Cenozoic foreland basin system in the central ... - Wiley Online Library

TECTONICS, VOL. 30, TC6013, doi:10.1029/2011TC002948, 2011

Cenozoic foreland basin system in the central Andes of northwestern Argentina: Implications for Andean geodynamics and modes of deformation P. G. DeCelles,1 B. Carrapa,1 B. K. Horton,2 and G. E. Gehrels1 Received 17 May 2011; revised 10 October 2011; accepted 12 October 2011; published 21 December 2011.

[1] Cenozoic strata in the central Andes of northwestern Argentina record the development and migration of a regional foreland basin system analogous to the modern Chaco‐Paraná alluvial plain. Paleocene‐lower Eocene fluvial and lacustrine deposits are overlain by middle‐upper Eocene hypermature paleosols or an erosional disconformity representing 10–15 Myr. This ‘supersol/disconformity’ zone is traceable over a 200,000 km2 area in the Andean thrust belt, and is overlain by 2–6 km of upward coarsening, eastward thinning, upper Eocene through lower Miocene fluvial and eolian deposits. Middle Miocene‐Pliocene fluvial, lacustrine, and alluvial fan deposits occupy local depocenters with contractional growth structures. Paleocurrent and petrographic data demonstrate westerly provenance of quartzolithic and feldspatholithic sediments. Detrital zircon ages from Cenozoic sandstones cluster at 470–491, 522–544, 555–994, and 1024–1096 Ma. Proterozoic‐Mesozoic clastic and igneous rocks in the Puna and Cordillera Oriental yield similar age clusters, and served as sources of the zircons in the Cenozoic deposits. Arc‐derived zircons become prominent in Oligo‐Miocene deposits and provide new chronostratigraphic constraints. Sediment accumulation rate increased from ∼20 m/Myr during Paleocene‐Eocene time to 200–600 m/Myr during the middle to late Miocene. The new data suggest that a flexural foreland basin formed during Paleocene time and migrated at least 600 km eastward at an unsteady pace dictated by periods of abrupt eastward propagation of the orogenic strain front. Despite differences in deformation style between Bolivia and northwestern Argentina, lithosphere in these two regions flexed similarly in response to eastward encroachment of a comparable orogenic load beginning during late Paleocene time. Citation: DeCelles, P. G., B. Carrapa, B. K. Horton, and G. E. Gehrels (2011), Cenozoic foreland basin system in the central Andes of northwestern Argentina: Implications for Andean geodynamics and modes of deformation, Tectonics, 30, TC6013, doi:10.1029/2011TC002948.

1. Introduction [2] The most impressive topographic feature of the Andean orogenic belt is the Central Andean Plateau, which encompasses an area of ∼500,000 km2 between 15°S and 27°S latitudes where the average surface elevation exceeds 3 km [Isacks, 1988; Allmendinger et al., 1997; Strecker et al., 2007]. It is now widely documented that the plateau is a result of crustal shortening, but the mode and history of shortening remain topics of active research. Allmendinger et al. [1983, 1997] and Kley and Monaldi [1998] pointed out that a number of major changes take place in the Central Andean Plateau at about 23–24°S. North of this latitude, horizontal 1 Department of Geosciences, University of Arizona, Tucson, Arizona, USA. 2 Institute for Geophysics and Department of Geological Sciences, Jackson School of Geosciences, University of Texas at Austin, Austin, Texas, USA.

Copyright 2011 by the American Geophysical Union. 0278‐7407/11/2011TC002948

tectonic shortening in the upper crust is on the order of several hundred km [Kley, 1999; McQuarrie, 2002a]; deformation is dominated by thin‐skinned thrusting above regional detachments that dip gently westward [Roeder, 1988; Sheffels, 1990; Dunn et al., 1995; Baby et al., 1997; McQuarrie, 2002a; Echavarria et al., 2003; Uba et al., 2009]; deformation propagated eastward through time [McQuarrie, 2002a; Horton, 2005; Gillis et al., 2006; Ege et al., 2007; McQuarrie et al., 2008; Barnes et al., 2008]; and the crust is relatively thick (ca. 65–70 km) [Beck and Zandt, 2002; Heit et al., 2008]. South of this latitude, documented shortening is only about 100 km and was accommodated by steeply dipping, bivergent thrust faults that cut deeply into the crust (>25 km) [Cristallini et al., 1997; Kley and Monaldi, 2002; Mortimer et al., 2007]; deformation is considered to have persisted throughout the orogenic belt during Miocene‐ Pleistocene time [Allmendinger and Gubbels, 1996]; and the crust is much thinner (45–55 km) and more variable [Yuan et al., 2002; Heit et al., 2008]. [3] Significant along‐strike differences in the pre‐existing geological structure and lithologic composition of Andean

TC6013

1 of 30

TC6013

DECELLES ET AL.: ANDEAN FORELAND BASIN

Figure 1. Cartoons illustrating differences between pure and simple shear modes of orogenic strain and lithospheric thickening, after Allmendinger [1986] and Allmendinger and Gubbels [1996]. Of particular importance is the lateral offset of flexural isostatic compensation required in the simple shear case, in contrast to the vertically juxtaposed compensation required by pure shear. upper crust may be responsible for some of the major differences noted above [Allmendinger et al., 1983, 1997; Allmendinger and Gubbels, 1996; McQuarrie, 2002b]. For example, changes in structural style and shortening coincide with a southward decrease in the thickness of Paleozoic cover strata and the presence of a Cretaceous rift system south of 23°S [Kley and Monaldi, 2002]. Still farther south (27°S), Andean crustal shortening is dominated by massive crystalline basement block uplifts of the Sierras Pampeanas, where the Paleozoic cover section is largely absent. Allmendinger [1986] and Allmendinger and Gubbels [1996] suggested that these differences in orogenic structure and history are manifestations of two distinct modes of lithospheric shortening: simple shear in the north, and pure shear in the south (Figure 1). Isacks [1988] proposed that the central Andes experienced a temporal transition from dominantly pure shear to simple shear during the Miocene. The idea of pure shear mountain building has had considerable influence on subsequent tectonics research in the central Andes [see, e.g., Sobolev and Babeyko, 2005; Hindle et al., 2005; Sobolev et al., 2006] and similar models have been applied to other major orogenic systems in order to explain first‐order features in the distribution of deformed lithosphere [England and Houseman, 1985; Houseman and England, 1993; Molnar et al., 1993; Ellis et al., 1995; Molnar and Houseman, 2004]. [4] Simple and pure shear modes of deformation should be accompanied by drastically different responses in laterally adjacent lithosphere [Allmendinger and Gubbels, 1996]. Pure shear lithospheric shortening and thickening should be compensated locally, whereas simple shear shortening and thickening should produce a flexural response in the foreland lithosphere (Figure 1). It follows that the Cenozoic stratigraphy of the central Andes should contain a well‐

TC6013

developed regional foreland basin system in Bolivia and northernmost Argentina, where the presence of a major thin‐ skinned thrust belt—attesting to orogen‐scale simple shear— is well documented [e.g., Dunn et al., 1995; McQuarrie, 2002a; Echavarria et al., 2003]. On the other hand, if pure shear deformation has been dominant south of 23°S, a flexural signal in Cenozoic strata should be absent or significantly diminished, insofar as pure shear lithospheric shortening does not require regional (flexural) compensation [Allmendinger and Gubbels, 1996]. [5] Previous work in the Bolivian Cenozoic record demonstrates that a regional scale foreland basin system developed during Paleocene time and migrated 800–1,000 km eastward to its present location in the Chaco plain [Horton and DeCelles, 1997; Sempere et al., 1997; Horton et al., 2001; DeCelles and Horton, 2003; Uba et al., 2006, 2009]. On the other hand, no consensus has been reached concerning the development and evolution of a foreland basin system south of 23°S in northwestern Argentina. Whereas some studies conclude that a foreland basin was present in northwestern Argentina from Eocene time onward [Jordan and Alonso, 1987; Starck and Vergani, 1996; Kraemer et al., 1999; Coutand et al., 2001; Carrapa et al., 2005; Hongn et al., 2007; Carrapa and DeCelles, 2008], others suggest that Paleocene‐Eocene strata in northwestern Argentina were accommodated by thermal subsidence following Cretaceous rifting, with foreland basin development delayed until late Oligocene‐early Miocene time [Jordan and Alonso, 1987; Salfity and Marquillas, 1994; Cominguez and Ramos, 1995; Marquillas et al., 2005; del Papa, 2006]. The late Miocene‐Quaternary history of basin evolution in northwestern Argentina is also a subject of debate, with some authors arguing for continued eastward propagation of deformation and migration of the foreland basin [Carrapa et al., 2011a, 2011b] while others advocate a change to localized basins associated with thick‐skinned deformation and regionally non‐systematic strain propagation [Hain et al., 2011; Strecker et al., 2011]. These debates are critical for accurately assessing the mechanisms and magnitude of shortening in the central Andes and for constraining models that link westward underthrusting of South American cratonic basement to arc magmatism and upper mantle processes [Kay et al., 1994; Beck and Zandt, 2002; McQuarrie et al., 2005; Sobolev et al., 2006; Haschke et al., 2006; Kay and Coira, 2009; DeCelles et al., 2009]. [6] In this paper we present new evidence from northwestern Argentina to support the view that orogenic shortening and foreland basin development began no later than late Paleocene time throughout the central Andes from northern Bolivia to at least as far south as 26°S. This interpretation is based on sedimentological observations from >14,000 m of detailed measured stratigraphic sections, 56 modal petrographic analyses, >2,500 U‐Pb ages of detrital and igneous zircons, and apatite fission track dating of cobbles in synorogenic conglomerates. Our results indicate that the central Andean foreland basin system in northwestern Argentina has migrated at least 600 km laterally at an unsteady pace varying from ∼5 mm/yr to more than 40 mm/yr to its present location east of the orogenic belt, and that this migration was superimposed upon paleotopographically and geologically complex lithosphere with pre‐existing Cretaceous rift‐

2 of 30

TC6013


related structures that locally influenced sediment accommodation and composition.

2. Geological Setting [7] Spanning the entire ∼7000 km long western margin of the South American plate, the Andean orogenic belt is Earth’s best example of an active Cordilleran‐style orogenic system [Coney and Evenchick, 1994; Ramos, 2009]. The central Andes in Bolivia and northern Chile and Argentina are where the orogen is highest and widest and where shortening is greatest [James, 1971; Isacks, 1988; Kley and Monaldi, 1998; Beck and Zandt, 2002; McQuarrie et al., 2005; Oncken et al., 2006]. Flanking the modern Andes on the east is a continental‐scale retroarc foreland basin system [Chase et al., 2009] that provides a rich archive of surficial and geodynamic processes that have shaped the Andes. [8] Eastward from the modern forearc region, the central Andes are divisible into four longitudinal tectonomorphic zones, including the Cordillera Occidental, Altiplano‐Puna plateau, Cordillera Oriental, and the frontal Subandean and Santa Bárbara ranges (Figure 2) [Allmendinger et al., 1997; Strecker et al., 2007]. The Cordillera Occidental is the locus of most active arc magmatism in this sector of the Andes since early Miocene time [Kay and Coira, 2009]. A complex of large, internally drained, late Cenozoic basins, local volcanic centers, and deformed bedrock outcrops occupies the Altiplano portion of the high plateau. Regional elevation in the Altiplano is ∼3800 m [Isacks, 1988; Masek et al., 1994]. The Puna part of the high plateau is approximately 400 m higher and more rugged than the Altiplano, and contains numerous smaller internally drained late Cenozoic basins separated by rugged mountain ranges composed of Precambrian and Paleozoic sedimentary, igneous, and low‐ grade metamorphic rocks. Neogene volcanic rocks, stratovolcanoes, and calderas of the magmatic arc extend eastward into the eastern part of the Puna (Figure 2). The Cordillera Oriental in northern Argentina is dominated by upper Proterozoic‐Carboniferous sedimentary rocks, and widespread exposures of Ordovician (ca. 470 Ma) and Cambrian (ca. 517 Ma) granitoid rocks (Figure 2). Within the study area, the Cordillera Oriental is composed of Proterozoic and Cambrian‐Carboniferous strata north of the northwest‐southeast–striking El Toro lineament, whereas only the Proterozoic metasedimentary and early Paleozoic igneous rocks are present to the south (Figure 2). [9] Major faults in the Cordillera Oriental are generally interpreted to have relatively steep trajectories in the uppermost crust, verging toward both the east and west and merging downward into a regional detachment [Cladouhos et al., 1994; Yuan et al., 2000; Kley and Monaldi, 2002]. Data bearing on the structure of the middle and lower crust are generally lacking, however, such that the presence of regional detachments is speculative. The Subandean ranges in Argentina are the southward continuation of the thin‐ skinned Subandean fold‐thrust belt that dominates the frontal part of the Bolivian Andes. These ranges are composed of Paleozoic through Cenozoic sedimentary rocks that are detached from structural ‘basement’ along a regional décollement below a thick Silurian shale unit [Allmendinger et al., 1983; Starck and Schulz, 1996; Echavarria et al., 2003].

TC6013

South of latitude 23°S, the frontal thrust belt is represented by the Santa Bárbara system, which is characterized by large uplifted blocks of Paleozoic‐Cenozoic sedimentary rocks that are bounded by relatively steep, generally west‐ verging reverse faults [Cristallini et al., 1997; Kley and Monaldi, 2002]. Cutting across the Santa Bárbara system, parts of the Cordillera Oriental, local areas in the eastern Puna, and also present in the subsurface beneath the modern foreland basin is the Salta rift, a complex branching array of extensional basins that formed during Aptian‐Campanian time [Salfity and Marquillas, 1994]. Reactivation of normal faults associated with the Salta rift partly controlled development of the Santa Bárbara ranges and structures within the Cordillera Oriental [Grier et al., 1991; Cladouhos et al., 1994; Cristallini et al., 1997; Kley and Monaldi, 2002; Monaldi et al., 2008]. Many previous workers have interpreted the lower Cenozoic deposits documented in this paper as the results of late‐stage thermal subsidence in the Salta rift system [e.g., Salfity and Marquillas, 1994; del Papa and Salfity, 1999], and petroleum company reflection seismic data from the Lomas de Olmedo arm of the rift indicate that the Santa Bárbara Subgroup was influenced by late‐stage thermal subsidence [Cominguez and Ramos, 1995]. East of the Andean topographic front lies the vast low‐elevation Chaco plain, which is the locus of upper Neogene to modern foreland basin sediment accumulation [Horton and DeCelles, 1997; Aalto et al., 2003; Chase et al., 2009].

3. Cenozoic Stratigraphy of Northwestern Argentina [10] Cretaceous through Cenozoic, predominantly clastic strata are widely distributed from the eastern Puna to the frontal Santa Bárbara ranges (Figures 2 and 3). These units are assigned to the Salta and Payogastilla (or Orán) Groups. The Salta Group is divided into three formal Subgroups: the Pirgua, Balbuena, and Santa Bárbara. Overlying the Salta Group is the Orán Group, which is composed of the Métan and Jujuy Subgroups (Figure 3) [Moreno, 1970; Reyes and Salfity, 1973]. The Neocomian‐lower Maastrichtian Pirgua Subgroup consists of conglomerate, sandstone, siltstone and alkaline volcanic/hypabyssal rocks that recorded the opening of the Salta rift [Galliski and Viramonte, 1988; Grier et al., 1991; Salfity and Marquillas, 1994]. Thickness of the Pirgua Subgroup ranges up to ∼6,000 m, and is controlled by the boundary faults of the Salta rift [Salfity and Marquillas, 1994; Marquillas et al., 2005]. The rift consists of three elongated sub‐basins arrayed about a central, structurally higher region referred to as the Salta‐Jujuy high. The Maastrichtian‐lower Paleocene Balbuena Subgroup comprises sandstone, shale, and limestone of the Lecho, Yacoraite, and Olmedo/Tunal Formations. These strata are more widespread than the Pirgua Subgroup, and consist of eolian, lacustrine, and fluvial deposits. Numerous workers have argued that the Balbuena Subgroup was deposited during an early phase of post‐rift thermal subsidence [e.g., Grier et al., 1991; Salfity and Marquillas, 1994; Marquillas et al., 2005]. [11] Above the Balbuena Subgroup lies the upper Paleocene‐Eocene Santa Bárbara Subgroup, which consists of the Mealla, Maiz Gordo, and Lumbrera Formations. These units are distributed throughout the eastern Puna, Cordillera Oriental, and Subandean zone (including the Santa Bárbara

3 of 30

TC6013


TC6013

Figure 2. Geological map of the study area in northwest Argentina and northern Chile, after Reutter et al. [1994]. Inset map (lower right) shows location of study area in the context of morphotectonic zones of the central Andes. Gray areas show the Santa Bárbara ranges (SBS). ranges). Deposition of the Santa Bárbara Subgroup took place in fluvial and shallow lacustrine environments, and paleosols are abundant, particularly in the Maiz Gordo and Lumbrera Formations. Starck and Vergani [1996] noted that

the Lumbrera Formation is divisible into a lower member that forms the upper part of the Santa Bárbara Subgroup, and an upper member that represents the fine‐grained distal equivalent of the Quebrada de los Colorados Formation in

4 of 30

TC6013


Figure 3. Stratigraphic chart for the eastern Puna, Cordillera Oriental, and Subandean zones of northwestern Argentina, based on Starck and Vergani [1996], Reynolds et al. [2000], and Marquillas et al. [2005]. Megasequences after Starck and Vergani [1996]. the overlying Métan Subgroup. The middle to upper Eocene Geste Formation crops out in local areas of the Puna and Eastern Cordillera [Jordan and Alonso, 1987; Alonso,

TC6013

1992], and consists of alluvial fan and fluvial deposits associated with local structural growth [Carrapa and DeCelles, 2008]. The Geste Formation rests upon Ordovician rocks and is not part of the regional Paleogene stratigraphic system of the Eastern Cordillera and Santa Bárbara Ranges (Santa Bárbara Subgroup). [12] The Métan Subgroup comprises several formations that have local names. For simplicity, we use the nomenclature from the Calchaqui Valley and Lerma Valley areas (Figure 3) for the western and eastern parts of the stratigraphic profile, respectively. In the Calchaqui Valley region, the lower unit of the Métan Subgroup is referred to as the Quebrada de los Colorados Formation, and the upper unit is named the Angastaco Formation (Figure 3) [Starck and Vergani, 1996]. In the northern part of the study area the Quebrada de los Colorados Formation is referred to as the Casa Grande Formation. Together these units form a several thousand‐meter‐thick upward coarsening succession, from mainly siltstone and sandstone in the lower part to cobble‐ boulder conglomerate in the upper part. In the Lerma Valley region, the Métan Subgroup is divided into the Rio Seco, Anta, and Jesús María Formations, which total ∼1,000 m in thickness and are collectively much finer‐grained than the Angastaco Formation. These three units are traceable into the frontal Subandean ranges [Gebhard et al., 1974; Hernández et al., 1996; Reynolds et al., 2000]. [13] The youngest units in the Cenozoic stratigraphic succession consist of the Palo Pintado, San Felipe, Guanaco, and Piquete Formations, which are generally younger than ∼9 Ma and contain abundant evidence for local tectonic activity in the form of angular and progressive unconformities [Starck and Anzótegui, 2001; Carrera and Muñoz, 2008; Carrapa et al., 2011a] and local provenance indicators [Hernández et al., 1999; Bywater‐Reyes et al., 2010; Carrapa et al., 2011a]. These units have been studied extensively by other groups [e.g., Gebhard et al., 1974; Hernández et al., 1996, 1999; Hernández and Echavarria, 2009; Reynolds et al., 2000; Starck and Anzótegui, 2001; Echavarria et al., 2003; Carrapa et al., 2011a; Hain et al., 2011] and must be taken into account in any regional analysis of the Cenozoic foreland basin system. [14] Starck and Vergani [1996] divided the Metán and Jujuy Subgroups into four unconformity‐bounded ‘megasequences’ that formed in response to progressive eastward migration of the foreland basin depositional framework (Figure 3). Megasequence I is represented by the Quebrada de los Colorados Formation and the eastward equivalent upper member of the Lumbrera Formation. Megasequence II consists of the Angastaco, Rio Seco, Anta, and Jesús María Formations, which form the bulk of the eastward tapering wedge of Cenozoic syntectonic strata exposed in the Cordillera Oriental and Subandean ranges. Megasequence III comprises the Palo Pintado, San Felipe, and Guanaco Formations, and Megasequence IV consists of the Piquete Formation.

4. Sedimentology [15] This work focuses on the Santa Bárbara and Métan Subgroups, which together form a dramatically eastward tapering prism of clastic syntectonic strata. Our sedimentological data include 16 detailed measured stratigraphic sections totaling >14 km in thickness and >1000 paleocur-

5 of 30

TC6013

TC6013


Table 1. Lithofacies (and Their Codes) Used in Logs of Measured Sections, and Interpretations in This Study, Modified After Miall [1978] and DeCelles et al. [1991] Lithofacies Code Fsl Fcl Fsm Sm Sr St Sp Sh Gcm Gcmi Gch Gchi Gct Gmm M

Description

Process Interpretation

Laminated black or gray siltstone Laminated gray claystone Massive, bioturbated, mottled (gleyed) siltstone; red, purple, gray; carbonate or iron‐hydroxide nodules common Massive medium‐ to fine‐grained sandstone; bioturbated Fine‐ to medium‐grained sandstone with small, asymmetric, 2D and 3D current ripples Medium‐ to very coarse‐grained sandstone with trough cross‐stratification Medium‐ to very coarse‐grained sandstone with planar cross‐stratification

Suspension‐settling in ponds and lakes Suspension‐settling in ponds and lakes Paleosols, usually calcic, histic, or vertic

Bioturbated sand, penecontemporaneous deformation Migration of small ripples under weak (∼20–40 cm/s), unidirectional flows in shallow channels Migration of large 3D ripples (dunes) under moderately powerful (40–100 cm/s), unidirectional flows in large channels Migration of large 2D ripples under moderately powerful (∼40–60 cm/s), unidirectional channelized flows; migration of sandy transverse bars Fine‐ to medium‐grained sandstone with plane‐parallel lamination Upper plane bed conditions under unidirectional flows, either strong (>100 cm/s) or very shallow Pebble to boulder conglomerate, poorly sorted, clast‐supported, Deposition from sheetfloods and clast‐rich debris flows unstratified, poorly organized Pebble to cobble conglomerate, moderately sorted, clast‐supported, Deposition by traction currents in unsteady fluvial flows unstratified, imbricated (long‐axis transverse to paleoflow) Pebble to cobble conglomerate, well sorted, clast‐supported, Deposition from shallow traction currents in longitudinal horizontally stratified bars and gravel sheets Pebble to cobble conglomerate, well sorted, clast‐supported, Deposition from shallow traction currents in longitudinal horizontally stratified, imbricated bars and gravel sheets (long‐axis transverse to paleoflow) Pebble conglomerate, well sorted, clast‐supported, Deposition by large gravelly ripples under traction flows trough cross‐stratified in relatively deep, stable fluvial channels Deposition by cohesive mud‐matrix debris flows Massive, matrix‐supported pebble to boulder conglomerate, poorly sorted, disorganized, unstratified Micritic massive gray and yellow marl Lacustrine carbonate mud

rent measurements. The measured sections are representative of the numerous other sections that we observed in the field over six field seasons, and are augmented by sections studied in Bolivia [Horton et al., 2001; DeCelles and Horton, 2003; Horton, 2005] and sections documented by other groups in Argentina (most notably Starck and Vergani [1996], Hernández et al. [1999], Reynolds et al. [2000, 2001], Echavarria et al. [2003], del Papa et al. [2010], and Carrapa et al. [2011a, 2011b]). In the following brief sedimentological descriptions, standard codes are used to denote lithofacies as summarized in Table 1. 4.1. Mealla and Maiz Gordo Formations [16] Sandy lithofacies of the Mealla and/or Maiz Gordo Formations overlie the upper Cretaceous Yacoraite Formation (and locally the Pirgua Subgroup). At many of the localities studied these two units are difficult to consistently separate using lithological criteria, so we find it useful to consider them together. The composite thickness of the Mealla and Maiz Gordo Formations ranges between ∼10 m and 121 m. [17] The Mealla Formation consists of medium‐ to coarse‐grained, pink and tan sandstone interbedded with red siltstone. Sandstone lithofacies are dominated by Sh, Sr, and St (Table 1) in beds generally less than 50 cm thick. In the Cachi and Susques regions, the Mealla Formation contains pebbly conglomerate and sandstone beds rich in milky‐ quartz clasts. Mealla Formation siltstones are massive (Fsm) and contain abundant carbonate nodules (Figure 4a). These nodules are micritic, with features characteristic of pedogenic glaebules, including sparry craze veins, mottling, and compound nodular textures. The nodules are commonly present in the lower parts of massive siltstone units. Paleo-

current data from trough cross‐stratification (method I of DeCelles et al. [1983]) indicate paleoflow direction generally toward the east‐northeast (Figures 5 and 6). [18] The Maiz Gordo Formation is composed of pink and purple siltstone and coarse‐grained to conglomeratic sandstone. The abundance of sandstone beds varies dramatically; some sections of the Maiz Gordo are composed almost exclusively of sandstone, whereas others are rich in siltstone. Lithofacies and primary sedimentary structures in the Maiz Gordo are almost completely overprinted by post‐ depositional, pedogenic processes. Pedogenic nodules of carbonate and iron‐manganese oxide are abundant. Large vertical burrows similar to Krausichnus trompitus [Hasiotis and Bown, 1992; DeCelles and Horton, 2003] are abundant at many localities (Figure 4b). 4.2. Lumbrera Formation [19] The Lumbrera Formation is composed of distinctive, bright brick‐red siltstone with variable amounts of thin‐ bedded sandstone and, in some sections, thin layers of gray/ white marl (Figure 4c) and stromatolitic limestone. One or two zones of laminated gray and green siltstone, typically 5– 10 m thick, are present in the lower part of the Lumbrera Formation in some areas (e.g., section AL); these are referred to informally as the “Faja Verde” beds, and are generally interpreted as perennial lacustrine deposits [del Papa et al., 2002]. Starck and Vergani [1996] divided the Lumbrera into informal lower and upper members separated by a disconformity, and correlated the upper member with the more proximal Quebrada de los Colorados Formation. The lower member ranges between 70 m and 485 m thick, is coarser‐grained than the upper member, contains perennial lacustrine deposits of the Faja Verde and, in some sections,

6 of 30

TC6013


numerous paleosol layers. The upper member is generally >100 m thick, and merges gradationally upward with coarser‐grained lithofacies more typical of the Quebrada de los Colorados Formation. The upper member also contains ephemeral lacustrine deposits [del Papa et al., 2002;

TC6013

del Papa, 2006]. Paleosols in the Lumbrera Formation are dominated by nodular Calcisols [Mack et al., 1993], but Vertisols, Gleysols, and nodular Histosols are also present (see section 4.3). 4.3. Zone of Intense Pedogenesis [20] A distinctive zone of extremely mature, multistory paleosols is present at various stratigraphic levels in the Maiz Gordo and Lumbrera Formations. The interval, which we refer to informally as the “supersol zone,” is one of the most striking features in the Cenozoic stratigraphy of northwestern Argentina (Figure 7a). The supersol zone is typically 50–100 m thick, and grades upward and downward into fluvial and lacustrine lithofacies that either lack paleosols or contain only less mature paleosols intercalated with fluvial channel deposits. Paleosol types in the supersol zone include Calcisols, Gleysols, Histosols, and Vertisols; Calcisols and Histosols are the most abundant (Figures 4a and 7b–7d). The zone of Histosols is present in nearly every section in which we documented the supersol zone. It is characterized by generally gray or purple colors with numerous vertically elongated nodules of rusty‐colored iron‐ manganese oxide (Figure 7c). Mottling (gleying) is ubiquitous. Paleosols in the supersol zone are stacked on top of each other in an unbroken succession, regardless of original grain‐size; original sedimentary structures are completely obliterated throughout the supersol zone. [21] In some places in the eastern Cordillera Oriental, the supersol zone is absent and the contact between Eocene Lumbrera Formation and lower Miocene strata is an erosional disconformity. In the Subandean zone and Santa Bárbara ranges the supersol is either not present or is formed on top of pre‐Cenozoic rocks (e.g., section CV) [Echavarria et al., 2003; Hernández and Echavarria, 2009]. [22] The stratigraphic level of the supersol/disconformity generally climbs eastward. In the southern transect, its level ascends from just a few meters above the Pirgua Subgroup in the Pucara section (section PU), to the Maiz Gordo and lower Lumbrera Formations in the Monte Nieva, Tin Tin, and Obelisco sections (MN, 2TT, OB), to the upper Lumbrera Formation in the Alemanía (AL) section (Figure 5). The supersol is not as obviously developed at locations east of Alemanía; however, in the Pampa Grande (PG) section, numerous moderately developed paleosol horizons in the upper Lumbrera Formation could represent the supersol. Although del Papa et al. [2010] did not observe the supersol zone in their Simbolar section, the stratigraphy they depict is similar to, but thicker than, what we documented at Alemanía and Pampa Grande (Figure 5). These authors also reported a U‐Pb zircon age of ∼40 Ma from multigrain fractions Figure 4. (a) Nodular Calcisol in the Mealla Formation. (b) Krausichnus trompitus trace fossils, attributed to colonial insects such as termites [see DeCelles and Horton, 2003]. Definitive aspects of the trace include the central vertical shaft, from which branching galleries spread outward. (c) Laminated marly siltstone in the Lumbrera Formation in the Tres Cruces section (TC). (d) Interbedded fluvial channel deposits (pale pink) and massive loessite (darker red) in the Monte Nieva (MN) section. Thickness of section shown is ∼60 m.

7 of 30

TC6013


Figure 5. Measured stratigraphic sections of the Paleocene‐Miocene strata in the southern part of the study area. 8 of 30

TC6013

TC6013


TC6013

Figure 6. Measured stratigraphic sections of the Paleocene‐Miocene strata in the northern part of the study area. Section CZ is from D. Gingrich (unpublished data, 2010). separated from a tuffaceous layer in the upper Lumbrera Formation. Comparing the Simbolar section of del Papa et al. [2010] with our nearby sections from Pampa Grande (PG) and Alemanía (AL) (Figure 5) suggests that if the supersol zone is present in the Simbolar area, it is either

younger than the ca. 40 Ma tuff or it resides directly below the tuff in the upper Lumbrera Formation. If the supersol is absent at Simbolar, then the contact between the basal Neogene and the Lumbrera Formation is a major disconformity, separating rocks that are as old as ∼20 Ma above

9 of 30

TC6013


from rocks that are approximately 40 Ma below. Farther east at the Rio González section [Reynolds et al., 2000] the Neogene rests disconformably upon the Lumbrera Formation and no supersol has been documented [Hernández et al., 1999].

TC6013

[23] In the northern transect (Figure 6), the supersol zone steps up section eastward from the Maiz Gordo Formation in the Susques section (SQ) to the upper Lumbrera Formation in the Tres Cruces section (TC). In the Cianzo syncline section (CZ), the supersol zone is located in the Maiz Gordo and upper part of the lower Lumbrera Formation. As in the southern transect, the supersol zone is either not as thickly developed or is accompanied by a major erosional disconformity in sections in the Subandean zone (section CV, Figure 6). In the Subandean zone of northernmost Argentina, Reynolds et al. [2001] interpreted the Tranquitas Formation as a zone of possible Paleogene or earliest Neogene stratigraphic condensation, which is capped by a significant disconformity. [24] Overall, progressively younger strata from west to east overlie the supersol/disconformity [Starck and Vergani, 1996; Hernández et al., 1999]. In the west the ca. 40 Ma Quebrada de los Colorados Formation sits above the supersol, whereas in the east the ca. 20 Ma Rio Seco Formation rests on top of the supersol or the disconformity. In addition, the time‐ span of the disconformity increases eastward [Hernández et al., 1999]. Together, the supersol and disconformity may be taken to represent different manifestations of similar conditions of extremely slow or zero net sediment accumulation. [25] A remarkably similar zone of intense pedogenesis was documented by Horton et al. [2001] and DeCelles and Horton [2003] in Eocene strata of the Bolivian Cordillera Oriental. Uba et al. [2006] extended the Bolivian supersol zone into the upper Oligocene‐lower Miocene section of the western Subandean zone in Bolivia. 4.4. Quebrada de los Colorados Formation (Casa Grande Formation) [26] The Quebrada de los Colorados Formation (Casa Grande Formation in the northern part of the study area) consists of an upward coarsening succession of sandstone and conglomeratic sandstone, with interbedded brick‐red silstone beds (Figure 4d). This unit is roughly correlative with the upper member of the Lumbrera Formation [Starck and Vergani, 1996], and where the two are present and distinguishable, the Quebrada de los Colorados Formation is much coarser‐grained than the Lumbrera Formation. Coarse‐ grained units are broadly lenticular with erosional basal surfaces and crude upward fining grain‐size trends in ∼2– 5 m‐thick compound depositional units. Trough cross‐ stratification (St) is most common in the lower parts of these upward fining sequences, giving way to ripples (Sr) and plane beds (Sh) in the upper parts. Imbricated (Gcmi), horizontally stratified (Gch) and trough cross‐stratified (Gct) conglomerates are common in the lower parts of these units as well. In the Monte Nieva section, lenticular sandstone and Figure 7. (a) Lower part of measured section 2TT, showing Mealla and Maiz Gordo Formations in the foreground and middle distance. Dark gray band is the Histosol interval within the supersol zone. (b) Thin bed of fluvial sandstone riddled with Krausichnus trompitus burrows, resting upon a gleyed nodular Calcisol with large root traces; section CZ. Hammer for scale. (c) Iron‐oxide nodules in blue‐gray Histosol in the Susques section (SQ). (d) Vertisol in Lumbrera Formation at the Alemanía section (AL).

10 of 30

TC6013


TC6013

data from trough cross‐stratification and imbricated clasts indicate eastward and northeastward paleoflow (Figures 5 and 6). [27] We interpret the Quebrada de los Colorados Formation as the deposits of a sandy to gravelly braided fluvial system. The upward coarsening and thickening intervals of stacked channel deposits that are common in this unit could represent distributary fluvial behavior [Hartley et al., 2010], such as seen on fluvial megafans in the modern Andean foreland basin [Horton and DeCelles, 2001].

Figure 8. (a) Eolian large‐scale cross‐stratification produced by dune slipface deposition in the Angastaco Formation in section MN. Height of cliff is ∼25 m. (b) Well‐ organized lithofacies Sh and Gch in lower Angastaco Formation at section MN. (c) Agujas Conglomerate in section QT, showing changes in dip of bedding from overturned dipping west (left) to upright dipping east (right). Person for scale highlighted in ellipse at bottom center. conglomerate bodies are stacked in 8–15 m thick packages that coarsen and thicken upward, and are capped by sharp, flat surfaces marking an abrupt transition to siltstone. The fine‐grained intervals between coarse‐grained units are typically massive siltstone and silty sandstone, usually red, with little internal sedimentary structure (Fsm). Paleocurrent

4.5. Angastaco Formation [28] The transition from the Quebrada de los Colorados Formation to the overlying Angastaco Formation is gradational [see also Carrapa et al., 2011a]. The contact is generally placed at the base of a prominent interval of well‐sorted medium‐grained sandstone containing large‐ scale cross‐stratification typical of eolian dune slipface deposits (Figure 8a) [Starck and Vergani, 1996; Carrapa et al., 2011a]. In the Monte Nieva section the eolian interval is approximately 800 m thick, and contains distinctive, orange‐colored, homogeneously structureless fine‐ to medium‐grained sandstone units that we interpret to be coarse‐grained loessite [e.g., Soreghan et al., 2008]. Individual units of the coarse‐grained loessite are up to ∼15 m thick, and are commonly interbedded with units of coarse‐ grained to pebbly trough cross‐stratified sandstone (Figure 4d). The loessite facies are gradually replaced up‐section by cross‐stratified dune facies (Figure 8a) that indicate eastward paleowind directions (Figure 5, section MN), suggesting that the core of the erg migrated eastward over the top of a fringing loess belt [e.g., Crouvi et al., 2010]. The eolian interval is much thinner (less than 100 m) in the Angastaco area. [29] The upper 2,000 m of the Angastaco Formation are dominated by pebble to cobble conglomerate and very coarse‐grained to pebbly sandstone beds (Figure 8b). Bedding is highly lenticular, with erosional basal contacts and crude upward fining sequences. The predominant lithofacies are well‐organized conglomerate (Gcm, Gct, Gcmi), and horizontally stratified (Sh) and trough cross‐stratified (St) sandstone (Table 1). The assemblage of lithofacies in the upper Angastaco Formation is typical of coarse‐grained low‐ sinuosity (braided) fluvial systems. Although we did not observe any clear‐cut evidence of sediment‐gravity flow deposits, it is plausible that the coarser‐grained portions of the Angastaco Formation were deposited in medial to distal alluvial fan environments [Carrapa et al., 2011a]. Imbricated conglomerates (Gcmi and Gchi; clast long‐axes perpendicular to paleoflow direction) indicate generally eastward paleoflow in the azimuthal range 000°–140° (Figures 5 and 6). [30] The Angastaco Formation tapers rapidly eastward from its type area in the Calchaqui valley to the Lerma and Los Conchas valleys (compare sections MN and AL, Figure 5), where roughly equivalent strata are represented by the Rio Seco, Anta, and Jesús María Formations. The Rio Seco Formation consists of a few tens of meters of medium‐grained, cross‐stratified sandstone, resting unconformably upon the Lumbrera Formation. The Rio Seco Formation has also been interpreted as eolian and fluvial deposits [Starck and Vergani, 1996]. In the Alemanía section (AL, Figure 5) the Rio Seco consists of sandy trough cross‐stratified fluvial facies. The

11 of 30

TC6013


Anta Formation is ∼265–720 m thick in the Alemanía area and the Santa Bárbara ranges [Reynolds et al., 2000], and consists of red fine‐grained sandstone and mudrocks, with occasional thin carbonate layers that have been interpreted as marginal marine deposits [Galli et al., 1996]. The Jesús María Formation consists of ∼1000–1200 m of fluvial sandstone and mudstone. Additional information about the Rio Seco, Anta, and Jesús María Formations may be found in the works by Gebhard et al. [1974], Galli et al. [1996], and Reynolds et al. [2000]. 4.6. Agujas Conglomerate, Quebrada del Toro Section [31] We measured and sampled an approximately 1200 m thick section of the Agujas Conglomerate [Marrett and Strecker, 2000] (section QT), which crops out along Quebrada Cerro Bayo, a tributary arroyo to the Quebrada del Toro (Figure 5). This unit is age‐equivalent to the upper part of the Angastaco Formation and the lowermost part of the Palo Pintado Formation (Figure 3). The predominant lithofacies include pebble to cobble conglomerate (Gcm) and massive, red, coarse‐grained sandstone (Sm) (Table 1). Several thin beds of light‐colored tuffaceous sandstone and tuff were observed and sampled for geochronological analysis. The section is folded into a north‐trending syncline, and our measured section is located in the western limb of the syncline. The western boundary of the section is marked by the Solá reverse fault, which juxtaposes phyllite and quartzite of the Proterozoic‐lower Cambrian Puncoviscana Formation with the Agujas Conglomerate [Marrett and Strecker, 2000]. Bedding in the western limb of the syncline is overturned and dipping westward, but becomes upright and decreases in magnitude from ∼70°E to 25°E upward through the measured section (Figure 8c). Faulting obscures the nature of the basal contact of the Agujas Conglomerate, but its stratigraphic context can be determined by tracing the interval along strike toward the south, where it rests directly above strata that correlate with the Quebrada de los Colorados Formation in the Tin Tin and Monte Nieva areas. We therefore infer that this interval is roughly equivalent to the upper part of the Angastaco Formation; this interpretation is borne out by U‐Pb zircon ages from tuffaceous layers (see section 5). [32] We tentatively suggest that the progressive up‐section decrease in dip magnitude reflects syndepositional structural growth in response to slip on the Solá fault. This interpretation is consistent with findings of Mazzuoli et al. [2008], who reported growth structures in the Agujas Conglomerate 10–15 km north of our measured section. The coarse grain‐ size and association of lithofacies in the Agujas Conglomerate suggest deposition in an alluvial fan system, and the presence of a progressive unconformity and clast‐compositions dominated by the Puncoviscana Formation indicate close proximity to the topographically rising source area in the hanging wall of the Solá fault.

5. Chronostratigraphy [33] The ages of Cenozoic stratigraphic units in northwestern Argentina are known from a locally rich fossil assemblage (including both vertebrate fossils and palynology) [Quattrocchio et al., 1997; del Papa and Salfity, 1999; Hongn et al., 2007; del Papa et al., 2010], paleomagnetic stratigraphy [Reynolds et al., 2000, 2001; Echavarria et al.,

TC6013

2003], radio‐isotopic ages from tuffaceous layers [Reynolds et al., 2000], and recently acquired U‐Pb ages from tuffaceous rocks and detrital samples [DeCelles et al., 2007; Carrapa et al., 2008, 2011a, 2011b; Bywater‐Reyes et al., 2010; Hain et al., 2011] (see also new ages reported in this study; Figures 9 and 10). The ages of the Mealla and Maiz Gordo Formations are based on paleontology and are well established as early–mid‐Paleocene and late Paleocene‐ early Eocene, respectively [see del Papa and Salfity, 1999, and references therein]. The Lumbrera Formation is also dated mainly by paleontology as middle to late Eocene [del Papa and Salfity, 1999; Hongn et al., 2007]. Our observations suggest that most of the Eocene is represented by the supersol. In the Tres Cruces section (section TC, Figure 6), a paleosol carbonate nodule from the 330 m level yielded a U‐Pb mean age of 47 ± 7 Ma (MSWD = 7.5, 95% confidence; M. N. Ducea, unpublished data, 2011). This sample was collected from just a few meters below the top of the supersol zone, and is interpreted as the approximate age of formation of the paleosol carbonate. A ca. 40 Ma U‐Pb age from zircon in a tuffaceous layer was reported by del Papa et al. [2010] in the upper Lumbrera Formation at Simbolar, and we obtained an age of ∼40.6 ± 0.3 Ma from a single detrital zircon in a sample from the Quebrada de los Colorados Formation (approximately equivalent to the upper member of the Lumbrera Formation) at the 755 m level of section MN, 230 m above the top of the supersol zone (Figures 5 and 9). An additional sample from the lower part of the Quebrada de los Colorados Formation in the Angastaco area yielded a cluster of detrital zircon U‐Pb ages with mean age of ∼37.6 ± 2.0 Ma [Carrapa et al., 2011a]. These ages are consistent with middle Eocene paleontological ages from the lower part of the Quebrada de los Colorados Formation in the la Poma area [Hongn et al., 2007]. Another sample from a cobble in the Quebrada de los Colorados Formation in section MN (1333 m level) produced an apatite fission track pooled age of 28.7 ± 1.9 Ma (Table 2). Because other cobbles sampled farther upsection yielded similar (within errors) apatite fission track pooled ages, we interpret this age as a true indication of the age of exhumation in the source terrane and not a result of partial annealing during burial. Therefore the depositional age of this sample must be younger than ∼28.7 Ma. [34] Samples of sandstone from the Geste Formation in the eastern Puna (section SP, Figure 5) yielded minimum U‐Pb zircon age clusters between ∼38 Ma and 34 Ma [DeCelles et al., 2007], and fission track and U‐Pb double dates from detrital apatites in the Geste Formation support a late Eocene depositional age [Carrapa et al., 2008] that is consistent with vertebrate paleontology from the same section [Pascual, 1983; Alonso, 1992]. [35] Sandstone samples from several sections produced early Miocene minimum detrital zircon U‐Pb age clusters, providing useful constraints on depositional ages of the Angastaco Formation and its lateral equivalents (Figure 9). Additional detrital zircon ages reported by [Carrapa et al., 2011a] provide constraints on ages of upper Miocene units in the region. Detrital zircons separated from sandstones in the Angastaco and Rio Seco Formations yielded numerous Miocene U‐Pb ages, supporting a depositional age range of ∼19–9 Ma for the Angastaco [Carrapa et al., 2011a], and a 5% and central when c2 < 5%. Error is one s, calculated using the zeta calibration method [Hurford and Green, 1983] and the following Zeta. For MN2936 and MN1333, 369.69 ± 6.82; samples analyzed using a Leica DMRM microscope (UW) at 1250 magnification. For MN3, 366.44 ± 17.31; sample analyzed using Olympus microscope (UA) at 1600 magnification. Analyses conducted using a drawing tube located above a digitizing tablet and kinetic computer‐controlled stage driven by the FTStage program [Dumitru, 1993]. Samples were irradiated at Oregon State University. Samples were etched in 5.5 molar nitric acid at 21°C for 20 seconds. Following irradiation, the mica external detectors were etched with 21°C in 40% hydrofluoric acid for 45 minutes. N is the number of individual crystals dated. Rho‐S and Rho‐I are the spontaneous and induced track density measured, respectively (tracks/cm2). NS and NI are the number of spontaneous and induced tracks counted, respectively. The chi‐square probability is c2 (%) [Galbraith and Green, 1990; Green, 1981]. Values greater than 5% are considered to pass this test and represent a single population of ages. Rho‐D is the induced track density in external detector adjacent to CN5 dosimetry glass (tracks/cm2). ND is the number of tracks counted in determining Rho‐D. Dpar is fission track etch pit measurements; SD is the related standard deviation. a

and volcanic clast types, mainly quartzite and andesite‐ dacite. In the Mealla and Maiz Gordo Formations milky quartz pebbles are common. Granitoid and granitic gneiss clasts are also present in some sections in the Quebrada de los Colorados and Angastaco Formations (e.g., section MN). [41] Apatite fission track data were collected from three granitic clasts from the MN section. Sample preparation followed standard procedures and analyses were performed using the zeta age calibration method [Hurford and Green, 1983]. More information can be found in Table 2. The three conglomerate clasts produced pooled ages of 28.7 ± 1.9 Ma, 29.7 ± 2.7 Ma, and 26.5 ± 2.3 Ma (Table 2). Clast sample MN1333, with a pooled AFT age of 28.7 ± 1.9 Ma, also produced a U‐Pb zircon mean age of 486.7 ± 6.1 Ma (MSWD = 1.2) (Table S1 in the auxiliary material).1 This age suggests that the clast (and the many others like it in the MN section) was derived from local Ordovician granitoid rocks such as those exposed in the nearby Nevado de Cachi. A sample from 6150 m elevation in the Nevado de Cachi produced an AFT pooled age of 28.2 ± 3.6 Ma (B. Carrapa, unpublished data, 2011), consistent with rapid exhumation at that time near the orogenic front.

more precise for older ages, we rely on 206Pb/238U ages up to 1000 Ma and 206Pb/207Pb ages if the 206Pb/238U ages are >1000 Ma [Gehrels et al., 2008]. These analyses are plotted on relative age‐probability diagrams (Figures 9 and 12), which represent a sum of the probability distributions of all analyses from a sample, normalized such that the areas beneath the probability curves are equal for all samples. Age peaks on these diagrams are considered robust if defined by several analyses. Table 3. Modal Petrographic Point‐Counting Parameters Symbol Qm Qp Qpt Qms C S Qt K P

7. Detrital Zircon Geochronology [42] Twenty‐two samples of medium‐ to coarse‐grained sandstone were processed by standard methods for dense minerals, and detrital zircon grains were separated from these concentrates. Zircons were mounted in epoxy, polished and analyzed for U‐Pb ages by laser ablation multicollector inductively coupled plasma mass spectrometry (LA‐MC‐ ICPMS) at the University of Arizona LaserChron Center. Details of the method are described by Gehrels et al. [2008]. A total of 1,925 grains from the Cenozoic samples produced data of sufficient precision for geochronological interpretation. We also obtained 524 U‐Pb zircon ages from six samples of key Paleozoic and Mesozoic rock units in the Puna and Cordillera Oriental, in order to provide a basis for interpreting the provenance of the zircons in the Cenozoic samples. Analyses that yielded isotopic data of acceptable discordance, in‐run fractionation, and precision are shown in Tables S1 and S2. Because 206Pb/238U ages are generally more precise for younger ages whereas 206Pb/207Pb ages are

F Lvma Lvf Lvv Lvx Lvl Lv Lsh Lph Lsm Lc Lm Ls Lt L Accessory minerals

1 Auxiliary materials are available in the HTML. doi:10.1029/ 2011TC002948.

15 of 30

Description monocrystalline quartz polycrystalline quartz foliated polycrystalline quartz monocrystalline quartz in sandstone or quartzite lithic grain chert siltstone total quartzose grains (Qm + Qp + Qpt + Qms + C + S) potassium feldspar (including perthite, myrmekite, microcline) plagioclase feldspar (including Na and Ca varieties) total feldspar grains (K + P) mafic volcanic grains (epidote ± pyx ± plag) felsic volcanic grains (sericite + qtz ± feldspar) vitric volcanic grains microlitic volcanic grains lathwork volcanic grains total volcanic lithic grains (Lvm + Lvf + Lvv + Lvx + Lvl) Mudstone Phyllite schist (mica schist) carbonate lithic grains total metamorphic lithic grains (Lph + Lsm + Qpt) total sedimentary lithic grains (Lsh + Lc + C + S + Qms) total lithic grains (Ls + Lv + Lm + Qp) total nonquartzose lithic grains (Lv + Ls + Lph + Lsm + Lc) Epidote/Zoisite, Chlorite, Muscovite, Biotite, Zircon, Sphene, Clinopyroxene, Orthopyroxene, Monazite, Magnetite

TC6013

TC6013


Table 4. Recalculated Modal Petrographic Data From Sections MN and 1TC Sample

Qm%

F%

Lt%

Qt%

F%

L%

Qm%

P%

K%

Lmet%

LV%

Lsed%

MN1 MN13 MN29 MN93 MN176 MN318 MN491 MN595 MN610 MN659 MN723 MN758 MN765 MN801 MN893 MN1026 MN1123 MN1262 MN1371 MN1405 MN1663 MN1665 MN1772 MN1986 MN2094 MN2179 MN2226 MN2290 MN2552 MN2702 MN2886 MN3492 MN3559 MN3908 MN4198 MN4340 Averages

59.8 56.8 57.6 64.7 53.3 75.7 54.4 56.3 62.9 55.9 59.1 74.8 49.8 41.3 59.9 44.2 47.0 75.5 57.8 51.9 66.4 44.8 71.6 46.5 46.1 44.3 50.0 55.6 42.6 43.2 48.2 46.1 41.1 42.9 48.0 39.4 53.8

19.1 24.6 27.5 22.3 31.1 9.1 29.5 19.4 23.0 26.2 34.2 15.9 34.0 34.5 23.0 27.0 28.0 20.7 29.5 30.8 24.8 44.8 18.5 31.4 25.3 30.1 29.6 26.5 24.0 27.5 26.9 29.0 26.6 27.4 29.5 25.9 26.6

21.1 18.6 15.0 12.9 15.6 15.1 16.1 24.3 14.2 17.9 6.7 9.3 16.2 24.2 17.2 28.8 25.0 3.8 12.7 17.2 8.8 10.3 9.9 22.0 28.6 25.6 20.4 17.9 33.4 29.3 24.9 24.9 32.4 29.7 22.5 34.6 19.6

74.4 68.3 67.4 72.3 64.6 86.1 65.6 73.1 70.8 64.6 62.9 80.5 58.0 54.5 67.0 56.4 56.5 78.3 63.9 58.9 70.8 48.2 76.4 56.0 60.1 57.1 60.7 63.9 55.4 51.3 57.8 53.9 50.6 56.7 57.7 47.1 63.0

18.1 23.9 25.9 21.8 29.9 8.8 28.5 18.4 22.6 25.8 34.0 15.7 33.0 33.5 22.2 25.8 26.9 20.6 28.5 29.8 24.6 44.1 18.3 30.3 24.1 28.8 28.5 25.4 23.0 26.6 26.2 28.3 25.8 26.1 29.0 25.6 25.8

7.5 7.8 6.7 5.9 5.5 5.2 5.9 8.5 6.6 9.6 3.1 3.8 9.1 12.0 10.8 17.8 16.6 1.1 7.6 11.4 4.6 7.7 5.3 13.7 15.8 14.2 10.8 10.8 21.7 22.1 16.0 17.8 23.6 17.1 13.4 27.4 11.2

75.8 69.8 67.7 74.4 63.2 89.2 64.9 74.4 73.2 68.0 63.4 82.5 59.4 54.5 72.3 62.1 62.7 78.5 66.2 62.7 72.8 50.0 79.5 59.7 64.6 59.5 62.8 67.8 64.0 61.1 64.2 61.4 60.7 61.0 61.9 60.4 66.6

17.9 9.3 29.7 11.5 31.8 6.3 22.4 6.8 17.8 16.8 23.6 14.7 40.3 24.1 24.9 32.2 17.1 9.3 20.2 33.7 16.9 31.8 20.5 31.7 16.9 10.6 13.0 18.3 15.2 16.8 10.2 13.5 25.1 26.5 16.4 23.5 19.9

6.3 20.9 2.6 14.1 5.0 4.5 12.7 18.8 8.9 15.2 13.0 2.9 0.3 21.4 2.8 5.7 20.2 12.3 13.6 3.6 10.4 18.2 0.0 8.6 18.5 29.9 24.2 13.9 20.9 22.2 25.6 25.1 14.2 12.5 21.7 16.1 13.5

48.4 33.3 67.4 36.7 52.6 40.5 38.5 37.6 32.5 37.3 26.3 66.7 46.0 50.0 64.7 64.6 54.8 28.6 52.3 74.1 56.5 32.4 54.2 55.7 66.7 71.1 60.3 63.0 57.3 65.1 76.6 72.1 61.2 64.9 52.9 66.7 53.6

4.8 2.1 2.3 0.0 0.0 9.5 7.7 7.1 15.0 19.6 36.8 22.2 26.0 22.9 23.5 5.1 15.5 14.3 11.4 6.9 0.0 54.1 33.3 28.6 11.9 9.2 24.1 13.0 18.2 18.3 15.6 9.3 17.4 13.8 25.0 24.6 15.8

46.8 64.6 30.2 63.3 47.4 50.0 53.8 55.3 52.5 43.1 36.8 11.1 28.0 27.1 11.8 30.3 29.8 57.1 36.4 19.0 43.5 13.5 12.5 15.7 21.4 19.7 15.5 24.1 24.5 16.5 7.8 18.6 21.5 21.3 22.1 8.7 30.6

1TC104 1TC194 1TC378 1TC380 1TC416 1TC456 1TC462 1TC468 1TC502 1TC530 1TC568 1TC651 1TC675 1TC695 1YC725 1TC852 1TC948 1TC1069 1TC1260 1TC1295 Averages

64.6 74.8 87.5 89.7 93.3 70.7 79.5 81.7 72.0 75.9 70.3 62.8 56.0 59.4 54.4 46.4 52.6 46.5 49.4 48.4 66.8

7.1 16.9 7.1 7.4 5.4 11.1 8.0 8.5 13.2 15.3 15.4 16.7 18.9 7.1 15.8 17.3 18.5 22.3 14.6 13.7 13.0

28.3 8.2 5.3 2.9 1.3 18.2 12.5 9.8 14.8 8.8 14.3 20.5 25.2 33.5 29.8 36.3 29.0 31.2 36.0 37.8 20.2

64.6 81.1 90.7 90.8 93.9 77.1 87.2 87.6 78.2 79.9 75.9 67.9 63.2 66.1 62.8 54.6 59.5 52.7 57.9 57.3 72.4

7.1 16.9 7.1 7.4 5.4 10.8 7.8 8.2 13.1 15.2 15.1 16.2 18.6 7.0 15.5 16.8 18.1 22.1 14.1 13.2 12.8

28.3 2.0 2.2 1.8 0.7 12.1 5.0 4.1 8.7 4.9 9.0 15.9 18.2 26.9 21.7 28.6 22.4 25.2 28.0 29.5 14.8

90.1 81.6 92.5 92.4 94.5 86.4 90.8 90.6 84.5 83.2 82.0 79.0 74.8 89.3 77.5 72.9 74.0 67.5 77.2 77.9 82.9

9.9 8.0 6.4 5.5 2.7 12.2 4.8 6.4 15.0 15.8 16.7 18.4 17.7 9.7 15.5 24.3 11.6 31.8 15.4 18.5 13.3

0.0 10.4 1.2 2.1 2.7 1.4 4.3 3.0 0.5 1.0 1.3 2.6 7.5 1.0 7.0 2.8 14.4 0.7 7.4 3.6 3.7

0.0 22.2 12.5 10.0 0.0 29.7 65.4 65.2 26.0 75.0 56.8 44.2 36.5 13.4 27.0 38.7 41.1 28.7 63.1 50.3 35.3

0.0 44.4 37.5 70.0 50.0 29.7 19.2 17.4 12.0 4.2 22.7 5.8 27.1 17.9 38.5 44.4 25.9 55.7 15.6 28.2 28.3

100.0 33.3 50.0 20.0 50.0 40.6 15.4 17.4 62.0 20.8 20.5 50.0 36.5 68.7 34.4 16.9 33.0 15.6 21.3 21.5 36.4

[43] Almost all of the Cenozoic samples exhibit three age clusters at ∼500 Ma, 555–994 Ma, and ∼1050 Ma (Figure 9). In detail, the ca. 500 Ma peak is divided into two subpopulations with ages of 470–491 Ma (latest Cambrian‐Early Ordovician) and 522–544 Ma (latest Proterozoic‐Early Cambrian). The ∼1050 Ma population ranges between 1024 Ma and 1096 Ma, with a few outliers in the 1124– 1149 Ma range. The late Proterozoic (555–994 Ma) population has a concentration of ages in the 620–675 Ma range.

Scattered mid‐Cretaceous, Triassic, Carboniferous, Paleoproterozoic, and Archean ages are also present. [44] Six samples yielded small populations of Oligocene and Miocene zircons, and at least five of these are of probable depositional age (see section 5). A seventh sample, MNDZ3‐755, yielded a single Cenozoic grain with an age of 40.6 ± 0.3 Ma, which is also probably close to the depositional age of the sample.

16 of 30

TC6013


TC6013

Figure 11. (top) Standard ternary diagrams illustrating modal framework grain compositions of sandstones from the Tres Cruces and Monte Nieva sections. Standard provenance fields of Dickinson [1983] labeled as follows: CB, Continental Block, RO, Recycled Orogen, MA, Magmatic Arc. (bottom) Vertical trends in total lithic grains (Lt) normalized to monocrystalline quartz (Qm) and feldspar (F). See Table 3 for definitions of parameters and Table 4 for recalculated data. [45] From a regional standpoint, detrital zircons from the sections in the southern part of the study area (sections MN, 1TT, 2TT, AL, and OB) are dominated by the ca. 500 Ma and ca. 1050 Ma age clusters, with lesser amounts of late Proterozoic ages. One exception is the Rio Seco Formation sample AL177, which closely resembles samples from farther north. Age spectra from sections in the northern study

area (1TC, SQ, and CV) are more complex, with many grains in the Cretaceous‐Devonian and middle to late Proterozoic ranges. [46] First‐cycle sources of detrital zircons with early Paleozoic ages are abundant in felsic rocks of the Famatina magmatic belt (sample CPLoire, Figure 12) and the Santa Rosa de Tastil granite [Pankhurst et al., 1998; Kirschbaum

17 of 30

TC6013


TC6013

from granitoid plutons and Proterozoic‐lower Cambrian (Puncoviscana) strata of the Puna and Cordillera Oriental.

8. Sediment Accumulation History

Figure 12. U‐Pb detrital zircon age probability plots for samples from the upper Proterozoic‐Cambrian Puncoviscana Formation, Cambrian Meson Group, Ordovician Complejo Eruptiva Oire, Ordovician Copalaya Formation at the base of section SP, Ordovician Santa Victoria Group, and the Cretaceous Lecho Formation. See Table S1 for data. et al., 2006; Hongn and Riller, 2007; Ramos, 2009; Hongn et al., 2010; Lucassen et al., 2011] of the Cordillera Oriental. Detrital zircons with ages in all three of the major populations recovered from the Cenozoic samples (∼500 Ma, 555–994 Ma, ∼1050 Ma) are present in Cambrian, Ordovician, and Cretaceous samples from the eastern Puna and Cordillera Oriental (Figure 12). The ∼1050 Ma grains were ultimately derived from orogenic rocks of the Sunsas belt which formed a prominent fringing terrane along the western margin of most of South America [e.g., Ramos, 2008, 2009, and references therein]. The relative abundance of middle Paleozoic and Cretaceous grains in the northern samples (e.g., samples CV1, CV2, and all of the 1TC samples) is consistent with derivation from Paleozoic‐ Mesozoic strata and Cretaceous igneous rocks [Coira, 1979; Gemmell et al., 1992] that are generally absent south of the El Toro lineament (Figures 2 and 13c). In contrast, samples from the southern sections (OB50, AL32, AL177, and all of the MN and 2TT samples) contain very few grains younger than 470 Ma (Figure 9), consistent with derivation

[47] The Monte Nieva section is the most complete and densely dated of those that we measured, rendering it useful for assessing the long‐term sediment accumulation history of the Cenozoic foreland basin system (Figure 14). Moreover, this section is representative of the regional stratigraphy throughout the study area. Chronostratigraphic control is provided by the youngest U‐Pb ages from detrital zircons [e.g., DeCelles et al., 2007], and apatite fission track pooled ages from granite clasts. We assume that (1) the depositional age of a detrital zircon sample cannot be older than its youngest grain, and (2) the AFT ages were not partially or entirely reset during burial and subsequent exhumation. The first assumption is legitimate because U‐Pb systematics of zircons are not reset at the burial temperatures experienced by these rocks. The second assumption is based on the fact that AFT ages of the analyzed clasts do not decrease down‐ section, as would be expected for progressively annealed samples with increasing depth. Moreover, these ages are within error of each other and similar to AFT ages obtained from bedrocks in nearby mountain ranges [Deeken et al., 2006]. Of the three AFT ages depicted on the diagram (Figure 14) only the age from sample MN1333 is likely to be a constraint on depositional age. The other two AFT ages are older than minimum ages from detrital zircons in their respective parts of the section, and they are older than the minimum AFT age from sample MN1333 despite coming from many hundreds of meters up‐section. All of the ages depicted in Figure 14 are from the Monte Nieva section except that from sample TC330 (Ducea, unpublished data, 2011), which was collected from a few meters below the top of the supersol zone in the Tres Cruces section (Figure 6). Although the Tres Cruces section is far to the north of the Monte Nieva section, the two sections are approximately along‐strike of each other, and it is plausible that the Tres Cruces sample provides a useful estimate of the age of the upper part of the supersol at Monte Nieva. [48] Geochronological control on the earliest portion of the accumulation history (Mealla and Maiz Gordo Formations) is lacking because samples did not yield Cenozoic zircon populations or apatite grains of sufficient quality for fission track analysis. However, robust paleontological ages from these units establish their late Paleocene‐middle Eocene ages [Quattrocchio and Volkheimer, 1990; Quattrocchio et al., 2000; del Papa et al., 2002; Marquillas et al., 2005]. The Rio González section of Reynolds et al. [2000] is also well‐ dated by paleomagnetic stratigraphy and isotopic ages on tuffs, and provides an opportunity to compare sediment accumulation histories in the frontal and interior parts of the orogenic belt during the late Miocene (Figure 14). [49] At Monte Nieva, the accumulation curve is upward convex and exhibits an overall acceleration of accumulation from a rate of ∼25 m/Myr from 60 to 47 Ma, to ∼200 m/Myr over the period 17–7 Ma (Figure 14). At the Rio González section, sediment accumulation rates range from 300 to 600 m/Myr, and the trend decreases slightly through time from ∼15 Ma to ∼9 Ma [Reynolds et al., 2000] (Figure 14). In the la Porcelana section of the northernmost Argentine

18 of 30

TC6013


Figure 13. (a–d) Block diagrams illustrating the temporal and spatial evolution of the Cenozoic foreland basin system in northwestern Argentina based on results of this study. Outlines of the Salta rift paleogeography shown in Figure 13a are from Salfity and Marquillas [1994]. See text for discussion. 19 of 30

TC6013

TC6013


TC6013

Figure 14. Thickness versus time plot for the Monte Nieva section (solid line), Rio González section [after Reynolds et al., 2000], and la Porcelana section [after Reynolds et al., 2001]. Arrows pointing toward the right signify that ages are maximum depositional ages; the actual depositional age could be younger. Subandes, Reynolds et al. [2001] documented accumulation rates of ∼580 m/Myr from ∼8 Ma to ∼2 Ma. The three sediment accumulation curves depicted in Figure 14 show a systematic eastward migration of the onset of accelerated accumulation: in the Monte Nieva section rapid accumulation commenced between ∼47 and 40 Ma; at Rio González, rapid accumulation began ∼15 Ma; and at la Porcelana accumulation accelerated at ∼8 Ma. Several other stratigraphic sections in the Subandes of northernmost Argentina exhibit rapid accumulation rates commencing at ∼8.5–9 Ma [Echavarria et al., 2003].

9. Regional Synthesis [50] Key features of the Cenozoic stratigraphic record in northwestern Argentina that must be explained by any holistic basin model include the regionally thin but widespread Paleogene clastic lithofacies of the Santa Bárbara Subgroup; the supersol zone and its eastward continuation as a major disconformity; the overlying eastward tapering, eastward onlapping wedge of upward‐coarsening Neogene fluvial and alluvial fan deposits, which are locally incorporated into syncontractional growth structures in the upper Miocene part of the section; the consistent westerly provenance of all of these deposits; and increasing sediment accumulation rates in Paleocene‐Pliocene strata, with progressive eastward migration of the temporal onset of rapid accumulation.

[51] We propose that the Cenozoic strata of northwestern Argentina were deposited in a regional foreland basin system that formed and migrated eastward under the influence of crustal shortening that originated in Chile and expanded eastward throughout the last ∼60 Myr. The modern foreland basin system east of the central Andes provides an analog for the ancient system preserved in outcrops in the modern orogenic belt. Chase et al. [2009] used geoid and flexural modeling to analyze gravity and topographic data for the modern foreland region, and showed that the modern basin contains a 300 km wide foredeep, a mostly buried 450 km wide forebulge, and a several hundred km wide diffuse back‐bulge depozone. Recent studies show that the foredeep is dominated by fluvial megafans [Horton and DeCelles, 1997, 2001] (or distributary fluvial systems [Hartley et al., 2010]); the forebulge is marked by shallow incision and a transition from well‐drained to poorly drained floodplain environments [McGlue and Cohen, 2006]; and the back‐bulge region is characterized by fluvial, lacustrine, and wetland environments that are strongly controlled by the seasonal hydrological cycle associated with the South American monsoon [Iriondo, 2004; Assine and Soares, 2004; McGlue and Cohen, 2006]. The modern back‐bulge and forebulge regions contain most of the morphological features that are represented in the Mealla, Maiz Gordo, and lower Lumbrera Formations, including lacustrine (perennial and ephemeral), wetland, and high‐sinuosity fluvial lithofacies, as well as intensely developed paleosols.

20 of 30

TC6013


TC6013

Figure 15. Regional time‐stratigraphic chart (or Wheeler diagram) for northwestern Argentina based on measured sections reported in this paper, plus the Angastaco (ANG) section from Carrapa et al. [2011a] and the Rio González section (GZ) of Reynolds et al. [2000]. [52] Notwithstanding local control by paleotopographic elements (Figure 13b), Paleocene‐lower Miocene strata of the Santa Bárbara and Métan Subgroups form an integrated, overall upward coarsening foreland basin succession that is readily interpreted in terms of foreland basin depozones [DeCelles and Giles, 1996]. An important feature of the succession is the supersol in the Maiz Gordo and Lumbrera Formations (Figure 7). This supersol zone represents at least 10–15 Myr of geological time during which the average sediment accumulation rate was near zero (Figure 14), and is therefore a zone of intense stratigraphic condensation. Eastward in the Subandean zone, the supersol zone is either absent or rests upon a regional erosional unconformity that truncates strata from Permian to Cretaceous age. Combined with the Bolivian supersol zone, its original palinspastic extent exceeded 200,000 km2. The great aerial extent of the supersol zone requires an extrinsic control on its origin (such as climate or geodynamics) rather than local processes that were intrinsic to the depositional system (e.g., normal fluvial overbank processes). The prolonged duration (ca. 10–15 Myr) and extreme variability in paleosol types in the supersol zone argue against a climatic control. Indeed, the paleosol types in the supersol zone require a complete range of hydrological and climatological conditions, from water‐ logged to arid (Figure 7). More importantly, paleoclimatic processes cannot explain the migration of the supersol zone/ basal Neogene disconformity up‐section toward the east (Figure 15). [53] Possible geodynamic mechanisms to explain the supersol/disconformity zone include regional isostatic rebound of the foreland basin during a period of tectonic inactivity, mantle‐driven (dynamic) uplift of the foreland

region, and forebulge migration. We reject isostatic rebound as a cause of the supersol zone because the contemporaneous thrust belt was located several hundred kilometers toward the west, beyond any reasonable wavelength of flexural isostatic rebound. Moreover, Paleocene‐Eocene foredeep deposits, temporally equivalent to the supersol zone, are well documented in northern Chile [Mpodozis et al., 2005; Arriagada et al., 2006] and demonstrate that the proximal part of the Paleogene foreland was actively subsiding in response to nearby crustal thickening during the development of the supersol zone. Dynamic uplift of the foreland region owing to processes driven by circulation in the mantle wedge beneath the western South American continental plate [e.g., Mitrovica et al., 1989; Gurnis, 1992] is a plausible mechanism for development of the supersol zone; it remains difficult, however, to reconcile dynamic uplift in the distal foreland with relatively rapid subsidence in the proximal region and eastward younging in the age of the supersol/disconformity. It is conceivable, however, that dynamic subsidence at a wavelength on the order of 103 km from the trench helped to create the minor accommodation above the forebulge required to preserve the supersol zone [e.g., Catuneanu, 2004; Allen and Allen, 2005]. [54] We interpret the supersol/disconformity zone as the result of migration of the flexural forebulge through this region (Figures 13 and 16) [DeCelles and Horton, 2003]. The up‐section migration and increased lacuna of the supersol/disconformity toward the east (Figure 15) [Starck and Vergani, 1996; Hernández et al., 1999] are predicted by geodynamic models of forebulge unconformity behavior [Sinclair et al., 1991; Coakley and Watts, 1991; Crampton and Allen, 1995; Schlunegger et al., 1997; Burkhard and

21 of 30

TC6013


Figure 16 22 of 30

TC6013

TC6013


Sommaruga, 1998]. Analogous supersols mark forebulge‐ related disconformities and condensation zones in the Himalayan, North American Cordilleran, North Alpine, and many other major foreland basins (see review by DeCelles [2011]). The Mealla and Maiz Gordo Formations locally contain pebbly conglomeratic sandstones that could have been partly derived from local granitic rocks associated with paleotopographic features associated with the margins of the Salta rift (Figure 13b). As the flexural forebulge swept through the region now occupied by the Cordillera Oriental, remnant paleotopographic features along the old rift shoulders were likely rejuvenated [e.g., Blisniuk et al., 1998]. However, the thinness of the Mealla and Maiz Gordo Formations, together with the localized distribution of the coarse‐grained material, suggests that these rejuvenated topographic elements were not large enough to generate a large amount of clastic sediment. Instead paleocurrent data from these units indicate east‐northeastward sediment transport from source terranes located to the west (Figures 5 and 6). We interpret the strata located beneath the supersol zone, including the Mealla and part of the Maiz Gordo Formation, as back‐bulge deposits that accumulated in the region east of the flexural forebulge, probably several hundred km to the east of the Paleocene‐early Eocene thrust belt. [55] Above the supersol/disconformity, the Quebrada de los Colorados Formation, uppermost part of the Lumbrera Formation, and the overlying lower Angastaco, Rio Seco, Anta, and Jésus María Formations represent a classic, eastward tapering foredeep wedge [Starck and Vergani, 1996; Hernández et al., 1999; Reynolds et al., 2000] (Figures 15). Evidence for local tectonic shortening in the Cordillera Oriental becomes abundant in the upper part of the Angastaco Formation and the Agujas Conglomerate in the form of angular and progressive unconformities [Carrera and Muñoz, 2008; Carrapa et al., 2011a, 2011b; this study]. The previously unbroken regional foreland basin was disrupted and actively exhumed between ∼14 Ma and 3 Ma as indicated by apatite (U‐Th)/He ages of Cretaceous and Cenozoic sedimentary rocks in the Angastaco area. These data are interpreted as the result of eastward migration of the strain front through this part of the Cordillera Oriental [Carrapa et al., 2011a, 2011b], and are consistent with previously reported evidence for the onset of syndepositional deformation in the Subandean zone during the latest Miocene and Pliocene [Reynolds et al., 2000; Echavarria et al., 2003; Uba et al., 2009; Hain et al., 2011]. We therefore interpret the Cenozoic succession in terms of foreland basin depozones that migrated eastward in response to the progressive growth of the Andean orogenic wedge (Figures 15 and 16). Although evidence for out‐of‐sequence deformation and local rapid exhumation during late Oligocene‐early Miocene time is documented in the Puna and Cordillera Oriental [Marrett et al., 1994; Coutand et al., 2001; Carrapa et al., 2005, 2011a; Deeken et al., 2006], the overall pattern is an eastward migration of the orogenic strain front and foreland flexural wave, albeit at an unsteady

TC6013

rate. The combined Bolivian‐northern Argentine foreland basin system was of regional scale, measuring at least 1000 km in a north‐south direction, and ∼300 km in an east‐west direction. Palinspastic restoration of shortening [Kley and Monaldi, 1998; McQuarrie, 2002a] in the Andean thrust belt would significantly increase the width of this paleo‐ foreland basin. [56] The documented sediment accumulation rates (Figure 14) are also typical of foreland basins worldwide [Xie and Heller, 2009; DeCelles, 2011]. The increasing rate of accumulation from Eocene time onward is interpreted to represent increasing subsidence rate, which results from the migration of the exponential flexural subsidence profile past a given location in the foreland. Early accumulation is slow because the flexural forebulge reduces or eliminates sediment accommodation. In this case, the part of the section marked by forebulge accumulation is strongly overprinted by pedogenic processes. Farther east, the passage of the forebulge is marked in the Rio González section by an erosional disconformity at the base of the foredeep deposits [Reynolds et al., 2000]. The erosional disconformity at the base of the foredeep strata is present at all of the sections documented in the Santa Bárbara Ranges and in the easternmost Cordillera Oriental (e.g., Alemanía and Campamento Vespucio sections [see also Starck and Vergani, 1996; Hernández et al., 1999; Reynolds et al., 2000, 2001]). The eastward migration of the onset of rapid subsidence as illustrated in Figure 14 is consistent with progressive migration of the flexural subsidence wave. [57] Isolated outcrops of Eocene strata (including the Geste Formation) that contain evidence for local syndepositional tectonic shortening are present in the eastern Puna and western Cordillera Oriental (Figure 15) [Carrapa et al., 2005; Carrapa and DeCelles, 2008]. These strata are generally correlative with the Quebrada de los Colorados Formation, but they are not spatially contiguous with the continuous Paleocene‐Pliocene succession of the regional foreland basin system. The Geste Formation rests depositionally upon Paleozoic basement rocks, was locally derived from orogenic highlands, and yields detrital thermochronological evidence of rapid exhumation, probably related to local shortening [Carrapa et al., 2008]. We tentatively interpret the Geste Formation as a remnant of the wedge‐top portion of the early foreland basin system. It is also plausible that Geste Formation depocenters were never integrated into the regional foreland basin system, and instead represent local intermontane basins [e.g., Horton, 2005]. Hongn et al. [2007] and Bosio et al. [2009] also reported evidence for late Eocene structural growth in the Eastern Cordillera.

10.

Discussion

[58] Three questions may be addressed using the results presented herein: (1) What information does the foreland basin succession provide about the distance and tempo of flexural wave migration—what was the rate of migration of

Figure 16. (a–e) Schematic paleogeographic maps for the study area showing the evolution of the orogenic belt and adjacent foreland basin system from latest Cretaceous to present. Paleogeography depicted in Figure 16a is based on Salfity and Marquillas [1994] and Mpodozis et al. [2005]. Barbed lines represent approximate locations of the deformation front at times (in Ma) corresponding to associated numbers. 23 of 30

TC6013


TC6013

Figure 17. Diagram illustrating the migration of the foreland basin flexural wave through the region now occupied by the Andean orogenic belt in northwestern Argentina, in a fixed flexural wave frame of reference. Four flexural profiles based on flexural rigidities listed at the top are shown to illustrate the plausible distribution of foreland depozones. The flexural profiles were generated by flexing an unbroken plate under a rectangular load with a half‐width of 300 km and a topographic elevation of 2.5 km. These profiles are anchored to the modern front of the Andes, and held fixed while the thrust belt is palinspastically restored at 10 Myr intervals back to 60 Ma. Each thrust belt panel is labeled with the progressively restored locations and sizes of the Cordillera de Domeyko (CD), Cordillera Occidental (COC), Puna and Cordillera Oriental (COR), and the Santa Bárbara ranges (SB), based on total shortening estimates from Kley and Monaldi [1998] and kinematic timing as discussed in this paper and by Carrapa et al. [2011a, 2011b]. Shaded zone represents average location of the forebulge zone for the two intermediate flexural rigidities, which are based on Tassara et al. [2007] and Chase et al. [2009]. Curves for flexural rigidity of 1023 Nm and 1024 Nm are also shown. Placement of each thrust belt panel is dictated by our interpretation of foreland basin depozones. On the right is a Wheeler diagram showing stratigraphic information used to reconstruct the foreland depozones. For example, between 60 and 50 Ma, the Puna and Cordillera Oriental were occupied by the flexural forebulge (based on the presence of the supersol zone), and by ∼40 Ma these regions were within the foredeep depozone. Velocities of migration of the system at different times are labeled along double‐arrowed lines. the foreland basin through time? (2) To what extent did the vestigial Salta rift interact with the distal flexural signal of the Andean orogenic wedge? and (3) What are the implications of this work for pure‐shear versus simple‐shear deformation in the central Andes? [59] 1. Figure 17 illustrates a spatial‐temporal reconstruction of the foreland basin flexural wave, calibrated to available shortening estimates in the Andean thrust belt and the stratigraphic record documented in this paper. Shortening estimates are from Kley and Monaldi [1998, and references therein], and the horizontal locations through time of palinspastic panels are fixed according to the interpreted foreland depozone in a given region, plotted against modeled flexural subsidence profiles based on flexural rigidities for central South American foreland lithosphere [Tassara et al., 2007; Chase et al., 2009]. Sequential palinspastic restoration of the thrust belt panel is based on available kinematic timing constraints [Kraemer et al., 1999; Reynolds et al., 2000; Echavarria et al., 2003; Coutand et al., 2006; Carrapa et al., 2005, 2008, 2011a, 2011b; Mazzuoli et al., 2008; this study]. For example, based on our interpretation

of the supersol zone, the eastern Puna and Cordillera Oriental were located within the region occupied by the forebulge during late Paleocene through middle Eocene time. The palinspastic thrust belt panel is therefore adjusted horizontally to a location consistent with these regions straddling the forebulge (in a fixed flexural wave frame of reference). Thus, through time, we can estimate the distance and rate of flexural wave migration [e.g., DeCelles and Horton, 2003]. [60] Important caveats in such an analysis include the possibility that flexural rigidity and orogenic loading changed through time, and that the amounts of shortening used in the reconstruction are extremely inaccurate. Changes in flexural rigidity are inscrutable from the viewpoint of the stratigraphic record, but we might safely assume that rigidity would have increased through time as older, stronger South American lithosphere became involved in the flexural profile. The effect of this would be that through time the wavelengths of foreland depozones would have increased. The values of flexural rigidity used in Figure 17 are based on present estimates for the central Andean foreland [Tassara et al., 2007; Chase et al., 2009], and are probably

24 of 30

TC6013


maxima. The issue of tectonic shortening is especially problematic, inasmuch as the total shortening documented in this region (no more than ca. 110 km [Allmendinger, 1986; Grier et al., 1991; Cladouhos et al., 1994; Monaldi and Kley, 1997; Kley and Monaldi, 1998]) is far below that required to explain the great thickness of the crust [Kley and Monaldi, 1998]. However, increased shortening would only increase the distance of flexural wave migration interpreted in this manner, so that our reconstruction is somewhat conservative. [61] The reconstruction (Figure 17) suggests that the flexural wave in northwestern Argentina migrated unsteadily eastward over a total distance of approximately 600 km, with two episodes of rapid migration (ca. 23–48 mm/yr) separated by periods of relatively slow migration (ca. 4– 5 mm/yr). The periods of rapid flexural wave migration occurred at ∼50–40 Ma and ∼5–0 Ma, and are consistent with major eastward jumps in the position of the contractional strain front as indicated by thermochronologically determined exhumation events [Deeken et al., 2006; Coutand et al., 2006; Carrapa et al., 2008, 2011b], growth structures and angular unconformities in reliably dated stratigraphic units [Hongn et al., 2007; Mazzuoli et al., 2008; Carrapa and DeCelles, 2008; Carrera and Muñoz, 2008; Carrapa et al., 2011a; Hain et al., 2011; this study], and patterns of provenance and paleotopographic reconstructions [Carrapa et al., 2011a; Hain et al., 2011; this study]. The first of these kinematic jumps took place as deformation propagated from the Cordillera de Domeyko in northern Chile during late Paleocene‐early Eocene time into the western part of the Cordillera Oriental by about 40 Ma [Maksaev and Zentilli, 1999; Deeken et al., 2006; DeCelles et al., 2007; Carrapa and DeCelles, 2008]. The second major jump is expressed as rapid eastward propagation of strain from the eastern Cordillera Oriental into the frontal ranges no earlier than ∼5 Ma [Reynolds et al., 2000; Echavarria et al., 2003; Carrera and Muñoz, 2008; Carrapa et al., 2011a; Hain et al., 2011]. Major deformation in the Santa Bárbara Ranges commenced after 2 Ma [Reynolds et al., 2000; Hain et al., 2011]. It is important to note that no evidence exists in either stratigraphic or structural records for complete cessation of shortening or long‐term stasis of the strain front in the time interval between these periods of rapid eastward strain propagation. Instead, deformation occurred more or less continuously in the Andean orogenic wedge from Eocene time onward [Coutand et al., 2006; Deeken et al., 2006; Carrera et al., 2006; Carrapa et al., 2009, 2011a, 2011b]. [62] Several workers [Strecker et al., 2007; Hain et al., 2011] have argued that a fundamental change has occurred in the eastern retroarc region of northwestern Argentina since late Miocene time, during the transition from an unbroken foreland to a topographically complex frontal thrust belt dominated by the structural inversion of fault systems that originated during Cretaceous extension (the Salta rift). In this view, upper Miocene‐Quaternary synorogenic deposits in the eastern Cordillera Oriental and Santa Bárbara system were strongly influenced by local structural uplift and formed in an intramontane tectonic setting isolated from the regional foreland basin system. However, we note that the modern Andean foreland basin system at these latitudes is typical of regional‐scale unbroken forelands, despite the presence of the Lomas de Olmedo arm of the Salta rift in the

TC6013

subsurface beneath the modern foredeep depozone [Cominguez and Ramos, 1995]. In our view, the late Miocene‐Quaternary deformation in the easternmost Cordillera Oriental and frontal Santa Bárbara ranges highlighted by Strecker et al. [2007] and Hain et al. [2011] is reconcilable with an eastward propagating orogenic wedge, whether or not the deformation has taken place along relatively shallow (200,000 km2) accumulation of extremely mature paleosols in the Eocene record, which we interpret to be the result of stratigraphic condensation during passage of the flexural forebulge through the region. [66] 2. The Cenozoic foreland basin record of northern Argentina is remarkably similar to that in Bolivia. Palinspastic restoration of minimum shortening estimates and paleogeographic restoration of the foreland basin system using modeled flexural profiles demonstrates that the system migrated at least 600 km eastward. This migration took place in two large jumps at rates of ∼20–50 mm/yr (50–40 Ma, and 5–0 Ma), separated by periods during which migration was much slower, only ∼5 mm/yr. [67] 3. The presence of a regional foreland basin system in northwestern Argentina and Bolivia throughout the Cenozoic effectively nullifies the notion that these two regions behaved significantly different in terms of simple versus pure shear; in particular, there is no support in the foreland basin history for dominantly pure shear thickening in the Puna and Cordillera Oriental of northwestern Argentina north of 26°S. Nor is there a rationale for viewing the late Miocene‐Pliocene kinematic history of the northwestern Argentina portion of the orogenic belt and foreland basin system as fundamentally different from its northerly counterpart in Bolivia, aside from the style of deformation. Instead, the Paleocene through Pliocene stratigraphic record of northwestern Argentina and Bolivia documents the development and expansion of the Andean orogenic belt and its associated foreland basin flexural wave. [68] Acknowledgments. We are grateful to D. Starck, R. M. Hernandez, J. H. Reynolds, L. Schoenbohm, and R. N. Alonso for numerous insightful discussions and assistance in locating stratigraphic sections in the field. T. P. Ojha helped with graphics, and D. Gingrich kindly provided details for the measured section at Cianzo. J. McNabb assisted with sample preparation for U‐Pb zircon analyses. The U.S. National Science Foundation– Tectonics and ExxonMobil provided funding. This paper represents part of an ongoing collaboration with ExxonMobil scientists on convergent orogenic systems. Careful reviews by F. Schlunegger, an anonymous reviewer, and Associate Editor Andrew Carter helped us to substantially improve the manuscript.

References Aalto, R., L. Maurice‐Bourgoin, T. Dunne, D. R. Montgomery, C. A. Nittrouer, and J. L. Guyot (2003), Episodic sediment accumulation on Amazonian flood plains influenced by El Nino/Southern Oscillation, Nature, 425, 493–497, doi:10.1038/nature02002. Allen, P. A., and J. R. Allen (2005), Basin Analysis, 2nd ed., 549 pp., Blackwell, Malden, Mass. Allmendinger, R. W. (1986), Tectonic development, southeastern border of the Puna Plateau, northwestern Argentine Andes, Geol. Soc. Am. Bull., 97, 1070–1082, doi:10.1130/0016-7606(1986)972.0. CO;2. Allmendinger, R. W., and T. Gubbels (1996), Pure and simple shear plateau uplift, Altiplano ‐ Puna, Argentina and Bolivia, Tectonophysics, 259, 1–13, doi:10.1016/0040-1951(96)00024-8. Allmendinger, R. W., V. A. Ramos, T. E. Jordan, M. Palma, and B. L. Isacks (1983), Paleogeography and Andean structural geometry, northwest Argentina, Tectonics, 2, 1–16, doi:10.1029/TC002i001p00001.

26 of 30

TC6013


Allmendinger, R., T. Jordan, S. Kay, and B. Isacks (1997), The evolution of the Altiplano‐Puna Plateau of the Central Andes, Annu. Rev. Earth Planet. Sci., 25, 139–174, doi:10.1146/annurev.earth.25.1.139. Alonso, R. N. (1992), Estratigrafía del Cenozoico de la cuenca Pastos Grandes (Puna Salteña) con énfasis en la Formación Sijes y sus boratos, Asoc. Geol. Argent. Rev., 47, 189–199. Arriagada, C., P. R. Cobbold, and P. Roperch (2006), Salar de Atacama basin: A record of compressional tectonics in the central Andes since the mid‐Cretaceous, Tectonics, 25, TC1008, doi:10.1029/ 2004TC001770. Assine, M. L., and P. C. Soares (2004), Quaternary of the Pantanal, west‐ central Brazil, Quat. Int., 114, 23–34, doi:10.1016/S1040-6182(03) 00039-9. Baby, P., P. Rochat, G. Mascle, and G. Hérail (1997), Neogene shortening contribution to crustal thickening in the back arc of the Central Andes, Geology, 25, 883–886, doi:10.1130/0091-7613(1997)0252.3.CO;2. Barnes, J. B., T. A. Ehlers, N. McQuarrie, P. B. O’Sullivan, and S. Tawackoli (2008), Thermochronometer record of central Andean Plateau growth, Bolivia (19.5°S), Tectonics, 27, TC3003, doi:10.1029/2007TC002174. Beck, S., and G. Zandt (2002), The nature of orogenic crust in the Central Andes, J. Geophys. Res., 107(B10), 2230, doi:10.1029/2000JB000124. Blisniuk, P. M., L. J. Sonder, and R. J. Lillie (1998), Foreland normal fault control on NW Himalayan thrust front development, Tectonics, 17, 766–779, doi:10.1029/98TC01870. Bond, G. C., and M. A. Kominz (1984), Construction of tectonic subsidence curves for the early Paleozoic miogeocline, southern Canadian Rocky Mountains: Implications for subsidence mechanisms, age of breakup, and crustal thinning, Geol. Soc. Am. Bull., 95, 155–173, doi:10.1130/0016-7606(1984)952.0.CO;2. Bosio, P. P., J. Powell, C. del Papa, and F. Hongn (2009), Middle Eocene deformation– sedimentation in the Luracatao Valley: Tracking the beginning of the foreland basin of northwestern Argentina, J. South Am. Earth Sci., 28, 142–154, doi:10.1016/j.jsames.2009.06.002. Bradley, D. C., and W. S. F. Kidd (1991), Flexural extension of the upper continental crust in collisional foredeeps, Geol. Soc. Am. Bull., 103, 1416–1438, doi:10.1130/0016-7606(1991)1032.3. CO;2. Burkhard, M., and A. Sommaruga (1998), Evolution of the western Swiss Molasse basin: Structural relations with the Alps and the Jura belt, in Cenozoic Foreland Basins of Western Europe, edited by A. Mascle et al., Geol. Soc. Spec. Publ., 134, 279–298. Bywater‐Reyes, S., B. Carrapa, M. Clementz, and L. Schoenbohm (2010), The effect of late Cenozoic aridification on sedimentation in the Eastern Cordillera of NW Argentina (Angastaco basin), Geology, 38(3), 235–238, doi:10.1130/G30532.1. Carrapa, B., and P. G. DeCelles (2008), Eocene exhumation and basin development in the Puna of Northwestern Argentina, Tectonics, 27, TC1015, doi:10.1029/2007TC002127. Carrapa, B., D. Adelmann, G. E. Hilley, E. Mortimer, E. R. Sobel, and M. R. Strecker (2005), Oligocene uplift and development of plateau morphology in the southern central Andes, Tectonics, 24, TC4011, doi:10.1029/2004TC001762. Carrapa, B., M. R. Strecker, and E. R. Sobel (2006), Cenozoic orogenic growth in the central Andes: Evidence from sedimentary rock provenance and apatite fission track thermochronology in the Fiambalá basin, southernmost Puna plateau margin (NW Argentina), Earth Planet. Sci. Lett., 247, 82–100, doi:10.1016/j.epsl.2006.04.010. Carrapa, B., B. Hauer, L. Schoenbohm, M. R. Strecker, A. K. Schmitt, A. Villanueva, and A. S. Gomez (2008), Dynamics of deformation and sedimentation in the northern Sierras Pampeanas: An integrated study of the Neogene Fiambalá basin, NW Argentina, Geol. Soc. Am. Bull., 120, 1518–1543, doi:10.1130/B26111.1. Carrapa, B., P. G. DeCelles, P. Reiners, and G. E. Gehrels (2009), Apatite triple dating and white mica 40Ar/39Ar thermochronology of syn‐tectonic detritus in the Central Andes: A multi‐phase tectono‐thermal history, Geology, 37, 407–410, doi:10.1130/G25698A.1. Carrapa, B., S. Bywater‐Reyes, P. G. DeCelles, E. Mortimer, and G. E. Gehrels (2011a), Eocene‐Miocene synorogenic basin evolution in the Eastern Cordillera of northwestern Argentina (25°–26°S): Regional implications for Andean orogenic wedge development, Basin Res., 23, 1–20. Carrapa, B., J. D. Trimble, and D. F. Stockli (2011b), Patterns and timing of exhumation and deformation in the Eastern Cordillera of NW Argentina revealed by (U‐Th)/He thermochronology, Tectonics, 30, TC3003, doi:10.1029/2010TC002707. Carrera, N., and J. A. Muñoz (2008), Thrusting evolution in the southern Cordillera Oriental (northern Argentine Andes): Constraints from growth strata, Tectonophysics, 459, 107–122, doi:10.1016/j.tecto.2007.11.068.

TC6013

Carrera, N., J. A. Muñoz, F. Sàbat, R. Mon, and E. Roca (2006), The role of inversion tectonics in the structure of the Cordillera Oriental (NW Argentinean Andes), J. Struct. Geol., 28, 1921–1932, doi:10.1016/j. jsg.2006.07.006. Catuneanu, O. (2004), Retroarc foreland systems—Evolution through time, J. Afr. Earth Sci., 38(3), 225–242, doi:10.1016/j.jafrearsci.2004.01.004. Chase, C. G., A. J. Sussman, and D. D. Coblentz (2009), Curved Andes: Geoid, forebulge, and flexure, Lithosphere, 1(6), 358–363, doi:10.1130/L67.1. Cladouhos, T. T., R. W. Allmendinger, B. Coira, and E. Farrar (1994), Late Cenozoic deformation in the central Andes; fault kinematics from the northern Puna, northwestern Argentina and southwestern Bolivia, J. South Am. Earth Sci., 7, 209–228, doi:10.1016/0895-9811(94)90008-6. Coakley, B. J., and A. B. Watts (1991), Tectonic controls on the development of unconformities, North Slope Alaska, Tectonics, 10, 101–130, doi:10.1029/90TC01982. Coira, B. L. (1979), Descripción geológica de la Hoja 3c Abra Pampa, provincia de Jujuy, Bol. Argent. Serv. Geol. Nac., 170, 1–85. Cominguez, A. H., and V. A. Ramos (1995), Geometry and seismic expression of the Cretaceous Salta rift of northwestern Argentina, in Petroleum Basins of South America, edited by A. J. Tankard, R. Suarez, and H. J. Welsink, AAPG Mem., 62, 325–339. Coney, P. J., and C. A. Evenchick (1994), Consolidation of the American Cordilleras, J. South Am. Earth Sci., 7, 241–262, doi:10.1016/08959811(94)90011-6. Coutand, I., P. R. Cobbold, M. de Urreiztieta, P. Gautier, A. Chauvin, D. Gapais, E. A. Rossello, and O. Lòpez‐Gamundí (2001), Style and history of Andean deformation, Puna plateau, Northwestern Argentina, Tectonics, 20, 210–234, doi:10.1029/2000TC900031. Coutand, I., B. Carrapa, A. Deeken, A. K. Schmitt, E. Sobel, and M. R. Strecker (2006), Orogenic plateau formation and lateral growth of compressional basins and ranges: Insights from sandstone petrography and detrital apatite fission‐track thermochronology in the Angastaco Basin, NW Argentina, Basin Res., 18, 1–26, doi:10.1111/j.1365-2117. 2006.00283.x. Crampton, S. L., and P. A. Allen (1995), Recognition of flexural forebulge unconformities in the geologic record, AAPG Bull., 79, 1495–1514. Cristallini, E., A. H. Cominguez, and V. A. Ramos (1997), Deep structure of the Metán‐Guachipas region; tectonic inversion in northwestern Argentina, J. South Am. Earth Sci., 10(5–6), 403–421, doi:10.1016/ S0895-9811(97)00026-6. Crouvi, O., R. Amit, Y. Enzel, and A. R. Gillespie (2010), Active sand seas and the formation of desert loess, Quat. Sci. Rev., 29, 2087–2098, doi:10.1016/j.quascirev.2010.04.026. Dávila, F. M., and R. A. Astini (2007), Cenozoic provenance history of synorogenic conglomerates in western Argentina (Famatina Belt): Implications for Central Andean foreland development, Geol. Soc. Am. Bull., 119, 609–622, doi:10.1130/B26007.1. DeCelles, P. G. (2011), Foreland basin systems revisited: Variations in response to tectonic settings, in Tectonics of Sedimentary Basins: Recent Advances, edited by C. Busby and A. Azor, Blackwell, Cambridge, Mass., in press. DeCelles, P. G., and K. N. Giles (1996), Foreland basin systems, Basin Res., 8, 105–123, doi:10.1046/j.1365-2117.1996.01491.x. DeCelles, P. G., and B. K. Horton (2003), Early to middle Tertiary foreland basin development and the history of Andean crustal shortening in Bolivia, Geol. Soc. Am. Bull., 115(1), 58–77, doi:10.1130/00167606(2003)1152.0.CO;2. DeCelles, P. G., R. P. Langford, and R. K. Schwartz (1983), Two new methods of paleocurrent determination from trough cross‐stratification, J. Sediment. Res., 53, 629–642. DeCelles, P. G., M. B. Gray, R. B. Cole, N. Pequera, D. A. Pivnik, K. D. Ridgway, and P. Srivastava (1991), Controls on synorogenic alluvial‐fan architecture, Beartooth Conglomerate, Wyoming and Montana, Sedimentology, 38, 567–590, doi:10.1111/j.1365-3091.1991. tb01009.x. DeCelles, P. G., B. Carrapa, and G. E. Gehrels (2007), Detrital zircon U‐Pb ages provide provenance and chronostratigraphic information from Eocene synorogenic deposits in northwestern Argentina, Geology, 35(4), 323–326, doi:10.1130/G23322A.1. DeCelles, P. G., M. N. Ducea, G. Zandt, and P. Kapp (2009), Cyclicity in Cordilleran orogenic systems, Nat. Geosci., 2, 251–257, doi:10.1038/ ngeo469. Deeken, A., E. R. Sobel, I. Coutand, M. Haschke, U. Riller, and M. R. Strecker (2006), Development of the southern Eastern Cordillera, NW Argentina, constrained by apatite fission track thermochronology: From early Cretaceous extension to middle Miocene shortening, Tectonics, 25, TC6003, doi:10.1029/2005TC001894.

27 of 30

TC6013


del Papa, C. (2006), Estratigrafía y Paleoambientes de la Formación Lumbrera, Grupo Salta, Noroeste Argentino, Rev. Asoc. Geol. Argent., 61(3), 15–29. del Papa, C. E., and J. Salfity (1999), Non‐marine Paleogene sequences, northwest Argentina, Acta Geol. Hisp., 34(2–3), 105–122. del Papa, C. E., V. García, and M. Quattrocchio (2002), Sedimentary facies and palynofacies assemblages in an Eocene perennial lake, Lumbrera Formation, northwest Argentina, J. South Am. Earth Sci., 15, 553–569, doi:10.1016/S0895-9811(02)00081-0. del Papa, C., A. Kirschbaum, J. Powell, A. Brod, F. Hongn, and M. Pimentel (2010), Sedimentological, geochemical and paleontological insights applied to continental omission surfaces: A new approach for reconstructing Eocene foreland basin in NW Argentina, J. South Am. Earth Sci., 29, 327–345, doi:10.1016/j.jsames.2009.06.004. Dickinson, W. R. (1983), Provenance of North American Phanerozoic sandstones in relation to tectonic setting, Geol. Soc. Am. Bull., 94, 222–257, doi:10.1130/0016-7606(1983)942.0.CO;2. Dumitru, T. A. (1993), A new computer automated microscope stage system for fission track analysis, Nucl. Tracks, 21, 575–580. Dunn, J. F., K. G. Hartshorn, and P. W. Hartshorn (1995), Structural styles and hydrocarbon potential of the Sub‐Andean thrust belt of southern Bolivia, in Petroleum Basins of South America, edited by A. J. Tankard, R. Suárez, and H. J. Welsink, AAPG Mem., 62, 523–543. Echavarria, R., R. Hernández, R. W. Allmendinger, and J. H. Reynolds (2003), Subandean thrust and fold belt of northwest Argentina: Geometry and timing of the Andean evolution, AAPG Bull., 87(6), 965–985, doi:10.1306/01200300196. Ege, H., E. R. Sobel, E. Scheuber, and V. Jacobshagen (2007), Exhumation history of the Central Andean plateau (southern Bolivia) constrained by apatite fission track thermochronology, Tectonics, 26, TC1004, doi:10.1029/2005TC001869. Ellis, S. M., P. Fullsack, and C. Beaumont (1995), Oblique convergence of the crust driven by basal forcing: Implications for length‐scales of deformation and strain partitioning in orogens, Geophys. J. Int., 120(1), 24–44, doi:10.1111/j.1365-246X.1995.tb05909.x. England, P., and G. Houseman (1985), Role of lithospheric strength heterogeneities in the Tectonics of Tibet and neighbouring regions, Nature, 315, 297–301, doi:10.1038/315297a0. Galbraith, R. F., and P. F. Green (1990), Estimating the component ages in a finite mixture, Nucl. Tracks, 17, 196–206. Galli, C. I., R. M. Hernández, and J. H. Reynolds (1996), Análisis paleoambiental y ubicación geocronológica del Subgrupo Metán (Grupo Orán, Neógeno) en el Río Piedras, departmento Metán: Salta, Argentina, Bol. Inf. Petrol., 12, 99–107. Galliski, M. A., and J. G. Viramonte (1988), The Cretaceous paleorift in northwestern Argentina; a petrologic approach, J. South Am. Earth Sci., 1, 329–342, doi:10.1016/0895-9811(88)90021-1. Gebhard, J. A., A. R. Guidice, and J. O. Gascon (1974), Geología de la comarca entre el Río Juramento y Arroyo las Tortugas, provincias de Salta y Jujuy, Asoc. Geol. Argent. Rev., 29, 359–375. Gehrels, G. E., V. Valencia, and J. Ruiz (2008), Enhanced precision, accuracy, efficiency, and spatial resolution of U‐Pb ages by laser ablation–multicollector–inductively coupled plasma–mass spectrometry, Geochem. Geophys. Geosyst., 9, Q03017, doi:10.1029/2007GC001805. Gemmell, J. B., H. Zantop, and L. D. Meinert (1992), Genesis of the Aguilar zinc‐lead‐silver deposit, Argentina: Contact metasomatic vs. sedimentary exhalative, Econ. Geol., 87, 2085–2112, doi:10.2113/ gsecongeo.87.8.2085. Gillis, R. J., B. K. Horton, and M. Grove (2006), Thermochronology, geochronology, and upper crustal structure of the Cordillera Real: Implications for Cenozoic exhumation of the central Andean plateau, Tectonics, 25, TC6007, doi:10.1029/2005TC001887. Green, P. F. (1981), A new look at statistics in fission‐track dating, Nucl. Tracks, 5, 77–86, doi:10.1016/0191-278X(81)90029-9. Grier, M. E., J. A. Salfity, and R. W. Allmendinger (1991), Andean reactivation of the Cretaceous Salta rift, northwestern Argentina, J. South Am. Earth Sci., 4(4), 351–372, doi:10.1016/0895-9811(91)90007-8. Gurnis, M. (1992), Rapid continental subsidence following the initiation and evolution of subduction, Science, 255, 1556–1558, doi:10.1126/science. 255.5051.1556. Hain, M. P., M. R. Strecker, B. Bookhagen, R. N. Alonso, H. Pingel, and A. K. Schmitt (2011), Neogene to Quaternary broken foreland formation and sedimentation dynamics in the Andes of NW Argentina (25°S), Tectonics, 30, TC2006, doi:10.1029/2010TC002703. Hartley, A. J., G. S. Weissmann, G. J. Nichols, and G. L. Warwick (2010), Large distributive fluvial systems: Characteristics, distribution and controls on development, J. Sediment. Res., 80, 167–183, doi:10.2110/ jsr.2010.016.

TC6013

Haschke, M., A. Günther, D. Melnick, H. Echtler, K.‐J. Reutter, E. Scheuber, and O. Onken (2006), Central and southern Andean tectonic evolution inferred from arc magmatism, in The Andes: Active Subduction Orogeny, edited by O. Oncken et al., pp. 337–353, Springer, Berlin. Hasiotis, S. T., and T. M. Bown (1992), Invertebrate trace fossils: The backbone of continental ichnology, in Trace Fossils, Short Courses Paleontol., vol. 5, edited by C.G. Maples and R. R. West, pp. 64–104, Paleontol. Soc., Knoxville, Tenn. Heit, B., I. Koulakov, G. Asch, X. Yuan, R. Kind, I. Alcozer, S. Tawackoli, and H. Wilke (2008), More constraints to determine the seismic structure beneath the Central Andes at 21°S using teleseismic tomography analysis, J. South Am. Earth Sci., 25, 22–36. Hernández, R. M., and L. Echavarria (2009), Faja plegada y corrida Subandina del noroeste Argentina: Estratigrafía, geometría y cronología de la deformación, Asoc. Geol. Argent. Rev., 65, 68–80. Hernández, R. M., J. Reynolds, and A. Di Salvo (1996), Análisis tectosedimentario y ubicación geocronológica del Grupo Orán en el río Iruya, Bol. Inf. Petrol., 45, 80–93. Hernández, R. M., C. I. Galli, and J. Reynolds (1999), Estratigrafía del Terciario en el noroeste Argentino, in Geologia del Noroeste Argentino: Relatorio del XV Congreso Geológico Argentino, edited by G. González Bonorino, R. Omarini, and J. Viramonte, pp. 316–328, Univ. Nac. de Salta, Salta, Argentina. Hindle, D., J. Kley, O. Oncken, and S. V. Sobolev (2005), Crustal balance and crustal flux from shortening estimates in the Central Andes, Earth Planet. Sci. Lett., 230, 113–124, doi:10.1016/j.epsl.2004.11.004. Hongn, F. D., and U. Riller (2007), Tectonic evolution of the western margin of Gondwana inferred from syntectonic emplacement of Paleozoic granitoid plutons in northwest Argentina, J. Geol., 115(2), 163–180, doi:10.1086/510644. Hongn, F. D., C. del Papa, J. Powell, I. Petrinovic, R. Mon, and V. Deraco (2007), Middle Eocene deformation and sedimentation in the Puna‐ Eastern Cordillera transition (23°–26°S): Control by preexisting heterogeneities on the pattern of initial Andean shortening, Geology, 35, 271–274, doi:10.1130/G23189A.1. Hongn, F. D., J. M. Tubía, A. Aranguren, N. Vegas, R. Mon, and G. R. Dunning (2010), Magmatism coeval with lower Paleozoic shelf basins in NW‐Argentina (Tastil batholith): Constraints on current stratigraphic and tectonic interpretations, J. South Am. Earth Sci., 29, 289–305, doi:10.1016/j.jsames.2009.07.008. Horton, B. K. (2005), Revised deformation history of the central Andes: Inferences from Cenozoic foredeep and intermontane basins of the Eastern Cordillera, Bolivia, Tectonics, 24, TC3011, doi:10.1029/2003TC001619. Horton, B. K., and P. G. DeCelles (1997), The modern foreland basin system adjacent to the Central Andes, Geology, 25, 895–898, doi:10.1130/0091-7613(1997)0252.3.CO;2. Horton, B. K., and P. G. DeCelles (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 Res., 13, 43–63, doi:10.1046/j.1365-2117.2001.00137.x. Horton, B. K., B. A. Hampton, and G. Waanders (2001), Paleogene synorogenic sedimentation in the Altiplano plateau and implications for initial mountain building in the central Andes, Geol. Soc. Am. Bull., 113, 1387–1400, doi:10.1130/0016-7606(2001)1132.0.CO;2. Houseman, G., and P. England (1993), Crustal thickening versus lateral expulsion in the Indian‐Asian continental collision, J. Geophys. Res., 98, 12,233–12,249, doi:10.1029/93JB00443. Hurford, A. J., and P. F. Green (1983), The zeta age calibration of fission‐ track dating, Chem. Geol., 41, 285–317, doi:10.1016/S0009-2541(83) 80026-6. Ingersoll, R. V., T. F. Bullard, R. L. Ford, J. P. Grimm, J. D. Pickle, and S. W. Sares (1984), The effect of grain size on detrital modes: A test of the Gazzi‐Dickinson point‐counting method, J. Sediment. Petrol., 54, 103–116. Iriondo, M. (2004), Large wetlands of South America: A model for Quaternary humid environments, Quat. Int., 114, 3–9, doi:10.1016/ S1040-6182(03)00037-5. Isacks, B. (1988), Uplift of the central Andean plateau and bending of the Bolivian orocline, J. Geophys. Res., 93, 3211–3231, doi:10.1029/ JB093iB04p03211. James, D. E. (1971), Plate‐tectonic model for the evolution of the central Andes, Geol. Soc. Am. Bull., 82, 3325–3346, doi:10.1130/00167606(1971)82[3325:PTMFTE]2.0.CO;2. Jordan, T. E., and R. W. Allmendinger (1986), The Sierras Pampeanas of Argentina: A modern analogue of Rocky Mountain foreland deformation, Am. J. Sci., 286, 737–764, doi:10.2475/ajs.286.10.737. Jordan, T. E., and R. N. Alonso (1987), Cenozoic Stratigraphy and Basin Tectonics of the Andes Mountains, 20°–28° South Latitude, AAPG Bull., 71, 49–64.

28 of 30

TC6013


Jordan, T. E., B. L. Isacks, R. W. Allmendinger, J. A. Brewer, V. A. Ramos, and C. J. Ando (1983), Andean tectonics related to geometry of subducted Nazca plate, Geol. Soc. Am. Bull., 94, 341–361, doi:10.1130/ 0016-7606(1983)942.0.CO;2. Kay, S., and B. Coira (2009), Shallowing and steepening subduction zones, continental lithospheric loss, magmatism, and crustal flow under the Central Andes, Mem. Geol. Soc. Am., 204, 229–259. Kay, S. M., B. Coira, and J. Viramonte (1994), Young mafic back‐arc volcanic rocks as indicators of continental lithospheric delamination beneath the Argentine Puna plateau, Central Andes, J. Geophys. Res., 99, 24,323–24,339, doi:10.1029/94JB00896. Kirschbaum, A., F. Hongn, and N. Menegatti (2006), The Cobres Plutonic Complex, eastern Puna (NW Argentina): Petrological and structural constraints for Lower Paleozoic magmatism, J. South Am. Earth Sci., 21(3), 252–266, doi:10.1016/j.jsames.2006.04.003. Kley, J. (1999), Geologic and geometric constraints on a kinematic model of the Bolivian orocline, J. South Am. Earth Sci., 12, 221–235, doi:10.1016/S0895-9811(99)00015-2. Kley, J., and C. R. Monaldi (1998), Tectonic shortening and crustal thickness in the Central Andes: How good is the correlation?, Geology, 26(8), 723–726, doi:10.1130/0091-7613(1998)0262.3.CO;2. Kley, J., and C. R. Monaldi (2002), Tectonic inversion in the Santa Bárbara system of the central Andean foreland thrust belt, northwestern Argentina, Tectonics, 21(6), 1061, doi:10.1029/2002TC902003. Kraemer, B., D. Adelmann, M. Alten, W. Schnurr, K. Erpenstein, E. Kiefer, P. van den Bogaard, and K. Görler (1999), Incorporation of the Paleogene foreland into the Neogene Puna plateau: The Salar de Antofalla area, NW Argentina, J. South Am. Earth Sci., 12, 157–182, doi:10.1016/ S0895-9811(99)00012-7. Londoño, J., and J. M. Lorenzo (2004), Geodynamics of continental plate collision during late tertiary foreland basin evolution in the Timor Sea: Constraints from foreland sequences, elastic flexure and normal faulting, Tectonophysics, 392, 37–54, doi:10.1016/j.tecto.2004.04.007. Lucassen, F., R. Becchio, and G. Franz (2011), The Early Palaeozoic high‐ grade metamorphism at the active continental margin of West Gondwana in the Andes (NW Argentina/N Chile), Int. J. Earth Sci., 100, 445–463, doi:10.1007/s00531-010-0585-3. Mack, G. H., W. C. James, and H. C. Monger (1993), Classification of paleosols, Geol. Soc. Am. Bull., 105, 129–136, doi:10.1130/00167606(1993)1052.3.CO;2. Maksaev, V., and M. Zentilli (1999), Fission track thermochronology of the Domeyko Cordillera, northern Chile: Implications for Andean tectonics and porphyry copper metallogenesis, Explor. Min. Geol., 8, 65–89. Marquillas, R. A., C. del Papa, and I. F. Sabino (2005), Sedimentary aspects and paleoenvironmental evolution of a rift basin: Salta Group (Cretaceous‐Paleogene), northwestern Argentina, Int. J. Earth Sci., 94, 94–113, doi:10.1007/s00531-004-0443-2. Marrett, R., and M. R. Strecker (2000), Response of intracontinental deformation in the central Andes to late Cenozoic reorganization of South American Plate motions, Tectonics, 19, 452–467, doi:10.1029/ 1999TC001102. Marrett, R. A., R. W. Allmendinger, R. N. Alonso, and R. E. Drake (1994), Late Cenozoic tectonic evolution of the Puna Plateau and adjacent foreland, northwestern Argentine Andes, J. South Am. Earth Sci., 7(2), 179–207, doi:10.1016/0895-9811(94)90007-8. Masek, J. G., B. L. Isacks, T. L. Gubbels, and E. J. Fielding (1994), Erosion and tectonics at the margins of continental plateaus, J. Geophys. Res., 99, 13,941–13,956, doi:10.1029/94JB00461. Mazzuoli, R., et al. (2008), Miocene magmatism and tectonics of the easternmost sector of the Calama–Olacapato–El Toro fault system in Central Andes at ∼24°S: Insights into the evolution of the Eastern Cordillera, Geol. Soc. Am. Bull., 120, 1493–1517, doi:10.1130/B26109.1. McGlue, M. M., and A. S. Cohen (2006), The evolution of topographic closure and foreland basin lakes: A modern time‐slice perspective from the Andes, Geol. Soc. Am, Abstr. Programs, 38(7), 368. McKenzie, D. (1978), Some remarks on the development of sedimentary basins, Earth Planet. Sci. Lett., 40, 25–32, doi:10.1016/0012-821X(78) 90071-7. McQuarrie, N. (2002a), The kinematic history of the central Andean fold‐ thrust belt, Bolivia: Implications for building a high plateau, Geol. Soc. Am. Bull., 114, 950–963, doi:10.1130/0016-7606(2002)1142.0.CO;2. McQuarrie, N. (2002b), Initial plate geometry, shortening variations, and evolution of the Bolivian orocline, Geology, 30, 867–870, doi:10.1130/ 0091-7613(2002)0302.0.CO;2. McQuarrie, N., B. K. Horton, G. Zandt, S. Beck, and P. G. DeCelles (2005), Lithospheric evolution of the Andean fold–thrust belt, Bolivia, and the origin of the central Andean plateau, Tectonophysics, 399, 15–37, doi:10.1016/j.tecto.2004.12.013.

TC6013

McQuarrie, N., J. B. Barnes, and T. A. Ehlers (2008), Geometric, kinematic, and erosional history of the central Andean Plateau, Bolivia (15–17°S), Tectonics, 27, TC3007, doi:10.1029/2006TC002054. Miall, A. D. (1978), Fluvial sedimentology: An historical review, in Fluvial Sedimentology, edited by A. D. Miall, Mem. Can. Soc. Pet. Geol., 5, 1–47. Mitrovica, J. X., C. Beaumont, and G. T. Jarvis (1989), Tilting of continental interiors by the dynamical effects of subduction, Tectonics, 8, 1079–1094, doi:10.1029/TC008i005p01079. Molnar, P., and G. A. Houseman (2004), The effects of buoyant crust on the gravitational instability of thickened mantle lithosphere at zones of intracontinental convergence, Geophys. J. Int., 158, 1134–1150, doi:10.1111/j.1365-246X.2004.02312.x. Molnar, P., P. England, and J. Martinod (1993), Mantle dynamics, the uplift of the Tibetan Plateau, and the Indian monsoon, Rev. Geophys., 31, 357–396, doi:10.1029/93RG02030. Monaldi, C. R., and J. Kley (1997), Balanced cross sections of the northern Santa Bárbara system and Sierra de Zapla, northwestern Argentina, paper presented at VIII Congreso Geológico Chileno, Congr. Geol. Chil., Antofagasta, Chile. Monaldi, C. R., J. A. Salfity, and J. Kley (2008), Preserved extensional structures in an inverted Cretaceous rift basin, northwestern Argentina: Outcrop examples and implications for fault reactivation, Tectonics, 27, TC1011, doi:10.1029/2006TC001993. Moreno, J. A. (1970), Estratigrafía y paleogeografía del Cretácico superior en la cuenca del noroeste argentino, con especial mención de los subgrupos Balbuena y Santa Bárbara, Asoc. Geol. Argent. Rev., 24, 9–44. Mortimer, E., B. Carrapa, C. Isabelle, L. Schoenbohm, E. R. Sobel, J. S. Gomez, and M. R. Strecker (2007), Fragmentation of a foreland basin in response to out‐of‐sequence basement uplifts and structural reactivation: El Cajón–Campo del Arenal basin, NW Argentina, Geol. Soc. Am. Bull., 119(5–6), 637–653, doi:10.1130/B25884.1. Mpodozis, C., C. Arriagada, M. Basso, P. Roperch, P. Cobbold, and M. Reich (2005), Late Mesozoic to Paleogene stratigraphy of the Salar de Atacama basin, Antofagasta, northern Chile: Implications for the tectonic evolution of the central Andes, Tectonophysics, 399, 125–154, doi:10.1016/j.tecto.2004.12.019. Oncken, O., D. Hindle, J. Kley, K. Elger, P. Victor, and K. Schemmann (2006), Deformation of the central Andean upper plate system—Facts, fiction, and constraints for plateau models, in The Andes: Active Subduction Orogeny, edited by O. Oncken et al., pp. 3–27, Springer, Berlin. Pankhurst, R. J., C. W. Rapela, J. Saavedra, E. Baldo, J. Dahlquist, I. Pascua, and C. M. Fanning (1998), The Famatinian magmatic arc in the southern Sierras Pampeanas, in The Proto‐Andean Margin of Gondwana, edited by R. J. Pankhurst and C. W. Rapela, Geol. Soc. Spec. Pub., 142, 343–367. Pascual, R. (1983), Novedosos marsupiales paleógenos de la Fm. Pozuelos de la Puna: Salta, Ameghiniana, 20, 265–280. Quattrocchio, M., and W. Volkheimer (1990), Paleogene paleoenvironmental trends as reflected by palynological assemblage types, Salta Basin, NW Argentina, Neues Jahrb. Geol. Palaeontol. Abh., 181(1–3), 377–396. Quattrocchio, M., W. Volkheimer, and C. del Papa (1997), Palynology and Paleoenvironment of the “Faja Gris”, Mealla Formation (Salta Group) at Garabatal Creek (NW Argentina), Palynology, 21, 231–247, doi:10.1080/01916122.1997.9989498. Quattrocchio, M., L. Ruiz, and W. Volkheimer (2000), Palynological zonation of the Paleogene of the Colorado and Salta Group basin, Argentina, Rev. Esp. Micropaleontol., 32(1), 61–78. Ramos, V. A. (2008), The basement of the central Andes: The Arequipa and related terranes, Annu. Rev. Earth Planet. Sci., 36, 289–324, doi:10.1146/annurev.earth.36.031207.124304. Ramos, V. A. (2009), Anatomy and global context of the Andes: Main geologic features and the Andean orogenic cycle, Mem. Geol. Soc. Am., 204, 31–66. Reutter, K.‐J., R. Döbel, T. Bogdanic, and J. Kley (1994) Geological map of the Central Andes between 20° and 26°S, in Tectonics of the Southern Central Andes, edited by K.‐J. Reutter, E. Scheuber, and P. J. Wigger, scale 1:1,000,000, Springer, Berlin. Reyes, F. C., and J. A. Salfity (1973), Consideraciones sobre la estratigrafía del Cretácico (subgrupo Pirgua) del noroeste Argentino, Actas Congr. Geol. Argent., 5, 355–385. Reynolds, J., C. Galli, R. Hernández, B. Idleman, J. Kotila, R. Hilliard, and C. Naeser (2000), Middle Miocene tectonic development of the Transition Zone, Salta Province, northwest Argentina: Magnetic stratigraphy from the Metán Subgroup, Sierra de González, Geol. Soc. Am. Bull., 112(11), 1736–1751, doi:10.1130/0016-7606(2000)1122.0. CO;2.

29 of 30

TC6013


Reynolds, J. H., R. M. Hernández, C. I. Galli, and B. D. Idleman (2001), Magnetostratigraphy of the Quebrada La Porcelana section, Sierra de Ramos, Salta Province, Argentina; age limits for the Neogene Orán Group and uplift of the southern Sierras Subandinas, J. South Am. Earth Sci., 14, 681–692, doi:10.1016/S0895-9811(01)00069-4. Roeder, D. (1988), Andean‐age structure of Eastern Cordillera (province of La Paz, Bolivia), Tectonics, 7, 23–39, doi:10.1029/TC007i001p00023. Salfity, J. A., and R. Marquillas (1994), Tectonic and sedimentary evolution of the Cretaceous–Eocene Salta Group, Argentina, in Cretaceous Tectonics of the Andes, Earth Evolution Sciences, edited by J. A. Salfity, pp. 266–315, Vieweg, Heidelberg. Schlunegger, F., T. E. Jordan, and E. M. Klaper (1997), Controls of erosional denudation in the orogen on foreland basin evolution: The Oligocene central Swiss Molasse Basin as an example, Tectonics, 16, 823–840, doi:10.1029/97TC01657. Sempere, T., R. F. Butler, D. R. Richards, L. G. Marshall, W. Sharp, and C. C. Swisher (1997), Stratigraphy and chronology of Late Cretaceous‐ early Paleogene strata in Bolivia and northwest Argentina, Geol. Soc. Am. Bull., 109, 709–727, doi:10.1130/0016-7606(1997)1092.3.CO;2. Sheffels, B. M. (1990), Lower bound on the amount of crustal shortening in the central Bolivian Andes, Geology, 18, 812–815, doi:10.1130/00917613(1990)0182.3.CO;2. Sinclair, H. D., B. J. Coakley, P. A. Allen, and A. B. Watts (1991), Simulation of foreland basin stratigraphy using a diffusion model of mountain belt uplift and erosion: An example from the central Alps, Switzerland, Tectonics, 10, 599–620, doi:10.1029/90TC02507. Sobel, E. R., and M. R. Strecker (2003), Uplift, exhumation and precipitation: Tectonic and climatic control of Late Cenozoic landscape evolution in the northern Sierras Pampeanas, Argentina, Basin Res., 15, 431–451, doi:10.1046/j.1365-2117.2003.00214.x. Sobolev, S. V., and A. Y. Babeyko (2005), What drives orogeny in the Andes?, Geology, 33, 617–620, doi:10.1130/G21557.1. Sobolev, S. V., A. Y. Babeyko, I. Koulakov, and O. Oncken (2006), Mechanism of the Andean orogeny: Insight from the numerical modeling, in The Andes: Active Subduction Orogeny, edited by O. Oncken et al., pp. 513–535, Springer, Berlin. Soreghan, G. S., M. J. Soreghan, and M. A. Hamilton (2008), Origin and significance of loess in late Paleozoic western Pangaea: A record of tropical cold?, Palaeogeogr. Palaeoclimatol. Palaeoecol., 268, 234–259, doi:10.1016/j.palaeo.2008.03.050. Starck, D., and L. M. Anzótegui (2001), The late Miocene climatic change —Persistence of a climatic signal through the orogenic stratigraphic

TC6013

record in northwestern Argentina, J. South Am. Earth Sci., 14, 763–774, doi:10.1016/S0895-9811(01)00066-9. Starck, D., and A. Schulz (1996), La configuración structural del limite entre Cordillera Oriental y Sierras Subandinas en el extremo norte de la Republica Argentina, Bol. Inf. Petrol., 52, 39–46. Starck, D., and G. Vergani (1996), Desarrollo tecto‐sedimentario del Cenozoico en el sur de la Provincia de Salta—Argentina, Congr. Geol. Argent., 8, 433–452. Strecker, M. R., R. N. Alonso, B. Bookhagen, B. Carrapa, G. E. Hilley, E. R. Sobel, and M. H. Trauth (2007), Tectonics and climate of the southern Central Andes, Annu. Rev. Earth Planet. Sci., 35, 747–787, doi:10.1146/annurev.earth.35.031306.140158. Strecker, M. R., G. E. Hilley, B. Bookhagen, and E. R. Sobel (2011), Structural, geomorphic and depositional characteristics of contiguous and broken foreland basins: Examples from the eastern flanks of the central Andes in Bolivia and NW Argentina, in Tectonics of Sedimentary Basins: Recent Advances, edited by C. Busby and A. Azor, Blackwell, Cambridge, Mass., in press. Tassara, A., C. Swain, R. Hackney, and J. Kirby (2007), Elastic thickness structure of South America estimated using wavelets and satellite‐derived gravity data, Earth Planet. Sci. Lett., 253, 17–36, doi:10.1016/j. epsl.2006.10.008. Uba, C. E., C. Heubeck, and C. Hulka (2006), Evolution of the late Cenozoic Chaco foreland basin, southern Bolivia, Basin Res., 18, 145–170, doi:10.1111/j.1365-2117.2006.00291.x. Uba, C. E., J. Kley, M. R. Strecker, and A. K. Schmitt (2009), Unsteady evolution of the Bolivian Subandean thrust belt; the role of enhanced erosion and clastic wedge progradation, Earth Planet. Sci. Lett., 281, 134–146, doi:10.1016/j.epsl.2009.02.010. Xie, X., and P. L. Heller (2009), Plate tectonics and basin subsidence, Geol. Soc. Am. Bull., 121, 55–64. Yuan, X., et al. (2000), Subduction and collision processes in the central Andes constrained by converted seismic phases, Nature, 408, 958–961. Yuan, X., S. V. Sobolev, and R. Kind (2002), Moho topography in the central Andes and its geodynamic implications, Earth Planet. Sci. Lett., 199, 389–402, doi:10.1016/S0012-821X(02)00589-7. B. Carrapa, P. G. DeCelles, and G. E. Gehrels, Department of Geosciences, University of Arizona, 1040 E. 4th St., Tucson, AZ 85721, USA. ([email protected]) B. K. Horton, Department of Geological Sciences, Jackson School of Geosciences, University of Texas at Austin, 1 University Station C1100, Austin, TX 78712, USA.

30 of 30