Tectonophysics 495 (2010) 78–92

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Tectonophysics

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Links between topography, erosion, rheological heterogeneity, and deformation in contractional settings: Insights from the central Andes

G.E. Hilley a,⁎, I. Coutand b,c a Department of Geological and Environmental Sciences, Stanford University, Stanford, CA 94305-2115, United States b Laboratoire Géosystèmes (UMR 8157 CNRS), Université de Lille 1, 59655 Villeneuve d'Ascq, France c Department of Earth Sciences, Dalhousie University, Halifax, NS, Canada B3H 4J1 article info abstract

Article history: Orogenic structure in the central Andes (15°S–34°S) systematically varies with mean annual precipitation, Received 30 May 2008 suggesting that erosional processes may be coupled to tectonic processes. We explore the range of possible Received in revised form 7 April 2009 interactions between deformation, erosional processes, changing geodynamic conditions, and preexisting Accepted 15 June 2009 geologic structures to assess the relative importance of each. Our review and synthesis indicates that the pre- Available online 23 June 2009 orogenic geologic history and changing geodynamic conditions leave a lasting imprint on the modern structure of the mountain belt. In contrast, changes in precipitation that result from regional-scale atmo- Keywords: spheric and oceanic circulation interact with local topography to control the climatic zoning across the Contractional orogens Erosion/tectonic coupling central Andes. Changes in erosion, which likely correspond to changes in climate, rock uplift rate, and the Rheological heterogeneity exposure of different lithologies, may in turn affect the tectonic development of the region at the scale of Climate each morphotectonic province. In this view, geologic and geodynamic factors set the stage for the nature and strength of coupling between erosional processes, tectonic deformation, and mountain belt structure in the different parts of the central Andes. © 2009 Elsevier B.V. All rights reserved.

1. Introduction The South American convergent plate margin mostly consists of the subduction of the Nazca Plate beneath South America along Over the last three decades, earth scientists have realized that an eastward-dipping subduction zone (Stauder, 1975; Barazangi processes that redistribute mass across Earth's surface may and Isacks, 1976, 1979; Hasegawa and Sacks, 1981; Bevis and Isacks, influence deep earth processes (e.g., Davis et al., 1983; Beaumont 1984; Cahill and Isacks, 1992; Allmendinger et al., 1997). Volcanic et al., 1992; Willett et al., 1993; Willett, 1999). For example, global and tectonic processes have deformed and uplifted this margin, and regional climate changes may enhance or subdue erosion which hosts a variety of landscapes, including volcanoes in the within mountain belts, and these changes result in mass redis- Western Cordillera, internally drained basins in the high-elevation tribution patterns that may affect the near-surface thermal state of Altiplano–Puna Plateau, thin- and thick-skinned fold-belts within the crust (e.g., Ehlers, 2005), the body forces within the crust (e.g., the Eastern Cordillera and Subandean ranges, and basin-and- Dahlen, 1984), and cause changes in mean elevations that result in range-like topography created by basement-cored uplifts of the isostatic adjustments (e.g., Molnar and England, 1990a). Many of Sierras Pampeanas (Fig. 1). In the central Andes, changes in theconceptslinkingsurfaceprocesses to those deep within the topography and deformation style have been attributed to a variety earth arose from observations from Taiwan (e.g., Davis et al., 1983; of tectonic factors, including the geometry of the down-going Dahlen and Suppe, 1988), the Himalayas (e.g., Hodges, 2000; subducting plate (Isacks, 1988; Gephart, 1994), the thickness of Thiede et al., 2005; Grujic et al., 2006), and geodynamic models the mantle lithosphere in the central segment of the mountain belt (e.g., Beaumont et al., 2004; Stolar et al., 2007). In the central (e.g., Kay et al., 1994; Whitman et al., 1996; Garzione et al., 2006), Andes, tectonic deformation and erosional patterns also correlate interplate coupling (e.g., Gephart, 1994; Lamb and Davis, 2003), with one another, suggesting that here too they may be linked (e.g., and pre-existing heterogeneities in the crust and mantle that Horton, 1999; Montgomery et al., 2001). reflect the pre-Cenozoic geologic history of the area (e.g., Allmendinger et al., 1983; Coutand et al., 2001; Ramos et al., 2002). However, changes in erosion related to climate, and the exposure and erosion of different rock types have been inferred to play a role in determining the patterns ⁎ Corresponding author. of deformation and topography observed within this orogen (e.g., Masek E-mail addresses: [email protected] (G.E. Hilley), [email protected] (I. Coutand). et al., 1994). For example, the intensity of erosion that likely varies

0040-1951/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2009.06.017 G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 79

Fig. 1. (A) Shaded relief map highlighting topography and morphotectonic provinces of the central Andes. White polygon shows areas that are internally drained. (B) Shaded relief map color-coded for mean annual precipitation; Mean annual precipitation taken from WMO (1975). Precipitation rates in excess of 2.6 m/yr are colored purple. Contours show depth to the downgoing Nazca plate, taken from Cahill and Isacks (1992), Bevis and Isacks (1984), and Barazangi and Isacks (1976).

because of climate patterns related to long-term global-scale rheology of the crust and its heterogeneities, the lithologies exposed at atmospheric circulation (e.g., Haselton et al., 2002; Parrish et al., the surface, and regional to local climatic conditions. In this contribution, 1982, Hartley et al., 1992; Masek et al., 1994; Strecker et al., 2007), we review geological and geophysical observations, as well as modeling may affect mountain-belt width (e.g., Willett, 1999; Montgomery studies to examine the strength and nature of coupling between tectonic et al., 2001) or control the interplate frictional coupling by and erosional processes in the central Andes. The spirit of this work modulating the sediment supply to the subduction trench (e.g., follows that of previous geological studies, in which correlations between Lamb and Davis, 2003). At a regional scale, studies have argued natural phenomena are documented and interpreted in the context of that erosion plays a key role in controlling the kinematics, experimental (e.g., Persson et al., 2004; Hoth et al., 2006)andtheoretical geometry and pattern of deformation within the subandean fold- (e.g., Dahlen, 1984) models that rationalize linkages between the and-thrust belt (e.g., Masek et al., 1994; Horton, 1999), the observed phenomena. Unfortunately, one cannot manipulate the under- Argentine Precordillera (Hilley et al., 2004), and the Sierras lying tectonic processes or the geological history of natural orogens as is Pampeanas (Sobel and Strecker, 2003; Hilley et al., 2005). the case in a controlled laboratory environment (e.g., Persson et al., 2004; While erosion may play a role in shaping the central Andes, the first- Hoth et al., 2006), and so while not wholly satisfying, we follow the path order characteristics of the topography and deformation of this orogen of others in which we seek correlations between various parameters that correlate strongly with changes in geodynamic features such as down- can be understood in terms of these simpler systems (e.g., Hoffmann- going slab geometry (e.g., Isacks, 1988; Cahill and Isacks, 1992; Gephart, Rothe et al., 2006). 1994; Allmendinger et al.,1997), and geologic heterogeneities in the crust that include inherited structures and paleogeographic features (e.g., 2. Conceptual framework linking tectonic and erosional processes Schmidt et al., 1995; Allmendinger et al., 1983, 1997; Ramos et al., 2002; Carlier et al., 2005). Based on expectations derived from the theoretical In the simplest conception, the state of stress within tectonic plates framework that links erosion to tectonic processes, the strength of the results from the sum of tractions at the plate boundaries due to plate– linkages between these factors should vary spatially depending on the tectonic interactions, the buoyancy stresses applied to the base of 80 G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 the crust that arise from density differences between the crust and strength of these erosional processes is likely moderated by factors mantle, and stresses produced by the overburden of crustal material. including climate (e.g., Knox, 1983; Bookhagen et al., 2005a,b; The rheology of the crust relates deformation and/or deformation rate Strecker et al., 2007) and the lithology of the rocks exposed at the to these loading conditions (Fig. 2). The rheology of rocks depends on surface (e.g., Gilbert, 1877; Hack, 1960; Stock and Montgomery, 1999; many factors including composition, temperature, loading stresses, Sklar and Dietrich, 2004). In turn, local climate is strongly influenced and grain size (e.g, Bürgmann and Dresen, 2008, and references by topographic relief that creates orographic gradients in precipitation therein), and so deformation may directly modify crustal rheology (e.g., Barry, 1981; Bookhagen and Burbank, 2006), and global climate through at least three mechanisms: 1) heat redistribution in the crust is perturbed by the uplift of large portions of Earth's surface such as due to magmatism, exhumation, and frictional failure in the upper the Tibetan Plateau (e.g., Raymo and Ruddiman, 1992; Ruddiman, crust, (e.g., Byerlee, 1968; Byerlee, 1978; Brace and Kohlstedt, 1980; 1997). Thus, surface uplift, climate, erosion, and deformation are Graham and England, 1976; Lachenbruch and Sass, 1980; Ryback and inextricably linked in actively deforming orogens. Cermak, 1982; Edman and Surdam, 1984; Molnar and England, 1990b; The lithology of rocks exposed at the surface partially controls Furlong et al., 1991; Royden et al., 1997; Beaumont et al., 2004; erosional efficiency and topographic slopes. Consequently, as orogens Byerlee, 1978) Lachenbruch and Sass, 1980; Ryback and Cermak, 1982; evolve, changing distribution of lithologies likely changes the way in 2) grain size reduction by dynamic recrystallization (e.g., Rybacki and which deformation is accommodated. For example, as contractional Dresen, 2000; Bystricky and Mackwell, 2001; Rutter and Brodie, 2004; deformation propagates toward the foreland, the exposure of poorly Dimanov and Dresen, 2005); and 3) spatial redistribution of consolidated basin sediments may enhance erosion, which allows rheologically distinct crustal units during deformation (e.g., Allmen- large quantities of material to be eroded evenwhere relief is subtle, and dinger et al., 1983). delays the accumulation of lithostatic stresses and advancement of Changes within the deep earth also influence the rheology of rock deformation to the orogen's external portions (e.g., Hilley et al., 2004). at shallower levels; for example, changes in mantle temperatures may Similarly, orogens in which surface processes are veracious due to eventually be manifest by temperature changes within the crust (e.g., abundant precipitation may exhume deep crystalline rocks whose Sleep, 2005, and references therein; Fig. 2). Convective removal of resistance to erosion is high. The surface exposure of such competent mantle lithosphere or magmatic underplating may alter the crust and rocks promotes the rapid steepening of orogenic slopes, increases mantle lithosphere's thicknesses and consequently, the thermal state lithostatic loads within the orogen's interior, and forces deformation to of the crust (e.g., Kay et al., 1994; Conrad and Molnar, 1997). Finally, migrate towards the external portions of the mountain belt (e.g., Hilley the rheological configuration of the crust prior to orogenisis et al., 2004; Hilley and Strecker, 2004). In addition, where erosion is potentially leaves a lasting imprint on the subsequent evolution of inefficient during the waxing stages of orogenesis, the balance the deforming crust. Thus, in active convergent orogens, the between lithostatic stresses at the base of the crust and the restoring distribution of deformation can be spatially non-uniform and evolve stresses that buoyancy provides may be altered, causing the thickened over time as the state of stress and rheology of the crust change in crust to partially sink into the mantle (e.g., Stüwe, 2002). As orogenesis space and time. wanes, lowering of the surface reduces the lithostatic stresses that Finally, erosion of Earth's surface can play a fundamental role in oppose buoyancy stresses, causing isostatic uplift of rocks (Molnar and deformation by modifying the distribution of lithostatic stresses in England, 1990a). Thus, erosional removal of mass from Earth's surface the crust (Fig. 2). For example, as surface elevations increase, so do the can modify the local lithostatic stresses acting along individual lithostatic stresses within the crust, and this may favor failure along structures and within the frictional medium of the crust, as well as the extremities of an orogen at the expense of internal shortening the balance between lithostatic loading and the restoring stresses (e.g., Dahlen, 1984; Hilley et al., 2005) and may elicit an isostatic provided by the relative buoyancy of the crust in mantle (Fig. 2). response where mean elevations change (Molnar and England, In the simple scenarios outlined above, deep earth processes 1990a). By implication, the state of stress and deformation within interact with surface processes through the alteration of crustal Earth's crust is, at least in part, controlled by erosional processes. The lithostatic loading and isostatic responses in complex ways that

Fig. 2. Schematic representation of potential feedbacks between crustal rheology, plate boundary forces, mantle processes, and erosion. G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 81 depend on the history of deformation, erosion, and heat transfer On average, elevations taper gently toward the foreland from within the crust. In addition, the pre-existing rheological state of 6000–5000 to 200–100m,althoughhighlocalslopesresultfrom the crust likely influences the deformation history during orogen- the partitioning of the foreland into a series of meridianally esis. Consequently, the initial distribution of physical properties of extensive, rhythmic and coherent ridges. the crust when combined with global and local climatic effects, establishes the strength of coupling between erosion and deforma- 3.2. Relationship of morphology to current plate geometry tion throughout the history of orogenesis. The width and height of the orogen mimic changes in the geometry 3. Physiographic, geologic, and climatic framework of the of the subducting Nazca plate. In the portion of the central Andes central Andes between 15° and 27°S, where the orogen reaches a maximum width of 650–700 km, the downgoing Nazca plate dips steeply (~30°) down 3.1. Physiography to 600 km beneath South America (Fig. 1B). By the time the Altiplano– Puna Plateau tapers both to the south and north, coinciding with a The central Andes are located between 70°W and 60°W, and significant narrowing of the orogenic system, the slab partially shoals 15°S and 35°S (Fig. 1) along an active subduction margin in which and dips shallowly (~5°) between depths of 100 and 125 km under- the oceanic Nazca Plate is translated ENE relative to stable South neath South America (e.g., Bevis and Isacks, 1984, Barazangi and America and is underthrust beneath the South American Plate Isacks, 1976; Cahill and Isacks, 1992). This phenomenon is reflected in along an eastward dipping subduction interface (Stauder, 1975; the symmetry of the Andean topography—the pole to the great circle Barazangi and Isacks, 1976; Hasegawa and Sacks, 1981; Bevis and about which topography is most symmetric matches that about which Isacks, 1984; Cahill and Isacks, 1992; Allmendinger et al., 1997). slab dip direction is also most symmetric (Gephart, 1994). In fact, this This geodynamic configuration has led to the formation of a broad pole also corresponds to the modern rotation pole of the Nazca plate swath of high topography that roughly parallels the subduction relative to a fixed South American plate, suggesting that the symmetry zone (Fig. 1). This topography reaches peak elevations of 6962 m of the orogen is tightly linked to the subduction of the Nazca plate at Mt. Aconcagua (Argentina), the highest point in the western beneath South America (Gephart, 1994). hemisphere. The elevation and extent of the high topography We show the topography of the central Andes projected about changes with latitude along the subduction zone: at ~18°S, the the great circle defining its symmetry axis in Fig. 3A. When looking meridianal extent of the Andes reaches widths of ~750 km, at the average difference between elevations located north and whereas it becomes narrower both north and south of this location south of this symmetry axis, the orogen is remarkably symmetric (Fig. 1). (Fig. 3B), although important excursions from this symmetric view A number of morphotectonic segments can be defined in the of the Andes exist. When considering the average difference central Andes based on topographic attributes. The forearc domain between elevations (e.g., Gephart, 1994), differences within the comprises, from west to east, the continental margin, the Coastal external portions of the margin are naturally de-emphasized, since Cordillera (up to 2000 m in elevation), the Central Depression, and they lie at low elevations and are unlikely to have the correspond- the North Chilean Precordillera reaching an average elevation of ing differences in elevation expected in steep terrains. To adjust for 3000 m (e.g., Reutter et al., 2006). To the east, the Western this, we normalized the difference between points north and south Cordillera consists of the Miocene to Quaternary volcanic arc atop of this symmetry axis to their mean elevation. This allows small the subduction zone (e.g., Isacks, 1988; Fig. 1). Between the differences in low relief areas to have similar weight as large latitudes of 15°S and 26°S, the Western Cordillera is flanked on differences at high elevation (Fig. 3C). Using this approach, we see its eastern side by the high-altitude internally drained Altiplano– that the center portion of the mountain belt constituted by the Puna Plateau (Fig. 1). Elevations within the Altiplano average Altiplano–Puna Plateau is highly symmetric; however, topography ~3700 m (Isacks, 1988), whereas the Puna is generally higher and along the margins and within the foreland displays marked has average elevations of ~4400 m (Whitman et al., 1996). Across asymmetries (Fig. 3C). In particular, the enclave into the eastern- the Altiplano, the Plateau surface is mostly flat; however, the Puna most portion of the Andes north of the symmetry axis that is is partitioned into a series of smaller basins by generally N–S- absent to the south constitutes a highly asymmetric feature that trending basement ranges that reach elevations of about 5000 m was obscured using the dimensional difference (compare Fig. 3B (Coutand et al., 2001; Fig. 1). South of the Puna (N26°S), the and C). In addition, the margins of the Eastern Cordillera, as well as Principal, Frontal, and Pre-Cordilleras, which consist of a series of the whole of the Pampean and Santa Barbara ranges are highly relatively narrow, N–S trending ranges that are externally drained asymmetric with respect to their counterparts to the north. Thus, (e.g., Jordan et al., 1993; Fig. 1). To the east at these latitudes, the while the extent and location of the Plateau appears highly central Andes consist of the Sierras Pampeanas, which is a zone symmetric, the external portions of the mountain belt, including of discontinuous and widely spaced mountain ranges that are the Eastern Cordillera, Principal, Frontal, and Pre-Cordilleras, the broadly distributed across the foreland (Fig. 1; Stelzner, 1923; Sierras Pampeanas and Santa Barbara system, and the Interandean González-Bonorino, 1950; Jordan and Allmendinger, 1986). Bound- and Subandean ranges are far less so. ing the eastern edge of the internally drained plateau north of ~26°S is the Eastern Cordillera, which consists of a series of high 3.3. Pre-Cenozoic geologic history ranges that generally serve as the drainage divide between the internally drained plateau and the externally drained foreland The oldest rocks exposed within the back-arc area, south of 15°S (Kennan et al., 1995; Kley, 1996). In a transitional area between are located within the Sierras Pampeanas (Fig. 1) and include ~24°S and ~26°S, the Eastern Cordillera and the Sierras Pampeanas Mesoproterozoic high-grade metamorphic basement consisting of intermingle, with some Pampean-style ranges located to the east of gneisses and migmatites (Mon, 1979; Ramos et al., 1986)thatmaybe the Eastern Cordillera. However, north of ~24°S, these vanish and locally intruded by several generations of late Neoproterozoic to early give way to a series of ranges that closely flank the Eastern Paleozoic granitoids (Halpern and Latorre, 1973; Lork and Bahlburg, Cordillera called the Santa Barbara system (Rolleri, 1976; Baldis 1993). Late Neoproterozoic to Early Cambrian greenschist-grade et al., 1976). Finally, north of the Santa Barbara system, a series of turbidites and pelagic clays of the Puncoviscana Formation appear ranges lying east of the Eastern Cordillera comprise the Inter- within the Puna and adjacent Eastern Cordillera (NW Argentina and andean and Subandean ranges (Allmendinger et al., 1983; Fig. 1). southernmost Bolivia) and were deposited in a deep marine basin 82 G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92

Fig. 3. Topography of the central Andes, highlighting the symmetry of the orogen. (A) GTOPO30 30 arc-second resolution shaded relief topography projected about a great circle whose pole coordinates are 63°N, 113°W (Gephart, 1994). (B) Difference between elevations in equi-angular locations from the great circle, as computed by Gephart (1994). (C) Fractional difference between points in equi-angular positions from the symmetry axis, which highlights the asymmetry of the external portions of the orogen. This asymmetry was somewhat masked by the analysis of Gephart (1994) due to the fact that large differences in absolute elevation occur mostly along the margins of the high relief plateau because of its high topography. In contrast, the fractional difference highlights large changes in elevation relative to each pair's average value. In the figure, z is the elevation value at each point north of the axis of symmetry, while z′ is the elevation at the corresponding symmetric location south of the axis of symmetry. Gray areas denote zones in which comparisons of the topographic symmetry were not possible due to the presence of the Pacific Ocean.

(thickness is ~2000 m; Ruiz Huidobro, 1960; Turner, 1960a,b; Cretaceous time, and perhaps during the Proterozoic to early Aceñolaza,1973; Omarini, 1983). Importantly, the transition between Paleozoic (Allmendinger et al., 1983). These units are locally the crystalline basement and the Puncoviscana Formation is abrupt, unconformably covered by Cambrian quartzites (Hausen, 1925) and corresponds to the boundary between the Sierras Pampeanas deposited in an extensional back-arc basin (Salfity et al., 1976) and Eastern Cordillera at ~26°S (Fig. 1). Whether this sharp transition associated with an extensional episode affecting the southwestern represents a buttress unconformity between a southerly crystalline margin of Gondwana (Ramos et al., 1986; de Wit and Ransome, 1992; basement high (potentially the Pampia terranes, e.g., Ramos, 2008) Bahlburg and Furlong, 1996). During the Ordovician, an up to 4000- within the Early Cambrian, or motion along a fault is unclear; how- m-thick turbiditic section is deposited across the Puna and Eastern ever, it certainly marks a persistent and profound change in the Cordillera in Argentina (Bahlburg and Breitkreuz, 1991; Bahlburg and paleogeography of the central Andes that occurred at least as early as Furlong, 1996). In the late Ordovician, these units were deformed and

Fig. 4. Correlations between Central Andean (A) physiography, (B) mean annual precipitation, (C) total horizontal displacements of the back arc region relative to fixed

SouthAmerica,(D)horizontaldisplacementswithineachmorphotectonicprovincerelativetofixed South America, (E) total shortening: (L/Lo)×100withLo, initial length and L, total displacement (see Table A1), (F) pre-Cenozoic foreland deposit thickness, (G) slab angle between 100 and 150 km, (H) mountain belt width, (I) structural style, and (J) precipitation distribution. Physiography modified after Isacks (1988) and Coutand (1999); numbered morphotectonic regions (1–3) are referenced to the text. Mean annual precipitation taken from WMO (1975); thick black lines show location of sections used to calculate displacements and shortening (Table A1; numbers on profiles reference specific sources given in Table A1). Calculations for total displacement, displacement by location, and total shortening are given in Table A1. Pre-Cenozoic foreland deposit thicknesses include the sediments that are undisrupted by Cretaceous rifting and are compiled in Table A2. Subduction angle calculated by measuring distance between slab depths of 100 and 150 km presented in Cahill and Isacks (1992). Mountain-belt width measured perpendicular to structures shown on Fig. 1. G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 83 84 G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 metamorphosed to greenschist grade (Turner and Méndez, 1979; 3.4. Cenozoic geologic history and deformation Coira et al., 1982; Mon and Hongn, 1987) and subsequently intruded by a granitoid and rhyodacitic belt known as the Faja Eruptiva de la The style, location, and amount of Cenozoic deformation correlates Puna Oriental (FEPO; Mendez et al., 1972). The FEPO is interpreted as with zones defined by the current plate-boundary geometry, as well as representing a late Ordovician to Silurian magmatic arc that resulted by pre-Cenozoic paleogeography of the central Andes. While the from a subduction zone located to the west of the Puna, which likely timing of Andean deformation within the Puna–Altiplano Plateau and dipped to the east (Coira et al., 1982). In northwestern Argentina, bounding Eastern Cordillera are somewhat disputed and certainly Silurian–Devonian sedimentation is mainly localized in the present- highly diachronous, the commencement of modern deformation in day Subandean ranges and eastern part of the Eastern Cordillera, and this area is broadly constrained to be between late-Eocene and mid- comprises about 2000 m of marine sandstones and shales (Turner, Miocene time (e.g., Allmendinger, 1986; Horton, 1997, 1998; Lamb and 1960a,b; Padula et al., 1967). To the west of the FEPO, Carboniferous Hoke, 1997, Marrett and Strecker, 2000; Coutand et al., 2001, 2006; and Permian plutons indicate that this margin was active during the Hongn et al., 2007). Late Paleozoic. The activity of the convergent margin during the Paleozoic is 3.4.1. Altiplano–Puna Plateau partially represented within the sedimentary sequence lying east of Apatite fission-track thermochronology data from the Alti- the current Altiplano–Puna Plateau margin. These Paleozoic strata plano–Puna Plateau indicate that exhumation of this part of the are present north of the crystalline basement/Puncoviscana transi- orogen has been limited (Carrapa et al., 2005; Deeken et al., 2006; tion around ~26°S (Allmendinger and Gubbels, 1996, their Fig. 6). In Ege et al., 2007; Letcher, 2007). The Bolivian Altiplano records parts of the Sierras Pampeanas to the south, lower Paleozoic strata exhumation in the Early Oligocene (33–27 Ma) at slow and consist of both marine and continental strata deposited in an constant rate of 0.2 mm/yr until the late Miocene (11–7Ma)(Ege extensional to trans-tensional tectonic environment, while Permian et al., 2007). In the Puna Plateau, the onset of exhumation is strata record filling and amalgamation of these sub-basins (Fernan- recorded by apparent AFT ages between 24 and 29 Ma (Deeken dez-Seveso and Tankard, 1995; Fisher et al., 2002). In total, these et al., 2006; Carrapa et al., 2005; Letcher et al., unpublished data). If Paleozoic strata generally reach thicknesses of b2.5 km. As an exhumation serves as a crude proxy for total tectonic shortening exception, Paleozoic sediments currently exposed within the and rock uplift, these data suggest that shortening within the Precordillera of Argentina (Fig. 1) reach thicknesses of up to 7 km interior of the Altiplano–Puna Plateau has been limited during the (e.g., Zapata and Allmendinger, 1996). Between ~26°S and ~24°S, last 20–25 Ma. The deformation that did occur within the Puna is Paleozoic sedimentary rocks are often present, but generally of more distributed than the Altiplano Plateau to the north and has limited thickness. However, north of ~24°S, a thick, eastwardly uplifted N–S-trending bedrock blocks that partition the plateau tapering section of Paleozoic clastic rocks reaches a thickness of into a series of ranges and adjacent basins. These basins contain the ~12 km in the Eastern Cordillera of Bolivia and thins to 4–9kmat detritus that presumably resulted from uplift and exhumation of the edge of deformation in the Chaco Plain (Allmendinger et al., the ranges in the Plateau interior (Alonso et al., 1991; Vandervoort 1983; Allmendinger and Gubbels, 1996). These rocks were appar- et al., 1995; Coutand et al., 2001). Basins that record the ently deposited in the retro-arc basin of the Paleozoic volcanic arc partitioning of the Puna into disconnected depocenters may have (e.g., Ramos, 2008). Thus, as with the Precambrian and earliest formed as early as the Oligocene, and certainly by the Miocene Cambrian rocks in the central Andes, the Paleozoic section starts to (Vandervoort et al., 1995). In fact, recent sedimentologic and thicken around ~26°S, reaches maximum thicknesses in the current structural data from the Puna have been interpreted as evidence for Bolivian foreland (e.g., Allmendinger and Gubbels, 1996; McQuarrie Eocene exhumation and depocenter development (Carrapa and and DeCelles, 2001, and references therein) and finally tapers again DeCelles, 2008). Some of these basins are up to 5-km thick northward to disappear around 17.5°S (Baby et al., 1994). (Coutand et al., 2001), and those basins whose tops define the low- Mesozoic strata exposed within the central Andes generally relief high elevation Plateau surface are internally drained (Isacks, record an episode of intracontinental rifting starting in the 1988; Vandervoort et al., 1995; Sobel et al., 2003). Shortening Neocomian (Salfity, 1982). This rift basin system has a complex estimates within the Puna are limited; however, those palinspatic geometry with N–StoE–W elongated depocenters (see Fig. 1 in reconstructions that exist suggest that less than 10–15% shortening Marquillas et al., 2005) filled with strata of the Salta Group (Turner, (~37 km) has occurred within the Puna during the Cenozoic 1959), which is Early Cretaceous to Late Eocene in age (e.g., (Coutand et al., 2001). Marquillas et al., 2005). Deformed remnants of these basins are The location and style of shortening within the interior of the preserved in the Altiplano–Puna Plateau, the Eastern Cordillera and Altiplano Plateau and along its margins differs from that of the Puna the foreland thrust belts. In the latter, Mesozoic strata are of limited (e.g., Allmendinger, 1986; Lamb and Hoke, 1997; Lamb et al., 1997; thickness or altogether absent south of ~27.5°S, reach a maximum Coutand et al., 2001; McQuarrie and DeCelles, 2001; Elger et al., 2005). thickeness of ~4–5 km between 25° and 26°S in the Santa Barbara The eastern margin of the Puna–Altiplano Plateau undergoes a system and disappear north of 23°S when the Lomas de Olmedo transition around ~23°S in which thick-skinned, high angle reverse branch diverges eastward (see Fig. 1 in Marquillas et al., 2005; Salfity faults accommodate deformation to the south, while a west-verging and Marquillas, 1994). back-thrust belt accommodates deformation to the north (Sempere A plot of pre-Cenozoic stratal thickness in the modern foreland et al., 1990; McQuarrie and DeCelles, 2001; Elger et al., 2005). with latitude shows the abrupt changes in thickness that accompany Within the Altiplano itself, ~65 km of shortening appears to be the widespread exposure of Precambrian basement south of ~26°S accommodated between the Oligocene and mid-Miocene, with (Fig. 4F). Importantly, changes in basin thickness correlate strongly deformation migrating to the orogen's margins after that time with the different morpho-tectonic provinces in the central Andes: (Lamb et al., 1997; Lamb and Hoke, 1997; Elger et al., 2005). Defor- the thick (up to 17 km) pre-Cenozoic foreland basin units are mation within the Plateau and along its margins has created up to a associated with the fold-and-thrust belts in the Subandean and the 12-km-thick depocenter (Sempere et al., 1990; Kennan et al., 1995; Precordillera fold-and-thrust belts, the irregular Cretaceous rift basins Lamb and Hoke, 1997; Horton et al., 2002)thatiscurrently are associated with thick-skinned deformation in the Santa Barbara internally drained, and may have been so since Oligocene time system, and relatively thin (or absent) pre-Cenozoic basin strata are (Horton et al., 2002; Sobel et al., 2003). Deformation within the often associated with the Sierras Pampeanas basement-cored uplift Altiplano did not partition the depocenter into discrete, isolated province (Fig. 4F). units to the same degree as in the Puna (Allmendinger et al., 1997, G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 85 and references therein), creating the broad, largely uninterrupted, the deep influence of pre-existing paleogeographic features on the low relief Plateau surface (Fig. 4A). modern structure of the range. Paleobotanical and stable-isotope data suggest that the broad, high-elevation portion of the central Andes, comprised mainly of 3.4.3. Southernmost central Andes the Puna–Altiplano Plateau abruptly rose sometimes after 10 Ma. In South of the termination of the Puna Plateau, the Aconcagua particular, paleobotanical evidence indicates that the Plateau had fold-and-thrust belt, the Frontal Cordillera, and Precordillera obtained no more than a third of its modern elevation at 20 Ma, and definetheeasternslopeoftheAndes(Fig. 4A). The emergence of had risen to no more than half of its modern elevation by ~11 Ma the thin- and thick-skinned fold-and-thrust belts at this latitude (Gregory-Wodzicki, 2000). Likewise, fractionation of oxygen iso- corresponds to the re-emergence of a thick pre-Tertiary sedimen- topes recorded by paleosol carbonates, lacustrine carbonates, tary section in the modern foreland (Allmendinger et al., 1983; sandstone cements, and carbonates formed in paleo-marshes, Ramos et al., 2002)thatappearstohavebeendeformed,andin indicate that the surface of the Bolivian Altiplano was uplifted some places continues to deform, from the mid-Miocene to present ~2.5–3.5 km between 10.3 and 6.8 Ma (Garzione et al., 2006). These (e.g., Allmendinger et al., 1983; Ramos et al., 2002). At ~33°S, results are consistent with paleothermometric data from carbo- geologic relationships indicate that deformation propagated east- nates of the area, which have been interpreted to reflect ~3700± ward with time, and is broadly constrained to have occurred 400 km of surface uplift between 10.3 and 6.7 Ma (Ghosh et al., between 15 and 9 Ma in the Aconcagua fold-and-thrust belt, 9– 2006). Thus, at least within the Plateau, elevations appear to have 6 Ma in the Frontal Cordillera, and 5–2 Ma in the Precordillera increased during the late Cenozoic in what appears to be a rela- (Ramos et al., 2002). Further north at ~30.5°S, deformation within tively short-lived phase of rapid surface uplift (Garzione et al., the Precordillera seems also to have migrated eastward with time, 2006). This rapid uplift has been attributed to the convective although the timing of deformation is slightly different than to the removal of eclogitized lower crust and lithosphere mantle beneath south (Allmendinger et al., 1990; Jordan et al., 1993). In particular, the Altiplano–Puna Plateau (Molnar and Garzione, 2007; Hoke and deformation within only the Precordillera appears to have begun Garzione, 2008). around ~20 Ma, and propagated eastward in a somewhat dis- continuous manner until the present (Jordan et al., 1993). In either 3.4.2. Eastern Cordillera case, deformation is broadly bracketed to have commenced in the Fission track data from the Eastern Cordillera record the earliest early- to mid-Miocene in this area, and currently continues to onset of Cenozoic exhumation in the orogen during Late Eocene–Early deform rocks at the boundary between the Precordillera and Oligocene in Bolivia (40 Ma) (Ege et al., 2007) and further south in the Sierras Pampeanas provinces (e.g., Zapata and Allmendinger, Argentina (40–38 Ma) (Coutand et al., 2001, 2006). Interestingly, the 1996). Despite its relatively narrow width, total Cenozoic displace- end of exhumation in the Eastern Cordillera is diachronous in the ment accommodated within the Precordillera ranges from ~108 North and the South. In Bolivia, it ceases in the Late Oligocene–Early to 135 km (relative to fixed South America) (Fig. 4C, D and E) Miocene, after which time exhumation resulting from tectonic rock (Allmendinger et al., 1990; Cristallini, 1996; Ramos et al., 2004, uplift is transferred to the foreland area (Ege et al., 2007), while in 1996a,b). Argentina it persists during the Late Miocene in the Angastaco area (Deeken et al., 2006; Coutand et al., 2006) until contractional 3.4.4. Sierras Pampeanas deformation is transferred eastward to the Santa Barbara System in The easternmost portions of the central Andes comprise, from the Pliocene (Kley and Monaldi, 2002). south to north, the Sierras Pampeanas, Santa Barbara system, and The Eastern Cordillera is a basement-involved thrust system Subandean fold-and-thrust belt, respectively (Figs. 1 and 4A). (Kley et al., 1996; Kley, 1996). In Argentina, Cenozoic deformation Within the Sierras Pampeanas, deformation is accommodated by is taken up along moderate- to steeply-dipping reverse faults that high-angle reverse faults whose locations often spatially corre- accommodateasmuchas55–85 km of horizontal displacement spond to terrain boundaries from the Paleozoic, or inverted normal relative to stable South America (Grier et al., 1991; Baby et al., 1997; faults formed during the Cretaceous extension (Allmendinger et Coutand et al., 2001). Nested in this range, intramontane basins al., 1983; Schmidt et al., 1995; Ramos et al., 2002). Movement preserve up to 6-km-thick Cenozoic sediments (Grier et al., 1991; along these thrust faults uplifts ranges of Late Proterozoic to Early Marrett and Strecker, 2000; Coutand et al., 2006) that have been Paleozoic metamorphic bedrock (Allmendinger et al., 1983; Jordan variously cannibalized, reworked and deformed. In the southern and Allmendinger, 1986). The utilization of pre-existing crustal Eastern Cordillera, the current margins of the Cenozoic basins weaknesses, whose along-strike extent is often limited, causes closely correspond with the paleo-borders of Cretaceous rift basins, these ranges to be discontinuous (e.g., Schmidt et al., 1995; emphasizing the impact of pre-existing structure in the modern Allmendinger and Gubbels, 1996). Intramontane depocenters asso- architecture of the range (Grier et al., 1991; Deeken et al., 2006; ciated with these ranges have been enduringly disconnected from Carrera et al., 2006; Carrera and Muñoz, 2008). Further north at each other and can preserve sediments up to 4000m in thickness ~22°S, the thickening of the Paleozoic section triggers a change in (Sosic, 1972; de Urreiztieta et al., 1996; Fisher et al., 2002). AFT the style and amount of deformation accommodated within this data indicate that the frontal part of the Sierras Pampeanas (Sierra unit with the development of a back-thrust belt along the western de Aconquija) was rapidly exhumed about 5 Ma ago (Sobel and flank of the Eastern Cordillera and a deformation largely accom- Strecker, 2003) and warping of erosional surfaces in the Northern modated along detachment horizons interlayered in the Paleozoic Sierras Pampeanas indicates that deformation remains active in section (e.g., Sempere et al., 1990; McQuarrie, 2002; Elger et al., this area (e.g., Strecker et al., 1989). The Pampean Ranges 2005). Correspondingly, the Eastern Cordillera absorbs a larger accommodate only about 10–20 km of horizontal displacement amount of deformation, with total displacement accommodated relative to fixed South America (Fig. 4A, C and D) (Costa, 1992; increasing to ~100–140 km (Baby et al., 1997; McQuarrie and Ramos et al., 2004). DeCelles, 2001; McQuarrie, 2002; Elger et al., 2005; McQuarrie et al., 2008a). Thus, while the Eastern Cordillera varies little in its 3.4.5. Santa Barbara system meridianal location, the style and amount of deformation accommo- At about 26°S, the northern boundary of the Sierras Pampeanas is dated by these ranges changes abruptly along the strike. These changes marked by an abrupt decrease in the wavelength of structures. There, spatially coincide with the distribution of pre-existing Late Paleozoic the Santa Barbara system (Figs. 1 and 4A) accommodates deformation detachment horizons (e.g., Allmendinger et al.,1983), again highlighting in the foreland largely by inversion of rift-related normal faults (Kley 86 G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 and Monaldi, 2002). The thickness of Cretaceous strata varies sig- due to the tapering of the Silurian and Devonian decollements in the nificantly according to the spatial distribution of the rift depocenters Paleozoic basin (Belotti et al., 1995). This leaves a gap of ~60 km bet- (see Fig. 1 in Monaldi et al., 2008; Fig. 4F). The thickness of Cenozoic ween the southern termination of the Subandean belt and the strata exposed within the Santa Barbara ranges varies between 2.5 and emergence of the Santa Barbara system to the south (Kley and Monaldi, 5.5 km, the lower part being comprised of sediments reflecting the 2002). Thus, the Subandean belt is closely associated with the presence post-rifting thermal subsidence stage and the upper part being foreland of mechanically weak detachment horizons in pre-existing basin basin strata related to the inversion of the rift (Kley and Monaldi, deposits (e.g., Allmendinger et al., 1983). 2002). While the utilization of high angle structures has created significant topography, it has also allowed only modest amounts of 3.5. Current and Cenozoic climate shortening to be accommodated: total displacement relative to stable South America is estimated to be between 25 and 30 km (Fig. 4B, C, and Today, the Andes constitute a major topographic barrier to D) (Kley and Monaldi, 2002). The timing of contractional deformation atmospheric circulation in South America and separate the humid is still controversial, but appears to have mostly taken place during the lowlands in the east from the arid Pacificmargininthewest Pliocene (Kley and Monaldi, 2002). The physiographic boundaries of (see Fig. 1b in Bookhagen and Strecker, 2008). The semiarid to the Santa Barbara system appear to be strongly controlled by the pre- hyperarid climates observed on the western slope of the orogen Cenozoic paleogeography of the Andes. The southern extent of the and across the high Plateau (Fig. 1B; Strecker et al., 2007,and Santa Barbara system lies at ~26°S, where the bedrock abruptly references therein), result from their proximity to upwelling of transitions from the greenschist-grade Neoproterozoic to early Cam- cold water along the Pacific coast due to the cold Humboldt brian Puncoviscana Formation to Meso-Neoproterozoic migmatites, current, which may have existed in some form since the early gneisses, and granitoids (Allmendinger et al., 1983) and both the Cenozoic (Zachos et al., 2001), and the fact that these regions southern and northern boundaries of these ranges coincide with the are located within a high-pressure sector with atmospheric sub- termination/bifurcation of the rift system (Kley and Monaldi, 2002). sidence (Houston and Hartley, 2003). Precipitation in the Central Specifically, the bifurcation of the Lomas de Olmedo depocenter to the Andes is largely derived from the westward penetration of Atlantic north, causes the strike of old Cretaceous normal faults to become moisture from the Chaco Plain and Argentine foreland (Rohmeder, roughly orthogonal with respect to the Cenozoic shortening direction 1943; Halloy, 1982; Bianchi and Yañez, 1992; Garreaud et al., 2003) (Kley et al., 1999), rendering their reactivation mechanically difficult, via the Andean low-level jet (Nogués-Paegele and Mo, 1997). As and together with the tapering of Mesozoic deposits, apparently leads moisture encounters the topography of the eastern Andes, to the termination of the Santa Barbara system. precipitation is orographically enhanced (e.g., Bianchi and Yañez, 1992; Sobel et al., 2003; Strecker et al., 2007; Bookhagen and 3.4.6. Subandean ranges Strecker, 2008). Both the relief structure of the eastern flank In Bolivia, the eastern margin of the central Andes consists of of the central Andes and the source of incoming precipitation cause the Interandean Zone and the Subandean fold-and-thrust belt, large latitudinal changes in the amount of rainfall reaching the which we consider together in this discussion for simplicity, eastern slope and interior of the orogen (Fig. 1; Strecker et al., although we acknowledge that important differences between 2007; Bookhagen and Strecker, 2008). In northern Bolivia at ~15– these two zones exist (e.g. Kley and Reinhardt, 1994). East of the 17°S, orographic rainfall peaks above 3.5 m/yr at a mean elevation Bolivian Eastern Cordillera, a thin-skinned fold-and-thrust belt has of 1.3 km once topographic relief of about 1 km is achieved formed within the thick Paleozoic sedimentary section (Allmen- (Bookhagen and Strecker, 2008). South of 17°S, orographic dinger et al., 1983). Deformation in this zone likely commenced precipitation decreases as the mountain belt widens (Montgomery during the early Miocene and has propagated eastward with time et al., 2001) but is still on the order of 1.5 m/yr at 26°S (Bookhagen (Kley, 1996). AFT data also document the migration of exhuma- and Strecker, 2008). As the long, linear ranges of the Subandean tionfromtheECaround30–20 Ma to the frontal part of the fold-and-thrust belt give way to the largely discontinuous, steep- subandesinthelateMiocene–Pliocene (8–2Ma) (Ege et al., flanked and higher elevated ranges of the Santa Barbara system and 2007). This eastward propagation has contributed to the deposi- Sierras Pampeanas, precipitation is strongly focused on the wind- tion of late Cenozoic foredeep sediments of up to ~5 km thick in ward sides of these ranges (Strecker et al., 2007; Hilley and front of the advancing tip of the fold-and-thrust belt that was, Strecker, 2005). Hence they shelter the eastern slope of the inner and continues to be cannibalized as the belt widens (e.g., Kley, orogen from precipitation derived from the Atlantic (Fig. 1B). As a 1996; Horton, 1997). Magnitude of relative displacement varies consequence, most areas of the Sierras Pampeanas, and the from north to south: in the north, it accommodates between 74 Principal, Frontal, and Pre-cordilleras receive little moisture from and 135 km of displacement relative to stable South America the east (Sobel and Strecker, 2003). (Roeder,1988; Baby et al.,1989,1997; Roeder and Chamberlain,1995), in Since ~18 Ma, the latitudinal location of the central Andes has its center ~163 km (McQuarrie, 2002), and in the south between 125 remained relatively stable (Scotese et al., 1988). In addition, the and 150 km (Kley, 1996; Baby et al., 1997)( Fig. 4B, C, and D). Mecha- atmospheric and oceanic circulation that exists today appears to nically weak horizons in the Silurian, Devonian, and Carboniferous have been established since at least the early Cenozoic (Parrish part of the section accommodate the formation of many of the et al., 1982), and perhaps the Mesozoic (Hartley et al., 1992). In the decollement surfaces of the fold-and-thrust belt (e.g., Baby et al., broadest sense, the persistent latitudinal position and stable 1990; Dunn et al., 1995). The geometry of these detachment horizons atmospheric circulation has created a generally arid climate within also changes from north to south—in the north, the basal decollement the central Andes, with the western portions of the orogen of the fold-and-thrust belt is ~4°, whereas it decreases to ~2° in the characterized by the most intense aridity, while the eastern flanks south—this southward shoaling corresponds to a widening of the fold- generally receive far more moisture (Fig. 1; Strecker et al., 2007). and-thrust belt (Roeder, 1988; Roeder and Chamberlain, 1995; Kley, While this overall pattern has persisted since the Miocene, 1996; Horton, 1999). Likewise, some interpretations of GPS data paleoclimate proxies record significant climate changes through- (Norabuena et al., 1998) suggest that the distribution of active defor- out this period (Strecker et al., 2007, and references therein). mation within the fold-and-thrust belt changes along its length, with Furthermore, along the western Andean margin, late Miocene to early out-of-sequence thrusting characterizing the northern part of Pliocene paleoclimate indicators are somewhat contradictory and the wedge (Horton, 1999), and deformation localized at its eastern vary latitudinally (Sáez et al., 1999; Gaupp et al., 1999; May et al., toe in the south. Finally, the fold-and-thrust belt disappears to the south 2005); nonetheless, generally arid to semi-arid conditions with G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 87 periods of greater moisture availability typify the western margin of the ments. In the foreland, shortening becomes accommodated by orogen since the mid-Miocene, and perhaps as early as the inversion of Cretaceous normal faults in the thick-skinned Santa Eocene (Blanco et al., 2003). The termination of supergene Barbara ranges (Jordan et al., 1983; Kley and Monaldi, 2002). alteration within the Atacama Desert and volcanic arc between 14 Commensurate with these changes in total displacement, total and 8 Ma (Alpers and Brimhall, 1988; Sillitoe et al., 1991) suggest that shortening, and structural style, the mountain-belt width also arid conditions commenced in these areas by mid-Miocene time. decreases (Fig. 4H).At~26°S,theSantaBarbarasystembecomes In addition, the deposition of thick halite and gypsum-bearing transitionally replaced by the high-angle reverse faulted Sierras sedimentary units between 24 and 15 Ma within the Puna basins, Pampeanas to the south. Both these ranges focus Atlantic-derived indicate that by this time, the climate was arid (Alonso et al., 1991; precipitation on their eastern margin, resulting in large precipita- Vandervoort et al., 1995). tion gradients between the humid flanks and arid interior of the In contrast, sediments found in areas east of the Puna and orogen (Fig. 1B; Hoffmann, 1975; Sobel et al., 2003; Sobel and Altiplano show evidence for a morecomplexandvariableclimate Strecker, 2003; Hilley and Strecker, 2005). history that appears strongly affected by the uplift of individual Finally, between 28°S and 34°S, pre-Cenozoic deposits are ranges and the resulting downwind aridification (Strecker et al., absent in the foreland belt, and upper crustal shortening is taken 2007). For example, intramontane basins in the Eastern Cordil- up by reactivated Paleozoic structures across the Sierras Pampeanas lera of northwestern Argentina preserve sediments that record a (Allmendinger et al., 1983; Schmidt et al., 1995; Ramos, 2000; shift from semi-arid/arid conditions to humid environment in Ramos et al., 2002) that accommodate only 10–20 km of horizontal thelateMiocene(Marshall and Patterson, 1981; Pascual et al., displacement (Fig. 4D). The distributed nature of the deformation 1985; Pascual and Ortiz Jaureguizar, 1990; Nasif et al., 1997; widens the mountain belt and consequently decreases the total Kleinert and Strecker, 2001; Starck and Anzótegui, 2001; Coutand shortening (Fig. 4E). The subduction angle is low and therefore et al., 2006; Strecker et al., 2007; Anzótegui, 2007; Hilley and active volcanism is absent (Fig. 4G). Previous geologic studies have Strecker, 2005). Further to the north in the Subandean ranges of suggested that tractions and/or thermal heating at the base of the Bolivia, climate also shifted towards more humid conditions South American lithosphere may have caused the eastward inthelateMiocene(Uba et al., 2005, 2006), a time during which migration of deformation in this area (Jordan et al., 1983; Ramos megafan were developed in this area (Horton and DeCelles, et al., 2002)(Fig. 4H). In the arid westernmost portion of this zone 2001). In most of these marginal basins, the late Miocene change (Figs. 1 and 4), the re-emergence of a thick pre-Cenozoic to more humid conditions was abruptly terminated as the uplift sedimentary section corresponds with the development of the of upwind ranges prevented moisture from penetrating into Principal, Frontal, and the presently active Precordillera fold-and- these areas (Sobel et al., 2003; Sobel and Strecker, 2003; thrust belt (Fig. 4A, D, and F), where a larger amount of dis- Coutand et al., 2006; Strecker et al., 1989, 2007, Hilley and placement is accommodated than in the Sierras Pampeanas (Fig. 4A Strecker, 2005). The timing of the onset of ensuing aridification and D). varies from basin to basin, depending on the local tectonic In the Central Andes, the structural style (thin-skinned versus history that uplifted the upwind orographic barriers (e.g., thick-skinned fold-and-thrust belt, basement-cored uplifts) of the Strecker et al., 2007). This tectonic control on local precipitation external portion of the orogenic system mirrors the paleogeo- patterns is observed in the modern rainfall distribution on the graphic zones that existed prior to the Andean orogenic event eastern slope of the Andes (Figs. 1 and 4). In summary, both past (Allmendinger et al., 1983). The fact that the amount of horizontal and current climate generally show an arid or-hyperarid western displacement absorbed by the various structural provinces is slope, an arid interior, and an eastern slope that varies from markedly different suggests that latitudinal changes in deforma- humid to arid, depending on local topography and latitudinal tion may be dominantly controlled by factors set in motion long position. before the construction of the modern mountain belt (Allmendin- ger et al., 1983). Thus, the paleogeographic heritage likely triggers 4. Relationships between paleogeography, climate, and tectonics marked dissimilarities in the topography of the eastern Andean in the central Andes margin. In contrast, the topography of the high Plateau itself is remarkably symmetric (Fig. 3; Gephart, 1994). However, the The location of paleogeographic provinces, total shortening, amount and style of deformation accommodated along the eastern downgoing slab geometry, mountain belt width, and climate boarder of the high plateau within the Eastern Cordillera appears correspond to varying degrees within the central Andes (Fig. 4). to show marked asymmetry along its length. In Bolivia, the Plateau In summarizing the above discussion, between ~15° and 22.5°– boundary is seated within a thick section of Paleozoic metasedi- 23°S total shortening is large (40%–70%) and is concentrated ments, and a large amount of shortening is accommodated within within the Eastern Cordillera and the Subandean ranges (Fig. 4 C, the Eastern Cordillera and its backthrust belt (McQuarrie and D, and E). The subducted Nazca plate dips steeply eastward DeCelles, 2001; Elger et al., 2005). In contrast, along the Argentine beneath continental South America (Fig. 4G) and the location of Puna margin, the Paleozoic deposits generally tend to become the foreland fold-and-thrust belt coincides with pre-existing thick thinner, high-angle reactivated structures accommodate deforma- Paleozoic deposits containing multiple stratigraphic horizons that tion and total shortening decreases (Baby et al., 1997; Kley et al., serve as decollements (e.g., Allmendinger et al., 1983; Kley et al., 2005). Thus, while the location of the eastern margin of the 1999). However, from 17.5°S, the abrupt increase of the mountain Plateau appears to reflect the geometry of the downgoing Nazca belt's width (Fig. 4A) spatially coincides with a decrease in mean plate, the nature of deformation and the amount of displacement annual precipitation (WMO, 1975; Montgomery et al., 2001) accommodated along the margin appears controlled by the paleo- (Fig. 4B), the shallowing of the basal decollement angle of the geography of those areas. orogenic wedge (Roeder, 1988; Roeder and Chamberlain, 1995; The above observations allow us to draw some inferences about Kley, 1996; Horton, 1999) and the distribution of active deforma- the roles that slab geometry, paleogeography, and climate may play tion, as documented by GPS data (Norabuena et al., 1998; Horton, in shaping the central Andes. First, we observe that the extent of 1999). Between ~22.5°S and 26.5°S, the central Andes accommo- the Plateau coincides with major changes in the downgoing Nazca dates less shortening (30 to 40%) (Figs. 4C, D and E) and this plate geometry, suggesting a causal link between the two (Gephart, corresponds spatially to the disappearance of the subandean 1994). As an alternative hypothesis, others have suggested that the ranges and Paleozoic foreland deposits with interlayered decolle- paleogeography of the South American craton may have influenced 88 G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 the slab geometry itself: in the south where pre-existing high- relief in landscapes as channels aggrade and peaks degrade, and slope angle structures are pervasive within the crust, the slab was reduction decreases erosion in the orogen's interior; (2) Internal flattened as it was pushed underneath a continent that could not drainage and uplifting upwind orographic barriers prevent mass yield to the degree that it could in the north, where the Paleozoic evacuation from this region, increase lithostatic loads, and reduce the basin allowed large amounts of shortening to be accommodated isostatically driven rock uplift that would otherwise result from mass within the foreland (Allmendinger et al., 1983; Allmendinger and removal. Maintenance of lithostatic stresses by restriction of erosional Gubbels, 1996). Nonetheless, a combination of the paleogeography mass export may conspire with geodynamic processes to cause and subduction zone geometry appears responsible for the large- deformation to migrate to the external parts of the orogen (Sobel et al., scale architecture of the interior part of the orogen, rather than 2003). Together, geodynamic processes and an arid climate decouples climate playing a primary role as some have previously suggested the Plateau's base-level from the foreland, allowing a broad, low- (e.g., Montgomery et al., 2001; Sobel et al., 2003). In addition, the relief, tectonically quiescent Plateau to persist over geologic time in a amount and nature of deformation accommodated within the manner similar to the hypothetical case when erosion is absent (e.g., interior and external parts of the orogen varies from north to south, Royden, 1996). leaving a strong imprint on the topography of the mountain range: The foreland fold-and-thrust belts of Bolivia and Argentina (region (1) Total W–E shortening decreases stepwise to the south; 2inFig. 4A) are consistently associated with thick pre-existing (2) These variations are focused on boundaries of tectonic sedimentary units that contain mechanically weak layers that served provinces characterized by different styles of deformation; and as decollement horizons. These regions are externally drained; thus, (3) The total width of the mountain belt changes within each of erosional mass efflux increases with mountain-belt width and/or with these provinces. Changes in the precipitation distribution that steeper mean orogenic slopes parallel to material transport direction result from both regional air-mass circulations and local orographic (e.g., Dahlen and Suppe, 1988; Willett et al., 1993). In contrast to effects broadly coincide with these tectonic and topographic plateau environments, field (Hilley et al., 2004), experimental (e.g., boundaries. However, the spatial coincidence of these boundaries Persson et al., 2004; Hoth et al., 2006) and theoretical (e.g., Dahlen, with factors established prior to the Cenozoic makes it unlikely 1984) studies suggest that the voracity of erosional processes likely that the current climate plays a dominant role in determining the play a role in controlling the long-term deformation and topography. large-scale architecture of the central Andes. Piggy-back basins in fold-and-thrust belts may act as significant storage areas for sediment over 105 year timescales (Allen, 2008; 5. What role might climate play in the tectonics and topography of Hilley and Strecker, 2005). However, the resistant bedrock exposed the central Andes? within the ranges of the fold-and-thrust belt likely limit their rate of erosion (Whipple et al., 1999). Slope adjustments within the fold- If the paleogeography and plate interface geometry have left such a and-thrust belt generally take ~106–107 years to balance erosional strong imprint on the morphology of the modern central Andes, then mass efflux with the flux of tectonically accreted material (e.g., Stock what role, if any, might climate have played in controlling different and Montgomery, 1999; Hilley et al., 2004). Both the Bolivian and tectonic features? The coincidence of pre-Andean paleogeography Argentine fold-and-thrust belts have been active since the Miocene, with the current Andean tectonic, climatic, and topographic attributes suggesting that erosion and deformation in these areas have at least argues that most of the orogen-scale features of the central Andes partially adjusted to one another. This adjustment may be reflected by have been shaped in large part by this paleogeography. However, it is changes in the geometry of the fold-and-thrust belt in space and/or possible that at a more local scale, climate, erosion, and tectonics may time. For example, in northern Bolivia, a southward decrease in mean be linked. To investigate this option, we further analyze three distinct annual precipitation coincides with a widening of the fold-and- morphotectonic provinces in the central Andes: (1) the internally thrust belt (Fig. 4A and B) and perhaps a transition from diffuse to drained Altiplano–Puna Plateau, (2) the foreland fold-and-thrust belts localized deformation along the topographic front of the Bolivian of Bolivia and Argentina, and (3) the Sierras Pampeanas basement- fold-and-thrustbelt(Horton, 1999). Theoretical considerations cored uplift province (Fig. 4; labeled region 1, 2, and 3 respectively). suggest that weaker erosion, presumably caused by decreased The Altiplano–Puna plateau typically receives b400 mm of annual precipitation rates, results in widening of the orogen to evacuate precipitation (WMO, 1975; Sobel et al., 2003; Bookhagen and Strecker, tectonically accreted foreland sediments (e.g., Whipple and Meade, 2008), and stratigraphic data indicate that the current western and 2004). Also, studies of the Argentine Principal, Frontal, and Pre- eastern Altiplano margins were uplifted in the Eocene–Oligocene, and cordilleras (Fig. 4) indicate that changes in erosion rates, prompted Oligocene–Miocene, respectively, leading to the formation of internal by the exposure of different rock types, may result in changing drainage in the area by ~25 Ma (Horton et al., 2002). Likewise, in the fold-and-thrust belt geometry and topography over time (Hilley et Puna, uplift of its present eastern margin and disparate basement al., 2004). Thus, theory predicts that externally drained areas uplifts within the plateau region may have led to the formation of presenting the proper geologic conditions for the formation of fold- internal drainage between Oligocene and Miocene time (e.g., Coutand and-thrust belts, may show a strong imprint of climatically- et al., 2001, Vandervoort et al., 1995; Coutand et al., 2006). While the moderated erosional processes on their morphology (Montgomery extent and symmetry of the Plateau appears to reflect the downgoing et al., 2001). In the central Andes, segments of the mountain belt Nazca plate geometry, erosion may help to reinforce the nature of the experiencing less erosion, either due to lower precipitation rates topography and tectonics of this area. In particular, the establishment (Montgomery et al., 2001) or to the exposure of more resistant of internal drainage decouples the local base-level of streams draining lithologies (Hilley et al., 2004), tend to be wider than more erosive the hinterland from their regional base-level, allowing large-scale areas. This is consistent with the view that the geometry of fold- aggradation to occur (Sobel et al., 2003; Hilley and Strecker, 2005). and-thrust belts depends on the efficiency of erosion (e.g., Dahlen Integrated field studies and theoretical models indicate that the and Suppe, 1988). However, within the Bolivian fold-and-thrust disconnection and reintegration of individual basins base-levels may belt, recent area balancing of structural sections found little occur over timescales of ~105 years (Allen, 2008; Hilley and Strecker, difference in shortening within this area, which may suggest that 2005). The fact that regional disconnection of the foreland and Plateau erosion may exert less of a control on orogenic structure here than base-levels have persisted for much longer (e.g., Vandervoort et al., previously thought (McQuarrie et al., 2008b). 1995) indicates that conditions that favor this disconnection have In contrast to fold-and-thrust belts, deformation along reacti- persisted over these longer timescales. This has two important effects vated crustal-scale heterogeneities such as in the Sierras Pampeanas on erosional–tectonic coupling: (1) Regional aggradation reduces (region 3 in Fig. 4A) involve steep, pre-existing crustal-scale G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92 89 weaknesses that may undergo frictional failure without necessitat- geologic history of the central Andes sets the stage for the nature and ing failure within the rest of the mountain range. An analysis of the strength of coupling between erosional processes, tectonic deformation, stresses expected in this situation suggests that the development of and mountain belt structure. This situation reflects the complex erosionally moderated surface slopes should be decoupled from the feedbacks that we might expect to arise between the pre-existing reactivation of these structures (Hilley et al., 2005), except when rheological state of the crust, the geodynamic boundary conditions, and mean slopes across the range exceed a threshold value. In this case, climate-moderated erosional processes (Fig. 2). the large lithostatic loads produced by these slopes oppose Several important lessons may be gathered from interactions horizontal stresses and favor failure of frictionally stronger struc- observed within the central Andes. Most importantly, the nature of tures without such loads. Because erosional processes moderate coupling between erosion, climate, pre-orogenic geologic history, plate- surface slopes, vigorous erosion may prevent the build-up of steep boundary forces, and mantle processes is complex, and the impact of the slopes, and hence the migration of deformation to other structures initial state of all of these variables may not easily be erased over the life of in low-lying areas. As in fold-and-thrust belts, bedrock erosion is the orogen. While our previous work has tried to view the orogenic likely the process that limits the timescale over which erosion structure in terms of simple interactions between these factors (Sobel et responds to tectonic uplift. However, the time it takes to build al., 2003, Hilley et al., 2004; Hilley and Strecker, 2004; Hilley et al., 2005) topography significant enough to cross this threshold slope depends and the work of others have attempted to understand overall orogenic on rates of uplift produced by fault motion, the erosional efficiency, development in terms of simple erosional and geodynamic feedbacks the value of the threshold slope, and the geometry of the underlying (e.g., Willett et al., 1993; Whipple and Meade, 2004), these approaches fault. This time-scale varies depending on local circumstances, but may not be altogether appropriate for an orogen such as the central Andes likely occurs over ranges of 104–106 years. These theoretical that has been subjected to a long and rich tectonic history. They may fail considerations shed light on the development of two mountain to reveal the complex interactions between tectonic and erosional ranges within the northern Sierras Pampeanas that have undergone processes that characterize these types of orogens. However, because similar geologic histories but have vastly different kinematics. Those factors affecting the Earth's mantle, crust, and surface at the onset of and ranges in wet areas tend to have deformation focused on a limited during mountain building profoundly influence the development of the number of structures with large amounts of displacement, while orogen throughout its life, we argue that the path to understanding these those in drier sectors display more widely distributed deformation orogens requires an integrative approach that combines geodynamic, along several faults with relatively small individual displacements geologic, and geomorphologic observations and models. (Sobel and Strecker, 2003). Unfortunately, the location of zones of pre-existing weakness, their frictional properties, and their local Acknowledgments loading conditions cannot be easily measured, and so a strict test of the idea that erosion may control a threshold in the development of This work has been supported by the France-Stanford Center for these ranges remains elusive. Nonetheless, the details of the Interdisciplinary Studies and by the Fulbright Scholar Program. development of individual ranges in the Sierras Pampeanas are consistent with qualitative theoretical expectations, and so it is Appendix A. Supplementary data possible that the details of the development of individual ranges can fl be in uenced by climate moderated erosional processes. Supplementary data associated with this article can be found, in While strict tests linking deformation to erosion remain elusive, the online version, at 10.1016/j.tecto.2009.06.017. detailed geologic observations are consistent with theory that links the topographic and tectonic development of different components of this References orogen. In some cases, erosion may be strongly coupled to tectonic deformation (e.g., fold-and-thrust belts), while it is less influential in Aceñolaza, F.G., 1973. Sobre la presencia de Oldhamia sp. en la Formación Puncoviscana others (e.g., internally-drained plateaus). However, even in regions de Cuesta Muñano, Provincia de Salta, República Argentina. Revista de la Asociación such as the internally drained plateau, the voracity of erosion may have Geológica Argentina, 28: 56. Allen, P.A., 2008. Time scales of tectonic landscapes and their sediment routing systems. left an important imprint on the landscape. 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