The western Andes at ~20–22°S: A contribution to the quantification of crustal shortening and kinematics of deformation.

. The Andes are an emblematic active Cordilleran orogen. Mountain-building in the Central Andes (~20°S) started by Late Cretaceous to Early Cenozoic along the subduction margin, and propagated eastward. In general, the structures 10 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 western Andes. Our results confirm the existence of two main structures, once our field observations are combined with regional data: (1) the Andean Basement 15 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. Once restored, we estimate that both structures accommodate together 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 along the whole western Andes. The timing of the main deformation recorded in the folded Mesozoic series can be bracketed between ~68 and ~29 Ma 20 – and possibly between ~68 and ~44 Ma – from dated deformed geological layers


Introduction 25
Along the western margin of South America (Figure 1), the oceanic Nazca plate plunges beneath the South American continent, with a present-day convergence rate of ~8 cm/yr at ~20ºS, 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 at 20°S (e.g. Brooks et al., 2011;Norabuena et al., 1998) -contributes to the deformation of the Andean mountain-building initiated by Late Cretaceous-Early Cenozoic along the western Andes of the Bolivian Orocline 45 (between 16-22°S), and proceeded since then with the progressive eastward propagation of deformation onto the South American continent (e.g. Anderson et al., 2017;Armijo et al., 2015;Barnes et al., 2008;Barnes & Ehlers, 2009;Charrier et al., 2007;DeCelles et al., 2015;Eichelberger et al., 2013;Elger et al., 2005;Faccenna et al., 2017;Kley & Monaldi, 1998;McQuarrie et al., 2005;Oncken et al., 2006;Sheffels, 1990; and references therein). Most local and mountain-wide previous studies have essentially focused on the structures of the Altiplano-Puna plateau and on those of the various cordilleras to the 50 east, but the structures located along the western flank of the orogen have remained up to now relatively under-studied, and their contribution to the significant topographic relief ( Figure 1) and crustal thickness (e.g. Allmendinger et al., 1990;Introcaso et al., 1992;Isacks, 1988) poorly understood.
In most classical models of Andean mountain-building, the western flank is described as a passive monoclinal-like crustalscale flexure (e.g., Isacks, 1988;Lamb, 2011Lamb, , 2016McQuarrie, 2002). However, in the late 1980's, Mpodozis and Ramos 55 (1989) described west-vergent thrusting along the western Andean margin. Later, other authors also described various thrusts, mostly west-vergent (but not only), at several localities along the western Andean flank (e.g. Charrier et al., 2007;Farías et al., 2005;Fuentes et al., 2018;Garcia & Hérail, 2005;Martínez et al., 2021;Muñoz & 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), a clear west-vergent fold-and-thrust-belt has been documented along the western 60 Andes (Armijo et al., 2010;Riesner et al., 2017Riesner et al., , 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 significant amount of crustal shortening (Riesner et al., 2017(Riesner et al., , 2018. At 33º30'S, the orogen is relatively younger and narrower than in the Bolivian Orocline ~1300 km further north. In contrast, 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 & 65 Ehlers, 2009;Eichelberger et al., 2013;Elger et al., 2005;Faccenna et al., 2017;Kley & Monaldi, 1998;McQuarrie et al., 2005;Oncken et al., 2006;Sheffels, 1990) across the entire >650 km wide orogen, but their role at the start 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 mid-upper Cenozoic deposits and volcanics (Armijo et al., 2015;Farías et al., 2005;SERNAGEOMIN, 2003;Victor et al., 2004). A quantitative analysis of this deformation and its 70 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 North Chile.
In this study, we provide quantitative 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 west Andean flank where the underlying deformed Mesozoic layers are exposed (Figure 1). The Pinchal area, at ~21°30'S, exhibits a west-vergent thrust that brings the 75 Paleozoic basement of the Cordillera Domeyko over folded Mesozoic units. In the Quebrada Blanca zone, ~80 km further https://doi.org/10.5194/egusphere-2022-629 Preprint. Discussion started: 14 July 2022 c Author(s) 2022. CC BY 4.0 License. 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 on the deformation of the whole western Andean flank (Figure 1). Despite these limitations, we find that the shortening of these structures is multi-kilometric, revealing that the contribution of the west Andean flank to Andean mountain building is not negligible. Additionally, we show 80 that the main deformation recorded by the folded Mesozoic units occurred sometime between ~68 and ~29 Ma (and possibly between ~68 Ma and ~44 Ma), further emphasizing that these structures mostly participated to the early stages of mountainbuilding.

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.
Starting from the East (oldest and deepest units, exposed at high altitudes) to the West (youngest units, lower altitudes), these are ( Figure 1): (1) Andean basement consisting of metamorphic rocks of Precambrian and Paleozoic ages; (2) folded volcanosedimentary deposits of Mesozoic age (Triassic-Cretaceous), unconformably overlain by (3) less-deformed mid-upper 120 Cenozoic (Oligocene -Quaternary) volcanics and sedimentary cover. Magmatic intrusions locally alter these different units, and are mostly Cenozoic (SERNAGEOMIN, 2003). This only pictures the first-order structuration of the western Andean flank, as Mesozoic strata may be locally trapped in between two basement units, and Cenozoic layers may be unconformably overlying older strata even to the east ( Figure 1). Laterally, and 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, and the description proposed here 125 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, intensive magmatic activity (volcanism and major 130 granite intrusions) welded together the basement during the Late Carboniferous to Early Permian (Charrier et al., 2007;Ramos, 2008;Vergara & Thomas, 1984).
The Mesozoic deposits (Triassic to Cretaceous), found today along the west Andean flank, formed in a proto-Andean arc and backarc basin system during the early period of the Andean cycle (e.g. Charrier et al., 2007;Mpodozis & Ramos, 1989).

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 (Figure 1), such as the Sierra del Medio to the east and the Sierra de Moreno to the west at ~22°S (e.g. Haschke 150 & Günther, 2003;Henriquez et al., 2019;Puigdomenech et al., 2020;Tomlinson et al., 2001) -not to be confused with the north-south trending strike-slip Domeyko Fault System, also called West Fissure System (or Falla Oeste) (e.g. Charrier et al., 2007;Reutter et al., 1996;Tomlinson & Blanco, 1997a, 1997b, 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 & Günther, 2003;SERNAGEOMIN, 2003;Skarmenta & Marinovic, 1981). 155 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. 160 Further west, a series of mostly west-vergent thrusts affecting the Mesozoic to Cenozoic series have been inferred from seismic profiles (Victor et al., 2004;Fuentes et al., 2018;Martinez et al., 2021). Victor et al. (2004) determined ~3 km of shortening of the syn-tectonic 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 165 disputable. Haschke and Günther (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 onto an eastward dipping master fault (Victor et al., 2004;Haschke and Gunther, 2003;Armijo et al, 2015) or whether their are steeply dipping single planar faults reactivated from the earlier Andean basins (Fuentes et al., 2018;Martinez et al., 2021) remains debated. It follows that even if published data document the existence of 170 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 mid-upper Cenozoic clastic sediments and ignimbrites commonly hide the folded Mesozoic layers and their contact with the basement (Figure 1). Field investigations are limited to the few sparse areas where the erosion 175 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 ( Figure 1): (1) at ~21°30'S, where the Paleozoic basement thrusts over the Mesozoic (Skarmenta and Marinovic, 1981). This zone will be referred to as the Pinchal area (next to Cerro Pinchal, 4193 m a.s.l.). (2) At ~20°45'S, where folded Mesozoic units can be observed. This zone is hereafter named Quebrada Blanca area, after its largest canyon. 180

Available Data
The most detailed -even though large-scale -existing geological map for the Pinchal area is the Quillagua map (1:250,000 scale, Skarmenta and Marinovic, 1981). For the Quebrada Blanca area, the recent Guatacondo map (1:100,000 scale, Blanco & 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 Blanco and Tomlinson (2013), Armijo 185 et al. (2015) and 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 Rupnik et al., 2016), we calculate high-resolution DEMs from tri-stereo Pléiades imagery, with a 0.5 m 190 resolution. These DEMs are down-sampled to a resolution of 2m to enhance data treatment and calculations (e.g. stratigraphic projection and image processing). Relative vertical accuracy may reach ~1m, depending on local slope.
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 200 allow ground verifications and provide supplementary details, such as minor thrusts and folds, 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 205 landslides or recent sediment deposits hide the underlying deformation pattern. Therefore, geometrical observations and detailed mapping of the structures may be locally difficult, 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 210
We use structural measurements, field observations and the obtained structural map to build cross-sections of 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 cross-sections mainly from field observations (strike and dip angles, polarity criteria, first-order stratigraphic column), with additional information taken from satellite imagery. 215 In contrast, in the Quebrada Blanca area, we build our 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 getting more precisely the large-scale structural geometries by averaging the usual local minor 220 variations in strikes and dips that derive from direct 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.
The limits of our interpretations mostly relate to the difficulty of unambiguously correlating stratigraphic layers, and to the 225 fact that layers may not keep constant thicknesses. As local topographic relief is reduced to a few hundred meters at most, the construction of cross-sections is mostly restricted to extrapolating derived average surface dip angles at depth.

Crustal shortening and kinematic modeling
We apply a line-length-balancing approach to the obtained 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 neither for penetrative deformation 230 nor for slip on underlying faults, and stands here as a conservative minimum.
From surface geology, we have no precise indication on the structure and geometry of faults and layers at depth. To make a step forward in our estimates of crustal shortening, we consider the simplest structural geometries where faults 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 235 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 a set of parameters that are here adjusted by trial and error to fit observed structural geometries. By adding sedimentary layers at various steps during ongoing deformation, we model syntectonic deposition and subsequent deformation, in order to reproduce deformation of Cenozoic layers.
Additional information on trishear modeling, together with the range of tested parameters, are provided later and in 240 supplementary material. We recognize that our best-fit model parameters may not be unique. This is not expected to impact much estimated total shortening for a specific structural interpretation as this result depends mostly on the modeled crosssections. We cannot discard the possibility that faults are steeper and root deeper. If this were the case, crustal shortening would be lower. These various points are further discussed in section 7.2 and in the supporting information.
Deformation is expressed in terms of shortening (in km) and of relative shortening at the scale of the investigated sites (in %). 245 Relative shortening is hereafter defined as the ratio of the estimated shortening by the initial length of the undeformed section.

Field observations
Because our observations are in contradiction with previous stratigraphic and structural interpretations of the folded Mesozoic series, we hereafter describe in detail our field observations. We subsequently discuss and compare them to previous 250 interpretations, and propose a solution reconciling these observations with regional stratigraphic knowledge.

Stratigraphic observations
In the landscape (Figure 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 Paleozoic basement ( Figure S1) dominates the eastern part of the Pinchal area, and is composed of mainly coarse-grain 260 granodiorites and diorites, as well as metamorphic rocks comprising gneisses, migmatites and mica-schist, consistent with previous descriptions in the area (Skarmenta & Marinovic, 1981).
The older part of the outcropping Mesozoic series consists of continental deposits ( Figure S2), with a high content of Paleozoic lithics and volcano-clastic and tuffitic low-rounded conglomerates, of greenish, beige and brownish colors. Clast sizes vary from a few millimeters to a few decimeters. At places, these rocks bear sedimentary polarity criteria such as grain-grading, 265 grain-sorting, cross-bedding and tangential beds ( Figure S3). In the eastern part of the Pinchal area, we locally observe below this series dark green detrital pelites (lutites) ( Figure S4). On the basis of petrographic and sedimentological correlations, these 275 detrital Mesozoic sediments recall units mapped as Triassic north of the Pinchal zone (between 21º-21º30'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 ( Figure 2) and can be easily  (Figures 4 and 5a) where the Mesozoic series seems to be most complete. Thicknesses of the stratigraphic units are not at scale on the figure, but are given in the main text (section 4.1.1). By analogy to regional descriptions, these layers are suspected to be Triassic at the base, and Jurassic in the case of the marine fossiliferous levels (see section 4.1.3 for additional details). The description of Cenozoic units is here completed based on the work of Victor et al. (2004). Color-code in line with maps ( Figures 1, 4 and S14) and cross-sections ( Figure 5). In 285 the Pinchal area, 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 (as depicted by "?"). See Figures S1-S12 and corresponding captions (in supplementary material) for detailed sedimentologic descriptions. 290 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 ( Figure S5). Upsection, numerous stromatolites ( Figure S6) and bivalves ( Figure S7) are found. Its thickness varies between 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 (cm-dm) limestone layers of rose-beige color ( Figure S8), over a thickness of ~50-100 m. Going up-section, the marine series becomes more marly, more beige, and with less limestone layers, evidencing 295 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 (cm to m) ( Figure 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 (Skarmenta & Marinovic, 1981) if applicable here. In this case they would be of Middle Jurassic age (Bajocian to Callovian). 300 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).
Upsection, the beige marls become more calcareous again, with thin limestone layers ( Figure S10). Finally, this marine sequence ends with black marls containing layers of beige sandstones (mm to few cm -rarely dm -thick) ( Figure S11), indicative of a detrital component in a probable deep-seated basin, comparable to the "flysch" series of the Alpine basins 305 (Homewood & 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 & Dingman, 1962;Galli-Olivier, 1967;Victor et al., 2004) (Figures 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 (Figure 310 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 section 2.2).     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.

Comparison to previous stratigraphic and structural interpretations 395
In the Pinchal area, a basement thrust was reported in the Quillagua 1:250,000 geological map (Skarmenta & Marinovic, 1981).
In this map, the Mesozoic units are interpreted as pertaining to the Jurassic Quinchamale formation, deposited in a backarc 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, Skarmenta & Marinovic (1981) interpreted the main structure of the Pinchal zone as an 400 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 405 detrital continental units are here stratigraphically below a marine sequence (Figure 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 410 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 supplementary material 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. 415

Structural Interpretation
The cross-section of Quebrada Tambillo (Figure 5c) summarizes our interpretation of the sub-surface structural geometry of the Pinchal area. Shortening estimates are summarized in Table 1.
Based on the dip angle of the C/S-fabric of the shear-zone ( Figure 6b) and on the mapping of the thrust on satellite imagery, we estimate that the Pinchal Thrust dips ~40ºE in the near-surface, locally less such as along Quebradas Tania and Martine 420 (Figures 5a-b). All secondary strands of the Pinchal Thrust are expected to root at depth onto the main shear-zone. The secondary thrust breaking the core of the syncline, roughly parallel to the Pinchal Thrust (Figure 4), may be described as an out-of-the-syncline thrust (Mitra, 2002) and probably also connects onto it at depth. A similar reasoning is proposed to all small-scale thrusts and décollements observed within the inverted synclinal limb.
Considering that the folds west of the Pinchal Thrust develop above underlying thrusts is a reasonable and classical assumption. 425 These thrusts are expected to root at least at the base of the outcropping folded Mesozoic series (Figure 5c), 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 (Figure 7b)   Shortening associated to folding is estimated from linelength 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. Trusting on the Pinchal Thrust (PT) is deduced from its sub-surface geometry and the minimum thickness of the folded Mesozoic series. See text for additional details.

440
Using the simplified geometry of layers along Quebrada Tambillo (Figure 5c), 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. Because of the pervasive presence of small-scale folding and thrusting (Figures 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 is surely to be 445 added, as well as the offsets on the interpreted underlying thrusts and on the Pinchal Thrust. The thrust offsets largely depend on the thrust geometry and therefore on the structural interpretation. An estimate of the contribution of thrusting will be provided below (section 6.2).
The minimum thickness of the Mesozoic series is estimated to ~2.2 km from the normal limb of the syncline along Quebrada Tambillo. Thus, it can be considered that the strict minimum exhumation of the basement is equally of ~2.2 km. Assuming a 450 constant 40°E dip angle, this yields a minimum displacement of ~2.6 km on the Pinchal thrust only.

Stratigraphy of the Quebrada Blanca area
The stratigraphy at ~20°45'S is well described in the Guatacondo geological map (Blanco & Tomlinson, 2013). Unlike in Pinchal, basement rocks do not crop out in the investigated zone (Figure 8), but larger scale maps (e.g. SERNAGEOMIN, 455 2003) show Paleozoic basement units further east and higher in the topography (Figure 1).

Structural observations
The structural map of the Quebrada Blanca area is reported on Figure 8. 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), three large-scale folds are observable: a wide syncline in the center (Higueritas syncline), bounded 475 by two anticlines to the west (Chacarilla anticline) and east (fold names from Blanco & Tomlinson, 2013;Fuentes et al. 2018).
The scale of these folds is multi-kilometric (Figure 8). The cross-section of Figure 9 illustrates the asymmetry of the folds: both anticlines have steeper western limbs (dip angles vary between ~50-80ºW), whereas their eastern limbs have more gentle dip angles (varying between ~20-50ºW) (Figure 9a). Despite the fact that the eastern flank of the eastern anticline is widely hidden by magmatic intrusions and hydrothermalism, its southern part is well observable in the field (Figure 8 Figure 1). Colored lines report mappable layers. For visibility, only major, well-correlated layer traces are represented here. Black boxes locate the swath profiles from which layers were projected for the construction of the structural east-west cross-section (Figure 9). The A-B section corresponds to the topographic profile used for this same cross-section. Strike and dip measurements are extracted from 3D-map-ping (see section 3.3) or observed in the field. Strike symbols without dip value are derived from satellite imagery. Thick black lines correspond to 490 major fold axes. 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 & Tomlinson, 2013). Letters C, D, E, F, G and H to the north-east (within the folded Chacarilla and Majala Formations) report the layers illustrated on Figure S17   (c) East-west cross-section based on (a) and (b). Interpretation at depth is indicated with transparent colors, in contrast with sub-surface 500 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 & Tomlinson, 2012). The ~27-29 Ma age of the basal deposits of the Altos de Pica formation is derived from regional considerations (Victor et al., 2004). In the field, we observe small-scale deformation within both anticlines (Figure 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) are observable on the western edge of the Chacarilla anticline (Figures 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 thrustaffected small-scale fold is observed (Figures 9c and Figure S17

Structural interpretations 525
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 -ie at the base of the Late Jurassic series - (Figure 9c), 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.) ( Figure   9c). In our cross-section of Figure 9c, 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 530 deeper.
The secondary frontal folds with westward decreasing wavelength (Figures 9c and 10) 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 thrust ( Figure   9c). Such a feature is also in good agreement with secondary (steeper) thrusts affecting these anticlines (Figure 9).
Line-length-balancing of the cross-section of Figure 9c results in ~3.8 km of shortening solely related to folding (Table 1). 535 This value is only a minimum as it does not account neither for slip on the related thrusts at depth nor for the observed smallscale deformation.

Timing of deformation
The projection of mapped strata indicates that the Mesozoic series is overall concordant in Quebrada Blanca (Figure 9). The 540 cross-section of the Guatacondo map (Blanco & Tomlinson, 2013) suggests however the presence of a minor angular unconformity (<10º ) 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 post-dates the deposition of all these series. In the Quebrada Blanca area, the youngest folded layers of the Cerro Empexa 545 Formation bear U-Pb ages of 68.9±0.6 Ma and 68±0.4 Ma (Blanco & Tomlinson, 2013) (Figures 8 and 9c). We can therefore conclude that the main deformation episode post-dates ~68 Ma, even though we cannot exclude earlier but minor deformation when compared to the observed large-scale folding (Figure 9c).
Magmatic intrusions dated at ~44 Ma intrude the folded Mesozoic units, and appear cartographically not affected by folding (Blanco & Tomlinson, 2013) (Figure 8). This suggests that the major part of the folding occurred during the ~68-44 Ma time 550 interval. However, without additional observations of the deformation -or not -of these intrusions (geometry of the contact with surrounding host units, mineral deformation…), we cannot unequivocally conclude here from this simple cartographic observation. Even though we suspect that the deformed series of the Pinchal zone are Triassic to Jurassic (section 4.1), we do not have any absolute ages of the folded units. Therefore, we postulate that the main deformation here also post-dates ~68 Ma by analogy 555 to our observations at the Quebrada Blanca.
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 few places in the Pinchal zone (Figure 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 560 recorded by folded Mesozoic units appears of greater intensity than that of the Cenozoic growth layers (Figures 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 Ma 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 Ma 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 565 thrusting took place prior to ~29 Ma.

Additional constraints on shortening
Because the underlying thrusts have not reached the surface (Figures 5c and 9c), 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 on the footwall and thrust geometries are needed. 570 The simplest 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 Figures 5c and 9c. 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 across the sections of Figures 5c and 9c, we do kinematic trishear modeling (e.g. Allmendinger, 1998;Erslev, 1991) of the westernmost anticlines documented at the Quebrada Tambillo 575 (Pinchal area) and Quebrada Blanca. This approach accounts for folding, slip on propagating thrust-faults, and 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 fit the proposed structural geometries of the anticlines. The values of these parameters are within the range considered in previous studies (e.g. Allmendinger, 1998;Allmendinger & Shaw, 2000;Cristallini & Allmendinger, 2002;Hardy & Ford, 1997;Zehnder & Allmendinger, 2000). Here we present our best-fitting model, which allows for 580 reproducing satisfactorily our structural sections, acknowledging that it is not unique. Further details are provided in supplementary material. Tables S1 to S3 provide the set of parameters used for our modeling.
The structural geometries of the westernmost anticlines of the study sites are reproduced (Figure 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 Figures S19 and S20 in the supplementary material. We find that the geometries of the western anticlines can be reproduced 585 with a cumulative shortening of 3.1 km for Quebrada Tambillo (Pinchal area), and of 6.6 km for Quebrada Blanca (Figure 11).   sites (a,b). Thicker lines outline model results, while transparent lines and colors refer to the cross-sections of Figures 9c and 5c. These lines are color-coded according to the stratigraphic level they represent, as in the original cross-sections. Black lines report the modeled thrusts and horizontal arrows report the model 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 595 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 (~44 Ma) that are cartographically discordant, both derived from the Guatacondo geological map (Blanco & Tomlinson, 2013 -see also Figures 8 and 9c), and to the ~29 Ma age of the oldest Cenozoic layer of the Altos de Pica Formation (APF) (Victor et al., 2004) above the Choja erosional surface. It is possible that most deformation occurred prior to ~44 Ma, as deduced from the age of the intrusions cartographically seemingly post-dating folding (Figure 8), even though this argument is to be taken with caution. Our 600 results underline two phases of deformation, with a slowing down of deformation since ~29 Ma at least, possibly even before. Intermediate stages of the trishear modeling are reported on Figures S19 and S20 (supplementary material) for the cross-sections of Quebrada Tambillo and Quebrada Blanca, respectively. Model parameters are indicated in Tables S1-S3 in supplementary material.
The above shortening values only account for the deformation (folding and thrusting) absorbed across the modeled 605 westernmost anticlines. Synclinal folding accounts for an additional minimum shortening of ~0.4 km as deduced by linelength-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 610 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 extent of the two sections are significantly different (~7 km long section for the Quebrada Tambillo vs. ~17 km long section for the Quebrada Blanca). The calculated shortenings similarly represent ~47% and ~34% of shortening when scaled to the extent of 615 the Quebradas Tambillo and Blanca sections, respectively. Differences between these sections may also relate to the depth at which the interpreted thrusts root (altitude of ~0.2 km for Quebrada Tambillo vs. depth of ~2 km for Quebrada Blanca, relative to sea level) (Figure 11). Lateral variations in deformation can also not be excluded.
The shortening values estimated above depend on the proposed sub-surface structural interpretations, and more specifically on the thrust geometries. They are to be considered lower bounds only within the considered structural framework. We favored 620 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 single planar faults that root deeper. Folding estimated from linelength balancing only and our favored structural interpretation therefore provide a lower and upper bounds on shortening estimates, respectively. However, as further discussed in section 7.2, large-scale considerations on the overall topography and geology at the scale of the whole Western Andean flank tend to favor our local structural interpretations of Figures 5c and 9c,  625 and from there the shortening estimates from trishear kinematic models.

Kinematics of shortening
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 (Figures S19 and S20). 630 Syntectonic surfaces and layers are prescribed an initial 3-6° W dipping angle, similar to the present-day regional topographic slope ( Figure 1). From there, we find that ~0.5 km and ~0.4 km of shortening are 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 km and 6.6 km of total shortening accumulated since ~68 Ma across the westernmost anticlines of these two sections, this indicates 635 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 640 of the two investigated sections, with account on the timing of deformation discussed in section 6.1 (Figure 11d). We find that shortening rates were on average of ~0.07-0.16 km/Myr over the time span ~68-29 Ma. They could have been even as high as ~0.11-0.26 km/Myr 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 after ~29 Ma, starting possibly earlier.
These average values are most probably minimum values, within the framework of our modeled structural interpretations. 645 Indeed, thrusting and folding are here only modeled for the westernmost anticlines of our study sites, and do not account for the shortening cumulated neither across the other structures nor on the Pinchal Thrust. Also, the main phase of deformation prior to ~29 Ma could have lasted less than the ~68-29 Ma or ~68-44 Ma considered time intervals, respectively ( Figure 11d).
In the case that the underlying faults are steeper and root deeper, these minimum values would be lower, but this interpretation is not favored here (section 7.2). 650 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 (Figure 11d), a general conclusion that is not dependent on the proposed sub-surface thrust geometries. 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 660 geological map of Skarmenta and Marinovic (1981) clearly documents this structure from ~21º15'S to 21º35'S, and possibly down to ~22°S with some structural complexities by ~21º35'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°S and ~22°S. North of the map by Skarmenta and Marinovic (1981), the Quehuita (up to ~21º11'S) and Choja (between ~21º08'S-21º01'S) Faults are westvergent thrusts bringing basement over folded Mesozoic sediments (Aguilef et al., 2019). North of ~21ºS, intrusions, 665 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) (Figure 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. 670 South of the map by Skarmenta and Marinovic (1981), in the Sierra de Moreno at ~21°45'S, Haschke and Günther (2003)'s section report 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 here called the Sierra de Moreno Thrust.
Together with the (1:1,000,000) Geological map of Chile (SERNAGEOMIN, 2003), Haschke and Günther (2003)'s map suggests that this basement thrust is cartographically continuous southward to the southern end of the Sierra de Moreno, at 675 ~22°05'S. This possibly documents its lateral termination.
As a conclusion, there exists a series of west-vergent basement thrusts all along the Western Andean flank, with various strands mapped as local basement faults, as in our study (Figure 4) or in other maps (Aguilef et al., 2019;Haschke & Günther, 2003;SERNAGEOMIN, 2003;Skarmenta & 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 680 folded Mesozoic units of the earlier Andean basins (Figure 1)-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 (Figure 1), that we propose to name here the Andean Basement Thrust (hereafter ABT) system.
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 685 ramp, as needed to sustain the large-scale uplift and topographic rise of the Western Andes (Figure 1), following the earlier ideas by Victor et al. (2004) and Armijo et al. (2015). Such 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 strict minimum of ~2.6 km of shortening 690 over a horizontal distance of ~1 km in Pinchal. This shortening would be associated with basement exhumation, but limited thermochronological data do not permit to evaluate this precisely. These data are presently absent locally in Pinchal, but sparsely exist at a regional scale when considering the ABT system over its whole extent. From apatite fission track dating in basement samples taken ~20 km east and south-east of our two study sites, Maksaev and Zentilli (1999) inferred at least 4-5 km of basement exhumation occurring between ~50-30 Ma. This is in good agreement with our results in terms of amount of 695 uplift that would result from basement overthrusting on the ABT, and above the WAT further east. Older thermochronological ages -(U-Th)/He zircon and apatite ages of ~91 Ma 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 on 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, 700 of thrusting over the ABT. In the absence of properly analyzed and modeled samples closer to the ABT, it is difficult to assess more precisely its timing or amount of exhumation, uplift and shortening.
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 Late Cretaceous to Early Cenozoic -and that its activity had ceased by Early Miocene. This suggests it may have been coeval with folding of the 705 Mesozoic units documented immediately further west -or starting slightly before.

Evidencing a west-vergent thrust system along the West Andean flank (~20-22°S)
The west-vergent folds described here as deforming Mesozoic units at ~20-22ºS are interpreted to form above thrusts that root at least at the base of the folded series. A similar system of folds and faults affecting Mesozoic units is expected to extend 710 further north and south than just the two sites described here, most probably over the entire zone of ~20-22ºS (Figure 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 & Günther, 2003;SERNAGEOMIN, 2003;Skarmenta & 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 (1:1,000,000) Geological map of Chile (SERNAGEOMIN, 2003) in the north-and south-ward 715 continuation of the two zones investigated here. or Martínez et al. (2021), show also a series of several blind mostly west-vergent thrust-faults. These west-verging thrust faults are interpreted by Fuentes et al. (2018) and Martínez et al. (2021) as single planar deep-reaching thrusts reactivating the faults 720 bounding the earlier Andean basins. However, even though such geometries cannot be discarded from local poorly resolved seismic data, 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 deeping eastward beneath the Western Cordillera can account for the overall large-scale topographic rise of the entire western plateau margin (Figure 1) (Victor et al., 2004, Armijo et al., 2015. The blind west-vergent thrust-faults found all along the western flank at ~20-22°S can therefore 725 reasonably be interpreted as connecting onto 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) andArmijo 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 North . We propose to name 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 (Figure 1).

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 735 structural interpretations of Figures 5c and 9c -and from there the associated shortening estimates. The WATS of northern Chile (~20-22°S) therefore 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 Günther (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 740 minimum shortening of ~4 km (i.e. a minimum of ~30% of shortening), a value consistent with our results. This study of Haschke and Günther (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. 745 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 under-estimates 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 find a possible crustal shortening of ~26-44 km, 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 750 crustal thickness. These estimates should however be taken with caution and as possible upper bounds, as deformation is localized on thrust faults (Haschke and Gunther, 2003;Victor et al., 2004;Fuentes et al., 2018;Martinez et al. 2021;this study) and not homogeneously distributed. 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) 755
Our investigations underline 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 ~68-29 Ma (possibly ~68-44 Ma); and (2) a subsequent period of moderate deformation from ~29-0 Ma (starting possibly earlier) (Figure 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 760 by post ~29 Ma Cenozoic layers can be qualitatively -but clearly -intuited from our field observations and cross-sections (Figures 5,9 and 11). Westward, deformation is mostly well-imaged on seismic profiles for Cenozoic post-~29 Ma growth https://doi.org/10.5194/egusphere-2022-629 Preprint. Discussion started: 14 July 2022 c Author(s) 2022. CC BY 4.0 License. strata and remains less well-resolved for underlying Mesozoic units (Victor et al, 2004;Labbé et al, 2016;Fuentes et al, 2018;Martinez et al, 2021), reflecting the fact that Mesozoic units could here also be much more deformed 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 765 investigated sections ( Figures 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 ( Figures S19, S20), 770 the post ~29 Ma shortening is clearly of limited importance.
The deformation slow-down, starting by ~29 Ma at 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 Early Cenozoic (Maksaev & Zentilli, 1999) or even Late Cretaceous -Early Cenozoic (deduced after Reiners et al., 2015), and had ceased by ~29 Ma (see discussion in section 7.1). 775 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 & Molnar, 1987;Steinmann, 1929). The simplest interpretation on the post ~29 Ma decline of the shortening rate is that it results from the slow-down 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 780 cannot exclude that this slow-down may have started before ~29 Ma -possibly as soon as ~44 Ma, or even before (section 6.3) -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 ~20ºS 785 (e.g. Anderson et al., 2017;Barnes & Ehlers, 2009;Eichelberger et al., 2013;Elger et al., 2005;Faccenna et al., 2017;Kley & Monaldi, 1998;McQuarrie et al., 2005;Oncken et al., 2012;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 ~68-29 Ma (possibly ~68-44 Ma) in the case of the WATS, starting possibly 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;790 Charrier et al., 2007;McQuarrie et al., 2005;Oncken et al., 2006) (see section 2.1), the early multi-kilometric shortening evidenced here represents a major contribution to initial Andean deformation, which has been most often neglected in orogenwide 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 et al., 1988;McQuarrie et al., 2005;Oncken et al., 2006). 795

Conclusion
In this study, we investigate and explore from two outcropping sites two major structural features within the western flank of the Andes at ~20-22°S: (1) the Andean Basement Thrust (ABT) system, which stands as a west-vergent, >120 km long system of ~north-south trending thrusts bringing Paleozoic basement over folded Mesozoic series; (2) the West Andean Thrust system (WATS), which is a west-vergent thrust system deforming Mesozoic and Cenozoic sediments, mostly covered by the Cenozoic 800 Altos de Pica Formation, but cropping out in few (up to ~10-20 km wide) places along the mountain flank. The WATS extends over at least ~200 km north-south and ~50 km across-strike. Even though our investigations only rely on two limited outcropping sites, our deductions have regional implications when compared and up-scaled 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 find a minimum shortening of >2.6 km on the ABT and of >3-9 km on the few exposed 805 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 ~20-22ºS. 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 ~29 810 Ma Altos de Pica Formation. We show that the outcropping WATS was mainly active between ~68-29 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 815 for the ABT. We therefore propose that the post ~29 Ma (or post ~44Ma) 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 820
Pleiades satellite imagery 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, however after approval from the CNES (contact: isis-pleiades@cnes.fr, with copy to lacassin@ipgp.fr and simoes@ipgp.fr, and referring to this manuscript). Numerical computations for the DEMs were performed using the free and open source MicMac software suite Rupnik et al., 2016) freely available at https://micmac.ensg.eu/index.php. For cartographic 825 mapping, we also used Google Earth imagery (Landsat 7, DigitalGlobe) freely accessible at https://earth.google.com. 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 supporting information. The trishear kinematic modeling was conducted using FoldFault Forward version 6 (Allmendinger, 1998), freely available at 830 http://www.geo.cornell.edu/geology/faculty/RWA/programs/faultfoldforward.html