icdataafterM

ancinetal.,2007).


N. Mancinet al.

STRATIGRAPHIC FRAMEWORK

The VFB represents a tectonic knot of three collisionalbelts: the eastern Southern Alps, the Dinarides and theNorthern Apennines. This area is located between thePo-Plain Basin and the Adriatic Basin (Fig. 2), the mostimportant sites of gas reservoirs in Italy (Casero, 2004).Hydrocarbon research provides an extensive amount ofsubsurface data for this reason, and, as a consequence,the lithostratigraphic units of the Cenozoic VFB sedi-ments have beenmainly de¢ned at and aroundwells (Don-di & D’Andrea, 1986; Venturini, 2002). The data are onlylocally integrated with outcrop sections exposed alongthe north and east margins of the basin (Massari et al.,1986; Tunis & Venturini, 1987, 1992; Stefani & Grandesso,1991;Mellere et al., 2000).

As shown in Figs 4 and 5, the CenozoicVFB ¢ll is com-posed of ¢ve depositional sequences (S) bounded by fourmajor unconformities and their correlative conformities(U), recognised regionally throughout the basin. Both thesequences and the sequence boundaries are describedbrie£y below.

Palaeoceneûpper Eocene sequence (S1)

Hemipelagic ¢ne-grained carbonate-rich sediments (Sca-glia Rossa) dominate the Palaeoceneûpper Eocene part of

the sequence. A somewhat di¡erent pattern is recorded inthe northwestern and northeastern parts of the basin. Inthe NW, the Scaglia Formation is progressively replacedby more terrigenous and coarsening-upward sediments(Vena d’OroMarl and Belluno Flysch), consistent with theincreasing detrital input from the rising Eastern Alps. Inthe NE, the accumulation of PalaeoceneÊocene Num-mulitic Limestone on top of the Mesozoic platform, fol-lowed by deep sea turbidites (Cormons Flysch), suggestsplatform subsidence and drowning triggered by the advan-cing Dinaric belt.

An increase in terrigenous input is recorded basin-wideby the transition from limy (Scaglia Rossa) to marly (Gal-lare group) sedimentation.This increase culminates in thenorthwestern part of the basinwith a progradational trendof deposition changing from deep-sea (Jesolo Flysch) todeltaic (PossagnoMarl) sediments.

U1unconformity

The U1 unconformity denotes the upper boundary of thePalaeoceneÊocene sequence; it is well developed in thenorthern part of the basin, and it passes southward to itscorrelative conformity running into hemipelagic marls.

The U1-related hiatus changes strongly in response tothe combined e¡ects of erosion of S1 sediments (NE endof the basin; Fig. 5) and onlapping termination of the over-

Fig. 5. Chronostratigraphic diagram of the Cenozoic subsurface succession of the Friulian^Venetian area along the BB0 transect.(biostratigraphic data afterMancin et al., 2007). Symbols are the same as in Fig. 4.



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Tab

le1.

Inpu

tdataused

tocalculatethesedimentation

ratecurves

ofthestud

iedwells

Stratigraphicun

its

Sedico

Volpago

SanDon

a'Eraclea

Jesolo

Amira

Agata

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

Astigroup

Eroded

1.8

EXP

1020

1.8

EXP

965

1.8EXP

998

1.3

EXP

597

1.8

EXP

1680

0.2

EXP

295

Con

egliano

unit

0.9

EXP

793

Santerno

group

1.7BB

629

2.2

BB

492

0.1

EXP

88Eraclea

Sand

ston

e2

EXP

223

2EXP

172

2EXP

84

Eroded

Mon

tello

Con

glom

er.

1.5

EXP

350

VittorioVenetoSand

.2.2

EXP

643

SanDon

a'Marl

1BB

320

2.8

BB

300

Upp

erGallaregroup

1.5

BB

598

1.3

BB

121

2.8

BB

125

2.8

BB

105

Cavanellagroup

1.5

RCP

362

4.5

RCP

139

4.5

RCP

110

4.5

RCP

205

4.7

RCP

112

8RCP

248

Possagno

Marl

1EXP

511

Hiatus

1EXP

212

1EXP

75JesoloFlysch

1.5

EXP

160

3.5

EXP

348

3EXP

532

Belluno

Flysch

8.5

EXP

1313

11.5

EXP

496

Not

sampled

Gallaregroup

6BB

350

10.5

BB

865

15BB

573

13.9

BB

573

3BB

316

2.6

BB

390

Scaglia

RossaFm.

2.5

BB

751

BB

4010

BB

218

Stratigraphicun

its

Gallare

Cod

igoro

Ada

Assun

taEraclea

Lavariano

Cargn

acco

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

Astigroup

1.8

EXP

1850

1.4

EXP

2068

1.3

EXP

882

1.4

EXP

543

1.8

EXP

965

1.6

EXP

327

Eroded

0.2

EXP

729

0.2

EXP

383

Santerno

group

0.2

BB

760.2

BB

135

0.3

BB

55PortoGaribaldiFm.

0.8

EXP

556

0.7

EXP

1318

PortoCorsiniFm.

0.7

EXP

584

Eraclea

Sand

ston

e1.3

RCP

622

EXP

172

0.5

EXP

495

Low

erSanterno

gr.

1.2

BB

268

2.2

BB

247

1BB

293

Borecaconglomer.

0.6

EXP

217

Eroded

Eroded

Gessoso

Sol¢feraFm.

0.5

EXP

86VittorioVenetoSand

.1

EXP

207

SanDon

a'Marl

1.8

BB

303

Upp

erGallaregroup

0.5

BB

279

0.7

BB

217

2.8

BB

125

3.8

BB

280

2.8

BB

216


N. Mancinet al.

lying S2 units (Fig. 5).The recorded stratigraphic gap lo-cally reaches the whole BartonianÂquitanian, due to botherosion of Bartonian^Priabonian strata and non-deposition of OligoceneÂquitanian units.

Chattian^Langhian sequence (S2)

Above the U1 surface, the Upper Oligocene^Middle Mio-cene depositional sequence (S2) is characterised by shal-low-water mixed carbonate-terrigenous platformsediments (Cavanella group); it accumulated on the wholebasin £oor previously smoothed byS1 ¢lling of accommo-dation space.

The S2 shallow-water sedimentation began earlier (lateChattianÂquitanian) to the north, where it reached theSouthalpine foothills (Mellere et al., 2000), and later (Bur-digalian) in the rest of the basin. It reaches part of the Car-nic Prealps and some external orogenic structures to theeast (Cargnacco anticline;Venturini, 2002), as well as partof the Northern Adriatic o¡shore to the south. Deeper¢ne-grained sedimentation persisted until the MiddleMiocene only in the southwesternmost portion of thebasin.

U2 unconformity

The second basin-wide unconformity (U2) forms the topof theChattian^Langhian sequence. It contains a gap gen-erally encompassing the late Langhian^SerravallianêarlyTortonian interval, and generally widens from north tosouth due to progressive onlap of Serravallian units onthe underlying Burdigalian Cavanella sediments (Fig. 4).Towards the SW, it passes to its correlative conformity intothe hemipelagic marls of the Gallare group (Fig. 5).

MiddleÛpper Miocene sequence (S3)

The MiddleÛpper Miocene depositional sequence (S3)covers the U2 surface. In the northern part of the basin, itis represented by thick, shallowing-upwards, clastic sedi-ments (from Gallare hemipelagic marls to Montello fandelta conglomerates) that record progressive ¢lling of theSouthAlpine foredeep.Conversely, to theSW,hemipelagicsedimentation continued until the early Messinian, whenevaporitic (Gessoso Sol¢fera Formation) to transitionalfan delta (Boreca Conglomerate) deposition was triggeredby the sea-level drop associated with the Messinian sali-nity crisis.

U3 unconformity

The unconformity forming the top of sequence U3 is themost prominent in the basin. It is the result of the well-known salinity crisis that occurred in the Mediterraneanduring the lateMessinian (e.g. Roveri et al., 2003 and refer-ences therein).

The erosional gap associated with sequence S3 re£ectsthe drainage pattern that developed along the margin oftheMediterranean basin during the lateMessinian sea-le-vel drop. Consequently, the stratigraphic gap associatedT

able1.

(Con

tinu

ed)

Stratigraphicun

its

Gallare

Cod

igoro

Ada

Assun

taEraclea

Lavariano

Cargn

acco

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

TI

(Myr)

DEC.

DT

(m)

Cavanellagroup

Not

sampled

0.5

RCP

764.5

RCP

110

4.5

RCP

247

1.4

RCP

811

RCP

118

Possagno

Marl

1EXP

212

Not

sampled

Eroded

JesoloFlysch

3.5

EXP

348

Gallaregroup

13BB

1028

14.5

BB

694

4BB

234

15BB

571

1.5

BB

387

10.5

BB

865

Not

sampled

3BB

337

8BB

279

15BB

260

14BB

726

Cormon

sFlysch

6EXP

874

Scaglia

RossaFm.

8BB

139

10.5

BB

114

10BB

218

Num

muliticLim

st.

9.5RCP

224

The

thickn

essesofstratigraphicun

itshave

been

correctedforcom

pactionusingtheBasinMod1D

software.

TI,depo

sition

timeintervalin

millions

ofyears;DEC,decom

pactionalgorithm;D

T,decompacted

thickn

essin

meters.



with the Messinian unconformity strictly depends on thedepth of the erosion su¡ered by the basin. Locally, itreaches down to upper Priabonian strata.

Pliocene sequence (S4)

The Pliocene depositional sequence (S4) records post-Messinian re-establishment of open marine conditions.It is documented by shallow-water sandstones (EracleaSandstone) locally overlain by small foramol carbonateplatforms (Amphistegina layers), grading southward into¢ne-grained hemipelagic sediments (Santerno group)covered locally by £ysch (PortoCorsini andPortoGarbaldiFormations).

The thickness of the sequence varies signi¢cantly as itmirrors the irregular shape of the underlying Messinianunconformity.

U4 unconformity

The unconformity bounding the top of the Pliocene se-quence is observed only in the central part of the basin,where a gap is recorded that encompasses the late Zan-clean to theEarly Pleistocene interval. In the rest of the ba-sin, Quaternary sediments conformably cover Plioceneunits without a detectable hiatus.

Pleistocene^Holocene sequence (S5)

The Pleistocene^Holocene depositional sequence (S5)mainly consists of shallow marine sands and sandstones(Asti group) that are spread across most of the basin. Lo-cally, towards the south, these pass into shallow water tur-bidites evolving upward to rapidly prograding deltaicdeposits. Over most of the basin, these units were replacedby continental deposits during the Late Pleistocene con-sistent with an overall regressive trend.

DEPOSITIONAL EVOLUTION: TECTONICFORCING

During theCenozoic, theVFB acted as a foreland basin forthree fold-and-thrust belts encompassing the easternSouthern Alps, the Dinarides and the Northern Apen-nines (Fig. 2; Fantoni et al., 2002 and references therein).

The resulting basin ¢ll evolution is described below. Forthe sake of simplicity, the description is organised intotime slices following the orogenic history experienced bythe surrounding collisional belts, i.e. the southwestwardadvancing Dinaric system during the PalaeoceneÊocene,the southward advancing SouthAlpine system in theMio-ceneÊarly Pliocene, and the northeastward advancingApennine system in the Pliocene^Pleistocene (e.g. Castel-larin et al., 1992; Tunis & Venturini, 1992; Castellarin &Cantelli, 2000; Fantoni et al., 2002).

Middle Palaeocene: the Mesozoic inheritance

Mid-Palaeocene deposition was governed by the basin£oor topography inherited from theMesozoic extensionalframework.This surface, known as the ‘top of carbonates’(Fantoni et al., 2002) is clearly recognisable in seismic pro-¢les at the basin scale and closely approximates the base ofthe Cenozoic succession in the subsurface (Fig. 6, greenseismic horizon).

Late PalaeoceneÊarly Eocene: the Dinaricbackbulge^forebulge system

During this time interval, the basin £oor topography re-mained basically unchanged; pelagic carbonate deposition(ScagliaRossa)withvery lowdepositional rates (usually notexceeding 0.05^0.1mmyear�1; Figs 7 and 8) continued inmost of the study area around the emergedMesozoic plat-form. This re£ects the backbulge position of most of thebasin with respect to the Dinaric collisional system, witha peripheral bulge centred between the Eraclea and Lavar-iano wells (Fig.9e).

A somewhat di¡erent evolution is recorded only at theeasternmost and northwesternmost ends of the basin. Inthe ¢rst area (Fig. 10e), shallow-water platform carbonates(Nummulitic Limestone) accumulated at low sedimenta-tion rates of about 0.05^0.012mmyear�1on top of Cretac-eous platform carbonates (Fig. 8); this evolution re£ectsthe onset of the eastward £exural subsidence of the Friu-lianPlatform in the forebulge region of theDinaric system.Conversely, at the northwestern margin of the basin (Fig.9e), hemipelagic marls (Vena d’Oro marl) followed by tur-bidites (BellunoFlysch) record the ¢rst arrival of the detri-tal input fed by theEasternAlps,which progressively ¢lledthe accommodation space from inheritedMesozoic topo-graphy.

Middle Eocene: the Dinaric forebulge^foredeep system

Since early Lutetian time (at about 49Ma), sedimentationrates have diachronously increased towards theNE along aSW^NE transect (Fig. 8). Higher values are recorded nearthe Dinaric front, where Middle Eocene turbidites accu-mulated at an average rate of 0.15mmyear�1 (CormonsFlysch). These form a clastic wedge (Merlini et al., 2002),that onlaps the eastward-tilted platform surface. Towardsthe south, the depositional rate progressively increasedup to 0.08^0.05mmyear�1 since ca. 48Ma, with the maxi-mum values remaining lower than near the Dinaric thrustfront (Fig. 8).This trend re£ects the £exural origin of thenew accommodation space created by the Dinaric load,which caused a foredeep to grow in the eastern part of thebasin.This new space was ¢lled by detrital sediments fedby the rising Dinaric belt (Stefani et al., 2004), and most ofthe ¢ll postdates £exure, consistentwith the parallel stack-ing of Eocene re£ectors (Merlini et al., 2002).

On the opposite side of the basin, in the NNW sector(Sedico well, Fig. 7), the depositional pattern is similar,


N. Mancinet al.

(a)

(b)

(c)

Fig. 6. (a)Mesozoic palaeogeography of theVenetian^Friulian area (modi¢ed after Cati et al., 1987); (b) seismically derived 3D image ofisodepth map of top carbonate surface, approximating Cretaceous top, clearly showing the morphology of the Friulian Platform; (c)example of seismic expression of the platform margin; note inherited accommodation space ¢lled by lowerTertiary sedimentsonlapping the platform slope.



Fig.7.

Sedimentation

ratesthrough

timeinthestud

iedwellsalon

gtheA

A0 ‘S

outh

Alpine’transect;the

timingofthemainorogeneticph

asesisalso

repo

rted

forcom

parison.Thesegraphs

show

the

distribu

tion

softhedi¡erent

stratigraphicun

its,aswellasthe

location

sthroughtimeofthedepo

centres.The

dashed

linemarks

theon

setofS

outh

Alpineforedeep

¢llin

g.Orogeneticph

ases

are

synthesisedafterTun

is&Venturini(1992),Castellarinetal.(1992)and

Castellarin&Cantelli

(2000).S

eethetextforfurtherexplanation.


N. Mancinet al.

Fig.8.

Sedim

entation

ratesthrough

timeinthestud

iedwellsalon

gtheB

B0 ‘dinaric’transect;tim

ingofmainorogeneticph

asesisalso

repo

rted

forcom

parison.Thesegraphs

show

thedistribu

tion

ofthedi¡erent

stratigraphicun

its,aswellasthe

location

sthrough

timeofthedepo

centres.The

dashed

linemarks

theon

setofthe

dinaric£yschesabovetheFriuliancarbon

ateplatform

(indarkgrey).

Orogeneticph

ases

aresynthesisedafterTun

is&Venturini(1992),Castellarinetal.(1992)and

Castellarin&Cantelli

(200

0).S

eethetextforfurtherexplanation.



(a)

(b)

(c)

(d)

(e)

Fig.9. Restored geometry of basin- ¢lling units along the AA0 transect during the main steps of basin evolution. Stratigraphic data ofthe Feltre and Follina outcropping sections are afterMellere et al. (2000). Note: (1) the strikingly di¡erent patterns of the Eocene (E),Miocene (D^C^B) and Quaternary (A) basin ¢ll steps, clearly re£ecting the di¡erent origins of ¢lled accommodation space (inherited,£exural, and possiblydue to buckling, respectively); and (2) the contrasting geometry of theMessinian erosional unconformity (sketchC)in the present-day onshore and o¡shore areas (NNWand SSE of Jesolo, respectively) probably due to the morphological thresholdcreated by the UpperMiocene South Alpine peripheral bulge.


N. Mancinet al.

but the tectonic context is completely di¡erent. Sedimen-tation rates increased, reaching values of 0.154mmyear�1

(at about 52.5Ma), and deposition changed from hemipe-lagic to turbiditic (Belluno Flysch) sediments gradingsouthward into distal marly facies (Gallare group). Bothprovenance (Stefani et al., 2004) and timing of this rein-forced clastic input point to tectonic control from the Di-naric thrust system. Nevertheless, the geometries ofdepositional bodies clearly show that the accommodationspace ¢lled by turbidites was still inherited fromMesozoictopography, since Eocene turbidites onlap the steep slopebounding the southwestern margin of the Friulian Plat-form (Figs 6 and 9e).

Late Eoceneêarly Oligocene: the transitionfrom Dinaric to South Alpine control

During the Late Eocene, the tectonic framework of theVFB changed fundamentally. In the western area, the resi-dual accommodation space was ¢lled by younger turbiditesediments (Jesolo Flysch and Possagno Marl) depositedabove the middle Eocene sequence. This depositionalphase was accompanied by a further increase in sedimen-tation rates reaching values of 0.2^0.5mmyear�1 (South-alpine Flysches in Fig. 7), along with an increasingcontribution from the Eastern Alps, mixed with still-dominant clastic sediment fed by theDinaric chain (Zattinet al., 2003; Stefani et al., 2004).This is consistent with theweak increase in sediment discharge from theEasternAlpsrecorded in the opposite, northern side of the orogenicsystem (Kuhlemann, 2000; Kuhlemann &Kempf, 2002).

It is worth noting that an increase in sedimentation rateby one order of magnitude or more (from 0.01 to 0.11^0.26mmyear�1 since the Bartonian to upper Priabonian) isalso recorded southward, even though the hemipelagicmarly facies remained apparently unchanged (Figs 7 and 8).

This pattern strongly supports the idea of a signi¢cantincrease in the terrigenous input to the basin, although lo-cal starvation (Assunta well; Fig. 8) occurred during theearliest Oligocene (between 34 and 32Ma). The peculiarbehaviour recorded in the Assunta well has two possibleexplanations: ¢rst, it could be due to the well location intheDinaric peripheral bulge area, consistentwith two-di-mensional (2D) £exural modelling (Fantoni et al., 2002; DiGiulio etal., 2005a), or second, it could re£ect a slope envir-onment of thewell site, due to a complex trend in planviewof the slope-basin boundary similar to that shown by theunderlying platform margin (Fig. 6).

To the NE, the area was a¡ected by subaerial erosionsince the late Bartonian, at about 38Ma, caused by activityof the Dinaric thrusts (Venturini & Tunis, 1991). Thisagrees with apatite ¢ssion track data recording exhuma-tion ages of 40^35Ma in the external part of the Dinaricbelt (Di Giulio et al., 2005b). In the central portion of thebasin, subaereal erosion is recorded since the late Priabo-nian.By contrast, towards theNNW(Sedicowell), the ero-sion of the post Lutetian succession is younger and is

related to combined South Alpine tectonics and Messi-nian sea-level drop (described below).

Chattian^Langhian: the South Alpinebackbulge^forebulge system

During Chattian^Langhian time, a mixed terrigenous-carbonate platform (Cavanella group sediments) devel-oped across most of the basin. These shallow marine de-posits accumulated very slowly; no faster than0.05mmyear�1 (Figs 7 and 8), at least in the lower part ofthe group. The older sediments are recorded to the NW(upper Chattian), and progressively become younger to-wards the SE (lowerMiocene), following the South Alpinevergence (Figs 7 and 9d). Conversely, in the southern partof the basin, hemipelagic ¢ne-grained sediments (Gallaregroup) continued to accumulate at a similar sedimentationrate (about 0.04mmyear�1). Only in the southernmostportion of the basin, a faint increase in the sedimentationrate is recordedwithvalues around 0.08mmyear�1 (Fig.8)since the Chattian/Aquitanian boundary (at ca. 23.8Ma).This increase can be interpreted as the distal e¡ect of thedynamics of the Northern Apennine collisional system,where thick coeval turbidite sequences (Marnoso Arena-cea; Gasperi et al., 1986; Vai, 1987) accumulated in the ad-vancing Northern Apennine foredeep, with their distalmuddy facies reaching the southernmost end of the studyarea. This is supported also by seismic pro¢les (Fig. 3),which show lower Miocene strata gently thickening to-wards the SW near the Northern Apennine margin. Thedepositional pattern during this time re£ects the de¢nitivesmoothing of previous basin £oor topography in the con-text of relative tectonic quiescence, since only a slight £ex-ure towards the north due to forebulge subsidence isrecorded by a moderate thickness increase of the Chat-tian^Langhian sediments towards the north (Figs 9 and10, step D).

This pattern contrasts quite strikingly with that re-corded on the northern side of the Eastern Alps, where asigni¢cant increase in detrital discharge is recorded since30Ma, with a peak around 22Ma, followed by a strong de-crease around 21^20Ma, a new increase at 18^17Ma andanother decrease at16Ma (Kuhlemann, 2000; Kuhlemann& Kempf, 2002). According to the cited authors, this pat-tern was mostly driven by tectonics acting in the northernpart of the belt, with compression and extension phasescausing detrital discharge to increase and decrease in theNorthern Alpine Foreland Basin. Thus, during this timeinterval, a strongly asymmetric behaviour of the collisionalsystem is evident, with the bulk of convergence accommo-dated by the pro-side of the belt.

Serravallianêarly Messinian: the SouthAlpine foredeep^forebulge system

The Serravallianêarly Messinian time span constitutesthe tectonically most active period in the Eastern South-ern Alps history (Castellarin et al., 1992). Rapid exhuma-tion of hinterland units, estimated at around 4 km,



(a)

(b)

(c)

(d)

(e)

Fig.10. Restored geometry of basin- ¢lling units along the BB0 transect during the main steps of basin evolution. Stratigraphic dataof the Cargnacco and Bernadia wells are after Venturini (2002) andMerlini et al. (2002), respectively. Note: (1) the striking asymmetryof the two margins of the Friulian Carbonate Platform during the Eocene pattern (E) due to the di¡erent origins of ¢lledaccommodation space, that was inherited from the Palaeocene to the SW, and on the contrary, created by £exure to the NE; (2) thecontrasting geometry of theMessinian erosional unconformity (sketch C) in the present-day onshore and o¡shore area (NNWandSSE from Jesolo, respectively) probably due to the morphological threshold created by the UpperMiocene South Alpine £exure; and(3) the di¡erent pattern of Pliocene and Quaternary steps of basin ¢lling which still re£ect strongly di¡erent distributions ofaccommodation space.


N. Mancinet al.

started at about 10Ma according to thermal modelling ofapatite ¢ssion track data (Zattin et al., 2003; Di Giulioet al., 2005b) and had a strong impact on theVFB in termsof both £exural subsidence and clastic input (Fantoni etal.,2002 and references therein; Zattin etal., 2003). In particu-lar, provenance data on clastic units prove that the mainsource change triggered by the tectonics along the Valsu-gana thrust system occurred since the late Tortonian(Zattin et al., 2003); at a basin scale, this allows to betterconstraint at about 7^8Ma the beginning of the most im-portant e¡ects of theValsugana tectonic activity on the det-rital £ux in theVFB.

Flexural subsidence is clearly recorded by the wedgegeometry of the Serravallian^Messinian clastic wedge(Fig. 9c), which passes from more than 3000m thick inthe sections exposed in the Venetian foothills (Massari etal., 1986; Mellere et al., 2000) to about 1500m in the Vene-tian plain, progressively decreasing to zero in the present-day o¡shore area.

This depositional trend is also clearly recorded by thedepositional rate curves (Fig. 7) that show an abrupt in-crease, which is stronger and occurs earlier close to theeastern Southern Alps belt (Volpago well) and progres-sively vanishes during latest Miocene times. Only in thesouthwesternmost part of the basin did ¢ne-grained,mostly hemipelagic, deposition continue until the earlyMessinian. Here slow sedimentation (0.073mmyear�1)stopped only when evaporitic (Gessoso Sol¢fera Forma-tion) to transitional fan delta deposition (BorecaConglomerate) was triggered by the onset of the Mediter-ranean salinity crisis.

At the scale of the orogenic system, this is another timespan during which a markedly di¡erent pattern of detritaldischarge can be seen by comparing the southern andnorthern sides of the collisional belt (present work andKuhlemann et al., 2002, respectively). Nevertheless, theimportance of tectonics as a driving process for detrital in-put in the foreland basins was inverted with respect to theEarly Miocene, when tectonics was dominant on thenorthern side of the system and subordinate in the south.

Late Messinian: the salinity crisis

During the late Messinian, subaerial erosion a¡ected thewhole basin due to the sea-level drop during theMediter-ranean salinity crisis (Roveri etal., 2003;Manzi etal., 2005).Erosion removed part of the studied successions, some-times reaching the upper Priabonian units, with an esti-mated eroded thickness locally exceeding1500m.

The Messinian erosion, however, is recorded in a strik-ingly di¡erent way in the Venetian region compared withthe Adriatic o¡shore area. First, erosion is channelled,with incisedvalleys not exceeding 300m in depth, whereasin theAdriatic o¡shore, erosion acted more strongly, creat-ing deeply incised canyons similar to those in the westernPo-Plain (Pieri &Groppi, 1981; Fantoni et al., 2001, 2004).

According to 3D £exural modelling (Barbieri et al.,2004; Di Giulio et al., 2005a), the boundary between these

two regions corresponds to the location of the peripheralbulge in the South Alpine foreland basin during LateMiocene £exure.Thus, the peculiar pattern of Messinianerosion can be reasonably linked to the occurrence of a tec-tonically driven morphological threshold caused by theSouth Alpine £exure, which partly protected the northernpart of the basin from retrogressive erosion (Figs 9 and10,step C).

Early Pliocene: vanishing South Alpine control

The re-establishment of open marine conditions in mostof the VFB during the post-Messinian sea-level rise fol-lowed a pattern guided by the palaeotopography of theMessinian subaerial surface and thrusting activity in theSouthern Alps.

In the central portion of the basin, £exural subsidence isrecorded by the wedge geometry of the Pliocene body (Fig.9b), passing from about180m thick in the central sector ofthe basin to about150 and 70m southwards, and then pro-gressively decreasing to zero in the Adriatic o¡shore area.This South Alpine depositional trend is also recorded bythe depositional rate (Fig. 7), which decreased from 0.11 to0.04mmyear�1. In the rest of the basin, the distribution oflowerPliocene sediments is strongly irregular due to ¢llingofMessinian incised valleys.

Late Pliocene^Pleistocene: Apennineforedeep^forebulge system

Since the Late Pliocene, coarse-grained £uvial deposits(Conegliano Conglomerate) were deposited at the north-ern margin of the basinwith averyhigh sedimentation rate(0.88mmyear�1; Fig. 7) sourced by the reactivation of thefrontal thrusts of the eastern Southern Alps.

On the opposite side of the basin, to the SW, the tec-tonic framework is very di¡erent due to Northern Apen-nine tectonics. Since the Piacenzian, the increasingaccommodation space created by the advancing NorthernApennine foredeep guided the deposition of turbidite se-quences (Porto Corsini and Porto Garibaldi Fomations)above the muddy deposits of the Santerno group.This ob-servation is supported by the wedge geometry of clasticunits, which pinch out towards the NE (Fig. 10b), coupledwith an increase in sedimentation rates to the SW (Fig. 8).

This general pattern changed dramatically during thePleistocene, when rapidly prograding shallow marinesands (Asti group) covered most of the basin, completelysmoothing the previous basin £oor topography. Thebroadly symmetric lens-shaped geometry of Pleistocenesediments (Figs 9 and 10), recording a basin-wide extre-mely high depositional rate (exceeding 0.5mmyear�1 overmost of the basin) implies that the e¡ect of £exural tec-tonics on basin ¢lling became minor at least since theMiddle^Late Pleistocene. Similarly, a vanishing tectoniccontrol on sedimentation is recorded in the Northern Al-pine Foreland Basin (Kuhlemann et al., 2002), pointing toan overall decrease in tectonic activity in the Alpine colli-sional system during this time span (Willett et al., 2006;



Champagnac et al., 2007; Vernon et al., 2008), even thoughstill-active deformation is shown by the present-day seis-micity in the Friuli area.

REGIONALTECTONICS VS. CLIMATEFORCING

Dominant regional tectonic forcing in the context of VFBevolution during the Cenozoic has been demonstrated interms of both subsidence and depositional patterns. Insuch a tectonically controlled setting, is climate forcing re-cognisable, and if so, is it signi¢cant?We attempt to answerthese questions below.

To this end, Fig. 11 provides the depositional rates dur-ing the whole Cenozoic and during the Pliocene^Pleisto-

cene, calculated for the VFB at 0.5 and 0.1Myr time steps,respectively. Results are discussed focusing on some keytime spans for which climate forcing can be better re-solved, with special emphasis on those intervals duringwhich major short-term climate anomalies were recordedworldwide (e.g. 35^33, 20^14, 14^12, 5.3^2.1 and1.8^0Ma).

35^33Myr (Late Eocene^Early Oligocene)

The ¢rst basin-wide event in the VFB that was possiblyimpacted by climate forcing is recorded within ZoneIFP18, at ca. 35Myr (Priabonian), when a clear increase interrigenous input into the foreland basin is evident from ajump in the depositional rate of at least one order of mag-nitude. It is worth noting that a similar coeval change in

EASTERN ALPS VENETIAN-FRIULIAN AREA

Fig.11. Curves showing, from left to right: the Cenozoic sediment discharge rate reconstructed for the northern side of the EasternAlps (redrawn after Kuhlemann et al., 2001), the Cenozoic depositional rate, calculated at 0.5Myr steps, and the Pliocene^Pleistocenedepositional rate, calculated at 0.1Myr steps. Data are shown for theVFB according to both the stratigraphic dataset provided by thestudiedwells and the methods described in the text.


N. Mancinet al.

sedimentation, with a strong decrease of the carbonate/terrigenous ratio in sediments and a consequent increasein ¢ne (clay and silt) terrigenous input, is recorded in thewestern portion of the Po Plain (Scaglia Fm./Gallare gr.boundary, Cerano and Settimo Milanese Eni wells), andin the North Alpine foreland basin (Nummulitic Lime-stone/Globigerina Marl boundary, Sinclair, 1997; Sissingh,1997).

Regarding the cause of this event, we argue that tec-tonics acted diachronously in the Dinaric belt (Fantoni etal., 2002), and in a di¡erent portion of the Alpine chain(Sissingh, 1997; Castellarin & Cantelli, 2000; Kuhlemannetal., 2001; and references therein).This can hardly explaina coeval increase in terrigenous input to the PoPlain and intheVFB andNorthAlpine foreland basins.Moreover, a si-milar increase of terrigenous material has been observedin areas unrelated to the Alpine orogenic system (Robert,1980; Lavier et al., 2001; Robert et al., 2002).

We also observe that a worldwide long-term globalcooling trend is documented since the Bartonian (about42Ma), culminating in a strong isotope shift (the most ob-vious of theTertiary) of the Oi-1 glaciation at the Eocene/Oligocene boundary (about 34Ma; Fig. 1). At a regionalscale, this was probably due to an increase in the seasonal-ity of rainfall, as documented for the Northern MolasseBasin (Germany; Utescher et al., 2000; Mosbrugger et al.,2005). The latter e¡ect was combined with the drop insea-level and the high-relief topography reached by theneighbouring mountain ranges. It resulted in an observa-ble increase in detrital input to theVFB basin.

20^14Myr (Early^Middle Miocene)

In the VFB, shallow-water mixed carbonate-terrigenousdeposits accumulated during the IFP22-Praeorbulina zonalinterval (Upper Chattian^Langhian), locally interruptedby small rhodalgal carbonate factories (Massari et al., 1986;Baccelle Scudeler&Reato,1988;Mellere etal., 2000;Fanto-ni etal., 2002).These are not exclusive, as similar coeval se-diments are widespread throughout the Mediterraneanregion and in the Pannonian basin.They occur in theMal-tese Islands (John etal., 2003), theValenciaGulf (Stoeckin-ger, 1976; Carannante et al., 1988; Seemann et al., 1990), theSouthern Apennines (Carannante et al., 1988), Sardinia(Cherchi &Montadert,1982; Cherchi etal., 2000),Monfer-rato (Vannucci et al., 1996; Piana et al., 1997), the NorthernApennines (Amorosi,1990; Amorosi etal.,1993) and the Si-lesia basin (Parathetys, Ney et al., 1974; Bicchi et al., 2003).

This evidence supports a generalised climate in£uencethat may have accelerated the genesis of these sedimentsdue to the growth of carbonate factories favoured by warm-ing after the Oligocene/Miocene transition (Fig. 1). Thiscaused a subtropical̂ tropical climate to develop at mid-la-titudes, with widespread transgressions and warm-watercarbonate sediments containing subtropical/tropical £oraand fauna, (Carannante et al., 1988; Vannucci et al., 1996;Cherchi etal., 2000).At the same time, from the lateOttnan-gian (late Burdigalian) to the middle Badenian (late Lan-

ghian), the climate pattern in Central Europe wascharacterised by warm andwet conditions with rainfall dis-tributed throughout the year (B˛hme et al., 2007). Thisprobably caused the decrease in runo¡ and the consequentdecrease in detrital input to the basin (Van der Zwan, 2002).

In the VFB, this event is clearly recorded. Owing to therelative tectonic quiescence during this time span, it isbracketed between the Dinaric orogenesis and the climaxof tectonic deformation in the eastern Southern Alps.Nevertheless, it is interesting that, conversely, on thenorthern side of theAlpine range (MolasseBasin), the det-rital input from the belts to the pro-foreland basin seemsto have been modulated mostly by tectonics, with increas-ing and decreasing steps matching compression and ex-tension phases in the belt (Frisch et al., 2000; Kuhlemannet al., 2001, 2002).

14^12Myr (Middle Miocene)

The regional context changed signi¢cantly at the end ofthe Langhian, at about 14Ma. In the VFB, as well as inmost of the Mediterranean area (John et al., 2003) and inthe Paratethyan domain (Bicchi et al., 2003), deposition ofthe lowerMiocene mixed terrigenous-carbonate platformstrata stopped almost synchronously within a depositionalhiatus spanning the Globorotalia peripheroronda ^ the lowerportion of Paragloborotalia mayeri zonal interval (late Lan-ghian^early Serravallian). The Serravallian sediments,which unconformably cover them, are mainly charac-terised by terrigenous muddy facies containing fauna and£ora indicative of colder waters (Bicchi et al., 2003; John etal., 2003).This change can be attributed to climate e¡ects,as it ¢ts the cooling recorded at the Langhian/Serravallianboundary (Mi-3 glaciation, Miller & Feigenson, 1991; Za-chos et al., 2001). It caused a regional climate change, dur-ing the early Serravallian, with the establishment of coolerand drier climate conditions (Utescher et al., 2000; Erdeietal., 2007).This in turn triggered the demise of carbonatefactories and an increase in terrigenous input.

12^10Myr (Tortonian)

During this time span, a weak cooling trend is recordedworldwide (Fig. 1), leading to an increase in seasonality inthe circum-Alpine area (Bruch et al., 2007); in the north-ern £ank of the Alps this resulted in a signi¢cant increaseof rainfall according to herpetological fossil assemblages(B˛hme et al., 2008). This regional climate change wasprobably accentuated by both the uplift of the Alpinerange and the reduction of the marine in£uence of theParatethys (Fauquette etal., 2006; Bruch etal., 2007;Marti-netto et al., 2007).

In the studied basin, sedimentary rocks depositedwith-in the Globorotalia menardii/Neogloboquadrina acostaensisZone (Tortonian s.l.) are sparsely preserved due to laterMessinian erosion. Basically, they are preserved only inthe depocentre of the eastern Southern Alps foredeep.Unfortunately, in this location, the eventual e¡ect ofclimate in terms of increased erosion rate in the belt and



consequent increase in depositional rate to the basinwould have acted in phase with intense South Alpine tec-tonics.The consequence of the combined poor preserva-tion and simultaneous change of climate and tectonicsmakes it hard to detect the eventual e¡ect of the LateMio-cene cooling on deposition in the basin.

Nevertheless, it can be observed that before LateTorto-nian^Messinian spikes of VFBdepositional rate, thatwerequite surely tectonically driven, the average VFB deposi-tional rate records a three-fold increase from 0.04 to0.13mmyear�1 since 12^11Ma; followed by a slight de-crease to 0.11mmyear�1 in the 10^9Ma interval and anew increase to 0.13mmyear�1 at around 9^7Ma. Even ifthese data must be handledwith care for the reasons men-tioned above, these variations perfectly ¢t those of themean annual precipitation obtained by B˛hme et al.(2008) from palaeontological data for SWEurope; this factseems something more than a pure coincidence.

5.3^2.1Myr (Pliocene)

Even during the Early Pliocene, climate forcing probablyconditioned sedimentation in the VFB and, more gener-ally, in Europe and the Mediterranean region, causing thedeposition of small foramol carbonate platforms withabundant Amphistegina assemblages over shallow-waterclastic deposits.This kind of facies is also found in Sicily(Roveri & Taviani, 2003), Central Italy (Di Bella et al.,2005) and the Northern Apennines (Capozzi & Picotti,2003; Roveri & Taviani, 2003).

Their growth was probably favoured by a warm climateand high sea-level during the Early Pliocene (Dowsett etal., 1996), as the Amphistegina assemblage distribution ismainly controlled by temperature and nutrient supply(Hollaus &Hottinger, 1997; Hallock, 2000a, b).The net re-sult is that modern Amphistegina assemblages occur in theMediterranean region only south of 391N; north of thislimit they have been absent since the Late Pliocene (about2.6^2.1Ma; Langer &Hottinger, 2000).

TheLate Pliocene cooling, causingAmphistegina extinc-tion and other signi¢cant biotic turnover across the wholeMediterranean region (Langer & Hottinger, 2000; Mone-gatti&Ra⁄, 2001;Capozzi&Picotti, 2003;Roveri&Tavia-ni, 2003; Di Bella et al., 2005), was also coupled withpeculiar sedimentary products, such as sapropel layers(Meyers & Negri, 2003). Furthermore, in the Mediterra-nean region the e¡ects of Pliocene warming^cooling cy-cles seem to be recorded by the on/o¡ switching ofAmphistegina-bearing foramol deposits.

In the Northern Apennine, Amphistegina deposits verysimilar to the described Venetian^Friulian strata show aregional unconformity at their top.This surface has beeninterpreted as due to the Late Pliocene cooling event andto the related sea level fall (Capozzi & Picotti, 2003) atabout 2.7Ma (Lisiecki & Raymo, 2007 and referencestherein).

From a more general point of view, the Pliocene was aperiod of accelerating global cooling, especially since

2.7Ma (Lisiecki & Raymo, 2007). Consequently, a longer-term climate e¡ect on the detrital £ux in the basin can beexpected, as suggested by data concerning terrigenous in-put at both the regional (northern Alpine Foreland Basin)and the global scale (Peizhen et al., 2001; Kuhlemann et al.,2002;Molnar, 2004).

Such a signal is evident in the VFB, although it is lessaccentuated than we might expect. In particular, an in-crease in the sedimentation rate is recorded since 2.6Ma(Fig. 11); which can be explained by the cooler and lessstable climate conditions in the region, also shown by fos-sil pollen data from the Po Plain basin (Bertini, 2001).Thesame data suggest that thePoPlain region, thanks to its pe-culiar humid environmental conditions, remained coveredby coniferous forests even during glacial phases (Fauquette& Bertini, 2003); this probably reduced erosion limitingthe sediment yield arriving in the basin.

1.8^0Myr (Pleistocene^Holocene)

Since the Pleistocene, a stratigraphic succession up to1000^1500m thick has been deposited in the VFB. Thecalculated sedimentation rate is on the order of 0.6^1.0mmyear�1 (with peaks of 1.9^3.6mmyear�1; Figs 8and 11). This rate is one order of magnitude higher thanthe previous metric and is approximately twice that re-corded during the Serravalian T̂ortonian interval(o0.5mmyear�1; Figs 7^11), which was the main exhu-mation period of the eastern South Alpine chain (Stefanietal., 2007). It is by far the intervalwith the highest deposi-tional rate in the basin during theTertiary.

The basin-scale distribution of the Pleistocene increaseof detrital input, as well as its incompatibility with all im-portant known tectonic events in the neighbouring belts,clearly points to tectonics playing a minor role as a processcontrolling the erosion^deposition system during thistime span. In addition, the detailed stratigraphic frame-work available in the VFB makes it possible to evaluatethe history of detrital input in the basin at 0.1Myr intervals(Fig. 11), one order of magnitude higher than those avail-able to date on the same and in other neighbouring basins(Kuhlemann et al., 2001).

The resolution is strictly limited by the stratigraphicdata and it is not su⁄cient to compare the depositionalrate with climate cyclicity, either before or after theMiddlePleistocene Revolution (around 40 and 100 ka, respec-tively). Nonetheless, it is possible to investigate which,among cooling, climate variability (in terms of both fre-quency and intensity) and widespread glaciations, is themost powerful climate-related process that modulateserosion rates in Alpine-type mountain ranges at mid-lati-tudes and the related detrital £uxes in depositional basins(e.g. Summer¢eld &Hulton, 1994; Liu et al., 2005).

In doing so, we must ¢rst de¢ne the concept of ‘majorglaciations’ and shortly discuss the age of their onset inthe Alps. It is clear that during cooling/warming cycles,Alpine glaciers grew and decayed, but at what point canwe begin to de¢ne Alpine glacier growth as a glaciation?


N. Mancinet al.

This is, surprisingly, a poorly de¢ned concept. In this pa-per, we use the term ‘major Alpine glaciation’ in the senseof the development of an ice cap covering most of the belt,due to largely merging glaciers, with its tongues reachingthe belt margin.The reasons for this choice are historical,practical and scienti¢c. This de¢nition ¢ts well with thecommon use of the term in the geological literature (e.g.Fiebig et al., 2004); the growth phases of Alpine glaciersare mostly known (and dated) by studying the morainesbordering their belts. Most important for this paper, thee¡ect of glacier growth and retreat on the erosion rate ofthe belt should be at a maximum when glaciers form anice cap like the one described above.

This de¢nition allows us to better discuss the age of the¢rst major glaciation in the Alps at least in their Southern£anks. Available literature suggesting a pre-Pleistoceneonset of glaciations in theAlps is sparse and local, possiblyreferring to minor advances of individual alpine glaciers,not ‘major glaciations’as de¢ned above (e.g. Bini&Zuccoli,2004 and references therein). Conversely,many papers (e.g.Castiglioni, 2004; Fiebig etal., 2004;VanHausen, 2004 andreferences therein) support the idea that major glaciationshave occurred in the southern £ank of the Alps since thePleistocene, probably since the Matuyama/Bruhnesboundary (at 0.78Ma) according to a study of morainesalong the belt border. Similar results were recently re-ported by Muttoni et al. (2003, 2007), that date the onsetof major glaciations in the southern £ank of the Alps at0.87Ma (Marine Isotope Stage 22) according to a study ofPo Plain sediments and their pollen contents. Therefore,it seems probable that major glaciations in the Alps didnot begin before the Middle Pleistocene, probably duringthe Middle Pleistocene Revolution which pushed climatecyclicity from 41 to 100 ka frequencies (e.g. Raymo, 1994).If this conclusion is correct, the glaciations started wellafter the onset of the acceleration of global cooling(3^4Ma), as well as after the increase in amplitude of cli-mate variability revealed by isotopic data since 1.5^1.0Ma(Molnar, 2004).

By comparing this chronology of regional and globalclimate change with the depositional rate recorded in theVFB,we argue that climate variability in a period of higherfrequency of variation is the main climate-related factorin£uencing the erosion rate in mountain ranges and detri-tal £ux in the neighbouring basins.This supports the ideathat a ‘predominant’disequilibrium erosional regime is themost e⁄cientway bywhich climate e¡ects may improve theerosion rate in mountain ranges, as suggested by Peizhenet al. (2001). One plausible alternative might be that majorglaciations in the Alps began earlier (around 1.5Ma) thanwhat is supposed by most recent authors (Muttoni et al.,2007). The record of oldest glaciations may be sparse andlocal due only to erasure following glacial advances.

CONCLUSIONS

The Cenozoic depositional history of the VFB demon-strates that the e¡ects of climate can be clearly recorded

in depositional basins whose evolution is strongly con-trolled by regional tectonics.

This is true even for time spans coeval with periodsof intense tectonic activity (as in the case of the Oi-1and Mi-3 glaciations), as well as for periods of relativetectonic quiescence (the Early Miocene). Nevertheless,as expected, some global events do not seem to be re-corded (e.g. Mi-1 glaciation), and others that worked inphase with regional tectonic forcing are di⁄cult to detect(e.g.Tortonian).

The way inwhich climate forcing is recorded in the de-positional record varies from case to case. As a generalrule, cooling events cause terrigenous input into thebasin to increase, resulting in a comparable increase inthe deposition rate and/or the terrigenous/indigenousgrain ratio (e.g. clay vs. carbonate content in hemipela-gic muds) in marine rocks.This e¡ect is mostly recogni-sable in relatively deep marine environments, butrequires decompacting pelitic sequences and identifyingdetailed chronostratigraphic constraints in order to beextracted.

Conversely, climatewarming (e.g. during the EarlyMio-cene) is commonly recorded by an increase in carbonateintrabasinal detritus, which is more evident in shallowmarine settings. Nevertheless, according to the calculateddepositional rates, this is mostly due to an increase in car-bonate production in the basin rather than to a relative de-crease of terrigenous input, suggesting that the e¡ect onthe erosion rate is minor.

Within this general framework, the trends of deposi-tion rate clearly highlight a revolution that occurredin theVenetian^Friulian region during the Early Pleisto -cene, when there was a basin-wide increase of detritalinput of an order of magnitude. Consistent with theavailable data, we argue that the most important pro-cess controlling this revolution was the combination ofhigh-frequency (41ka) and high-amplitude warming^cooling cycles, more so than the overall Pliocene^Pleis-tocene cooling trend or the onset of major glaciationsin the Alps. Therefore, the case of the VFB supportsthe idea that a predominant disequilibrium state of theerosional system, promoted by concomitant high-amplitude and high-frequency climate variability, is themost important process that modulates sediment £uxfrom mountain belts to sedimentary basins in mid-latitude regions.

ACKNOWLEDGEMENTS

We thank the ENI sta¡ (particularly R. Fantoni) for wellsamples and helpful suggestions, and G.Toscani for seis-mic pro¢les and 3D modelling.We also thank B. Carrapafor her comments on a preliminary version of the paper.P. van der Beek, J. Kuhlemann and two anonymous re-viewers are acknowledged for their constructive feedback.This study was ¢nancially supported by PRIN 2003 (Resp.A. Di Giulio).



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Manuscript received 18 June 2008;Manuscript accepted 23 Jan-uary 2009



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