Carrapa, B., Adelmann, D., Hilley, G., Mortimer, E., Strecker, M.R.

TECTONICS, VOL. 24, TC4011, doi:10.1029/2004TC001762, 2005
Oligocene range uplift and development of plateau morphology
in the southern central Andes
B. Carrapa,1 D. Adelmann,2 G. E. Hilley,3 E. Mortimer,1 E. R. Sobel,1 and M. R. Strecker1
Received 26 October 2004; revised 14 February 2005; accepted 11 April 2005; published 13 August 2005.
[1] The Puna-Altiplano plateau in South America is a
high-elevation, low internal relief landform that is
characterized by internal drainage and hyperaridity.
Thermochronologic and sedimentologic observations
from the Sierra de Calalaste region in the southwestern
Puna plateau, Argentina, place new constraints on
early plateau evolution by resolving the timing of
uplift of mountain ranges that bound present-day
basins and the filling pattern of these basins during
late Eocene-Miocene time. Paleocurrent indicators,
sedimentary provenance analyses, and apatite fission
track thermochronology indicate that the original
paleodrainage setting was disrupted by exhumation
and uplift of the Sierra de Calalaste range between
24 and 29 Ma. This event was responsible for
basin reorganization and the disruption of the
regional fluvial system that has ultimately led to the
formation of internal drainage conditions, which, in
the Salar de Antofalla, were established not later than
late Miocene. Upper Eocene-Oligocene sedimentary
rocks flanking the range contain features that suggest
an arid environment existed prior to and during its
uplift. Provenance data indicate a common similar
source located to the west for both the southern Puna
and the Altiplano of Bolivia during the late EoceneOligocene with sporadic local sources. This suggests
the existence of an extensive, longitudinally oriented
foreland basin along the central Andes during this
time. Citation: Carrapa, B., D. Adelmann, G. E. Hilley,
E. Mortimer, E. R. Sobel, and M. R. Strecker (2005), Oligocene
range uplift and development of plateau morphology in the
southern central Andes, Tectonics, 24, TC4011, doi:10.1029/
2004TC001762.
1. Introduction
[2] The central Andean Altiplano-Puna plateau is a
hyperarid, low internal relief, high-elevation region with
1
Institut für Geowissenschaften, Universität Potsdam, Potsdam,
Germany.
2
Institut für Geowissenschaften, Friedrich-Schiller-Universität Jena,
Jena, Germany.
3
Department of Earth and Planetary Science, University of California,
Berkeley, California, USA.
Copyright 2005 by the American Geophysical Union.
0278-7407/05/2004TC001762$12.00
average and peak elevations greater than 3700 and 6000 m,
respectively. Uplift of this high-elevation region has been
ascribed to processes such as lithospheric thinning [Isacks,
1988] following delamination [Kay et al., 1994], distributed
crustal shortening [Allmendinger et al., 1997], emplacement
of regional basement thrust sheets [Kley et al., 1997;
McQuarrie and DeCelles, 2001], and underthrusting of
the Brazilian craton [Isacks, 1988]. Whereas these processes
may have created much of the surface uplift in the area, the
low-relief morphology, typical of continental orogenic
plateaus, may result from simultaneous erosion of high
peaks and deposition within basins, as the incising power
of regional drainage systems is lost in such arid environments [e.g., Sobel et al., 2003; Hilley and Strecker, 2005].
Despite the importance of internal drainage conditions in
many of the world’s large continental plateaus and the
potential impact of sediment storage on their evolution as
well as adjacent foreland areas [Vandervoort et al., 1995;
Métivier et al., 1998; Tapponnier et al., 2001; Horton et al.,
2002; Sobel et al., 2003], the controls on their establishment
remain elusive.
[3] The clastic fill preserved in intramontane basins
within the plateau contains the unique record of the timing
and pattern of orogenic evolution and its relationship to
tectonics and climate. By constraining the sedimentary
evolution of such basins and the uplift of related ranges, a
better understanding of the processes leading to final
internal drainage can be achieved.
[4] Clastic sediments preserved in the Puna plateau
suggest that sedimentation within the plateau started in the
late Eocene in a broad foreland basin sourced from the west
[Jordan and Alonso, 1987]. However, detailed sedimentological investigations which could assess this hypothesis are
still limited and therefore its validity must be further tested.
During the Miocene, and possibly Oligocene, the appearance of evaporites in northern Argentina documents the
onset of hyperarid conditions and has been directly related
to the establishment of internal drainage in the area [e.g.,
Vandervoort et al., 1995; Alonso et al., 1991]. Thermochronologic and provenance data from the eastern Puna margin
suggest that some topography may have formed as early as
the late Eocene – early Oligocene [Muruaga, 2001; Coutand
et al., 2001; Deeken et al., 2004]. The eastern ranges
constituted a topographic high at least by middle Miocene
time [e.g., Strecker, 1987; Strecker et al., 1989; Grier and
Dallmeyer, 1990; Marrett and Strecker, 2000; Kleinert and
Strecker, 2001; Hilley and Strecker, 2005]. However, the
relationships between range exhumation and uplift, sediment dispersal and final internal drainage development
remain unclear.
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plateau or from climate conditions that may be largely
independent of plateau formation [e.g., Hartley, 2003;
Sobel et al., 2003].
[7] Combined apatite fission track thermochronology
(AFTT) and sedimentologic investigations in the southwestern Puna plateau (Figure 1) provide information on
the timing of exhumation and related uplift of basinbounding ranges, the sedimentary dynamics within adjacent basins, and climatic conditions during the tectonic
events that preceded plateau uplift. In the southern Puna,
sedimentary basins that contain upper Eocene to Quaternary sedimentary rocks are bounded by high-angle reverse
faults whose hanging wall rocks form bedrock-cored
mountain ranges (Figure 2). By constraining the timing
of exhumation and related uplift of a bounding range and
changes in the sedimentary dynamics of the basins, we
suggest that exhumation and uplift of ranges might have
been the trigger for basin compartmentalization which
eventually led to internal drainage development. Our
results support the hypothesis that the deformation driving
range uplift started at least by Oligocene time, contributing
to the establishment of internal drainage and to the
characteristic high elevation, low internal relief observed
today within the Puna plateau.
2. Regional Setting
2.1. Tectonic Evolution of the Central Southern Andes
Figure 1. General map of the central Andes including
different morphotectonic domains and area over 3 km
elevation (gray) (modified after Horton et al. [2001]).
[5] In addition, constraints on the timing of initiation of
deformation related to range uplift and basin compartmentalization within the present plateau area in northwestern
Argentina are limited. At present, a widely accepted model
proposes that deformation leading to crustal thickening and
subsequent uplift occurred during the middle-late Miocene
[e.g., Allmendinger, 1986; Isacks, 1988; Allmendinger et
al., 1997; Jordan et al., 1997, 2001]. However, subsequent
studies document pre-Miocene deformation in the present
plateau of Argentina [e.g., Coutand et al., 2001] and
Bolivia [McQuarrie and DeCelles, 2001; Horton et al.,
2001, 2002; DeCelles and Horton, 2003; Ege, 2004; Elger,
2004].
[6] Extensive studies exist for the Altiplano and Eastern
Cordillera of Bolivia documenting an initial pattern of
foreland basin development, followed by structural disruption, drainage internalization and compartmentalization of
sediment basins during Eocene through early Miocene
time [e.g., Horton et al., 2001; Horton et al., 2002;
Ege, 2004]. Comparable comprehensive investigations that
document the timing and pattern of such processes are
scarce in the Puna region. In particular, it remains unclear
if deformation driving marginal range uplift, basin compartmentalization and subsequent infill of related sedimentary basins result from tectonic processes that form the
[8] The Puna-Altiplano plateau is part of the central
Andes and represents the second largest plateau on Earth
after Tibet. The southern Puna of northwestern Argentina
constitutes the southern end of the intraorogenic plateau. It
is bounded to the west by a magmatic arc and to the
northeast by the Eastern Cordillera fold-and-thrust belt,
while the eastern border is transitional to the high-angle
reverse-fault bounded Sierra Pampeanas basement blocks
(Figure 1). The 3700-m-high southern Puna plateau is
characterized by Neogene contraction [Alonso, 1986;
Allmendinger et al., 1997; Coutand et al., 2001; Ege,
2004], meridionally trending mountain ranges (often in
excess of 6000 m elevation), and internally draining sedimentary basins.
[9] During the late Eocene, contractional to transpressional deformation (‘‘Incaic phase’’ [Steinmann, 1929])
involved Upper Cretaceous to Paleogene rocks of the
present Western Cordillera [Günther et al., 1998]. Deformation and uplift of this belt is inferred to have triggered
deposition of the earliest clastic sequences in an extensive
foreland basin spanning both the present-day plateau and
regions to the east [e.g., Jordan and Alonso, 1987; Horton
et al., 2001; DeCelles and Horton, 2003]. This model is
supported by apatite fission track thermochronology in the
Chilean cordillera, indicating considerable exhumation
between 50 and 30 Ma [Maksaev and Zentilli, 2000] and
by sedimentological data indicating westerly sourced
late Eocene-Oligocene sedimentary rocks in the Altiplano
[Horton et al., 2002] and possibly in the Puna [Jordan and
Alonso, 1987]. Sedimentation continued in isolated intramontane basins that received sediments from more local
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Figure 2. (a) Shaded relief map of the central Andes, based on GTOPO30 data (USGS). AB, Arizaro
Basin; AD, Atacama Desert; CA, Campo Arenal; HM, Salar de Hombre Muerto; QT, Quebrada del Toro;
H, Humahuaca; SA, Salar de Antofalla; SC, Sierra de Calalaste; Scu, Siete Curvas; Spg, Salar de Pastos
Grande; TC, Tres Cruces; VC, Valles Calchaquı́es basin. Faults are modified after Reutter et al. [1994],
Urreiztieta et al. [1996], and Coutand et al. [2001]. Dots indicate volcanoes. Star denotes 30.3 ± 3 Ma
apatite fission track (AFT) age, after Andriessen and Reutter [1994]; triangle marks AFT ages between
30 ± 2 Ma and 38 ± 3 Ma from Coutand et al. [2001]. (b) Simplified geological map of the central Andes
modified after McQuarrie [2002a]. For a more detailed geological map of the study area we refer to
Figure 3.
sources from at least Miocene until present time [e.g.,
Jordan and Alonso, 1987; Vandervoort, 1993; Horton et
al., 2001; DeCelles and Horton, 2003].
[10] Basins within the plateau contain thick sequences of
continental evaporites and clastic deposits that yield fundamental information as to the cooling/exhumation history of
hinterland sources, sediment dispersal, and provenance.
These basins are bounded structurally by high-angle reverse
faults [e.g., Jordan et al., 1997]. The timing of clastic
sedimentation in basins within the plateau and along the
eastern Puna border is relatively well known based on
magnetostratigraphy and 40Ar/39Ar dating on ash layers in
synorogenic deposits [Coira et al., 1993; Kay et al., 1994;
Marrett and Strecker, 2000; Coutand et al., 2001].
2.2. Geology of the Sierra de Calalaste Area
[11] The study area is located in the southern Puna
between the Salar de Antofalla and the Sierra de Calalaste
(Figures 2 and 3). Tertiary E-W to WNW-ESE shortening
produced a series of east and west vergent reverse and
thrust faults striking parallel to the present Salar de
Antofalla (Figure 4) [Voss, 2000; Adelmann, 2001]. The
Sierra de Calalaste constitutes low-grade metamorphic
basement rocks that were deformed during and after late
Eocene time [Adelmann, 2001]. Within the Sierra de
Calalaste, Paleozoic sedimentary rocks are thrust over
Tertiary sedimentary rocks along west and east verging
reverse faults (Figure 3). Crystalline basement rocks,
including migmatitic gneisses, metabasites, granitoids and
aplites, as well as Tertiary volcanics rocks crop out to the
west and southwest of the present-day Salar de Antofalla
area [e.g., Kraemer et al., 1999, and references therein]
(Figures 2 and 3). In contrast, Precambrian sedimentary
and low-grade metamorphic rocks are more widespread to
the east (Figures 2 and 3).
[12] In the study area, sedimentation started in late
Eocene – early Oligocene time with the deposition of the
Quiñoas Formation during the Incaic deformation phase
[e.g., Kraemer et al., 1999; Adelmann, 2001; Voss, 2002].
During the late Oligocene, thick-skinned compressive
deformation [Adelmann and Görler, 1998] triggered sedimentation of syntectonic coarse-grained alluvial fans constituting the Chacras Formation (Figure 4). This phase of
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Figure 3. Detailed geologic map of the Antofalla area [see Kraemer et al., 1999] with location of the
analyzed samples and sedimentological logs.
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Figure 4. Overview of the tectonostratigraphic and magmatic development of the Salar de Antofalla area
in the southern Puna modified after Kraemer et al. [1999], Voss [2000], and Adelmann [2001]. Undulated
lines represent angular unconformities of regional significance separating main lithostratigraphic units
present in the Salar de Antofalla area.
early deformation is also marked by an angular unconformity between the Quiñoas and the Chacras formations.
Immediately beneath the unconformity, a tuff layer yields
an 40Ar/39Ar age of 28.9 ± 0.8 Ma (ID-51) [Adelmann and
Görler, 1998]. The oldest strata of the overlying Chacras
Formation have been dated at 24.2 ± 0.9 Ma (ID-86) and
22.5 ± 0.6 Ma (ID-18) [Kraemer et al., 1999].
[13] During the early Miocene (20– 17 Ma), renewed
E-W to WNW-ESE shortening [Adelmann and Görler,
1998] reactivated the west vergent fault system that was
active during the preceding deformation phase. Additionally,
Miocene shortening produced new east and westward
directed basement thrusts onto tilted alluvial fan sediments
of the Potrero Grande Formation [Adelmann, 2001].
In the middle Miocene, west vergent thrusts affected
Lower Paleozoic, Permian and Tertiary rocks [Adelmann,
1997; Voss, 2000; Adelmann, 2001]. This deformation
triggered deposition of the syntectonic Juncalito Formation
(middle Miocene-Pliocene) which is characterized by thick
evaporites (dated as late Miocene [Kraemer et al., 1999])
typical of a hyperarid internal drainage environment
[Kraemer et al., 1999].
3. Methods
[14] Sediment logging, facies interpretation, paleocurrent
and provenance analyses were carried out on upper Eocene –
lower Miocene sedimentary rocks (Quiñoas and Chacras
formations) to reveal changes in sediment source and sedimentary environments (Figures 5 and 6 and Table 1). Fifteen
thin sections were analyzed using the Gazzi-Dickinson
method [Dickinson, 1970; Gazzi et al., 1973]. An average
of 400 framework grains per thin section was counted on
unstained thin sections. Petrographic counting parameters
and recalculated detrital modes are reported in Table 2.
Paleocurrent direction was determined by measuring at least
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Figure 5. Correlation panel of the investigated sedimentological logs indicated in Figure 3. For more
details, refer to Figure 6.
50 imbricated pebbles of clast-supported channelized streamflow and proximal sheet flood units in alluvial fan deposits.
[15] AFTT was conducted on six samples collected along
a vertical transect through the Sierra de Calalaste between
elevations of 3729 and 4455 m to constrain the cooling and
exhumation history of the central part of the range. Samples
were separated and analyzed following the procedure
described by Sobel and Strecker [2003] (Table 3). Raw
data were reduced using the Trackkey program (I. Dunkl,
Trackkey: Windows program for calculation and graphical
presentation of EDM fission track data, version 4.2, 2002,
available at http://www.sediment.uni-goettingen.de/staff/
dunkl/software/trackkey.html). Measurements of fission
track etch pits were made to assess annealing kinetic
variability [Donelick et al., 1999]. Length measurements
were attempted on all samples in order to gain information
on the degree of annealing [e.g., Wagner and Van der
Haute, 1992]. Exhumation rate was calculated using the
inverse slope of weighted least squares regression of the
AFT elevation versus ages.
4. Sedimentology
4.1. Quiñoas Formation
[16] The Quiñoas Formation records the onset of sedimentation in the study area in the late Eocene as determined
by an 40Ar/39Ar age of 37.6 ± 0.3 obtained from an ash
layer in the basal part of the formation [Kraemer et al.,
1999]. The Quiñoas Formation is divided into two members
based on facies associations and a change in the interpreted
depositional environments [Kraemer et al., 1999].
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Figure 6. Detailed sedimentological logs: (a) log R; (b) log U; (c) log G; (d) log H/I; (e) log S. Ages
indicated in log U are from Kraemer et al. [1999]. For legend, refer to Figure 5. The most typical facies
are indicated in italics. For a description of the facies we refer to Table 1.
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Figure 6. (continued)
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Figure 6. (continued)
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Figure 6. (continued)
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Figure 6. (continued)
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Table 1. Lithofacies Description and Interpretation Based on Work by Miall [1996]
Facies Codes
Gm
Lithofacies
Fl
conglomerates, matrix
supported
conglomerates, clast
supported
conglomerates, clast
supported
conglomerates, clast
supported
sand, fine to coarse
sand, fine to very coarse,
may be pebbly
sand, fine to very coarse,
may be pebbly
sand, fine to very coarse,
may be pebbly
sand, very fine to coarse
sand, fine to medium,
well sorted
sand, silt, mud
Fsm/m
P
silt, mud
paleosol carbonate
Gc
Gh
Gt/p
Sm
St
Sp
Sh/l
Sr
Seod
Sedimentary Structures
Interpretation
massive, faint gradation
high-strength (cohesive) debris flow
massive, faint horizontal laminations,
imbrications
horizontal laminations, imbrications
high-strength or low-strength (noncohesive)
debris flow
longitudinal gravel banks, lag deposits
trough and planar cross beds
minor channel fills, transverse bed forms
massive or faint lamination
trough cross beds
distal debris flow
subaqueous 3-D dunes, Transition or upper
part of a flow regime
subaqueous 2-D dunes
planar cross beds
horizontal laminations/low-angle (<15)
cross beds
ripple, cross lamination
large-scale cross lamination (>25)
fine laminations, very small ripples
massive, desiccation cracks
pedogenic features: nodules filaments
4.1.1. Quiñoas I
[17] This member is deposited above an erosional unconformity with the underlying Permo-Carboniferous rocks. At
its maximum recorded thickness, this member reaches
840 m. It is characterized by fine-grained, gypsiferous
siltstones and mudstones (Fsm/m, Fl) intercalated with
massive, horizontally stratified and imbricated conglomerates with undulating basal contacts (Gc); lenticular clastsupported conglomerates (Gt/p); and interbedded massive
and laminated sandstones (Sh/l, Sm) (Table 1). Combined
thicknesses of conglomerate units can exceed 250 m
(Figure 5). In the farthest sections to the west (logs U
and R), these fine grained facies dominate the sections,
while more proximal to the Sierra de Calalaste the section is
coarser and dominated by the conglomeratic facies, though
still with intercalated finer horizons (Figure 5).
[18] We interpret the gypsiferous siltstones and mudstones (Fsm/m, Fl) as being deposited in a playa mudflat
[e.g., Flint, 1985; Hartley et al., 1992]. The presence of fine
grained, laminated sediments, often bioturbated and with
rooting and soil remnants and gypsum, is indicative of an
arid, or at least episodically dry environment [e.g., Hardie
et al., 1978]. The coarse-grained lenticular conglomerates
(Gt/p) are interpreted as having been deposited in fluvial
channels [e.g., Hartley et al., 1992; Miall, 1996]. Massive,
bedded and imbricated conglomerates (Gc) are interpreted to
have been deposited under tractive flow [Rasmussen, 2000].
Horizontally stratified conglomerates (Gh) might reflect
discontinuous discharge and accretion during high-density
flows and sheet floods [Nemec and Steel, 1984]. The
interbedded, laminated and massive sandstones (Sh/l, Sm)
which occur both within the conglomerates and throughout
the finer-grained section represent waning flow conditions
and are probably distal derivatives of debris flows on the
alluvial fan [e.g., Lowe, 1979; Nemec and Steel, 1984]. This
scour fills
ripples (lower flow regime)
aeolian dunes
overbank, abandoned channel or waning
flood deposits
overbank, abandoned or drape deposits
soil with chemical precipitation
facies association is typical of ephemeral discharge in
semiarid to arid alluvial fan environment [e.g., Flint and
Turner, 1988; Sohn, 1997; Rasmussen, 2000]. Paleocurrent
data from member I, on the west side of the present-day
Sierra de Calalaste, suggest a provenance mainly from the
N-NE and S-SW (Figure 3).
4.1.2. Quiñoas II
[19] This member reaches a thickness of about 500 m
adjacent to the Sierra de Calalaste (Figures 5 and 6; log G)
and conformably overlies the underlying member. It comprises fine to very coarse-grained, trough cross-bedded
sandstones and pebbly sandstones (St). These occur in beds
between 0.5 and 5.0 m thickness. Vertically stacked beds of
this facies can combine to form significant thicknesses of
more than 100 m (e.g., log G, U). Alternating with these
larger coarse-grained units are finer-grained sandstones
(decimeter-scale beds) with scoured bases and occasional
climbing ripples occurring in many of the sandstone units
(Sr) and mudstones (Fsm/m, Fl) (Table 1). These mudstones
often contain desiccation cracks, and are bioturbated (e.g.,
Figure 6b, log U).
[20] We interpret the trough cross-bedded sandstones as
being derived from the migration of dune bed forms within
sand bed channels of a fluvial depositional environment
[Miall, 1996]. Scour based sandstones with climbing ripples
represent a lower flow regime within a fluvial environment
[e.g., Allen, 1963]. The mudstones represent the finest
component of the system, and are likely to have been
deposited onto a floodplain environment, possibly through
crevasse splay deposition. The close vertical association
between channel and floodplain deposits suggests deposition onto a fluvial plain [Miall, 1996]. Such a facies
association might be representative of either meandering
or braided river systems [Miall, 1978; Smith, 1987]. However, because of the significant thicknesses of trough cross-
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Quiñoas
Quiñoas
Quiñoas
Quiñoas
Quiñoas
Quiñoas
Quiñoas
Quiñoas
Fm
Fm
Fm
Fm
Fm
Fm
Fm
Fm
I
I
I
I
I
I
I
I
log
log
log
log
log
log
log
log
S
S
R
R
H/I
H/I
H/I
H/I
G
H/I
H/I
S
S
log
log
log
log
log
II
II
II
II
II
Quiñoas
Quiñoas
Quiñoas
Quiñoas
Quiñoas
Fm
Fm
Fm
Fm
Fm
log G
log G
mS
mS
mS
mS
mS
mS
mS
mS
mS/cS
mS
mS
mS
mS
mS
mS
Qp
21.8
39
22.5
19.1
18.4
16.3
18.4
28.3
26.7
21.0
25.7
42.8
30.0
29.1
14.0
6.9
6.9
6.1
11.9
9.6
19.5
9.9
13.6
13.5
13.9
11.8
20.5 5.6
15.6 11.5
Mean
Locationb Grain Sizec Qm
Chacras Fm
Chacras Fm
Stratigraphy
3.8
5.8
5.1
9.1
5.9
4.2
4.3
1.9
6.3
3.2
6.1
3.9
5.8
4.5
4.9
16.0
11.6
14.4
17.7
23.5
22.2
12.0
9.3
20.2
13.6
16.6
13.1
11.3
16.5
29.2
Plag K-Feldspar
Lv
2.6
3.6
2.9
1.7
2.7
2.9
12.8
11.3
0 6.4
1.1 1.9
3.2 4.4
2.4 6.7
6.7 5.1
8.6 2.9
1.9 9.3
1.7 11.0
1.7
0.0
0.2
0.0
0.8
0.9
0.8
0.6
4.1
5.0
7.6
8.4
4.3
7.9
3.2
1.7
3.7
2.1
4.4
4.2
5
2.7
1.9
0.9
1.1
0.5
1.0
1.9
4.0
1.3
0.8
2.3
0.8
1.4
0.3
0.5
0.8
1.9
0
0.8
0
0.5
0
1.1
0.5
1.9
0.6
0.5
1.4
0.3
0
1.1
0.8
4.1
3.6
14.7
16.7
9.3
7.3
6.1
5.2
2.6
7.2
4.1
2.2
8.4
6.9
4.4
9.6
12.7
17.6
10.0
15.5
10.1
20.0
6.9
3.4
9.4
7.5
14.2
18.2
8.8
0.6
32
29
31
2
40
42
46
68
57
51
12
60
65
44
36
33
36
39
56
46
12
Chert
Minor
Lp Fragments Constituents Mica Cement Matrix Q
6.7 0.5
7.7
0
Lm
2.6 14.8 5.1 2.0
1.3 10.6 13.3 3.7
6.9 5.0 6.1 1.4
1.7 0.8 1.9 0.8
3.4 1.6 3.2 0.8
11.2 14.4
4.1 17.5
Ls
25
37
31
8
29
20
26
20
24
24
4
23
21
29
37
40
34
22
13
27
9
F
43
34
39
6
31
38
28
11
19
25
11
17
14
27
27
27
30
39
31
27
8
L
25
17
21
6
29
25
30
52
41
35
11
26
47
34
27
25
21
25
33
30
8
Qm
25
37
31
8
29
20
26
20
24
24
4
23
21
29
37
40
34
23
13
28
9
F
50
46
48
3
42
54
43
28
35
40
10
51
32
37
36
35
45
52
54
43
9
24
38
31
10
44
57
60
88
77
65
17
93
80
70
79
53
63
47
62
68
15
43
50
47
5
48
38
17
4
7
23
19
0
5
16
12
34
28
7
5
13
12
33
12
23
15
8
5
23
8
16
12
7
7
15
14
9
13
9
47
33
18
14
Lt Qp Lvm Lsm
a
Recalculated detrital modes are reported for QFL, Qm, monocrystalline quartz; Qp, polycrystalline quartz; Plag, plagioclase; Ls, sedimentary rock fragments; Lv, volcanic rock fragments; Lm, metamorphic
rock fragments; Lp, plutonic rock fragments. Q, F, L, Lt, Lvm, and Lsm are parameters based on the classification of Dickinson [1970] and Graham et al. [1976]. SD, standard deviation.
b
Location is given in Figures 5 and 6.
c
Here mS is middle sandstone and cS is coarse sandstone.
A091
A092
Average
SD
A075
A081
A082
A443
A444
Average
SD
A442
A441
A412
A411
A090
A100
A103
A108
Average
SD
Sample
Table 2. Sandstone Petrography Parameters Based Methods Described by Ingersoll et al. [1984] and Dickinson [1985]a
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CARRAPA ET AL.: OLIGOCENE UPLIFT IN THE CENTRAL ANDES
TC4011
a
0.2
0.1
0.2
0.1
0.1
0.1
2.2
2.0
2.0
2.1
1.9
2.0
67.4215
67.4819
67.4516
67.4527
67.4606
67.4609
metasediments
metasediments
metasediments
metasediments
metasediments
metasediments
SCAD9
SCAD5
SCAD6
SCAD2
SCAD4
SCAD3
3729
3978
4248
4328
4335
4455
26.1436
26.1618
26.1356
26.1356
26.1526
26.1501
19
6
9
20
20
20
5.176
4.206
3.196
4.519
3.227
2.500
327
57
67
300
252
175
46.711
36.232
31.678
38.528
26.699
22.843
2951
491
664
2558
2085
1599
88
82
62
82
100
46
12.493
12.821
12.739
13.076
12.903
12.985
5166
5166
5166
5166
5166
5166
25.8
27.7
23.9
28.5
29.0
26.5
1.6
3.9
3.1
1.8
2.0
2.2
47
34
30
36
26
23
SD
Dpar,
mm
U,
ppm
±1s
Age,
Ma
NDg
Rho-D,f
105
P(c)2,e %
NId
Rho-I,c
105
NSd
Rho-S,c
105
Number
of Xlsb
Longitude,
decimal degrees
Latitude,
decimal degrees
Elevation,
m
Lithology
Sample
Table 3. AFT Dataa
Samples are analyzed with a Leica DMRM microscope with drawing tube located above a digitizing tablet and a Kinetek computer-controlled stage driven by the FTStage program [Dumitru, 1993].
Analysis is performed with reflected and transmitted light at 1250X magnification. Samples were irradiated at Oregon State University. Samples were etched in 5.5 molar nitric acid at 21C for 20 s. Following
irradiation, the mica external detectors were etched with 21C, 40% hydrofluoric acid for 45 min. The pooled age is reported for all samples as they pass the c2 test. Error is one sigma, calculated using the zeta
calibration method [Hurford and Green, 1983] with zeta of 373.2 ± 6.1 for apatite (B. Carrapa). Dpar, fission track etch pit measurements; SD, the related standard deviation.
b
Number of Xls is the number of individual crystals dated.
c
Rho-S and Rho-I are the spontaneous and induced track density measured, respectively (tracks/cm2).
d
NS and NI are the number of spontaneous and induced tracks counted, respectively.
e
P(c)2 is the chi-square probability [Galbraith, 1981; Green, 1981]. Values greater than 5% are considered to pass this test and represent a single population of ages.
f
Rho-D is the induced track density in external detector adjacent to CN5 dosimetry glass (tracks/cm2).
g
ND is the number of tracks counted in determining Rho-D.
CARRAPA ET AL.: OLIGOCENE UPLIFT IN THE CENTRAL ANDES
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TC4011
bedded sandstones, climbing ripples and minor interbedded
mud silt components we interpret this association as typical
of sheet flood and distal braided environments [Miall, 1978].
4.2. Chacras Formation
[21] Separated locally by an erosional unconformity, the
Quiñoas Formation is overlain by the 24.2 ± 0.9 to 22.5 ±
0.6 Ma Chacras Formation [Kraemer et al., 1999] that is
deposited to the west and east of the Sierra de Calalaste
(Figures 5 and 6). At its maximum recorded thickness, this
member reaches 650 m. At the southern end of Sierra de
Calalaste, the Chacras Formation is missing.
[22] The Chacras Formation is dominated by laterally
continuous, massive, stratified and imbricated conglomerates (Gc, Gh) and lenticular conglomerates (Gt/p). Clasts
within the conglomerates are angular to subangular, and
there are frequently boulders throughout. These conglomerates are interbedded with medium to coarse-grained crossbedded, and planar-bedded sandstones (St, Sp), and massive
sandstones (Sm) (Table 1). In the upper 150 m of log U
(Figure 6b) large-scale (5 – 10 m) cross sets of mediumgrained sandstones are preserved with foreset dips of 15–
25 that reach a combined thickness of up to 50 m.
[23] We interpret the lenticular conglomerates (Gt/p) as
resulting from deposition in shallow, gravely, bed load
channels, with planar stratification and imbrication (Gc,
Gh) occurring due to tractional flow at the channel bases
[e.g., Nemec and Steel, 1984]. Planar bedded conglomerates
also probably result from high-density flows during highdischarge events [Smith, 1987; Flint and Turner, 1988;
Adelmann, 2001] and variations in the amount of accumulation [Nemec and Steel, 1984]. The interbedded sandstones
are deposited from high-density flows during waning conditions [Rasmussen, 2000]. Such a close spatial relationship
between lenticular and planar conglomerates, and interbedded sandstones is typical of deposition through changing
flow regimes on a shallow gravel braided streams on
an alluvial fan [Miall, 1996]. The subangular clasts and
presence of boulders indicate a proximal source for
these deposits. The large-scale, cross-bedded sandstones
are interpreted as eolian dunes and exhibit geometries that
are typical of modern day examples [e.g., Hunter, 1977;
Reading, 1996], suggesting that at this time deposition
occurred in an arid environment. Paleocurrent data from
these sediments, on the east side of the present-day Sierra de
Calalaste, suggest a provenance mainly from the W-NW
(Figure 3).
4.3. Potrero Grande Formation
[24] The lower to middle Miocene Potrero Grande Formation unconformably overlies the Chacras Formation.
There are a series of interbedded tuffs that occur within
this formation, and the oldest tuff has a 40Ar/39Ar age of
18.0 ± 0.9 Ma [Kraemer et al., 1999]. A maximum
thickness of 300 m was estimated west of the Salar de
Antofalla [Adelmann, 2001]. Adjacent to the Sierra de
Calalaste, the Potrero Grande Formation consists of 150 m
of conglomerates, conglomeratic sandstones, and sandstones that are interpreted as having been deposited in an
14 of 19
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CARRAPA ET AL.: OLIGOCENE UPLIFT IN THE CENTRAL ANDES
alluvial fan environment [Kraemer et al., 1999]. Paleoflow
indicators within this formation show large variations in
direction with a scattered pattern (Figure 5) [Kraemer et al.,
1999]. Here, we focus on the early evolution of the Sierra de
Calalaste region. The Potrero Formation documents the later
evolution, and as such we do not document it in detail here
but refer the reader to Kraemer et al. [1999] for further
discussion.
5. Provenance
[25] Sandstone petrography for members I and II of the
Quiñoas Formation show a lithic-feldspatic to feldspaticlithic composition (Figure 7a) with an increase in quartz
fragments, shown in the QFL and QmFLt diagrams, an up
sequence typical of an evolution toward a more crystalline
source. The Qp, Lvm, Lsm diagram shows a generally
quartz rich composition for these sediments (Figure 7c).
[26] The only two samples available for the Chacras
Formation have a lithic-feldspatic composition suggesting
a different composition compared to the Quiñoas sediments
(Figures 7a and 7b). Despite the limited number of
samples available, the greater contribution from volcanic
and sedimentary-metamorphic rocks typical of the Sierra de
Calalaste range (Figure 7c) expressed by these two samples,
is consistent with paleocurrent data measured east of Sierra
de Calalaste, indicating an eastward direction and in turn a
source located in this range (Figure 3).
6. AFTT Data
[27] All analyzed fission track samples show comparable
results within one standard deviation, with ages ranging
between 24 ± 3 and 29 ± 1 Ma (Figure 8 and Table 3). Only
limited lengths were available in sample SCAD3 (Table 3),
giving a mean length of 12.98 ± 0.56 mm and suggesting
that partial annealing was not significant. Etch pit measurements indicate that the samples are monocompositional,
suggesting a homogeneous closure temperature. The ageelevation pattern of the Sierra de Calalaste vertical profile
points to an exhumation rate of 0.3mm/yr, which is
consistent with rates obtained in neighboring ranges to the
north [Deeken et al., 2004].
7. Discussion and Conclusions
[28] Our multiple data sets show that sedimentation
commenced during the late Eocene – early Oligocene with
the deposition of the Quiñoas sediments in a partially
segmented foreland basin with sediments derived mainly
from the west, and with contribution from proximal
sources. The general trend in the petrographical data for
the Quiñoas Formation indicates an increase in Qz up
sequence (Figure 7). This could be explained by an
evolution of the source toward more crystalline inputs.
Crystalline sources are generally typical of areas to the
west (Figure 7c). Also, the timing of the deposition of the
Quiñoas formation corresponds to the end of the exhuma-
TC4011
tion episode in the Domeyko Cordillera in northern Chile,
composed of granite and granodiorite, between 50 and
30 Ma [Maksaev and Zentilli, 2000]. This clear association
between the onset of exhumation of the Chilean Cordillera
and the deposition of the Quiñoas Formation, and the
sandstone provenance data would indicate that the dominant input to sedimentation was from the west. However,
the presence of coarse grained clastic sediments with more
variable paleocurrent directions in member I of the Quiñoas
Formation would suggest that the foreland basin system
seen during Quiñoas time was probably already partially
segmented by proximal highlands. These areas provided a
local, and minor sediment source, e.g., plutonic bodies
located immediately to the NW of the study area or even
part of the Sierra de Calalaste to the east (Figure 3)
[Kraemer et al., 1999]. Also, late Eocene – early Oligocene
apatite fission track cooling ages have been reported to the
east-southeast of the study area [Coutand et al., 2001]
introducing the possibility that there may have been some
topography to the east at this time that was responsible for
the compartmentalization of the foreland. In this respect
it is interesting that Horton et al. [2002] proposed that
the Eastern Cordillera of southern Bolivia was a source of
the Altiplano sediments during the Paleocene and OligoMiocene. The western flank of the Eastern Cordillera was
strongly deformed during the paleogene by west vergent
backthrusts [McQuarrie and DeCelles, 2001; McQuarrie,
2002b]. The Eastern Cordillera continues southward along
strike into the Puna plateau, and this range is a candidate
source for the investigated sediments (Figure 7).
[29] However, to date, no clear sedimentological and
thermochronological evidence exists that the Eastern Cordillera provided detritus to the west into the Puna at this time.
Furthermore, the metamorphic and volcanic lithic fragments
that would be characteristic of the Sierra de Calalaste and in
general of more eastern and southeastern sources [e.g.,
Reutter et al., 1994] are exceedingly scarce in the Quiñoas
sandstone, suggesting that such sources, if present, were not
contributing to sandstone detritus. Also, recent AFTT on
granitic rocks from eastern and northeastern areas in the
Puna suggest that the nearest ranges were exhuming later
than the deposition of the Quiñoas sediments by 20 Ma
[Deeken et al., 2004]. In light of the presently available data
we cannot rule out the possibility that the eastern boundary
of the Puna might have constituted some topographic
high already during the late Eocene –early Oligocene [e.g.,
Coutand et al., 2001; Horton et al., 2002].
[30] Interestingly, during the late Eocene, the local gypsum and anhydrite layers within the lower Quiñoas Formation (member I), and their association with evaporitic playa
mud flats, suggest that at least a seasonal arid environment
existed. This arid environment may have resulted from
deformation and uplift of the Chilean Cordillera [Maksaev
and Zentilli, 2000], which might have shielded the southern
Puna from occasional moisture incursions at a time when
the cold Humboldt current had not been fully established
and thus provided sufficient moisture to generate precipitation [Alpers and Brimhall, 1988]. In addition to regional
climatic effects, the prolonged aridity seen in the Calalaste
15 of 19
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CARRAPA ET AL.: OLIGOCENE UPLIFT IN THE CENTRAL ANDES
Figure 7. Main sandstone compositions of the investigated sediments: (a) QFL diagram based on the
technique described by Dickinson [1970]. Q, monocrystalline and polycrystalline quartz; F, feldspar;
L, sedimentary, metasedimentary, and volcanic lithic fragments including chert. Gray area represents the
compositional field of the late Eocene-Oligocene Potoco sedimentary rocks from the Altiplano sourced
from the Western Cordillera (sections 1– 2; after Horton et al. [2002]). (b) QmFLt diagram based on
the technique described by Graham et al. [1976]. Qm, monocrystalline quartz; F, feldspar;
Lt, sedimentary, metasedimentary, and volcanic, lithic fragments including polycrystalline quartz and
chert. (c) QpLvmLsm diagram based on the technique described by Graham et al. [1976]. Qp,
polycrystalline quartz; Lvm, volcanic lithics; Lsm, metamorphic lithics.
16 of 19
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CARRAPA ET AL.: OLIGOCENE UPLIFT IN THE CENTRAL ANDES
TC4011
Figure 8. AFT vertical profile and length histogram related to sample SCAD 3.
area may also have been further supported by the potential
topographic culminations of the Eastern Cordillera at least
by early Miocene time [Deeken et al., 2004].
[31] At a broader scale, time-equivalent clastic sedimentary rocks (Potoco Formation) deposited in the Bolivian
Altiplano and sourced from the Western Cordillera show
very similar facies associations in turn suggesting a similar
sedimentary environment [Horton et al., 2002]. It is
important to acknowledge however, that the Altiplano
provenance data were obtained from localities hundreds
of kilometers to the north of the study area. In any case,
similarities between these units may indicate that an
extensive longitudinally oriented basin, with local highlands, existed during the late Eocene – early Oligocene that
spanned the length of the central Andes. In such a
scenario, the Quiñoas Formation may represent a part of
a semicontinuous foreland basin located to the east of the
Eocene Incaic mountain belt.
[32] The overlying upper Oligocene – lower Miocene
Chacras Formation shows a change in sediment dispersal
and facies association, with a marked coarsening compared
to the Quiñoas Formation, and paleocurrent and petrographic
data that suggest input from the Sierra de Calalaste. Despite
the fact that paleocurrent measurements within the Chacras
Formation are only available along the eastern margin of the
range, these data unequivocally indicate a source from the
Sierra de Calalaste.
[33] AFTT data show that exhumation of the range
occurred between 24 and 29 Ma, which is the time of
deposition of the late Quiñoas and early Chacras formations
(28.9 ± 0.8 Ma; 24.2 ± 0.9 [Kraemer et al., 1999]).
Therefore we propose that the observed change in sediment
dispersal is a direct response to the uplift and erosion of the
Sierra de Calalaste range, which must have had a profound
effect on the fluvial systems within adjacent basins during
the Oligocene. Finally, the facies association with the
presence of extensive eolian dunes within this formation
suggests that the arid climate established prior to or during
Quiñoas time persisted during Chacras time.
[34] In Sierra de Calalaste and adjacent basins, uplift
resulting from Oligocene deformation led to reorganization
of the depositional systems within the present-day Salar de
Antofalla area. Paleocurrent and sandstone provenance data
show that the depositional system, originally mainly sourced
from western crystalline rocks, was reorganized during the
Oligocene. This reorganization was contemporaneous with
deformation within the Sierra de Calalaste, suggesting a
causal linkage between uplift of the range and response of
the adjacent basin. At least transient internally drained
conditions existed during depositions of the Quiñoas sediments though it is less clear if such conditions were present
during the deposition of the Chacras sediments. However,
the occurrence of thick evaporite units in the late Miocene
(Juncalito Formation [Kraemer et al., 1999]) indicates that
the Salar de Antofalla basin was internally drained by that
time. Therefore we suggest that the basin reorganization
seen between the Quiñoas and Chacras formations heralded
the beginning of the process of disruption of the regional
fluvial system that may have ultimately led to the formation
of internal drainage.
[35] Many workers have documented deformation possibly driving range uplift and a transition from external
to internal drainage within basins of the Puna plateau in
Oligo-Miocene time [e.g., Alonso, 1986; Jordan and
Alonso, 1987; Marrett, 1990; Alonso et al., 1991; Coira et
al., 1993; Vandervoort, 1993]. Around 23S latitude, the
Salinas Grandes and Tres Cruces basins contain evidence of
deformation beginning in the late Eocene to early Oligocene
[Coutand et al., 2001]. At 24S latitude, internal drainage at
Siete Curvas may have formed as early as the late Oligocene
and certainly by late Miocene time [Vandervoort et al.,
1995], whereas to the west internal drainage within the
Arizaro and Tolar Grande basins commenced no later than
early Miocene time [Donato, 1987; Coutand et al., 2001],
17 of 19
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CARRAPA ET AL.: OLIGOCENE UPLIFT IN THE CENTRAL ANDES
and early to middle Miocene [Vandervoort et al., 1995],
respectively. Within the Salar de Pastos Grandes to the west
of Siete Curvas, thick evaporites show that internal drainage
formed sometime between 11.2 Ma and perhaps as early as
the late Eocene – early Oligocene [Alonso et al., 1991].
Within the Salar de Hombre Muerto area at 25S latitude,
evaporite deposition started by 15.0 ± 1.2 Ma, and internal
drainage may have been established as early as the Oligocene [Alonso et al., 1991; Vandervoort et al., 1995]. In the
same region the termination of supergene copper mineralization indicates that hyperaridity was established after
14.7 Ma [Alpers and Brimhall, 1988]. Finally, in the Campo
Arenal area along the present Puna margin at 27S latitude,
AFTT data indicate that deformation had begun between 29
and 38 Ma [Coutand et al., 2001].
[36] In summary, these observations suggest that deformation, exhumation of basement ranges, and the establishment of internal drainage within the Puna plateau were
spatially diachronous [e.g., Vandervoort, 1993; Coutand et
al., 2001]. In particular, the timing of the onset of internal
drainage in the area may be dependent on the details of the
uplift of discrete mountain ranges that may be controlled by
local structural or volcanic conditions [Segerstrom and
Turner, 1972; Alonso, 1986]. Combined with our new data
TC4011
set, this suggests that deformation driving uplift within the
Puna may not have occurred progressively from west-toeast as previously suggested [e.g., Andriessen and Reutter,
1994] and that the establishment of internal drainage may
not only have been dependent on local details of marginal
range uplift but also on the history of the uplift of ranges
farther upwind.
[37] At a more regional scale, our new data show that
deformation and uplift of the southern Puna plateau started
already in Oligocene time, if not earlier, and contributed to
the fragmentation and paleodrainage reorganization of an
earlier semicontinuous foreland basin. This event occurred
in an already arid climate environment, and created the ideal
morphotectonic preconditions for the establishment of the
subsequent internal drainage environment that was in place
no later than middle Miocene time.
[38] Acknowledgments. Deutsche Forschungsgemeinschaft (DFG)
and the Alexander von Humboldt Foundation are kindly acknowledged
for financial support through grants to M. Strecker and B. Carrapa,
respectively. Peter DeCelles, Teresa Jordan, Brian Horton, and an anonymous reviewer are kindly thanked for constructive reviews of this manuscript. Also, we thank R. Marrett and R. Alonso for their help during
sample collection and logistics.
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D.
Institut für Geowissenschaften, FrieAdelmann,
drich-Schiller-Universität Jena, Burgweg 11, D-07749
Jena, Germany.
B. Carrapa, E. Mortimer, E. R. Sobel, and M. R.
Strecker, Institut für Geowissenschaften, Universität
Potsdam, Karl-Liebknecht-Str. 24, D-14476 Potsdam,
Germany. (carrapa@geo.uni-potsdam.de)
G. E. Hilley, Department of Earth and Planetary
Science, University of California, 377 McCone Hall,
Berkeley, CA 94720-4767, USA.