Early to middle Tertiary foreland basin development ... - GSA Journals

Early to middle Tertiary foreland basin development and the history of Andean crustal shortening in Bolivia

Peter G. DeCelles† Department of Geosciences, University of Arizona, Tucson, Arizona 85721, USA

Brian K. Horton Department of Earth and Space Sciences, University of California, Los Angeles, California, 90095 USA

ABSTRACT A .2.5-km-thick succession of Tertiary strata in the Eastern Cordillera of southern central Bolivia consists of predominantly fluvial and lacustrine deposits. The age of the base of the succession is middle Paleocene, and its upper part (which is erosionally truncated) is probably late Oligocene– early Miocene. The lower 50–130 m of the succession consists of interbedded fluvial sandstone and mudstone (including minor paleosols) of the Santa Lucıá Formation. These strata are overlain by up to 50 m of pervasively pedogenically altered mudstone and sandstone in the upper part of the Santa Lucıá Formation and lower part of the Impora Formation that may represent much of late Paleocene–Eocene time. Above the paleosol zone is a thin zone (;10 m) of lacustrine carbonate and clastic rocks in the uppermost Impora Formation, and this in turn is overlain by a .2000-m-thick, upwardcoarsening succession of clastic fluvial deposits in the Cayara, Camargo, and Suticollo Formations. Paleocurrent data indicate that fluvial channels that deposited the Santa Lucıá, Impora, and Cayara Formations flowed mainly westward, whereas channels responsible for the Camargo and Suticollo Formations flowed generally eastward. Modal sandstone petrographic data show a long-term evolution from subarkosic (during Santa Lucıá deposition) to quartzarenitic (during Impora and Cayara deposition) to sublitharenitic (during Camargo and Suticollo deposition) compositions. We argue that the lithofacies and sedimentaccumulation history of this succession are most consistent with deposition in an eastward-migrating foreland basin system. E-mail: [email protected].

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Critical to this interpretation is the zone of extreme stratigraphic condensation in the lower Impora Formation, the inferred upper Paleocene–Eocene part of the succession. The abrupt decrease in sediment accumulation represented by this zone is difficult to reconcile with the leading alternative model in which continuous, postrift thermal subsidence resulted in a gradual, continuous decay of early Tertiary sedimentaccumulation rates. Stratigraphic condensation is, however, consistent with passage of the forebulge through the region, an interpretation that also can be reconciled with the record of foreland basin development in the Altiplano to the west and gross sediment accumulation patterns in the Subandean zone and modern foreland basin to the east. Simple flexural modeling and palinspastic reconstruction of the complete Cenozoic migration history of the foreland basin system suggest that ;1000 km of foreland lithosphere has been underthrust westward beneath the Andean orogenic belt at the latitude of southern Bolivia. The amount of middle and lower crust that would have been added to the Andean infrastructure is sufficient to explain the present crustal thicknesses in Bolivia. Keywords: Andes, foreland basins, foldthrust belts, Tertiary, sedimentology, sedimentary petrology. INTRODUCTION The Andean orogenic belt is one of the most impressive topographic features on Earth, with a length of .7000 km, a maximum width of ;800 km, and peak elevations of .6.7 km. Crust up to 70 km thick (James, 1971; Wigger et al., 1994; Beck et al., 1996)

GSA Bulletin; January 2003; v. 115; no. 1; p. 58–77; 13 figures; 1 table; Data Repository item 2003008.

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For permission to copy, contact [email protected] q 2003 Geological Society of America

isostatically supports the Central Andean Plateau between latitude 158 and 278 S (Isacks, 1988), the widest, highest part of the orogenic belt (Fig. 1). Two major unresolved problems in Andean geology are (1) the timing of the onset of regional contraction and mountain building, and (2) the kinematic history of this contraction. Both issues lie at the heart of current debates about the construction of the Central Andean Plateau (Isacks, 1988; Gubbels et al., 1993; Wdowinski and Bock, 1994; Allmendinger and Gubbels, 1996; Allmendinger et al., 1997; Lamb and Hoke, 1997; Beck and Zandt, 2002; McQuarrie, 2002). With respect to the first problem, although evidence for Cretaceous rifting in the central Andes is present in northern Argentina and locally in southern Bolivia (e.g., Salfity and Marquillas, 1994; Viramonte et al., 1999), the timing of the transition from extension to regional shortening remains uncertain. Some authors have maintained that the transition occurred during Late Cretaceous or early Paleocene time (e.g., Sempere et al., 1997; Horton et al., 2001), whereas others have contended that regional shortening did not begin until the Eocene (Welsink et al., 1995; Lamb et al., 1997; Lamb and Hoke, 1997; Viramonte et al., 1999) or possibly later (Jordan et al., 1997, 2001). The extreme crustal thickening is generally attributed to horizontal shortening (Lyon-Caen et al., 1985; Isacks, 1988; Allmendinger et al., 1997; Lamb et al., 1997), but current estimates of total shortening are insufficient to explain the actual crustal thickness in much of the Andes (see summaries in Kley and Monaldi, 1998; McQuarrie, 2002). Most shortening estimates are based on balanced regional cross sections in the eastern half of the orogenic belt (the Eastern Cordillera and Subandean zone), where the structure of the foldthrust belt is well exposed. Shortening in this

FORELAND BASIN DEVELOPMENT AND HISTORY OF ANDEAN CRUSTAL SHORTENING, BOLIVIA

Figure 1. Generalized geologic map of southern central Bolivia, showing locations of measured sections and major tectonomorphic zones as discussed in the text. Inset shows the location of the study area in the context of the central Andes. Numbers 1–7 refer to locations of measured stratigraphic sections shown in Figure 4. SYA—Sama-Yunchara anticlinorium. region occurred from Eocene time to the present (Roeder, 1988; Baby et al., 1990, 1997; Sheffels, 1990; Gubbels et al., 1993; Schmitz, 1994; Dunn et al., 1995; Roeder and Chamberlain, 1995; Kley, 1996; Kley et al., 1996, 1997; Schmitz and Kley, 1997; Allmendinger et al., 1997; Jordan et al., 1997; McQuarrie and DeCelles, 2001; McQuarrie, 2002). On the other hand, reliable shortening estimates from the western half of the orogenic belt have not been obtained. Nevertheless, syntec-

tonic sedimentary rocks in the Bolivian Altiplano indicate that the region west of the Altiplano was topographically elevated and capable of driving regional flexural subsidence by early Tertiary time (Horton et al., 2001). Moreover, structural relationships in the modern arc and forearc regions provide evidence for early Tertiary crustal shortening (Chong and Reutter 1985; Hartley et al., 1992; Hammerschmidt et al., 1992; Scheuber and Reutter 1992; Scheuber et al., 1994; Amilibia et al.,

1999; Nicolas et al., 1999) and tectonic (verticalaxis) rotations (Roperch et al., 2000a, 2000b). Therefore, it is plausible that substantial crustal shortening occurred in the western central Andes prior to the Neogene. With regard to the second problem, the kinematic history of the Andes, many authors view the mountain belt as the result of a series of (originally interpreted as only three, but now viewed as six or more) brief orogenic phases (Steinmann, 1929; Me´gard, 1984; Me´-

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gard et al., 1984; Se´brier et al., 1988; Noble et al., 1990) or ‘‘crises’’ (Sempere et al., 1990) that were separated by periods of ‘‘quiescence’’ or regional extension. Other workers have recognized a more continuous progression of deformation (Marocco et al., 1987; Sempere, 1991; Noblet et al., 1996; Horton and DeCelles, 1997), and work in the frontal part of the Andean fold-thrust belt shows that folding and thrusting during Miocene–Pliocene time occurred continuously, not sporadically, as the orogenic front migrated eastward (Gubbels et al., 1993; Jordan et al., 1993, 1997; Hernańdez et al., 1996, 1999; Reynolds et al., 2000). A similar pattern of temporally continuous forward propagation of major thrusting events (punctuated by relatively minor synchronous and out-of-sequence thrusting events) has been documented in most major fold-thrust belts. Thus, we pose the question: Was the kinematic history of the central Andean fold-thrust belt fundamentally sporadic or continuous? In this paper we assess the regional kinematic history of the central Andes by tracking the development of the Andean foreland basin system. Unlike the structural details of a foldthrust belt, foreland basin deposits are often well preserved and provide a regionally integrated geodynamic signature of orogenic tempo (Jordan et al., 1988). We document the sedimentology and provenance of Paleocene to lower Miocene strata in the Eastern Cordillera and develop a predictive model for the evolution of the central Andean foreland basin system. GEOLOGIC SETTING The geology of the central Andes in Bolivia and northern Chile is divisible into six longitudinal tectonomorphic zones (Fig. 1). East of the Chilean forearc lies the Western Cordillera, the location of the present magmatic arc, which consists of a belt of high (5–6.5 km) Miocene–Quaternary stratovolcanoes and ignimbrite sheets. East of the arc lies the central Andean plateau, an ;300,000 km2 topographic plateau with an average elevation of ;4 km (Isacks, 1988). The central part of the plateau is occupied by the Altiplano, a vast internally drained Cenozoic basin. Beneath a veneer of Quaternary sediment, as well as locally cropping out along the northern and eastern parts of the Altiplano, are Cenozoic, Mesozoic, and Paleozoic sedimentary rocks that have been subject to significant horizontal shortening (Lamb and Hoke, 1997; McQuarrie and DeCelles, 2001; McQuarrie, 2002). To the east, the Eastern Cordillera is a bivergent fold-

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Figure 2. Photograph (view toward the west) of the Maastrichtian–Tertiary section in the eastern limb of the Camargo syncline. The high point on the horizon is Cerro Tonka. Vertical relief is ;1600 m. thrust belt composed of Ordovician–Devonian sedimentary rocks, overlain locally by Permian, Jurassic, and Cretaceous strata and local accumulations of Cenozoic strata. Permian– Triassic and middle Tertiary granitic intrusions form several of the highest (.6 km) peaks in the Eastern Cordillera. The western part of the Eastern Cordillera is a west-vergent backthrust belt (the Huarina–San Vicente foldthrust belt; Roeder, 1988; Baby et al., 1990; Kley et al., 1997; McQuarrie and DeCelles, 2001). The eastern part of the Eastern Cordillera is referred to as the Interandean zone and consists of Silurian–Devonian strata carried by east-directed thrust faults (Kley, 1996; Schmitz and Kley, 1997; McQuarrie, 2002). East of the Interandean zone lies the Subandean zone, which is the active frontal part of the fold-thrust belt (Dunn et al., 1995). Farther east lies the Chaco Plain, comprising most of the wedge-top and foredeep depozones of the modern, central Andean foreland basin system (Horton and DeCelles, 1997). We have concentrated on an ;300-kmlong, north-trending belt of Paleocene through lower Miocene rocks exposed in the Camargo and Otavi synclines and in the region northwest of Potosı´ in the Eastern Cordillera of southern Bolivia (Figs. 2, 3, 4). In addition, we draw upon results of recent studies in the Tupiza region ;30–80 km west of the axial trace of the Camargo syncline (He´rail et al., 1996; Tawackoli et al., 1996; Horton, 1996, 1998a, 1998b; Kley et al., 1997). The steeply dipping (up to 908) western limb of the Camargo syncline is locally broken by the Camargo-Tojo thrust, which places Ordovician phyllite against the Upper Cretaceous El Molino For-

mation and, locally, the Paleocene Santa Lucıá Formation. The syncline involves the Cretaceous Aroifilla (or Torotoro), Chaunaca, and El Molino Formations and the Tertiary Santa Lucıá, Impora, Cayara, and Camargo Formations (Fig. 2; Sempere et al., 1997). In some publications (e.g., Sempere et al., 1997), the lower approximately two-thirds of the Camargo Formation in the Camargo syncline has been correlated with the Potoco Formation in the eastern Altiplano. Although these two units may be partly age equivalent, detailed sedimentologic work demonstrates that the Potoco Formation in the Altiplano is much finer grained and differs sedimentologically and compositionally from the strata in the Camargo syncline (Horton and DeCelles, 2001; Horton et al., 2001). In addition, the rocks above the Cayara Formation in the syncline gradually coarsen upward, and a distinct formational boundary does not exist between the lower sandstone-rich and upper conglomeratic strata (Fig. 2). We therefore refer to the entire upward-coarsening succession of mudstone, sandstone, and conglomerate above the Cayara Formation in the Camargo syncline as the Camargo Formation (Troe¨ng et al., 1993; Geobol, 1995). We also have studied Tertiary rocks in the Potosı´ region that have been variably referred to as the Suticollo (Geobol, 1962a), Mondragoń (Sempere et al., 1989), and Tusque and Cayara Formations (Geobol, 1994). Our observations lead us to interpret these rocks to be more or less equivalent to the Camargo Formation as just defined herein (see also Sempere et al., 1997). The upper boundary of the Tertiary section in the Camargo and Otavi synclines is erosional, and in



Figure 3. Chronostratigraphy of the Cretaceous–Tertiary sedimentary rocks in the study area. See text for discussion of sources. the Potosı´ region the Suticollo Formation is overlain in angular unconformity by Miocene ignimbrites. This paper focuses on the Santa Lucıá, Impora, Cayara, Camargo, and Suticollo Formations. CHRONOSTRATIGRAPHY The chronostratigraphy of Tertiary units in the Eastern Cordillera of southern Bolivia (Fig. 3) is defined by a regionally extensive, well-dated, lower stratigraphic interval (Maastrichtian–lower Paleocene El Molino and Santa Lucıá Formations) and a discontinuous, volcanic rock– bearing upper interval (lower Miocene Mondragoń and Bolı´var Formations and strata in basins of the Tupiza area). The interval unconformably overlies deformed Paleozoic through Paleogene strata. Although the Tertiary strata between these lower and upper intervals are not directly dated in Bolivia, age estimates based on apatite fission-track data and reasonable rates of sedimentation for observed lithofacies suggest a primarily late Eocene to early Miocene time span of deposition. For the lower Tertiary succession, the El Molino Formation is defined as Maastrichtian to lower Paleocene on the basis of magnetostratigraphy, biostratigraphy, and 40Ar/39Ar ages on two interbedded tuffs (see summary by Sempere et al., 1997). The overlying Santa Lucıá Formation has been dated at ca. 60–58 Ma (middle Paleocene) on the basis of correlation of an exclusively reversed magnetic po-

larity stratigraphy with chron 26r of the geomagnetic polarity time scale (Sempere et al., 1997). At the upper end of the Tertiary column, the Mondragoń Formation is a clastic unit that crops out along the eastern edge of the Los Frailes volcanic field, including the type locality near the town of Mondragoń (Fig. 1). The Mondragoń Formation is a gently deformed, subhorizontal unit that has been dated as early Miocene (ca. 20–19 Ma) by KAr ages on volcanic tuffs in its lower part (Evernden et al., 1977; Kennan et al., 1995). Gubbels et al. (1993) reported an 40Ar/39Ar biotite age of 17.4 6 0.11 Ma from a tuff bed in the Bolı´var Formation, a probable lateral equivalent along strike ;150 km to the northnorthwest. Lying stratigraphically between the relatively well-dated middle Paleocene and lower Miocene levels of the Tertiary column is a thick succession of Tertiary strata that remains largely undated (Fig. 3). These strata have been assigned to numerous different units, including the Impora, Cayara, Potoco, Camargo, Incapampa, Mondragoń, and Suticollo Formations (Geobol, 1962a, 1962b, 1994, 1995, 1996; Troe¨ng et al., 1993; Sua´rez and Diaz, 1996; Sempere et al., 1989, 1997). These units concordantly overlie the Maastrichtian–Paleocene El Molino and Santa Lucıá Formations. For example, in a partial erosional window beneath the Los Frailes field (Challa Mayu and Ayo Ayo areas; Fig. 1), a moderately deformed Tertiary section originally referred to

as the Suticollo Formation (Geobol, 1962a) rests concordantly upon early Tertiary rocks. Some investigators have reassigned these strata to the Mondragoń Formation (Sempere et al., 1989, 1997; Kennan et al., 1995). However, given the differing basal stratigraphic contacts and intensities of deformation, we consider the Suticollo Formation to be an older unit that predates the lower Miocene Mondragoń Formation. Therefore, the early Miocene is a minimum age constraint for both the deposition and subsequent moderate deformation of the stratigraphic units that constitute the intermediate part of the Tertiary column. For the Camargo syncline, Sempere et al. (1997) suggested that deposition of the Impora and Cayara Formations (totaling ;250 m in thickness) spanned ;0.7 m.y. during the late Paleocene on the basis of a tentative magnetostratigraphic correlation. This interpretation would require that these two units were deposited rapidly at an undecompacted sedimentaccumulation rate of ;0.36 mm/yr, which is inconsistent with the abundant lithologic evidence for extremely slow deposition of these units (discussed subsequently). Sempere et al. (1997) summarized biostratigraphic evidence from units in northern Argentina that are lithostratigraphically similar to the Impora and Camargo (Potoco) Formations and suggest Paleocene–early Eocene ages. In the Camargo syncline, Branisa et al. (1969) reported charophytes from a stratigraphic level that Sempere et al. (1997) later assigned to the ‘‘La Cabanã Member’’ of the Potoco Formation. This interval lies in what we refer to as the lower part of the Camargo Formation. Whereas Branisa et al. (1969) considered the age of the charophytes to be late Eocene or perhaps Oligocene (admitting that middle Eocene cannot be excluded), Mourier et al. (1988) interpreted the fauna as Paleocene–Eocene. We sampled a zone of laminated shale and siltstone at approximately the same stratigraphic level (at the town El Rancho; lat 65813.89 W, long 20848.59 S) and retrieved numerous ostracodes, conchostracans, charophytes, and a single fish tooth. The tooth is of the subfamily Serrasalminae with similarities to the modern genera Colossoma, Piaractus, and Mylossoma (J.G. Lundberg, 2002, personal commun.; Lundberg et al., 1986; Lundberg, 1998). These modern species are restricted to lowland rivers that flow toward the Atlantic Ocean. Unfortunately, none of the material that we collected has proven to be chronostratigraphically definitive. Recent apatite fission-track (AFT) studies of rocks in the Eastern Cordillera provide important constraints on the minimum age of the


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Tertiary succession in the Camargo syncline. Ege et al. (2001) reported AFT ages that range between 38 6 7 Ma and 30 6 5 Ma from Ordovician rocks in this part of the Eastern Cordillera. One of these ages is from a sample located in the hanging wall of the CamargoTojo thrust, which cuts, and therefore postdates, the strata in the western limb of the Camargo syncline. These ages may be reasonably interpreted to indicate the timing of thrusting-induced erosional exhumation of this part of the Eastern Cordillera (Ege et al., 2001). Therefore, the majority of the Tertiary succession in the Camargo syncline is probably older than the Miocene. Combining this constraint with the limited biostratigraphic information and estimated sedimentation rates (discussed in the next section), we tentatively interpret the age of the Camargo Formation to be latest Eocene through Oligocene. This interpretation suggests that the Impora and Cayara Formations lie between the late Paleocene and latest Eocene, an interval of ;20–25 m.y. As we subsequently discuss in detail, the bulk of this time interval is probably archived in a zone of intense, regionally persistent pedogenesis in the transition between the Santa Lucıá and Impora Formations (Figs. 3, 4). SEDIMENTOLOGY The following sedimentological descriptions and interpretations are based on detailed lithofacies data from seven measured sections (Figs. 1 and 4) and observations of the Tertiary succession at many outcrops throughout the study area. We also draw upon previous work by Sempere et al. (1997). Santa Lucıá Formation The Santa Lucıá Formation is a regionally persistent siliciclastic unit throughout the Eastern Cordillera and eastern Altiplano of southern Bolivia (Sempere et al., 1997; Horton et al., 2001; Fig. 4). The unit consists of 50–130 m of quartzose, lenticular sandstone bodies alternating with red siltstones. The sandstone bodies are 1–6 m thick and contain trough and planar cross-stratification, ripples, and plane-parallel stratification with primary

Figure 4. Logs of measured stratigraphic sections of Tertiary rocks in the Eastern Cordillera. Inset shows locations of the sections relative to each other in the context of Tertiary outcrops of the Eastern Cordillera (see also Fig. 1).

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Figure 5. Photographs of the pedogenically altered interval in the transition zone between the Santa Lucıá and Impora Formations. (A) Uppermost part of the Santa Lucıá Formation exposed in canyon just north of Camargo town site (Rio Camargo). The doubleheaded arrow highlights the massive, pedogenically altered sandstone that caps the Santa Lucıá. Arrow in upper left points to the lacustrine facies in the upper Impora Formation. (B) Densely packed pedogenic carbonate nodules in the lower part of the Impora Formation in the Camargo syncline. Scale (10 cm) rests on the transition between two stacked paleosol profiles. (C) Double-headed arrow highlights the massive, nodular sandstone interval that caps the Santa Lucıá Formation in the Challa Mayu section. (D) The ;6-mthick, massive, concretionary sandstone interval at the top of the Santa Lucıá Formation in the El Puente section.

current lineations. Sandstone bodies have erosional bases and typically fine upward from medium- to fine-grained sandstone. Thinner tabular beds of fine-grained, rippled sandstone are also present. Interbedded with the sandstones are red, massive siltstones with carbonate nodules, root traces, and various small tubular burrows. The siltstones have a

distinctive spheroidal weathering habit, clay skins, and superposition of noncalcareous (leached) and nodular calcareous (accumulation) zones. In every section of the Santa Lucıá Formation that we have observed, the uppermost part of the unit is capped by a several-meter-thick zone of massive sandstone or siltstone with pervasive mottling and abun-


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dant carbonate nodules (Fig. 5; see also Sempere et al., 1997). This massive zone grades upward into the lower Impora Formation. In the Otavi syncline, an ;1140-m-thick red-bed unit overlying the El Molino Formation and underlying the gray siltstones of the Impora Formation has been included entirely within the Santa Lucıá Formation. This unit consists of lithofacies that are similar to the Santa Lucıá Formation throughout the study area, with a greater abundance of mudstone. However, a well-developed pencil cleavage and numerous small faults suggest that the Otavi section may be structurally inflated by intraformational faulting and penetrative strain. We therefore do not regard our Otavi section of the Santa Lucıá Formation as a reliable representative of the true thickness of the unit. We interpret the Santa Lucıá Formation to be composed of fluvial channel and overbank deposits. The lenticular, cross-stratified sandstones are typical fluvial channel deposits. The tabular, thin-bedded, rippled sandstones are probably crevasse-splay deposits, and the massive, nodular, spheroidally weathering siltstones are considered to be paleosols. Sempere et al. (1997) reported that the Santa Lucıá Formation contains abundant evaporites locally, which they interpreted as lacustrine playa deposits. We interpret the capping massive zone to be a series of sandy paleosols that grades into the overlying Impora Formation. Impora Formation The Impora Formation is exposed only in the Otavi and Camargo synclines. It is 60–80 m thick in these exposures (Fig. 4), and comprises five main lithofacies: (1) massive, gray and red mudrock with abundant pedogenic carbonate nodules; (2) massive, nodular limestone, locally with red and gray chert nodules; (3) thin, rooted, fine- to medium-grained, rippled and trough cross-stratified sandstone beds, some with abundant intraformational red chert and carbonate clasts; (4) laminated gray siltstone (Fig. 6A); and (5) gray marly limestone (Fig. 6A). Fossil fish, turtles, and crocodiles have been reported from the Impora (Gayet et al., 1991). The lower ;40–50 m of the Impora Formation are composed of massive, mottled siltstone with abundant calcareous nodules and clay skins. Pebbly sandstone beds are locally concentrated in the lower 25 m. The marly limestones and gray laminated siltstones are concentrated in the upper 11 m of the formation. All of the ;25 samples of the gray laminated siltstone facies that we processed for palynology were barren.

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We interpret the Impora Formation as a mixture of calcareous paleosols, lacustrine deposits, and minor fluvial deposits (see also Sempere et al., 1997). The lower ;50 m of the unit are dominated by calcareous paleosols and thin fluvial deposits. The upper marly and laminated siltstone-rich part of the unit is interpreted as profundal-lake, suspension fallout deposits (e.g., Carroll and Bohacs, 1999). The locally abundant chert nodules are interpreted as Magadi-type cherts (Eugster, 1967), probably indicating that lake waters were at least occasionally highly alkaline. High alkalinity could explain the paucity of palynomorphs and other microfossils in the laminated siltstones. Cayara Formation In the Camargo syncline, the Cayara Formation rests along a sharp contact on top of the Impora Formation (Figs. 2, 4). We place the base of the formation at a break in slope at the top of a laterally continuous gray, nodular and cherty limestone bed (Fig. 2). Above this bed is a 23-m-thick red siltstone interval with mottling but no carbonate nodules. The top of the unit is marked in the eastern limb of the Camargo syncline by a tabular, ledgeforming 12-m-thick trough cross-stratified sandstone unit. Where this prominent sandstone is absent along strike from our measured section south of Camargo (Cerro Tonka section), the top of the Cayara Formation is marked by a subtle change in mudrock color from lavender-red to the bright brick red of the overlying Camargo Formation. The total thickness of the Cayara Formation as defined here is 192 m. A distinguishing feature of the Cayara Formation is the presence of white to gray, fineto medium-grained quartzarenitic sandstone beds with abundant trough cross-strata and distinctive vertical burrows as long as 50 cm with prominent, horizontally arrayed chambers (Figs. 6B, 6C). These traces resemble Krausichnus trompitus, a diffuse termite nest composed of vertically stacked, elongate, and flattened chambers that form spindle-shaped structures (Hasiotis and Bown, 1992, their Fig. 7). Associated with these nests is a diffuse gallery system. K. trompitus is known to occur in Eocene–Oligocene paleosols developed in alluvial-channel and overbank sandstones in Egypt (S.T. Hasiotis, 2002, personal commun.). Some Cayara sandstone beds are completely riddled by these traces (Fig. 6C). Cayara sandstones are composed internally of multiple laterally overlapping lenticular bodies that amalgamate into 2–12-m-thick units

that can be traced laterally for many hundreds of meters. Whereas similar sandstones are present sporadically in the lower several hundred meters of the overlying Camargo Formation and locally in the Impora Formation, they are thickest and most abundant in the Cayara Formation. Mudcracks and siltstone intraclastic conglomerates are common in Cayara sandstones. Fine-grained rocks in the Cayara are dark red, purple, and lavender siltstones. These siltstones exhibit pervasive mottling and a lack of primary laminations. Nodular gypsum is locally abundant in Cayara mudrocks (Fig. 6D), and veins of gypsum are abundant in both the sandstones and siltstones. Like Sempere et al. (1997), we interpret the Cayara Formation as fluvial channel deposits, paleosols, and playa-lake deposits. Sempere et al. (1997) suggested that the Cayara Formation accumulated in a vast (5 3 105 km2) depocenter that received sediment mainly from the east. Our paleocurrent measurements indicate paleoflow strictly from the east (Fig. 4), which is consistent with the quartzarenitic composition and cratonal provenance of the unit. The nodular gypsum probably represents pedogenic accumulations in a highly evaporative environment, and the veins of gypsum formed during early diagenesis, perhaps by remobilization of disseminated primary gypsum and/or anhydrite. The abundant siltstone intraclastic conglomerates may have formed by erosional scouring of mud-crack drapes, and the pervasive deep burrowing in Cayara sandstones suggests long periods of channel abandonment and colonization by insects. The abundant evidence for sporadic flow between long periods of abandonment and high evaporation rates suggests that the Cayara was deposited in ephemeral streams and shallow playa lakes. Camargo and Suticollo Formations The Camargo Formation is .2000 m thick and holds up enormous cliffs along both limbs of the Camargo syncline (Figs. 2, 7A). Its total original thickness is unknown because the top is not preserved. In the two sections where we have measured the Suticollo Formation, it is 320 m (Ayo Ayo) and 840 m (Challa Mayu) thick, and it is capped in both locations by Miocene ignimbrites (Fig. 3). The Camargo and Suticollo Formations contain four principal lithofacies: (1) broad, lenticular, upward-fining, trough cross-stratified, medium- to coarse-grained or conglomeratic sandstone; (2) 0.5–10.0-m-thick beds of lenticular, well-organized, cross-stratified or subhorizontally stratified and imbricated, pebble



Figure 6. (A) Steeply dipping beds in the Impora Formation in the Otavi syncline. The lighter, resistant beds are lacustrine marl, and the darker, more recessive beds are laminated organic-rich siltstone. Section is ;6 m thick. (B) and (C) Krausichnus trompitus traces in trough cross-stratified sandstone of the Cayara Formation. Note the vertical shafts and subhorizontally arrayed chambers and connecting galleries. Lens cap is ;7 cm in diameter. (D) Gypsum nodules in red mudstone of the Cayara Formation. Knife (arrow) is 12 cm long.

to cobble conglomerate (Fig. 7B); (3) 0.25– 2.5-m-thick units of massive, mottled red siltstone, locally with pedogenic carbonate nodules; and (4) thin (0.05–0.20 m) laminated, clayey siltstones. The conglomerates are generally in the lower parts of upward-fining conglomeratesandstone units. The lower ;250 m of the Camargo Formation are dominated by brick-red siltstone

with thin, lenticular sandstone bodies that become increasingly abundant upward (Figs. 2, 4). A few beds of laminated claystone, rich in ostracodes and conchostracans, are present in the lower part of the Camargo Formation in the east limb of the Camargo syncline. The middle part of the Camargo Formation consists of packages of stacked, multistory sandstones (or pebbly sandstones) that form units

up to 50 m thick. The Suticollo Formation begins with medium- to coarse-grained sandstone and gradually coarsens upward into conglomeratic sandstone with isolated conglomerate bodies. As in the Camargo Formation, sandstone bodies are commonly amalgamated into laterally persistent cliff-forming units that are 10–40 m thick. In the Challa Mayu section, the uppermost part of the unit consists of an ;20-m-thick pebble to cobble conglomerate with well-developed imbrication (Fig. 7B). The typical Camargo and Suticollo Formation sandstone unit fines upward from an erosional base and contains a vertical sequence of trough cross-stratification, plane-parallel laminations, and ripples, followed gradationally upward by massive bioturbated sandstone (Horton and DeCelles, 2001). The massive sandstone interval grades upward into massive, mottled, red siltstone and then into a thin laminated siltstone unit below the next thick sandstone unit. A few of the massive siltstone units contain pedogenic carbonate nodules, but most are thoroughly leached of carbonate. In contrast, the laminated siltstones contain abundant disseminated carbonate. The Camargo and Suticollo Formations contain abundant evidence for deposition by large, bedload-dominated fluvial channels that were flanked by floodplains and zones of soil development (Horton and DeCelles, 2001). The dimensions of the channel fills and the types of sedimentary structures that they contain suggest peak discharges of ;103–104 m3/s. These were clearly large river systems. In a previous paper (Horton and DeCelles, 2001) we suggested that the Camargo Formation was deposited by fluvial megafans similar to those that drain the modern central Andean thrust belt. The ostracode-bearing laminated mudstones in the lower Camargo Formation were deposited in small, probably ephemeral ponds. Provenance and Paleocurrent Directions We obtained .1300 measurements of paleocurrent indicators from the Tertiary section of the study area (Figs. 4, 8). Most of the data are from the limbs of medium- to large-scale trough cross-strata (method I of DeCelles et al., 1983); a few tens of primary current lineations, erosional furrows, and clast imbrications were also measured. In the Santa Lucıá, Impora, and Cayara Formations, fluvial channels generally flowed westward and northwestward. In contrast, paleocurrent data from the Camargo and Suticollo Formations indicate a strong eastward mode. The presence of lacustrine deposits in the upper Impora For-


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Figure 7. (A) Photograph of the upper several hundred meters of the Cerro Tonka stratigraphic section, showing high cliffs in the thick sandstone and conglomerate beds of the upper Camargo Formation. (B) Well-organized pebble conglomerate in the upper Suticollo Formation. Hammer (center) is 35 cm long.

Figure 8. Paleocurrent data from the Tertiary units of the Camargo and Otavi synclines and the Potosı´ area. The data are all from limbs of large- to medium-scale trough crossstrata, with ;20 measurements per station according to method 1 of DeCelles et al. (1983).

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mation suggests that the basin was ponded briefly during roughly late Eocene time. The modal petrographic composition of Tertiary sandstone units in the Camargo syncline was determined by point-counting 59 thin sections of medium-grained sandstone. We counted 450 grains per slide by using a modified Gazzi-Dickinson method (Ingersoll et al., 1984). The principal modification that we used is in the identification of monocrystalline quartz grains that are part of sedimentary lithic fragments, such as quartzite or quartz sandstone. All thin sections were stained for potassium feldspar (K-feldspar) and calcium plagioclase. Table 1 summarizes the different grain types that were identified and the parameters that were normalized in ternary diagrams (Fig. 9). Table DR11 lists the recalculated data. Sandstones in the Santa Lucıá Formation are distinguished by their subarkosic composition, with average Qm/F/Lt 5 85/08/07 and Qt/F/L 5 89/08/03 (Table DR1 [see footnote 1]; Fig. 9). The feldspar fraction is dominated by orthoclase grains, many of which are microperthitic. The quartz fraction is dominated by unstrained monocrystalline quartz (Qm). The lithic grain fraction mostly consists of vitric volcanic grains and a few chert and mudstone grains. The sandstones of the Impora Formation are nearly pure quartzarenites, with average Qm/ F/Lt 5 98/00/02 and Qt/F/L 5 99/00/01. Sandstones from the Cayara Formation are only slightly less quartzose; they have average Qm/F/Lt 5 94/00/06 and Qt/F/L 5 97/00/03 (Table DR1; Fig. 9). Trace amounts of limestone, shale, chert, and vitric volcanic grains are the only lithic grains in Impora and Cayara sandstones. Pebbles of intraformational chert are common in the lower part of the Impora Formation. Sandstones in the Camargo Formation have average Qm/F/Lt 5 83/01/16 and Qt/F/L 5 93/01/06 (Table DR1; Fig. 9). These sandstones contain significant quantities of lithic grains, mainly quartz tectonite, polycrystalline quartz, quartzose sandstone/quartzite fragments, chert, limestone, phyllite, and volcanic lithic fragments. The latter consist of felsic, vitric, and microlitic varieties. Conglomerates in the Camargo Formation contain clasts of gray phyllite, pink/purple/gray sandstone, white quartzite, fossiliferous gray limestone, 1 GSA Data Repository item 2003008, Table DR1. Implications of Foreland basin development for the history of Andean orogenesis, is available on the Web at http://www.geosociety.org/pubs/ft2003.htm. Requests may also be sent to editing@geosociety. org.


FORELAND BASIN DEVELOPMENT AND HISTORY OF ANDEAN CRUSTAL SHORTENING, BOLIVIA TABLE 1. PETROGRAPHIC PARAMETERS Quartzose grains Qm Qp Qpt Qss Silt Qt Feldspar grains K Kper Kmic Kmy P F Lithic grains Metamorphic Lsp Lm Sedimentary Lc Lsm Lcht Ls Volcanic Lvl Lvx Lvf Lvv Lvm Lv Lt Accessory Grains

Monocrystalline quartz Polycrystalline quartz Polycrystalline quartz with foliation Monocrystalline quartz in sandstone fragment Siltstone Total quartzose grains (5 Qm 1 Qp 1 Qpt 1 Qss 1 silt) Potassium feldspar Potassium feldspar with perthitic intergrowths Microcline Myrmekitic potassium feldspar Plagioclase feldspar Total feldspar grains (5 K 1 Kper 1 Kmic 1 Kmy 1 P)

Phyllite grains Total metamorphic lithic grains (5 Lsp 1 Qpt) Carbonate grains (including fossil fragments) Shale/mudrock grains Chert grains Total sedimentary lithic grains (5 Qss 1 silt 1 Lc 1 Lsm 1 Lcht) Lathwork volcanic grains Microlitic volcanic grains Felsic volcanic grains Vitric volcanic grains Mafic volcanic grains Total volcanic lithic grains (5 Lvl 1 Lvx 1 Lvf 1 Lvv 1 Lvm) Total lithic grains (5 Lm 1 Ls 1 Lv) Alterite (Fe-oxide aggregates) Tourmaline Epidote Zircon Magnetite

porphyritic andesite/dacite, and gray and red chert. Sandstones from the Suticollo Formation are similar in composition to those from the Camargo Formation; they have average Qm/F/Lt 5 83/02/15 and Qt/F/L 5 93/02/05 and a lithic-grain population dominated by low-grade metasedimentary and sedimentary grains (Table DR1; Fig. 9). The key petrographic feature of Santa Lucıá Formation sandstones is their high K-feldspar content (Fig. 10). Santa Lucıá fine-grained rocks also contain large amounts of feldspar (Deconinck et al., 2000). Possible sources of these grains include Precambrian cratonic basement rocks to the east of the Andean foldthrust belt, granitic igneous rocks of the Mesozoic magmatic arc to the west of the Altiplano, and granitic Precambrian basement rocks that are exposed in the Arequipa terrane of coastal southern Peru and scattered along the Western Cordillera or buried beneath the Tertiary of the western Altiplano (Isaacson, 1975; Lehmann, 1978; Wo¨rner et al., 2000). The western sources, however, are not supported by the paleocurrent data and would have required a complex dispersal path to deliver sediment into the generally westwardflowing Santa Lucıá fluvial system. The rocks

of the Subandean zone and Eastern Cordillera contain no obvious candidates for a source of K-feldspar, with the possible exceptions of an Ordovician–Silurian diamictite (Cancan˜iri Formation) that contains clasts of coarsegrained plutonic rocks (Sua´rez, 1995) and local Permian–Triassic granites in the Eastern Cordillera of northwestern Bolivia. Small granitic and granodioritic intrusions of Cretaceous age are scattered along the Eastern Cordillera in southern Bolivia and northern Argentina (Viramonte et al., 1999), but these may not have been exposed in early Tertiary time. Thus, it seems most likely that K-feldspar in the Santa Lucıá Formation was derived from the Brazilian Shield to the east. Deconinck et al. (2000) suggested that abundant swelling clay in the Santa Lucıá Formation in the Potosı´ region could have been derived from the Paleocene magmatic arc to the west. This possibility is consistent with the presence of volcanic lithic grains in the Santa Lucıá sand fraction. It is conceivable that airborne ash derived from the west could have mixed with fluvial bedload derived from the Brazilian Shield. The simplest explanation for the provenance of the sand-sized volcanic lithic grains in the Santa Lucıá Formation is that

they were derived from the magmatic arc to the west, however, this explanation is not supported by our paleocurrent data. Only two of our 22 paleocurrent stations in the Santa Lucıá Formation yielded an eastward paleoflow direction (Fig. 8). The highly quartzose compositions of sandstones in the Impora and Cayara Formations are consistent with the lithofacies data that indicate intense weathering (particularly during Impora deposition) over a prolonged period of time. In addition, although the petrographic data by no means prove the point, they support our contention that the Impora and Cayara Formations were deposited over a prolonged period of time (Fig. 4). Modal sandstone compositions of the Camargo and Suticollo Formations are typical of foreland basin sandstones, plotting along the quartz-lithics binary with abundant recycled sedimentary lithic grains (Fig. 9; e.g., Dickinson and Suczek, 1979; Schwab, 1986; Lawton, 1986; DeCelles and Hertel, 1989). The sedimentary and metasedimentary lithic grains and conglomerate clasts (sandstone/ quartzite, limestone, chert, and phyllite) in the Camargo Formation were derived from Ordovician, Jurassic, and Cretaceous rocks in the Eastern Cordillera of the developing Andean fold-thrust belt. The volcanic lithic grains could have been derived from Tertiary volcanic rocks of the magmatic arc and hypabyssal intrusions and flows of Jurassic–Cretaceous age that are locally present in the Altiplano and Eastern Cordillera. For example, a volcanic clast from the upper Camargo Formation at Cerro Tonka yielded an 40Ar/39Ar age of ca. 91 Ma (Horton, 1998b). To summarize, the overall trend of sandstone compositions in the Tertiary succession in the southern part of the Bolivian Eastern Cordillera is from subarkosic (Santa Lucıá Formation) to quartzarenitic (Impora and Cayara Formations) to sublitharenitic (Camargo and Suticollo Formations) (Figs. 9, 10). Potassium feldspar and volcanic lithic grains are relatively common in the Santa Lucıá Formation, but from Impora deposition onward, feldspar is rare, and the lithic fraction is dominated by sedimentary and low-grade metasedimentary grains (Fig. 9). In light of the paleocurrent data (Figs. 4, 8), the overall compositional trend can be interpreted as a change from dominantly cratonic provenance (with possible minor additions from the magmatic arc) during deposition of the Santa Lucıá Formation to a fold-thrust belt provenance during deposition of the Camargo and Suticollo Formations.


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ANDEAN FORELAND BASIN DYNAMICS The Foreland Basin System Model

Figure 9. Ternary diagrams illustrating modal framework-grain compositions of 59 sandstone samples. See Table 1 for definitions of symbols and Table DR1 [see footnote 1] for recalculated data. Individual samples from the Santa Lucıá, Cayara, and Impora Formations are plotted, and dotted areas encompass the distribution of Camargo and Suticollo Formation samples. The base of the QmFLt diagram represents the 50% Qm value, whereas the base of the LmLvLs diagram is 0% Lm. Raw data are available from the authors.

Many modern foreland basin systems, including the one along the eastern margin of the modern Andean fold-thrust belt (Horton and DeCelles, 1997; Ussami et al., 1999), comprise four discrete depositional zones (Fig. 11; DeCelles and Giles, 1996). The frontal part of the orogenic wedge is buried by syndepositionally deformed wedge-top sediments, characterized by coarse texture and growth structures. The wedge-top depozone merges toward the craton with the foredeep depozone, the extremely thick, rapidly subsiding part of the system that decreases in thickness toward the craton. The sediments of the distal foredeep merge with the increasingly condensed (i.e., time-rich), thin veneer of sediments of the forebulge depozone, or with the unconformity that is produced by forebulge uplift and lateral migration (e.g., Crampton and Allen, 1995; White et al., 2002). Cratonward of the forebulge, some foreland basin systems include a region of minor but widespread subsidence and aggradational sedimentation (up to or above the crest of the forebulge); this region is referred to as the

Figure 10. Diagrams illustrating the up-section trends in percent feldspar relative to the total monomineralic fraction in Tertiary sandstones of the Eastern Cordillera.

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Figure 11. (A) Schematic diagram illustrating the depozones of a foreland basin system (after DeCelles and Giles, 1996). Lower surface is an ideal elastic flexural profile; dotted area is basin fill; and upper surface is a logarithmic aggradational surface tangent to the crest of the forebulge. In actual foreland basin systems, this surface may be above or below the crest of the forebulge. (B) Schematic vertical juxtaposition of depozones in a complete foreland basin system that would result from steady, lateral migration of the flexural profile shown in A. The accompanying curve (dS/dt, where S 5 subsidence distance and t 5 time) shows the idealized subsidence history calculated from the first derivative of the flexural profile in A. To the left in B is shown a schematic log of the stratigraphic section in the Camargo syncline, and connecting lines show our interpretations of corresponding foreland depozones. Note that only the upper and lower age constraints are depicted on the Camargo succession, and no attempt is made to assign real ages to the diagram at right, which is an idealized model. back-bulge depozone (e.g., Giles and Dickinson, 1995; Currie, 1997). Sediment in the back-bulge depozone may be derived from both the fold-thrust belt (and magmatic arc if one is present) and the craton. The foredeep, forebulge, and back-bulge depozones are predicted by elastic flexural modeling, which shows that depozone widths are generally subequal and depend upon the flexural rigidity of the foreland lithosphere. Gravity data suggest that the flexural rigidity of the Brazilian Shield that is underthrusting the An-

dean fold-thrust belt is 3 3 1023 to 5 3 1023 N·m (Lyon-Caen et al., 1985; Goetze and Schmidt, 1999), corresponding to predicted widths of each depozone of 200–350 km. In southern Bolivia, the wedge-top depozone of the modern foreland basin system is ;50–100 km wide, and the foredeep, forebulge, and back-bulge depozones are each ;250–350 km wide (Horton and DeCelles, 1997). The thickness of sediment in the back-bulge depozone may be controlled by several factors. Whereas flexural models predict only ;100–250 m

(depending on the flexural rigidity) for a backbulge depozone that is graded to the crest of the forebulge, ‘‘overfilling’’ of the foredeep, along-strike input, and regional dynamic subsidence may produce thicker back-bulge accumulations (DeCelles and Currie, 1996). As it underthrusts the fold-thrust belt, the foreland lithosphere moves through the standing wave set up by the orogenic load. If the distance of underthrusting of the foreland lithosphere is equal to or greater than the wavelength of the flexural profile, then the deposits of the four depozones will be stacked vertically in an upward-coarsening, several-kilometersthick succession (DeCelles and Currie, 1996; Sinclair, 1997). A long-term subsidence (or sediment-accumulation) curve from this succession will exhibit a sigmoidal shape (Fig. 11B) with initial slow subsidence in the backbulge depozone, followed by development of a major forebulge unconformity or zone of stratigraphic condensation, followed by increasingly rapid subsidence as the foredeep moves through the region, and finally, by reduced and erratic subsidence and uplift with the arrival of the frontal orogenic wedge. Given the range of known flexural rigidities and velocities of underthrusting in continental fold-thrust belts (DeCelles and DeCelles, 2001), the duration of the forebulge unconformity or condensed interval would be expected to be 10–30 m.y. We emphasize that this concept of a foreland basin system is highly simplified and does not consider the short-term changes in subsidence rate and local structural features that complicate many foreland regions (e.g., Schwartz and DeCelles, 1988; Zaleha et al., 2001). Reactivation of older faults and fractures, as well as development of new faults due to tensile fiber stresses along the upper arc of the forebulge, are common in forebulge depozones (e.g., Jacobi, 1981; Bradley and Kidd, 1991; Meyers et al., 1992; Crampton and Allen, 1995; Lihou and Allen, 1996; Lorenzo et al., 1998; Ussami et al., 1999; Zaleha et al., 2001). In addition, some foreland regions lack a well-developed back-bulge depozone, and even in cases where a back-bulge depozone is present, it may not be preserved owing to erosion as it passes over the crest of the forebulge (Allen et al., 2001). Foreland basin depozones may change markedly through time as the character of the foreland lithosphere entering the flexed region of the system changes (e.g., Cardoza and Jordan, 2001). For example, it might be expected that the backbulge depozone could be relatively well developed on outboard, less rigid lithosphere, but poorly developed as stiffer cratonic litho-


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sphere enters the flexed region. Finally, it should be pointed out that the kinematic and erosional histories of fold-thrust belts may be highly unsteady, producing spatial and temporal changes in the subsidence history of the foreland basin system. Applying the Model to the Eastern Cordillera Tertiary Succession The Tertiary stratigraphic section in the Eastern Cordillera of southern Bolivia, particularly in the Camargo syncline region, compares favorably with the predicted pattern of sediment-accumulation rates and lithofacies for the back-bulge, forebulge, and foredeep depozones (Fig. 11B). The distal fluvial and local lacustrine deposits of the Santa Lucıá Formation can be interpreted as back-bulge deposits. The Santa Lucıá forms a widespread sheet of fluvial and lacustrine deposits with a palinspastic distribution of .100,000 km2 (Sempere et al., 1997), including the Eastern Cordillera and eastern Altiplano. Its thickness over this broad region changes little, with the possible exception of the Otavi anomaly. The depositional axis of this saucer-shaped (in cross section) depozone was generally located in the western half of the Eastern Cordillera (Cherroni, 1977; Sempere, 1994; Sempere et al., 1997). Most of the Santa Lucıá sand was transported from the east in rivers that presumably headed in the western Brazilian Shield and perhaps in Paleozoic and Mesozoic rocks that are now involved in the Subandean part of the fold-thrust belt (Fig. 12A). Minor amounts of Santa Lucıá sand may have been derived from the magmatic arc to the west, suggesting that some rivers flowing from the west were capable of transecting, or bypassing along strike, the crest of the forebulge. The uppermost Santa Lucıá Formation and the majority of the Impora Formation are overprinted by intense pedogenesis. In the context of the chronostratigraphic scheme proposed in Figure 3, we suggest that this zone of pedogenesis may have developed as the flexural forebulge migrated through the region over a period of ;15–20 m.y. (Fig. 12B). As the Santa Lucıá lithosome was carried up onto the crest of the forebulge, the rate of deposition was reduced to values as low as ;0.0025 mm/yr, two orders of magnitude slower than typical accumulation rates in foredeep depozones. Although it is plausible that some net erosion took place, the upward transition from the uppermost Santa Lucıá Formation into the lower Impora Formation is gradational, suggesting that in the study area the forebulge depozone was not a surface of significant ero-

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sional beveling. However, the observation that the Cayara Formation overlaps pre-Tertiary rocks in parts of the Eastern Cordillera (Sempere et al., 1997) suggests that the forebulge crest was locally beveled. This scenario is remarkably similar to the stratigraphic record in the Altiplano to the west, where ;15–20 m.y. is represented within a 20–100-m-thick paleosol zone of middle Paleocene to middle Eocene age (Horton et al., 2001). The lacustrine facies in the uppermost Impora Formation may represent the onset of more rapid subsidence as the distal end of the foredeep depozone migrated into the Camargo region (Fig. 12C). Structural, stratigraphic, and thermochronologic data indicate that the Eastern Cordillera began to be deformed and exhumed during late Eocene–Oligocene time (McBride et al., 1987; Farrar et al., 1988; Sempere et al., 1990; Lamb and Hoke, 1997; Horton, 1998a; McQuarrie and DeCelles, 2001; Horton et al., 2002). Crustal thickening in the Eastern Cordillera may have driven rapid flexural subsidence that temporarily outpaced the ability of regional depositional systems to fill the resulting accommodation space. The sandstones in the Cayara Formation were also derived from the east, suggesting that the distal foredeep continued to receive sediment from the forebulge and/or craton for some time after passage of the forebulge through the Camargo region. The Cayara fluvial system was eventually overwhelmed by eastward progradation of the Camargo Formation, which heralded the reversal of regional drainage directions and arrival of the central to proximal foredeep depozone in the Camargo region (Fig. 12D). The Camargo and Suticollo Formations also contain the first clearcut provenance evidence for a major orogenic source terrane to the west (Fig. 9). We have found no evidence for wedge-top deposits (i.e., growth structures) in the Tertiary strata of the study area, possibly because they have been eroded. The absence of a wedge-top depozone in the study area implies that the structural front of the orogenic wedge lay to the west of this longitude until after deposition of the preserved parts of the Camargo and Suticollo Formations (i.e., earliest Miocene time). In the modern fluvial megafans of the Himalayan foreland basin system, the gravel front generally does not extend farther than ;75 km from the topographic front of the fold-thrust belt (Wells and Dorr, 1987; DeCelles and Cavazza, 1999). Thus, the topographic front probably lay a few tens of kilometers to the west of Camargo at the time of deposition of the upper Camargo and Su-

ticollo Formations. This scenario is consistent with previous interpretations that the western part of the Eastern Cordillera was structurally elevated and shedding voluminous sediment eastward into the foreland basin and westward into the Altiplano basin by Oligocene time (Sempere et al., 1990; Horton et al., 2002). A candidate for the preserved wedge-top depozone during Oligocene–early Miocene time is the Tupiza region (Fig. 1), where thick coarsegrained alluvial-fan facies and growth structures are present (Horton, 1998a). By Oligocene–early Miocene time, thrusts in the Tupiza region had produced high local topographic relief (He´rail et al., 1996; Tawackoli et al., 1996; Horton, 1996, 1998a; Kley et al., 1997). Thrusting in the Subandean zone was under way by late Miocene time (Gubbels et al., 1993; Reynolds et al., 1996, 2000; Jordan et al., 1997). DISCUSSION Comparison with Regional Stratigraphic Patterns The interpretation shown in Figures 11 and 12 differs in two aspects from that of Sempere et al. (1997), who broadly interpreted the Upper Cretaceous and Tertiary succession in the Eastern Cordillera as synorogenic foreland basin deposits. First, we are suggesting a refinement of this general interpretation by differentiating the succession in terms of separate foreland depozones. Second, our placement of the late Paleocene–Eocene condensed interval and the duration that we infer it represents are different from the interpretations offered by Sempere et al. (1997) and are crucial for the model we are advocating. Sempere et al. (1997) proposed a major unconformity at the base of the Cayara Formation, a few meters up section from where we locate the top of the zone of stratigraphic condensation in the uppermost Santa Lucıá and lower Impora Formations; however, Sempere et al. assigned no time gap to the unconformity. Sempere et al. (1997) interpreted the Santa Lucıá and Cayara Formations as having been deposited in rapid succession on the west side of a forebulge, whereas we suggest that they are separated by ;20 m.y. of time and were deposited on opposite sides of the forebulge before (Santa Lucıá Formation) and after (Cayara Formation) it swept through the region over that time span (Fig. 12). Although either interpretation could be regarded as valid until better age data become available, the abundant lithologic evidence for extreme condensation at this stratigraphic level, the widespread distribution of this interval, and documentation of ;15–20


Figure 12. Series of schematic, east-west cross sections across the Andean fold-thrust belt and the adjacent foreland basin system from Paleocene through early Miocene time, illustrating the proposed history of foreland basin migration as discussed in the text. Vertical dotted line indicates approximate location of Camargo syncline. Large arrows indicate paleocurrent polarity, solid where documented and dashed where inferred (see Figs. 3 and 8 for data).



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m.y. of greatly reduced sedimentation within a similar paleosol zone in the Altiplano (Horton et al., 2001) lead us to argue that the condensed interval is attributable to a large-scale geodynamic mechanism. Our interpretation is consistent with previous work that suggests that the western part of the Eastern Cordillera was undergoing uplift during late Eocene and Oligocene time. This eastward advance of the front of the orogenic wedge, from a proposed location west of the Altiplano, involved the growth of the west-vergent Central Andean backthrust belt, uplift of the crestal region of the present Eastern Cordillera, and development of the Altiplano basin (Roeder, 1988; Sempere et al., 1990, 1997; Lamb and Hoke, 1997; Horton et al., 2001, 2002; McQuarrie and DeCelles, 2001; Ege et al., 2001; McQuarrie, 2002). The foredeep shifted eastward to the eastern part of the Eastern Cordillera during deposition of the Camargo and Suticollo Formations. The stratigraphic manifestation of the major eastward propagation of the thrust front is the abrupt change from forebulge deposition (uppermost Santa Lucıá and majority of the Impora Formations) to lacustrine, ‘‘underfilled’’ foredeep deposition in the uppermost Impora Formation (Fig. 12C). If our interpretation is correct, it should be reconcilable with the middle to upper Tertiary stratigraphic record of foreland basin development in the Subandean zone and Chaco Plain (to the east of the Subandean zone; Figs. 1, 13). In the western Subandean zone of southern Bolivia, a several-kilometers-thick, upward-coarsening fluvial succession rests disconformably upon Cretaceous strata (e.g., Dunn et al., 1995; Jordan et al., 1997). The lowest part of the succession is probably of late Oligocene age (Erikson and Kelley, 1995; Jordan et al., 1997). The Tertiary section of the eastern Subandean zone consists of a similar upward-coarsening fluvial succession that is interrupted by a middle Miocene marine interval (Yecua Formation; Marshall and Sempere, 1991; Gubbels et al., 1993; Dunn et al., 1995; Moretti et al., 1996; Erikson and Kelley, 1995; Jordan et al., 1997). Some authors include in the upper part of this succession the late Miocene Intra-Chaco Discordance, a subtle, regionally extensive angular unconformity (Gubbels et al., 1993, and references therein; Fig. 13) possibly related to structural growth in the subsurface (Dunn et al., 1995). The age of the base of the Tertiary succession in the eastern Subandean zone is late Oligocene–early Miocene (29–21 Ma), based on Deseadan mammal fossils in its lower part (Marshall et al., 1993). The Tertiary section beneath the

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Figure 13. Diagram illustrating the regional evolution of the Andean foreland basin system plotted on a palinspastic base, with the foreland basin depozones anchored by our interpretations of the Camargo syncline succession. Stratigraphies of the Eastern Cordillera (EC), western Subandes (WSA), eastern Subandes (ESA), and Chaco Plain are shown at right. ICD represents the Intra-Chaco Discordance. The two flexural profiles shown in the graph at the top left are derived from flexing elastic lithosphere of differing rigidities, as indicated. The palinspastic panels show locations of Tupiza (T), Camargo (C), and the Subandean zone. Diagonally ruled area is the forebulge depozone through time. In this depiction, the flexural profile is held fixed while the palinspastic panel is moved progressively westward through time. Note that the Tupiza-Camargo panel begins to shorten after ca. 30 Ma. See text for discussion.

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Chaco Plain is similar to, but slightly younger than, that exposed in the eastern Subandean zone (Dunn et al., 1995; Fig. 13). From the viewpoint of our model for the Andean foreland basin system, the basal Tertiary disconformity in the Subandean zone and beneath the Chaco Plain may be reasonably interpreted as a result of forebulge migration within the resolution of the available chronostratigraphic data. If this is the case, then no back-bulge depozone is preserved in the Subandean zone of southern Bolivia. Possible explanations for the absence of a back-bulge depozone include erosion during passage of the forebulge or damping out of the flexural wave during Oligocene time owing to a lateral increase in flexural rigidity of the Brazilian Shield lithosphere. The upward-coarsening succession above the basal disconformity is considered to be typical foredeep deposits (Jordan et al., 1997). By analogy with the Subandean record from northern Argentina (e.g., Hernańdez et al., 1996, 1999; Reynolds et al., 2000), we interpret the Bolivian record to reflect eastward migration of initial foredeep conditions during late Oligocene–early Miocene time. Structural evidence and angular stratigraphic relationships in seismic data from the eastern Subandean zone and Chaco Plain suggest that development of growth structures associated with the wedge-top depozone commenced during late Miocene time (Gubbels et al., 1993; Dunn et al., 1995; Moretti et al., 1996; Horton and DeCelles, 1997). A test of this proposed scheme of regional tectono-sedimentary correlation is shown in Figure 13. We calculated a pair of flexural profiles by using a 3-km-thick rectangular load with a 150 km half-width on an unbroken elastic plate, with flexural rigidities of 3 3 1023 N·m and 5 3 1023 N·m. These rigidities correspond to elastic thicknesses of 35 and 45 km, respectively, and are comparable to the modern values obtained in gravity modeling studies of the Subandean zone and Chaco Plain (Lyon-Caen et al., 1985; Watts et al., 1995; Goetze and Schmidt, 1999). McQuarrie (2002) provided palinspastic widths of the Subandean zone (206 km) and the eastern half of the Eastern Cordillera (330 km). The locations of the Tupiza and Camargo areas in the Eastern Cordillera are plotted at their locations on this undeformed panel, and the east-west location of the panel was allowed to slide progressively westward across the modeled flexural profiles, simulating the eastward migration of the flexural profile through the Andean foreland region prior to arrival of the fold-thrust belt. The locations of the panel in the context of the flexural profile at 10 m.y.

time steps are fixed by the interpretations of the Tertiary succession where it is most complete in the Camargo area (Fig. 13). For example, the Camargo area was located in the forebulge depozone at the 50 Ma time step because the upper Santa Lucıá–Impora Formation transition is interpreted as forebulge deposits (Fig. 12). The Camargo section was located in the foredeep depozone during deposition of the Camargo Formation. By 30– 20 Ma the tip of the orogenic wedge had arrived in the Tupiza region (Horton, 1998a), and the Camargo section lay a few tens of kilometers to the east. Between ca. 20 and 10 Ma, the front of the fold-thrust belt migrated into the Subandean zone (Gubbels et al., 1993; McQuarrie, 2002; Fig. 12E). The palinspastic reconstruction shows that the forebulge depozone could have migrated from the western Subandean zone into the eastern Subandean zone during late Eocene to early Miocene time, producing the basal Tertiary disconformity in the process. By late Miocene time, the forebulge had migrated eastward into the Chaco Plain (Fig. 1, inset), ultimately arriving where it is today ;350 km from the front of the fold-thrust belt (Horton and DeCelles, 1997) in a region of minor topography and extensive lateritic soil development (Litherland and Pitfield, 1983; Litherland et al., 1986). The model is deliberately simplistic because there is no way to reliably estimate past values of flexural rigidity and load size. Some minor temporal increase in the flexural rigidity of the underthrusting Brazilian lithosphere is expected to have occurred as older, colder cratonic lithosphere encountered the front of the fold-thrust belt. This increase in rigidity would have had the effect of increasing the wavelength of the flexural profile and possibly damping out the back-bulge depozone. Alternative Model: Regional Extension It has been recently suggested that all of the central Andes was a region of extension during late Oligocene–early Miocene time, possibly triggered as a complex response to an increase in the rate and orthogonality of convergence between the Nazca and South American plates (Jordan et al., 2001). This interpretation is based in part on sedimentological and geochronological data from the arc and forearc regions in Argentina and Chile .1500 km south of our study area (e.g., Burns and Jordan, 1999). Insofar as extension and basin development are common attributes of magmatic arcs (e.g., Smith and Landis, 1995; Marsaglia, 1995), it is neither paradoxical nor con-

trary to our model that extension in the coastal belt may have taken place contemporaneously with shortening in the fold-thrust belt to the east. Rather, we expect that the middle Tertiary arc and forearc region was subject to various amounts of extension, contraction, and strike-slip deformation as a result of factors that do not bear simple relationships to convergence rates and orthogonality (e.g., Hartley et al., 2000). At the latitude of Bolivia, the unequivocal compositional and lithofacies evidence for Oligocene–Miocene foreland basin development in the Eastern Cordillera (Tawackoli et al., 1996; Sempere et al., 1997; Horton, 1998a; this paper), results of structural studies (McQuarrie, 2002), and late Eocene– Oligocene apatite fission-track ages from the Eastern Cordillera (Ege et al., 2001) provide strong evidence that regional shortening and uplift had commenced in the Eastern Cordillera by late Eocene–Oligocene time. The evidence reported in this paper for the existence of a forebulge in the Eastern Cordillera, when combined with the record from the eastern Altiplano (Horton et al., 2001), pushes the onset of foreland basin development into late Paleocene time. The argument about timing of foreland basin development (and associated thrusting) is less clear-cut for Late Cretaceous and early Paleocene time in southern Bolivia (see discussion in Sempere, 1994). Particularly difficult to assess is the tectonic regime during deposition of the El Molino and Santa Lucıá Formations; data presented in this paper, as well as those presented by Sempere et al. (1997) and Deconinck et al. (2000), do not provide direct evidence for the existence of a foreland basin system prior to development of the late Paleocene–Eocene condensed interval. A possible alternative interpretation is that Late Cretaceous–Paleocene deposition in the Eastern Cordillera was controlled by postrift thermal subsidence in an extensional setting north of the Salta rift system of Argentina (Salfity and Marquillas, 1994; Welsink et al., 1995; Viramonte et al., 1999). It is conceivable that thermal subsidence controlled deposition of the ;1–2-km-thick Upper Cretaceous–Paleocene section (Aroifilla through Santa Lucıá Formations, representing ;35 m.y.) in Bolivia, but the small amount of documented extension (;20 km of maximum regional extension; Grier et al., 1991) and limited geochemical evidence for crustal thinning (Lucassen et al., 1999) in the Salta rift itself suggest that crustal stretching was minor. Moreover, the gradually decreasing subsidence rates predicted by simple rift models (e.g., McKenzie, 1978) are in direct conflict


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with variable sedimentation rates of the El Molino–Santa Lucıá section (as calculated by Sempere et al., 1997) and with the abrupt cessation of subsidence in the Altiplano (Horton et al., 2001) and inferred for the upper Santa Lucıá–Impora paleosol zone in the Eastern Cordillera (this paper). A number of recent studies have documented Paleogene crustal shortening and tectonic (vertical-axis) rotations in the arc and forearc regions to the west of the Altiplano (Chong and Reutter 1985; Hartley et al., 1992; Hammerschmidt et al., 1992; Scheuber and Reutter, 1992; Scheuber et al., 1994; Amilibia et al., 1999; Nicolas et al., 1999; Roperch et al., 2000a, 2000b), lending support to the Paleogene time frame of our model. Ultimately, the question settles on the issue of the back-bulge depozone. Because the back-bulge region is in many cases isolated from direct sediment input from the foldthrust belt, its sediments may be characterized by cratonic provenance and therefore may be difficult to distinguish from passive-margin deposits or postrift sediments. In some cases, back-bulge deposits can be directly tied to the fold-thrust belt by virtue of their provenance and contemporaneity with known dynamothermal events in the orogenic belt (e.g., Camilleri et al., 1997; DeCelles et al., 1998). In the case of the Santa Lucıá Formation, it is plausible that multiple mechanisms operated in concert to generate sediment accommodation in the back-bulge depozone, including minor flexural subsidence (e.g., Flemings and Jordan, 1989; DeCelles and Giles, 1996), dynamic subsidence driven by viscous coupling of South American lithosphere with the subducting Nazca plate (e.g., Mitrovica et al., 1989; Gurnis, 1992), and limited thermal subsidence associated with extension that was better developed in northern Argentina and northern Chile during the Late Cretaceous (e.g., Mpodozis and Allmendinger, 1993; Viramonte et al., 1999). Tectonic Implications We point to two major implications revealed by the preceding modeling exercise. First, the migration of the foreland basin system as discussed in the context of the Camargo section is regionally consistent with the Tertiary stratigraphic record of the eastern part of the Eastern Cordillera, the Subandean zone, and the Chaco Plain of southern Bolivia. Data from the Altiplano also support the overall geodynamic model (Horton et al., 2001). For the entire Tertiary section in the Camargo syncline to have been deposited in the foredeep, as suggested by previous workers (e.g., Lamb

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and Hoke, 1997; Sempere et al., 1997), the foredeep would have to have remained stationary for ;40 m.y. In addition, such a foredeep would have to have encompassed the entire palinspastic distribution of the Tertiary section in the Eastern Cordillera and Altiplano, an original east-west distance of ;450 km (McQuarrie, 2002). On the basis of the continuity of convergence between the Nazca and South American plates (Pardo-Casas and Molnar, 1987; Somoza, 1998; Norabuena et al., 1999) and reasonable flexural rigidities of Brazilian lithosphere, we regard both of these possibilities as unrealistic. The second, and more significant, implication of the modeling exercise is that the total lateral migration distance of the foreland basin system must have been ;1000 km. This value derives from the placement of the palinspastically restored panels on the flexural profile, which indicates that the eastern Subandean zone lay ;1000 km from the edge of the orogenic load in late Paleocene time (Fig. 13). Although this distance may have been less if the flexural rigidity during Paleocene time was lower, it is difficult to reduce the distance by more than ;200 km while still maintaining reasonable loads and rigidities and matching the regional stratigraphic patterns. The potential uncertainties in the palinspastic panels are on the order of a few tens of kilometers (McQuarrie, 2002). Thus, the length of Brazilian Shield lithosphere that has underthrust the Andean fold-thrust belt since late Paleocene time is also ;1000 km. This amount of underthrusting includes both the distance of actual crustal ‘‘subduction’’ and the distance of eastward propagation of the Andean foldthrust belt over the top of the underthrust Brazilian Shield (DeCelles and DeCelles, 2001). Therefore, the actual amount of shortening is approximately half of the estimated distance of foreland basin migration, or ;500 km. The estimated rates of flexural wave migration (;12–20 mm/yr; Fig. 13) should equal the propogation rate plus the rate of convergence between the fold-thrust belt and the Brazilian craton. Although we do not place much confidence in these rates because of the uncertainties in the flexural profile and the chronostratigraphy, their values are consistent with present rates of convergence between the foldthrust belt and the Brazilian Shield (Norabuena et al., 1998). The recent literature on Andean tectonics is peppered with attempts to explain an apparent deficit of crustal shortening necessary to account for the present thickness of the crust across the central Andes (Isacks, 1988; Schmitz, 1994; Baby et al., 1997; Schmitz and Kley, 1997; Kley and Monaldi,

1998; James and Sacks, 1999; McQuarrie, 2002). To produce the crustal cross section of Beck et al. (1996) by horizontal shortening alone would require ;500 km of shortening, depending on the original thickness of the crust in the Brazilian Shield (see also McQuarrie, 2002). Our analysis suggests that the amount of crust (and South American mantle lithosphere) that has been underthrust beneath the Andes is sufficient to explain the current crustal thicknesses. Excess South American lithosphere could have been disposed of by corner-flow erosion in the mantle wedge (Owens and Zandt, 1997), delamination (Beck and Zandt, 2002), and/or subduction erosion (‘‘ablative subduction’’ of Pope and Willett, 1998) along its contact with the subducting Nazca plate. The apparent mismatch between our estimate of total shortening and estimates based on balanced cross sections can be explained as a result of the fact that potentially large amounts of Paleogene shortening in the western Altiplano, Western Cordillera, and farther west are simply not included in total shortening estimates because the necessary structural information does not yet exist. That large amounts of shortening could have occurred in such hinterland regions comes as no surprise when the Andes are compared with other major retroarc fold-thrust belts, particularly the North American Cordillera (Price, 1981; Camilleri et al., 1997). CONCLUSIONS 1. A migratory foreland basin system, formed in response to Andean crustal shortening and loading, can be traced back at least into late Paleocene time in the Eastern Cordillera of southern Bolivia. The existence of a fully developed foreland basin system by that time suggests that major amounts of early Tertiary shortening took place in the central Andes. 2. Extrapolation of the tectonostratigraphic foreland basin scheme developed for the Eastern Cordillera into the Subandean zone and Chaco Plain to the east demonstrates that the Andean foreland basin system has migrated progressively eastward to its present location at long-term rates of ;12–20 mm/yr over a time span of .55 m.y. The early history of foreland basin migration is uncertain, but a major eastward jump of the front of the foldthrust belt into the Eastern Cordillera occurred by late Eocene–Oligocene time, as indicated by this study and previous studies. Eastward migration of the fold-thrust belt and foreland



basin appears to have been more continuous during Neogene time. 3. The total lateral distance of foreland basin migration is on the order of 1000 km. Therefore, a comparable length of South American crust and mantle lithosphere has been underthrust beneath the Andean orogenic belt, enough to account for the present 60–70 km thickness of central Andean crust in southern Bolivia. Resolution of the details of the crustal-budget problem in the Andes will depend upon acquisition of better estimates of the amount and chronology of upper-crustal shortening in the Altiplano and Western Cordillera. However, the foreland basin record strongly suggests that the amount of crustal mass needed to thicken the Andean crust has been subducted from the east beneath the Andes. ACKNOWLEDGMENTS The U.S. National Science Foundation (EAR9804680) provided funding. Rich Fink provided assistance in the field. We thank N. McQuarrie, T. Sempere, R.F. Butler, S. Beck, G. Zandt, J. Reynolds, R.M. Hernańdez, G.L. Waanders, and L. Echavarria for informative discussions, preprints, and comments on early drafts of the manuscript. E. Lindsey, J.G. Lundberg, and S.T. Hasiotis kindly identified fossils and trace fossils. S. Tawackoli helped with logistical arrangements. Thorough reviews by T.E. Jordan, R. Dorsey, R. Astini, and T. Sempere helped us to substantially improve the paper. REFERENCES CITED Allen, P.A., Burgess, P.M., Galewsky, J., and Sinclair, H.D., 2001, Flexural-eustatic numerical model for drowning of the Eocene perialpine carbonate ramp and implications for Alpine geodynamics: Geological Society of America Bulletin, v. 113, p. 1052–1066. Allmendinger, R.W., and Gubbels, T., 1996, Pure and simple shear plateau uplift, Altiplano-Puna, Argentina and Bolivia: Tectonophysics, v. 259, p. 1–13. Allmendinger, R.W., Jordan, T.E., Kay, S.M., and Isacks, B.L., 1997, The evolution of the Altiplano-Puna plateau of the central Andes: Annual Review of Earth and Planetary Sciences, v. 25, p. 139–174. Amilibia, A., Sa`bat, F., Chong, G., Munõz, J.A., Roca, E., and Rodriguez-Perea, A., 1999, Evolution of Domeyko range, northern Chile, in Extended abstracts, Fourth International Symposium on Andean Geodynamics, Go¨ttingen, Germany: Paris, Institut de Recherche pour le De´veloppement, p. 25–29. Baby, P., Sempere, T., Oller, J., Barrios, L., He´rail, G., and Marocco, R., 1990, A late Oligocene–Miocene intermontane foreland basin in the southern Bolivian Altiplano: Comptes Rendus de l’Acade´mie des Sciences, Se´rie II, v. 311, p. 341–347. Baby, P., Rochat, P., Mascle, G., and He´rail, G., 1997, Neogene shortening contribution to crustal thickening in the back arc of the central Andes: Geology, v. 25, p. 883–886. Beck, S., and Zandt, G., 2002, The nature of orogenic crust in the central Andes: Journal of Geophysical Research (in press). Beck, S., Zandt, G., Myers, S.C., Wallace, T.C., Silver, P.G., and Drake, L., 1996, Crustal-thickness variations in the central Andes: Geology, v. 24, p. 407–410. Bradley, D.C., and Kidd, W.S.F., 1991, Flexural extension

of the upper continental crust in collisional foredeeps: Geological Society of America Bulletin, v. 103, p. 1416–1438. Branisa, L., Grambast, L., and Hoffstetter, R., 1969, Quelques prećisions nouvelles, d’apre´s les charophytes, sur l’aˆge du Groupe Puca (Cre´tace´–Paleóce`ne, Bolivie): Compte Rendu Sommaires de Seánces de la Socie´te´ Geólogique de France, v. 8, p. 321–322. Burns, W.M., and Jordan, T.E., 1999, Extension in the southern Andes as evidence by an Oligo–Miocene age intraarc basin, in Extended abstracts, Fourth International Symposium on Andean Geodynamics, Go¨ttingen, Germany: Paris, Institut de Recherche pour le De´veloppement, p. 115–118. Camilleri, P., Yonkee, W.A., Coogan, J.C., DeCelles, P.G., McGrew, A., and Wells, M., 1997, Hinterland to foreland transect through the Sevier orogen, northeast Nevada to southwest Wyoming: Structural style, metamorphism, and kinematic history of a large contractional orogenic wedge: Brigham Young University Geology Studies, v. 42, pt. 1, p. 297–309. Cardoza, N., and Jordan, T., 2001, Causes of spatially variable tectonic subsidence in the Miocene Bermejo foreland basin, Argentina: Basin Research, v. 13, p. 335–357. Carroll, A.R., and Bohacs, K.M., 1999, Stratigraphic classification of ancient lakes: Balancing tectonic and climatic controls: Geology, v. 27, p. 99–102. Cherroni, C., 1977, El Sistema Cretacico en la parte Boliviana de la cuenca Cretacica Andina: Revista Tećnica de Yacimientos Petrolı´feros Fiscales Bolivianos, v. 6, no. 1–2, p. 5–46. Chong, G., and Reutter, K.-J., 1985, Fenomenos de tectonica compresiva en las Sierras de Varas y de Argomedo, Precordilleran Chilena, en el ambito del paralelo 258 sur: IV Congreso Geologica Chile Actas, v. 2, p. 2-219–2-238. Crampton, S.L., and Allen, P.A., 1995, Recognition of flexural forebulge unconformities in the geologic record: American Association of Petroleum Geologists Bulletin, v. 79, p. 1495–1514. Currie, B.S., 1997, Sequence stratigraphy of nonmarine Jurassic–Cretaceous rocks, central Cordilleran forelandbasin system: Geological Society of America Bulletin, v. 109, p. 1206–1222. DeCelles, P.G., and Cavazza, W., 1999, A comparison of fluvial megafans in the Cordilleran (Late Cretaceous) and modern Himalayan foreland basin systems: Geological Society of America Bulletin, v. 111, p. 1315–1334. DeCelles, P.G., and Currie, B.S., 1996, Long-term sediment accumulation in the Middle Jurassic–early Eocene Cordilleran retroarc foreland basin system: Geology, v. 24, p. 591–594. DeCelles, P.G., and DeCelles, P.C., 2001, Rates of shortening, propagation, underthrusting, and flexural-wave migration in continental orogenic systems: Geology, v. 29, p. 135–138. DeCelles, P.G., and Giles, K.N., 1996, Foreland basin systems: Basin Research, v. 8, p. 105–123. DeCelles, P.G., and Hertel, F., 1989, Petrology of fluvial sands from the Amazonian foreland basin, Peru and Bolivia: Geological Society of America Bulletin, v. 101, p. 1552–1562. DeCelles, P.G., Langford, R.P., and Schwartz, R.K., 1983, Two new methods of paleocurrent determination from trough cross-stratification: Journal of Sedimentary Petrology, v. 53, p. 629–642. DeCelles, P.G., Gehrels, G.E., Quade, J., Ojha, T.P., Kapp, P.A., and Upreti, B.N., 1998, Neogene foreland basin deposits, erosional unroofing, and the kinematic history of the Himalayan fold-thrust belt, western Nepal: Geological Society of America Bulletin, v. 110, p. 2–21. Deconinck, J.F., Blanc-Valleron, M.M., Rouchy, J.M., Camoin, G., and Badaut-Trauth, D., 2000, Palaeoenvironmental and diagenetic control of the mineralogy of the Upper Cretaceous–lower Tertiary deposits of the Central palaeo-Andean basin of Bolivia (Potosı´ area): Sedimentary Geology, v. 132, p. 263–278. Dickinson, W.R., and Suczek, C.A., 1979, Plate tectonics

and sandstone compositions: American Association of Petroleum Geologists Bulletin, v. 63, p. 2164–2182. Dunn, J.F., Hartshorn, K.G., and Hartshorn, P.W., 1995, Structural styles and hydrocarbon potential of the SubAndean thrust belt of southern Bolivia, in Tankard, A.J., Sua´rez, R., and Welsink, H.J., eds., Petroleum basins of South America: American Association of Petroleum Geologists Memoir 62, p. 523–543. Ege, H., Jacobshagen, V., Scheuber, E., Sobel, E., and Vietor, T., 2001, Thrust related exhumation revealed by apatite fission track dating, central Andes (southern Bolivia): EGS XXVI General Assembly, Geophysical Research Abstracts, Nice, France, v. 3, p. 624. Erikson, J.P., and Kelley, S.A., 1995, Late Oligocene initiation of foreland-basin formation in Bolivia based on newly dated volcanic ash: IX Congreso Latinoamericano de Geologıá, Caracas, Venezuela, Ministerio de Energıá y Minas. Eugster, H.P., 1967, Hydrous sodium silicate from Lake Magadi: Precursors of bedded chert: Science, v. 157, p. 1177–1180. Evernden, J.F., Kriz, S.J., and Cherroni, C., 1977, Potassiumargon ages of some Bolivian rocks: Economic Geology, v. 72, p. 1042–1061. Farrar, E., Clark, A.H., Kontak, D.J., and Archibald, D.A., 1988, Zongo–San Gabań zone: Eocene foreland boundary of the Central Andean orogen, northwest Bolivia and southeast Peru: Geology, v. 16, p. 55–58. Flemings, P.B., and Jordan, T.E., 1989, A synthetic stratigraphic model of foreland basin development: Journal of Geophysical Research, v. 94, p. 3851–3866. Gayet, M., Marshall, L.G., and Sempere, T., 1991, The Mesozoic and Paleocene vertebrates of Bolivia and their stratigraphic context: A review: Revista Tećnica de Yacimientos Petrolı´feros Fiscales Bolivianos, v. 12, no. 3–4, p. 393–433. Geobol (Servicio Geolo´gico de Bolivia), 1962a, Carta geologica de Bolivia: La Paz, Servicio Geolo´gico de Bolivia, Lagunillas (Hoja 6336), scale 1:100,000. Geobol (Servicio Geolo´gico de Bolivia), 1962b, Carta geologica de Bolivia: La Paz, Servicio Geolo´gico de Bolivia, Potosı´ (Hoja 6435), scale 1:100,000. Geobol (Servicio Geolo´gico de Bolivia), 1994, Mapas tematicos de recursos minerales de Bolivia: La Paz, Servicio Geolo´gico de Bolivia, Hoja Uncia (SE 19–12), Ser. II-MTB-2B, scale 1:250,000. Geobol (Servicio Geolo´gico de Bolivia), 1995, Mapas tematicos de recursos minerales de Bolivia: La Paz, Servicio Geolo´gico de Bolivia, Hoja Camargo (SF 20– 1), Ser. II-MTB-9B, scale 1:250,000. Geobol (Servicio Geolo´gico de Bolivia), 1996, Mapas tematicos de recursos minerales de Bolivia: La Paz, Servicio Geolo´gico de Bolivia, Hoja Sucre (SE 20–13), Ser. II-MTB-8B, scale 1:250,000. Giles, K.A., and Dickinson, W.R., 1995, The interplay of eustasy and lithospheric flexure in forming stratigraphic sequences in foreland settings: An example from the Antler foreland, Nevada and Utah, in Dorobek, S.L., and Ross, G.M., eds., Stratigraphic evolution of foreland basins: SEPM (Society for Sedimentary Geology) Special Publication 52, p. 187–211. Goetze, H.J., and Schmidt, S., 1999, Rigidity and isostasy of the central Andes (208S to 308S): Eos (Transactions, American Geophysical Union), v. 80, p. 1052. Grier, M.E., Salfity, J.A., and Allmendinger, R.W., 1991, Andean reactivation of the Cretaceous Salta rift, northwestern Argentina: Journal of South American Earth Sciences, v. 4, p. 351–372. Gubbels, T.L., Isacks, B.L., and Farrar, E., 1993, High-level surfaces, plateau uplift, and foreland development, Bolivian central Andes: Geology, v. 21, p. 695–698. Gurnis, M., 1992, Rapid continental subsidence following the initiation and evolution of subduction: Science, v. 255, p. 1556–1558. Hammerschmidt, K., Do¨bel, R., and Friedrichsen, H., 1992, Implication of 40Ar/39Ar dating of early Tertiary volcanic rocks from the north-Chilean Precordillera: Tectonophysics, v. 202, p. 55–81. Hartley, A.J., Jolley, E.J., and Turner, P., 1992, Paleomagnetic evidence for rotation in the Precordillera of northern Chile: Structural constraints and implications


75

DECELLES and HORTON for the evolution of the Andean forearc: Tectonophysics, v. 205, p. 49–64. Hartley, A.J., May, G., Chong, G., Turner, P., Kape, S.J., and Jolley, E.J., 2000, Development of a continental forearc: A Cenozoic example from the central Andes, northern Chile: Geology, v. 28, p. 331–334. Hasiotis, S.T., and Bown, T.M., 1992, Invertebrate trace fossils: The backbone of continental ichnology, in Maples, C.G., and West, R.R., eds., Trace fossils: Paleontological Society, Short Courses In Paleontology no. 5, p. 64–104. He´rail, G., Oller, J., Baby, P., Bonhomme, M., and Soler, P., 1996, Strike-slip faulting, thrusting and related basins in the Cenozoic evolution of the southern branch of the Bolivian orocline: Tectonophysics, v. 259, p. 201–212. Hernańdez, R.M., Reynolds, J., and Disalvo, A., 1996, Ana´lisis tectosedimentario y ubicacioń geocronolo´gica del Grupo Orań en el Rıó Iruya: Boletıń de Informaciones Petroleras (Yacimientos Petrolı´feros Fiscales), v. 12, no. 45, p. 80–93. Hernańdez, R.M., Galli, C.I., and Reynolds, J., 1999, Estratigrafıá del Terciario en el noroeste Argentino, in Gonzalez, B., Bonorino, R., and Viramonte, J., eds., Geologia del noroeste Argentino: Relatorio XIV Congreso Geologico Argentino, Salta, Argentina, v. 1, p. 316–328. Horton, B., 1996, Sequence of late Oligocene–Miocene fold-thrust deformation and development of piggyback basins in the Eastern Cordillera, southern Bolivia, in Extended abstracts, Third International Symposium on Andean Geodynamics, St. Malo, France: Paris, ORSTOM, p. 383–386. Horton, B.K., 1998a, Sediment accumulation on top of the Andean orogenic wedge:Oligocene to late Miocene basins of the Eastern Cordillera, southern Bolivia: Geological Society of America Bulletin, v. 110, p. 1174–1192,p. 1513. Horton, B.K., 1998b, Late Cretaceous to Holocene evolution of the foreland basin system and associated foldthrust belt in the central Andes of Bolivia [Ph.D. thesis]: Tucson, University of Arizona, 208 p. Horton, B.K., and DeCelles, P.G., 1997, The modern foreland basin system adjacent to the central Andes: Geology, v. 25, p. 895–898. Horton, B.K., and DeCelles, P.G., 2001, Modern and ancient fluvial megafans in the foreland basin system of the central Andes, southern Bolivia: Implications for drainage network evolution in fold-thrust belts: Basin Research, v. 13, p. 43–63. Horton, B.K., Hampton, B.A., and Waanders, G., 2001, Paleogene synorogenic sedimentation in the Altiplano plateau and implications for initial mountain building in the central Andes: Geological Society of America Bulletin, v. 113, p. 1387–1400. Horton, B.K., Hampton, B.A., LaReau, B.N., and Baldelloń, E., 2002, Tertiary provenance history of the northern and central Altiplano plateau (central Andes, Bolivia): A detrital record of plateau-margin tectonics: Journal of Sedimentary Research, v. 72, p. 711–726. Ingersoll, R.V., Bullard, T.F., Ford, R.L., Grimm, J.P., Pickle, J.D., and Sares, S.W., 1984, The effect of grain size on detrital modes: A test of the Gazzi-Dickinson point-counting method: Journal of Sedimentary Petrology, v. 54, p. 103–116. Isaacson, P.E., 1975, Evidence for a western extracontinental land source during the Devonian Period in the central Andes: Geological Society of America Bulletin, v. 86, p. 39–46. Isacks, B.L., 1988, Uplift of the central Andean plateau and bending of the Bolivian orocline: Journal of Geophysical Research, v. 93, p. 3211–3231. Jacobi, R.D., 1981, Peripheral bulge—A causal mechanism for the Lower Ordovician unconformity along the western margin of the northern Appalachians: Earth and Planetary Science Letters, v. 56, p. 245–251. James, D.E., 1971, Plate-tectonic model for the evolution of the central Andes: Geological Society of America Bulletin, v. 82, p. 3325–3346. James, D.E., and Sacks, S., 1999, Cenozoic formation of the central Andes: A geophysical perspective, in Skinner, B.J., ed., Geology and ore deposits of the central

76

Andes: Society of Economic Geologists Special Publication 7, p. 1–25. Jordan, T.E., Flemings, P.B., and Beer, J.A., 1988, Dating thrust-fault activity by use of foreland-basin strata, in Kleinspehn, K.L., and Paola, C., eds., New perspectives in basin analysis: New York, Springer-Verlag, p. 307–330. Jordan, T.E., Allmendinger, R.W., Damanti, J.F., and Drake, R.E., 1993, Chronology of motion in a complete thrust belt: The Precordillera, 30–318S, Andes Mountains: Journal of Geology, v. 101, p. 135–156. Jordan, T.E., Reynolds, J.H., and Erikson, J.P., 1997, Variability in age of initial shortening and uplift in the central Andes, 16–338309 S, in Ruddiman, W.F., ed., Tectonic uplift and climate change: New York, Plenum Press, p. 41–61. Jordan, T.E., Burns, W.M., Veiga, R., Pańgaro, F., Copeland, P., Kelley, S., and Mpodozis, C., 2001, Extension and basin formation in the southern Andes caused by increased convergence rate: A mid-Cenozoic trigger for the Andes: Tectonics, v. 20, p. 308–324. Kennan, L., Lamb, S., and Rundle, C., 1995, K-Ar dates from the Altiplano and Cordillera Oriental of Bolivia: Implications for Cenozoic stratigraphy and tectonics: Journal of South American Earth Sciences, v. 8, p. 163–186. Kley, J., 1996, Transition from basement-involved to thinskinned thrusting in the Cordillera Oriental of southern Bolivia: Tectonics, v. 15, p. 763–775. Kley, J., and Monaldi, C.R., 1998, Tectonic shortening and crustal thickness in the central Andes: How good is the correlation?: Geology, v. 26, p. 723–726. Kley, J., Gangui, A.H., and Kruger, D., 1996, Basementinvolved blind thrusting in the eastern Cordillera Oriental, southern Bolivia: Evidence from cross-sectional balancing, gravimetric and magnetotelluric data: Tectonophysics, v. 259, p. 179–184. Kley, J., Mu¨ller, J., Tawackoli, S., Jacobshagen, V., and Manutsoglu, E., 1997, Pre-Andean and Andean-age deformation in the Eastern Cordillera of southern Bolivia: Journal of South American Earth Sciences, v. 10, p. 1–19. Lamb, S., and Hoke, L., 1997, Origin of the high plateau in the central Andes, Bolivia, South America: Tectonics, v. 16, p. 623–649. Lamb, S., Hoke, L., Kennan, L., and Dewey, J., 1997, Cenozoic evolution of the central Andes in Bolivia and northern Chile, in Burg, J.P., and Ford, M., eds., Orogeny through time: Geological Society [London] Special Publication 121, p. 237–264. Lawton, T.F., 1986, Compositional trends within a clastic wedge adjacent to a fold-thrust belt: Indianola Group, central Utah, U.S.A., in Allen, P.A., and Homewood, P., eds., Foreland basins: International Association of Sedimentologists Special Publication 8, p. 411–423. Lehmann, B., 1978, A Precambrian core sample from the Altiplano/Bolivia: Geologische Rundschau, v. 67, p. 270–278. Lihou, J.C., and Allen, P.A., 1996, Geometry of the underfilled North Alpine foreland basin, Switzerland, and the importance of inherited rift margin structures: Basin Research, v. 8, p. 425–442. Litherland, M., Annells, R.N., Appleton, J.D., Berrange, J.P., Bloomfield, K., Burton, C.C.J., Darbyshire, D.P.F., Fletcher, C.J.N., Hawkins, M.P., Klinck, B.A., Llanos, A., Mitchell, W.I., O’Connor, E.A., Pitfield, P.E.J., Power, G., and Webb, B.C., 1986, The geology and mineral resources of the Bolivian Precambrian shield: British Geological Survey Overseas Memoir, v. 9, 153 p. Litherland, M., and Pitfield, P.E.J., 1983, The Mesozoic– Cenozoic history of eastern Bolivia and the recognition of four ages of duricrust, in Proceedings of the Second International Seminar on Lateritisation Processes: Sao Paulo, Brazil, Sao Paulo University, p. 281–294. Lorenzo, J.M., O’Brien, G.W., and Tandon, K., 1998, Inelastic yielding and forebulge shape across a modern foreland basin: North West Shelf of Australia, Timor Sea: Geophysical Research Letters, v. 25, p. 1455–1458. Lucassen, F., Lewerenz, S., Franz, G., Viramonte, J., and Mezger, K., 1999, Metamorphism, isotopic ages and

composition of lower crustal granulite xenoliths from the Cretaceous Salta Rift, Argentina: Contributions to Mineralogy and Petrology, v. 134, p. 325–341. Lundberg, J.G., 1998, The temporal context for diversification of neotropical fishes, Chapter 2, in Malabarba, L.R., Reis, R.E., Vari, R.P., Lucena, C.A.S., and Lucena, Z.M.S., eds., Phylogeny and classification of neotropical fishes: Porto Alegre, Brazil, Museu de Cieˆncias e Tecnologia, Pontificia Universidade Cato´lica do Rio Grande do Sul (PUCRS). Lundberg, J.G., Machado-Allison, A., and Kay, R.F., 1986, Miocene characid fishes from Colombia: Evidence for evolutionary stasis and extirpation in the South American ichthyofauna: Science, v. 234, p. 208–209. Lyon-Caen, H., Molnar, P., and Sua´rez, G., 1985, Gravity anomalies and flexure of the Brazilian Shield beneath the Bolivian Andes: Earth and Planetary Science Letters, v. 75, p. 81–92. Marocco, R., Sempere, T., Cirbian, M., and Oller, J., 1987, Evidence for Paleocene deformation in southern Bolivia: Importance for geodynamic evolution of central Andes: Comptes Rendus Acade´mie des Sciences, Se´r. II, v. 304, p. 1139–1143. Marsaglia, K., 1995, Interarc and backarc basins, in Busby, C.J., and Ingersoll, R.V., eds., Tectonics of sedimentary basins: Oxford, Blackwell Science, p. 299–329. Marshall, L.G., and Sempere, T., 1991, The Eocene to Pleistocene vertebrates of Bolivia and their stratigraphic context: A review: Revista Tećnica de Yacimientos Petrolı´feros Fiscales Bolivianos, v. 12, no. 3– 4, p. 631–652. Marshall, L.G., Sempere, T., and Gayet, M., 1993, The Petaca (late Oligocene–middle Miocene) and Yecua (late Miocene) formations of the Subandean-Chaco basin, Bolivia, and their tectonic significance: Documents des Laboratoires de Geólogie de Lyon, v. 125, p. 291–301. McBride, S.L., Clark, A.H., Farrar, E., and Archibald, D.A., 1987, Delimitation of a cryptic Eocene tectono-thermal domain in the Eastern Cordillera of the Bolivian Andes through K-Ar dating and 40Ar-39Ar step-heating: Geological Society [London] Journal, v. 144, p. 243–255. McKenzie, D.P., 1978, Some remarks on the development of sedimentary basins: Earth and Planetary Science Letters, v. 40, p. 25–32. McQuarrie, N., 2002, The kinematic history of the central Andean fold-thrust belt, Bolivia: Implications for building a high plateau: Geological Society of America Bulletin, v. 114, p. 950–963. McQuarrie, N., and DeCelles, P.G., 2001, Geometry and structural evolution of the central Andean backthrust belt, Bolivia: Tectonics, v. 20, p. 669–692. Me´gard, F., 1984, The Andean orogenic period and its major structures in central and northern Peru: Geological Society [London] Journal, v. 141, p. 893–900. Me´gard, F., Noble, D.C., McKee, E.H., and Bellon, H., 1984, Multiple pulses of Neogene compressive deformation in the Ayacucho intermontane basin, Andes of central Peru: Geological Society of America Bulletin, v. 95, p. 1108–1117. Meyers, J.H., Suttner, L.J., Furer, L.C., May, M.T., and Soreghan, M.J., 1992, Intrabasinal tectonic control on fluvial sandstone bodies in the Cloverly Formation (Early Cretaceous), west-central Wyoming, U.S.A.: Basin Research, v. 4, p. 315–333. Mitrovica, J.X., Beaumont, C., and Jarvis, G.T., 1989, Tilting of continental interiors by the dynamical effects of subduction: Tectonics, v. 8, p. 1079–1094. Moretti, I., Baby, P., Meńdez, E., and Zubieta, D., 1996, Hydrocarbon generation in relation to thrusting in the Sub Andean Zone from 18 to 228S, Bolivia: Petroleum Geoscience, v. 2, p. 17–28. Mourier, T., Bengtson, P., Bonhomme, M., Buge, E., Cappetta, H., Crochet, J.-Y., Feist, M., Hirsch, K.F., Jaillard, E., Laubacher, G., Lefranc, J.P., Moullade, M., Noblet, C., Pons, D., Rey, J., Sige´, B., Tambareau, Y., and Taquet, P., 1988, The Upper Cretaceous–lower Tertiary marine to continental transition in the Bagua basin, northern Peru: Paleontology, biostratigraphy, radiometry, correlations: Newsletter on Stratigraphy, v. 19, p. 143–177. Mpodozis, C., and Allmendinger, R.W., 1993, Extensional


FORELAND BASIN DEVELOPMENT AND HISTORY OF ANDEAN CRUSTAL SHORTENING, BOLIVIA tectonics, Cretaceous Andes, northern Chile (278S): Geological Society of America Bulletin, v. 105, p. 1462–1477. Nicolas, C., Wilke, H.-G., Schneider, H., and Lucassen, F., 1999, New constraints on the Cretaceous–Tertiary deformation in Sierra de Moreno, Precordillera of the Segunda region de Antofagasta, northern Chile, in Extended abstracts, Fourth International Symposium on Andean Geodynamics, Go¨ttingen, Germany: Paris, Institut de Recherche pour le De´veloppement, p. 543–546. Noble, D.C., McKee, E.H., Mourier, T., and Me´gard, F., 1990, Cenozoic stratigraphy, magmatic activity, compressive deformation, and uplift in northern Peru: Geological Society of America Bulletin, v. 102, p. 1105–1113. Noblet, C., Lavenu, A., and Marocco, R., 1996, Concept of continuum as opposed to periodic tectonism in the Andes: Tectonophysics, v. 255, p. 65–78. Norabuena, E., Leffler-Griffin, L., Mao, A., Dixon, T., Stein, S., Sacks, I.S., Ocola, L., and Ellis, M., 1998, Space geodetic observations of Nazca–South America convergence across the central Andes: Science, v. 279, p. 358–362. Norabuena, E., Dixon, T.H., Stein, S., and Harrison, C.G.A., 1999, Decelerating Nazca–South America and Nazca-Pacific plate motions: Geophysical Research Letters, v. 26, p. 3405–3408. Owens, T., and Zandt, G., 1997, The implications of crustal property variations for models of Tibetan plateau evolution: Nature, v. 387, p. 37–43. Pardo-Casas, F., and Molnar, P., 1987, Relative motion of the Nazca (Farallon) and South American plates since Late Cretaceous time: Tectonics, v. 6, p. 233–248. Pope, D.C., and Willett, S.D., 1998, Thermal-mechanical model for crustal thickening in the central Andes driven by ablative subduction: Geology, v. 26, p. 511–514. Price, R.A., 1981, The Cordilleran foreland thrust-and-fold belt in the southern Canadian Rocky Mountains, in McClay, K.J., and Price, N.J., eds., Thrust and nappe tectonics: Geological Society of London Special Publication 9, p. 427–448. Reynolds, J.H., Hernandez, R.M., Idleman, B.D., Naeser, C.W., and Guerstein, P.G., 1996, Rio Iruya revisited: The Neogene tectonic development of the western Sierras Subandinas, Salta province, northwest Argentina: Geological Society of America Abstracts with Programs, v. 28, no. 7, p. 59. Reynolds, J.H., Galli, C.I., Hernańdez, R.M., Idleman, B.D., Kotila, J.M., Hilliard, R.V., and Naeser, C.W., 2000, Middle Miocene tectonic development of the transition zone, Salta province, northwest Argentina: Magnetic stratigraphy from the Metań Subgroup, Sierra de Gonza´lez: Geological Society of America Bulletin, v. 112, p. 1736–1751. Roeder, D., 1988, Andean-age structure of Eastern Cordillera (province of La Paz, Bolivia): Tectonics, v. 7, p. 23–39. Roeder, D., and Chamberlain, R.L., 1995, Structural geology of Sub-Andean fold and thrust belt in northwestern Bolivia, in Tankard, A.J., Sua´rez, R., and Welsink, H.J., eds., Petroleum basins of South America: American Association of Petroleum Geologists Memoir 62, p. 459–479. Roperch, P., Arriagada, C., and Coutand, I., 2000a, Tectonic rotations in the central Andes: Implications for the geodynamic evolution of the Altiplano-Puna plateau: Eos (Transactions, American Geophysical Union), v. 81, p. 1083. Roperch, P., Fornari, M., He´rail, G., and Parraguez, G.V., 2000b, Tectonic rotations within the Bolivian Altiplano: Implications for the geodynamic evolution of the central Andes during the late Tertiary: Journal of Geophysical Research, v. 105, p. 795–820.

Salfity, J.A., and Marquillas, R.A., 1994, Tectonic and sedimentary evolution of the Cretaceous–Eocene Salta Group basin, Argentina, in Salfity, J.A., ed., Cretaceous tectonics of the Andes: Wiesbaden, Germany, Vieweg Publishing, p. 266–315. Scheuber, E., and Reutter, K.-J., 1992, Magmatic arc tectonics in the central Andes between 218 and 258S: Tectonophysics, v. 205, p. 127–140. Scheuber, E., Bogdanic, T., Jensen, A., and Reutter, K.-J., 1994, Tectonic development of the North Chilean Andes in relation to plate convergence and magmatism since the Jurassic, in Reutter, K.-J., Scheuber, E., and Wigger, P.J., eds., Tectonics of the southern central Andes: Berlin, Springer-Verlag, p. 121–140. Schmitz, M., 1994, A balanced model of the southern central Andes: Tectonics, v. 13, p. 484–492. Schmitz, M., and Kley, J., 1997, The geometry of the central Andean backarc crust: Joint interpretation of cross-section balancing and seismic refraction data: Journal of South American Earth Sciences, v. 10, p. 99–110. Schwab, F.L., 1986, Sedimentary ‘‘signatures’’ of foreland basin assemblages: Real or counterfeit?, in Allen, P.A., and Homewood, P., eds., Foreland basins: International Association of Sedimentologists Special Publication 8, p. 395–410. Schwartz, R.K., and DeCelles, P.G., 1988, Foreland basin evolution and synorogenic sedimentation in response to interactive Cretaceous thrusting and reactivated foreland partitioning, in Schmidt, C.J., and Perry, W.J. Jr., eds., Interaction of the Rocky Mountain foreland and the Cordilleran thrust belt: Geological Society of America Memoir 171, p. 489–513. Se´brier, M., Lavenu, A., Fornari, M., and Soulas, J.-P., 1988, Tectonics and uplift in central Andes (Peru, Bolivia and northern Chile) from Eocene to present: Geódynamique, v. 3, p. 85–106. Sempere, T., 1991, Cenozoic tectonic ‘‘phases’’ in Bolivia: Some needed clarifications: Vinã del Mar, VI Congreso Geologico Chileno, Actas, v. 1, p. 877–881. Sempere, T., 1994, Kimmeridgian? to Paleocene tectonic evolution of Bolivia, in Salfity, J.A., ed., Cretaceous tectonics of the Andes: Wiesbaden, Germany, Vieweg Publishing, p. 168–212. Sempere, T., Cha´vez, H., and Pe´rez, M., 1989, Estudio sedimentologico preliminar de la parte superior de la Formacion Mondragon en las cercanias de Lenas (Departamento de Potosı´) consecuencias geodinamicas: Memorias del VIII Congreso Geolo´gico de Bolivia, La Paz, p. 273–278. Sempere, T., He´rail, G., Oller, J., and Bonhomme, M.G., 1990, Late Oligocene–early Miocene major tectonic crisis and related basins in Bolivia: Geology, v. 18, p. 946–949. Sempere, T., Butler, R.F., Richards, D.R., Marshall, L.G., Sharp, W., and Swisher, C.C., 1997, Stratigraphy and chronology of Late Cretaceous–early Paleogene strata in Bolivia and northwest Argentina: Geological Society of America Bulletin, v. 109, p. 709–727. Sheffels, B.M., 1990, Lower bound on the amount of crustal shortening in the central Bolivian Andes: Geology, v. 18, p. 812–815. Sinclair, H.D., 1997, Tectonostratigraphic model for underfilled peripheral foreland basins: An Alpine perspective: Geological Society of America Bulletin, v. 109, p. 324–346. Smith, G.A., and Landis, C.A., 1995, Intraarc basins, in Busby, C.J., and Ingersoll, R.V., eds., Tectonics of sedimentary basins: Oxford, Blackwell Science, p. 263–298. Somoza, R., 1998, Updated Nazca (Farallon)–South America relative motions during the last 40 My: Implications for mountain building in the central Andean region: Journal of South American Earth Sciences, v. 11, p. 211–215.

Steinmann, G., 1929, Geologie von Peru: Heidelberg, Germany, Karl Winter, 448 p. Sua´rez, R., 1995, Comentarios sobre la edad de la Formacion Cancan˜iri: Revista Tećnica de Yacimientos Petrolı´feros Fiscales Bolivianos, v. 16, no. 1–2, p. 51–54. Sua´rez, R., and Diaz, E., 1996, Lexico estratigrafico de Bolivia: Revista Tećnica de Yacimientos Petrolı´feros Fiscales Bolivianos, v. 17, no. 1–2, p. 7–227. Tawackoli, S., Jacobshagen, V., Wemmer, K., and Andriessen, P.A.M, 1996, The Eastern Cordillera of southern Bolivia: A key region to the Andean backarc uplift and deformation history, in Extended abstracts, Third International Symposium on Andean Geodynamics, St. Malo, France: Paris, ORSTOM, p. 505–508. Troe¨ng, B., Claure, H., Oliveira, L., Ballun, R., and Walser, G., 1993, Thematic maps of the mineral resources of Bolivia: Tarija and Villazoń sheets (1:250,000 scale): Boletıń del Servicio Geolo´gico de Bolivia, v. 3, 178 p. Ussami, N., Shiraiwa, S., and Landim Dominguez, J.M., 1999, Basement reactivation in a sub-Andean foreland flexural bulge: The Pantanal wetland, southwest Brazil: Tectonics, v. 18, p. 25–39. Viramonte, J.G., Kay, S.M., Becchio, R., Escayola, M., and Novitski, I., 1999, Cretaceous rift related magmatism in central-western south America: Journal of South American Earth Sciences, v. 12, p. 109–121. Watts, A.B., Lamb, S.H., Fairhead, J.D., and Dewey, J.F., 1995, Lithospheric flexure and bending of the central Andes: Earth and Planetary Science Letters, v. 134, p. 9–21. Wdowinski, S., and Bock, Y., 1994, The evolution of deformation and topography of high elevated plateaus: 2. Application to the central Andes: Journal of Geophysical Research, v. 99, p. 7121–7130. Wells, N.A., and Dorr, J.A., 1987, A reconnaissance of sedimentation on the Kosi alluvial fan of India, in Flores, R.M., and Harvey, M.D., eds., Recent developments in fluvial sedimentology: Society of Economic Paleontologists and Mineralogists, Special Publication 39, p. 51–61. Welsink, H.J., Martıńez, E., Aranı´bar, O., and Jarandilla, J., 1995, Structural inversion of a Cretaceous rift basin, southern Altiplano, Bolivia, in Tankard, A.J., Sua´rez, R., and Welsink, H.J., eds., Petroleum basins of South America: American Association of Petroleum Geologists Memoir 62, p. 305–324. White, T., Furlong, K., and Arthur, M., 2002, Forebulge migration in the Cretaceous Western Interior basin of the central United States: Basin Research, v. 14, p. 43–54. Wigger, P.J., Schmitz, M., Araneda, M., Asch, G., Baldzuhn, S., Giese, P., Heinsohn, W., Martıńez, E., Ricaldi, E., Rower, P., and Viramonte, J., 1994, Variation in the crustal structure of the southern central Andes deduced from seismic refraction investigations, in Reutter, K.-J., Scheuber, E., and Wigger, P.J., eds., Tectonics of the southern central Andes: New York, Springer-Verlag, p. 23–48. Wo¨rner, G., Lezaun, J., Beck, A., Heber, V., Lucassen, F., Zinngrebe, E., Ro¨ssling, R., and Wilke, H.G., 2000, Precambrian and early Paleozoic evolution of the Andean basement at Belen (northern Chile) and Cerro Uyarani (western Bolivia Altiplano): Journal of South American Earth Sciences, v. 13, p. 717–737. Zaleha, M.J., Way, J.N., and Suttner, L.J., 2001, Effects of syndepositional faulting and folding on Early Cretaceous rivers and alluvial architecture (Lakota and Cloverly Formations, Wyoming, U.S.A.): Journal of Sedimentary Research, v. 71, p. 880–894. MANUSCRIPT RECEIVED BY THE SOCIETY 15 NOVEMBER 2000 REVISED MANUSCRIPT RECEIVED 10 JULY 2002 MANUSCRIPT ACCEPTED 12 JULY 2002 Printed in the USA


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