9c960522eba62937f6

Evolution of the late Cenozoic Chaco foreland basin,Southern BoliviaCornelius Eji Uba1, Christoph Heubeck and Carola Hulka

Institut fˇr GeologischeWissenschaften, Freie Universit�t Berlin, Berlin, Germany

ABSTRACT

Eastward Andean orogenic growth since the late Oligocene led to variable crustal loading, £exuralsubsidence and foreland basin sedimentation in the Chaco basin.To understand the interactionbetween Andean tectonics and contemporaneous foreland development, we analyse stratigraphic,sedimentologic and seismic data from the Subandean Belt and the Chaco Basin.The structuralfeatures provide a mechanism for transferring zones of deposition, subsidence and uplift.These canbe reconstructed based on regional distribution of clastic sequences. Isopach maps, combinedwithsedimentary architecture analysis, establish systematic thickness variations, facies changes anddepositional styles.The foreland basin consists of ¢ve stratigraphic successions controlled byAndeanorogenic episodes and climate: (1) the foreland basin sequence commences between �27 and14Mawith the regionally unconformable, thin, easterly sourced £uvial Petaca strata. It represents asigni¢cant time interval of low sediment accumulation in a forebulge-backbulge depocentre. (2) Theoverlying �14^7Ma-old Yecua Formation, deposited in marine, £uvial and lacustrine settings,represents increased subsidence rates from thrust-belt loading outpacing sedimentation rates. Itmarks the onset of active deformation and the under¢lled stage of the foreland basin in a distalforedeep. (3) The overlying �7^6Ma-old, westerly sourcedTariquia Formation indicates a relativelyhigh accommodation and sediment supply concomitant with the onset of deposition of Andean-derived sediment in the medial-foredeep depocentre on a distal £uvial megafan. Progradation ofsyntectonic, wedge-shaped, westerly sourced, thickening- and coarsening-upward clastics of the(4) �6^2.1Ma-old Guandacay and (5) �2.1Ma-to-Recent Emborozu¤ Formations represent thepropagation of the deformation front in the present Subandean Zone, thereby indicating selectivetrapping of coarse sediments in the proximal foredeep andwedge-top depocentres, respectively.Overall, the lateCenozoic stratigraphic intervals record the easterly propagation of the deformation

front and foreland depocentre in response to loading and £exure by the growing Intra- andSubandean fold-and-thrust belt.

INTRODUCTION

Foreland basin systems develop as a result of £exuralwarp-ing of the lithosphere in response to supralithosphericand sublithospheric orogenic wedging (DeCelles & Giles,1996; P¢¡ner etal., 2002). Lithospheric £exure under staticloads generates down-bending £exure proximal to theorogen, which migrates as the load advances. Forelandbasins therefore exhibit a characteristic asymmetriccross-section. Their sedimentary ¢ll generally preservesand records a detailed £exural response of the continentallithosphere to orogenic loading (Beaumont, 1981; Jordan,1981; Tankard, 1986).The lithospheric response to thrust-ing varies between and within the foreland basin system

but is mainly controlled by the elastic thickness of thelithosphere and the applied loads (Watts, 1992, 2001). De-Celles &Giles (1996) characterized foreland basin systemsinto four di¡erent depocentres:wedge-top, foredeep, fore-bulge and backbulge. Each depocentre exhibits distinctiveinternal architecture, sedimentology and structure. Ac-commodation space is created by combined static anddynamic subsidence (DeCelles & Giles, 1996; Catuneanuet al., 1997).

The Chaco foreland basin of the central South Americais a classic example of a foreland basin system in a retro-arc position. It can be subdivided into the InterandeanZone, the Subandean Zone and the Chaco plain tectono-morphologic units (Uba et al., 2005) (Fig. 1). The basinformed during the late Cenozoic (Sempere et al., 1990; De-Celles&Horton, 2003) in response toNazca-SouthAmer-ican plate convergence and its related eastward interactionwith theBrazilian shield.Detailed structural studies in theInterandean and Subandean Zones documented structur-al styles and timing of deformation (Sempere et al., 1990;

Correspondence: Cornelius Eji Uba, Institut fˇr GeologischeWissenschaften, Freie Universit�t Berlin, Malteserstrasse 74-100, 12249 Berlin, Germany. E-mail: [email protected] Present address: Institut fˇr Geowissenschaften, Universit�tPotsdam,Karl-LiebknechtStr. 24/25,14476Potsdam/Golm,Ger-many.

BasinResearch (2006) 18, 145–170, doi: 10.1111/j.1365-2117.2006.00291.x

r 2006 The Authors. Journal compilation r 2006 Blackwell Publishing Ltd 145

Baby et al., 1992, 1997; Belotti et al., 1995; Dunn et al., 1995;Roeder & Chamberlain, 1995; Welsink et al., 1995; Morettiet al., 1996; Kley et al., 1996, 1999; Mˇller et al., 2002; Echa-varria et al., 2003).

The west-to-east variation in heat £ow and thermalgradient, coupledwith the hydrocarbon potential, thermalmaturity and exhumation history of the Subandean Zoneand Chaco plain, have also been examined (Baby et al.,1995; Dunn et al., 1995; Moretti et al., 1996; Somoza, 1998;Husson & Moretti, 2002; Echavarria et al., 2003; Ege,2004). Furthermore, the uplift history of the central Andeshas been analysed and documented (Isacks, 1988; Gubbelset al., 1993; Kennen et al., 1997; Gregory-Wodzicki, 2001).Echavarria et al. (2003) examined the accumulation ratesof Neogene strata in the Subandean Zone and postulatedtwo-phase deformation, similar to the documentedthree-phase tectonic subsidence history in the Chaco ba-sin by Coudert et al. (1995). However, not much has beendone on the foreland basin evolution and the dynamic in-teraction between the Andean fold^thrust-belt and theforeland basin.

The internal architecture of a foreland basin ¢ll, as wellas its age, lithofacies and thickness, re£ects how the lead-ing edge of the orogen advances over the foreland basin intime and space. These geometries and ages are best stu-died through a combination of outcrop, seismic, and welldata that provide, in their combination, a unique archiveof basin response to orogenic growth (e.g. Schlunegger etal., 1997; Alves et al., 2003; Echavarria et al., 2003). Our

study utilizes 2-D industry seismic data, logs from se-lected exploration wells, and stratigraphic pro¢les fromoutcrops to establish such history for the Chaco forelandbasin of southeastern Bolivia.

Our objectives are (1) to examine the spatial and tem-poral changes in facies, texture and sedimentation rate ofthe foreland strata, (2) to explore links between stages ofAndean structural evolution and basin-stratigraphic re-sponse and (3) to de¢ne and characterize temporal andspatial distribution and migration of the depocentre as re-lated to the deformational front.These objectiveswill con-tribute to a better understanding of tectonic processesbecause the stratigraphic sections record the basinwardpropagation of thrust sheets and their leading depositio-nal systems. Furthermore, a deformational sequence isstraightforward to examine and to interpret in situationswhere variations in sediment composition, palaeo-drai-nage network, £uvial patterns and thickness can be takenas a proxy to accumulation and sedimentation rate (Wad-worth et al., 2003; Jones et al., 2004).

GEOLOGICAL HISTORY

The Subandean fold-and-thrust belt and theChaco foreland basin

The Andean orogeny created and subsequently progres-sively deformed the Chaco foreland basin since the late

Fig.1. Tectonic map of Bolivia modi¢ed after Baby et al. (1995, 1997), Kley et al. (1997), andMˇller et al. (2002) showing majormorphotectonic divisions and location of the study area (in square).

r 2006 The Authors. Journal compilation r 2006 Blackwell Publishing Ltd,Basin Research, 18, 145^170146

C. E. Ubaet al.

Oligocene (Oller, 1986; She¡els, 1988; Sempere et al., 1990;Baby et al., 1992; He¤ rail et al., 1996). This foreland basinforms the easternmost part of the Andean orogen, whichdeveloped from the Cretaceous in the Altiplano (Horton& DeCelles, 2001; DeCelles & Horton, 2003) as a result ofsubduction of the Nazca plate below the South Americanplate and the simultaneous subduction of the BrazilianShield at an initially low rate of 5.8 cmyear�1 to a subse-quent maximum rate of up to 15.2 cmyear�1 (in the lateOligocene; Somoza, 1998). Andean deformation com-menced in the west with formation of the Altiplano basinand foreland sedimentation,with the depocentre probablyat the present Eastern Cordillera (Sempere et al., 1997; De-Celles & Horton, 2003; Elger et al., 2005). During the Cre-taceous-Eocene, the area of present-day southern Bolivia

was already part of a foreland system but was presumablyincluded in a large intracontinental plain of non-deposi-tion.

The geology of the Bolivian Andes is classi¢ed into sixtectonomorphic units, of which three units participate inthe late Cenozoic foreland system (Fig. 1). Sedimentaryunits pertaining to the Chaco basin occur (west to east)from the Inter-Andean Fault (IAT), through the Suban-dean Zone, and below the Chaco plain to its onlap on theBrazilian Shield and the Alto de Izozog basement high.The western part of this basin is deformed by the Suban-dean fold-and-thrust-belt and is still undergoing activeshortening at its leading edge (Fig. 2; Oller, 1986; She¡els,1988; Baby etal., 1992; He¤ rail et al., 1996). Late Cenozoic se-dimentary strata are commonly well exposed along £anks

Fig. 2. Geological and structural map of the study area (modi¢ed from Suarez-Soruco, 2000) showing the data set and localities ofmeasured sections mentioned in the text: 1, Abapo; 2,Tatarenda; 3, Saipuru; 4, Piriti; 5, SanAntonio; 6, Oquitas; 7, Choreti; 8, Itapu; 9,Ivoca; 10, Cuevo; 11, Boyuibe; 12, Iguamirante; 13,Machareti; 14, Angosto del Pilcomayo (Villamontes); 15, Puesto Salvacion; 16,Zapaterimbia; 17, Rancho Nuevo; 18, Sanadita; 19, SanTelmo; 20, Nogalitos; 21, Emborozu¤ .

r 2006 The Authors. Journal compilation r 2006 Blackwell Publishing Ltd,Basin Research, 18, 145^170 147

Evolution of the late Cenozoic Chaco foreland basin

of the major leading syn- and anticlines near the wedge-tip of the Subandean zone.This region, between �641300and 621300E and 181450 and 221300S, forms the principalstudy area.

Chaco foreland basin sedimentation is assumed to havebegun approximately 27Ma ago near the Eastern Cordil-lera (Sempere et al., 1990), as a result of eastward migrationof the deformation front ahead of the Interandean andSubandean Zones (Sempere etal., 1990;Husson&Moretti,2002; DeCelles & Horton, 2003; Echavarria et al., 2003;Ege, 2004). Late Cenozoic strata show characteristic west-ward thickening. It stands to reason that Chaco forelandbasin strata had also been deposited in considerable thick-ness in the region occupied by the present-daySubandeanZone before its uplift and incorporation into the eastward-migrating orogenic wedge. Although these deposits arewell preserved in the Subandean Zone, in some areas thecoarse-grained proximal foreland basin deposits havebeen eroded.The eroded proximal basin sedimentation isas a result of erosion removal after deformation and subse-quent propagation of the fold^thrust belt (Burbank&Ray-nolds, 1988). A rough estimate of their original thickness(ca. 3^5 km) can be obtained by reconstructing thermal-gradient-calibrated sedimentary thickness from AFTsamples of the youngestMesozoic strata in the SubandeanZone (Ege, 2004).

Magnitude and timing of shortening

Lithospheric thickening and corresponding shortening inthe fold-and-thrust belt of the Subandean zone, recon-structed from structural balanced cross-sections (e.g.Sempere et al., 1990; Kley et al., 1996, 1997), began east oftheEasternCordillera in the lateMiocene.However, wide-spread shortening there started only in the Oligocene(Baby et al., 1992; Gubbels et al., 1993; Dunn et al., 1995;Kley, 1996; Jordan et al., 1997; Kley et al., 1997; McQuarrie,2002). Since then, continuous eastward propagation ofthrusting, accompanied by large-scale folding, produceda generally eastward-younging synorogenicwedge (Moret-ti et al., 1996; DeCelles & Horton, 2003; Echavarria et al.,2003).

During the late Oligocene, the Eastern Cordillera wasthe focus of pronounced shortening (Kley etal., 1997;Hor-ton, 1998). Figure 3 shows structural styles and majorthrust sheets, illustrating that the Subandean Zone is de-formed by mostly in-sequence, thin-skinned thrustsheets that include north-northeast-trending ramp anti-clines and passive roof duplexes (Baby etal., 1992,1997; Be-lotti et al., 1995; Dunn et al., 1995; Kley et al., 1996, 1999;Echavarria et al., 2003).This progressive thin-skinned de-formation is recorded in a suite of angular unconformitiesand stratigraphically distinct foreland packages. A totalshortening of 210^336 km is postulated for the Central An-des (Baby et al., 1992;Moretti et al., 1996;McQuarrie &De-Celles, 2001;Mˇller et al., 2002; Elger et al., 2005) together.The Interandean and Subandean Zones take up 140 and86 km shortening at 201S and 221S, respectively (Baby etal., 1997). This matches well with a total shortening of�140 km at 211S in the Interandean and Subandean zonestogether (Kley et al., 1997).Moretti et al. (1996) calculated apeak shortening rate between 6 and 2.1Ma, followed by aminimum shortening rate between 2.1Ma and the presentin the Subandean Zone. Their values, however, disagreewith the estimates byEchavarria etal. (2003),who postulatetwo periods of high shortening rates (11 and 8mmyear�1)at 9^7 and 2^0Ma, respectively, separated by an in-be-tween low of 0^5mmyear�1 at 221300 latitude.These con-tradictions may be due to the paucity of direct age dates forthe deformation and the inherent variability of geologiccross-section construction and interpretation.

The variations in shortening values (Moretti et al., 1996;Echavarria etal., 2003) and the resulting inferred time per-iods of uplift (Sempere et al., 1990; Baby et al., 1992, 1997;Dunn et al., 1995, Kley et al., 1997) suggest diachronousmovement on individual thrust sheets. For example, Echa-varria et al. (2003) attribute the 2-Ma-shortening event tothrust reactivation in the south-western Subandean Zone( �221300), whereas Moretti et al. (1996) interpreted the2.1Ma shortening as a major displacement event synchro-nous with folding and uplift of the leading Aguaragˇerange (Fig. 2). No data are available to constrain the timeof formation of the ramp anticlines of the western Chacoplain, which are clearly visible on industry re£ection-seis-

Fig. 3. Structural balanced cross-section of the Central Andes fromAltiplano to Chaco plain at 211S.Modi¢ed after Kley et al. (1999)and Elger et al. (2005). UKFZ,Uyuni-Khenayani Fault Zone; SVT, SanVicenteThrust; CYT, Camarga-Yavi Thrust; IAT, InterandeanThrust; SAT, SubandeanThrust; PF, Pajonal Fault; PBF, Palos Blanco Fault. See Fig. 2 for location.


C. E. Ubaet al.

mic pro¢les but have as yet only an indistinct and low to-pographic expression. Anticline cores of these structuresare morphologically expressed in a N^S trend of low hills,cut by east^west-trending gullies.These structures appearto be actively forming and indicate the continuous east-ward growth of the Andes onto the Brazilian Shield.

DATA ANDMETHODS

The data set compiled for this study includes measuredsections, seismic data and well logs. Twenty-one strati-graphic sections along major rivers, small streams androad cuts in the Subandean foothills were measured andsampled for lithologic, sedimentologic and biostrati-graphic data (Fig. 2) to document architectural style andbasin geometry. In addition, we interpreted 45 wire-linelogs and their well reports from hydrocarbon industry ex-ploration wells and tied them to42800 km of 2-D indus-try seismic pro¢les. Wire-line logs (g-ray, resistivity andsonic), combined with well reports, provide ¢ne verticaldetails of wells and lithology resolution and thus comple-ment seismic data for a better understanding of thesubsurface geology. Similar methods were used by Schlu-negger etal. (1997) andAlves etal. (2003) to study theUpperMarineMolasseGroup of theNorthAlpine foreland basinand the Lusitanian rift basin of West Iberia, respectively.Wire-line log analysis and well reports were used in com-bination with seismic facies attributes to delineate di¡er-ent stratigraphic packages and to correlate them to the¢ve late Cenozoic formations.The seismic data, wire-linelogs andwell reports were provided byChaco S.A. and Ya-cimientos Petroleros Fiscales de Bolivia (YPFB), SantaCruz.

The seismic lines cover mostly the Chaco plain whereoutcrop is poor or absent, and partially extend into the

foothills of the Subandean Zone.They de¢ne the regionalstratigraphic architecture of the late Cenozoic basin ¢ll.We used regional isopach trends as a proxy for accommo-dation space (e.g.Wadworth et al., 2003), and traced theirthickness variations fromvertical facies associations. Syn-thetic seismogram and check-shots from well logs wereused to perform time-to-depth conversion from two-way-travel time (TWT, in ms).

FORELAND LITHOSTRATIGRAPHY

The up to 7.5-km-thick (Emborozu¤ section), eastward-thinning strata of the Chaco foreland basin are largelycomposed of siliciclastic non-marine redbeds with minorshallow-marine strata. We used a detailed stratigraphyafter Suarez Soruco (2000) that is principally based onlithology,with only minor modi¢cations (Fig.4).The basin¢ll includes (from base to top) the Petaca, Yecua,Tariquia,Guandacay and Emborozu¤ Formations. Age dating ofthese formations has proved di⁄cult and principally relieson a combination of mammal biostratigraphy and radio-metric dating of rare tu¡s (e.g. Marshall et al., 1993; Mar-shall & Sempere, 1991; Moretti et al., 1996; Echavarria etal., 2003; Hulka, 2005). Notwithstanding the recent bypublished new 40Ar/39Ar radiometric ages for the lateCenozoic units in the Bolivian Subandean zone by Hulka(2005), no complete and precise chronology for the basin¢ll is yet available. Therefore, we used published ages todocument the late Cenozoic lithostratigraphy of thesouthernBolivia.However, the ages should be appliedwithcaution. Parts of the formations are suspected to be dia-chronous, not only younging west-to-east, as could be ex-pected, but possibly also north-to-south (Echavarria etal.,2003). In addition, the stratigraphy is complicated by sev-eral nearly basinwide low-angle unconformities.

Fig.4. Stratigraphy of the Subandean Zone and Chaco basin ¢ll and rose diagrams summarize the palaeocurrent directions. Ages arebased onMarshall et al. (1993), Moretti et al. (1996), Echavarria et al. (2003), Hulka (2005), andHulka et al. (in press).



Petaca formation

Cenozoic sedimentation in the Chaco basin commencedduring the late Oligocene (assumed ca. 27Ma; Marshallet al., 1993; Moretti et al., 1996) with the deposition of the

up to 250-m-thick Petaca Formation (Gubbels et al., 1993;Sempere, 2000). This formation unconformably overliesMesozoic eolian strata (Sempere, 1995).The lower part ofthe Petaca Formation consists of greenish grey, white andlight purple basal calcrete.The calcrete consist of isolated

Fig. 5. Selected outcrop photographs showing (a) clast- supported reworked pedogenic conglomerate facies of the Petaca Fm (seehammer in circle for scale). (b) Shallow-marine-lacustrine mudstone-dominated facies with thin-bedded ooid-, shell hash-dominatedsandstone bed of theYecua Fm (arrow). (c) Channelized sandstone beds with desiccation cracks (arrow) of theTariquia Fm. (d)Alternation of conglomerate and sandstone beds of the Guandacay Fm. (e) Sheet-like cobble-boulder-dominated conglomerate bedwith thin beds of sandstone of the Emborozu Fm.


C. E. Ubaet al.

to clustered, blocky to massive and bracciated nodules(Uba et al., 2005).The calcrete body is overlain byhorizon-tal to disorganized clast- supported reworked-pedogenicconglomerate (Fig. 5a) composing of poorly sorted,denselypacked clasts of poorly rounded intraformational re-worked calcrete nodules and subordinate chert. The con-glomerates show sharp and erosive bases. Medium tovery-coarse-grained sandstone and sandy to ¢ne-grainedmudstone mark the upsection lithology of the Petaca For-mation.The calcareous, red to grey, bioturbated sandstoneis characterized by tabular to lenticular beds, troughcross-, planar and horizontal strati¢cations, as well asrip-up clasts at the base. The massive, laminated mud-

stone bodies have bioturbation, minor desiccation cracksand padogenesis. The formation thins towards the centreof the study area (Villamontes-Camiri axis, Fig. 6).The re-worked pedogenic conglomerate and sandstone bodiesshow channel and bedform architectural elements and anoverall ¢ning-upward sequence. Cross-strati¢cation insandstones indicates awestward-directed drainage (Fig.4).

Uba etal. (2005) attributed the thick calcrete horizons towell-developed palaeosols, indicating 0 or low sedimenta-tion in an arid to semiarid climate in which evaporationgenerally exceeded precipitation (e.g. Cecil, 1990). Thelithofacies and architectural elements in the Petaca con-glomerate and sandstone indicate variable high-energy

Fig. 6. Correlated pro¢les of stratigraphic sections of the late Cenozoic strata in the northern study area �191S. See Fig. 2 for location.



stream £ows in a channelized setting. Uba et al. (2005) andMarshall et al. (1993) interpret the Petaca strata as havingbeen deposited by braided stream. The ¢ning-upwardtrend and changes in bedform represent a decrease in £owstrength or depth as a result of waning of £ood intensity(Miall, 1996).The occasional occurrence of successions ofpalaeosols indicates predominantly non-deposition andsurface exposure.This is supported by desiccation marks,bioturbation and purple colour (e.g.Miall,1996; Retallack,1997).The contact between the Petaca Formation and un-derlying eolian strata is a regional erosional unconformitythat may have formed as a far- ¢eld response to early An-dean tectonics (Sempere etal.,1990;Dunn etal.,1995).Mar-shall et al. (1993) reported reptilian and mammal bone

fragments of late Oligocene to late Miocene age, foundclose to the Aguarague range in conglomerate (Sempere etal.,1990;Marshall &Sempere,1991).However, as the age ofthe basal calcretes has not been ascertained, the onset ofdeposition is poorly constrained.

Yecua formation

The up to 600-m-thickYecua Formation (Padula &Reyes,1958) overlies the Petaca Formationwith an indistinct low-angle erosional unconformity.TheYecua Formation showsa west-to-east and northeast-to-southwest facies varia-tion. North of Camiri, it consists of red-green to brownsandstone^mudstone couplets (Fig. 5b) showing herring-

Fig.7. Correlated pro¢les of stratigraphic sections of the late Cenozoic rocks in the southern study area �211S. See Fig. 2 for location.


C. E. Ubaet al.

bone cross-strati¢cation, laminated, convolute, £aser,wavy and lenticular bedding in ¢ning- and coarsening-upward successions.This lithofacies also consists of gyp-sum veins, syndepositional structures, bioturbation anddesiccation marks (see Fig. 5b). Fossils include bivalves,the foraminifera Globigerinacea and Corbicula, the ostra-codes genera Cypridelis and Heterocypris, pelecypods,gastropods, cirripeds, decapods, crabs, ¢sh skeleton frag-ments, ooids, shell hash and terrestrial plants (Marshall &Sempere, 1991; Marshall et al., 1993; Hulka et al., in press).In the western and southern part of the study area (southof Camiri), the Yecua Formation consists of red to lightbrown, lenticular andvery ¢ne- to medium-grained sand-stone interbedded with red to light-brown, ripple-lami-nated sandstone couplets. The proportion of mudstonebodies dominate over the sandstone (Fig. 5b). The sand-stones show erosive channel structure and ¢ning-upwardtrends.These sand bodies contain cross-bedding, climb-ing ripples, gypsum veins, rip-up clasts and burrows.Mudstone^sandstone couplets contain mottled soil, de-siccation cracks and extensive burrows. In general, thesandstone proportion, bed thickness and the channel pro-portion of theYecua Formation increase upsection and to-wards the west (Figs 6 and 7).

We interpret the fossiliferous and varicoloured Yecuafacies in the north of the study area as deposits of lacus-trine, tidal, shoreline and brackish to shallow marine en-vironments, in agreement with the previous work byMarshall et al. (1993), Hulka et al. (in press) and Uba et al.(2005), and is supported by the presence of lacustrine-shallow marine fossils and the lithofacies.The mudstone-dominated terrestrial facies of the Yecua Formation to thewest and south are products of £uvial overbank and chan-nel processes, with occasional lacustrine and mud£at set-tings. Hulka et al. (in press) placed these variations in aregional context and argued that the marginal marine fa-cies of the Yecua Formation represented a marine incur-sion from the northeast along the axis of the developingforeland basin as far as Camiri.The age of theYecua stratahas variably been estimated based on ostracodes and fora-minifera to be �14^7Ma (Padula & Reyes, 1958; Marshalletal., 1993;Hulka etal., in press) and11^7Ma (Moretti etal.,1996). Recently published 40Ar/39Ar radiometric ages of10.49 � 0.33 and 9.41 � 0.52Ma (Hulka, 2005) on inter-bedded tu¡s in the Yecua Formation £uvial facies fromthe Emborozu¤ and Nogalitos sections matches the esti-mated biostratigraphic age from the marine facies. By ana-logy, these ages are considered herein to correlate with the£uvial-and-lacustrine Yecua-equivalent strata (TariquiaFormation of Bolivian nomenclature; Russo, 1959; Ayaviri,1964; Moretti et al., 1996; Suarez Soruco, 2000) near Ar-gentina’s border with Bolivia that yielded an age of9.95 � 0.34Ma (Echavarria et al., 2003).

Tariquia formation

TheTariquiaFormation (Russo,1959;Ayaviri,1964) is up to3800-m-thick and overlies theYecua Formation with gra-

dational contact.TheTariquia Formation is characterizedby thick- and thin-bedded sandstone bodies interbeddedby laminated mudstone and very ¢ne-grained sandstone(Uba et al., 2005). The light brown, light yellow and red,well- sorted, very ¢ne- to medium-grained sandstonebodies range between 0.5 and 15m thickness and consistof sharp erosional base, ribbon and channel geometry(Fig. 5c), and extend laterally for hundreds of meters.Sandstone units have massive bedding, planar, troughcross- and climbing ripple sedimentary structures. Intra-formational rip-up clasts and reworked calcareous no-dules are common. The sandstone bodies have multi-storey channel architecture and an overall coarsening-and thickening-upward trend (Fig. 5d).There is an upwardincrease in the degree of vertical stacking, bed thicknessand lateral interconnectedness in the sandstone unit.Overall, the mean grain size, channel interconnectedness,sandstone proportion and thickness of the TariquiaFormation increase towards the west (Figs 6 and 7). Themassive, laminated- or ripple-strati¢ed interbeddedmudstone^sandstone couplets show sheet geometry andare laterally extensive.Taenidium barreti trace fossils (Bua-tois et al., in press) in the thick-bedded channelized sand-stone and in mudstone and sandstone couplets arecommon.The trace fossils disrupt the primary sedimen-tary structures. A distinguishing feature of the TariquiaFormation is the presence of abundant mudcracks (Fig.5c, arrow), occasional syndepositional deformation, andpoorly developed palaeosols that are more dominant inthe mudstone^sandstone couplets. Palaeocurrent mea-surements indicate a mean transport towards the east(Fig. 4). In theTariquia Formation, the sandstone propor-tion and size and bed thickness increase towards the west(Figs 6 and 7).

The Tariquia Formation is interpreted to represent arange of processes that operate in a large £uvial system(Uba et al., 2005). Thick-bedded sandstones were depos-ited within major channels, whereas the thin-beddedsandstone units indicate deposits from crevasse channels.In overbank areas,mudstone and sheet sandstonewere de-posited by crevassing splay and suspension fallout. Follow-ingUba etal. (2005), we interpret theTariquia Formation asa product of a low-gradient, high-sedimentation, channe-lized anastomosing stream and associated thick £ood-plains on a distal £uvial megafan. This interpretation issupported by the laterally extensive channel geometry, ag-grading thick £oodplain deposit (ponded area), verticalchannel stacking, frequent crevassing and avulsion, and ageneral lack of lateral channel migration architecture (e.g.Smith, 1986; Makaske et al., 2002; Bridge, 2003). The in-ferred depositional processes and lithofacies are similarto active modern meganfans in the Chaco plain that re-ceive sediment from large £uvial networks (Horton & De-Celles, 2001). Avulsion and crevassing result in thedevelopment of new channels on the overbank, whereasthe active channels are abandoned (Smith, 1986; Makaskeetal., 2002).The abundant rip-up clasts that mayhave beenformed by erosional scouring of overbank sediments and



high degree of bioturbation in theTariquiaFormation sug-gest long periods of channel abandonment and coloniza-tion by insects (Buatois et al., in press). The well-developed upward coarsening and thickening trend sug-gests a systematic stratigraphic development governed byeither long-term eastward propagation of the fold^thrust-belt and/or the expansion of drainage networks. Uba et al.(2005) postulated a shift in climate from a semi-arid to ahumid condition during the deposition of the Tariquiastrata. The Tariquia Formation age is late Miocene(Chasicoan-Huaquerian) based on biostratigraphy (Mar-shall & Sempere, 1991), in agreement with a singleapatite ¢ssion-track age of 7Ma fromMica (Moretti et al.,1996). In addition, Moretti et al. (1996) assumed 6Ma asthe upper age limit of theTariquia Formation. In the ab-sence of a well-constrained age for this unit, we use theimprecise age of 7^6Ma for the deposition duration forthis formation.

Guandacay formation

The up-to-1500-m-thick Guandacay Formation consistsof conglomerate, sandstone and mudstone (Jimenez-Mir-anda&Lopez-Murillo,1971) (Fig. 5d).The granule-cobbleconglomerate shows sheet-like and lenticular geometry,clast-supported, polymictic, a coarsening- and thicken-ing-upward trend, massive to inversely graded, well-de-veloped imbrication and basal scour surfaces. Gravelbedforms and poorly developed lateral accretion surfacesare common architectural elements. The dominantlymedium- to very-coarse-grained sheet-like sandstonesare moderately to well sorted, and are laterally extensivefor several hundreds of metres (Fig. 5d). The sandstonebodies consist of trough cross-, planar, ripple and horizon-tal strati¢cation, and occasional stringers of pebbles.Thick interbedded mudstones and sandstone are massiveto laminated and laterally continuous for several tens orhundreds of metres.Lenses of thin coal seams, poorly pre-served bioturbation, and weakly developed mottled soilsare present (Uba et al., 2005).The conglomerates generallythicken and coarsen upsection and to the west. The con-glomerate and sandstone bodies show, like the TariquiaFormation, an upward increase in stacked packages, lateralinterconnectedness, and multi- to single-storey channelsystems that grade into the interbedded mudstone andsandstone. Palaeocurrent measurements indicate a north-east-to-southeast-directed £ow (Fig. 4).

The conglomerate and sandstone lithofacies provideevidence of deposition in £uctuating, high-energy, bed-load-dominated large £uvial channels, £anked by £ood-plains, and zones of incipient soil development (Uba etal., 2005). The dimensions of channel ¢lls and the typesof sedimentary structures in the Guandacay Formationsuggest large discharges (e.g. Horton & DeCelles, 2001).Consequently,Uba etal. (2005) envision a proximal braidedsetting on a medial £uvial megafan, similar to those thatdeposited the Camargo Formation and that drain themodern central Andean (Horton & DeCelles, 2001; De-

Celles & Horton, 2003). Lenses of coal suggest the pre-sence of a ponded area and vegetation, and therefore, ahumid palaeoclimate (Uba et al., 2005). Vertical stackingand aggradation of channels into overbank deposits implycrevassing and avulsion, indicating periodic abandonmentof active channels. The weakly developed palaeosol andpoorly preserved bioturbation may suggest a high over-bank aggradation rate (Bridge, 2003).The contact betweenthe Tariquia and the overlying Guandacay Formation isunconformable (Moretti et al., 1996; Echavarria et al.,2003), approximately 6Ma in age (Moretti et al., 1996), andis marked by a distinct increase in mean grain size. Hulka(2005) estimated the top of this formation at 2.1 � 0.2Mabased on 40Ar/39Ar dating of tu¡ at its contact to the Em-borozu¤ Formation in the Abapo¤ section (Fig. 2).The age ofthe Guandacay Formation is therefore late Miocene toEarly Pliocene (6^2.1Ma).

Emborozu¤ formation

The Emborozu¤ Formation (Ayaviri, 1967) is exposed onlyin the northeast (Abapo¤ Section) and within synclines inthe southwestern (Emborozu¤ and Nogalitos; Fig. 2) studyarea. The up-to-2000-m-thick, conglomeratic upward-coarsening strata of this formation cap the foreland strati-graphic succession in the Chaco Basin. In outcrops nearthe present Subandean topographic front, growth struc-tures occur (Echavarria et al., 2003), documenting asyndeformational origin. The Emborozu¤ Formation isdominated by an up-to-60-m-thick, cobble-boulder con-glomerate that reaches at least 153 cm in diameter (Fig. 5e;Uba et al., 2005).This laterally extensive (several hundredsof metres) conglomerate shows sharp erosive scoursurfaces, sheet-like to lenticular single-channel geometry,inverse and normal grading, and moderately to poorlydeveloped imbrications. This conglomerate lithofaces isassociated with up-to 6-m-thick, coarse- to very-coarse-grained, sheet-like sandstone with horizontal, troughcross-, ripple- and planar strati¢cation. The single-storey, vertically stacked conglomerate and sandstonebodies grade into medium- to very coarse-grained,rippled, massive-laminated interbedded sandstone andmudstone in which poorly preserved burrows and coallenses occur.Upsection and to the west, the thickness, lat-eral continuity, amalgamation and maximum grain size ofthe conglomerate and sand bodies increase and the per-centage of overbank ¢nes decreases. The palaeodrainagepattern shows a northeast-to-southeast-directed £ow(Fig. 4).

Uba et al. (2005) interpret the Emborozu¤ Formation as a£uctuating-energy, bedload, proximal £uvial system ofsuccessions of large, isolated to amalgamated channels.The presence of a thick to subordinate £oodplain and thelateral extent suggests deposits on a proximal £uvial mega-fan (Horton &DeCelles, 2001; Uba et al., 2005).The sharpscour surfaces may represent discrete channels or an amal-gamation of scour as a result of avulsion events and chan-nel abandonment. The Emborozu¤ Formation overlies the


C. E. Ubaet al.

Guandacay Formation with a well-de¢ned regional angu-lar unconformity that marks its base in seismic sections(Moretti et al., 1996; Echavarria et al., 2003). Moretti et al.(1996) previously estimated the age of this contact basedon ArÂr on mica to be 3.3Ma. However, a new2.1 � 0.2Ma (tu¡; Abapo¤ section; ArÂr on mica) estimateby Hulka (2005) agrees relatively well with 1.8Ma docu-mented by Echavarria et al. (2003) for correlative strata inArgentina. Consequently, 2.1� 0.2Ma is used herein asthe basal age of the formation.

SEISMIC STRATIGRAPHYOF THE LATECENOZOIC DEPOSITS

A good well-to-seismic tie and the lateral continuity ofhorizons allowed interpretation of the visible geometricfeatures on the seismic lines. After interpreting seismiclines andwire-line logs, we subdivided the foreland-basin¢ll into ¢ve regionally mappable packages, numbered se-quentially N1 to N5.These are delineated by discontinu-ities that coincide with changes in seismic facies and thatcan be correlated with wire-line logs. Seismic facies attri-butes include prominent re£ectors, termination geometry(onlap, toplap, downlap and truncation), re£ection con¢g-uration, and external form. Not all onlap and truncationgeometries could be mapped in the seismic sections dueto limited vertical resolution combined with small unitthickness. Figures 8 and 9 illustrate the most characteristicseismic facies features and g-ray (GR), resistivity (ILD)and sonic (DT) log responses of the ¢ve packages.The cor-responding lithofacies and depositional environments arecalibrated by well data.

Package N1

The base ofN1is a prominent, readily traceable re£ector ofhigh amplitude,medium frequency and medium continu-ity across most of the study area, marking the contact be-tween the late Cenozoic foreland basin and underlyingMesozoic strata (Fig. 8a). In some seismic sections, the un-derlying Mesozoic strata show di¡use toplap and trunca-tions with low angular geometry. In wire-line logs, thiscontact shows an abrupt increase from �100 to 170 Omin ILD and an immediate drop from �90 to 40 in DTre-sponses (e.g. Fig. 8a). A clear di¡erentiation can be madebetween the top of N1 and the base of N2 as a result of apronounced medium- to high-amplitude, continuousand medium- to high-frequency re£ector that is easilyidenti¢ed and correlated throughout all seismic sections(Fig. 10).Wire-line logs show a sharp increase from �30to 120 API in the GR curve and from �50 to 70 in theDT curve, coupled with an abrupt increase to �90 in theILD curve.The thinness of this package does not allow adetailed seismic facies characterization. However, in someareas, N1displays internally lateral extensive, low- to med-ium-amplitude, subparallel, discontinuous, low-fre-quency re£ectors (Fig. 10). Among them, a wedge-shapedset reaches up to ca. 200m ( �0.2ms) thickness across abroad area in the western Chaco plain and graduallypinches out to 0 with onlap terminations upon reachingthe Alto de Izozog high (Fig. 10).Wire-line logs throughN1 (Fig. 11) generally indicate low GR value (30^60 API),average 90^55 DT, and 90^170 Om ILD values. The GRcurves show cylindrical shape characteristic.

The toplap re£ection terminations and truncation ofN1 on the underlying Mesozoic strata indicate an uncon-

Fig. 8. Major package boundaries and their characteristics recognized on seismic sections, well logs, and interpreted lithology.Thebase and top of each package is de¢ned.



formity. It could also be a condensed or non-depositionsurface with minor erosion (e.g.Mitchum etal., 1977; Sher-i¡ & Geldart, 1995). However, the quality of the seismicdata and the overall low thickness ofN1do not allow a cleardi¡erentiation between toplap and erosional truncation.Distal onlap and the overall wedge form of N1on the Altode Izozog High indicate basin progradation.We interpretpackageN1as a sand-dominated aggradational £uvial sys-tem.This interpretation is based on seismic facies charac-teristic (variable-amplitude, discontinuous and low-

frequency re£ection) combined with low GR and DTandthe cylindrical shape of the GR curve, implying that thispackage consists mainly of relatively high-energy £uvialdeposits (Badley, 1985; Cant, 1992; Emery & Myers, 1996).The cylindrical shape of GR logs suggests an aggradingbraided £uvial system (Cant, 1992; Emery & Myers, 1996).The high acoustic impedance variation between the over-lying mudstone-dominated N2 and the underlying, sand-stone-dominated Mesozoic rocks also suggests a changein lithology and probably a high degree of cementation or

Fig.9. g-ray, resistivity, and sonic records for IGR 01well and the interpreted lithology correlated to the Angosto del Pilcomayo sectionlocated approximately 40 km farther south.


C. E. Ubaet al.

pedogenesis. We interpret the strong seismic facies andsharp wire-line log relations at the contact of N1 and N2to re£ect an unconformity.We correlate N1 based on theseismic facies andwire-line log characteristics mentionedabove as the subsurface equivalent of the Petaca Forma-tion.

Package N2

The top of N2 is a laterally continuous, high-to-moder-ate-amplitude re£ector (Fig. 8a). The GR, ILD and DTlogs do not show a sharp di¡erence but rather a gradual re-sponse at the contact to N3 (Fig. 8a). N2 is overall wedge-shaped, with a maximum thickness of more than �450m(0.20ms) in the west.To the east,N2 terminates with onlapgeometries on the Alto de Izozog High, where it overliesN1 and pinches out onto the Mesozoic strata (Figs 10 and11). Its internal seismic facies are: to the west, N2 displaysinternal seismic re£ections that show variable-amplitude,discontinuous, subparallel, low-frequency, low verticalspacing and chaotic pattern. Low-scale hummocky clino-forms dip at variable angles ( �1^21); to the east, low-scalecomplex sigmoid-oblique seismic re£ections also occur.In contrast, to the east, low-angled clinoforms, discontin-uous, low-amplitude, semi transparent and chaotic re£ec-tions, coupledwith low acoustic impedance, occur (Figs 8aand 10).Wire-line logs show �60^120 API in GR, 95^85Om in ILD and 65^85 in DTvalues. However, GR and

DTvalues decrease and increase upsection, respectively.GR indicate marked thin spikes and large percent of highto low values (80 : 20) in the N2 package (Fig. 9). However,the low GR and high DTvalues increase upsection. GRlogs show an irregular or serrated response (Cant, 1992;Emery &Myers, 1996) and small- scale variability in valuesas indicated by numerous thin cycles with ¢ning-upwardtrends (Fig.11).

The variable-amplitude, discontinuous, semi-trans-parent, internal structure, higher GR and lower DTvaluesare typical of a poorly strati¢ed mudstone-dominated sys-tem, deposited mainly by suspension settling and subordi-nate channel settings (Cant, 1992; Alves et al., 2003). Basedon the seismic facies and well-log characteristics, we caninterpret the N2 package as deposition in varied settingssuch as shallow marine, lacustrine and £uvial environ-ments (e.g. Badley, 1985; Cant, 1992; Alves et al., 2003;Hofmann et al., 2006), with aggrading £uvial settingdominantly in thewest and south of the study area.The in-terpretation of varied-depositional settings is furthersupported by thin sandstone intervals in wire-line logs, arelativelyhigh and serratedGRresponse, sigmoid-obliqueand hummocky clinoforms, varied-amplitude and low fre-quency (Sangree&Widmier,1977;Badley,1985;Cant,1992;Sheri¡ & Geldart, 1995). As N2 thickens westward towardsits depocentre, it develops varied-amplitude and low-continuity re£ections and de¢ned-clinoforms. Thesmall- scale variability observed in the GR, DTand ILD

Fig.10. Segment of aW-E uninterpreted and interpreted migrated seismic line along approximately 201450S showing the ¢ve lateCenozoic sequences and thrusting and folding of the foreland sequences.The location of the line is shown in Fig. 2.



pro¢les largely represents variation in depositional energyassociated with high-frequency cyclicity (Cant, 1992; Em-ery & Myers, 1996).The gradual GR and DTresponses atN2^N3 contact suggest a fairly steady change of deposi-tional environment between these two packages.The east-ward thinning indicates the presence of a palaeo-high nearthe eastern border of the study area before the depositionof N2.The upsection decrease in the GR and increase inthe DT values, which can imply upsection increase insandstone proportion, re£ects a basinwide shift in facies.The characteristic mudstone-dominated seismic faciesand wire-line log attributes of N2 package are analogousto theYecua Formation.

Package N3

The base ofN3 is a variable continuous and varied-ampli-tude re£ector; theGR andDTshowdecrease and decreasein values, respectively (Figs 8 and 9).The laterally continu-ous, moderate- to high-amplitude re£ector marks the topof this package and a transitional contact to N4. In wire-line logs, this contact is marked by a relatively sharp lowGR and high ILD and DTresponses (Figs 8b and 9). N3shows a maximum thickness of �1500m ( �0.75ms) inthe western study area, thinning gradually eastward to

pinch out at the Alto de Izozog basement high, where itonlaps and overlies Mesozoic strata (Fig. 10). Internally,N3 displays varied-seismic facies; in the lower portion ofthe section, it shows low- to medium-amplitude, discon-tinuous, subparallel, low-frequency, semi-transparentand hummocky re£ectors (Figs 10 and 12). However, theseismic facies changes upward to more moderate-low con-tinuous, varied-amplitude, less chaotic and less hum-mocky re£ectors and wedge-sheet external forms. Theproportion of clinoform, hummocky, low-frequency re-£ectors increases to the east. The clinoforms show east-oriented downlap onto, and appear to coalesce with, med-ium-amplitude re£ectors. The log character of the N3package is distinguished from the underlying N2 packagebecause it contains relatively lowerGR(30^90API), higherDT (70^100) and higher ILD (95^160) responses. In addi-tion, it shows a thicker and larger percent of a low GR re-sponse compared with the N2 package, with the percentand thickness of lowGR andDTresponses increasing sig-ni¢cantly upward and to the west, where it reaches tens ofmetres in thickness (Figs 9, 12 and14). GR and DT curveshave both serrated and occasional bell shapes.

We interpret the varied-amplitude, subparallel, moder-ately to discontinuous seismic facies, coupledwith lowGR

Fig.11. Segment of aW-E uninterpreted and interpreted migrated seismic line along approximately191300S showing pinch out of thelate Cenozoic sequences.The location of the line shown is in Fig. 2.


C. E. Ubaet al.

and high DT response of package N3, as alternations ofsandstone and mudstone deposited by sandstone-domi-nated £uvial deposition (e.g. Sangree & Widmier, 1977;Cant, 1992; Sheri¡ & Geldart, 1995; Emery & Myers,1996).The low to high amplitude, moderate^low continu-ity, lens-sheet external forms, and serrated-bell shape GRand DTresponses suggest channels aggrading into £ood-plains (Cant, 1992). This interpretation is further sup-ported by the upsection change in seismic facies (e.g.chaotic, hummocky, semi-transparent, combinedwith up-section decrease in theGRvalues and thickness of lowGRresponse) that indicate an upsection increase in the pro-portion of sandbodies and bed thickness. The serrated-and bell- shaped GR responses suggest multiple ¢ning-upward trends and variable depositional energy. Thesemi-transparent and chaotic seismic features representlack of strati¢cation. The westward-thickening wedge-shaped geometry of N3, the eastwardly oriented clino-forms, and the wedge form suggest deposition by pro-

gradation from the west.The upsection increase in clino-forms re£ectors, variable seismic characteristics and pro-portion of sandstone and GR thickness within N3 suggesta strong progradational pulse concomitant with a basin-ward shift in facies and depocentre location. We assigntheN3package to theTariquia Formation because of inter-pretation of characteristic of the seismic facies expressionsandwire-line logs identi¢ed in this package.

Package N4

The top of package N4 is characterized by a prominent,high-amplitude, continuous re£ector that can be mappedand correlated throughout all seismic sections. This topcontact is marked by local toplap and truncation termina-tions of N4 re£ectors on N5 (Figs 8d and 13, inset photo),accompanied by abrupt breaks on GR, DTand ILD logs(Figs 9, 12 and14). Figures 8d and13 show that this contactis a well-de¢ned angular unconformity.The base of N4 is

Fig.12. W-Ewire-line logs at about 211S illustrating aspects of sequence boundaries.The depositional sequences identi¢ed can becorrelated to near outcrops. See Fig. 2 for location.



delimited by a moderate to high-amplitude, continuousre£ector.Wire-line logs indicate a decrease and an increasein the values of GR and DT, respectively. The thicknessvariation of Package N4 is similar to that of N3, with abroad area in the west, where the thickness exceeds ap-proximately1500m ( �750ms) thinning to the east to zerolat the Alto de Izozog high. At this basement high, N4 alsoonlaps and overlies N2, N3 and Mesozoic strata (Fig. 11).Internally, N4 shows parallel to subparallel, variable-am-plitude and -frequency, and moderate^low continuity re-£ectors (Figs 8c, 9, 10 and 13). Package N4 shows a wedge-shaped, vertically spaced re£ections, suggesting severaltens of metre-scale bedding. In the southern part of thestudy area,N4 includes a growth structure near theLaVer-tiente Fault (Fig.12), showing an upsection decrease in theinclination of re£ectors and onlap geometry (Fig.12 inset).However, this fault is limited to the southern part of thestudy area and is not recognized further north (e.g. Fig.10). The N4 package is identi¢ed in wire-line logs by lowGR (30^70 API), high DT (105^140) and high ILD (95^130) responses (Figs 12 and 14). Figure 9 show that boththe GR and SP log curves show an upsection decreaseand increase in response, respectively, and have cylindrical

and bell shapes (Figs 9, 12 and14). However, the percent oflow GR values in N4 are relatively higher and thicker thanin the underlying N3 package (Figs 9 and 14).The low GRandDT intervals have a serrated shape.

The high GR and DT values, cylindrical shape andthickness, combined with parallel- subparallel, varied-amplitude, moderate^low continuous and wedge-shapedseismic facies suggest a thick intercalation of sandstonewith a conglomerate-dominated £uvial environment,probably in a braided setting (Cant, 1992; Emery &Myers,1996).The upward increase in the highGR log response atthe base of each cylindrical- or bell- shaped unit coupledwith its heterogeneity and vertical spacing in seismic linesmay indicate conglomerate lithofacies.We interpret the in-tercalated-thin-serratedGR andDTresponse as alternat-ing sand- and mudstone-bodies of overbank deposits(Cant, 1992; Emery & Myers, 1996). The overall decreasein GR and increase in DT curves and the upsection in-crease in thickness indicate a thickening- and coarsen-ing-upward trend.The seismic facies attributes and wire-line log characteristics of package N4 described above areanalogous to the sandstone-conglomerate-dominatedGuandacay Formation.The top of N4 marks a local angu-

Fig.13. Segment of aW-E uninterpreted and interpreted migrated seismic line along approximately 211S showing the structural stylesand the ¢ve sequences.The location of the line is shown in Fig. 2.


C. E. Ubaet al.

lar unconformity.Thewestward thickening, wedge geome-try of the package suggests that the depocentre is locatedto the west.The pinch out and onlap of N4 at the Alto deIzozog High indicate that of this basement palaeohigh ex-isted before deposition of N4.The onlap and thinning ofthe N4 re£ectors on the La Vertiente Fault are attributedto syndepositional deformation. The pronounced top re-£ector implies a strong acoustic impedance contrast andlikely represents the erosional surface or an angular regio-nal unconformity (Fig.12, inset; Dunn et al., 1995; Moretti

et al., 1996; Horton & DeCelles, 1997; Echavarria et al.,2003).

Package N5

The base ofN5 is de¢ned by a pronounced thick high-am-plitude, continuous, high-frequency re£ector, with onlapson the underlying N4 package.The top of the N5 packageis not well de¢ned and consists of medium- to variable-amplitude and moderately continuous, high frequency

Fig.14. W-Ewire-line logs at about191S illustrating aspects of sequence boundaries.The depositional sequences identi¢ed can becorrelated to near outcrops. See Fig. 2 for location.



re£ectors. However, well-log data are not available forthe top of N5.This package shows a maximum thicknessof �2000m (�1.0ms) in the west and thins eastwardto �0^200m (�0.1ms) at the Alto de Izozog high, whereit occasionally shows a progressive onlap on N4. Theinternal seismic facies includes laterally extensive, sub- toparallel re£ectors ofvariable amplitude, low continuity andlow frequency (Figs 8d, 9 and 12).The external forms pre-sent are sheet andwedge geometries.The seismic sectionsshow a large vertical spacing of re£ectors (Figs 8d and 13).Growth structures above the tips of major fault-propaga-tion folds display characteristic fan-shaped re£ectors withupward-decreasing dip. Only one of these faults (Man-deyepecuaFault) can be mapped throughout the study area(Figs10 and13).Nowire-line data are available to ascertainthe N5 lithological and sedimentological characteristics.

Notwithstanding the absence of wire-line logs, we in-terpret the N5 package to consist of an alternation of con-glomerate and sandstone deposited in a conglomerate-dominated £uvial setting (e.g. Cant, 1992), probably a largealluvial fan. This interpretation is supported by the sub-parallel-to-parallel, variable-amplitude, and low continu-ous, sheet- to wedge-shaped, and high vertical spacing ofre£ectors, which probably suggest thick conglomeratelithofacies. The overall thickening- and coarsening-up-ward trend, wedge-shaped external form, and pronouncedupsection increase in vertical spacing in seismic sectionare interpreted as a continuing basinward shift in grainsize.The seismic facies expressions and alluvial setting in-terpretation for theN5package permit its correlationwiththe surface Emborozu¤ Formation.

DISCUSSION

Overall stratigraphic pattern

In constructing the isopach maps (Fig. 15), we compiledthickness information for each formation derived fromoutcrop, depth-converted seismic, wire-line logs andwellreport data.The thickness values are not corrected for thee¡ect of compaction due to limited postdepositional bur-ial.We constructed isopach maps for ¢ve time periods (oc-casionally poorly) constrained by age estimates for theformations (Marshall & Sempere, 1991; Marshall et al.,1993; Moretti et al., 1996; Echavarria et al., 2003; Hulka,2005; Hulka et al., in press). We take thickness variations(Fig.15) through time as a proxy for available accommoda-tion space to infer that the Chaco foreland basin experi-enced variations in creation of accommodation spacesince the late Oligocene (e.g.Wadworth et al., 2003). Theearly basin history is illustrated in a single map (Fig. 15a)spanning more than 13Ma. A second map (Fig. 15b) spansa 7Ma- time period. In contrast, the ¢nal three time inter-vals represent only 3^1Ma each (Fig. 15cê). As expected,these ¢ve maps show distinctive thickness patterns.

The isopachmap of the 27^14Ma-oldPetacaFormation(Fig. 15a) shows a regional and broad area along the Villa-montes-Camiri axis witho50m of strata, possibly re£ect-

ing a structural high.This ‘ridge’ is £anked by up to 250mof strata cratonward and4100m of strata orogenward, re-spectively. Further east, the Petaca thins to 0 at the Alto deIzozog structural high. This ¢nding updates a previousview expressing a lack of thickness variations for this for-mation (e.g. Gubbels et al., 1993; Moretti et al., 1996). Thevery low available accommodation space during this peri-odwas probably a result of a long time interval of basin sta-bility and non- to low subsidence in this distal part of thebasin, augmented by scarce sediment supply inferred fromsuccessions of palaeosols (Gubbels etal.,1993;Horton etal.,2001).

Gubbels et al. (1993) and Moretti et al. (1996) reported asimilar lack of thickness variation for theYecuaFormation.In contrast, Fig.15b shows a distinctive westward thicken-ing, reaching a maximum thickness of �600m in outcrop(e.g. Emborozu¤ section).This westward thickening and in-crease in sandstone proportion of the Yecua strata is alsodocumented in seismic sections (Figs 6, 7, 10 and13). Dur-ing 14^7Ma, £exural foreland basin subsidence as a resultof thrusting episode that was probably centred in the pre-sent-day Interandean- or Subandean Zone (Coudert et al.,1995;Moretti etal.,1996; Echavarria etal., 2003) led to crea-tion of accommodation space in thewestern part of the ba-sin. Flemings& Jordan (1989) and Sinclair etal. (1991) showthat thrusting event results in an increase of the ratiobetween tectonic subsidence and sediment £ux. Thedominantly ¢ne-grained sediments in the more distaland relatively sandy facies in the west suggest that the de-positional slope was probably too low to produce largecoarse-grained deposits or there was a long lag-time be-tween erosion and more coarse-grained sedimentation asdocumented in other foreland basins (e.g. Blair & Bilo-deau, 1988; Jones et al., 2004).The ¢rst appearance of oro-genward increases in thickness and sandstone proportionindicates a change in locus of deposition.

The three subsequent isopach maps (7^6, 6^2.1and 2.1^0; Fig. 15cê) show a similar and regular westward-thick-ening trend, with a maximum thickness of 3800, 2000and 1500m, respectively, near the western limit of ourstudy area in the Subandean Ranges. The thick Tariquiastrata exhibit a relatively high sediment accumulation rateof �1mmyear�1 corresponding during this time in theSubandean Zone (Coudert et al., 1995; Echavarria et al.,2003).The high accommodation space and sedimentationrate could be linked to the elastic £exural model fromFlemings & Jordan (1989). According to their model, ero-sion of uplifted area and subsequent transportation anddeposition of sediment in foreland basin decrease the oro-genic load and increase the sediment load within the ba-sin, thereby resulting in an increased basin wavelengthand an increase in sediment supply, causing migration ofbasin-magin facies.We proposed that the high sedimenta-tion rate was not only as a result of high topography, as wellas a shift in climate from a semi-arid to a humid condition(Uba et al., 2005), which led to high precipitation and highdenudation, thus high sediment supply.The high denuda-tion in the west and high sedimentation in the central part


C. E. Ubaet al.

of the basin, coupledwith a humid climate,might have alsoresulted in the migration of the proximal Guandacay con-glomerate facies into the region as rapid unloading

outpaced loading (e.g. Blair & Bilodeau, 1988; Catuneanu,2004). The Tariquia and Guandacay Formations bothclearly exhibit a regional asymmetrical geometry and thin

Fig.15. Isopach maps of the ¢ve late Cenozoic units of the Chaco foreland basin in the study area based on measured surface sections,interpreted industry seismic data, andwell logs. (a) 27^14Ma Petaca Fm. (b) 14^7MaYecua Fm. (c) 7^6MaTariquia Fm. (d) 6^2.1MaGuandacay FM. (e) 2.1Ma Emborozu¤ Fm.



to 0 at the Alto de Izozog basement high. In contrast, theseismic facies of the Emborozu¤ Formation shows a combi-nation of symmetrical and asymmetrical geometries.

Underfilled and overfilled stages of the Chacoforeland basin

The sedimentary and seismic interpretations, as summar-ized in Fig. 15, show an overall asymmetrically westward-thickening wedge, a decrease in depositional energy withdistance from the deformational front, variable sedimentsupply and axial to transverse sediment dispersal.

Deposits of the Petaca Formation, displaying sedimen-tation on a very low topographic gradient (Fig.15a), minorerosion, interbasinal-reworked pedogenic conglomerate,terrestrial condition and a low sediment thickness (�250^0m), suggest an over¢lled stage of the embryonic and stillextremely distal Chaco foreland basin because of its con-sistent easterly transverse sediment dispersal and sedi-mentary style (e.g. Flemings & Jordan, 1989; Jordan, 1995).

During the deposition of the Yecua Formation, basindrainage changed from a transverse to an axial pattern(Fig. 4). This, together with the deposition of mudstone-dominated lacustrine and marginal marine facies, indicatean under¢lled stage (e.g. Flemings & Jordan, 1989), resem-bling the under¢lled phase of theWesternTaiwan and Ca-margo Basins, Bolivia (Covey, 1986; DeCelles & Horton,2003). Because the sediment-accumulation rate decreased(�600m in �7Ma) and the £uvial pattern was modi¢ed

once more, the dominance of ¢ne-grained rocks alsoagrees with models of facies patterns on the distal marginsof under¢lled foreland basin models (e.g. Blair &Bilodeau,1988; Sinclair, 1997).

During the deposition of theTariquia, Guandacay andEmborozu¤ Formations, the predominantly £uvial depos-its, coarsening-upward trend, increase in single-intercon-nected-channel geometry and high avulsion frequencyindicate that the Chaco foreland basin shifted to an over-¢lled stage. Furthermore, this stage is expressed by a pre-dominance of a transverse sediment supply from themountain belt, and a gradual decrease in accommodationspace (e.g. Sinclair & Allen, 1992; Jordan, 1995; Catuneanu,2004).The transition from an under¢lled to an over¢lledstage in a foreland basin system is controlled by a decreasein the rate of £exural subsidence, a decrease in sedimentbypass and an increase in exhumation (Flemings& Jordan,1989; Sinclair & Allen, 1992; Catuneanu, 2004).

Alto de Izozog

The Alto de Izozog is a large, topographical high between550 and 800m elevation and a width of ca. 300 km (Horton& DeCelles, 1997). It forms a NNE-SSW-trending struc-tural high bordering the eastern limit of the study area. Itsuplift mechanism and timing is debated.The Alto de Izo-zog has been interpreted as a recent forebulge depocentre(e.g. Coudert et al., 1995; Moretti et al., 1996; Horton & De-Celles, 1997; DeCelles & Horton, 2003). However, the in-terpreted seismic lines show that all late Cenozoic,Mesozoic and even post-Carboniferous strata onlap andpinch out on this structure (Fig. 10), thereby suggesting apre-Mesozoic origin.

Husson &Moretti (2002) reported a general geothermalgradient of up to 50 1C km�1 and a heat £ow of more than100mWm� 2 at the Alto de Izozog.These values are extre-melyhigh comparedwith the gradient and heat £ow valuesof 26 1C km�1 and 52mWm� 2 from wells further to thewest (Husson & Moretti, 2002). Husson & Moretti (2002)also pointed out that these high heat values are abnormalfor a forebulge depocentre.

In addition, the distance from the Alto de Izozog to thedeformation front is rather short. At its minimum, only ca.70 km separate the exposed basement rocks from the to-pographic front of the Subandean ranges, implying thatthe combinedwidth of the wedge-top and foredeep depo-centre reaches barely100 km (Figs11and15).Theoretically,this short distance is possible, but will imply a very lowelastic thickness and require a thicker basin sedimentary¢ll (e.g.Watts, 2001). In contrast, a high elastic thicknessof 460 km (Stewart & Watts, 1997; Tassara, 2005) is ob-served in the southern Central Andes. The low cross-strike width between assumed forebulge location and de-formation front strongly disagrees with values of ca.4300 km for most other foreland basin systems (e.g. De-Celles &Giles, 1996; DeCelles &Horton, 2003).

We found no evidence of forebulge migration since thelate Miocene from our interpretation of the seismic data,

Fig.15. Continued


C. E. Ubaet al.

althoughCoudert etal. (1995) estimated 90 km of forebulgemigration, based on limited seismic data. A back and forth‘jump’ in forebulge location (Waschbusch&Royden,1992),as observed in the late Devonian/early Mississippian An-tler orogeny of the westernUnited States (Giles &Dickin-son, 1995), would require such a role of the Alto de Izozogsince Mesozoic time, as observed from the onlap andpinch out relationships clearly visible on seismic lines.However, no equivalent pre-Mesozoic foreland basin sys-tem is known below theChaco foreland basin.We thereforeconsider the Alto de Izozog an unlikely candidate for a re-cent forebulge but rather advocate a yet-to-be-de¢ned,pre-Mesozoic continental-interior uplift mechanism.

Depocentremigration through time

The late Cenozoic strata express the foreland basin geo-metry and sedimentation pattern in four depocentres(backbulge, forebulge, foredeep and wedge-top; DeCelles& Giles, 1996). We interpret the Petaca Formation as anOligo-Miocene backbulge depocentre east of its Villa-montes-Camiri structural high and the axis itself, withonly ca. 50m thickness of thePetacaFormation, as the fore-bulge (Fig.16a).The forebulge was likely very low in topo-graphic relief and was therefore subjected only to minorerosion. Its preservation is probably a result of forebulge£exural migration through the study area between �20

Fig.16. Structural cross- sections (modi¢ed after Dunn et al., 1995;Moretti et al., 1996; Baby et al., 1997; Kley et al., 1999) of theevolutionary model illustrating the eastward migration of the deformation front and foreland basin depocentres in time and space withevolution of the Andean fold^thrust belt. EC, Eastern Cordillera; IA, Interandes; SZ, Subandean Zone; PF, Pajonal Fault; PBF, PalosBlanco Fault.



and10Ma (e.g. Crampton &Allen,1995;White etal., 2002),similar to the forebulge migration of the Camargo basin(Horton &DeCelles, 1997; DeCelles &Horton, 2003).Thepresence of a forebulge-backbulge depocentre during Pe-taca time ( �27^14Ma) is also supported by the westward-directed palaeocurrent directions. This interpretationsupports the previously predicted backbulge in this regionby Mcquarrie et al. (2005). The £exural migration of theforebulge into the backbulge area led to minor uplift andmoderate erosion. This may explain the presence of anerosional unconformity (seeMoretti etal., 1996; Echavarriaet al., 2003).

The marginal marine, lacustrine and £uvial facies of theYecua Formation indicate a low ratio between sedimentsupply and accommodation space, and implies the pro-gressive migration of the foredeep into the study area by�14^7Ma.During the deposition of this formation,Cou-dert et al. (1995) and Echavarria et al. (2003) documented asubsidence rate of 1mMa�1. Exhumation and structuraldata indicate that the Subandean Zone began to be de-formed and exhumed in this interval (Kley et al., 1996;Moretti et al., 1996; Echavarria et al., 2003; Ege, 2004).Theunder¢lled stage of the basin is due to the several-million-years time lag between loading to the west of the wideningbasin and its subsequent in¢ll by prograding sedimentwedges (e.g. Blair & Bilodeau, 1988). Our interpretation ofthe Yecua Formation as the ¢ll of a distal foredeep con-trasts with its interpretation as a backbulge depocentre byMarshall et al. (1993), but agrees with the interpretation ofDeCelles & Horton (2003). The distal foredeep develop-ment in the study area may be time-correlative to thewedge-top depocentre in the Camargo basin (DeCelles &Horton, 2003).

The thickening- and coarsening-upward Tariquia For-mation represents a medial-foredeep depocentre ¢ll (Fig.16c). This interpretation is supported by the long-lastingand substantial creation of accommodation space and a

high accumulation rate (e.g. Echavarria et al., 2003), a re-sulting westward increase in large-scale sandstone-domi-nated facies, and its variable £uvial pattern. Palaeocurrentsclearly indicate for the ¢rst time a signi¢cant Andean pro-venance (Uba et al., 2005). During Guandacay time(6^2.1Ma), the proximal foredeep depocentre had arrivedin the study area (Fig. 16d).Westward thickening, consis-tent eastward-directed palaeocurrents, and awestward in-crease in the proportion of conglomerates provide furtherevidence for the presence of the proximal foredeep. Strik-ingly similar facies and geometries of foredeep depocen-tres have been documented by Flemings & Jordan (1989),Sinclair & Allen (1992), DeCelles &Horton (2003) and Ca-tuneanu (2004) for other foreland basins worldwide.

We interpret the Emborozu¤ Formation as representingthe wedge-top depocentre (Fig. 16e). These thickening-and coarsening-upward strata are apparently regionallyrestricted, related to speci¢c thrusts, and show wedge-shaped, high-amplitude re£ectors and growth structuresabove active blind thrusts (e.g. Fig. 12). The contact be-tween the Guandacay and the Emborozu¤ formations is aprogressive regional angular unconformity that marks thetransfer from foredeep to wedge-top depocentre.

Regional tectonic implications

The propagation of a foreland basin system depocentres isrelated to the migration of the orogenic load and to thelithospheric £exural response to crustal load and erosionalunloading (Jordan et al., 1988; Sinclair & Allen, 1992;DeCelles & Giles, 1996; P¢¡ner et al., 2002; DeCelles &Horton 2003; Catuneanu, 2004). Figure 17a shows a com-pilation of totalshortening and shortening rates in theSubandean (Echavarria etal., 2003; Elger etal., 2005; Onck-en et al., in press), whereas Fig. 17b displays the timing ofdeformation and the propagation of the exhumation frontbased on apatite ¢ssion track analysis (Ege, 2004) between

Fig.17. Diagram illustrating (a) timing and total shortening rate compiled from Echavarria et al. (2003) and Oncken et al. (in press); andthe possible major thrusting episode. (b) The rate and propagation of exhumation front of the Central Andes (after Ege, 2004).


C. E. Ubaet al.

Tupiza (EasternCordillera) andVillamontes (westernCha-co plain) corresponding to eastward propagation of the de-formation front since the late Oligocene.The ¢gure showsthat the shortening rate increased markedly around �27(?), 10 and 2.1Ma.

The basal palaeosols of the Petaca Formation indicatelong periods of low sediment accumulation and suggestlittle to no structural activity (e.g. Gubbels et al., 1993).However, the lack of age constraints makes it di⁄cult to re-late it to Andean tectonics.We speculate that the succes-sion of palaeosols may be older than the Andean orogeny(Cretaceous? or Eocene?). The subsequent reworking ofthe palaeosols and the sand-mudstone deposition may re-present the ¢rst in£uence of distant Andean tectonics.Werelate this tectonic episode to a major thrusting event thatis represented byhigh shortening and exhumation rates inFig.17 (Oncken etal., in press; Ege, 2004). It is expressed bythe onset of thrusting to the west in theTupiza region inthe Eastern Cordillera (He¤ rail et al., 1996; Kley et al., 1997)that produced 55 km of shortening and a low crustal load(Gregory-Wodzicki, 2000). Consequently, this shorteningand low crustal load produced low topography that re-sulted in a small-magnitude £exural wavelength, whichprobably caused the foreland basin system to migrate intothe study area.This situation supports a correlation of thelate Oligocene-late Miocene forebulge/backbulge devel-opment to the Cayara and Camargo foredeep depocentrefarther to the west, as proposed by DeCelles & Horton(2003) and Mcquarrie et al. (2005). According to Horton(1998), Mcquarrie (2002) andMˇller et al. (2002), the sedi-mentary basins of the Tupiza region are associated withfold^thrust deformation,whereas apatite ¢ssion track agesfrom the Eastern Cordillera show a decrease in coolingages from �38 to 17Ma (Ege, 2004). During the deposi-tion of the Petaca Formation, the structural, sedimentolo-gical and thermochronologic data indicate majorstructural growth and crustal thickening within the East-ern Cordillera.

Figure 17 shows a pronounced increase in shorteningrate, coupled with an increase in exhumation rate duringthe lateMiocene, which has been attributed to low-angle-basement thrusting and the arrival of the deformationfront in the westernmost part of the Subandean Zone (e.g.Gubbels et al., 1993; Coudert et al., 1995; Moretti et al. 1996;Kley et al., 1997; Echavarria et al., 2003; Ege, 2004). Yecuastrata record basement-imbricated uplift in the Interan-dean and eastern propagation of the fold^thrust system inthe eastern Interandean or Subandean Zone. Coudert et al.(1995) and Echavarria et al. (2003) suggest that the sedi-mentation rate increased rapidly during this time. How-ever, Echavarria et al. (2003) suggest that during this time,the rate of uplift in the Subandean Zone must have beenless than the rate of sedimentation.

We attribute the deposition of theTariquia and Guan-dacayFormations to £exural response due to the high sedi-mentation as a result of high erosion of shortened anduplifted regions in the Interandean and Subandean rangesduring tectonic quiescence, notwithstanding the occur-

rence of minor thrusting (e.g. Echavarria et al., 2003) inthe basin, such as the development of the local, 6^2.1Ma-old La Vertiente structure (Moretti et al., 1996; Fig. 12). Itsuggests that the14^7Mamajor thrust episode might haveproduced a very large crustal load and therefore, a largewavelength shortly before the onset of deposition of theTariquia Formation, which resulted in high denudation inthe western part of the Subandean Zone and increasedaccommodation space and deposition of more than 3500-m-thick-sediments in 7^6Ma in the central part of theSubandean Zone. During this time, the young structureswere just beginning to grow in the study area (e.g. Morettiet al., 1996; Kley et al., 1999; Ege, 2004).

The Emborozu¤ Formation marks both the reactivationof thrusting in the west (Emborozu¤ section) and the arrivalof the deformation front at the western Chaco plain(Aguarague range). However, the timing of this thrusting(Gubbels et al., 1993; Moretti et al., 1996; Echavarria et al.,2003; Ege, 2004) remains debated. Moretti et al. (1996)and Gubbels et al. (1993) estimate the age of the thrustingto postdate the formation and uplift of the leading largeanticline (Aguaragua range; Fig. 2) at 3.3Ma age of micaon tu¡,whereasEchavarria etal. (2003) postulated ayoung-er age of approximately 2.5Ma for the in-sequence thrust,with out-of-sequence reactivation of older structures inthe west at 2^2.2Ma. Our study agrees with the results ofEchavarria et al. (2003) that 2.1Ma (herein constrained)Emborozu¤ strata to the east re£ect in-sequence fold^thrust propagation into the basin, which led to Aguaraguerange uplift, although the equivalent strata to the west re-present the reactivation of the older Nogalitos range (Fig.2), thus forming an out-of sequence intermontane basin(e.g. Echavarria et al., 2003).

CONCLUSIONS

The combination of seismic stratigraphy and outcrop fa-cies interpretation argues for a close interaction betweenAndean fold^thrust belt deformation and Chaco forelandbasin development since the late Oligocene, resulting in agroup of eastward-migrating foreland system depocen-tres, driven primarily by crustal shortening and tectonicloading. Signi¢cant £exural subsidence developed sincethe late Miocene. Variably created accommodation spacewas predominantly ¢lled by non-marine siliciclastics.

The sedimentary ¢ll of the Chaco foreland basin can beassigned to di¡erent depocentres starting with the PetacaFormation in the backbulge and forebulge, through distal-,medial- and proximal foredeep by Yecua, Tariquia andGuandacay Formations, respectively, and ¢nally to wedge-top deposition of the Emborozu¤ Formation.

Three major tectonic episodes are expressed in the fore-land strata: (1) the lateOligocene uplift of the EasternCor-dillera initiated the foreland development and is expressedas relatively steep to low-angle basement thrusts; (2) lateMiocene formation of the Intrandean/Subandean fold-and-thrust belt led to a pronounced under¢lled stage;



and (3) late Pliocene shortening generated the over¢lled,coarse clastic wedges of the Emborozu¤ Formation. Fore-land development and depocentre migration agree wellwith fold-and-thrust belt exhumation rates.

The late Cenozoic Chaco basin is a classical example ofa foreland basin system.The overall coarsening- and thick-ening-upward trend and stratigraphic architectures docu-ment a propagating fold^thrust belt and correspondingforeland basin depocentres (DeCelles & Giles, 1996). Si-milar migration of depocentre with time and foreland ba-sin architecture has been recorded for numerous otherbasins worldwide, such as the Karoo foreland basins (Ca-tuneanu et al., 1999), Taiwan (Covey, 1986) and North Al-pine (Schlunegger et al., 1997; P¢¡ner et al., 2002).

ACKNOWLEDGEMENTS

This paper is part of a PhD thesis by the ¢rst author at theFreie Universit�t Berlin, Germany. The authors were sup-ported ¢nancially by the DFG through the Sonder-forschungsbereich (SFB) 267 and logistically by ChacoS.A., SantaCruz,Bolivia.We are indebted toOscarAranibar,Fernando Alegria and Nigel Robinson of Chaco S.A. fortheir assistance.Thanks are also due toDavidTu¢no Ba¤ nzerof Yacimientos Petroleros Fiscales de Bolivia (YPFB), SantaCruz, Bolivia, for providing some of the seismic lines.Wealso thankHarald Ege andAndresTassara (Freie Universit�tBerlin) for contributing shortening and AFT data and forhelpful and stimulating discussions on Andean geody-namics.We are grateful for comments provided by the re-viewers Jonas Kley, Fritz Schlunegger and Patrice Baby,which greatly improved the early revisions of this manu-script.

REFERENCES

Ayaviri,A. (1964)Geolog|¤ a del a¤ rea deTarija, entre los r|¤ os Pilaya^ Pilcomayo y R|¤ o Bermejo. Informe InternoYPFB (GXG-996).

Ayaviri, A. (1967) Estratigraf|¤ a del Subandino meridional. In-forme InternoYPFB (GXG-1215).

Alves, T.M., Manuppella, G., Gawthorpe, R.L., Hunt,D.W. &Montiero, J.H. (2003) The depositional evolution ofdiaper- and fault-bounded rift basins: examples from the Lu-sitanian Basin ofWest Iberia. Sediment. Geol., 162, 273^303.

Baby,P.,He¤ rail,G., Salinas,R.&Sempere,T. (1992)Geome-try and kinematic evolution of passive roof duplexes deducedfrom cross-section balancing: example from the forelandthrust system of the southern Bolivian Subandean Zone.Tec-tonics, 11, 523^536.

Baby, P.,Moretti, I.,Guillier, B., Limachi, E.,Mendez, E.,Oller, J.&Specht,M. (1995) Petroleum system of the north-ern and central Bolivian Subandean Zone. In: Petroleum Basinsof South America (Ed. by A.J.Tankard, R. Sua¤ rez Soruco & H.J.Welsink),Am. Assoc. Pet. Geol. Mem., 62, 445^458.

Baby, P.,Rochat,P.,Mascle,G.&He¤ rail,G. (1997)Neogeneshortening contribution to crustal thickening in the back-arcof the Central Andes.Geology, 25, 883^886.

Badley, M.E. (1985) Practical Seismic Interpretation. Intern. Hu-man Res. Develop. Corp., Boston.

Beaumont, C. (1981) Foreland basins. Geophys. J. Roy. Astron.Soc., 65, 291^329.

Belotti, H.J., Saccavino, L.L. & Schachner, G.A. (1995)Structural styles and petroleum occurrence in the Subandeanfold and thrust belt of northern Argentina. In: PetroleumBasinsof South America (Ed. by A.J.Tankard, R. Sua¤ rez Soruco & H.J.Welsink),Am. Assoc. Pet. Geol. Mem., 62, 545^555.

Blair, T.C. & Bilodeau, W.L. (1988) Development of tecto-nic cyclotherms in rift, pull-apart, and foreland basins:sedimentary response to episodic tectonism. Geology, 16,517^520.

Bridge, J.S. (2003) Rivers and Floodplains: Forms, Processes, andSedimentary Records. Blackwell, Oxford, 491pp.

Buatois,L.A.,Uba,C.E.,Ma' ngano,M.G,Hulka,C.&Heu-

beck,C. (in press). Deep bioturbation in continental environ-ments: evidence from Miocene £uvial deposits of Bolivia. In:Ichnology at the Crossroads: A Multidimensional Approach to theScience of Organism^Substrate Interactions (Ed. by R. Bromley,L.A.Buatois, J.J.Genise,M.G.Ma¤ ngano&R.Melchor),SEPMSpecial Publ..

Burbank,D.W. & Raynolds, R.G.H. (1988) Stratigraphic keysto the timing of thrusting in terrestrial foreland basins; appli-cations to the northwestern Himalaya. In: New Perspectives inBasin Analysis (Ed. by K.L. Kleinspehn & C. Paola) Springer-Verlag, NewYork.

Cant,D.J. (1992) Subsurface facies analysis. In:FaciesModels:Re-sponse to Sea Level Changes (Ed. by R.G.Walker & N.P. James),Geol. Assoc., 27^45.

Catuneanu, O. (2004) Retroarc foreland systems ^ evolutionthrough time. J. Afr. Earth Sci., 38, 225^242.

Catuneanu, O., Beaumont, C. & Waschbusch, P. (1997) In-terplay of static loads and subduction dynamics in foreland ba-sins: reciprocal stratigraphies and ‘‘missing’’ peripheral bulge.Geology, 25, 1087^1090.

Catuneanu, O., Sweet, A.R. & Miall, A.D. (1999) Conceptand styles of reciprocal stratigraphies: western Canada fore-land basin.TerraNova, 11, 1^8.

Cecil, C.B. (1990) Paleoclimate controls on stratigraphicrepetition of chemical and siliciclastic rocks. Geology, 18,533^536.

Coudert, L., Frappa,M.,Viguier, C. & Arias, R. (1995) Tec-tonic subsidence and crustal £exure in the Neogene ChacoBasin of Bolivia.Tectonophysics, 243, 277^292.

Covey,M. (1986) The evolution of foreland basin to steady state:evidence from the westernTaiwan foreland basin. In: ForelandBasins (Ed. by P. Allen & P. Homewood), Spec. Publ. Int. Assoc.Sediment., 8, 77^90.

Crampton, S.L.&Allen, P.A. (1995) Recognition of forebulgeunconformities associated with early stage foreland basin de-velopment: example from north Alpine foreland basin. Am.Assoc. Pet. Geol. Bull., 79, 1495^1514.

Decelles, P.G. & Giles, K.A. (1996) Foreland basin systems.Basin Res., 8, 105^123.

Decelles,P.G.&Horton,B.K. (2003) Early toMiddleTertiaryforeland basin development and the history of Andean crustalshortening in Bolivia.Geol. Soc. Am. Bull., 115, 58^77.

Dunn, J.F., Hartshorn, K.G. & Hartshorn, P.W. (1995)Structural styles and hydrocarbon potential of the Subandeanthrust belt of Southern Bolivia. In: Petroleum Basins of SouthAmerica (Ed. byA.J.Tankard, R. Suarez Soruco &H.J.Welsink),Am. Assoc. Pet. Geol.Mem., 62, 523^543.


C. E. Ubaet al.

Echavarria, L., Herna¤ ndez, R., Allmendinger, R. & Rey-

nolds, J. (2003) Subandean thrust and fold belt of northwes-ternArgentina: geometry and timing of the Andean evolution.Am. Assoc. Pet. Geol. Bull., 87, 965^985.

Ege,H. (2004) Exhumations- und Hebungsgeschichte der zen-tralen Anden in Sˇdbolivien (211S) durch Spaltspur-Thermo-chronologie an Apatit. PhDThesis, Freie Universit�t, Berlin.

Elger,K., Oncken, O. &Glodny, J. (2005) Plateau-styles ac-cumulation of deformation: the southern Altiplano.Tectonics,24,TC4020, doi:10.1029/2004TC001675.

Flemings, P.B. & Jordan,T.E. (1989) A synthetic stratigraphicmodel of foreland basin development. J. Geophys. Res., 94,3851^3866.

Emery, D. & Myers, K.J. (1996) Sequence Stratigraphy. BlackwellPublishers, Oxford, 297pp.

Giles,K.A.&Dickinson,W.R. (1995) The interplay of eustasyand lithospheric £exure in forming stratigraphic sequences inforeland settings: an example from the Antler foreland, Neva-da andUtah. In:Stratigraphic Evolution of ForelandBasins (Ed. byG.M. Ross), Spec. Publ. SEPM, 52, 187^211.

Gregory-Wodzicki, K.M. (2000) Uplift history of the centraland northernAndes: a review.Geol.Soc.Am.Bull.,112,1091^1105.

Gubbels,T.L., Isacks,B.L.&Farrar,E. (1993)High-level sur-face, plateau uplift, and foreland development, Bolivian cen-tral Andes.Geology, 21, 695^698.

He¤ rail,G.,Oller, J.,Baby, P.,Bonhomme,M.G.& Soler, P.(1996) Strike-slip faulting, thrusting and related basins inCenozoic evolution of the southern branch of theBolivian oro-cline.Tectonophysics, 259, 201^212.

Hofmann,M.H.,Hendrx,M.S.,Moore, J.N.&Sperazza,M.

(2006) Late Pleistocene and Holocene depositional history ofsediments inFlatheadLake,Montana: evidence fromhigh-re-solution seismic re£ection interpretation. Sediment. Geol., 184,111^131.

Horton, B.K. (1998) Sediment accumulation on top of the An-dean orogenic wedge: Oligocene to lateMiocene basins of theEastern Cordillera, southern Bolivia. Geol. Soc. Am. Bull., 110,1174^1192.

Horton,B.K.&Decelles,P.G. (1997)The modern foreland ba-sin system adjacent to theCentral Andes.Geology, 25, 895^898.

Horton, B.K. & Decelles, P.G. (2001) Modern and ancient£uvial megafans in the foreland basin system of the centralAn-des, southern Bolivia: implications for drainage network evo-lution in fold^thrust belts.Basin Res., 13, 43^63.

Horton, B.K.,Hampton, B.A., Lareau, B.N. & Baldello¤ n,E. (2001)Tertiary provenance history of the northern and cen-tral Altiplano (Central Andes, Bolivia): a detrital record of pla-teau-margin tectonics. J. Sediment. Res., 72, 711^726.

Hulka, C. (2005) Sedimentary and tectonic evolution of theCenozoic Chaco foreland basin, Southern Bolivia. PhD The-sis, Freie Universit�t Berlin.

Hulka,C.,Gr�fe,K.U.,Sames,B.,Heubeck,C.&Uba,C.E.(in press). Depositional setting of the middle to late MioceneYecua Formation of the central Chaco foreland basin, Bolivia.J. South Am. Earth Sci.

Husson, L. & Moretti, I. (2002) Thermal regime of fold andthrust belts ^ an application to the Bolivian sub Andean zone.Tectonophysics, 345, 253^280.

Isacks, B.L. (1988) Uplift of the central Andean plateau andbending of the Bolivian orocline. J.Geophys.Res., 93, 3211^3231.

Jimenez-Miranda, F. & Lopez-Murillo, R. (1971) Estratigraf|¤ade algunas secciones del subandino sur y centro. Informe InternoYPFB.

Jones,M.A.,Heller, P.L., Roca, E.,Garce¤ s,M. & Cabrera,L. (2004) Time lag of syntectonic sedimentation across an al-luvial basin: theory and example fromEbroBasin, Spain.BasinRes., 16, 467^488.

Jordan,T.E. (1981) Thrust loads and foreland basin evolution,Cretaceous, western United States. Am. Assoc. Pet. Geol. Bull.,65, 2506^2520.

Jordan,T.E. (1995) Retroarc foreland and related basins. In:Tec-tonics of Sedimentary Basins (Ed. by C.J. Busby & R.V. Ingersoll),pp. 331^362. Blackwell Scienti¢c, Oxford.

Jordan, T.E., Flemings, P.B. & Beer, J.A. (1988) Dating ofthrust-fault activity by use of foreland basin strata. In: Newperspectives in basin analysis (Ed. by K. Kleinspehn & C. Paola),Springer^Verlag, NewYork, 307^330.

Jordan,T.E.,Reynolds, J.H.&Erikson, J.P. (1997) Variabilityin age of initial shortening and uplift in the central Andes, 16^331300S. In:Tectonic Uplift and Climate Change (Ed. byW.F. Rud-diman), pp. 41^61. Plenum Press, NewYork.

Kennen, L., Lamb, S.H. &Hoke, L. (1997) High-altitude pa-leosurfaces in the Bolivian Andes: evidence for late Cenozoicsurface uplift. In: Palaeosurfaces: Recognition, Reconstruction, andPalaeoenvironmental Interpretation (Ed. by M.Widdowson), Spec.Publ. Geol. Soc. Am., 120, 307^323.

Kley, J. (1996) Transition from basement-involved to thin-skinned thrusting in the Cordillera Oriental of southern Boli-via.Tectonics, 15, 763^775.

Kley, J.,Gangui,A.H.&Krˇger,D. (1996)Basement-involvedblind thrusting in the eastern Cordillera Oriental, southernBolivia: evidence from cross-sectional balancing, gravimetricand magnetotelluric data.Tectonophysics, 259, 171^184.

Kley, J., Monaldi, C.R. & Salfity, J.A. (1999) Along-strikesegmentation of the Andean foreland: causes and conse-quences.Tectonophysics, 301, 75^94.

Kley, J.,Mˇller, J.,Tawackoli, S., Jacobshagen,V. &Man-

utsoglu, E. (1997) Pre-Andean and Andean-Age Deforma-tion in the Eastern Cordillera of Southern Bolivia. J. SouthAm. Earth Sci., 10, 1^19.

Makaske, B., Smith, D.G. & Berendsen, H.J.A. (2002) Avul-sions, channel evolution and £oodplain sedimentation ratesof the anastomosing upper Columbia River, British Columbia,Canada. Sedimentology, 49, 1049^1071.

Marshall, L.G. & Sempere,T. (1991) The Eocene to Pleisto-cene vertebrates of Bolivia and their stratigraphic context: areview. In:Fosiles y facies deBolivia:Vertebrados (Ed. byR.Suarez-Soruco),Rev.Te¤ cn. deYPFB, 11, 631^652.

Marshall, L.G., Sempere,T. & Gayet,M. (1993) The Petaca(late Oligocene-middle Miocene) and Yecua (late Miocene)formations of the Subandean-Chaco basin, Bolivia, and theirtectonic signi¢cance.Doc. Lab. Ge¤ ol. Lyon, 125, 291^301.

Mcquarrie,N. (2002) The kinematic history of the central An-dean fold^thrust belt, Bolivia: implications for building a highplateau.Geol. Soc. Am. Bull., 114, 950^963.

Mcquarrie, N. & Decelles, P.G. (2001) Geometry and struc-tural evolution of the central Andean backthrust belt, Bolivia.Tectonics, 20, 669^692.

Mcquarrie,N.,Horton,B.K.,Zandt,G.,Beck, S.&Decel-

les, P.G. (2005) Lithospheric evolution of the Andean fold^thrust belt, Bolivia, and the origin of the central Andean pla-teau.Tectonophysics, 399, 15^37.

Miall, A.D. (1996) The Geology of Fluvial Deposits. Springer-Ver-lag, Berlin.

Mitchum Jr, R.M., Vail, P.R. & Sangree, J.B. (1977)Seismic stratigraphy and global changes of sea level, part 6:



stratigraphic interpretation of seismic re£ection patterns indepositional sequences. In: Seismic Stratigraphy ^ Application toHydrocarbon Exploration (Ed. by C.E. Payton), Am. Assoc. Pet.Geol.Mem., 26, 516.

Moretti, I.,Baby, P.,Mendez,E.&Zubieta,D. (1996)Hydro-carbon generation in relation to thrusting in the SubandeanZone from181 to 221S, South Bolivia. Pet. Geosci., 2, 17^28.

Mˇller, J.P., Kley, J. & Jacobshagen,V. (2002) Structure andCenozoic kinematics of theEasternCordillera.Tectonics, 21,1^27.

Oller, J. (1986) Consideraciones generales sobre la geologia yestratigra¢a de la Faja Subandina norte. MSc Thesis Univer-siteMayor de San Andres, La Paz.

Oncken, O., Hindle, D., Kley, J., Elger, K. & Victor, P. (inpress). Deformation of the Central Andean upper plate sys-tem^facts, ¢ction, and constraints for plateau models. In: De-formation Processes in the Andes (Ed. by O. Oncken, G. Chong, G.Franz,P.Giese,H.J.G˛tze,V.Ramos,M.Strecker&P.Wigger),Front. Earth Sci.Monogr. Ser. 1.

Padula, L.E. & Reyes, F.C. (1958) Contribucio¤ n al Le¤ xico Es-tratigra¤ ¢co de las Sierras Subandinas, Repu¤ blica de Bolivia.Bolet|¤ nTe¤ cn.YPFB, 1, 9^70.

Pfiffner, O.A., Schlunegger, F. & Buiter, S. (2002) TheSwiss Alps and their peripheral foreland basin: stratigraphicresponse to the deep crustal processes.Tectonics, 21, 3.1^3.16.

Retallack,G.J. (1997) A Color Guide to Paleosols.Wiley, Chiche-ster, 175pp.

Roeder,D. & Chamberlain, R.L. (1995) Structural geology ofSub Andean fold and thrust belt in northwestern Bolivia. In:PetroleumBasins of South America (Ed. by A.J.Tankard, S. Suarez&H.J.Welsink),Am. Assoc. Pet. Geol.Mem., 62, 459^479.

Russo,A. (1959) Estructura y estratigraf|¤ a del a¤ rea de Agua Sala-da.Bolet|¤ nTe¤ c.YPFB, 3, 13^35.

Sangree, J.B. &Widmier, J.M. (1977) Seismic stratigraphy andglobal changes of sea level, part: seismic interpretation of clas-tic depositional facies. In: Seismic Stratigraphy ^ Applications toHydrocarbon Exploration (Ed. by C.E. Payton), Am. Assoc. Pet.Geol.Mem. 26 (pp. 516.

Schlunegger, F., Leu,W. & Matter, A. (1997) Sedimentarysequences, seismic facies, subsidence analysis, and evolutionof the Burdigalian upper Marine Grop, central Switzerland.Am. Assoc. Pet. Geol. Bull., 81, 1185^1207.

Sempere,T. (2000) Sediment accumulation on top of the An-dean orogenic wedge; Oligocene to lateMiocene basins of theEastern Cordillera, southern Bolivia; discussion and reply.Geol. Soc. Am. Bull., 112, 1752^1759.

Sempere,T. (1995) Phanerozoic evolution ofBolivia and adjacentregions. In: Petroleum Basins of South America (Ed. by A.J. Tan-kard, R. Sua¤ rez Soruco & H.J. Welsink), Am. Assoc. Pet. Geol.Mem., 62, 207^230.

Sempere,T.,He¤ rail,G.,Oller, J.&Bonhomme,M.G. (1990)Late OligoceneÊarly Miocene major tectonic crisis and re-lated basins in Bolivia.Geology, 18, 946^949.

Sempere, T., Butler, R.F., Richards, D.R., Marshall, L.G.,Sharp,W.&Swisher,C.C., III (1997)Stratigraphy and chronol-ogy of the upperCretaceous-lowerPaleogene strata inBolivia andnorthwest Argentina.Geol. Soc. Am.Bull., 109, 709^727.

Sheffels, B. (1988) Structural Constraints on Crustal Shortening inthe Bolivian Andes. PhD Thesis, Massachusetts Institute ofTechnology, Boston.

Sheriff, R.E. & Geldart, L.P. (1995) Exploration Seismology.Cambridge University Press, Cambridge.

Sinclair, H.D. (1997) Tectonostratigraphic model for under-¢lled peripheral foreland basins: an Alpine perspective. Geol.Soc. Am. Bull., 109, 324^346.

Sinclair,H.D. &Allen, P.A. (1992) Vertical vs horizontal mo-tions in the Alpine orogenic wedge: stratigraphic response inthe foreland basin.Basin Res., 4, 215^232.

Sinclair, H.D., Coakley, B.J., Allen, P.A. & Watts, A.B.(1991) Simulation of foreland basin stratigraphy using a di¡u-sion model of mountain belt uplift and erosion: an examplefrom the central Alps, Switzerland.Tectonics, 10, 599^620.

Smith,D.G. (1986) Anastomosing river deposits, sedimentationrates and basin subsidence, Magdalena River, northwesternColumbia, South America. Sediment. Geol., 46, 177^196.

Somoza, R. (1998) Updated Naszca (Farallon)-South Americarelation motions during the last 40 My: implication for mountainbuilding in the central Andean region. J. South Am. Earth Sci., 11,211^215.

Stewart, J. &Watts, A.B. (1997) Gravity anomalies and spatialvariations of £exural rigidity at mountain ranges. J. Geophys.Res., 102, 5327^5352.

Suarez Soruco, R. (2000) Le¤ xico estratigra¤ ¢co de Bolivia. Rev.Tec.YPFB, 17, 1^213.

Tankard, A.J. (1986) On the depositional response to thrustingand lithospheric £exure: examples from the Appalachian andRocky Mountain basins. In: Foreland Basins (Ed. by P. Allen &P. Homewood), Spec. Publ. Int. Assoc. Sediment., 8, 369^392.

Tassara, A. (2005) Interaction between the Nazca and SouthAmerican plates and formation of theAltiplano-Puna plateau:review of a £exural analysis along the Andean margin (151^341S).Tectonophysics, 399, 39^57.

Uba,C.E.,Heubeck,C.&Hulka,C. (2005) Facies analysis andbasin architecture of the Neogene Subandean synorogenicwedge, southern Bolivia. Sediment. Geol., 180, 91^123.

Wadworth, J., Boyd, R., Diessel, C.F.K. & Leckie, D.A.(2003) Stratigraphic styles of coal and non-marine strata inhigh-accommodation setting: Falher Member and Gates For-mation Lower Cretaceous), western Canada. Can. Assoc. Pet.Geol. Bull, 51, 275^303.

Waschbusch, P.J. & Royden, L.H. (1992) Spatial and tem-poral evolution of foredeep basins: lateral strength variationsand inelastic yielding in continental lithosphere. Basin Res., 4,179^196.

Watts, A.B. (1992) The e¡ective elastic thickness of the litho-sphere and evolution of foreland basin.Basin Res., 4, 169^178.

Watts, A.B. (2001) Isostasy and Flexure of the Lithosphere. Cam-bridge University Press, Cambridge, UK, 458 pp.

White,T.,Furlong,K.&Arthur,M. (2002) Forebulge migra-tion in the Cretaceous Western Interior basin of the centralUnited States.Basin Res., 14, 43^54.

Welsink, H.J.,Martinez, E., Aranibar, O. & Jarandilla, J.(1995) Structural inversion of a Cretaceous rift basin, southernAltiplano, Bolivia. In: Petroleum Basins of South America (Ed. byA.J.Tankard, R. Sua¤ rez-Soruco & H.J.Welsink), Am. Assoc. Pet.Geol.Mem., 62, 305^324.

Manuscript received 7 July 2005;Manuscriptaccepted17April 2006


C. E. Ubaet al.

Documents

9c960522eba62937f6