The Andes are an emblematic active Cordilleran orogen.
Mountain building in the Central Andes (∼20∘ S)
started by the Late Cretaceous to early Cenozoic along the subduction margin
and propagated eastward. In general, the structures sustaining the uplift of
the western flank of the Andes are dismissed, and their contribution to
mountain building remains poorly constrained. Here, we focus on two sites
along the Western Andes at ∼20–22∘ S in the
Atacama desert, where structures are well exposed. We combine mapping from
high-resolution satellite images with field observations and numerical
trishear forward modeling to provide quantitative constraints on the
kinematic evolution of the investigated field sites. When upscaling our
local field interpretations to the regional scale, we identify two main
structures: (1) the Andean Basement Thrust, a west-vergent thrust system
placing Andean Paleozoic basement over Mesozoic strata, and (2) a series of
west-vergent thrusts pertaining to the West Andean Thrust System, deforming
primarily Mesozoic units. From our interpreted sections, we estimate that
both structures together accommodate at least ∼6–9 km of
shortening across the sole investigated ∼7–17 km wide field
sites. This multi-kilometric shortening represents only a fraction of the
total shortening accommodated across the whole Western Andes. The timing of
the main deformation recorded in the folded Mesozoic series can be bracketed
between ∼68 and ∼29 Ma – and possibly between
∼68 and ∼44 Ma – from dated deformed
geological layers, with a subsequent significant slowing-down of shortening
rates. Even though the structures forming the Western Andes only absorbed a
small fraction of the total shortening across the whole orogen, their
contribution was relatively significant at the earliest stages of
Andean mountain building before deformation proceeded eastward.
Introduction
Along the western margin of South America (Fig. 1), the oceanic Nazca
plate plunges beneath the South American continent, with a present-day
convergence rate of ∼8 cm yr-1 at ∼20∘ S according to the NUVEL-1A model (Demets et al.,
1994). The subduction megathrust absorbs most of this convergence in the
form of large earthquakes (magnitude Mw≥8). A small fraction of it –
presently ∼1 cm yr-1 at 20∘ S (e.g., Norabuena et
al., 1998; Brooks et al., 2011) – has contributed to the deformation of the
upper plate over millions of years and to the formation of one of the
largest reliefs at the Earth's surface: the Andean Cordilleras and the
Altiplano–Puna Plateau in between.
Simplified geological and structural map of the western
Central Andes at ∼20–22∘ S (northern Chile) and
average topographic profile. Map modified from Armijo et al. (2015); vertical exaggeration (ve) of 5 for topographic profile. The
Marginal Block and the Western Cordillera constitute the two main structural
ensembles. The Western Cordillera includes the West Andean Thrust System
(WATS), a basement high (Cordillera Domeyko), and the modern volcanic arc.
The structures forming the WATS are mostly hidden beneath blanketing
Cenozoic deposits and only outcrop in a few places. The Andean Basement Thrust
separates the basement high of the Western Cordillera from the WATS. The
locations of Figs. 4 (Pinchal area) and 8 (Quebrada Blanca area) are
reported. Inset: location of the map (red box) within the South American
continent.
Andean mountain building initiated by Late Cretaceous–early Cenozoic along
the Western Andes of the Bolivian Orocline (between 16–22∘ S)
and has proceeded since then with the progressive eastward propagation of
deformation onto the South American continent (e.g., Armijo et al., 2015;
Charrier et al., 2007; Decelles et al., 2014; Eichelberger et al., 2013;
Oncken et al., 2006; Jaillard et al., 2000; and references therein). Most
local and mountain-wide studies have essentially focused on the
Altiplano–Puna Plateau and on the various cordilleras to the east. In
comparison, the structures located along the western flank of the orogen
have remained relatively understudied up to now.
In most classical models of Andean mountain building, the western flank is
described as a passive monoclinal-like crustal-scale flexure (e.g.,
Isacks, 1988; Lamb, 2011, 2016; Mcquarrie, 2002). However, in the late
1980s, Mpodozis et al. (1989) described west-vergent thrusting
along the Western Andean margin. Later, other authors also described various
thrusts, mostly west-vergent (but not only) (e.g., Charrier et al., 2007;
Farías et al., 2005; Fuentes et al., 2018; Garcia and Hérail, 2005;
Martinez et al., 2021; Muñoz and Charrier, 1996; Victor et al., 2004),
but they generally gave these thrusts a minor role in the building of the
western flank of the orogen. Only further south, at the latitude of Santiago
de Chile (∼33∘30′ S), has a clear west-vergent
fold-and-thrust belt been documented along the Western Andes (Armijo
et al., 2010; Riesner et al., 2017, 2018). This thrust belt
emerges at the active San Ramon Fault in front of the capital city of
Santiago de Chile and has absorbed a relatively significant amount of
crustal shortening (Riesner et al., 2017, 2018).
At 33∘30′ S, the orogen is relatively young and narrow. In
contrast, ∼1300 km further north, at ∼20–22∘ S where the Andes–Altiplano system is much wider and
structurally more complex, the contribution of structures along the Western
Andes is probably small compared to the >300 km total shortening
(e.g., Anderson et al., 2017; Barnes and Ehlers, 2009; Eichelberger et
al., 2013; Elger et al., 2005; Kley and Monaldi, 1998; Mcquarrie et al.,
2005; Sheffels, 1990) across the entire >650 km wide orogen, but
their role at the onset of orogenic building may have been significant
(Armijo et al., 2015). One of the difficulties in better
quantifying the contribution of these structures is that a large part of the
deformation is hidden under blanketing middle to upper Cenozoic deposits and
volcanics (Fig. 1) (Armijo et al., 2015; Farías et al., 2005;
Sernageomin, 2003; Victor et al., 2004). A quantitative analysis of this
deformation and its kinematics is only possible at the few sites along the
western flank where deformed Mesozoic series crop out and which are
accessible despite the hostile desert conditions in northern Chile.
In this study, we provide quantitative structural data to better constrain
the geometry of structures, the shortening they accommodated, and their
kinematics of deformation over time in two of the few areas along the Western
Andean flank where the underlying deformed Mesozoic layers are exposed
(Fig. 1). The Pinchal area, at ∼21∘30′ S,
exhibits a west-vergent thrust that brings the Paleozoic basement of the
Cordillera Domeyko over folded Mesozoic units. In the Quebrada Blanca zone,
∼80 km further north, the excellent exposure of folded
Mesozoic series allows for a more quantitative estimate of the shortening
and of the timing of the main deformation episode. These two study areas
only give a limited view of the deformation of the whole Western Andean
flank (Fig. 1). Despite these limitations, we find that the shortening of
these structures is multi-kilometric, revealing that the contribution of the
Western Andean flank to Andean mountain building is not negligible.
Additionally, we show that the main deformation recorded by the folded
Mesozoic units occurred sometime between ∼68 and
∼29 Ma (and possibly between ∼68 and
∼44 Ma), further emphasizing that these structures mostly
participated in the early stages of mountain building.
Geological context of the Central Andes (∼20–22∘ S)General geological framework
At ∼20–22∘ S, from west to east, the Andean margin
(Fig. 1) is constituted of (1) the subduction margin, including the
Peru–Chile Trench, the oceanward forearc, and the Coastal Cordillera that
reaches elevations >1 km and that corresponds to the former
Mesozoic volcanic arc; (2) the Atacama Bench or Central Depression, at an
altitude of ∼1 km, corresponding to a modern continental
forearc basin well expressed in the morphology and topography of northern
Chile; and (3) the strictly speaking Andean orogen (e.g., Charrier et al.,
2007; Mcquarrie et al., 2005; Oncken et al., 2006). The morpho-tectonic
units located west of the Andean orogen constitute the Marginal Block (i.e.,
the oceanward forearc, the Coastal Cordillera, and the Atacama Bench; Fig. 1) (Armijo et al., 2015, 2010).
At ∼20–22∘ S latitude, the Andes are characterized
by their largest width (>650 km), highest average elevation
(∼4–4.5 km above sea level, hereafter a.s.l., Fig. 1),
thickest crust (70–80 km, e.g., Heit et al., 2007; Tassara et al., 2006;
Wölbern et al., 2009; Yuan et al., 2000; Zandt et al., 1994), and
greatest total shortening (>300 km, e.g., Anderson et al.,
2017; Barnes and Ehlers, 2009; Eichelberger et al., 2013; Elger et al.,
2005; Kley and Monaldi, 1998; Mcquarrie et al., 2005; Sheffels, 1990). Here,
the Andean orogen is composed, from west to east, of (1) the Western
Cordillera, including the Cordillera Domeyko and the modern volcanic arc
(Fig. 1) (e.g., Armijo et al., 2015; Eichelberger et al., 2013; Garzione
et al., 2017; Mcquarrie, 2002; Oncken et al., 2006); (2) the Altiplano
Plateau, a high-elevation internally drained low-relief basin; (3) the
Eastern Cordillera (or Cordillera Oriental); (4) the Interandean zone; and
(5) the Subandean ranges, east of which the South American craton
underthrusts the Andes (e.g., Armijo et al., 2015; Isacks, 1988; Mcquarrie
et al., 2005; Oncken et al., 2006).
The building of the Andean mountain belt stricto sensu proceeded since the Late
Cretaceous–early Cenozoic at ∼20–22∘ S and
was associated with crustal shortening and thickening (e.g., Armijo et
al., 2015; Charrier et al., 2007; Decelles et al., 2014; Eichelberger et
al., 2013; Oncken et al., 2006; Jaillard et al., 2000; and references
therein). Based on regional syntheses (e.g., Armijo et al., 2015; Charrier
et al., 2007; Garzione et al., 2017; Horton, 2018; Mcquarrie et al., 2005;
Oncken et al., 2006), the across-strike growth of the orogen is summarized
as follows: (1) by the Late Cretaceous, the Mesozoic arc and back-arc basin
(formed during the early Andean cycle) were located at the position of the
present-day forearc, and most of the Andes showed mainly flat topography;
(2) by the Late Cretaceous–early Cenozoic, orogenic growth initiated primarily
along the western margin of the present-day Altiplano; (3) by
∼45–30 Ma, shortening vanished along the western flank of
the Andes and was transferred to the Eastern Cordillera; (4) by
∼25 Ma, deformation ended in the Eastern Cordillera and
migrated to the Interandean Belt; (5) from ∼10 Ma until
present, deformation within the Subandean Belt proceeded with the
underthrusting of the Brazilian Craton beneath the Andes. It is therefore
clear that the Andean shortening started along the Western Andes and
subsequently propagated eastward, progressively enlarging the orogen to form
the different cordilleras and the Altiplano Plateau in between.
Geological setting of the western flank of the Andes at ∼20–22∘ S
The Andean western flank is formed of three tectono-stratigraphic units at
∼20–22∘ S, aside from the present-day volcanic
arc. These units are hereafter described following a stratigraphic logic,
starting from the oldest and deepest units exposed to the east at higher
altitudes and following to the youngest units observed mostly to the west at lower
altitudes (Fig. 1). These units are (1) the Andean basement consisting of
metamorphic rocks of Precambrian and Paleozoic ages and (2) the folded
volcano-sedimentary deposits of Mesozoic age (Triassic–Cretaceous),
unconformably overlain by (3) less-deformed middle to upper Cenozoic (Oligocene–Quaternary) volcanics and sedimentary cover (Sernageomin, 2003).
Magmatic intrusions locally alter these different units and are mostly
Cenozoic. This only pictures the first-order structuration of the Western
Andean flank, as Mesozoic strata may be locally trapped between two
basement units, and Cenozoic layers may be unconformably overlying older
strata even to the east (Fig. 1). Laterally, in particular further
south (i.e., south of the city of Calama, ∼22∘27′),
the structural organization of the western flank of the Andes is more
complex, most probably because of the variable lateral structuration of the
earlier Mesozoic Andean basins, and the description proposed here does not
directly apply.
Stratigraphic and geologic background
The pre-Andean basement rocks formed during the late Proterozoic and
Paleozoic, when the Amazonian craton was progressively assembled from
various terranes (e.g., Charrier et al., 2007; Lucassen et al., 2000;
Ramos, 1988; Rapela et al., 1998). At the end of this period of subduction
and continental accretion, intense magmatic activity (volcanism and major
granite intrusions) welded together the basement during the late
Carboniferous to early Permian (Charrier et al., 2007; Ramos, 1988;
Vergara and Thomas, 1984).
The Mesozoic deposits (Triassic to Cretaceous), found today along the Western
Andean flank, formed in a proto-Andean arc and back-arc basin system during
the early period of the Andean cycle (e.g., Charrier et al.,
2007; Mpodozis et al., 1989). Marine and continental sediments are
interbedded with volcano-magmatic rocks (Aguilef et al., 2019;
Sernageomin, 2003). These Mesozoic units locally attain thicknesses up to
≥10 km (e.g., Buchelt and Cancino, 1988; Charrier et al., 2007;
Mpodozis et al., 1989).
A regional erosional surface called the Choja Pediplain
(Galli-Olivier, 1967) developed during the Eocene to early Oligocene
(∼50–30 Ma) (e.g., Armijo et al., 2015;
Victor et al., 2004; and references therein). Above this angular
unconformity, the up to ∼1600 m thick (Labbé et
al., 2019) Cenozoic deposits of the Altos de Pica Formation are composed of
continental clastic sediments, interbedded with volcanic layers
(Victor et al., 2004). The oldest documented age within the Altos
de Pica Formation is ∼24–26 Ma from dated ignimbrites
(Farías et al., 2005; Victor et al., 2004). From there,
an age of ∼27–29 Ma for the base of the formation is
inferred regionally when extrapolated to the basal erosional surface. The
youngest ignimbrites within the Altos de Pica Formation are dated at
∼14–17 Ma (Middle Miocene) (Vergara and
Thomas, 1984; Victor et al., 2004). From there and from other younger dated
ignimbrites (Baker, 1977; Vergara and Thomas, 1984),
Victor et al. (2004) deduced from stratigraphic correlations that
the development of the Altos de Pica Formation finished by ∼5–7 Ma (late Miocene) at ∼20–22∘ S.
Structural and kinematic context
The Paleozoic basement of the Western Cordillera is disrupted at places in
the form of various basement highs boarded by reverse faults (e.g.,
Haschke and Gunther, 2003; Henriquez et al., 2019; Puigdomenech et al.,
2020; Tomlinson et al., 2001) (Fig. 1) – not to be confused with the
north–south-trending strike-slip Domeyko Fault System, also called the West
Fissure System (e.g., Charrier et al., 2007; Reutter et al., 1996;
Tomlinson and Blanco, 1997b, a), which is east and out of
our field study area. At ∼20–22∘ S, various maps
describe west-vergent thrusts in overall structural continuity, bringing the
Paleozoic basement westward over folded Mesozoic units (Aguilef et al.,
2019; Haschke and Gunther, 2003; Sernageomin, 2003; Skarmeta and Marinovic,
1981). Using apatite fission-track dating, Maksaev and Zentilli (1999)
proposed significant exhumation of the basement units between 50 Ma and 30
Ma, possibly related to basement overthrusting. Older exhumation ages (Late
Cretaceous to early Cenozoic (U–Th) / He zircon and apatite ages) are, however,
provided by Reiners et al. (2015) for the Western Andean
basement at ∼21∘42′ S, but from only one sample and
without modeling. Together, these ages indicate that data remain missing to
better quantify the exhumation, uplift, and timing of deformation of the
basement thrusts reported along this part of the Western Andean flank.
Further west, a series of mostly west-vergent thrusts has been inferred,
essentially from seismic profiles. These thrusts result from the tectonic
inversion of the previous Mesozoic basins and affect the Mesozoic to
Cenozoic series (Fuentes et al., 2018; Martinez et al., 2021; Victor et
al., 2004; Armijo et al., 2015). Victor et al. (2004) determined
∼3 km of shortening of the syntectonic Altos de Pica
Formation layers, but they did not take into account the deformation of the
underlying more deformed Mesozoic units. Other authors propose limited
shortening on these older deeper layers (Fuentes et al., 2018;
Martinez et al., 2021), but the poor quality of the seismic profiles at
these depths renders these interpretations quite tenuous and disputable.
Haschke and Gunther (2003) estimated that >9 km of
shortening across the western flank in the outcropping Sierra de Moreno area
(∼21∘45′ S) occurred since the Late Cretaceous to
Eocene on a west- and east-verging thrust system. Whether these various
faults are connected at depth to an east-dipping master fault (Armijo
et al., 2015; Haschke and Gunther, 2003; Victor et al., 2004) or whether
they are steeply dipping single planar faults (Fuentes et al.,
2018; Martinez et al., 2021) remains debated: these considerations are
documented nowhere by data and only rely on a priori concepts or structural
reasonings at a larger regional scale. It follows that even if published
data document the existence of various faults along the Western Andean front
at ∼20–22∘ S, their geometry, kinematics, and total
amount of shortening have not yet been satisfactorily evaluated.
Data and methods
Unconformable slightly deformed middle to upper Cenozoic clastic sediments and
ignimbrites commonly hide the folded Mesozoic layers and their contact with
the basement (Fig. 1). Field investigations are limited to the few sparse
areas where the erosion of the Cenozoic cover has exposed the underlying
structures (Aguilef et al., 2019; Sernageomin, 2003). In this
study, we focus on two relatively accessible outcrop sites (Fig. 1): (1) one is at ∼21∘30′ S, where the Paleozoic basement
overthrusts Mesozoic units (Skarmeta and Marinovic, 1981). This
zone will be referred to as the Pinchal area (next to Cerro Pinchal, 4193 m a.s.l.). (2) The other is at ∼20∘45′ S, where folded Mesozoic
units can be observed. This zone is hereafter named the Quebrada Blanca area
after its largest canyon.
Available data
The most detailed – even though large-scale – existing geological map for
the Pinchal area is the 1:250000 Quillagua map (Skarmeta and
Marinovic, 1981). For the Quebrada Blanca area, the recent 1:100000
Guatacondo map (Blanco and Tomlinson, 2013) provides detailed and updated
information on the stratigraphy and structure. There, the structure of the
folded Mesozoic rocks has been preliminarily mapped and qualitatively
described by other authors (Armijo et al., 2015; Blanco and Tomlinson,
2013; Fuentes et al., 2018).
Enhanced cartographic details can be deduced from high-resolution satellite
imagery. We use Google Earth imagery (Landsat 7, DigitalGlobe) whose
resolution varies from a few meters to a few tens of meters depending on the
zones. In addition, this work benefits from very high-resolution imagery
from the European Pléiades satellites. Using the MicMac software suite
(Rosu et al., 2014; Rupnik et al., 2016), we calculate high-resolution
digital elevation models (DEMs) from tri-stereo Pléiades imagery with a 0.5 m resolution. These
DEMs are down-sampled to a resolution of 2 m to enhance data treatment and
calculations (e.g., stratigraphic projection and image processing). Relative
vertical accuracy may reach ∼1 m, depending on local slope.
Field observations were acquired during two field surveys in March 2018 and
January 2019. Difficult accessibility and field logistics in the remote and
desert Pinchal area only allow detailed complementary field observations of
a relatively limited area. Observation points and the off-road track
followed to reach our field site in the Pinchal area are provided in the
Supplement.
Structural maps
To establish structural maps, we conducted 3D mapping of stratigraphic layers
after Armijo et al. (2010) and Riesner et al. (2017). Layers are traced and correlated on Google Earth satellite images.
The so-obtained georeferenced traces are projected on the DEM-derived
topographic map to obtain their altitude and on geological maps for
stratigraphic referencing. Field observations allow ground verifications and
provide supplementary details, such as minor thrusts and folds, as well as the
observation of polarity criteria or local dip angles.
The approach used here is mainly limited by local geological complications.
Continuous mapping of Mesozoic strata is locally complicated where incision
of Cenozoic strata is limited, where magmatic intrusions and associated
hydrothermalism alter the structural geometries, where layers with no
well-expressed bedding such as marls are present (ex: Pinchal area), or
where landslides or recent sediment deposits hide the underlying deformation
pattern. Therefore, geometrical observations and detailed mapping of the
structures may be locally difficult and in some zones impossible. These
difficulties cause some uncertainties in precisely correlating mapped
layers but only result in limited metric to decametric errors and do not
modify our large-scale (km) results and interpretations.
Structural cross-sections
We use structural measurements, field observations, and the obtained
structural map to propose surface cross-sections across the two investigated
areas.
In the Pinchal area, because of limited canyon incision, marls, and frequent
blanketing of the structures by Cenozoic cover, we build our structural
surface cross-sections mainly from field observations (strike and dip
angles, polarity criteria, first-order stratigraphic column), with
additional information taken from satellite imagery.
In contrast, in the Quebrada Blanca area, we build our surface cross-section
mostly from mapping on satellite imagery. Here, we follow the approach
proposed in Armijo et al. (2010) and described in detail in
Riesner et al. (2017). The mapped georeferenced horizons are
projected on the high-resolution Pléiades DEMs. Using a 3D modeler, we
project these layers along swath profiles chosen where Mesozoic strata crop
out the best, where folds appear cylindrical, and where topographic relief is
most significant. This approach allows for more precise
large-scale structural geometries by averaging the usual local minor
variations in strikes and dips that derive from multiple field measurements.
From there, we successfully obtain the overall sectional geometry of folded
layers, and by comparing with the structural map, we determine the
approximate locations of the major synclinal and anticlinal axes. By
respecting the classical rule of constant layer thickness, we derive fold
geometries.
We recognize the difficulty of unambiguously correlating stratigraphic
layers in some cases and the fact that layers may not keep constant
thicknesses. As local topographic relief is reduced to a few hundred meters
at most, the construction of surface cross-sections is mostly restricted to
extrapolating derived average surface dip angles at depth. However, the main
limitation relies on the fact that we can only draw the subsurface
sectional geometry of folds from surface geology. Indeed, we do not have
any constraints from field observations on the geometry of the associated
thrusts at depth, nor on the footwall structure of these thrusts. We propose
a possible structural interpretation at depth and discuss its implications.
Crustal shortening and kinematic modeling
We apply a line-length balancing approach to the obtained surface
cross-sections to determine shortening related to folding only. This result
is independent of the geometry of the associated faults at depth, does not
account for penetrative deformation or for slip on underlying
faults, and stands as a conservative minimum here.
From surface geology, we have no precise indication about the structure and
geometry of faults and layers at depth, in particular within the footwall of
the inferred faults. To make a step forward in our estimates of crustal
shortening, we consider the simplest structural geometries where underlying
thrusts are parallel to the anticlinal backlimbs and root at least at the
base of the folded series. From there, we model anticlinal geometries using
a numerical trishear approach (e.g., Allmendinger, 1998;
Erslev, 1991). We use the code FaultFold Forward (version 6)
(Allmendinger, 1998) in order to jointly model thrust displacement
and anticlinal folding. Trishear models the deformation distributed within a
triangular zone located at the tip of a propagating fault (Erslev,
1991). This forward modeling relies on testing a set of parameters, namely
the position of the fault tip, the angle of the propagating fault ramp, the
slip on the fault, the propagation-to-slip ratio (P / S) of the fault, the
angle of the triangular zone at the tip of the fault where distributed
deformation occurs (i.e., the trishear angle), and the inclined shear angle
controlling the backlimb kinematics. We assume here the case of linear
symmetric trishear to keep models as simple as possible, meaning that
folding of the backlimb occurs parallel to the fault. Initial layers are
assumed to be sub-horizontal, with a slight eastward tilt (3∘ E at
Quebrada Tambillo, 2∘ E at Quebrada Blanca) as expected in the
initial Andean basin. We tested various combinations of parameters within
the range considered in the literature (e.g., Allmendinger, 1998;
Allmendinger and Shaw, 2000; Cristallini and Allmendinger, 2002; Hardy and
Ford, 1997; Zehnder and Allmendinger, 2000) and with regards to our field
observations such as the possible geometry of faults at depth as constrained
from surface geology. Parameters are adjusted by trial and error to visually
fit observed and interpreted structural geometries. By adding sedimentary
layers at various steps during ongoing deformation, with initial geometries
similar to the present-day regional topographic slope, we model syntectonic
deposition and subsequent deformation in order to reproduce deformation of
Cenozoic layers. Additional information on trishear modeling (initial
conditions, values of tested parameters, etc.) is provided later and in
the Supplement (Text S1, Tables S1–S3). We recognize that parameters
of our preferred model may not be unique. This is not expected to significantly impact estimated total shortening, as this result depends mostly on the
modeled cross-sections (and therefore on the structural interpretation to be
modeled) rather than on the preferred set of parameters. For instance, we
cannot discard the possibility that faults are steeper and root deeper. If
this were the case, crustal shortening would be lower.
Deformation is expressed in terms of shortening (in kilometers) and of relative
shortening at the scale of the investigated sites (in percent). Relative
shortening is hereafter defined as the ratio of the estimated shortening by
the initial length of the undeformed section.
Basement thrust and deformed Mesozoic series at Pinchal (∼21∘30′ S)
Because our observations are in contradiction with previous stratigraphic
and structural interpretations of the folded Mesozoic series, we hereafter
describe our field observations in detail. We subsequently discuss and
compare them to previous interpretations and propose a solution reconciling
these observations with regional stratigraphic knowledge.
Landscape field view of the Pinchal area depicting the
main tectono-stratigraphic units. The Paleozoic (Pz) basement clearly stands
out in the background, characterized by its darker color and higher
elevations. The Mesozoic (Mz) series in the central part and in the
foreground bear a marine part and a volcano-detrital part, respectively,
separated by an outstanding calcareous (Calc.) crest. Unconformable Cenozoic
erosional surfaces, with limited fluvial deposits, can also be observed.
First-order stratigraphic column of the Pinchal area.
Data are derived from field observations, mainly obtained along Quebrada
Tania (Figs. 4 and 5a). The column only depicts the stratigraphic
succession, and thicknesses are not to scale (see text for information on
thicknesses). The description of Cenozoic units is completed here based on
the work of Victor et al. (2004). The color code is in line with maps
(Figs. 1, 4, and S14) and cross-sections (Fig. 5). The Paleozoic basement
overthrusts folded Mesozoic series along the Pinchal Thrust so that part of
the deeper and older Mesozoic series may be missing here. See main text and
Figs. S1–S12 for detailed sedimentological
descriptions.
Stratigraphic observations
In the landscape (Fig. 2), the three main tectono-stratigraphic units are (1) the metamorphic basement, (2) the Mesozoic sedimentary series (with a
continuum from continental upward to marine facies), and (3) the continental
Cenozoic cover. From field observations, we propose a first-order
stratigraphic column (Fig. 3). Field pictures of sedimentary formations
are provided in the Supplement (Figs. S1–S12) to complement the
forthcoming descriptions. We acknowledge that we do not have any constraint on the
absolute ages of these series, but the relative stratigraphic ages are
deduced from the kilometer-scale structural geometry and from clear
sedimentary or structural polarity criteria observed in the field.
Thicknesses are inferred only locally, and thickness variations cannot be
excluded.
The Paleozoic basement (Fig. S1) dominates the eastern part of the Pinchal
area and is composed of mainly coarse-grain granodiorites and diorites, as
well as metamorphic rocks comprising gneisses, migmatites, and mica–schists,
consistent with previous descriptions in the area (Skarmeta and
Marinovic, 1981).
The older part of the outcropping Mesozoic series consists of continental
deposits, with a high content of Paleozoic lithics and volcano-clastic and
tuffitic low-rounded conglomerates, of greenish, beige, and brownish colors
(Fig. S2). Clast sizes vary from a few millimeters to a few decimeters. At
places, these rocks bear sedimentary polarity criteria such as
grain sorting, cross-bedding, and tangential beds (Fig. S3). In the eastern
part of the Pinchal area, we locally observe below this series dark green
detrital pelites (lutites) (Fig. S4). On the basis of petrographic and
sedimentological correlations, these detrital Mesozoic sediments recall units mapped as Triassic north of the Pinchal zone
(between 21–21∘30′ S) in the Quehuita
area (Aguilef et al., 2019).
In paraconformity, a characteristic limestone layer marks the beginning of a
marine sequence, evidencing a marine transgression. We refer to this layer
as the “calcareous crest” as it is prominent in the landscape (Fig. 2) and
can be easily used as a reference layer in the field or on satellite images.
The base of the calcareous crest is characterized by silex layers or nodules
(Fig. S5). Up-section, numerous stromatolites (Fig. S6) and bivalves
(Fig. S7) are found. Its thickness varies from a few meters (less than
10 m) in the eastern part to ∼10–20 m to the west.
The calcareous crest is overlain by thin-bedded (centimeters–decimeters) limestone layers of
rose–beige color (Fig. S8) over a thickness of ∼50–100 m.
Going up-section, the marine series becomes more marly, more beige, and with
fewer limestone layers, evidencing a deeper marine paleo-environment.
Belemnite fossils were encountered in the lower part of this
limestone-to-marl sequence. Characteristic calcareous oval concretions of
variable diameter (centimeters to meters) (Fig. S9) are pervasive at the transition
from marly limestones to marls. The marls bear ammonites, which we have not
precisely identified. These ammonites could be Perisphinctes, Euaspidoceras,
Mirosphinctes, and Gregoryceras, according to the notice of the Quillagua
geological map (Skarmeta and Marinovic, 1981) if applicable here.
In this case they would be of Middle Jurassic age (Bajocian to Callovian).
The series from the thin-bedded limestones to the top of the beige marls is
∼200 m thick along one of the canyons (Quebrada Tania).
Structural map of the Pinchal area (location on Fig. 1). White thin lines highlight Mesozoic layers mappable on satellite images.
The thick blue line depicts the calcareous crest, which is used as a marker
layer (Fig. 2). A–A' and B–B' segments locate the surface cross-sections
of Quebrada Tania and Quebrada Martine, respectively (Fig. 5a–b). In the
case of the Quebrada Tambillo cross-section, a topographic swath profile was
used along C–C' (dashed box). The fold axes are relatively well defined for
the synclinal fold but less well constrained for the anticlinal fold
because the latter is only observable along Quebrada Tambillo. Field pictures are numbered
according to the corresponding figures. Background hillshaded DEM produced
from tri-stereo Pléiades imagery. Q: Quebrada (Spanish word for
“canyon”).
Up-section, the beige marls become more calcareous again, with thin limestone
layers (Fig. S10). Finally, this marine sequence ends with black marls
containing layers of beige sandstones (millimeters to a few centimeters – rarely decimeters – thick)
(Fig. S11), indicative of a detrital component in a probable deep-seated
basin, comparable to the “flysch” series of the Alpine basins
(Homewood and Lateltin, 1988). This unit is hereafter called “black
flysch” and has a minimum thickness of ∼50 m.
Continental-clastic Cenozoic deposits (Altos de Pica Formation)
unconformably overlie this folded Mesozoic series over the Choja erosional
surface (Galli-Olivier, 1967; Victor et al., 2004) (Figs. 2
and 3). They are mainly composed of alluvial fan deposits sourced from the
mountain front immediately to the east, locally interlayered or covered by
ignimbrites. We encountered red arenites at the base of the Cenozoic series
in the western part of the Pinchal area (Fig. S12). The age of the oldest
sedimentary deposits above this erosional surface is regionally inferred to
be ∼27–29 Ma (Victor et al., 2004) (see also
Sect. 2.2).
Surface cross-sections along (a) Quebrada Tania, (b) Quebrada Martine, and (c) Quebrada Tambillo (sections A–A', B–B', and C–C' in
Fig. 4, respectively). Reported dip angles have been measured in the
field or deduced from 3D mapping. Faults are outlined in black and dashed
when they are only observable at a local spatial scale. Only larger faults
(continuous lines) are mapped in Fig. 4. Fold axes are depicted above
their surface trace based on our field observations, and their orientation
illustrates the deduced orientation of the corresponding axial plane. Gray
numbers with arrows point to field pictures and indicate the associated
figure. In the case of Quebrada Tania (a), the sedimentary polarity
criterion (β) indicated to the west of the section has been observed
∼1 km further downstream than reported here. For Quebrada
Martine (b), note the stripe of continental Mesozoic rocks trapped
between two strands of the Pinchal Thrust. The subsurface interpretation from
surface observations is reported with transparent colors in the case of
Quebrada Tambillo (c). Note the different spatial scales of the three
sections. PT: Pinchal Thrust.
Structural observations
A structural map of the area is built after satellite and field observations
(Fig. 4). Two approximately east–west cross-sections show detailed
surface observations along two accessible representative canyons: Quebrada
Tania and Quebrada Martine (Fig. 5a, b). The Quebrada Tambillo incises
deeper into folded units so that surface structural observations can be
further extrapolated at depth (Fig. 5c).
The easternmost part of our study area is marked by a west-vergent thrust
bringing the metamorphic basement over folded Mesozoic units (Figs. 2, 4,
and 6a). This basement thrust is hereafter named the Pinchal Thrust. The
C–S fabric (Cisaillement–Schistosité) observed within the thrust shear
zone indicates top-to-the-west thrusting (Fig. 6b). The Pinchal Thrust
roughly follows a north–south direction (Fig. 4). This contact often
resumes to a single basement thrust (Fig. 5a, c), but may also show local
geometrical complexities, with secondary thrusts and branches, eventually
involving basement with stripes of trapped Mesozoic units, as along Quebrada
Martine (Fig. 5b).
Folded Mesozoic units are observable west of the Pinchal Thrust (Figs. 2
and 4). From east to west, an asymmetric and overturned syncline (Fig. 7a)
is followed by a relatively symmetric anticline (Fig. 7b).
Field characteristics of the Pinchal Thrust.
(a) Field view of the Pinchal Thrust (PT), with dark grayish Paleozoic
basement thrusting over the greenish folded Mesozoic units. Reddish rocks on
the hanging wall to the east-northeast correspond to the thrust shear zone
(hatched area in picture). Location in Fig. 4; note that the scale here is
only approximative. The non-interpreted photograph can be found in the
Supplement (Fig. S15).
(b) Shear band with characteristic C–S fabric (for
Cisaillement–Schistosité) indicative of top-to-the-west thrusting.
Observation within the shear zone in the hanging wall of the Pinchal Thrust
(deformed metamorphic basement).
The eastern limb of the syncline is inverted and locally highly faulted and
folded (Figs. 5 and S13). Within this inverted limb, the series goes
westward (and up-section) from sheared lutites beneath the Pinchal Thrust,
followed by Mesozoic detrital series with conglomerates, to the Mesozoic
marine series from the calcareous crest up-section to the marly limestones.
The overturned strata dip steeply (50–70∘ E). Penetrative
small-scale deformation is observed pervasively within the marine Mesozoic series in the form of numerous local small folds, kinematically
indicative of an inverted fold limb (used here as a structural polarity
criterion) (Fig. S13a), and local secondary shear zones and thrusts
(Fig. S13b).
Going westward, as observed in detail along Quebrada Tania (Fig. 5a), the
eastern part of the black flysch bears small-scale folds characteristic of
an inverted fold limb, whereas normal limb folds (also used here as
structural polarity criteria) are observed slightly further west: the axis
of the overturned west-vergent syncline therefore passes through the black
flysch. Part of the Mesozoic series is missing, as overthrusting within the
flysch and (marly) limestones is observed frequently along Quebrada Tania
(Fig. 5a). The overturned syncline is therefore broken by a secondary
thrust fault striking approximately parallel to the Pinchal Thrust and roughly coinciding with the synclinal
fold axis (Figs. 4–5). Westward, the normal western limb of the syncline
encompasses the whole Mesozoic series from the black flysch down-section to
the Mesozoic volcano-detrital series, with more gentle dip angles
(20–40∘ E) (Figs. 2 and 5). Penetrative deformation is
observed to be limited here.
Field pictures of the two major folds within the Pinchal
area (location on Fig. 4). Non-interpreted photos can be found in the
Supplement (Fig. S16). Scales are only approximative because
of perspective.
(a) Panoramic view over the northeastern part of the Pinchal area. The
Paleozoic basement (red crosses) overthrusts the Mesozoic units (blue and
purple horizons) along the Pinchal Thrust (red line with triangles). The
topographic low locates the synclinal axis. The calcareous crest on both
sides is highlighted by the thick blue lines.
(b) Panoramic view along Quebrada Tambillo in the southern part of the
Pinchal area. The ∼200 m deep incised canyon reveals the
geometry of the large western anticline affecting Mesozoic layers (purple)
underneath the unconformable Cenozoic strata (yellow). The fold axis (black
line) probably coincides with an approximately vertical fault that is clearly
observable on satellite imagery. Note also the repetition of smaller folds
with westward-decreasing amplitude and wavelength discernable beneath the
westward-thickening Cenozoic growth strata to the right of the picture. The
Mesozoic calcareous crest (blue) and the Paleozoic basement (red crosses)
over the Pinchal Thrust (red) appear in the far eastern background.
The continental Mesozoic layers of the normal limb of the syncline flatten
toward the west. The section along Quebrada Tambillo (Fig. 5c) shows a
broad, overall symmetrical, anticlinal fold (Fig. 7b). Its axial plane is
steep, dipping ∼80∘ E. Smaller, secondary folds
with westward-decreasing wavelength and amplitude are found at the western
front of this large anticline. Field logistics did not permit further
detailed structural observations.
The folded Mesozoic units are unconformably covered by sheet-like,
river-incised Cenozoic fluvial deposits, forming aggradational terraces
deposited above erosional surfaces at different elevations, of varying
spatial extent and probably of different ages (Fig. 2). The majority of
these erosional surfaces show a westward tilt (Fig. 5c). Further west, the
Cenozoic deposits become thicker and bury the westward extent of the folded
Mesozoic units. Westward thickening of the Cenozoic
layers is observed along Quebrada Tambillo and indicates growth strata at
the front of the western anticline (Figs. 5c and 7b).
Comparison to previous stratigraphic and structural interpretations
In the Pinchal area, a basement thrust was reported in the 1:250000
Quillagua geological map (Skarmeta and Marinovic, 1981). In this
map, the Mesozoic units are interpreted as pertaining to the Jurassic
Quinchamale formation, deposited in a back-arc basin context and composed of
an Oxfordian (∼157–163 Ma) and a younger Kimmeridgian
(∼152–157 Ma) sub-unit. Based on this age interpretation and
relying on a regionally established Mesozoic stratigraphy where marine
sequences are followed upward by younger clastic deposits,
Skarmeta and Marinovic (1981) interpreted the main structure of
the Pinchal zone as an anticline.
Our field investigations confirm the existence of a basement thrust but
contradict the earlier interpretation of the folded Mesozoic series and of
the local Mesozoic stratigraphy. Even though we do not know the absolute
ages of the folded sedimentary series, our structural and sedimentary field
observations allow for clearly constraining the relative stratigraphic ages
of the folded Mesozoic units, from either structural or sedimentary polarity
criteria, and unambiguously indicate that detrital continental units are
stratigraphically below a marine sequence here (Fig. 3). In the case that
the marine strata are Jurassic in age from their likely fossiliferous
content, the older continental clastic units could be Triassic by
comparison to recent observations not far from the Pinchal area
(Aguilef et al., 2019).
Given this, even though the Pinchal stratigraphic sequence may look in
contradiction with the regionally known stratigraphy, it may rather be
viewed as complementary: the detrital component observed below marine series
may be older than the Jurassic and Cretaceous marine-to-continental upward
succession that has been well described regionally. In this sense, the
Pinchal area may provide a key outcrop to refine our knowledge of older
series, possibly Triassic.
Detailed field pictures of the various stratigraphic and sedimentological
observations are provided in the Supplement for reference. In any
case, we recall that relative ages are only needed here for the scope of
this study to decipher the general structure and deformation pattern.
Structure of the folded Mesozoic series at the Quebrada Blanca
(∼20∘45′ S)Stratigraphy of the Quebrada Blanca area
The stratigraphy at ∼20∘45′ S is well described in
the Guatacondo geological map (Blanco and Tomlinson, 2013). Unlike in
Pinchal, basement rocks do not crop out in the investigated zone (Fig. 8),
but larger-scale maps (e.g., Sernageomin, 2003) show Paleozoic
basement units further east and higher in the topography (Fig. 1).
Structural map of the Quebrada Blanca zone, refined from
Armijo et al. (2015) (location on Fig. 1). Colored lines report
mappable layers. For visibility, only major, well-correlated layer traces
are represented here. Black boxes locate where mapped layers were considered
and projected for the construction of the structural east–west surface
cross-section (Fig. 9). The A–B section corresponds to the topographic
profile used for this cross-section. Strike and dip measurements are
extracted from 3D mapping (see Sect. 3.3) or observed in the field. Strike
symbols without dip value are derived from satellite imagery. Field pictures
are located (with view direction) and numbered according to the associated
figure. Ages from uranium–lead (U / Pb) radioisotope dating on zircon are
taken from the Guatacondo geological map (Blanco and Tomlinson, 2013).
Letters C, D, E, F, G, and H to the northeast (within the folded Chacarilla
and Majala formations) report the layers illustrated on Fig. S17 in
the Supplement. Background hillshaded DEM produced from tri-stereo
Pléiades imagery. Cz: Cenozoic; K: Cretaceous; Jr: Jurassic; Q: Quebrada.
East–west surface cross-section of the Quebrada Blanca
site, established from the projection of selected well-expressed layers
mapped on satellite imagery. APF: Altos de Pica Formation.
(a) Observations reporting the geometry of projected layers and associated
dip angles, together with their stratigraphic ages (color code).
(b) Subsurface interpretation and extrapolation of observations.
(c) East–west subsurface section A–B based on (a) and (b). Interpretation
at depth is indicated with transparent colors, in contrast to surface
observations. Extrapolation above the topographic surface is drawn with
dashed lines. Ages from uranium–lead (U / Pb) radioisotope dating on zircon
are taken from the Guatacondo geological map (Blanco and Tomlinson,
2013). The ∼27–29 Ma age of the basal deposits of the Altos
de Pica Formation is derived from regional considerations (Victor
et al., 2004).
The Mesozoic units of the Quebrada Blanca are of Jurassic to Cretaceous age
(Blanco and Tomlinson, 2013). They were deposited in a back-arc
basin context in successive transgression–regression sequences
(Charrier et al., 2007) and are subdivided into three formations:
(1) the late Oxfordian Majala Formation, a clastic unit of sandstones,
shales, and stromatolitic limestones of transitional marine origin; (2) the
Late Jurassic–Early Cretaceous Chacarilla Formation, a fluvial clastic
sequence; and (3) the Late Cretaceous Cerro Empexa Formation, an andesitic
volcanic and continental sedimentary unit (Blanco and Tomlinson, 2013;
Blanco et al., 2000, 2012; Dingman and Galli, 1965; Dingman
and Galli Olivier, 1962; Tomlinson et al., 2001). The Majala and Chacarilla
formations both bear detritic reddish and beige sediments. The Cerro Empexa
Formation appears grayish and massive in the field. In the Quebrada Blanca
area, uranium–lead (U / Pb) dated zircons from this formation bear ages
between ∼75 and ∼68 Ma (Blanco and
Tomlinson, 2013; Blanco et al., 2012; Tomlinson et al., 2015) (Fig. 8).
Field picture of the western limb of the western
anticline in the Quebrada Blanca area. Non-interpreted photographs are
provided in the Supplement (Fig. S18). Location in Figs. 8 and 9.
(a) Series of folds with westward-decreasing amplitude and wavelength
(hundreds to tens of meters) observed at the front of the western anticline.
(b) Detailed view of the westernmost outcropping small-scale anticlines,
located in (a) by the black box.
Magmatic intrusions and hydrothermalism occur locally and hide the eastern
continuation of the folded Mesozoic series. Some of these intrusions are
dated by uranium–lead (U / Pb) on zircons at ∼44 Ma (Blanco
and Tomlinson, 2013) (Fig. 8).
The Cenozoic deposits of the Altos de Pica Formation here also overlie the
Mesozoic series, over the Choja Pediplain angular unconformity (see also
Sect. 2.2). The age of the basal deposits of the Altos de Pica Formation
is regionally estimated to be ∼27–29 Ma (e.g., Victor
et al., 2004).
Structural observations
Although the cartography of the folds is complicated by the blanketing
Cenozoic cover (notably in the west and south) and by magmatic intrusions
and hydrothermalism (particularly to the east) (Fig. 8), three large-scale
folds are observable: a wide syncline in the center (Higueritas syncline)
bounded by two anticlines to the west (Chacarilla anticline) and east (fold
names from Blanco and Tomlinson, 2013; Fuentes et al., 2018). The
scale of these folds is multi-kilometric (Fig. 8). Both anticlines are
asymmetric: they have steeper western limbs (dip angles vary between
∼50 and 80∘ W), whereas their eastern limbs have
more gentle dip angles (varying between ∼20 and 50∘ E) (Fig. 9). Despite the fact that the eastern
flank of the eastern anticline is widely hidden by magmatic intrusions and
hydrothermalism, its southern part is clearly observable in the field (Fig. 8). The central Higueritas syncline is wider and more symmetric, with dip
angles of ∼40–50∘ on both limbs. The anticlines
involve the Majala and Chacarilla formations, while the core of the syncline
bears the Cerro Empexa Formation. Overall, the documented folds show a clear
west vergence (Fig. 9c).
In the field, we observe small-scale deformation within both anticlines
(Fig. 9). A series of anticlines with westward-decreasing amplitude and
wavelength (of a few tens to a few hundreds of meters – to be compared to
the ∼4 km wavelength of the main anticline) is observable on
the western edge of the Chacarilla anticline (Figs. 9c and 10). In the
field, at least one of these small-scale folds is affected by a minor
thrust. Additionally, within the eastern large-scale anticline, a
thrust-affected small-scale fold is observed (Figs. 9c and S17) and
confirms the west vergence at this smaller scale.
The Cenozoic detrital units are unconformably deposited above the folded
Mesozoic series. Thin sheet-like river-incised Cenozoic surfaces remain in
the central part, becoming more dominant to the south and west (Fig. 8).
These superficial erosional surfaces show an overall westward tilt (Fig. 9). Westward thickening of the Cenozoic layers deposited above the erosional
Choja surface is clearly observed at the front of the western anticline
(Fig. 10) and reveals the presence of growth strata.
Kinematics of shortening of the Pinchal and Quebrada Blanca areasTiming of deformation
The projection of mapped strata indicates that the Mesozoic series is
overall concordant in Quebrada Blanca (Fig. 9). The cross-section of the
Guatacondo map (Blanco and Tomlinson, 2013), however, suggests the presence
of a minor angular unconformity (<10∘) at the base
of the Cerro Empexa Formation, not observed here from our large-scale
high-resolution
mapping. As this local unconformity does not produce any major change in the
geometry of layers from Jurassic to Cretaceous, we consider it to be minor,
in particular with respect to the main large-scale folding documented here.
Folding therefore mostly postdates the deposition of all these series. In
the Quebrada Blanca area, the youngest folded layers of the Cerro Empexa
Formation bear U–Pb ages of 68.9±0.6 and 68±0.4 Ma
(Blanco and Tomlinson, 2013) (Figs. 8 and 9c). We can therefore
conclude that the main deformation episode postdates ∼68 Ma,
even though we cannot exclude earlier but minor deformation when compared to
the observed large-scale folding (Fig. 9c).
Magmatic intrusions dated at ∼44 Ma intrude the folded
Mesozoic units and appear cartographically unaffected by folding
(Blanco and Tomlinson, 2013) (Fig. 8). This suggests that the major
part of the folding occurred during the ∼68–44 Ma time
interval. However, without additional observations of the deformation – or
not – of these intrusions (geometry of the contact with surrounding host
units, mineral deformation, etc.), we cannot unequivocally draw conclusions from this simple cartographic observation.
Even though we suspect that the deformed series of the Pinchal zone are
Triassic to Jurassic, we do not have any absolute ages of the folded units.
Therefore, we postulate that the main deformation here also postdates
∼68 Ma by analogy to our observations at the Quebrada Blanca.
Cross-sections and kinematics of folding
of the Quebrada Blanca and Pinchal zones. Cross-sections are built from
field observations considering that faults root at the base of the folded
series. Modeling was performed with FaultFold Forward v.6
(Allmendinger, 1998). Intermediate stages of the trishear modeling
are reported in Figs. S19 and S20 for the cross-sections of Quebrada
Tambillo and Quebrada Blanca, respectively. Model parameters are reported in
Tables S2–S3. (a–c) Proposed cross-sections and final stages of our preferred trishear
models in the case of (a) the Quebrada Blanca area and (b) the Quebrada
Tambillo (Pinchal area) shown here at the same scale as (a). (c) Detailed
and enlarged view of our results for the Quebrada Tambillo. Note the large
scale difference between the sections of the two investigated sites (a, b).
Thicker lines outline model results, while transparent lines and colors
refer to the proposed cross-sections. Black lines report the modeled thrusts,
and horizontal arrows report the modeled total shortening. PT: Pinchal
Thrust.
(d) Shortening vs. time, as deduced from trishear modeling of the western
anticlines of the Quebrada Blanca and Pinchal areas, and the ages of
deformed layers. The three temporal benchmarks correspond to the age of the
youngest folded Cretaceous (Kr) unit (∼68 Ma), to the age of
magmatic intrusions possibly postdating folding (∼44 Ma),
both derived from the Guatacondo geological map (Blanco and Tomlinson,
2013), and to the ∼29 Ma age of the oldest Cenozoic layer of
the Altos de Pica Formation (APF) (Victor et al., 2004). It is
possible that most deformation occurred prior to ∼44 Ma
(steeper lines). Our results underline two phases of deformation, with a
slowing-down of deformation since ∼29 Ma at least, possibly
even before.
The folded Mesozoic units are unconformably covered by the Cenozoic Altos de
Pica Formation at both investigated sites. This is also the case for the
Pinchal Thrust and secondary thrusts at a few places in the Pinchal zone
(Fig. 4). The presence of growth strata at the front of the westernmost
anticlines in both study areas, over the erosional Choja Pediplain, suggests
that some deformation proceeded after ∼29 Ma during
deposition of the Altos de Pica Formation. However, the deformation recorded
by folded Mesozoic units appears of greater intensity than that of the
Cenozoic growth layers (Figs. 5c and 9c).
Given this, we propose that the main folding of the Mesozoic layers
documented here can be bracketed to a maximum time span of ∼40 Myr, sometime between ∼68 and ∼29 Ma,
with additional relatively minor deformation after ∼29 Ma.
Possibly, the main deformation period could be shorter (∼24 Myr at most), sometime between ∼68 and ∼44 Ma, with minor shortening after the Eocene intrusions. In the case of the
Pinchal Thrust, we can only propose from our observations that thrusting
took place prior to ∼29 Ma.
Structural interpretationsSection across the Pinchal zone
The cross-section proposed along Quebrada Tambillo (Figs. 11b–c)
summarizes our preferred interpretation of the subsurface structural
geometry of the Pinchal area.
Considering that the folds west of the Pinchal Thrust developed above
underlying thrusts is a reasonable and classical assumption (Davis et
al., 1983; Suppe, 1983). These thrusts have dip angles parallel to the
layers forming the backlimb of the overlying anticlines (as proposed in
Fig. 11b–c) or can be steeper. They are expected to root at least at the
base of the outcropping folded Mesozoic series (as proposed in Fig. 11b–c) or deeper. From there, it can be extrapolated that the thrusts root
at least 2 km beneath the topographic surface (i.e., at ∼0.2 km a.s.l.), assuming that the layer thickness is constant over our study
area. We cannot discard the possibility that the thrusts root deeper and are
steeper from surface geology alone. To the west of our field area, at the
front of the anticline, the small-scale folds with westward-decreasing
wavelength and amplitude (Fig. 7b) are interpreted as the possible
expression of disharmonic folding within the forelimb of the anticline
and/or of a thrust ramping up toward the subsurface (Fig. 11b–c).
Using the simplified geometry of layers along Quebrada Tambillo (Figs. 5c
and 11c), line-length balancing results in a minimum of ∼1 km
of shortening absorbed by folding only, from the Pinchal Thrust to the front
of the anticline (Table 1). Because of the pervasive presence of small-scale
folding and thrusting (Figs. 5a–b and S13), in particular within the
inverted limb of the overthrusted syncline, this estimate represents a
minimum value. A significant – but unconstrained – amount of shortening by
folding is surely to be added. In addition to folding, the offsets on the
interpreted underlying thrusts are to be considered when quantifying
shortening. However, the thrust offsets largely depend on the interpreted
thrust geometry and on the structure of the footwall, which cannot be
precisely determined from surface geology.
Based on the dip angle of the C–S fabric of the shear zone (Fig. 6b) and
on the mapping of the thrust on satellite imagery, we estimate that the
Pinchal Thrust dips ∼40∘ E in the
near surface, locally less, such as along Quebradas Tania and Martine
(Fig. 5a–b). All secondary strands of the Pinchal Thrust are expected to
root at depth to the main shear zone. The secondary thrust breaking the
core of the syncline, roughly parallel to the Pinchal Thrust (Fig. 4), may
be described as an out-of-the-syncline thrust (Mitra, 2002) and probably
also connects onto it at depth. Similar reasoning is proposed for all
small-scale thrusts and décollements observed within the inverted
synclinal limb. The minimum thickness of the Mesozoic series is estimated to be
∼2.2 km from the normal limb of the syncline along Quebrada
Tambillo. Thus, the strict minimum exhumation of
the basement can be considered to equally be ∼2.2 km. Assuming a constant
40∘ E dip angle and taking exhumation as a proxy for structural
uplift, this yields a minimum displacement of ∼2.6 km on the
Pinchal thrust only.
Section across the Quebrada Blanca
As for the Pinchal area, we interpret the Quebrada Blanca folds as related
to ramp thrusts rooting at least at the base of the folded Mesozoic units –
i.e., at the base of the Late Jurassic series – or deeper. Assuming constant
layer thicknesses, it can be deduced that the thrusts root at least 4 km
beneath the current topographic surface (i.e., at least at -2 km a.s.l.)
(Fig. 11a). In our cross-section of Fig. 11a, the interpreted thrust
needs to deepen eastward to balance the proposed section. From surface
geology alone, we cannot discard the possibility that the thrust is steeper
below the documented anticlines and roots deeper.
The secondary frontal folds with westward-decreasing wavelength (Figs. 10
and 11a) can be explained as disharmonic folds within the forelimb of the
large western anticline and/or be interpreted as reflecting the existence of
a shallow blind thrust (Fig. 11a). Such a feature is also in good
agreement with secondary (steeper) thrusts affecting these anticlines
(Fig. 9).
Line-length balancing of the cross-section of Fig. 11a results in
∼3.8 km of shortening solely related to folding (Table 1).
This value is only a minimum as it does not account for the observed
small-scale deformation or for slip on the related thrusts at depth. As for
Pinchal, the fault geometries and the structure of the footwall of the
thrusts are not constrained from surface geology to provide a complete
cross-section and the associated shortening.
Additional constraints on shortening: a possible structural
interpretation modeled by trishear
The underlying thrusts have not reached the surface and remain blind
(Figs. 5c and 9c). They are also associated with disharmonic folding at
their probable tip, with small-scale folds at the front of major anticlines.
Because of these observations, we assume fault propagation folding to be the
dominant mode of deformation in both study areas. To estimate the amount of
shortening that is taken by thrusting at depth, some assumptions for the
footwall and thrust geometries are needed.
One possible interpretation is to consider a thrust ramp parallel to the
layers of the backlimb of the anticlines, rooting at the base of the series
involved in folding, as in the sections of Fig. 11. Disharmonic folding at
the front of the anticlines is likely related to the local thickening of
layers at the front of a shallow upward-propagating ramp, as in the frontal
triangular zone of trishear folds. To further quantify the shortening
associated with this chosen structural interpretation, we conduct kinematic
trishear modeling (e.g., Allmendinger, 1998; Erslev, 1991) of
the westernmost anticlines documented at the Quebrada Tambillo (Pinchal
area) and Quebrada Blanca. This approach accounts for folding and slip on
propagating thrust faults, and it models the deformation distributed at the tip of these evolving faults. The trishear
formalism relies on a set of parameters that are adjusted here by trial and
error so as to visually fit the proposed structural geometries of the
anticlines. The values of these parameters are within the range considered
in previous studies (Tables S1–S3) (e.g., Allmendinger, 1998; Allmendinger
and Shaw, 2000; Cristallini and Allmendinger, 2002; Erslev, 1991; Hardy and
Ford, 1997; Zehnder and Allmendinger, 2000), and 65–100 combinations of
these parameters have been tested here. Further details are provided in the
Supplement (Text S1).
Shortening values on the various structures documented in
this study. Shortening associated with folding is estimated from line-length
balancing on the various folds of the two investigated sites. Additional
constraints on shortening are provided for the western anticlines from
trishear modeling; these include folding of these anticlines and thrusting
on the associated ramp. Thrusting on the Pinchal Thrust (PT) is deduced from
its subsurface geometry and the minimum thickness of the folded Mesozoic
series. See text for additional details.
Here we present our preferred models, which allow for satisfactorily reproducing
the proposed structural geometries, acknowledging that these
solutions are not unique. The structural geometries of the westernmost
anticlines of the two study sites are reproduced (Fig. 11), and the
evolution of deformation is modeled over time taking into account the
Cenozoic growth strata. The various stages of deformation are shown in
Figs. S19 and S20. We find that the
geometries of the western anticlines can be reproduced with a cumulative
shortening of 3.1 km for Quebrada Tambillo (Pinchal area) and 6.6 km for
Quebrada Blanca (Fig. 11). Based on the range of tested acceptable models,
we estimate that these shortening values are determined here with an
uncertainty of ∼0.2 km.
The above shortening values only account for the deformation (folding and
thrusting) absorbed across the modeled westernmost anticlines. Synclinal
folding accounts for an additional minimum shortening of ∼0.4 km as deduced by line-length balancing in the Pinchal area, leading to a
total of >3.5 km of shortening across the Mesozoic units along
the Quebrada Tambillo section. This includes folding, as well as slip on the
detachment and western thrust ramp (Table 1). When adding the minimum
∼2.6 km of thrusting deduced on the Pinchal Thrust, we get
>6.1 km of total shortening across the whole Pinchal area.
Similarly, in the Quebrada Blanca area, the easternmost anticline and
syncline take up ∼2 km of shortening by folding deduced by
line-length balancing, leading to a minimum amount of shortening of
∼8.6 km across the whole Quebrada Blanca section, including
folding and slip on the underlying detachment and western ramp (Table 1).
The two investigated sites take up different amounts of shortening. This may
relate to the fact that the across-strike extents of the two sections are
significantly different (∼7 km long section for the Quebrada
Tambillo vs. ∼17 km long section for the Quebrada Blanca)
(Fig. 11). The calculated shortenings similarly represent ∼47 % and ∼34 % of shortening when scaled to the extent of
the Quebradas Tambillo and Blanca sections, respectively. Differences
between these sections may also relate to the depth at which the interpreted
thrusts root (elevation of ∼0.2 km for Quebrada Tambillo vs.
depth of ∼2 km for Quebrada Blanca relative to sea level)
(Fig. 11), which is a probable result of the varying structural and
stratigraphic inheritance from the earlier Andean basins. Lateral variations
in deformation also cannot be excluded.
The shortening values estimated above depend on the proposed subsurface
structural interpretations, more specifically on the thrust geometries –
much more than on the detailed model parameters. They are to be considered
lower bounds, but only within the considered structural framework. We
favored the simplest geometry where the thrusts root at the base of the
folded series and connect at depth but cannot discard from local field
observations alone the possibility that they are steeper and root deeper. If
this were the case, shortening related to thrusting would be lower than
proposed here. Folding estimated from line-length balancing only and our
favored structural interpretation (Fig. 11) therefore provide a lower and
upper bound on shortening estimates, respectively.
Kinematics of shortening
Within the structural framework proposed in the sections of Fig. 11,
trishear modeling allows for simulating the evolution of thrust slip and
folding in the case of the westernmost anticlines of the two investigated
sites. By adding syntectonic layers while deformation proceeds, we also
reproduce the overall geometry of the base of the Cenozoic Altos de Pica
Formation deposits and of the subsequent growth strata (Figs. S19 and S20). Syntectonic surfaces and layers are prescribed an initial
3–6∘ W dipping angle, similar to the present-day regional
topographic slope (Fig. 1). From there, we find that ∼0.5
and ∼0.4 km of shortening is needed to reproduce the
first-order geometry of the base of the Altos de Pica Formation at the front
of the Quebrada Tambillo (Pinchal area) and Quebrada Blanca sections,
respectively, using the previous trishear models adjusted to our final
cross-sections. When compared to the minimum 3.1 and 6.6 km of total
shortening accumulated since ∼68 Ma across the westernmost
anticlines of these two sections, this indicates that the ∼29 Ma old basal Cenozoic layers above the Choja surface record at most only
16 % and 6 % of this total shortening, respectively. We have tested the
possibility of initial horizontal Cenozoic syntectonic layers. In this case,
a post-∼29 Ma shortening of 0.8 km at most is needed to best
adjust the observed geometry of the basal Altos de Pica Formation layers,
even though a good fit to both the geometry of the growth strata and of the
finite fold structure cannot be satisfactorily found.
These results are then used to quantitatively describe the evolution of
shortening over time across the westernmost anticlines of the two
interpreted sections, with account of the timing of deformation discussed in
Sect. 6.1 (Fig. 11d). We find that shortening rates were on average
∼0.07–0.16 km Myr-1 over the time span ∼68–29 Ma. They could have even been as high as ∼0.11–0.26 km Myr-1
if considering that the main deformation phase is confined to
∼68–44 Ma. Subsequently, deformation rates decreased to an
average value of ∼0.015 km Myr-1 after ∼29 Ma,
possibly starting earlier.
These average values are most probably minimum values within the framework
of our modeled structural interpretations. Indeed, thrusting and folding are
only modeled here for the westernmost anticlines of our study sites and do
not account for the shortening cumulated across the other structures
or on the Pinchal Thrust. Also, the main phase of deformation prior to
∼29 Ma could have lasted less than the ∼68–29 or ∼68–44 Ma considered time intervals
(Fig. 11d). In the case that the underlying faults are steeper and root
deeper, these minimum values would both be similarly lower.
Our results therefore quantitatively emphasize our former qualitative
conclusion that the major phase of deformation occurred sometime between
∼68 and ∼29 Ma, with a significant subsequent
slowing-down of deformation rates afterwards, possibly as soon as
∼44 Ma or earlier (Fig. 11d), a general conclusion that is
not dependent on the proposed subsurface thrust geometries.
DiscussionThe Andean Basement ThrustEvidencing a major basement thrust system along the Western Andean flank
(∼20–22∘ S)
Here, we have further documented the Pinchal Thrust, which brings basement
units of the Sierra de Moreno westward over folded Mesozoic units. Our study
in the Pinchal area suggests that this thrust bears local complexities with
several strands and minor splays, most probably related to the reactivation
of structures in the initial pre-Andean back-arc basins. Laterally, the
geological map of Skarmeta and Marinovic (1981) clearly documents
this structure from ∼21∘15′ S to
21∘35′ S and possibly down to ∼22∘ S with some structural complexities by ∼21∘35′ S with the junction of two possible strands of this basement thrust.
Similar basement thrusts have been described all along the Cordillera
Domeyko between ∼20 and ∼22∘ S. North of the map by Skarmeta and Marinovic (1981), the Quehuita (up to ∼21∘11′ S) and
Choja (between ∼21∘08′ S–21∘01′ S) faults are west-vergent thrusts bringing basement over folded Mesozoic
sediments (Aguilef et al., 2019). North of ∼21∘ S, intrusions, hydrothermalism, and surface volcanics
hamper clear observation of similar basement thrusts. Such basement thrusts,
if existent, would, however, provide a reasonable mechanism for the exhumation
and exposure of basement rocks east of the folded Mesozoic units and at
higher elevations, at the latitude of Quebrada Blanca (∼20∘45′ S) (Fig. 1). For these reasons, we cannot tell with any
certainty whether a thrust contact similar to that described in Pinchal
(this study) and further north (Aguilef et al., 2019) exists at this
latitude, but such structure is to be suspected.
South of the map by Skarmeta and Marinovic (1981), in the Sierra
de Moreno at ∼21∘45′ S, Haschke and Gunther (2003) report in their section a basement thrust over folded Mesozoic units, in
agreement with the style of deformation documented here, but with a
relatively minor displacement compared to our results in Pinchal. This
thrust is called the Sierra de Moreno Thrust here. Together with the
1:1000000 geological map of Chile (Sernageomin, 2003), the
Haschke and Gunther (2003) map suggests that this basement thrust is
cartographically continuous southward to the southern end of the Sierra de
Moreno, at ∼22∘05′ S. This possibly documents its
lateral termination.
As a conclusion, a series of west-vergent basement thrusts exists all
along the Western Andean flank, with various strands mapped as local
basement faults, as in our study (Fig. 4) or on other maps (Aguilef et
al., 2019; Haschke and Gunther, 2003; Sernageomin, 2003; Skarmeta and
Marinovic, 1981; Tomlinson et al., 2001). Altogether, these thrusts appear
as a major structural boundary all along the Western Andean flank, bringing
the basement of the Cordillera Domeyko westward over folded Mesozoic units
of the earlier Andean basins (Figs. 1 and 12) – and therefore contributing
to the uplift of the western margin of the Altiplano. They form a segmented
thrust system extending laterally over at least ∼120 km
north–south (Fig. 1) that we propose naming the Andean Basement
Thrust (hereafter ABT) system.
Schematic section across the Western Cordillera
(20–22∘ S), modified after Armijo et al. (2015). This section is
only meant to be illustrative of the general structural framework of the
Western Andes at the investigated latitudes. Colored rectangular frames locate
the structures equivalent to the Quebrada Blanca folds (blue) and to the
Pinchal Thrust (red) studied here. Shortenings estimated at these sites are
reported. Shortening values in gray are tentatively estimated by upscaling
our local findings to the whole West Andean Thrust System. The various
thrusts need to root at depth onto an east-dipping décollement (West Andean
Thrust), whose depth is not constrained. See text for further details.
We interpret the ABT to dip eastward beneath the Western Cordillera, at
least >2 km (Pinchal area) or >4 km (Quebrada Blanca
area) beneath the present-day topographic surface. Deeper and eastward, this
thrust probably connects to a crustal-scale ramp, as needed to sustain the
large-scale uplift and topographic rise of the Western Andes (Figs. 1 and 12), following the earlier ideas by Victor et al. (2004) and
Armijo et al. (2015). Such a crustal-scale structure has been termed
the West Andean Thrust (or WAT) by Armijo et al. (2015).
Shortening and timing of deformation of the Andean Basement Thrust
We estimated that the Pinchal Thrust (as part of the ABT system)
accommodated a minimum of ∼2.6 km of shortening over a
horizontal distance of ∼1 km in Pinchal. This estimate is
deduced from the geometry of the thrust and from the minimum ∼2.2 km of exhumation needed to erode Mesozoic series and expose basement at
the surface, considering that exhumation is a proxy for structural
uplift here. Thermochronological data are too limited to evaluate the amount of
basement exhumation more precisely, as well as its timing. These data are
presently absent locally in Pinchal but sparsely exist at the regional
scale when considering the ABT system over its whole extent. From apatite
fission-track dating in basement samples taken ∼20 km east
and southeast of our two study sites, Maksaev and Zentilli (1999)
inferred at least 4–5 km of basement exhumation occurring between
∼50 and 30 Ma. Such exhumation is consistent with our results
when considering the exhumation that may have accompanied the uplift
expected from overthrusting on the ABT and on the WAT further east. Older
thermochronological ages – (U–Th) / He zircon and apatite ages of
∼91 and ∼57 Ma, respectively – were found
by Reiners et al. (2015) from the basement of the Quebrada
Arcas, ∼30 km south of Pinchal, in a structural setting
equivalent to that documented here. These ages do not contradict the
previous estimates of total exhumation by Maksaev and Zentilli (1999),
even though modeling would be needed here to precisely test this. However,
they question the exact timing of basement exhumation and, from there, of
thrusting over the ABT. In the absence of properly analyzed and modeled
samples closer to the ABT, it is difficult to more precisely assess its
timing or amount of exhumation, uplift, and thrusting.
At a few places, the Pinchal segment of the ABT is covered by Cenozoic
deposits. Given this observation and with existing thermochronological ages,
we postulate that the ABT was most probably active sometime by the Late
Cretaceous to early Cenozoic – and that its activity had ceased by the early
Miocene. This suggests it may have been coeval with folding of the Mesozoic
units documented immediately further west – or starting slightly before.
The West Andean Thrust System at ∼20–22∘ SEvidencing a west-vergent thrust system along the Western Andean flank
(∼20–22∘ S)
The west-vergent folds described here as deforming Mesozoic units at
∼20–22∘ S are interpreted to form above
faults. A similar system of folds and faults affecting Mesozoic units is
expected to extend further north and south than just the two sites described
here, most probably over the entire zone of ∼20–22∘ S (Fig. 1), even though a large part north of
Quebrada Blanca is covered by Cenozoic strata. This is deduced from existing
maps and previous works (e.g., Aguilef et al., 2019; Haschke and Gunther,
2003; Sernageomin, 2003; Skarmeta and Marinovic, 1981). It therefore
probably spreads out over a north–south distance of at least ∼200 km – and possibly more as folded Mesozoic sediments are mapped on the
1:1000000 geological map of Chile (Sernageomin, 2003) in the
northward and southward continuation of the two zones investigated here.
Further west, structures at depth are hidden beneath Cenozoic deposits
(Fig. 1). Seismic profiles from the Chilean Empresa Nacional del Petroleo
(ENAP), as re-interpreted by various authors (Victor et al.,
2004; Fuentes et al., 2018; Labbé et al., 2019;
Martinez et al., 2021), also show a series of several blind mostly
west-vergent thrust faults. The faults and folds documented here across two
∼7–17 km wide sites therefore most probably pertain to a
thrust system that extends across-strike over a much wider region
(∼50 km, maybe locally more).
West-verging thrust faults along the Western Andes at ∼20–22∘ S most probably derive from the inversion of the normal
faults that bounded the earlier Andean basins. Fuentes et al. (2018)
and Martinez et al. (2021) interpreted these faults as single planar
deep-reaching thrusts, possibly rooting into basement. However, even though
such geometries cannot be discarded from local poorly resolved seismic data
or from scarce field observations, they cannot satisfactorily explain the
large-scale geometry of the Western Andean flank, as noted earlier by
Victor et al. (2004). Indeed, only a ramp–flat–ramp geometry of a
basal master fault deepening eastward beneath the Western Cordillera, whatever
its rooting depth, can account for the overall large-scale continuous
topographic rise of the entire western plateau margin (Figs. 1 and 12)
(Armijo et al., 2015; Victor et al., 2004). The blind
west-vergent thrust faults found all along the western flank at
∼20–22∘ S can therefore reasonably be interpreted
as connecting to such an east-dipping master fault (or detachment). By
integrating our local observations into these regional large-scale
considerations, we favor the earlier interpretations of Victor et al. (2004) and Armijo et al. (2015).
Altogether, these data suggest that all these thrust faults, either hidden
below Cenozoic deposits or deduced from outcropping folds, most possibly
pertain to a common west-vergent thrust system found all along the Western
Cordillera of northern Chile (20–22∘ S). We propose naming this
thrust system the West Andean Thrust System (or WATS). The WATS at
∼20–22∘ S therefore extends laterally over at
least ∼200 km and across-strike over a much wider region
(∼50 km, maybe locally more) than the two ∼7–17 km wide sites investigated in this study (Figs. 1 and 12).
Shortening across the West Andean Thrust System (∼20–22∘ S)
By excluding the possibility of steep deep-rooting single faults from the
above large-scale considerations, we favor our local structural
interpretations of Fig. 11 – and from there the associated shortening
estimates – where the faults possibly root at the base of the folded series,
at least locally. The WATS of northern Chile (∼20–22∘ S) therefore probably accommodates a minimum shortening of
∼3–9 km, as quantified from the ∼7–17 km
wide investigated areas (not including the contribution of the ABT in
Pinchal). At ∼21∘45′ S, ∼30 km south
of the Pinchal area, Haschke and Gunther (2003) report a minimum
shortening of >9 km from a ∼50 km wide
cross-section in the Sierra de Moreno area and further east. Within the
∼8–10 km wide area encompassing an equivalent of the WATS
and ABT, they estimate a minimum shortening of ∼4 km (i.e., a
minimum of ∼30 % of shortening), a value consistent and at
scale with our results. This study of Haschke and Gunther (2003) is to
our knowledge the only other work attempting to estimate the minimum total
shortening absorbed by the WATS. It becomes obvious that the various
structures of the WATS in northern Chile, wherever they are (Quebrada
Blanca, Pinchal, or Sierra Moreno areas), all absorb multi-kilometric
shortening at the scale of only one to three major folds and thrusts.
This conclusion further emphasizes that the ∼3–9 km of
shortening proposed here from the folds of the Quebrada Blanca and Pinchal
areas (when excluding the contribution of the ABT in Pinchal) are in any
case underestimates of the total shortening across the whole WATS. When
applying the minimum ∼34 %–47 % shortening estimated across
our two investigated sites to the ∼50 km across-strike extent
of the whole WATS, we tentatively propose a possible crustal shortening of
∼26–44 km (Fig. 12), a value consistent – even though in
the high range – with the ∼20–30 km qualitatively estimated
by Armijo et al. (2015) by scaling with structural relief and
crustal thickness. These estimates should, however, be taken with caution and
only as possible upper bounds, as deformation is localized on thrust faults
(e.g., Fuentes et al., 2018; Haschke and Gunther, 2003; Martinez et al.,
2021; Victor et al., 2004; this study) and not homogeneously distributed.
The precise geometry of faults at depth, in particular their rooting
depth, is also in no case constrained. A precise quantification of the
deformation recorded by buried folded Mesozoic units west of our study sites
is at the moment not possible from available seismic profiles.
Temporal evolution of deformation along the Western Andes
(∼20–22∘ S)
Our investigations underline the fact that the deformation of the Quebrada Blanca and
Pinchal areas is not linearly distributed over time and can be assigned to
two main periods: (1) a period of major deformation sometime between
∼68 and 29 Ma (possibly ∼68–44 Ma) and (2) a
subsequent period of moderate deformation from ∼29–0 Ma
(starting possibly earlier) (Fig. 11d). This is deduced for the
westernmost anticline of both study sites from trishear modeling, but the
reduction in deformation rates is expected at the scale of the whole
investigated sites as the difference in the deformation cumulated by
Mesozoic units and by post-∼29 Ma Cenozoic layers can be
qualitatively – but clearly – intuited from our field observations and
cross-sections (Figs. 5, 9, and 11). Westward, deformation is mostly
well imaged on seismic profiles for Cenozoic post-∼29 Ma
growth strata and remains less well resolved for underlying Mesozoic units
(Fuentes et al., 2018; Labbé et al., 2019; Martinez et al., 2021;
Victor et al., 2004), reflecting the fact that Mesozoic units could
also be much more deformed here than Cenozoic layers.
In this study, we find ∼0.4–0.5 km of post-∼29 Ma shortening on one single most frontal fault and fold in the case of
our two investigated sections (Figs. S19, S20): that is, over a distance of
∼5–8 km. Based on the ENAP seismic profiles in the westward
prolongation of our study areas, Victor et al. (2004) determined
a post-∼29 Ma shortening of ∼3 km,
accommodated by several west-vergent thrusts within the ∼40 km wide Atacama Bench. All these values are in overall good agreement when
setting them to the same spatial scale, as they consistently represent
∼6 %–8 % of shortening. Compared to the minimum
∼3–6 km of ante-∼29 Ma shortening (or
∼34 %–47 % of shortening) quantified on one single structure
in this study (Figs. S19, S20), the post-∼29 Ma shortening
is clearly of limited importance.
The deformation slowdown, starting by ∼29 Ma at the latest and
possibly earlier by ∼44 Ma, could therefore be regional
across the entire WATS. This reasoning applies to the WATS but may also hold
for the ABT. If the age of basement thrusting is not precisely known, it
most probably occurred by the early Cenozoic (Maksaev and Zentilli, 1999)
or even Late Cretaceous–early Cenozoic (deduced after
Reiners et al., 2015) and had ceased by ∼29 Ma (see discussion in Sect. 7.1).
This proposed time window for major folding and possibly for thrusting over
the ABT is generally consistent with the main Incaic phase of deformation
inferred by various authors as the main period of Andean mountain building
stricto sensu (e.g., Charrier et al., 2007; Cornejo et al., 2003; Pardo-Casas and
Molnar, 1987; Steinmann, 1930). The simplest interpretation of the post-∼29 Ma decline of the shortening rate is that it results from
the slowdown of the same protracted regional compressional event which
caused the formation of the west-vergent WATS and ABT. With the presently
available data at 20–22∘ S, we cannot exclude the possibility that this slowdown
may have started before ∼29 Ma – possibly as soon as
∼44 Ma or even before (Sect. 6.4) – but definitely not
afterwards.
Regional implications
Even though multi-kilometric, the shortening accommodated by the
west-vergent structures of the Western Andes outlined in this study
represents a modest contribution to the total crustal shortening of
>300 km across the entire Central Andes at ∼20∘ S (e.g., Anderson et al., 2017; Barnes and Ehlers,
2009; Eichelberger et al., 2013; Elger et al., 2005; Kley and Monaldi, 1998;
Mcquarrie et al., 2005; Sheffels, 1990). It should, however, be recalled that
the deformation absorbed across the Western Andes took place mostly in the
early stages of the Andean orogeny, sometime between ∼68 and 29 Ma (possibly ∼68–44 Ma) in the case of the WATS, possibly starting
earlier for the ABT – in any case during the Incaic phase. In
fact, when replaced within the temporal evolution of Andean
mountain building at these latitudes (e.g., Armijo et al., 2015; Charrier
et al., 2007; Mcquarrie et al., 2005; Oncken et al., 2006) (see sect. 2.1), the early multi-kilometric shortening evidenced here represents a
major contribution to initial Andean deformation, which has most often been
neglected in orogen-wide studies. The slowing-down of deformation across the
Western Andean flank by ∼29 Ma – and possibly starting after
∼44 Ma – may have accompanied the jumping and transfer of
deformation towards the east, i.e., towards the eastern Altiplano and further
east (e.g., Isacks, 1988; Mcquarrie et al., 2005; Oncken et al., 2006).
Conclusions
In this study, we investigate and explore from two outcropping sites two
major structural features within the western flank of the Chilean Andes at
∼20–22∘ S: (1) the Andean Basement Thrust (ABT)
system, which stands as a system of west-vergent thrusts bringing Paleozoic
basement over folded Mesozoic series, and (2) the West Andean Thrust System
(WATS), which is a west-vergent thrust system deforming Mesozoic and
Cenozoic sediments. The WATS is mostly hidden by the Cenozoic Altos de Pica
Formation, but structures crop out in few (up to ∼10–20 km
wide) places along the mountain flank. Even though our investigations only
rely on two limited outcropping sites, our deductions have regional
implications when compared and upscaled with previous results.
Using field and satellite observations, we build structural cross-sections
and quantify the recorded shortening at two key sites along the western
mountain flank. We estimated a minimum shortening of >2.6 km on
the ABT and >3–9 km on the few exposed structures of the
WATS. This shortening – derived from outcrop areas of limited extent –
corresponds only to a fraction of the entire deformation at the scale of the
whole Western Cordillera at ∼20–22∘ S. When
set on scale with the extent of the investigated structures, it implies the
possibility of multi-kilometric shortening across the western flank of the
Andes, possibly up to 26–44 km.
We further exploit the differential deformation recorded by folded Mesozoic
layers and Cenozoic growth strata of the post-∼29 Ma Altos de
Pica Formation. We show that the outcropping WATS was mainly active between
∼68 and 29 Ma (possibly ∼68–44 Ma) and that its
deformation rates significantly decreased after ∼29 Ma (a
decrease that may have started earlier, e.g., by ∼44 Ma). By
comparison to previous studies of the blind portions of the WATS west of our
study sites, we propose that such slowing-down of deformation rates was
regional rather than local. In addition, field observations and published
thermochronological results of basement exhumation suggest that this
temporal evolution of deformation rates may also hold for the ABT. We
therefore propose that the post-∼29 Ma (or post-∼44 Ma) decline in shortening rates resulted from the
regional slowing-down of the same protracted compressional event that caused
the formation of the west-vergent WATS and ABT, most probably accompanying
the transfer of Andean deformation towards the Altiplano Plateau, Eastern
Cordillera, and further eastward.
Data availability
Pléiades satellite imagery
(https://earth.esa.int/eogateway/missions/pleiades, last access: March 2018) was obtained through the
ISIS program of the CNES under an academic license and is not available for
open distribution. On request, the DEMs calculated from this imagery can be
provided to any academic researcher, but only after approval from the CNES
(contact: isis-pleiades@cnes.fr, with copy to lacassin@ipgp.fr and
simoes@ipgp.fr referring to this paper). Numerical computations
for the DEMs were performed using the free and open-source MicMac software
suite (Rosu et al., 2014; Rupnik et al., 2016) freely available at
https://micmac.ensg.eu/index.php (last access: April 2018). For cartographic mapping, we also used
Google Earth imagery (Landsat 7, DigitalGlobe) freely accessible at
https://earth.google.com (last access: July 2021). All geological maps used in this work are cited in
the main text and in the reference section. Our own maps are provided in the
main text. All field measurements and observations have been collected by us
during our field missions (March 2018, January 2019) and are provided in the
main text, in the figures, and in the Supplement. The trishear
kinematic modeling was conducted using FoldFault Forward version 6
(Allmendinger, 1998), freely available at
http://www.geo.cornell.edu/geology/faculty/RWA/programs/faultfoldforward.html (last access: October 2020).
The supplement related to this article is available online at: https://doi.org/10.5194/se-14-17-2023-supplement.
Author contributions
RL and MS designed the study, and TH carried it out. TH designed all figures.
The paper was prepared by TH, MS, and RL and revised by MS and RL –
with the contribution of all co-authors. All authors participated in fieldwork and in the various scientific discussions.
Competing interests
The contact author has declared that none of the authors has any competing interests.
Disclaimer
Publisher's note: Copernicus Publications remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
Acknowledgements
This study was supported by grants from the CNRS-INSU (program TELLUS-SYSTER)
and from the Institut de physique du globe de Paris (IPGP) (PI: Robin Lacassin). Fieldwork was also funded by the Andean Tectonics Laboratory of the Advanced
Mining Technology Center, University of Chile (PI: Daniel Carrizo). Earlier work on this
zone by Robin Lacassin and Daniel Carrizo was supported by ANR project MegaChile (grant
ANR-12-BS06-0004-02) and the LABEX UnivEarthS project. Tania Habel benefited from a PhD
grant provided by the French Ministry of Higher Education and Research.
Pléiades satellite imagery was obtained through the ISIS program of the CNES
under an academic license. The authors thank Arthur Delorme for his technical
assistance in producing the DEMs using the free and open-source MicMac
software. Numerical computations for the DEMs were performed on the S-CAPAD
platform, Institut de physique du globe de Paris (IPGP). The kinematic
modeling was made using FoldFault Forward version 6. Rolando Armijo and the late
Ricardo Thiele are warmly thanked for the fruitful discussions that led over the
years to this work and paper. We also benefited from discussions with
Christian Creixell, Nicolas Blanco, Andrew Tomlinson, and Fernando Sepulveda (SERNAGEOMIN), from
the valuable help of Magali Riesner for the 3D mapping, and that of Laurie Barrier
for facies and polarity identifications. Laurie Barrier and Nicolas Bellhasen are
also thanked for inspiring discussions. Comments by Laura Giambiagi, Constantino Mpodozis, and Rick Allmendinger on an earlier version of this paper are
acknowledged. Constructive reviews by Benjamin Gérard and Patrice Baby helped
improve this paper. This study was partly supported by IdEx Université
de Paris ANR-18-IDEX-0001.
Financial support
This research has been supported by the Institut national des sciences de l'Univers (TELLUS-SYSTER), the Institut de Physique du Globe de Paris (Projet Lacassin), the Centro Avanzado de Tecnología para la Minería (Projecto Carrizo), the Agence Nationale de la Recherche (grant no. ANR-12-BS06-0004-02), and IdEx Université de Paris (ANR-18-IDEX- 0001).
Review statement
This paper was edited by Federico Rossetti and reviewed by Patrice Baby and Benjamin Gérard.
References
Aguilef, S., Franco, C., Tomlinson, A., Blanco, N., Álvarez, J.,
Montecino, D., Gardeweg, M., Campos, V., Rodríguez, C., and Maksaev,
V.: Geología Del Área Quehuita-Chela, Regiones De Tarapacá Y
Antofagasta, 1:100.000, 2019.Allmendinger, R. W.: Inverse and forward numerical modeling of trishear
fault-propagation folds, Tectonics, 17, 640–656, 10.1029/98TC01907, 1998.Allmendinger, R. W. and Shaw, J. H.: Estimation of fault propagation
distance from fold shape: Implications for earthquake hazard assessment,
Geology, 28, 1099–1102, 10.1130/0091-7613(2000)28<1099:EOFPDF>2.0.CO;2, 2000.Anderson, R. B., Long, S. P., Horton, B. K., Calle, A. Z., and Ramirez, V.:
Shortening and structural architecture of the Andean fold-thrust belt of
southern Bolivia (21 S): Implications for kinematic development and crustal
thickening of the central Andes, Geosphere, 13, 538–558, 10.1130/GES01433.1,
2017.Armijo, R., Rauld, R., Thiele, R., Vargas, G., Campos, J., Lacassin, R., and
Kausel, E.: The West Andean thrust, the San Ramon Fault, and the seismic
hazard for Santiago, Chile, Tectonics, 29, TC2007, 10.1029/2008TC002427, 2010.
Armijo, R., Lacassin, R., Coudurier-Curveur, A., and Carrizo, D.: Coupled
tectonic evolution of Andean orogeny and global climate,
Earth-Sci. Rev., 143, 1–35, 2015.Baker, M.: Geochronology of Upper Tertiary volcanic activity in the Andes of
North Chile, Geol. Rundsch., 66, 455–465, 10.1007/BF01989588, 1977.Barnes, J. B. and Ehlers, T. A.: End member models for Andean Plateau
uplift, Earth-Sci. Rev., 97, 105–132,
10.1016/j.earscirev.2009.08.003, 2009.
Blanco, N. and Tomlinson, A.: Carta Guatacondo, Región de Tarapacá,
Servicio Nacional de Geología y Minería, Carta Geológica de
Chile, Serie Geología Básica, 2013.
Blanco, N., Tomlinson, A. J., Moreno, K., and Rubilar, D.: Importancia
estratigráfica de las icnitas de dinosaurios presentes en la
Formación Chacarilla (Jurásico-Cretácico Inferior), Región
de Tarapacá, Chile, 9e Congreso geológico chileno, Servicio Nacional de Geología y Minería, 2000.
Blanco, N., Vásquez, P., Sepúlveda, F., Tomlinson, A., Quezada, A.,
and Ladino, M.: Levantamiento geológico para el fomento de la
exploración de recursos minerales e hídricos de la Cordillera de la
Costa, Depresión Central y Precordillera de la Región de
Tarapacá (20–21 S), Servicio Nacional de Geología y Minería,
Santiago, Chile, 2012.Brooks, B. A., Bevis, M., Whipple, K., Ramon Arrowsmith, J., Foster, J.,
Zapata, T., Kendrick, E., Minaya, E., Echalar, A., and Blanco, M.:
Orogenic-wedge deformation and potential for great earthquakes in the
central Andean backarc, Nat. Geosci., 4, 380–383, 10.1038/ngeo1143,
2011.Buchelt, M. and Cancino, C. T.: The Jurassic La Negra Formation in the area
of Antofagasta, northern Chile (lithology, petrography, geochemistry), in:
The Southern Central Andes, edited by: Bahlburg, H., Breitkreuz, C., and
Giese, P., Lecture Notes in Earth Sciences, Springer, Berlin, 169–182,
10.1007/BFb0045181, 1988.Charrier, R., Pinto, L., and Rodrigues, M.P., Tectonostatigraphic evolution of the Andean Orogen in Chile, in: The Geology of Chile, edited by: Moreno, T. and
Gibbons, W., The Geological Society, London, 21–114, 10.1144/GOCH.3, 2007.Cornejo, P., Matthews, S., and Perez, C.: The “K-T” compresive
deformation event in northern Chile (24∘–27∘ S), 10th
Congreso Geologico Chileno, Concepcion, Chile, 6–10 October 2003.Cristallini, E. O. and Allmendinger, R. W.: Backlimb trishear: a kinematic
model for curved folds developed over angular fault bends,
J. Struct. Geol., 24, 289–295, 10.1016/S0191-8141(01)00063-3, 2002.
Davis, D., Suppe, J., and Dahlen, F.: Mechanics of fold-and-thrust belts and
accretionary wedges, J. Geophys. Res.-Sol. Ea., 88,
1153–1172, 1983.DeCelles, P., Zandt, G., Beck, S., Currie, C., Ducea, M., Kapp, P., Gehrels,
G., Carrapa, B., Quade, J., and Schoenbohm, L.: Cyclical orogenic processes
in the Cenozoic central Andes, Geological Society of America Memoirs, 212,
MWR212–222, 10.1130/2015.1212(22), 2014.DeMets, C., Gordon, R. G., Argus, D. F., and Stein, S.: Effect of recent
revisions to the geomagnetic reversal time scale on estimates of current
plate motions, Geophys. Res. Lett., 21, 2191–2194,
10.1029/94GL02118, 1994.Dingman, R. J. and Galli, O. C.: Geology and ground-water resources of the
Pica area, Tarapaca Province, Chile, Geological Survey Bulletin, 1188–1189,
US Department of Interior, 10.3133/b1189, 1965.
Dingman, R. J. and Galli Olivier, C.: Cuadrángulos Pica, Alca, Matilla y
Chacarilla, con un estudio sobre los recursos de agua subterránea:
Provincia de Tarapacá, Escala 1:50.000, Instituto de Investigaciones
Geologicas, Carta Geologica de Chile, 1962.Eichelberger, N., McQuarrie, N., Ehlers, T. A., Enkelmann, E., Barnes, J.
B., and Lease, R. O.: New constraints on the chronology, magnitude, and
distribution of deformation within the central Andean orocline, Tectonics,
32, 1432–1453, 10.1002/tect.20073, 2013.Elger, K., Oncken, O., and Glodny, J.: Plateau-style accumulation of
deformation: Southern Altiplano, Tectonics, 24, TC4020, 10.1029/2004TC001675, 2005.Erslev, E. A.: Trishear fault-propagation folding, Geology, 19, 617–620,
10.1130/0091-7613(1991)019<0617:TFPF>2.3.CO;2, 1991.Farías, M., Charrier, R., Comte, D., Martinod, J., and Hérail, G.:
Late Cenozoic deformation and uplift of the western flank of the Altiplano:
Evidence from the depositional, tectonic, and geomorphologic evolution and
shallow seismic activity (northern Chile at 19 30' S), Tectonics, 24,
TC4001, 10.1029/2004TC001667, 2005.Fuentes, G., Martínez, F., Bascuñan, S., Arriagada, C., and
Muñoz, R.: Tectonic architecture of the Tarapacá Basin in the
northern Central Andes: New constraints from field and 2D seismic data,
Geosphere, 14, 2430–2446, 10.1130/GES01697.1, 2018.Galli-Olivier, C.: Pediplain in northern Chile and the Andean uplift,
Science, 158, 653–655, 10.1126/science.158.3801.653, 1967.Garcia, M. and Hérail, G.: Fault-related folding, drainage network
evolution and valley incision during the Neogene in the Andean Precordillera
of Northern Chile, Geomorphology, 65, 279–300,
10.1016/j.geomorph.2004.09.007, 2005.Garzione, C. N., McQuarrie, N., Perez, N. D., Ehlers, T. A., Beck, S. L.,
Kar, N., Eichelberger, N., Chapman, A. D., Ward, K. M., and Ducea, M. N.:
Tectonic evolution of the Central Andean plateau and implications for the
growth of plateaus, Annu. Rev. Earth Pl. Sc., 45,
529–559, 10.1146/annurev-earth-063016-020612, 2017.Hardy, S. and Ford, M.: Numerical modeling of trishear fault propagation
folding, Tectonics, 16, 841–854, 10.1029/97TC01171, 1997.Haschke, M. and Gunther, A.: Balancing crustal thickening in arcs by
tectonic vs. magmatic means, Geology, 31, 933–936, 10.1130/G19945.1, 2003.Heit, B., Sodoudi, F., Yuan, X., Bianchi, M., and Kind, R.: An S receiver
function analysis of the lithospheric structure in South America,
Geophys. Res. Lett., 34, L14307, 10.1029/2007GL030317, 2007.Henriquez, S., DeCelles, P. G., and Carrapa, B.: Cretaceous to middle
Cenozoic exhumation history of the Cordillera de Domeyko and Salar de
Atacama basin, northern Chile, Tectonics, 38, 395–416, 10.1029/2018TC005203,
2019.Homewood, P. and Lateltin, O.: Classic swiss clastics (flysch and molasse)
The alpine connection, Geodin. Acta, 2, 1–11,
10.1080/09853111.1988.11105150, 1988.Horton, B. K.: Sedimentary record of Andean mountain building, Earth-Sci. Rev., 178, 279–309, 10.1016/j.earscirev.2017.11.025, 2018.Isacks, B. L.: Uplift of the central Andean plateau and bending of the
Bolivian orocline, J. Geophys. Res.-Sol. Ea., 93,
3211–3231, 10.1029/JB093iB04p03211, 1988.
Jaillard, E., Hérail, G., Monfret, T., Diaz-Martinez, E., Baby, P.,
Lavenu, A., and Dumont, J. F.: Tectonic evolution of the Andes of Ecuador,
Peru, Bolivia and northern Chile, in: Tectonic evolution of South America,
edited by: Cordani, U. G., Milani, E. J., Thomaz Filho, A., and Campos, D.
A., Rio de Janeiro, Brazil, 481–559, 2000.Kley, J. and Monaldi, C. R.: Tectonic shortening and crustal thickness in
the Central Andes: How good is the correlation?, Geology, 26, 723–726,
10.1130/0091-7613(1998)026<0723:TSACTI>2.3.CO;2, 1998.Labbé, N., García, M., Simicic, Y., Contreras-Reyes, E., Charrier,
R., De Pascale, G., and Arriagada, C.: Sediment fill geometry and structural
control of the Pampa del Tamarugal basin, northern Chile, GSA Bulletin, 131,
155–174, 10.1130/B31722.1, 2019.Lamb, S.: Did shortening in thick crust cause rapid Late Cenozoic uplift in
the northern Bolivian Andes?, J. Geol. Soc., 168,
1079–1092, 10.1144/0016-76492011-008, 2011.Lamb, S.: Cenozoic uplift of the Central Andes in northern Chile and
Bolivia–reconciling paleoaltimetry with the geological evolution,
Can. J. Earth Sci., 53, 1227–1245, 10.1139/cjes-2015-0071, 2016.Lucassen, F., Becchio, R., Wilke, H., Franz, G., Thirlwall, M., Viramonte,
J., and Wemmer, K.: Proterozoic–Paleozoic development of the basement of
the Central Andes (18–26∘ S)–a mobile belt of the South
American craton, J. S. Am. Earth Sci., 13, 697–715,
10.1016/S0895-9811(00)00057-2, 2000.
Maksaev, V. and Zentilli, M.: Fission track thermochronology of the Domeyko
Cordillera, Northern Chile: implications for Andean tectonics and porphyry
copper metallogenesis, Explor. Min. Geol., 8, 65–89, 1999.Martinez, F., Fuentes, G., Perroud, S., and Bascuñan, S.: Buried thrust
belt front of the western Central Andes of northern Chile: Style, age, and
relationship with basement heterogeneities, J. Struct. Geol.,
147, 104337, 10.1016/j.jsg.2021.104337, 2021.McQuarrie, N.: The kinematic history of the central Andean fold-thrust belt,
Bolivia: Implications for building a high plateau, Geological Society of
America Bulletin, 114, 950–963, 10.1130/0016-7606(2002)114<0950:TKHOTC>2.0.CO;2, 2002.McQuarrie, N., Horton, B. K., Zandt, G., Beck, S., and DeCelles, P. G.:
Lithospheric evolution of the Andean fold–thrust belt, Bolivia, and the
origin of the central Andean plateau, Tectonophysics, 399, 15–37,
10.1016/j.tecto.2004.12.013, 2005.Mitra, S.: Structural models of faulted detachment folds, AAPG Bull., 86,
1673–1694, 10.1306/61EEDD3C-173E-11D7-8645000102C1865D, 2002.
Mpodozis, C., Ramos, V., Ericksen, G., Cañas Pinochet, M., and
Reinemund, J.: The Andes of Chile and Argentina, in: Geology of the Andes
and its Relation to Hydrocarbon and Mineral Resources, edited by: Ericksen,
G. E., Canas Pinochet, M. T., and Reinemund, J. A., Earth Sciences Series,
Circum-Pacific Council for Energy and Mineral Resources Houston, Texas,
59–90, 1989.Muñoz, N. and Charrier, R.: Uplift of the western border of the
Altiplano on a west-vergent thrust system, northern Chile, J. S. Am. Earth Sci., 9, 171–181, 10.1016/0895-9811(96)00004-1, 1996.Norabuena, E., Leffler-Griffin, L., Mao, A., Dixon, T., Stein, S., Sacks, I.
S., Ocola, L., and Ellis, M.: Space geodetic observations of Nazca-South
America convergence across the central Andes, Science, 279, 358–362,
10.1126/science.279.5349.358, 1998.Oncken, O., Hindle, D., Kley, J., Elger, K., Victor, P., and Schemmann, K.:
Deformation of the central Andean upper plate system–Facts, fiction, and
constraints for plateau models, in: The Andes – Active subduction orogeny,
edited by: Oncken, O., Chong, G., Franz, G., Giese, P., Götze, H.-J.,
Ramos, V. A., Strecker, M. R., and Wigger, P., Frontiers in Earth Sciences,
Springer, Berlin, Germany, 3–27, 10.1007/978-3-540-48684-8_1,
2006.Pardo-Casas, F. and Molnar, P.: Relative motion of the Nazca (Farallon) and
South American plates since Late Cretaceous time, Tectonics, 6, 233–248,
10.1029/TC006i003p00233, 1987.Puigdomenech, C., Somoza, R., Tomlinson, A., and Renda, E.: Paleomagnetic
data from the Precordillera of northern Chile: A multiphase rotation history
related to a multiphase deformational history, Tectonophysics, 791, 228569,
10.1016/j.tecto.2020.228569, 2020.
Ramos, V. A.: Late Proterozoic-early Paleozoic of South America-a
collisional history, Episodes Journal of International Geoscience, 11,
168–174, 1988.Rapela, C., Pankhurst, R. J., Casquet, C., Baldo, E., Saavedra, J., and
Galindo, C.: Early evolution of the Proto-Andean margin of South America,
Geology, 26, 707–710, 10.1130/0091-7613(1998)026<0707:EEOTPA>2.3.CO;2, 1998.Reiners, P. W., Thomson, S. N., Vernon, A., Willett, S. D., Zattin, M.,
Einhorn, J., Gehrels, G., Quade, J., Pearson, D., and Murray, K. E.:
Low-temperature thermochronologic trends across the central Andes, 21 S–28
S, in: Geodynamics of a Cordilleran Orogenic System: The Central Andes of
Argentina and Northern Chile, edited by: DeCelles, P. G., Ducea, M. N.,
Carrapa, B., and Kapp, P. A., Memoir, Geological Society of America,
212, 215–249, 10.1130/2015.1212(12), 2015.Reutter, K.-J., Scheuber, E., and Chong, G.: The Precordilleran fault system
of Chuquicamata, northern Chile: Evidence for reversals along arc-parallel
strike-slip faults, Tectonophysics, 259, 213–228, 10.1016/0040-1951(95)00109-3, 1996.Riesner, M., Lacassin, R., Simoes, M., Armijo, R., Rauld, R., and Vargas,
G.: Kinematics of the active West Andean fold-and-thrust belt (Central
Chile): structure and long-term shortening rate, Tectonics, 36,
287–303, 10.1002/2016TC004269, 2017.Riesner, M., Lacassin, R., Simoes, M., Carrizo, D., and Armijo, R.:
Revisiting the crustal structure and kinematics of the Central Andes at
33.5∘ S: implications for the mechanics of Andean
mountain-building, Tectonics, 1347–1375, 10.1002/2017TC004513, 2018.Rosu, A.-M., Pierrot-Deseilligny, M., Delorme, A., Binet, R., and Klinger,
Y.: Measurement of ground displacement from optical satellite image
correlation using the free open-source software MicMac,
ISPRS J. Photogramm., 100, 48–59,
10.1016/j.isprsjprs.2014.03.002, 2014.Rupnik, E., Pierrot-Deseilligny, M., Delorme, A., and Klinger, Y.: Refined
satellite image orientation in the free open-source photogrammetric tools
Apero/Micmac, ISPRS Annals of Photogrammetry, Remote Sensing and Spatial
Information Sciences, 3, 83–90, 10.5194/isprs-annals-III-1-83-2016, 2016.
SERNAGEOMIN: Mapa Geológico de Chile: versión digital, Servicio
Nacional de Geología y Minería, Publicación Geológica
Digital, Santiago, Chile, 2003.Sheffels, B. M.: Lower bound on the amount of crustal shortening, in the
central Bolivian Andes, Geology, 18, 812–815,
10.1130/0091-7613(1990)018<0812:LBOTAO>2.3.CO;2, 1990.
Skarmeta, M. J. and Marinovic, S. N.: Geologia de la hoja Quillagua: Region
de Antofagasta, escala 1:250.000, Instituto de Investigaciones Geologicas
(Chile), Carta Geologica de Chile, 1981.Steinmann, G.: Geologie von Peru, J. Geol., 38,
190, 10.1086/623704, 1930.
Suppe, J.: Geometry and kinematics of fault-bend folding, Am. J.
Sci., 283, 684–721, 1983.Tassara, A., Götze, H.-J., Schmidt, S., and Hackney, R.:
Three-dimensional density model of the Nazca plate and the Andean
continental margin, J. Geophys. Res.-Sol. Ea., 111,
B09404, 10.1029/2005JB003976, 2006.
Tomlinson, A. J. and Blanco, N.: Structural evolution and displacement
history of the West Fault System, Precordillera, Chile: Part 2, postmineral
history, VIII Congreso Geoloìgico Chileno, Antofagasta, Chile, 13–17 October 1997, 1878–1882,, 1997a.
Tomlinson, A. J. and Blanco, N.: Structural evolution and displacement
history of the West Fault system, Precordillera, Chile: Part 1, synmineral
history, VIII Congreso Geoloìgico Chileno, Antofagasta, Chile, 13–17 October 1997, 1873–1877,, 1997b.Tomlinson, A. J., Blanco, N., Maksaev, V., Dilles, J., Grunder, A. L., and
Ladino, M.: Geología de la Precordillera Andina de Quebrada Blanca e
Chuquicamata, Regiones I y II (20∘ 30' – 22∘ 30'S),
Servicio Nacional de Geología y Minería, Santiago, Chile, 444,
2001.
Tomlinson, A. J., Blanco, N., and Ladino, M.: Carta Mamiña, Región
de Tarapacá, Servicio Nacional de Geología y Minería, Serie
Geología Básica, 2015.
Vergara, H. and Thomas, A.: Hoja Collacagua: región de Tarapaca: carta
geológica de Chile 1: 250.000, Servicio Nacional de Geología y
Minería, 1984.Victor, P., Oncken, O., and Glodny, J.: Uplift of the western Altiplano
plateau: Evidence from the Precordillera between 20∘ and
21∘ S (northern Chile), Tectonics, 23, TC4004, 10.1029/2003TC001519, 2004.
Wölbern, I., Heit, B., Yuan, X., Asch, G., Kind, R., Viramonte, J.,
Tawackoli, S., and Wilke, H.: Receiver function images from the Moho and the
slab beneath the Altiplano and Puna plateaus in the Central Andes,
Geophys. J. Int., 177, 296–308,
10.1111/j.1365-246X.2008.04075.x, 2009.Yuan, X., Sobolev, S. V., Kind, R., Oncken, O., Bock, G., Asch, G., Schurr,
B., Graeber, F., Rudloff, A., Hanka, W., Wylegalla, K., Tibi, R., Haberland,
C., Rietbrock, A., Giese, P., Wigger, P., Röwer, P., Zandt, G., Beck,
S., Wallace, T., Pardo, M., and Comte, D.: Subduction and collision
processes in the Central Andes constrained by converted seismic phases,
Nature, 408, 958–961, 10.1038/35050073, 2000.Zandt, G., Velasco, A. A., and Beck, S. L.: Composition and thickness of the
southern Altiplano crust, Bolivia, Geology, 22, 1003–1006,
10.1130/0091-7613(1994)022<1003:Catots>2.3.Co;2, 1994.Zehnder, A. T. and Allmendinger, R. W.: Velocity field for the trishear
model, J. Struct. Geol., 22, 1009–1014,
10.1016/S0191-8141(00)00037-7, 2000.