TECTONICS, VOL. 24, TC4020, doi:10.1029/2004TC001675, 2005
Plateau-style accumulation of deformation: Southern Altiplano
Kirsten Elger, Onno Oncken, and Johannes Glodny
GeoForschungsZentrum Potsdam, Potsdam, Germany
Received 5 May 2004; revised 17 December 2004; accepted 23 March 2005; published 31 August 2005.
[1] Employing surface mapping of syntectonic
sediments, interpretation of industry reflectionseismic profiles, gravity data, and isotopic age dating,
we reconstruct the tectonic evolution of the southern
Altiplano (20 – 22S) between the cordilleras
defining its margins. The southern Altiplano crust
was deformed between the late Oligocene and the late
Miocene with two main shortening stages in the
Oligocene (33– 27 Ma) and middle/late Miocene
(19 – 8 Ma) that succeeded Eocene onset of
shortening at the protoplateau margins. Shortening
rates in the southern Altiplano ranged between 1 and
4.7 mm/yr with maximum rates in the late Miocene.
Summing rates for the southern Altiplano and the
Eastern Cordillera, we observe an increase from
Eocene times to the late Oligocene to some 8 mm/yr,
followed by fluctuation around this value during the
Miocene prior to shutoff of deformation at 7–8 Ma and
transfer of active shortening to the sub-Andean fold
and thrust belt. Shortening inverted early Tertiary
graben and half graben systems and was partitioned
in three fault systems in the western, central, and
eastern Altiplano, respectively. The east vergent fault
systems of the western and central Altiplano were
synchronously active with the west vergent Altiplano
west flank fault system. From these data and from
section balancing, we infer a kinematically linked
western Altiplano thrust belt that accumulated a
minimum of 65 km shortening. Evolution of this belt
contrasts with the Eastern Cordillera, which reached
peak shortening rates (8 mm/yr) in between the above
two stages. Hence local shortening rates fluctuated
across the plateau superimposed on a general trend of
increasing bulk rate with no trend of lateral
propagation. This observation is repeated at the
shorter length and time scales of individual growth
structures that show evidence for periods of enhanced
local rates at a timescale of 1–3 Myr. We interpret
this irregular pattern of deformation to reflect a
plateau-style of shortening related to a self-organized
state of a weak crust in the central South American back
arc with a fault network that fluctuated around the
critical state of mechanical failure. Tuning of this state
may have occurred by changes in plate kinematics,
Copyright 2005 by the American Geophysical Union.
0278-7407/05/2004TC001675$12.00
during the Paleogene, initially reactivating crustal
weak zones and by thermal weakening of the crust
with active magmatism mainly in the Neogene stage.
Citation: Elger, K., O. Oncken, and J. Glodny (2005), Plateaustyle accumulation of deformation: Southern Altiplano, Tectonics,
24, TC4020, doi:10.1029/2004TC001675.
1. Introduction
[2] Although considerable advance has been made in
recent years in understanding the processes involved in
the formation of orogenic plateaus, the precise temporal
and spatial patterns of uplift and lateral progradation of
deformation in plateaus beyond the last few million years is
still shrouded in darkness. This situation is largely due to a
lack of systematic data with the potential to date the timing
of deformation and uplift, as well as of detailed structural
data that reveal the internal architecture of the world’s
large orogenic plateaus, Tibet and Altiplano [e.g., Dewey
et al., 1988; Isacks, 1988; England and Molnar, 1997;
Allmendinger et al., 1997; Shen et al., 2001]. Young,
syntectonic and posttectonic sediments largely cover both
plateaus and restrict access to the subsurface. Yet, models of
large-scale orogeny and plateau formation heavily rely on
assumptions that relate to the early stages of plateau
evolution and the conditions governing these. This is
particularly true for the Andean Altiplano, which, due to
its setting in an ocean-continent convergence system, has
been called a plate tectonic paradox [e.g., Allmendinger et
al., 1997]. Because of the unsatisfactory data situation,
published ideas on accumulation of deformation in the
Altiplano range from various inferred stages in plateau
evolution [e.g., James, 1971; Pardo-Casas and Molnar,
1987; Isacks, 1988; Sempere et al., 1990; Allmendinger
and Gubbels, 1996; Kley, 1996; Allmendinger et al., 1997;
Lamb and Hoke, 1997; Lamb et al., 1997; Lamb, 2000;
Hindle and Kley, 2002] to suggestions of continuous eastward progradation of deformation, starting at the Western
Cordillera, with a propagation style much like that observed
in classic foreland fold and thrust belts [e.g., Horton et al.,
2001; McQuarrie, 2002]. To date, the available data in the
Altiplano do not provide a complete quantitative record of
the spatial and temporal pattern of strain accumulation
during plateau formation that also sufficiently resolves the
earlier stages.
[ 3 ] In the present paper we attempt to provide a
first nearly complete record of magnitude and timing of
shortening across the southern Altiplano plateau (compare
Figure 1) by the mapping of syntectonic sediments, by
identifying structures, their relative ages and magnitudes of
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Figure 1
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deformation from the interpretation of industry reflectionseismic profiles, and by isotopic age dating of deformation,
including a reevaluation of published ages. Our analysis
focuses on the southern Altiplano at 20– 22S. This segment
is characterized by the highest density of geophysical and
geological data in the central Andes, all provided by a
variety of recent projects [e.g., Wigger et al., 1994; Beck
et al., 1996; Swenson et al., 2000; Yuan et al., 2000;
Haberland and Rietbrock, 2001; Brasse et al., 2002; Götze
and Krause, 2002; ANCORP Working Group, 2003]. In
addition, high-resolution industry reflection seismic sections
covering the entire plateau were made available to this
project by the Bolivian national oil industry (courtesy of
YPFB).
2. Geological Framework
[4] In spite of ongoing subduction for more than 200 Myr,
large-scale deformation of the upper plate and formation of
the Andes high plateau only commenced during early to
mid-Tertiary times [Sempere et al., 1990; Lamb and Hoke,
1997; Allmendinger et al., 1997; Horton et al., 2001]. The
resulting central Andean mountain belt is built from a series
of morphotectonic units (Figure 1d): the Longitudinal
Valley in the forearc with an average elevation of some
1000 m and thin sedimentary infill; the 4500 – 5500 m high
Chilean Precordillera with Paleozoic and Mesozoic basement, overlain by a thin Neogene cover; the Western
Cordillera built by the Neogene to recent magmatic arc
with peak elevations of up to 6000 m; the flat Altiplano
Plateau at some 3800 m average elevation forming a
200-km-broad intramontane basin; the Eastern Cordillera
reaching 5000 –6000 m altitude, a doubly vergent thickskinned thrust system; the inter-Andean and sub-Andean
zones representing an east vergent, ‘‘classical’’ thin-skinned
foreland fold-thrust belt that accumulated most of the back
arc shortening; and finally the down flexed, but otherwise
undeformed Chaco at some 500 m elevation, a 200-kmwide foreland basin with a wedge-shaped Neogene clastic
fill of up to 5 km thickness.
[5] Recent work has shown that deformation and uplift of
the western flank of the Plateau and Precordillera occurred
between the late Oligocene to late Miocene at very low,
mildly fluctuating strain rates [Victor et al., 2004] following
an Eocene/?early Oligocene precursor shortening stage
[Haschke and Guenther, 2003]. In contrast, the Eastern
Cordillera and inter-Andean belt accumulated substantially
more shortening starting in the late Eocene, with most
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shortening during the early to middle Miocene [e.g., Kley,
1996; Allmendinger et al., 1997; Lamb and Hoke, 1997;
Müller et al., 2002]. The adjoining sub-Andean fold-thrust
belt accumulated most shortening since the Miocene [Baby
et al., 1997; Kley and Monaldi, 1998] and still shows active
growth at its eastward deformation front as evidenced from
active seismicity and GPS data [Lamb, 2000; Hindle and
Kley, 2002; Bevis et al., 1999].
[6] In contrast to the plateau flanks, the southern Altiplano evolution is only recorded by few structural and age
data [e.g., Baby et al., 1997; Rochat et al., 1999; Welsink et
al., 1995] (see Figure 1a). Most of the area is covered by
Quaternary fluvial and lacustrine sediments, including several large saltpans (e.g., Salar de Uyuni, Figure 1), and by
posttectonic volcanic rocks from the active magmatic arc of
the Western Cordillera. Most of the exposed pre-Cenozoic
basement is made up of mainly unmetamorphic and mildly
folded Ordovician to Devonian clastic sediments while the
western margin of the plateau is built of deformed Late
Paleozoic and Mesozoic igneous and sedimentary rocks. In
southern Bolivia, Cretaceous sediments unconformably
overlie this basement. These sediments were partly deposited in a shallow marine environment [Fiedler et al., 2003]
and mark the last marine ingression prior to plateau formation. This postrift sequence is related to the Potosi basin
[Fiedler et al., 2003], a branch of a major rift system in the
south (‘‘Salta Rift’’), the site of which largely matches the
extent of the later Eastern Cordillera.
[7] The establishment of internal drainage early in the
Andean evolution with two cordilleras confining the protoplateau area was responsible for storing sediment that
reflects the entire deformation history (compare Vilque well
[Welsink et al., 1995]; see Figure 1a). In the southern
Altiplano, the late Paleocene Santa Lucia formation and
the Eocene/Oligocene Potoco formation [Jordan and
Alonso, 1987; Sempere et al., 1997; Horton et al., 2001],
a series of coarse to fine-grained clastics and red beds
conformably overlie the Cretaceous/Paleocene deposits
(Figure 2). Rochat et al. [1999] and Jordan et al. [2001]
suggested that these sediments were partly deposited in
local extensional basins. This sequence is in turn overlain
by the Miocene continental clastic San Vicente Formation
[Martı́nez et al., 1994; Mertmann et al., 2003]. To the east
this formation is delimited by the San Vicente Thrust zone
(Figure 1a) marking the western topographic limit of the
Eastern Cordillera. The San Vicente Formation is characterized by significant lateral variations in thickness, grain
size, and detrital composition as well as by unconformable
Figure 1. (a) Geological sketch map of the southern Altiplano (section of map of central Andes by Reutter et al. [1994]
including our own mapping results) superimposed on shaded relief of the topography showing the distribution of pre-late
Neogene rocks, lower-hemisphere equal-area stereonet projections with bedding and regional fold axes, and the location of
reflection seismic lines. (b) Bouguer anomaly map (provided by H. J. Götze; see Götze and Krause [2002] for description
of gravity field) of area shown in Figure 1a depicting positive anomalies closely related to faults and folds mapped at the
surface and observed in the seismic sections. Note the chain of anomalies in the nearly completely sediment-covered
western Altiplano with similar style as obvious from the exposed central Altiplano unit. (c) Location of the working area.
(d) Major morpho-tectonic units of the southern central Andes. See color version of this figure at back of this issue.
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and age with those identified by Lamb and Hoke [1997]
farther north in the Altiplano.
3. Plateau Architecture
3.1. Structural Style
Figure 2. Schematic stratigraphic columns of the
(a) central and (b) eastern parts of the southern Altiplano
with thickness as mapped in the field, in the Vilque well
[Welsink et al., 1995], and in seismic sections (Foldout 1).
Age data collected in this study and by Silva-González
[2004] are shown at respective sample location in the
stratigraphic column. Superscript 1 indicates that ages of
base of Potoco and Molino Formations are from Sempere et
al. [1997, and references therein]; asterisks indicate that
ages are from Silva-González [2004].
onlaps on the Paleozoic basement. Isotopic ages from
intercalated volcanic rocks of the southern Altiplano range
between some 27.4 and 5 Ma, interpreted to bracket the
timing of deposition of the San Vicente Formation and the
supposedly contemporaneous deformation [Grant et al.,
1979; Sempere et al., 1990; Soler et al., 1993; Lamb and
Hoke, 1997; Horton et al., 2002]. Late Miocene to recent
volcano-sedimentary deposits unconformably overlie all
preexisting units. The above sequences, summarized in
Figure 2, show significant similarities in depositional style
[8] In outcrop, major folds often associated with thrust
faults characterize the southern Altiplano structural style.
The folds have subvertical axial planes, a parallel geometry,
and are nearly cylindrical with NNE-SSW oriented fold axes
plunging less than 5 north or south (see Figure 1a).
Bedding surfaces on the fold limbs have well developed
slickensides usually oriented downdip (flexural-slip folds).
Sediments at the limbs of folds often exhibit growth
structures. Apart from folding, most of the shortening is
recorded by thrusts rooting below the anticlines. Strike-slip
faults are of very minor importance. Major extensional
structures are not observed on the southern Altiplano,
except for rare small-scale faults, which usually show
displacements of no more than a few decimeters to meters.
[9] At surface, the moderately eastward dipping thrust
system of the San Vicente Fault Zone (SVFZ) separates the
Eastern Cordillera in the east from the sediment-covered
Altiplano in the west (Figure 1a). In the plateau area to the
west, we subdivide the southern Altiplano into three tectonostratigraphic domains based on its internal architecture: the
eastern, the central, and the western southern Altiplano
domain. These are also particularly well imaged by the
high-resolution Bouguer map (Figure 1b). In the areas
exposing deformed Tertiary rocks positive Bouguer anomalies perfectly trace anticlines and fault hanging walls. Hence
the Bouguer map can be seen to highlight the structural trend
of folds and faults beneath the younger cover.
[10] The eastern domain, which also includes the Tertiary/
Quaternary Lı́pez Basin, is the buried tip of the Eastern
Cordillera’s west vergent, thin-skinned western margin with
widely spaced open anticlines. The central Altiplano domain is formed by a thrust system separated from the
Eastern Cordillera by the undeformed Lı́pez basin: The
Khenayani-Uyuni Fault Zone (KUFZ) represents its key
part and is the most important structural element of the
southern Altiplano [Sempere et al., 1990; Martı́nez et al.,
1994; Welsink et al., 1995; Rochat et al., 1999]. It represents
an east vergent leading imbricate fan of up to four major,
N-S to NNE-SSW trending, emergent thrusts (dips are 20–
30 at frontal thrust to subvertical at westernmost thrust).
Along the thrusts, Paleozoic to Oligocene sediments were
displaced eastward over the Cenozoic fill of the Lı́pez Basin
(Figure 3). The hanging walls show folded and faulted
Paleozoic sedimentary basement with remnant Late Mesozoic, and a thinned to incomplete Cenozoic cover
(Figure 3a). The footwalls of the individual thrusts are
characterized by thrust front synclines containing growth
structures with steep western limbs and up to 2 km long,
gently dipping eastern limbs. Because thrust hanging wall
cutoffs are eroded, we assess displacement from downplunge projection of cutoff geometries mapped in the
southern parts of the area (Figure 1a). The western parts
of the central Altiplano system is built from west vergent
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the identical gravity signature as compared to the exposed
central Altiplano, appears to have a similar internal architecture. Its western part is also exposed farther north (18 –
19S) in the Western Cordillera [Charrier et al., 2002;
Garcı́a et al., 2002].
3.2. Reflection Seismic Data
Figure 3. Structures from Khenayani-Uyuni Fault Zone
system. (a) Unconformity between San Vicente Formation
(Pilkhaua Subsequence?) and Silurian deposits in hanging
wall of Khenayani-Uyuni Fault. (b) Steeply inclined San
Cristóbal Fault with subvertical hanging wall structure
(Potoco Formation) intruded by San Cristóbal Volcano. (c)
Growth strata in footwall of San Cristóbal Fault (Pilkhaua
Subsequence).
fault propagation folds with 6 – 10 km wavelength overlying
blind thrusts. The folds have associated small thrust top
basins filled with wedge-shaped, onlapping deposits. This
entire central Altiplano system is continued to the north in
the exposed Tambo-Tambillo area (Figure 1a) [cf. Lamb and
Hoke, 1997].
[11] The largely sediment-covered western Altiplano domain only exposes deformed rocks at the Yazón Peninsula
in the southern Salar de Uyuni (21S, 68 W), where
folded strata of the San Vicente Formation crop out
(Figure 4). The structure of the Yazón Peninsula is a west
vergent, NNE-SSW trending fold with a long eastern back
limb, a steep west facing forelimb and an associated thrust
front basin filled with growth strata. In spite of the isolated
nature of this exposure, the Bouguer map (Figure 1d) shows
that the Yazón structure is part of a major NNE-SSW
trending system of folds and/or faulted and uplifted rocks
under the volcanic and Salar cover (Figure 4). This nearly
unknown structural unit in the western Altiplano significantly contributes to its deformation architecture and, from
[12] Surface observations are complemented by industry
reflection data that image the crust down to some 4 s twoway time (TWT) depth. The definition of seismic facies and
their correlation with mappable stratigraphic formations and
well data mainly is shown in Figure 9 in the auxiliary
material1.
[13] Seismic lines 10007 and 10010 (Foldout 1) illustrate
the lateral (north-south) variability of the Lı́pez Basin and
deformation front of the Eastern Cordillera. Toward depth
(line 10007 in Foldout 1), the Potoco Formation shows a
westward thinning wedge with conformable top lap reflector geometries at both the upper and lower sequence
boundaries. The overlying San Vicente Formation is built
from three subsequences, defined here as the OMsv1,
OMsv2 (Oligo-Miocene San Vicente Formation), and
Pilkhaua Subsequence. The OMsv1 and OMsv2 Subsequences are characterized by relative uniform thickness
across the entire section. The Pilkhaua Subsequence is
absent in line 10007, most probably due to erosion (compare, however, its occurrence on the west flank of Vilque
Anticline in line 10010). In the syncline east of the Vilque
Well small normal faults are local features, related to the
collapse of the forelimb of the eastern Anticline above
evaporite layers within the OMsv2 Subsequence (600 m
of evaporites drilled in the Vilque Well). The most prominent unconformity of line 10007 is the truncation at the
base of the OMsv1 Subsequence near the deformation front
(DF, see Foldout 1e, 1f) above eroded top lap terminations
of the Potoco and El Molino Formations. A second structural disconformity in the western part of the Vilque
Anticline between the OMsv1 and OMsv2 Subsequences
(Foldout 1e) is related to the blind tip of the hanging wall
ramp of the Vilque fault bend fold.
[14] At its eastern border, the southern part of the Lı́pez
basin (line 10010, Foldout 1f) exhibits thinning and progressive onlaps of reflectors of the Potoco Formation and
OMsv1 Subsequence against a basement high. As in line
10007, the unconformity at the base of the OMsv1 Subsequence overlies well-developed top lap terminations from
the Potoco Formation With the onset of deposition of the
OMsv2 Subsequence the downlapping reflectors in the
western part of the basin indicate a westward prograding
system of foresets. Locally developed top lap terminations
below the eastern OMsv2 Subsequence show that this
sequence boundary was also partly erosive (Foldout 1e, 1f).
[15] The most obvious difference between the northern
section (10007) and the southern section (10010, compare
Foldout 1) is the smaller basin width in the southern section.
A second difference between lines 10007 and 10010 is a
1
Auxiliary material is available at ftp://ftp.agu.org/apend/tc/
2004TC001675.
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Figure 4. Exposed part of western Altiplano thrust system (see Figure 1a for location, Yazón
Peninsula). (a) Satellite image. (b) Geological map with lower-hemisphere equal-area stereonet projection
showing regional fold axis and bedding. (c) Cross section with local age data. (d) SW-ward view of roof
and frontal limb of Yazón Anticline with onlapping growth sediments of Pilkhaua Subsequence.
southward decrease of thrust displacement related to formation of the Vilque Anticline, a classical fault-bend
anticline, in the north. The most significant observation of
Foldout 1, however, is the strong similarity between both
sections in terms of the characteristics of seismic facies,
similar positions of unconformities, and the structural style
allowing correlation between lines.
[16] In the central Altiplano domain (lines 10023 and
2593, see Foldout 1) surface outcrops at sites of poor quality
of the seismic image correlate with steeply inclined strata or
with footwall units buried beneath deformed imbricates.
The strongly folded and imbricated sequences of the central
Altiplano can be observed in nearly continuous outcrop 5–
10 km south of the seismic section (Figure 1a). Line 10023
shows an east vergent imbricate fan building the eastern part
of the system; west vergent fault bend folds dominate the
western part. The analysis of seismic facies and their
correlation with surface geology exhibits the same, albeit
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Foldout 1. Cross section of the southern Altiplano based on seismic lines 2553 (stacked section), 10023, 2593, and 10007
(migrated sections), courtesy of YPFB (see Figure 1 for location) and their time-depth conversion using average interval
velocities of the lines as shown in inset (bold line, average; thin lines, variation along line). Western Altiplano profile is projected
into the common section. Structures are shown by name or numbers (eastern Altiplano) with nearby isotopic ages projected
along strike. Restored section shows late Oligocene stage prior to onset of sedimentation of San Vicente Formation Details show
(a) western deformation front of western Altiplano thrust system; (b) wedge-shaped westward thinning of Potoco Formation and
youngest fold of central Altiplano thrust system folding the onlapping Pilkhaua Subsequence; (c) fault-propagation folds of San
Vicente Formation with piggyback basins filled with Pilkhaua Subsequence that onlap neighboring folds; (d) thrust units
bounded by Corregidores (right) and San Cristóbal Faults (left) with Potoco beds truncated by San Vicente Formation and
growth strata of the Pilkhaua Subsequence; (e) buried western deformation front of the Eastern Cordilleran Thrust System
truncated by the base of San Vicente Formation (OMsv1); (f) western deformation front of the Eastern Cordilleran Thrust System
along the southern line 10010 (see Figure 1 for location) with truncation of older structures by the basal San Vicente Formation.
The latter shows internal truncations in the east (between OMsv1 and OMsv2), westward prograding sedimentary foresets in the
west (OMsv2), and growth structures at the top of the sequence (Pilkhaua Subsequence). Numbers indicate sequence of fault
activation. (See enlarged version of this figure in the HTML.)
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Figure 5. (e) Surface outcrop and (a and b) interpreted part of seismic line 10023 (see Foldout 1 for
location of detail) with backlimb of Allka Orkho fold structure (in the west) and Ines fold (in the east) and
intervening basin filled with Pilkhaua Subsequence. The basin fill shows stacked generations of
onlapping reflector terminations onto both folds and decreasing intensity of folding upsection in
discontinuous steps marked as (c and d) individual sequences (see text for details; PT, pretectonic; ST,
syntectonic; PST, posttectonic after ST1 and erosion of Allka Orkho fold). Bars in San Vicente Formation
in Figure 5a) indicate reflector sequences bounded by continuous high-amplitude reflections.
more deformed and dissected, seismic sequences as detected
in the eastern Altiplano (Foldout 1 and Figure 9 in the
auxiliary material). This includes the prominent double
reflector of the El Molino Formation, the seismically
transparent facies of the Potoco Formation, and the continuous high-amplitude reflectors of the San Vicente Formation In the central Altiplano, subdivision of the San Vicente
Formation into the OMsv1 and OMsv2 Subsequences is not
possible any more. The overlying Pilkhaua Subsequence
is most distinct as infill of local thrust top basins (e.g.,
Foldout 1c and Figure 5).
[17] The Potoco Formation forms a continuously eastward thickening wedge that reaches its maximum thickness
in the hanging wall of the San Cristóbal Fault and again
exhibits conformable sequence boundaries. Its thickness is
significantly smaller east of the San Cristóbal Fault. From
seismic analysis, map interpretation, and field observations,
the lower sequence boundary of the overlying San Vicente
Formation evolves from conformable in the footwall of the
individual faults to erosive in their hanging wall positions
(top lap terminations of the Potoco reflectors, see arrow in
Foldout 1d; see also Figure 3c).
[18] The main body of the San Vicente Formation has a
more constant thickness, but also thins westward (line
10023 in Foldout 1) and is very thin on the KhenayaniUyuni Fault Zone as compared to the immediately adjacent
areas. Contacts with underlying strata are conformable in
most cases except for the above erosional unconformity at
in hanging wall positions (Corregidores Fault, KhenayaniUyuni Fault, and Allka Orkho Fault). In the uppermost San
Vicente Formation, thrust top basins filled with onlapping
deposits show wedge-shaped sedimentary bodies defining
the Pilkhaua Subsequence. These Pilkhaua-filled basins
generally do not extend below 0.5 –0.8 s TWT depth. West
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of the westernmost anticline, the seismic profile was acquired over undeformed lacustrine sediments south of the
Salar de Uyuni (see Foldout 1) that cover eroded top laps
of the Pilkhaua Subsequence at flanks of major folds
(Foldout 1b).
[19] In the western Altiplano, the structural architecture
was only accessible at the Yazón peninsula and in a seismic
stack section of poor quality crossing the Salar de Uyuni
(example and tentative interpretation of the line drawing of
line 2553 is shown in Foldout 1a). Interestingly, structural
analysis exhibits a doubly vergent imbricate fan and fault
bend fold system quite similar to the central Altiplano
structure. In addition, this unit shows a similar style of
Bouguer anomalies as the central Altiplano (Figure 1d).
Judging from reflection styles, there is restricted evidence
for the presence of the El Molino and Potoco Formations
(see Figure 9 in the auxiliary material) or they are indistinguishable from the San Vicente Formation In contrast, the
Pilkhaua Subsequence is clearly identified by its reflector
architecture. Both San Vicente Subsequences are also observed at surface (Yazón Peninsula) and show the same
onlapping relationship as in the central and eastern Altiplano domains (Figure 4). All units are covered and onlapped
by younger sediments and volcanics of the Salar de Uyuni
basin.
3.3. Section Balancing
[20] Using surface data and the depth-converted interpreted line drawing (depth conversion based on averaged
interval velocities of the seismic lines, see inset in Foldout 1),
we iteratively constructed and restored a balanced section
with the software packages GEOSEC1 (Paradigm) and
2D-MOVE1 (Midland Valley Exploration Ltd.) using
flexural flow as the dominant deformation mechanism as
indicated from field observations. We determined fault
geometry at depth from seismic interpretation and depthto-detachment calculations using the fold geometries of
various fault-related folds [cf. Epard and Groshong, 1993].
[21] The eastern Altiplano domain is shown to be part of
a west vergent thin-skinned fold-and-thrust belt with widely
spaced ramp anticlines and open fault-bend folds (Foldout 1,
line 10007). Here, most shortening was accumulated prior
to deposition of the San Vicente Formation in the western
part of the section (4.5 km shortening). All thrusts join a
gently eastward dipping detachment, which is characterized
by a double flat ramp geometry, descending eastward in two
steps from six to nine kilometers depth within lower
Paleozoic sediments and probably linking with the San
Vicente fault. Minor reactivation (1.5 km shortening)
during a later stage (post-OMsv2 deposition) was responsible for the formation of two wide ramp anticlines east of the
earlier deformation front. This interpretation is in contrast to
the extensional model of Welsink et al. [1995] but is similar
to that of Baby et al. [1997] and Müller et al. [2002]
showing, however, significantly more detail.
[22] In the central Altiplano domain, the KhenayaniUyuni Fault Zone (KUFZ) forming the eastern part of this
doubly vergent thrust system is built from four westward
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dipping thrusts (KUF, CF, SCF, IPF, see Foldout 1) with
steep to moderately dipping hanging wall strata underlain
by a gently west dipping detachment (Foldout 1, line
10023). In comparison to earlier interpretations [e.g., Baby
et al., 1997; Lamb and Hoke, 1997; Rochat et al., 1999;
McQuarrie and DeCelles, 2001], this geometry, as deducted
from surface as well as from seismic data (compare
Foldout 1), accommodates significantly more shortening.
Large parts of the hanging walls were eroded during
thrusting. Eroded hanging wall cutoffs were therefore
reconstructed from downplunge projection of the map.
The less deformed, west vergent part of the central Altiplano Fault System exhibits four fault propagation folds above
eastward dipping, mostly blind, thrusts (Foldout 1, line
10023). The folds have associated syntectonic basins filled
with sediments of the Pilkhaua Subsequence. Initial shortening prior to deposition of the San Vicente Formation, as
evidenced from its erosive basis, only affected the external
parts of the doubly vergent thrust system, namely the
Khenayani-Uyuni and the Corregidores imbricates in the
east and the Allka Orkho Fault in the west. The eroded
hanging wall cutoffs were reconstructed from extrapolating
stratigraphic thickness from the footwall to the hanging
wall. Where thickness changes are evidenced, like the
reduced Potoco Formation thickness in the KUFZ area,
we used erosion estimates derived from fission track modeling. Ege [2004] calculated some 0.9– 2.4 km of erosion
for the Potoco Formation prior to the onset of San Vicente
deposition in this area. Restoration of the central Altiplano
revealed a total of 40 km horizontal contraction (13 km of
which were accumulated in the early stage) with a predominance of east directed thrusting (see Foldout 1).
[23] The main body of the San Vicente Formation below
the Pilkhaua Subsequence exhibits a uniform thickness
unaffected by the west vergent structures. In contrast, in
the dominantly east verging thrusts, the same formation is
very thin or absent, or represented only by Pilkhaua-style
deposits (compare Foldout 1, line 10023). This fact requires
either continuous or repeated activity of the KUFZ, in order
to maintain local relief, or very high relief formed during the
initial stage that was not downgraded before the late stage of
deposition of the San Vicente Formation In both cases, local
relief with respect to neighboring domains persisted until
the sedimentary surface in these reached the surface elevation of the eroded KUFZ area and then started to transgress
the latter. In summary, deformation in the central Altiplano
locally started prior to deposition of the San Vicente
Formation and continued throughout San Vicente deposition in the KUFZ area, while all of the central Altiplano
structures were active during deposition of the Pilkhaua
Subsequence.
[24] For the western Altiplano thrust system, we only
aimed at a semiquantitative model interpreting mainly the
seismic section based on reflection styles and geometries.
The doubly vergent style of deformation, again with a
predominance of east directed thrusting, is similar to that
of the central Altiplano system. However, for lack of
resolution of structures, the measured shortening of 21 km
(see Foldout 1) is a minimum estimate only. A separation of
9 of 19
ELGER ET AL.: ALTIPLANO DEFORMATION
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shortening into two stages is not detectable due to the poor
quality of the seismic line and lack of exposure.
[25] Balancing shows that all thrusts in the central and
western Altiplano thrust systems merge into a basal detachment, which lies at depths between 9 and 12 km in the
Paleozoic sediments (compare eastern Altiplano) and dips
1 – 2 westward (Foldout 1). Apart from the dominant east
directed thrusting, indications for this detachment geometry
were derived from the following two observations:
[26] 1. The Late Cretaceous El Molino Formation forms a
subhorizontal regional near sea level in the east rising to
1500 m elevation in the western Altiplano. Exceptions to
this are the strongly deformed external parts of the central
Altiplano doubly vergent thrust system, i.e., the KhenayaniUyuni Fault Zone and the Allka Orkho Fault. There, the El
Molino Formation was uplifted along faults and locally
crops out at an altitude of 4000 m (Foldout 1).
[27] 2. The floor of the basins filled with Pilkhaua
Subsequence has a constant elevation of some 3200 m west
of the Khenayani-Uyuni Fault dropping to 2100 m east of
the fault. Because the Pilkhaua Subsequence has a nearly
identical age throughout the southern Altiplano (see below
and Figure 2) and lacks lateral facies changes, this indicates
(1) joint uplift of the latter basins in the west and (2) a
planar decollement geometry below this part of the plateau
without major variation of dip angle that would have
triggered formation of ramp anticlines and associated differential uplift of the above units.
[28] The restored section (see Foldout 1) reveals that
shortening was preceded by an extensional stage that led
to the formation of wedge-shaped half graben fills of the
Eocene/early Oligocene Potoco Formation [cf. Jordan et al.,
2001]. The eastern border of the central half graben is
formed by a normal fault that was inverted as the San
Cristóbal reverse fault. A similar half graben structure may
build the Lı́pez Basin and also appears plausible for explaining the seismic architecture of the western Altiplano.
Estimated extension of the protoplateau domain was less
than 10– 15%.
4. Geochronology: The
Age Data
40
Ar/39Ar and K-Ar
[29] To further decipher the Cenozoic deformation history
of the central and western Altiplano, we sampled volcaniclastic intercalations at key locations along the cross section
for 40Ar/39Ar and K/Ar isotopic age determinations. Linking
age data with structural constraints and stratigraphy aimed
at providing time brackets for specific deformation increments. Samples have been selected to cover the entire
sedimentation and deformation history. Sample material
mostly consisted of felsic, rhyolithic tuffs with abundant
indications for textural and physico-chemical disequilibria.
An outline of data acquisition procedures and a detailed
evaluation of analytical data are given in the auxiliary
material.
[30] The sedimentation history of the San Vicente Formation is surveyed with five crystal tuff samples. Samples
TC4020
KE06-99 and KE10-99 were taken from the eastern flank of
the Vilque Anticline in the eastern Altiplano, sample KE1999 is from the eastern flank of the Ines Anticline in the
central Altiplano, and samples KE11-99 and KE12-99 come
from the northern part of the Yazón Peninsula in the western
Altiplano (Figure 1).
[31] The history of three syntectonic basins of the Pilkhaua Subsequence is constrained as follows: For the basin
west of the Yazón Anticline (western Altiplano), a late
pretectonic tephra sample is from the youngest predeformational strata of the San Vicente Formation, from the vertically inclined forelimb of the Yazón Anticline (KE28-00).
KE24-00 was taken only 50 m farther west, in the basal part
of the steeply inclined (48) syntectonic Pilkhaua sediments,
5 m above the basal conglomerate (Figure 4). The samples
KE01-00 and KE30-00 were interbedded in the syntectonic
basin west of the Ines Anticline, central Altiplano (Figure 5).
KE39-00, KE40-00, and KE41-00 were sampled from the
syntectonic basin west of the central Anticline (Foldout 1).
The crystal tuff KE40-00 is only slightly inclined, while
KE39-00 lies horizontally. KE41-00 is a volcanic boulder in
a volcaniclastic layer within the steeper parts of the basin.
The latter two tuffs were deposited during or after the latest
stages of thrusting. Sample KE13-00 is a posttectonic
ignimbrite of the southern central Altiplano (southwest of
Julaca, Figure 1 and Foldout 1).
[32] The 40Ar/39Ar analytical data are summarized in
Table 1. An important observation is the presence of
variable amounts of excess Ar in almost all samples, both
in hornblende and biotite. We therefore evaluated the data
using both the apparent age spectra and the information
from Ar-Ar isochrons [cf. Miller et al., 1998]. In some
samples, excess Ar is restricted to the first, low-temperature
degassing steps, and high-temperature degassing steps constitute plateau ages, which we regard as valid extrusion ages
for the respective tuffs (KE13-00 biotite, KE24-00 biotite,
KE40-00 biotite, KE28-00 amphibole, KE39-00 amphibole,
and KE41-00 amphibole; Table 2). Samples with persistence of excess Ar throughout the degassing spectra allowed
determination of reasonably precise Ar-Ar isochron ages
(samples KE01-00 biotite, KE39-00 biotite; Table 2), similarly interpreted as extrusion ages.
[33] The K/Ar analytical data are presented in Table 2. All
K/Ar apparent ages for pre-Pilkhaua tuffs of the San Vicente
Formation are higher than 18 Ma. The observation of
abundant presence of excess Ar in the Ar/Ar data set sets
important limitations on the geological significance of the
K-Ar ages from similar rocks. Apparently, the melts that
formed the felsic tuffs and ignimbrites were rich in argon
with elevated 40Ar/36Ar ratios. Correspondingly, Ar solubility is known to be particularly high in felsic, rhyolithic melts
[Kelley, 2002]. It is very likely that the presence of excess Ar
is a general phenomenon in phenocrysts from Altiplano tuffs
and ignimbrites that are petrologically and genetically similar to the rocks that were analyzed by the Ar/Ar method in
the present study. The conventional K/Ar approach depends
on the assumption that all Ar trapped during the magmatic
stage is of atmospheric composition. The K-Ar dates are thus
equivalent to the total gas ages calculated from Ar/Ar data
10 of 19
ELGER ET AL.: ALTIPLANO DEFORMATION
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Table 1. Sample Details and Summary of
Sample
40
TC4020
Ar/39Ar Data Collected by Incremental Heatinga
Sieve
Fraction
(mm)
UTM-x,
UTM-y
Strike/
Dip
Tectonic
Position
Plateau
Age, Ma
TGA,
Ma
Isochron
Age, Ma
—
29.7 ± 5.3
posttectonic
9.4 ± 0.3
10.1 ± 0.4
288/48
syntectonic
17.1 ± 0.7
17.1 ± 1.1
605967, 7730893
323/88
pretectonic
20.9 ± 2.1
24.0 ± 2.9
0.355 – 0.5
675418, 7678968
horizontal
syntectonic
—
15.8 ± 2.6
hbl
0.355 – 0.5
675418, 7678968
horizontal
syntectonic
8.3 ± 1.2
9.2 ± 1.4
crystal tuff
biotite
0.5 – 1
679815, 7684634
272/15
syntectonic
7.9 ± 0.2
8.1 ± 0.3
crystal tuff
hbl
0.125 – 0.25
680936, 7682507
257/60
syntectonic
29.3 ± 2.0
32.2 ± 2.2
15.8 ± 1.2
CF: 0.9878
9.8 ± 0.2
CF: 0.9990
17.0 ± 0.9
CF: 0.9933
20.6 ± 0.1
CF: 0.9999
7.6 ± 0.9
CF: 0.9710
8.5 ± 0.2
CF: 0.9991
8.2 ± 0.1
CF: 0.9997
29.2 ± 0.6
CF: 0.9994
Rock Type
Material
KE01-00
tephra
biotite
>0.25
662592, 7677470
290/60
syntectonic
KE13-00
ignimbrite
biotite
>0.355
644433, 7669689
315/15
KE24-00
tephra
biotite
0.160 – 0.355
605882, 7731505
KE28-00
tephra
hbl
0.2 – 0.5
KE39-00
crystal tuff
biotite
KE39-00
crystal tuff
KE40-00
KE41-00
Initial
Ar/36Ar
40
366.1 ± 20.6
242.7 ± 47.2
330.5 ± 47.8
447.9 ± 31.8
400.6 ± 13.2
307.0 ± 45.5
279.1 ± 20.9
317.6 ± 75.2
a
Isochron age is age from 40Ar/36Ar versus 39Ar/36Ar isotope correlation; TGA is total gas age. Errors are given at the 1s level. Abbreviations are cl,
clast; hbl, hornblende. Preferred ages are in bold type. CF is correlation coefficient of 40Ar/36Ar versus 39Ar/36Ar isochron. Detailed analytical data are
presented in the auxiliary material.
that are systematically too high if excess Ar is present (see
the Ar-Ar data of Table 2). Since presence of excess Ar must
also be suspected for the samples analyzed with the K-Ar
method, all our K-Ar data as well as previously published
K-Ar data from the same area and rocks have to be
interpreted as maximum ages for eruption.
[35] The formation of structures related to deposition of
the Pilkhaua Subsequence occurred between 19 and
7 Ma. In the western Altiplano, the earliest onset of
shortening was found at the Yazón Anticline (youngest
pretectonic tuff, 20.6 ± 0.1 Ma, Ar-Ar; sample ke28-00
overlain by the oldest syntectonic tuff, 17.1 ± 0.7 Ma, ArAr, sample ke24-00, Figure 4). However, interpretation of
the seismic section farther north indicates that these samples
may be underlain by as much as several 100 m of onlapping
Pilkhaua Subsequence that records an earlier onset of
deformation. The oldest undeformed andesitic lava covering
the western Altiplano thrust system (at northern margin of
Salar de Uyuni) was dated at 11.1 ± 0.4 Ma [Baker and
Francis, 1978] marking a minimum age for the end of
shortening here (compare similar observations in Western
Cordillera farther north at 18 – 19S by Charrier et al.
[2002] and Garcı́a et al. [2002]).
[36] Syntectonic sedimentation in the doubly vergent
thrust system of the central Altiplano began at a similar
5. Age and Pattern of Deformation
5.1. Constraints on Deformation Stages
[34] The youngest structures observed are small normal
faults that dissect the syntectonic sediments of the Pilkhaua
Subsequence in the eastern Altiplano (to the east of the
Vilque Anticline, Foldout 1) but are covered by undeformed
Quaternary sediments. These faults are local features related
to underlying evaporates (see chapter on seismic interpretation) and were only observed in the eastern part of the
main cross section (line 10007).
Table 2. Sample Details and Summary of K-Ar Dataa
Sample
Rock
Type Material
KE06-99 crystal
tuff
KE10-99 crystal
tuff
KE11-99 crystal
tuff
KE12-99 crystal
tuff
KE19-99 crystal
tuff
UTM-x,
UTM-y
Strike/
Dip
Tectonic
Position
Ar0,b
ppm
40
Ar0/
Artotal
40
40
Average
Ar0, ppm
40
K, %
Average
K, %
Age,
Ma
biotite
750589, 7664554 108/20 pretectonic 0.007841, 0.008093 0.302, 0.112
0.007967 6.242, 6.291
6.266
18.2 ± 0.5
biotite
751833, 7668609 069/03 pretectonic 0.005046, 0.005200 0.227, 0.222 0.005123 4.124, 4.137
4.130
17.8 ± 0.5
biotite
607111, 7757827 270/55 pretectonic 0.010510, 0.010850 0.274, 0.298 0.010680 6.298, 6.065
6.182
23.6 ± 0.6
fsp
607111, 7757827 275/55 pretectonic 0.000626, 0.000550 0.066, 0.077 0.000588 0.343, 0.353
0.348
24.2 ± 1.2
glass
667540, 7669542 120/55 pretectonic 0.005949, 0.005676 0.325, 0.281 0.005812 3.348, 3.380
3.367
24.7 ± 0.6
a
Abbreviation is fsp, feldspar; 40Ar refers to in situ radiogenic Ar. Constants used are lb = 4.962 1010/yr; l0 + l00 = 0.581 1010/yr; 40K/Ktotal =
1.193 104 g/g. Errors given at the 1s level.
b
Replicate analyses of the same sample (every sample was measured twice).
11 of 19
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ELGER ET AL.: ALTIPLANO DEFORMATION
time: The oldest syntectonic age (15.8 ± 1.2 Ma, Ar-Ar,
ke01-00) was obtained from a 60 WNW dipping tephra in
the thrust front basin west of the west vergent Ines Anticline
(see Figure 5, note, however, that the sample position is
underlain by 200 m of Pilkhaua deposits at depth). In the
east vergent part, the oldest K-Ar maximum age was
derived from a vertically inclined tuff in the footwall of
the San Cristóbal Fault (13.7 ± 0.4 Ma, sample ps99-17
[Silva-González, 2004]; note again underlying Pilkhaua
deposits of 400– 600 m thickness). The end of deformation
is indicated by undeformed dacitic lava, which unconformably overlies steeply inclined San Vicente strata in the
hanging wall of the westernmost fault of the KUFZ
system, dated at 11.0 ± 0.5 Ma (K-Ar, sr99-10 [SilvaGonzález, 2004]). Subhorizontal to horizontal tuffs in
growth sequences yielded 7.9 ± 0.2 and 8.3 ± 1.2 Ma
(ke41-00, Ar-Ar, bedding 272/15; ke39-00, Ar-Ar, horizontal bedding) in the central parts of the doubly vergent
system, and 11.0 ± 0.3 Ma (K-Ar maximum age, ps99-26
[Silva-González, 2004]) in the footwall of the San Cristóbal
Fault. Subhorizontal tuffs sealing the Khenayani-Uyuni
Fault (10 ± 0.3; sample SR99-15 [Silva-González, 2004])
indicate that deformation at the KUFZ was shut off between
10 and 11 Ma (compare similar observations on TamboTambillo structure toward the north [Lamb and Hoke,
1997]). The progressively tilted deposits of growth strata
in the hanging wall of the KUFZ indicate continuous
thrusting since 14 Ma (oldest dated Pilkhaua deposit).
However, the parallel and folded reflectors of several 100 m
of San Vicente deposits underlying this sample may indicate
slow or absent deformation earlier. On the other hand, the
absence of San Vicente deposits older than 15– 18 Ma in the
KUFZ structure, their unconformable basis covering all
deeper units in the hanging wall anticlines, and the thickness of 3 – 3.5 km for the San Vicente deposits in the
footwall require continuous or repeated fault activation
during San Vicente deposition in order to maintain local
relief and to provide accommodation space (see above).
[37] In the eastern Altiplano domain, the laterally uniform
thickness and overall conformable boundaries of the OMsv1
and OMsv2 Subsequences indicate a late Miocene age of
deformation for the Vilque and the eastern Anticlines (see
Foldout 1). The youngest OMsv2 age in the eastern Altiplano (16 – 17 Ma, extrapolated from samples ke10/99,
17.8 ± 0.5 Ma, and ke 06/99, 18.2 ± 0.5, 200 m below
top OMsv2) is a maximum age of formation of these
structures. Onlaps and divergent reflectors to the west of
the Vilque Anticline (line 10010, see Foldout 1) indicate
that its growth was coeval with the deposition of the
Pilkhaua Subsequence. The youngest mildly tilted tuff in
the Pilkhaua Subsequence north of the Vilque anticline has
an age of 7.6 ± 0.2 Ma [Silva-González, 2004].
[38] Except for the observations in the KUFZ area,
undisturbed deposition of the San Vicente Formation in
the entire southern Altiplano indicates a stage of local
tectonic quiescence. In contrast, its more complex internal
sequence architecture toward the margin of the Eastern
Cordillera (with westward prograding foresets) must have
witnessed nearby uplift east of the study area. From the
TC4020
sedimentary record in the San Vicente Formation with
material provenance from the east [Horton et al., 2002],
deformation in the adjoining parts of the Eastern Cordillera
(i.e., San Vicente thrust and hanging wall to the east) must
have peaked between 25 and 17 Ma, during deposition of
the westward prograding OMsv2. According to Müller et al.
[2002], the nearby San Vicente thrust was sealed by
volcanics with an age of 17 Ma marking the end of the
main stage of deformation and uplift of the margin of the
Eastern Cordillera. This matches the end of deposition of
the OMsv2 Subsequence of the San Vicente Formation.
[39] The locally developed truncation of tilted hanging
wall sequences below the San Vicente/OMsv1 across major
parts of the southern Altiplano is evidence for pre-San
Vicente deformation. The almost complete lack of volcanic
intercalations in the Potoco Formation, which was deposited
prior to this deformation increment, prevents deriving
precise age constraints from syntectonic deposits. A single
tuff age within intensely folded Potoco sediments in the
footwall of the Corregidores Fault yielded a K-Ar maximum
age of 40.4 ± 1.1 Ma (sr99-03 [Silva-González, 2004]). The
oldest dated volcanic rock of the San Vicente Formation in
the central Altiplano gave 29.3 ± 1.9 Ma (volcanic boulder
in volcaniclastic deposit; sample ke41-00, amphibole ArAr). Moreover, near sites of early shortening (eastern
margin of KUFZ) the basal conglomerate overlying Eocene
to Paleozoic sediments contains Silurian clasts from the
adjacent eroded hanging walls and is overlain by a basalticandesitic tuff with a maximum age of 27.4 ± 0.6 Ma (K-Ar,
sample sr99-12 [Silva-González, 2004]). Finally, Ege [2004]
reports apatite fission track ages from the uplifted Paleozoic
basement rocks of the KUFZ faults, ranging between 28 and
33 Ma. In conclusion, this early shortening stage in the
central Altiplano is bracketed between 35 and 27 Ma.
[40] In the eastern Altiplano domain, the anticlinal hinges
formed during this stage were eroded prior to the deposition
of the lowermost OMsv1 sediments. Some 700 m above the
base of the latter, a K/Ar age of 25.3 ± 1 Ma was determined
(sample an14-2000 [Silva-González, 2004]). Considering
average sedimentation rates (compare Figure 2b) the base
of the San Vicente Formation, and hence the period of this
early deformation increment, is constrained to between 28
and 32 Ma at the western margin of the Eastern Cordillera.
5.2. Evolution of Shortening
[41] Using the data on horizontal shortening as derived
from section construction and the newly reported as well as
published ages from synkinematic deposits, we analyze the
evolution of shortening in more detail with the strategy
outlined by Victor et al. [2004]: Shortening related to each
fault is distributed over the period of its activity as constrained from the sedimentary record in the immediately
adjoining syntectonic deposits. The age estimate for the
oldest beds recording the onset of deformation and of the
youngest deformed are based on the above isotopic data or
were estimated from extrapolation using average sedimentation rates (the latter are 100 m/Myr ± 20 m/Myr as derived
for the Pilkhaua Subsequence across the Altiplano with
12 of 19
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ELGER ET AL.: ALTIPLANO DEFORMATION
local maximum of 200 m/Myr in the center of the Lı́pez
basin) and some 150– 250 m/Myr or 150 – 200 m/Myr for
the deeper members of the San Vicente Formation in the
eastern and central Altiplano domain. respectively; compare
Figure 2). To assess the maximum or minimum period of
fault activity we included the respective errors in isotopic
age dating and the sedimentation rate estimate used to
calculate the ages of the base and top of syntectonic
deposits. Potential error in determining heave is estimated
at ±10% (+20% in less well constrained sections). Using
these estimates, we assessed the highest and lowest possible
shortening rates for each fault over its activity period.
Summing these numbers for all kinematically linked faults
provides minimum shortening rates (minimum shortening
and maximum activity period) and maximum shortening
rates (maximum shortening and minimum activity period)
for a thrust system.
[42] Figure 6 shows the summed results merged with
equivalent data from the literature for the western plateau
flank and the Eastern Cordillera at 21S. As mentioned
above, two main stages of shortening can be identified
across the Altiplano up to the western border of the
Eastern Cordillera. An earlier Oligocene stage (between
35 and 27 Ma) has averaged local shortening rates
between 0.1 and some 3 mm/yr (not resolved in western
Altiplano for lack of exposure). This stage is followed by
an early Miocene lull in deformation. Shortening resumed
with a distinct local acceleration to some 1.7 – 3 mm/yr
(west and center) in the middle to late Miocene (20 –
10 Ma). In most of the Altiplano, rates were higher
during the later stage while the eastern Altiplano shows
the opposite relationship. At 7 –8 Ma, deformation ceased
everywhere. From these figures, bulk Cenozoic strain
rates in the southern Altiplano were at an average of
3.5 1016 s1 with a middle Miocene maximum of 2 1014 s1 (Figure 7c).
5.3. Deformation Partitioning
[43] Inspection of the data at higher temporal and
spatial resolution underlines the complex mode of strain
accumulation discussed in the following with two examples. In the symmetrical syntectonic basin (Pilkhaua
Subsequence) separating the adjacent Ines and Allka
Orkho faults (Figure 5), growth sediments record alternating and contemporaneous activity of both faults:
stage 1, activity of Allka Orkho Fault and minor uplift
of Ines fold (ST1); stage 2, growth of Ines fold, quiescent
Allka Orkho Fault (ST2); stage 3, growth of Allka Orkho
Fault dominates over Ines fold (ST3); and stage 4, the
Ines Fold again takes over. At the scale of the entire
central Altiplano unit, we observe Miocene initiation of
first the east vergent KUFZ structures, succeeded by
formation of new thrusts (14 – 11 Ma) in all parts of this
thrust system. Finally, active deformation between 11 and
8 Ma prograded westward with minor shortening in the
internal parts of the system and a quiescent eastern
deformation front.
[44] The eastern Altiplano shows a similar evolution for
the deformation front of the Eastern Cordillera (section
TC4020
10010 in Foldout 1f). Shortly before or during initial
deposition of OMsv1, shortening propagated westward into
the Lı́pez basin and, subsequently, retreated back east to the
San Vicente thrust. The latter accommodated subsequent
shortening at the Eastern Cordillera front between 27 and
17 Ma (i.e., during OMsv1 and OMsv2 [Müller et al.,
2002]). Following shutoff of the San Vicente Thrust,
deformation again propagated westward, forming the Vilque
Anticline, until its end at 7 Ma.
[45] Across the entire southern Altiplano, the Pilkhauafilled basins exhibit 3 to 5 internal subsequences based on
observation of reflector terminations and reflector packages
(compare Foldout 1 and Figures 5c and 5d). These packages
show a trend toward long high-amplitude reflections at their
respective tops. This trend in reflection signal reflects a
progressive change in grain size, porosity, and/or average
bed thickness. In the Vilque well, Welsink et al. [1995]
describe the San Vicente Formation drilled to consist of two
major fining upward sequences matching the two seismic
sequences in line 10007 that crosses the well site (compare
Foldout 1). Mertmann et al. [2003] also describe sequences
of several 10 to 500 m thickness showing internal fining
upward cycles in the San Vicente Formation While these
sequences are very probably also influenced by climatic
fluctuations the nature of the sediments as syntectonic
deposits next to tectonically created relief indicates that
changes of the balance between relief formation and relief
degradation have the most important impact on the local
sedimentation.
[46] From the age span recorded in the Pilkhaua Subsequence (10 Ma, Figure 2), the above number of internal
sequences amounts to substages of locally enhanced rates of
formation of relief at recurrence intervals of 2 ± 1 Ma for an
individual structure. The San Vicente Formation shows a
laterally varying number of 4 – 7 reflector packages of
similar properties as mentioned above spanning the
time between 29 and 19 Ma (Figure 5a and Foldouts 1c,
1e, and 1f). This time span yields the same recurrence
interval as for the Pilkhaua Subsequence. Hence the major
deformation periods in the Altiplano identified in Figures 6c
and 6d that typically encompass some 5 – 15 Myr are broken
down locally into shortening rate peaks at a typical
timescale of 1 to 3 Myr for individual structures. Because
of the variable number of these internal sequences between
neighboring structures, the structures within a coupled fault
system may at best exhibit a partial synchronization. More
probably, they record random, mostly unsynchronized
distribution of these fluctuations of shortening rate.
6. Discussion
6.1. Deformation Age and Pattern
[47] The new data in this study together with published
data complete a cross section of the central Andean Plateau
and its margins at 20– 22S, except for a 40-km-long
segment of the Western Cordillera (the Recent magmatic
arc), where Miocene to Quaternary volcanic deposits prevent access to the subsurface and no high-resolution reflection seismic data are available. However, observations from
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Figure 6. Distribution of deformation ages, shortening, and magmatism across the southern central
Andes (21S) based on published and own data. (a) Compilation of deformation ages: Western Flank
[Victor et al., 2004], Precordillera [Haschke and Guenther, 2003], Altiplano [this study; Ege, 2004; SilvaGonzález, 2004], Eastern Cordillera [Gubbels et al., 1993; Müller et al., 2002], inter-Andean [Kley, 1996;
Ege, 2004], and sub-Andean [Kley, 1996]. (b) Balanced cross section at 21S compiled from Victor et al.
[2004] for Altiplano West Flank; own data (Altiplano), and Müller et al. [2002] for Eastern Cordillera and
sub-Andean; Moho and Andean Low-Velocity Zone (ALVZ) from receiver function data [Yuan et al.,
2000]. Line drawing in the middle crust indicates locations of strong reflectivity in ANCORP seismic line
[cf. ANCORP Working Group, 2003]. The western Altiplano cross section was projected along strike of
gravity anomalies (Figure 1) 50 km southward. (c) Estimates of maximum and minimum periods of
active faulting and related folding in the units building the Altiplano (see Foldout 1 for names and
location of structures). (d) Cumulative shortening rates of the western, central, and eastern Altiplano
thrust systems, calculated from fault activity periods and heave (see text for explanation; result for
western Altiplano is extrapolated from observation of ages at Yázon peninsula to all faults observed in the
seismic line considering the maximum possible error). Bold line delineates average shortening rates based
on sliding average of 3 Myr sampling width. Shortening rate of the Precordillera from Victor et al. [2004]
(complemented by Haschke and Guenther [2003]). For the Eastern Cordillera we estimated average
shortening rates for individual subunits based on sections and age data summarized by Müller et al.
[2002]. Because of the lack of resolution, there are both very few preserved piggyback basins and
incomplete age dating of individual structures; the uncertainty here is higher than in the plateau area. See
color version of this figure at back of this issue.
exposed Neogene deposits of the Western Cordillera area
farther north (18 – 19S [Garcı́a et al., 2002; Charrier et al.,
2002]) show similar age properties as the Precordillera and
the western Altiplano at 21S and may be projected to the
unexposed part of the section at 21S. The structural style of
the entire southern Altiplano Plateau is here identified as a
series of up to four doubly vergent thrust systems, which
vary in size and in amount of shortening (Figure 6). Most of
these resulted from inversion of earlier extensional structures. Inversion commenced in the Eocene in the west
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(Precordillera domain [Haschke and Guenther, 2003]) and
only affected the protoplateau area farther east in the
Oligocene.
[48] Accumulation of shortening in the past 45 Myr (see
Figures 6 and 7) initiated at the protoplateau margins, i.e.,
the Eastern Cordillera and Chilean Precordillera at around
40– 45 Ma ago [e.g., Sempere et al., 1997; Lamb and Hoke,
1997; Müller et al., 2002; Haschke and Guenther, 2003]
followed by a break in the west. The subsequent evolution
shows peak activity to repeatedly shift between the Eastern
Cordillera and the more central and western parts of the
evolving plateau area. In the final stage at 11 –8 Ma,
contraction affected the entire southern Altiplano plateau
area dying out everywhere at around 7 – 8 Ma. Shortening
then migrated to the eastern flank of the plateau where
eastward migrating deformation led to the formation of the
broad, thin-skinned thrust belt of the inter-Andean and subAndean Ranges [Kley, 1996].
[49] Synchronization of peak shortening periods in the
western systems (Altiplano west flank, western and central
Altiplano) and their anticorrelation with those from the
Eastern Cordillera, exhibiting repeated advance and retreat
of its western deformation front that is unrelated to the
remaining Altiplano, indicates kinematic coupling of
various thrust systems: One is rooted in the Chilean
Precordillera/western Altiplano and the other in the Eastern
Cordillera. The existence of the here newly defined western
Altiplano Thrust Belt (see Figures 8 and 6b), coupling the
three western thrust units, is also corroborated by the here
identified gently westward dipping detachment underlying
the systems. These findings make the more eastern parts of
this belt (western and central Altiplano thrust systems) a
broad east facing, low-tapered prowedge accumulating most
shortening (60 km). This prowedge is paired by the
steeply inclined, slowly moving and narrow retrowedge
forming the Western Flank Monocline/Precordillera of the
Altiplano (3 – 4 km shortening [cf. Victor et al., 2004]).
[50] Although the data imply significant synchronicity of
deformation stages across parts of the plateau at time frames
>5 Myr, more detailed inspection of the seismic sequence
architecture of the syntectonic basins reveals significant
short-term complexity with repeated retreat of the deformation front (Eastern Cordillera) or repeated alternate thrust
activation in the individual doubly vergent systems, and
last, but not least, variations in local fault slip rates at local
timescales of 1– 3 Myr. Hence tectonic deformation rates
would be seen to peak at various timescales between the
isotopically resolved scales of several million years down to
at least the subisotopic geological record exhibiting discon-
Figure 7. Cumulative shortening and shortening rate
evolution of the southern Altiplano (sub-Andean plateau
flank not included). (a) Summed minimum and maximum
shortening rates across the southern Altiplano. Bold line
depicts average based on sliding 3 Myr average. (b) Stacked
average shortening rates of individual thrust systems
building the southern Altiplano (from Figure 6).
(c) Cumulative shortening.
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Figure 8. Outline of western Altiplano thrust belt and of
the western deformation front of the Eastern Cordillera
thrust belt drawn on shaded relief map of Altiplano Plateau.
Note narrow intervening basin between both belts (Lı́pez
Basin (LB) in the south and Toledo Basin (TB) in the north)
as well as restricted outcrop of prowedge of western
Altiplano Thrust Belt in Corque (C), Tambo-Tambillo (TT),
and the here analyzed San Cristóbal (SC) and Yazón (Y)
areas.
tinuous deformation of the southern Altiplano crust over a
range of scales.
[51] The pattern of plateau-wide bulk pure shear contraction followed at around 10 Ma by a switch to bulk overthrusting of the plateau toward the east as identified by
Allmendinger and Gubbels [1996] is here refined by resolving the pure shear stage into several substages with alternate
or overlapping activity of two major thrust systems. The
observed lack of any large-scale eastward or westward
migrating deformation front over the entire plateau, i.e.,
between the Western Flank and the Eastern Cordillera, is in
contrast to recently published models of continuous eastward plateau expansion based on an interpretation of the
sedimentological record [McQuarrie and DeCelles, 2001;
Horton et al., 2002; McQuarrie, 2002].
[52] The summed rates of all subsystems building the
Altiplano highlight and reflect the important individual
stages in Altiplano deformation that were previously repeatedly identified across the plateau as key deformation
phases [e.g., Sempere et al., 1990; Lamb and Hoke, 1997;
Allmendinger et al., 1997]. Their reinterpretation in the light
of the here presented data would suggest these to be
fluctuations of shortening rates superimposed on the general
trend of increase of bulk shortening velocity of the upper
plate identified by Hindle and Kley [2002]. Moreover, the
widespread occurrence of these periods of enhanced shortening rates throughout the plateau would suggest that the
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here reported details along the section at 21S may well be
representative for most of the Altiplano to the north (compare to observations by Lamb and Hoke [1997]). Toward
the south, the Puna plateau currently exhibits seismic
activity across its entire area, bulk GPS-based shortening
at some 6 mm/yr distributed over the entire Puna [Klotz et
al., 2001], an increase of shortening rate since the Pliocene
[Ramos et al., 2002], and the growth of three to four
regularly spaced thrust systems across strike that confine
intervening basins. This situation is highly similar to the late
Miocene stage of synchronous plateau-wide shortening in
the southern Altiplano prior to eastward expansion with an
associated acceleration of rates.
[53] Last, but not least, it is interesting to note that this
evolution of shortening rates has a complex relation with the
evolution of plate convergence rates between the South
American and the oceanic Nazca plates [cf. Pardo-Casas
and Molnar, 1987; Somoza, 1998]. Only the early, Eocene
to Oligocene, evolution of upper plate shortening may be
seen to correlate with two stages of enhanced plate convergence rates at 50 to 44 Ma and 25 to 20 Ma. In contrast, the
decrease of Neogene convergence rate from some 15 cm/yr
at 20 Ma to the present value of 6.5 cm/yr [Klotz et al.,
2001] is anticorrelated with a shortening rate increase [cf.
also Hindle and Kley, 2002]. We therefore speculate that
plate convergence rate was only a relevant factor controlling
upper plate shortening rate in the initial stages.
6.2. Shortening, Crustal Thickening, and Surface Uplift
[54] The distribution of upper crustal shortening concluded
from the two thrust systems building the plateau (western
Altiplano thrust system, 65 km; Eastern Cordillera,
85 km [cf. Müller et al., 2002]) does not match an
equivalent thickness distribution of the deeper crust beneath
these (compare Figure 6b). While there is slightly thicker
crust underlying the plateau margins, various geophysical
data (see Figure 6) [Wigger et al., 1994; Beck et al., 1996;
Yuan et al., 2000; ANCORP Working Group, 2003] indicate
that the plateau crust everywhere exceeds 60 km thickness.
Moreover, plateau surface uplift as recorded by GregoryWodzicki [2002] with enhanced rates since the late Miocene
is succeeding the main stage of upper crustal shortening in
the plateau domain. Hence upper crustal shortening as
discussed here is geometrically as well as temporally
decoupled from deeper crustal and mantle lithosphere
deformation. This is in line with late Neogene to Recent
underthrusting of the plateau domain from the east and
deep crustal flow as suggested by various authors [i.e.,
Allmendinger and Gubbels, 1996, and references therein].
From the lack of a significant near surface deformation
response to this stage, we conclude a near complete mechanical decoupling of the upper crust from the deeper crust
since at least the late Miocene.
6.3. Magmatism and Deformation
[55] The evolution of the Altiplano was paralleled by arc
and back arc volcanic activity since the late Oligocene,
largely coeval therefore to the period of shortening and
largely matching the lateral extent of the plateau. This link
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has been suggested to reflect a general genetic relationship
between magmatic activity and plateau-forming processes
in terms of thermally controlled weakening of the upper
plate [e.g., Isacks, 1988; Allmendinger et al., 1997; Kay et
al., 1999; Babeyko et al., 2002; ANCORP Working Group,
2003].
[56] In the southern Altiplano, the initiation of volcanic
activity (27 – 29 Ma) slightly postdates the onset of the first
shortening increment (35 – 27 Ma) and the early stage of
deposition of the San Vicente Formation (compare
Figures 6a and 7a) (compare similar observations in other
parts of the plateau [e.g., Kay et al., 1999; Riller et al.,
2001; Ramos et al., 2002]). Volcanism continued until 10–
8 Ma in the back arc plateau with ongoing magmatism in
the present arc [Evernden et al., 1977; Fornani et al.,
1993; Soler et al., 1993; Kennan et al., 1995], largely
overlapping therefore with the main stage of shortening of
the entire southern Altiplano domain. Hence it would
appear that near surface magmatism in the southern
Altiplano locally followed the onset of deformation with
a minor time lag, but that it paralleled the general trend of
increased shortening rates of the southern Altiplano plateau (compare Figure 7a).
[57] On the basis of these observations, it appears plausible that the thermal evolution of the lithosphere resulting
in transient presence of melts may have sufficiently mechanically weakened the upper crust of the Altiplano to
enhance global deformation rates through time as observed
for the plateau (compare classical model by Isacks [1988]).
Modeling results by Babeyko et al. [2002] showed that
advective melt transfer through the crust with associated
crustal melting requires some previous crustal shortening
and thickening, which is equivalent to the correlation
observed here.
[59] This conclusion and the above observations throw
additional light on the dynamics of plateau formation. On
the basis of numerical modeling, the mechanical evolution
underlying plateau formation (see Wdowinski and Bock
[1994], England and Molnar [1997], Pope and Willett
[1998], Shen et al. [2001], and Vanderhaeghe et al.
[2003] for details) involves heating and related weakening
of the lower crust as a result of continuous crustal thickening and melting. However, most of these models invoke an
initial domain of localized deformation, which expands
laterally forming classical self-similar growing wedges until
mechanical decoupling at the base entails surface flattening
and the formation of a high plateau with two topographically distinct margins bounding a hot, viscous deeper
lithosphere. The resulting plateau lacks a lateral mass
gradient forcing lateral migration of deformation. Subsequent shortening within the plateau would be delocalized
with an irregular pattern in time and space coincident with
the here reported observations. In contrast to the model
predictions, however, our results show that plateau initiation
in the Altiplano area occurred by growth of two independent thrust systems at the site of the later plateau margins
during the early stages. These do not show much evidence
for lateral expansion until the sub-Andean ranges developed
with concordant shutdown of deformation of the plateau. In
the case of the central Andes, first-order anisotropies
probably controlled this initial localization of deformation
and, hence the subsequent evolution: the hot magmatic arc
crust in the kinematic axis of the western Altiplano thrust
system, and large-scale mechanical heterogeneity resulting
from repeated rifting of the crust at the site of the later
Eastern Cordillera.
6.4. Plateau-Style Deformation
[60] The southern Altiplano structure formed during a
complex shortening history that peaked in the early Oligocene (<35– 27 Ma) and middle/late Miocene (19 – 7 Ma).
These stages were preceded by an extensional stage that
resulted in formation of a series of grabens across the entire
Altiplano, which later localized the formation of several
thrust systems. The three western systems, the west facing
Precordillera/Western Cordillera and the east facing western
and central Altiplano systems probably form a coupled
retroprowedge system that is separate from the eastern
Altiplano/Eastern Cordillera system. Both cordilleras delimiting the later plateau area started shortening earlier during
the Eocene (40 – 45 Ma). Subsequently, their deformation
history was decoupled with joint activity only in the middle
Miocene. Horizontal contraction on the southern Altiplano
gradually subsided between 11 and 7 Ma as indicated by the
age of undeformed volcanic rocks sealing the contractional
structures.
[61] Deformation of the Andean Plateau at 21S is
characterized by a spatially and temporally irregular fault
activity pattern, at a generally increasing bulk rate of plateau
shortening during the Oligocene/Miocene. Largely synchronized peaks at the 5– 15 Myr timescale within a kinematically coupled thrust system were complemented by local
[58] Along with the above observation of fluctuating
shortening rates, the highly irregular spatial and temporal
pattern of strain accumulation at all scales hints at a selfsimilar distribution of deformation during plateau growth
since the Eocene with a lack of characteristic length and
timescales in the processes shaping the plateau. Fault slip,
or periods of enhanced local strain rates, on a fault network
that is in failure equilibrium and has an approximately
fractal size distribution (compare observations by Marrett
and Allmendinger [1991] in the Puna) would be expected to
generate the type of scale-independent response found in the
syntectonic sedimentary record of the Altiplano. It is
obvious that the number of small tectonic events, at least
partly reflected in individual layers deposited in the syntectonic basins, is significantly larger than the number of
reflection seismic sequences within each basin. Their number, in turn, is lower than the few larger ‘‘displacement
events’’ building the Andes, like the rapid strong shortening
of the Eastern Cordillera in the Miocene. These observations suggest that plateau deformation in the Andes is
essentially accommodated by a series of interconnected
faults, which are close to failure equilibrium helping to
distribute strain.
7. Conclusions
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fluctuations of strain accumulation at individual faults at a
typical timescale of 1 – 3 Myr that show no evidence of
lateral synchronization. These observations preclude systematically propagating deformation. The spatial and temporal pattern of deformation indicates that the plateau
probably was self-organized fluctuating about a critical state
of failure throughout its evolution. This style of deformation
is fundamentally different from smaller-scale orogenic
wedges that exhibit more continuous lateral growth. The
Altiplano plateau has probably had low relief since its
initiation and evolved from partly independent activity of
two thrust belt systems rather than from a single initially
doubly vergent orogen. Previously suggested thermal weakening of the plateau crust is reflected in a loose spatial and
temporal relationship between deformation and, slightly
later onset of, volcanism. Eocene to Oligocene crustal
shortening with some relation to plate convergence rates
preceded thermal weakening of the crust. The latter, in
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turn, may have been responsible for discontinuous acceleration of upper plate shortening rates during slowdown of
convergence.
[62] Acknowledgments. Research for this project was funded by the
German Science Foundation (DFG) under the auspices of the Collaboration
Research Center 267 ‘‘Deformation Processes in the Andes’’ and the
Leibniz-Program (On7/10-1). This support is gratefully acknowledged. A
great number of people have helped through discussion of our results as
well as in finalizing the manuscript. First and foremost, the former Bolivian
Oil Industry (YPFB) provided the seismic sections in the context of the
ANCORP cooperation: David Tufiño and Ramiro Suárez (YPFB) as well as
Oscar Aranibar and Eloy Martı́nez (Chaco SA) provided valuable support
in seismic data exchange and interpretation; K. Munier and P. Gôni
processed the satellite images; M. Dziggel has drawn most of the figures;
W. Frank (Vienna University) has provided the Ar/Ar age data; T. Vietor,
K.-J. Reutter, D. Mertmann, and Hernán Castillo have supported field work;
J. Herwig has prepared samples for age dating; and H. Ege, E. Scheuber,
and P. Silva-González provided important additional age data. Last, but not
least, G. Wörner and an anonymous reviewer provided thoughtful reviews.
We thank all these colleagues for their collaboration and help.
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Figure 1
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Figure 1. (a) Geological sketch map of the southern Altiplano (section of map of central Andes by Reutter et al. [1994]
including our own mapping results) superimposed on shaded relief of the topography showing the distribution of pre-late
Neogene rocks, lower-hemisphere equal-area stereonet projections with bedding and regional fold axes, and the location of
reflection seismic lines. (b) Bouguer anomaly map (provided by H. J. Götze; see Götze and Krause [2002] for description
of gravity field) of area shown in Figure 1a depicting positive anomalies closely related to faults and folds mapped at the
surface and observed in the seismic sections. Note the chain of anomalies in the nearly completely sediment-covered
western Altiplano with similar style as obvious from the exposed central Altiplano unit. (c) Location of the working area.
(d) Major morpho-tectonic units of the southern central Andes.
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Figure 6. Distribution of deformation ages, shortening, and magmatism across the southern central
Andes (21S) based on published and own data. (a) Compilation of deformation ages: Western Flank
[Victor et al., 2004], Precordillera [Haschke and Guenther, 2003], Altiplano [this study; Ege, 2004; SilvaGonzález, 2004], Eastern Cordillera [Gubbels et al., 1993; Müller et al., 2002], inter-Andean [Kley, 1996;
Ege, 2004], and sub-Andean [Kley, 1996]. (b) Balanced cross section at 21S compiled from Victor et al.
[2004] for Altiplano West Flank; own data (Altiplano), and Müller et al. [2002] for Eastern Cordillera and
sub-Andean; Moho and Andean Low-Velocity Zone (ALVZ) from receiver function data [Yuan et al.,
2000]. Line drawing in the middle crust indicates locations of strong reflectivity in ANCORP seismic line
[cf. ANCORP Working Group, 2003]. The western Altiplano cross section was projected along strike of
gravity anomalies (Figure 1) 50 km southward. (c) Estimates of maximum and minimum periods of
active faulting and related folding in the units building the Altiplano (see Foldout 1 for names and
location of structures). (d) Cumulative shortening rates of the western, central, and eastern Altiplano
thrust systems, calculated from fault activity periods and heave (see text for explanation; result for
western Altiplano is extrapolated from observation of ages at Yázon peninsula to all faults observed in the
seismic line considering the maximum possible error). Bold line delineates average shortening rates based
on sliding average of 3 Myr sampling width. Shortening rate of the Precordillera from Victor et al. [2004]
(complemented by Haschke and Guenther [2003]). For the Eastern Cordillera we estimated average
shortening rates for individual subunits based on sections and age data summarized by Müller et al.
[2002]. Because of the lack of resolution, there are both very few preserved piggyback basins and
incomplete age dating of individual structures; the uncertainty here is higher than in the plateau area.
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