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
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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‐
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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‐
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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].
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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
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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
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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,
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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-
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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,
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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;
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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.
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Figure 5. Measured stratigraphic sections of the Paleocene‐Miocene strata in the southern part of the
study area.
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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
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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].
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[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).
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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
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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.,
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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 <19 Ma age for the Rio Seco Formation. The detrital
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Figure 9. U‐Pb detrital zircon age probability plots and mean youngest‐age plots for samples from the
Cenozoic deposits. Mean age plots incorporate both random and systematic errors. Ages shown on mean
age plots are not included in the probability plots. See Table S2 for data.
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and by magnetostratigraphy [Kraemer et al., 1999; Reynolds
et al., 2000, 2001; Coutand et al., 2001; Hain et al., 2011;
Carrapa et al., 2011a, 2011b]. Three tuffaceous beds were
sampled in the Agujas Conglomerate (section QT, Figure 5),
and yielded mean U‐Pb zircon ages of 10.49 ± 0.56 Ma,
9.1 ± 1.6 Ma, and 8.85 ± 0.35 Ma (Figure 10). Evidence
for structural growth in this section (see section 4.6) provides a basis for locating the orogenic strain front directly
west during late Miocene time.
6. Provenance Data: Sedimentary Petrology
and Apatite Fission Track Data
Figure 10. Mean U‐Pb ages for three tuffs collected in section QT. Mean ages include random and systematic errors;
“n” represents number of grains used in each calculation.
zircon ages demonstrate that eolian lithofacies recognized
widely from the Puna to the Cordillera Oriental [Starck and
Vergani, 1996] were deposited between ∼22 and 18 Ma
[Carrapa et al., 2011a; this study].
[36] Upper Miocene strata in the Puna, Cordillera Oriental,
and Santa Bárbara ranges are relatively well‐dated by U‐Pb
zircon and fission track methods on volcanogenic deposits,
[37] Fifty‐six standard petrographic thin sections from the
Monte Nieva and Tres Cruces sections were stained for
Ca‐ and K‐feldspars and point‐counted (450 counts per slide)
according to the Gazzi‐Dickinson method [Ingersoll et al.,
1984] in order to assess bedrock provenance of the Cenozoic deposits of the study area. Petrographic parameters are
listed in Table 3, and recalculated modal data are listed in
Table 4. Standard ternary diagrams are shown in Figure 11.
[38] These sandstones are quartzolithic and feldspatholithic arenites, containing abundant monocrystalline quartz
(Qm) grains, with lesser amounts of polycrystalline quartz
(Qp), tectonized polycrystalline quartz (Qpt), and quartzose
sedimentary lithic fragments (Qss). Plagioclase is slightly
more abundant than K‐feldspar. K‐feldspar types include
orthoclase (dominant), microcline, and perthite. Sedimentary
and low‐grade metasedimentary lithic fragments comprise
limestone, shale, chert, quartzose sandstone (Qss), siltstone,
phyllite, and fine‐grained quartz‐mica schist. Volcanic grains
include lathwork (Lvl), microlitic (Lvm), vitric (Lvv), felsic
(Lvf), and rare mafic (Lvma) varieties. Trace amounts of
micas, tourmaline, cordierite, epidote and zoisite, zircon,
magnetite, and Fe‐oxide alterite fragments are present.
[39] Modal sandstone compositions are regionally consistent, with monomineralic fractions dominated by quartz,
and lithic compositions dominated by low‐grade metasedimentary grains (Figure 11). Average values of key ternary
modes are given in Table 4. These data are consistent with
deposition in a basin dominated by upper crustal source
terranes, and indicate hybrid Recycled Orogenic and Continental Block provenance (Figure 11) [Dickinson, 1983],
with some influence from magmatic arc sources. In both
sections the ratio Lt/QmFLt increases upward, and Qm/
QmFLt decreases upward. Increased influence of magmatic
arc provenance is evident in both sections as the proportion
of volcanic lithic grains increases upward. Increased arc
influence is also indicated by influx of detrital zircons with
ages that are roughly equivalent to depositional age (see
section 7). The major sandstone compositional difference
between northern (Tres Cruces) and southern (Monte Nieva)
localities is in the amount of feldspar present. Sandstones in
the Monte Nieva locality contain approximately twice as
much feldspar as those from Tres Cruces (Table 4). The
increase in feldspar toward the south includes both plagioclase and K‐feldspar.
[40] Conglomerate clast counts (at least 100 per station,
counted on a regular grid spacing) in the Quebrada de los
Colorados and Angastaco Formations and the Agujas Conglomerate are dominated by sedimentary‐metasedimentary
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Table 2. Apatite Fission Track Dataa
Cobble
Irridiation Number
N
Rho‐S (e5)
NS
Rho‐I (e5)
NI
U (ppm)
P(c)2
Rho‐D (e6)
ND
Age (Ma)
Error
MN2936
MN3(1835 m)
MN1333
UW09‐03‐42
UW10‐05‐14
UW09‐03‐35
12
18
25
2.785
4.415
4.255
166
203
361
19.509
32.711
29.211
1163
1504
2478
25.25
33.89
35.45
24.95
25.11
24.22
1.0081
1.2024
1.0681
4018
3146
4018
26.5
29.7
28.7
2.3
2.7
1.9
Age is pooled when c2 > 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.
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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
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DECELLES ET AL.: ANDEAN FORELAND BASIN
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.
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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
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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
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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.
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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.
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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
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Figure 16
22 of 30
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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
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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.
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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
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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
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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 (<10 km) low‐angle thrust décollements. In this
context, the upper Miocene‐Quaternary deposits represent
wedge‐top accumulations that were depositionally integrated
into the regional foreland basin system, as in the case of the
Quaternary to modern record of the frontal Andes. This view
is consistent with the regional kinematic history of the central
Andes and reflection seismic data and balanced regional
cross‐sections that show faults of the Santa Bárbara system
merging with a regional detachment at 15–20 km depth
[Grier et al., 1991; Cristallini et al., 1997; Mortimer et al.,
2007].
[63] 2. As discussed in the Introduction, previous workers
have interpreted the Santa Bárbara Subgroup as the result of
post‐Salta rift thermal subsidence [e.g., Marquillas et al.,
2005, and references therein]. This interpretation is problematic for several reasons. First, the Santa Bárbara
Subgroup and its stratigraphic equivalents in Bolivia blanket
the entire Cordillera Oriental, Subandes, Santa Barbara
system, and the eastern Puna and Altiplano [Sempere et al.,
1997; DeCelles and Horton, 2003], far exceeding what
would be expected for post‐rift deposits related exclusively
to the Salta rift. Control by the Salta rift on the spatial
distribution of the Santa Bárbara Subgroup is not evident,
with the exception of the Lomas de Olmedo arm of the
rift, where subsurface data show axial thickening of Santa
Bárbara strata [Cominguez and Ramos, 1995] (see also
unpublished petroleum company seismic data). Second, it is
not clear how the Santa Bárbara Subgroup relates to the
underlying Balbuena Subgroup in terms of standard models
for rift‐basin subsidence [e.g., McKenzie, 1978; Bond and
Kominz, 1984]. Salfity and Marquillas [1994] proposed
that these two packages of rocks represent discrete pulses of
‘post‐rift’ thermal subsidence. Whereas we concur with the
post‐rift thermal subsidence interpretation for the fine‐
grained carbonate‐rich Balbuena Subgroup—indeed the
Balbuena Subgroup lies in a depositional continuum on top
of the synrift Pirgua Subgroup and exhibits many of the
classic features of post‐rift sedimentation—models for rift‐
basin subsidence provide no explanation for a second, post‐
rifting pulse of thermally driven subsidence that would be
necessary to accommodate the Santa Bárbara Subgroup and
its equivalents in Bolivia. Third, the post‐rift interpretation
provides no explanation for the abrupt coarsening in the
Mealla and Maiz Gordo Formations; in many sections these
units are dominated by sandstone and are locally conglomeratic. Standard models for rift‐ to post‐rift sediment accumulation predict continued fining and decreased sediment
accumulation rates with time. We suggest that paleotopographic elements associated with the Salta rift were locally
prominent during Paleocene‐early Eocene time in northwestern Argentina, such as in the Angastaco and surrounding regions [Sobel and Strecker, 2003; Mortimer et al.,
2007; Carrera and Muñoz, 2008; Carrapa et al., 2011a]
(Figure 13), but that the overall distribution of the Santa
Bárbara Subgroup was dictated by the existence of the
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nascent Andean orogenic wedge, which was located several
hundred km to the west in northern Chile during Paleocene‐
early Eocene time [Maksaev and Zentilli, 1999; Arriagada
et al., 2006]. Fourth, paleocurrent data from fluvial sandstones in these units indicate consistent paleoflow from the
west (in Bolivia) and southwest (Argentina), and sedimentary petrological data and detrital zircon ages indicate a
cosmopolitan provenance in Proterozoic and Paleozoic
sedimentary and igneous rocks distributed over a wide
region in the Puna, Altiplano, and the orogenic hinterland.
Were the Mealla and Maiz Gordo related exclusively to post‐
rift thermal subsidence, we would expect to see stronger
local provenance and paleoflow control by remnant rift‐
related topographic highlands. Fifth, the post‐rift thermal
subsidence interpretation does not provide an explanation for
the supersol/disconformity and its progressive eastward
migration. Indeed, the supersol has not been recognized in
literature that attributes the Santa Bárbara Subgroup to post‐
rift thermal subsidence, suggesting that its existence has not
heretofore been taken into account. Some authors have
highlighted minor (ca. 10 m) fault offsets in the Santa
Bárbara Subgroup in reflection seismic profiles as evidence
that the interval accumulated while extension continued
[Cominguez and Ramos, 1995; Monaldi et al., 2008].
However, these minor offsets could be explained as well by
flexural extension during passage through the region of the
flexural forebulge [e.g., Bradley and Kidd, 1991; Blisniuk
et al., 1998; Londoño and Lorenzo, 2004].
[64] 3. Finally, we note that the persistence of a regional,
eastward migrating foreland basin system in northwestern
Argentina since Paleocene time is not consistent with a
dominantly pure shear mode of crustal thickening in this
part of the Andes [Allmendinger, 1986; Allmendinger and
Gubbels, 1996]. Instead, it appears that shortening and
crustal thickening in this part of the orogenic belt produced a
load that was supported by the flexural strength of adjacent
foreland lithosphere from Paleocene time to the present.
South of latitude 27°S, however, a Paleogene foreland basin
has not been documented [Carrapa et al., 2008], large
basement blocks are involved in the orogenic belt (the
northern Sierras Pampeanas) [Jordan et al., 1983; Jordan
and Allmendinger, 1986], and the timing of deformation
may be more regionally diffuse [Mortimer et al., 2007;
Dávila and Astini, 2007; Carrapa et al., 2006, 2008]. This
leaves open the possibility that pure shear shortening may be
operating farther south. In addition, our argument against
dominantly pure shear thickening does not preclude the
possibility that the orogenic belt thickened by pure shear in
its hinterland while the frontal part of the orogenic belt
operated in a simple shear mode.
11.
Conclusions
[65] 1. The Paleogene‐Neogene stratigraphy of northwestern Argentina records the early to mature central
Andean foreland basin system. Previously interpreted
Paleocene‐Eocene post‐rift sag deposits are reinterpreted to
contain evidence of nascent Andean thrust belt development, far to the west in northern Chile. Accommodation for
these deposits was likely a result of combined distal back‐
bulge subsidence, some component of dynamic subsidence
associated with the subducting oceanic Nazca plate, and
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perhaps local late‐stage thermal subsidence in easternmost
regions of Cretaceous rifting. However, post‐rift thermal
subsidence alone cannot account for the dominantly
westerly provenance and widespread distribution of the
Paleocene‐Eocene deposits from northwestern Argentina
into central Bolivia. Nor is post‐rift thermal subsidence
consistent with the presence of a regional (>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.
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