evolutia bazinului de foreland chaco

8/3/2019 Evolutia Bazinului de Foreland Chaco

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

Institut fr GeologischeWissenschaften, FreieUniversitt Berlin, Berlin, Germany

ABSTRACTEastward Andean orogenic growth since the late Oligocene led to variable crustal loading, exural

subsidence and foreland basin sedimentation in the Chaco basin. To understand the interaction

between Andean tectonics and contemporaneous foreland development, we analyse stratigraphic,

sedimentologic and seismic data from the Subandean Belt and the Chaco Basin.The structural

features provide a mechanism for transferring zones of deposition, subsidence and uplift.These can

be reconstructed based on regional distribution of clastic sequences. Isopach maps, combined with

sedimentary architecture analysis, establish systematic thickness variations, facies changes and

depositional styles.The foreland basin consists ofve stratigraphic successions controlled by Andean

orogenic episodes and climate: (1) the foreland basin sequence commences between $27 and 14 Ma

with the regionally unconformable, thin, easterly sourced uvial Petaca strata. It represents a

signicant time interval of low sediment accumulation in a forebulge-backbulge depocentre. (2) The

overlying $14 7 Ma- old Yecua Formation, deposited in marine, uvial and lacustrine settings,

represents increased subsidence rates from thrust-belt loading outpacing sedimentation rates. It

marks the onset of active deformation and the underlled stage of the foreland basin in a distal

foredeep. (3) The overlying $7^6 Ma- old, westerly sourced Tariquia Formation indicates a relatively

high 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 of

syntectonic, wedge- shaped, westerly sourced, thickening- and coarsening-upward clastics of the

(4) $6^2.1 Ma- old Guandacay and (5) $2.1 Ma-to -Recent Emborozu Formations represent the

propagation of the deformation front in the present Subandean Zone, thereby indicating s elective

trapping of coarse sediments in the proximal foredeep and wedge-top depocentres, respectively.

Overall, the late Cenozoic stratigraphic intervals record the easterly propagation of the deformationfront and foreland depocentre in response to loading and exure by the growing Intra- and

Subandean fold-and-thrust belt.

INTRODUCTION

Foreland basin systems develop as a result of exuralwarp-

ing of the lithosphere in response to supralithospheric

and sublithospheric orogenic wedging (DeCelles & Giles,

1996; Pner etal., 2002). Lithospheric exure under static

loads generates down-bending exure proximal to the

orogen, which migrates as the load advances. Forelandbasins therefore exhibit a characteristic asymmetric

cross-section. Their sedimentary ll generally preserves

and records a detailed exural response of the continental

lithosphere 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 the

lithosphere and the applie d loads (Watts, 1992, 2001). De-

Celles & Giles (1996) characterized foreland basin systems

into fourdierent depocentres: wedge-top, foredeep, fore-

bulge and backbulge. Each depocentre exhibits distinctive

internal architecture, sedimentology and structure. Ac-

commodation space is created by combined static anddynamic subsidence (DeCelles & Giles, 1996; Catuneanu

etal., 1997).

The Chaco foreland basin of the central South America

is a classic example of a foreland basin system in a retro-

arc position. It can be subdivided into the Interandean

Zone, the Subandean Zone and the Chaco plain tectono-

morphologic units (Uba et al., 2005) (Fig. 1). The basin

formed during the late Cenozoic (Sempere etal.,1990; De-

Celles& Horton, 2003) in response to Nazca-South Amer-

ican plate convergence and its related eastward interaction

with the Brazilian shield.Detailed structural studies in the

Interandean and Subandean Zones documented structur-

al styles and timing of deformation (Sempere et al., 1990;

Correspondence: Cornelius Eji Uba, Institut fr GeologischeWissenschaften, Freie Universitt Berlin, Malteserstrasse 74 -100, 12249 Berlin, Germa ny. E-mail: [email protected] -po tsdam.de1 Present address: Institut fr Geowissenschaften, UniversittPotsdam, Karl-Liebknecht Str. 24/25,14476 Potsdam/Golm, Ger-

many.

BasinResearch (2006) 18, 145170, doi: 10.1111/j.1365-2117.2006.00291.x

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


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Oligocene (Oller, 1986; Sheels, 1988; Sempere etal., 1990;

Baby et al., 1992; He rail et al., 1996). This foreland basinforms the easternmost part of the Andean orogen, which

developed from the Cretaceous in the Altiplano (Horton

& DeCelles, 2001; DeCelles & Horton, 2003) as a result of

subduction of the Nazca plate below the South American

plate and the simultaneous subduction of the Brazilian

Shield at an initially low rate of 5.8 cm year1 to a subse-

quent maximum rate of up to 15.2 cm year1 (in the late

Oligocene; Somoza, 1998). Andean deformation com-

menced in the west with formation of the Altiplano basin

and foreland sedimentation, with the depocentre probably

at the present Eastern Cordillera (Sempere etal.,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 presumably

included in a large intracontinental plain of non-deposi-tion.

The geology of the Bolivian Andes is classied into six

tectonomorphic units, of which three units participate in

the late Cenozoic foreland system (Fig. 1). Sedimentary

units 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 the

Brazilian 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 active

shortening at its le ading edge (Fig. 2; Oller, 1986; Sheels,

1988; Baby etal.,1992; He rail etal.,1996). Late Cenozoic se-

dimentary strata are commonly well exposed along anks

Fig. 2. Geological and structural map of the study area (modied from Suarez-Soruco, 2000) showing the data s et and localities of

measured sections mentioned in the text:1, Abapo; 2,Tatarenda; 3,Saipuru; 4, Piriti; 5,San Antonio; 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, San Telmo ; 20, Nogalitos ; 21, Emborozu .

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

Evolutionof thelateCenozoicChaco foreland basin


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of the major leading syn- and anticlines near the wedge-

tip of the Subandean zone. This region, between $641300

and 621300E and 181450 and 221300S, forms the principal

study area.

Chaco foreland basin sedimentation is assumed to have

begun approximately 27 Ma ago near the Eastern Cordil-

lera (Sempere etal., 1990), as a result of eastward migration

of the deformation front ahead of the Interandean and

Subandean 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 foreland

basin strata had also been deposited in considerable thick-

ness in the region occupied by the present-daySubandean

Zone before its uplift and incorporation into the eastward-

migrating orogenic wedge. Although these deposits are

well preserved in the Subandean Zone, in some areas thecoarse-grained proximal foreland basin deposits have

been eroded.The eroded proximal basin sedimentation is

as 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 AFT

samples of the youngest Mesozoic strata in the Subandean

Zone (Ege, 2004).

Magnitude and timing of shortening

Lithospheric thickening and c orresponding shortening in

the 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 of

the Eastern Cordillera in the late Miocene. 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 of

thrusting, accompanied by large-scale folding, produced

a generallye astward-younging synorogenic wedge (Moret-

ti et al., 1996; DeCelles & Horton, 2003; Echavarria et al.,

2003).

During the late Oligocene, the Eastern Cordillera was

the focus of pronounced shortening (Kleyetal.,1997; Hor-

ton, 1998). Figure 3 shows structural styles and major

thrust sheets, illustrating that the Subandean Zone is de-

formed by mostly in-sequence, thin- skinned thrust

sheets 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 unc onformities

and stratigraphically distinct foreland packages. A total

shortening of 210 336 km is postulated for the Central An-

des (Baby etal., 1992; Moretti etal.,1996; McQuarrie & De-

Celles, 2001; Mlleretal., 2002; Elger etal., 2005) together.

The Interandean and Subandean Zones take up 140 and

86 km shortening at 201S and 221S, respectively (Baby et

al., 1997). This matches well with a total shortening of$140 km at211S in the Interandean and Subandean zones

together (Kleyetal.,1997). Moretti etal. (1996) calculated a

peak shortening rate between 6 and 2.1Ma, followed by a

minimum shortening rate between 2.1 Ma and the present

in the Subandean Zone. Their values, however, disagree

with the estimates by Echavarria etal. (2003), who postulate

two periods of high shortening rates (11 and 8 mm year1)

at 9^7 and 2^0 Ma, respectively, separated by an in-be-

tween low of 0^5 mm year1 at 221300 latitude.These con-

tradictions may be due to the paucity of direct age dates for

the deformation and the inherent variability of geologic

cross-section construction and interpretation.The variations in shortening values (Moretti etal., 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 diachronous

movement on individual thrust sheets. For example,Echa-

varria et al. (2003) attribute the 2-Ma-shortening event to

thrust reactivation in the south-western Subandean Zone

( $221300), whereas Moretti et al. (1996) interpreted the

2.1 Ma shortening as a major displacement event synchro -

nous with folding and uplift of the leading Aguarage

range (Fig. 2). No data are available to constrain the time

of formation of the ramp anticlines of the western Chaco

plain, which are clearly visible on industry reection- seis-

Fig. 3. Structural balanced cross- section of the Central Andes from Altiplano to Chaco plain at 211S. Modied after Kley etal. (1999)

and Elger et al. (2005). UKFZ, Uyuni-Khenayani Fault Zone ; SVT, San VicenteThrust ; CYT, Cama rga-Yavi Thrust ; IAT, Inte randean

Thrust ; SAT, Subandean Thrust ; PF, Pajonal Fault; PBF, Palos Blanco Fault. See Fig. 2 for location.

r 2006 The Authors. Journal compilat ion r 2006 Blackwell Publishing Ltd, Basin Research, 18, 145 170148

C. E. Ubaet al.


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mic proles but have as yet only an indistinct and low to-

pographic expression. Anticline cores of these structures

are morphologically expressed in a N^S trend of low hills,

cut by east^west-trending gullies.These structures appear

to be actively forming and indicate the continuous east-

ward growth of the Andes onto the Brazilian Shield.

DATA AND METHODS

The data set compiled for this study includes measuredsections, seismic data and well logs. Twenty- one strati-

graphic sections along major rivers, small streams and

road cuts in the Subandean foothills were measured and

sampled for lithologic, sedimentologic and biostrati-

graphic data (Fig. 2) to document architectural style and

basin geometry. In addition, we interpreted 45 wire-line

logs and their well reports from hydrocarbon industry ex-

ploration wells and tied them to42800 km of 2-D indus-

try s eismic proles. Wire-line logs (g-ray, resistivity and

sonic), combined with well reports, provide ne vertical

details 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)and Alves etal. (2003)to study the Upper

Marine Molasse Group of the North Alpine foreland basin

and the Lusitanian rift basin of West Iber ia, respectively.

Wire-line log analysis and well reports were used in com -

bination with seismic facies attributes to delineate dier-

ent stratigraphic packages and to correlate them to the

ve late Cenozoic formations.The seismic data, wire-line

logs and well reports were provided by Chaco S.A. and Ya-

cimientos Petroleros Fiscales de Bolivia (YPFB), Santa

Cruz.

The seismic lines cover mostly the Chaco plain where

outcrop is poor or absent, and partially extend into the

foothills of the Subandean Zone. They dene the regional

stratigraphic 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 their

thickness variations from vertical facies associations. Syn-

thetic seismogram and check-shots from well logs were

used 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 largely

composed of siliciclastic non-marine redbeds with minor

shallow-marine strata. We used a detailed stratigraphy

after Suarez Soruco (2000) that is principally based on

lithology, with only minor modications(Fig. 4).The basin

ll includes (from base to top) the Petaca, Yecua, Tariquia,

Guandacay and Emborozu Formations. Age dating of

these formations has proved dicult and principally relies

on a combination of mammal biostratigraphy and radio-

metric dating of rare tus (e.g. Marshall et al., 1993; Mar-shall & Sempere, 1991; Moretti et al., 1996; Echavarria et

al., 2003; Hulka, 2005). Notwithstanding the recent by

published new 40Ar/39Ar radiometric ages for the late

Cenozoic 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 to

document the late Cenozoic lithostratigraphy of the

southernBolivia.However, the ages should be appliedwith

caution. 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 on Marshall etal. (1993), Moretti etal. (1996), Echavarria etal. (2003), Hulka (2005), and Hulka etal. (in press).




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Petaca formation

Cenozoic sedimentation in the Chaco basin commenced

during the late Oligocene (assumed ca. 27Ma; Marshall

et 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 overlies

Mesozoic eolian strata (Sempere, 1995). The lower part of

the Petaca Formation consists of greenish grey, white and

light 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 (see

hammer in circle for scale). (b) Shallow-marine-lacustrine mudstone- dominated facies with thin-bedded ooid-, shell hash-dominated

sandston e bed of the Yecua Fm (arrow). (c) Channeli zed sandstone beds with de siccatio n cracks (arrow) of theTariquia Fm. (d)

Alternation of conglomerate and s andstone beds of the Guandacay Fm. (e) Sheet-like cobble-boulder-dominated conglomerate bed

with thin beds of sandstone of the Emborozu Fm.


C. E. Ubaet al.


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to clustered, blocky to massive and bracciated nodules

(Uba etal., 2005).The calcrete body is overlain by horizon-

tal to disorganized clast-supported reworked-pedogenic

conglomerate (Fig.5a) composing of poorly sorted, densely

packed clasts of poorly rounded intraformational re-

worked calcrete nodules and subordinate chert. The con-

glomerates show sharp and erosive bases. Medium to

very-coarse-grained sandstone and sandy to ne-grained

mudstone mark the upsection lithology of the Petaca For-

mation.The calcareous, red to grey, bioturbated sandstone

is characterized by tabular to lenticular beds, trough

cross-, planar and horizontal stratications, as well as

rip-up clasts at the base. The massive, laminated mud-

stone bodies have bioturbation, minor desiccation cracks

and padogenesis. The formation thins towards the centre

of the study area (Villamontes-Camiri axis, Fig. 6).The re-

worked pedogenic conglomerate and sandstone bodies

show channel and bedform architectural elements and an

overall ning-upward sequence. Cross- stratication in

sandstones indicates a westward- directed drainage( Fig.4).

Uba etal. (2005) attributed the thick calcrete horizons to

well-developed palaeosols, indicating 0 or low sedimenta-

tion in an arid to semiarid climate in which evaporation

generally exceeded precipitation (e.g. Cecil, 1990). The

lithofacies and architectural elements in the Petaca con-

glomerate and sandstone indicate variable high-energy

Fig. 6. Correlated proles of stratigraphic sections of the late Cenozoic strata in the northern study area $191S. SeeFig.2 for location.




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stream ows in a channelized setting. Uba etal. (2005) andMarshall et al. (1993) interpret the Petaca strata as having

been deposited by braided stream. The ning-upward

trend and changes in bedform represent a decrease in ow

strength or depth as a result of waning of ood intensity

(Miall,1996). The occasional occurrence of successions of

palaeosols indicates predominantly non-deposition and

surface 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 unc onformity

that may have formed as a far- eld response to early An-

dean tectonics (Sempereetal.,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 et

al.,1990;Marshall & Sempere,1991). However, as the age of

the basal calcretes has not been ascertained, the onset of

deposition is poorly constrained.

Yecua formation

The up to 600 -m-thick Yecua Formation (Padula & Reyes,

1958) overlies the Petaca Formation with an indistinct low-

angle erosional unconformity. The Yecua Formation shows

a west-to-east and northeast-to-southwest facies varia-

tion. North of Camiri, it consists of red-green to brown

sandstone^mudstone couplets (Fig. 5b) showing herring-

Fig. 7. Correlated proles of stratigraphic sections of the late Cenozoic rocks in the southern study area $211S. See Fig. 2 for location.


C. E. Ubaet al.


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bone cross-stratication, laminated, convolute, aser,

wavy and lenticular bedding in ning- and coarsening-

upward successions.This lithofacies also consists of gyp-

sum veins, syndepositional structures, bioturbation and

desiccation 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 (south

of Camiri), the Yecua Formation consists of red to light

brown, lenticular andvery ne- to medium- grained sand-

stone interbedded with red to light-brown, ripple-lami-

nated sandstone couplets. The proportion of mudstone

bodies dominate over the sandstone (Fig. 5b). The sand-

stones show erosive channel structure and ning-upward

trends.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, the

sandstone proportion, bed thickness and the channel pro-

portion of the Yecua Formation increase upsection and to-

wards the west (Figs 6 and 7).

We interpret the fossiliferous and varicoloured Yecua

facies 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 by

Marshall 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 the

west and south are products of uvial overbank and chan-nel processes, with occasional lacustrine and mudat set-

tings. Hulka et al. (in press) placed these variations in a

regional 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 developing

foreland basin as far as Camiri.The age of the Yecua strata

has variably been estimated based on ostracodes and fora-

minifera to be $14 7 Ma (Padula & Reyes, 1958; Marshall

etal.,1993; Hulka etal., in press) and11 7Ma (Moretti etal.,

1996). Recently published 40Ar/39Ar radiometric ages of

10.49 0.33 and 9.41 0.52 Ma ( Hulka, 2005) on inter-

bedded tus 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 (Tariquia

Formation of Bolivian nomenclature; Russo,1959; Ayaviri,

1964; Moretti et al., 1996; Suarez Soruco, 2000) near Ar-

gentinas border with Bolivia that yielded an age of

9.95 0.34 Ma ( Echavarria etal., 2003).

Tariquia formation

TheTariquia Formation (Russo,1959; Ayaviri,1964) is up to

3800- m- thick and overlies the Yecua Formation with gra-

dational contact. The Tariquia Formation is characterized

by thick- and thin-bedded sandstone bodies interbedded

by 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 sandstone

bodies range between 0.5 and 15 m thickness and consist

of sharp erosional base, ribbon and channel geometry

(Fig. 5c), and extend laterally for hundreds of meters.Sandstone units have massive bedding, planar, trough

cross- 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 upward

increase in the degree of vertical stacking, bed thickness

and lateral interconnectedness in the sandstone unit.

Overall, the mean grain size, channel interconnectedness,

sandstone proportion and thickness of the Tariquia

Formation increase towards the west (Figs 6 and 7). The

massive, laminated- or ripple-stratied interbedded

mudstone^sandstone couplets show sheet geometry and

are laterally extensive. Taenidium barreti trace fossils (Bua-

tois etal., in press) in the thick-bedded channelized sand-

stone and in mudstone and sandstone couplets are

common.The trace fossils disrupt the primary sedimen-

tary structures. A distinguishing feature of the Tariquia

Formation is the presence of abundant mudcracks (Fig.

5c, arrow), occasional syndepositional deformation, and

poorly developed palaeosols that are more dominant in

the 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 a

range 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-bedded

sandstone units indicate deposits from crevasse channels.

In overbank areas,mudstone and sheet sandstone were de-

posited by crevassing splayand suspension fallout. Follow-

ingUba etal. (2005),we interpret theTariquia Formation as

a 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), vertical

channel stacking, frequent crevassing and avulsion, and a

general lack of lateral channel migration architecture (e.g.

Smith, 1986; Makaske et al., 2002; Bridge, 2003). The in-

ferred depositional processes and lithofacies are similar

to active modern meganfans in the Chaco plain that re-

ceive sediment from large uvial networks (Horton & De-

Celles, 2001). Avulsion and crevassing result in the

development of new channels on the overbank, whereas

the active channels are abandoned (Smith, 1986; Makaske

etal., 2002).The abundant rip-up clasts that may have been

formed by erosional scouring of overbank sediments and




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high degree of bioturbation in theTariquia Formation 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 by

either 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 Tariquia

strata. The Tariquia Formation age is late Miocene

(Chasicoan-Huaquerian) based on biostratigraphy (Mar-

shall & Sempere, 1991), in agreement with a single

apatite ssion-track age of 7 Ma from Mica (Moretti et al.,

1996). In addition, Moretti et al. (1996) assumed 6 Ma as

the upper age limit of the Tariquia Formation. In the ab-

sence of a well-constrained age for this unit, we use the

imprecise age of 7^6 Ma for the deposition duration for

this formation.

Guandacay formation

The up-to-1500-m-thick Guandacay Formation consists

of conglomerate, sandstone and mudstone (Jimenez-Mir-

anda & Lopez-Murillo,1971) (Fig. 5d).The granule-cobble

conglomerate 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 s cour surfaces. Gravel

bedforms and poorly developed lateral accretion surfaces

are common architectural elements. The dominantly

medium- to very-coarse-grained sheet-like sandstones

are moderately to well sorted, and are laterally extensive

for several hundreds of metres (Fig. 5d). The sandstonebodies consist of trough cross-,planar, ripple and horizon-

tal stratication, and occasional stringers of pebbles.

Thick interbedded mudstones and sandstone are massive

to laminated and laterally continuous for several tens or

hundreds of metres. Lenses of thin coalseams,poorly pre-

served bioturbation, and weakly developed mottled soils

are present (Uba etal., 2005).The conglomerates generally

thicken and coarsen upsection and to the west. The con-

glomerate and sandstone bodies show, like the Tariquia

Formation, an upward increase in stacked packages, lateral

interconnectedness, and multi- to single-storey channel

systems that grade into the interbedded mudstone andsandstone. Palaeocurrent measurements indicate a north-

east-to - southeast-directed ow (Fig. 4).

The conglomerate and sandstone lithofacies provide

evidence of deposition in uctuating, high-energy, bed-

load-dominated large uvial channels, anked by ood -

plains, and zones of incipient soil development (Uba et

al., 2005). The dimensions of channel lls and the types

of sedimentary structures in the Guandacay Formation

suggest large discharges (e.g. Horton & DeCelles, 2001).

Consequently, Uba etal. (2005) envision a proximalbraided

setting on a medial uvial megafan, similar to those that

deposited the Camargo Formation and that drain the

modern central Andean (Horton & DeCelles, 2001; De-

Celles & Horton, 2003). Lenses of coal suggest the pre-

sence of a ponded area and vegetation, and therefore, a

humid palaeoclimate (Uba et al., 20 05). Vertical stacking

and aggradation of channels into overbank deposits imply

crevassing and avulsion, indicating periodic abandonment

of active channels. The weakly developed palaeosol and

poorly preserved bioturbation may suggest a high over-

bank aggradation rate( Bridge, 2003).The contact betweenthe Tariquia and the overlying Guandacay Formation is

unconformable (Moretti et al., 1996; Echavarria et al.,

2003), approximately 6 Ma in age (Moretti etal., 1996), and

is marked by a distinct increase in mean grain size. Hulka

(2005) estimated the top of this formation at 2.1 0.2Ma

based on 40Ar/39Ar dating of tu at its contact to the Em-

borozu Formation in the Abapo section (Fig. 2).The age of

the Guandacay Formation is therefore late Miocene to

Early Pliocene (6^2.1Ma).

Emborozu formation

The Emborozu Formation (Ayaviri, 1967) is exposed only

in the northeast (Abapo Section) and within synclines in

the southwestern (Emborozu and Nogalitos ; Fig. 2) study

area. 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 near

the present Subandean topographic front, growth struc-

tures occur (Echavarria et al., 2003), documenting a

syndeformational origin. The Emborozu Formation is

dominatedby an up-to- 60-m -thick,cobble-bouldercon-

glomerate that reaches at least 153 cm in diameter (Fig. 5e;

Uba et al., 2005).This laterally extensive (several hundreds

of metres) conglomerate shows sharp erosive scoursurfaces, sheet-like to lenticular single-channel geometry,

inverse and normal grading, and moderately to poorly

developed imbrications. This conglomerate lithofaces is

associated with up -to 6 -m-thick, coarse- to very-coarse-

grained, sheet-like sandstone with horizontal, trough

cross-, ripple- and planar stratication. The single-

storey, vertically stacked conglomerate and sandstone

bodies grade into medium- to very coarse-grained,

rippled, massive-laminated interbedded sandstone and

mudstone in which poorly preserved burrows and coal

lenses occur. Upsection and to the west,the thickness,lat-

eral continuity, amalgamation and maximum grain s ize ofthe conglomerate and sand bodies increase and the per-

centage of overbank nes decreases. The palaeodrainage

pattern shows a northeast-to- southeast-directed ow

(Fig. 4).

Uba etal. (2005) interpret the Emborozu Formation as a

uctuating-energy, bedload, proximal uvial system of

successions of large, isolated to amalgamated channels.

The presence of a thick to subordinate oodplain and the

lateral extent suggests deposits on a proximal uvial mega-

fan (Horton & DeCelles, 2001; Uba etal., 2005). The sharp

scour surfaces may represent discretechannels 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.


11/26

Guandacay Formation with a well- dened 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 based

on Ar Ar on mica to be 3.3 Ma. However, a new

2.1 0.2 Ma (tu; Abapo section; Ar^Ar on mica) estimate

by Hulka (2005) agrees relatively well with 1.8Ma docu-

mented by Echavarria et al. (2003) for correlative strata in

Argentina. Consequently, 2.1 0.2Ma is used herein as

the basal age of the formation.

SEISMIC STRATIGRAPHY OF THE LATECENOZOIC DEPOSITS

A good well-to-seismic tie and the lateral continuity of

horizons allowed interpretation of the visible geometric

features on the seismic lines. After interpreting seismic

lines and wire-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 that

can be correlated with wire-line logs. Seismic facies attri-

butes include prominent reectors, termination geometry

(onlap, toplap, downlap and truncation), reection cong-

uration, and external form. Not all onlap and truncation

geometries could be mapped in the seismic sections due

to limited vertical resolution combined with small unit

thickness.Figures 8 and 9 illustrate the most characteristic

seismic facies features and g-ray (GR), resistivity (ILD)

and sonic (DT)log responses of the ve packages.The cor-

responding lithofacies and depositional environments are

calibrated by well data.

Package N1

The base of N1is a prominent, readily traceable reector of

high amplitude, medium frequencyand medium continu-

ity across most of the study area, marking the contact be-

tween the late Cenozoic foreland basin and underlying

Mesozoic strata (Fig. 8a). In some seismic sections, the un-

derlying Mesozoic strata show diuse toplap and trunca-tions with low angular geometry. In wire-line logs, this

contact shows an abrupt increase from $100 to 170 Om

in ILD and an immediate drop from $90 to 40 in DTre-

sponses (e.g. Fig. 8a). A clear dierentiation can be made

between the top of N1 and the base of N2 as a result of a

pronounced medium- to high- amplitude, continuous

and medium- to high-frequency reector that is easily

identied and correlated throughout all seismic sections

(Fig. 10). Wire-line logs show a sharp increase from $30

to 120 API in the GR curve and from $50 to 70 in the

DT curve, coupled with an abrupt increase to $90 in the

ILD curve.The thinness of this package does not allow a

detailed seismic facies characterization. However, in some

areas, N1displays internally lateral extensive, low- to med-

ium-amplitude, subparallel, discontinuous, low-fre-

quency reectors (Fig. 10). Among them, a wedge- shaped

set reaches up to ca. 200 m ( $0.2 ms) thickness across a

broad area in the western Chaco plain and gradually

pinches out to 0 with onlap terminations upon reaching

the Alto de Izozog high (Fig. 10). Wire-line logs through

N1 (Fig. 11) generally indicate low GR value (30^60 API),

average 90^55 DT, and 90^170 Om ILD values. The GR

curves show cylindrical shape characteristic.

The toplap reection terminations and truncation of

N1 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.The

base and top of each package is dened.




12/26

formity. It could also be a condensed or non- deposition

surface with minor erosion (e.g. Mitchum etal.,1977; Sher-

i & Geldart, 1995). However, the quality of the seismic

data and the overall low thickness ofN1do not allowa clear

dierentiation between toplap and erosional truncation.

Distal onlap and the overall wedge form of N1on the Alto

de Izozog High indicate basin progradation. We interpret

package N1 as a sand- dominated aggradational uvial sys-

tem.This interpretation is based on seismic facies charac-

teristic (variable-amplitude, discontinuous and low-

frequency reection) combined with low GR and DTand

the cylindrical shape of the GR curve, implying that this

package consists mainly of relatively high-energy uvial

deposits (Badley, 1985; Cant, 1992; Emery & Myers, 1996).

The cylindrical shape of GR logs suggests an aggrading

braided 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 change

in 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 section

located approximately 40 km farther south.


C. E. Ubaet al.


13/26

pedogenesis. We interpret the strong seismic facies and

sharp wire-line log relations at the contact of N1 and N2

to reect an unc onformity. We correlate N1 based on theseismic facies and wire-line log characteristics mentioned

above as the subsurface equivalent of the Petaca Forma-

tion.

Package N2

The top of N2 is a laterally continuous, high-to-moder-

ate-amplitude reector (Fig. 8a). The GR, ILD and DT

logs do not show a sharp dierence 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

(0.20 ms) in the west.To the east, N2 terminates with onlapgeometries on the Alto de Izozog High, where it overlies

N1 and pinches out onto the Mesozoic strata (Figs 10 and

11). Its internal seismic facies are: to the west, N2 displays

internal seismic reections that show variable-amplitude,

discontinuous, subparallel, low-frequency, low vertical

spacing and chaotic pattern. Low-scale hummocky clino-

forms dip at variable angles ( $1^21);to the east, low-scale

complex sigmoid-oblique seismic reections also occur.

In contrast, to the east, low-angled clinoforms, discontin-

uous, low-amplitude, semi transparent and chaotic reec-

tions,coupledwith low acoustic impedance,occur (Figs 8a

and 10). Wire-line logs show $60^120 API in GR, 95^85

Om 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 high

to low values (80 : 20) in the N2 package (Fig. 9). However,the low GR and high DTvalues increase upsection. GR

logs show an irregular or serrated response (Cant, 1992;

Emery & Myers, 1996) and small-s cale variability in values

as indicated by numerous thin cycles with ning-upward

trends (Fig.11).

The variable-amplitude, discontinuous, semi-trans-

parent, internal structure, higher GR and lower DTvalues

are typical of a poorly stratied mudstone-dominated sys-

tem,deposited mainlyby suspension settling and subordi-

nate channel settings (Cant, 1992; Alves etal., 2003). Based

on the seismic facies and well-log characteristics, we can

interpret 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 setting

dominantly in the west and south of the study area.The in-

terpretation of varied-depositional settings is further

supported by thin sandstone intervals in wire-line logs, a

relatively high and serrated GR response, sigmoid-oblique

and hummocky clinoforms, varied-amplitude and low fre-

quency (Sangree & Widmier,1977; Badley, 1985; Cant, 1992;

Sheri & Geldart, 1995). As N2 thickens westward towards

its depocentre, it develops varied-amplitude and low-

continuity reections and dened-clinoforms. The

small-scale variability observed in the GR, DT and ILD

Fig. 10. Segment of a W-E uninterpreted and interpreted migrated seismic line along approximately 201450S showing the ve late

Cenozoic sequences and thrusting and folding of the foreland sequences.The location of the line is shown in Fig. 2.




14/26

proles largely represents variation in depositional energy

associated with high-frequency cyclicity (Cant, 1992; Em-

ery & Myers, 1996). The gradual GR and DTresponses at

N2^N3 contact suggest a fairly steady change of deposi-

tional environment between thes e two packages.The e ast-

ward thinning indicates the presence of a palaeo -high near

the eastern border of the study area before the deposition

of N2. The upsection decrease in the GR and increase in

the DT values, which can imply upsection increase in

sandstone proportion, reects a basinwide shift in facies.

The characteristic mudstone-dominated seismic facies

and wire-line log attributes of N2 package are analogous

to the Yecua Formation.

Package N3

The base of N3 is a variable continuous and varied-ampli-

tude reector; the GR and DTshow decrease and decrease

in values, respectively (Figs 8 and 9).The laterally continu-

ous, moderate- to high-amplitude reector marks the top

of this package and a transitional contact to N4. In wire-

line logs, this contact is marked by a relatively sharp low

GR and high ILD and DTresponses (Figs 8b and 9). N3

shows a maximum thickness of $1500m ( $0.75 ms) in

the western study area, thinning gradually eastward to

pinch out at the Alto de Izozog basement high, where it

onlaps and overlies Mesozoic strata (Fig. 10). Internally,

N3 displays varied-seismic facies; in the lower portion of

the section, it shows low- to medium-amplitude, discon-

tinuous, subparallel, low-frequency, semi-transparent

and hummocky reectors (Figs 10 and 12). However, the

seismic facies changes upward to more moderate-low con-

tinuous, varied-amplitude, less chaotic and less hum-

mocky reectors and wedge-sheet external forms. The

proportion 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 reectors. The log character of the N3

package is distinguished from the underlying N2 package

because it contains relatively lower GR (30^90API),higher

DT (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 percent

and thickness of low GR and DTresponses increasing sig-

nicantly upward and to the west, where it reaches tens of

metres in thickness (Figs 9, 12 and 14). GR and DT curves

have both serrated and occasional bell shapes.

We interpret the varied- amplitude, subparallel, moder-

ately to discontinuous seismic facies, coupledwith low GR

Fig.11. Segment of a W-E uninterpreted and interpreted migrated seismic line along approximately 191300S showing pinch out of the

late Cenozoic sequences.The location of the line shown is in Fig. 2.


C. E. Ubaet al.


15/26

and high DT response of package N3, as alternations of

sandstone and mudstone deposited by s andstone-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 GR

and 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, combined withup -

section decrease in the GR values and thickness of lowGR

response) 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. The

semi-transparent and chaotic seismic features represent

lack of stratication. 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 reectors, variable seismic characteristics and pro-

portion of sandstone and GR thickness within N3 suggest

a strong progradational pulse concomitant with a basin-

ward shift in facies and depocentre location. We assign

the N3 package to theTariquia Formation because of inter-

pretation of characteristic of the seismic facies expressions

and wire-line logs identied in this package.

Package N4

The top of package N4 is characterized by a prominent,

high-amplitude, continuous reector that can be mapped

and correlated throughout all seismic sections. This top

contact is marked by local toplap and truncation termina-

tions of N4 reectors 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 contact

is a well-dened angular unconformity.The base of N4 is

Fig.12. W-Ewire-line logs at about 211S illustrating aspects of sequence boundaries.The depositional sequences identied can b e

correlated to near outcrops. See Fig. 2 for location.




16/26

delimited by a moderate to high-amplitude, continuous

reector.Wire-line logs indicate a decrease and an increase

in the values of GR and DT, respectively. The thickness

variation of Package N4 is similar to that of N3, with a

broad area in the west, where the thickness exceeds ap-

proximately1500 m ( $750 ms) thinning to the eastto zerol

at the Alto de Izozog high. At this basement high, N4 also

onlaps 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 reections, suggesting several

tens of metre-scale bedding. In the southern part of the

study area, N4 includes a growth structure near the LaVer-

tiente Fault (Fig.12), showing an upsection decrease in the

inclination of reectors and onlap geometry (Fig.12 inset).

However, this fault is limited to the southern part of the

study area and is not recognized further north (e.g. Fig.

10). The N4 package is identied in wire-line logs by low

GR (30^70 API), high DT (105^140) and high ILD (95^

130) responses (Figs 12 and 14). Figure 9 show that both

the GR and SP log curves show an upsection decrease

and increase in response, respectively, and have cylindrical

and bell shapes (Figs 9, 12 and 14). However, the percent of

low GR values in N4 are relatively higher and thicker than

in the underlying N3 package (Figs 9 and 14). The low GR

and DT intervals have a s errated shape.

The high GR and DT values, cylindrical shape and

thickness, combined with parallel-subparallel, varied-

amplitude, moderate^low continuous and wedge-shaped

seismic facies suggest a thick intercalation of sandstone

with a conglomerate-dominated uvial environment,

probably in a braided setting (Cant, 1992; Emery & Myers,1996).The upward increase in the high GR log response at

the base of each cylindrical- or bell-shaped unit coupled

with its heterogeneity and vertical spacing in seismic lines

may indicate conglomerate lithofacies.We interpret the in-

tercalated-thin-serrated GR and DTresponse as alternat-

ing sand- and mudstone-bodies of overbank deposits

(Cant, 1992; Emery & Myers, 1996). The overall decrease

in 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 are

analogous to the sandstone-conglomerate-dominated

Guandacay Formation.The top of N4 marks a local angu -

Fig.13. Segment of a W-E uninterpreted and interpreted migrated seismic line along approximately 211S showing the structural styles

and the ve sequences.The lo cation of the line is shown in Fig. 2.


C. E. Ubaet al.


17/26

lar unconformity.The westward thickening,wedge geome-

try of the package suggests that the depocentre is located

to the west.The pinch out and onlap of N4 at the Alto de

Izozog High indicate that of this basement palaeohigh ex-

isted before deposition of N4. The onlap and thinning of

the N4 reectors on the La Vertiente Fault are attributed

to syndepositional deformation. The pronounced top re-

ector implies a strong acoustic impedance contrast and

likely represents the erosional surface or an angular regio-

nal unconformity (Fig.12, inset; Dunn etal., 1995; Moretti

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

2003).

Package N5

The base ofN5 is dened bya pronouncedthick high-am-

plitude, continuous, high-frequency reector, with onlaps

on the underlying N4 package.The top of the N5 package

is not well dened 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 s equences identied can be

correlated to near outcrops. See Fig. 2 for location.




18/26

reectors. However, well-log data are not available for

the top of N5. This package shows a maximum thickness

of $2000m ($1.0ms) in the west and thins eastward

to $0^200 m ($0.1ms) at the Alto de Izozog high, where

it occasionally shows a progressive onlap on N4. The

internal seismic facies includes laterally extensive, sub- to

parallel reectors ofvariable amplitude,low continuity and

low frequency (Figs 8d, 9 and 12). The external forms pre-sent are sheet andwedge geometries.The seismic sections

show a large vertical spacing of reectors (Figs 8d and 13).

Growth structures above the tips of major fault-propaga-

tion folds display characteristic fan- shaped reectors with

upward-decreasing dip. Only one of these faults (Man-

deyepecuaFault) can be mapped throughoutthe study area

(Figs 10 and13). No wire-line data are available to ascertain

the 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 large

alluvial 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 of

reectors, which probably suggest thick conglomerate

lithofacies. The overall thickening- and coarsening-up-

ward trend, wedge- shaped external form, and pronounced

upsection increase in vertical spacing in seismic section

are interpreted as a continuing basinward shift in grain

size.The seismic facies expressions and alluvial setting in-

terpretation for the N5 package permit its correlationwith

the surface Emborozu Formation.

DISCUSSION

Overall stratigraphic pattern

In constructing the isopach maps (Fig. 15), we compiled

thickness information for each formation derived from

outcrop, depth- converted seismic, wire-line logs and well

report data.The thickness values are not corrected for the

eect of compaction due to limited postdepositional bur-

ial.We constructed isopach maps for ve time periods (oc-

casionally poorly) constrained by age estimates for the

formations (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 space

since the late Oligocene (e.g. Wadworth et al., 2003). The

early basin history is illustrated in a single map (Fig. 15a)

spanning more than 13 Ma. A second map (Fig. 15b) spans

a 7 Ma- time period. In contrast, the nal three time inter-

vals represent only 3^1 Ma each (Fig. 15c^e). As expected,

these ve maps show distinctive thickness patterns.

The isopach map of the 27 14 Ma-old PetacaFormation

(Fig. 15a) shows a regional and broad area along the Villa-

montes-Camiri axiswitho50 m of strata, possibly reect-

ing a structural high. This ridge is anked by up to 250 m

of strata cratonward and4100 m of strata orogenward, re-

spectively. Further east, the Petaca thins to 0 at the Alto de

Izozog structural high. This nding updates a previous

view expressing a lack of thickness variations for this for-

mation (e.g. Gubbels et al., 1993; Moretti et al., 1996). The

very 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 the

basin, augmented by scarce sediment supply inferred from

successions of palaeosols (Gubbels etal.,1993; Horton etal.,

2001).

Gubbels etal. (1993) and Moretti etal. (1996) reported a

similar lack of thickness variation for the Yecua Formation.

In contrast, Fig.15b shows a distinctive westward thicken-

ing, reaching a maximum thickness of $600 m in outcrop

(e.g. Emborozu section). This westward thickening and in -

crease in sandstone proportion of the Yecua strata is also

documented in seismic sections (Figs 6, 7, 10 and13). Dur-

ing 14 7 Ma, exural foreland basin subsidence as a result

of thrusting episode that was probably centred in the pre-

sent-day Interandean- or Subandean Zone (Coudert etal.,

1995; Moretti etal.,1996; Echavarria etal., 2003) led to crea-

tion ofaccommodation space in thewestern part of the ba-

sin. Flemings & Jordan (1989) and Sinclair etal. (1991) show

that thrusting event results in an increase of the ratio

between tectonic subsidence and sediment ux. The

dominantly ne-grained sediments in the more distal

and relatively sandy facies in the west suggest that the de-

positional slope was probably too low to produce large

coarse-grained deposits or there was a long lag-time be-

tween erosion and more coarse-grained sedimentation as

documented 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 proportion

indicates a change in locus of deposition.

The three subsequent isopach maps (7^6, 6^2.1and 2.1^

0; Fig. 15c^e) show a similar and regular westward-thick-

ening trend, with a maximum thickness of 3800, 2000

and 1500 m, respectively, near the western limit of our

study area in the Subandean Ranges. The thick Tariquia

strata exhibit a relatively high sediment accumulation rate

of $1mmyear1 corresponding during this time in the

Subandean Zone (Coudert et al., 1995; Echavarria et al.,

2003).The high accommodation space and sedimentationrate could be linked to the elastic exural model from

Flemings & Jordan (1989). According to their model, ero -

sion of uplifted area and subsequent transportation and

deposition of sediment in foreland basin decrease the oro -

genic load and increase the s ediment load within the ba-

sin, thereby resulting in an increased basin wavelength

and an increase in sediment supply, causing migration of

basin-magin facies.We proposed that the high sedimenta-

tion rate was not only as a result of high topography, as well

as a shift in climate from a semi-arid to a humid condition

(Uba etal., 2005), which led to high precipitation and high

denudation, thus high sediment supply.The high denuda-

tion in the west and high sedimentation in the central part


C. E. Ubaet al.


19/26

of thebasin, coupledwith a humid climate,might have also

resulted 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 both

clearly exhibit a regional asymmetrical geometry and thin

Fig.15. Isopach maps of the ve lateCeno zoic units of the Chaco foreland basin in the study area based on measured surface sections,

interpre ted industry seismic data, and well logs. (a) 27^14 Ma Petaca Fm. (b) 14^7 Ma Yecua Fm. (c) 7^6 Ma Tariquia Fm. (d) 6^2.1 Ma

Guandacay FM. (e) 2.1 Ma Emborozu Fm.




20/26

to 0 at the Alto de Izozog basement high. In contrast, the

seismic 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 with

distance from the deformational front, variable sediment

supply and axial to transverse sediment dispersal.

Deposits of the Petaca Formation, displaying sedimen-

tation on a very low topographic gradient (Fig.15a), minor

erosion, interbasinal-reworked pedogenic conglomerate,

terrestrial condition and a low sediment thickness ($

250^0 m), suggest an overlled stage of the embryonic and still

extremely 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, basin

drainage changed from a transverse to an axial pattern

(Fig. 4). This, together with the deposition of mudstone-

dominated lacustrine and marginal marine facies, indicate

an underlled stage (e.g. Flemings & Jordan, 1989), resem-

bling the underlled phase of the Western Taiwan and Ca-

margo Basins, Bolivia (Covey, 1986; DeCelles & Horton,

2003). Because the sediment-accumulation rate decreased

($600m in $7 Ma) and the uvial pattern was modied

once more, the dominance of ne-grained rocks also

agrees with models of facies patterns on the distal margins

of underlled foreland basin models(e.g. Blair & Bilodeau,

1988; Sinclair, 1997).

During the deposition of theTariquia, Guandacay and

Emborozu 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 the

mountain belt, and a gradual decrease in accommodation

space (e.g. Sinclair & Allen, 1992; Jordan, 1995; Catuneanu,

2004).The transition from an underlled to an overlled

stage in a foreland basin system is controlled by a decrease

in the rate of exural subsidence, a decrease in sediment

bypass 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 between

550 and 800 m elevation and a width ofca. 300 km (Horton

& DeCelles, 1997). It forms a NNE- SSW-trending struc-

tural high bordering the e astern limit of the study area. Its

uplift mechanism and timing is debated.The Alto de Izo-

zog has been interpreted as a recent forebulge depocentre

(e.g. Coudert etal., 1995; Moretti etal., 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 and

pinch out on this structure (Fig. 10), thereby suggesting a

pre-Mesozoic origin.Husson & Moretti (2002) reported a general geothermal

gradient of up to 50 1C km1 and a heat ow of more than

100 mWm 2 at the Alto de Izozog.These values are extre-

mely high compared with the gradient and heat ow values

of 26 1C km1 and 52mWm 2 from wells further to the

west (Husson & Moretti, 2002). Husson & Moretti (2002)

also pointed out that these high heat values are abnormal

for a forebulge depocentre.

In addition, the distance from the Alto de Izozog to the

deformation front is rather short. At its minimum, onlyca.

70km separate the exposed basement rocks from the to -

pographic front of the Subandean ranges, implying thatthe combined width of the wedge-top and foredeep depo -

centre reaches barely100 km (Figs 11 and 15).Theoretically,

this short distance is possible, but will imply a very low

elastic thickness and require a thicker basin sedimentary

ll (e.g.Watts, 2001). In contrast, a high elastic thickness

of460 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 the

late Miocene from our interpretation of the seismic data,

Fig.15. Continued


C. E. Ubaet al.


21/26

althoughCoudert etal. (1995) estimated 90 km of forebulge

migration, 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 western United States (Giles & Dickin-

son, 1995), would require such a role of the Alto de Izozog

since Mesozoic time, as observed from the onlap and

pinch out relationships clearly visible on seismic lines.

However, no equivalent pre-Mesozoic foreland basin sys-

tem is known below the Chaco foreland basin.We therefore

consider the Alto de Izozog an unlikely candidate for a re-

cent forebulge but rather advocate a yet-to-be-dened,

pre-Mesozoic continental- interior uplift mechanism.

Depocentre migration 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 an

Oligo -Miocene backbulge depocentre east of its Villa-

montes-Camiri structural high and the axis itself, with

only ca. 50m thickness of thePetaca Formation,as thefore-

bulge (Fig.16a). The forebulge was likely very low in topo -

graphic relief and was therefore subjected only to minor

erosion. Its preservation is probably a result of forebulge

exural migration through the study area between $20

Fig.16. Structural cross- sections (modied after Dunn etal., 1995; Moretti etal., 1996; Baby etal., 1997; Kley etal., 1999) of the

evolutionary model illustrating the eastward migration of the deformation front and foreland basin depocentres in time and space with

evolution of the Andean fold^thrust belt. EC, Eastern Cordillera; IA, Interandes; SZ, Subandean Zone; PF, Pajonal Fault; PBF, Palos

Blanco Fault.




22/26

and10 Ma (e.g. Crampton & Allen, 1995;White etal., 2002),

similar to the forebulge migration of the Camargo basin

(Horton & DeCelles, 1997; DeCelles & Horton, 2003). The

presence of a forebulge-backbulge depocentre during Pe-

taca time ( $27 14 Ma) is also supported by the westward-

directed palaeocurrent directions. This interpretation

supports the previously predicted backbulge in this region

by Mcquarrie et al. (2005). The exural migration of the

forebulge into the backbulge area led to minor uplift and

moderate erosion. This may explain the presence of an

erosional unconformity (see Moretti etal.,1996; Echavarria

etal., 2003).The marginal marine, lacustrine and uvial facies of the

Yecua Formation indicate a low ratio between sediment

supply and accommodation space, and implies the pro-

gressive migration of the foredeep into the study area by

$14 7 Ma. During the deposition of this formation, Cou-

dert et al. (1995) and Echavarria et al. (2003) documented a

subsidence rate of 1m Ma1. Exhumation and structural

data indicate that the Subandean Zone began to be de-

formed and exhumed in this interval (Kley et al., 1996;

Moretti etal., 1996; Echavarria et al., 2003; Ege, 2004). The

underlled stage of the basin is due to the several-million-

years time lag between loading to the west of the wideningbasin and its subsequent inll by prograding sediment

wedges (e.g. Blair & Bilodeau, 1988). Our interpretation of

the Yecua Formation as the ll of a distal foredeep con-

trasts with its interpretation as a backbulge depocentre by

Marshall etal. (1993), but agrees with the interpretation of

DeCelles & Horton (2003). The distal foredeep develop-

ment in the study area may be time-correlative to the

wedge-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 suppor ted by the long-lasting

and 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.Palaeocurrents

clearly indicate for the rst time a signicant Andean pro-

venance (Uba et al., 2005). During Guandacay time

(6^2.1 Ma), the proximal foredeep depocentre had ar rived

in the study area ( Fig. 16d). Westward thickening, consis -

tent eastward-directed palaeocurrents, and a westward in-

crease in the proportion of conglomerates provide further

evidence for the prese nce of the proximal foredeep. Strik-

ingly similar facies and geometries of foredeep depocen-

tres have bee n documented by Flemings & Jordan (1989),Sinclair & Allen (1992), DeCelles & Horton (2003) and Ca-

tuneanu (2004) for other foreland basins worldwide.

We interpret the Emborozu Format ion as representing

the wedge-top depocentre (Fig. 16e). These thickening-

and coarsening- upward strata are apparently regionally

restricted, related to specic thrusts, and show wedge-

shaped, high-amplitude reectors and growth structures

above active blind thrusts (e.g. Fig. 12). The contact be-

tween the Guandacay and the Emborozu formations is a

progressive regional angular unconformity that marks the

transfer from foredeep to wedge-top depocentre.

Regional tectonic implications

The propagation of a foreland basin system depocentres is

related to the migration of the orogenic load and to the

lithospheric exural response to crustal load and erosional

unloading ( Jordan et al., 1988; Sinclair & Allen, 1992;

DeCelles & Giles, 1996; Pner et al., 2002; DeCelles &

Horton 2003; Catuneanu, 2004). Figure 17a shows a com-

pilation of total-shortening and shortening rates in the

Subandean (Echavarria etal., 2003; Elgeretal., 2005; Onck-

en et al., in press), whereas Fig. 17b displays the timing of

deformation and the propagation of the exhumation front

based on apatite ssion track analysis (Ege, 2004) between

Fig. 17. Diagram illustrating (a) timing and total shortening rate compiled from Echavarria etal. (2003) and Oncken etal. (in press); and

the possible major thrusting episode. (b) The rate and propagation of exhumation front of the Central Andes (after Ege, 2004).


C. E. Ubaet al.


23/26

Tupiza (Eastern Cordillera)and Villamontes (western Cha-

co plain) corresponding to eastward propagation of the de-

formation front since the late Oligocene.The gure shows

that the shortening rate increased markedly around $27

(?), 10 and 2.1 Ma.

The basal palaeosols of the Petaca Formation indicate

long periods of low sediment accumulation and suggest

little to no structural activity (e.g. Gubbels et al., 1993).However, the lack of age constraints makes it dicult 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 of

the palaeosols and the sand-mudstone deposition may re-

present the rst inuence of distant Andean tectonics. We

relate this tectonic episode to a major thrusting event that

is represented by high shortening and exhumation rates in

Fig.17 (Oncken etal.,in press; Ege, 2004). It is expressedby

the onset of thrusting to the west in the Tupiza region in

the Eastern Cordillera (He rail etal., 1996; Kley etal., 1997)

that produced 55 km of shortening and a low crustal load

(Gregory-Wodzicki, 2000). Consequently, this shortening

and low crustal load produced low topography that re-

sulted in a small-magnitude exural wavelength, which

probably caused the foreland basin system to migrate into

the study area.This situation supports a correlation of the

late Oligocene-late Miocene forebulge/backbulge devel-

opment to the Cayara and Camargo foredeep depocentre

farther to the west, as proposed by DeCelles & Horton

(2003) and Mcquarrie et al. (2005). According to Horton

(1998), Mcquarrie (2002) and Mller et al. (2002), the sedi-

mentary basins of the Tupiza region are associated with

fold^thrust deformation, whereas apatite ssion track ages

from the Eastern Cordillera show a decrease in coolingages from $38 to 17 Ma (Ege, 2004). During the deposi -

tion of the Petaca Formation, the structural, sedimentolo -

gical and thermochronologic data indicate major

structural growth and crustal thickening within the East-

ern Cordillera.

Figure 17 shows a pronounced increase in shortening

rate, coupled with an increase in exhumation rate during

the late Miocene, which has been attributed to low-angle-

basement thrusting and the arrival of the deformation

front in the westernmost part of the Subandean Zone (e.g.

Gubbels etal., 1993; Coudert etal., 1995; Moretti etal.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 in

the eastern Interandean or Subandean Zone. Coudert etal.

(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 been

less than the rate of sedimentation.

We attribute the deposition of theTariquia and Guan-

dacay Formations to exural response due to thehigh sedi-

mentation as a result of high erosion of shortened and

uplifted regions in the Interandean and Subandean ranges

during tectonic quiescence, notwithstanding the occur-

rence of minor thrusting (e.g. Echavarria et al., 2003) in

the basin, such as the development of the local, 6^2.1 Ma-

old La Vertiente structure (Moretti et al., 1996; Fig. 12). It

suggests that the 14^7Ma major thrust episode might have

produced a very large crustal load and therefore, a large

wavelength shortly before the onset of deposition of the

Tariquia Formation, which resulted in high denudation in

the western part of the Subandean Zone and increasedaccommodation space and deposition of more than 3500-

m-thick- sediments in 7^6 Ma in the central part of the

Subandean Zone. During this time, the young structures

were just beginning to grow in the study area (e.g. Moretti

etal., 1996; Kley etal., 1999; Ege, 2004).

The Emborozu Formation marks both the reactivation

of thrusting in the west( Emborozu section)and the arrival

of 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 etal. (1993) estimate the age of the thrusting

to postdate the formation and uplift of the leading large

anticline (Aguaragua range ; Fig. 2) at 3.3 Ma age of mica

on tu, whereas Echavarria etal. (2003)postulated a young-

er age of approximately 2.5 Ma for the in-s equence thrust,

with out-of-sequence reactivation of older structures in

the west at 2^2.2Ma. Our study agrees with the results of

Echavarria et al. (2003) that 2.1Ma (herein constrained)

Emborozu strata to the east reect in- sequence fold^

thrust propagation into the basin, which led to Aguarague

range 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 etal., 2003).

CONCLUSIONS

The combination of seismic stratigraphy and outcrop fa-

cies interpretation argues for a close interaction between

Andean fold^thrust belt deformation and Chaco foreland

basin development since the late Oligocene, resulting in a

group of eastward-migrating foreland system depocen-

tres, driven primarily by crustal shortening and tectonic

loading. Signicant exural subsidence developed since

the late Miocene. Variably created accommodation spacewas predominantly lled by non-marine siliciclastics.

The sedimentary ll of the Chaco foreland basin can be

assigned to dierent depocentres starting with the Petaca

Formation in the backbulge and forebulge, through distal-,

medial- and proximal foredeep by Yecua, Tariquia and

Guandacay 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 late Oligocene uplift of theEasternCor-

dillera initiated the forelanddevelopment and is expressed

as relatively steep to low-angle basement thrusts; (2) late

Miocene formation of the Intrandean/Subandean fold-

and-thrust belt led to a pronounced underlled stage;




24/26

and (3) late Pliocene shortening generated the overlled,

coarse clastic wedges of the Emborozu Formation. Fore-

land development and depocentre migration agree well

with fold-and- thrust belt exhumation rates.

The late Cenozoic Chaco basin is a classical example of

a 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 other

basins worldwide, such as the Karoo foreland basins (Ca-

tuneanu et al., 1999), Taiwan (Covey, 1986) and North Al-

pine (Schlunegger etal., 1997; Pner etal., 2002).

ACKNOWLEDGEMENTS

This paper is part of a PhD thesis by the rst author at the

Freie Universitt Berlin, Germany. The authors were sup-

ported nancially by the DFG through the Sonder-

forschungsbereich (SFB) 267 and logistically by Chaco

S.A., SantaCruz, Bolivia.We are indebtedto Oscar Aranibar,

Fernando Alegria and Nigel Robinson of Chaco S.A. for

their assistance.Thanks are also due to David Tuno Ba nzer

of Yacimientos Petroleros Fiscales de Bolivia (YPFB), Santa

Cruz, Bolivia, for providing some of the seismic lines. We

also thank Harald Ege and AndresTassara (Freie Universitt

Berlin) for contributing shortening and AFT data and for

helpful 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.

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