links between topography, erosion, rheological...

Tectonophysics 495 (2010) 78–92

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Tectonophysics

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Links between topography, erosion, rheological heterogeneity, and deformation incontractional 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 Statesb Laboratoire Géosystèmes (UMR 8157 CNRS), Université de Lille 1, 59655 Villeneuve d'Ascq, Francec Department of Earth Sciences, Dalhousie University, Halifax, NS, Canada B3H 4J1

⁎ Corresponding author.E-mail addresses: [email protected] (G.E. Hilley),

0040-1951/$ – see front matter © 2009 Elsevier B.V. Aldoi:10.1016/j.tecto.2009.06.017

a b s t r a c t
a r t i c l e i n f o
Article history:Received 30 May 2008Received in revised form 7 April 2009Accepted 15 June 2009Available online 23 June 2009

Keywords:Contractional orogensErosion/tectonic couplingRheological heterogeneityClimate

Orogenic structure in the central Andes (15°S–34°S) systematically varies with mean annual precipitation,suggesting that erosional processes may be coupled to tectonic processes. We explore the range of possibleinteractions between deformation, erosional processes, changing geodynamic conditions, and preexistinggeologic structures to assess the relative importance of each. Our review and synthesis indicates that the pre-orogenic geologic history and changing geodynamic conditions leave a lasting imprint on the modernstructure of the mountain belt. In contrast, changes in precipitation that result from regional-scale atmo-spheric and oceanic circulation interact with local topography to control the climatic zoning across thecentral Andes. Changes in erosion, which likely correspond to changes in climate, rock uplift rate, and theexposure of different lithologies, may in turn affect the tectonic development of the region at the scale ofeach morphotectonic province. In this view, geologic and geodynamic factors set the stage for the nature andstrength of coupling between erosional processes, tectonic deformation, and mountain belt structure in thedifferent parts of the central Andes.

© 2009 Elsevier B.V. All rights reserved.

1. Introduction

Over the last three decades, earth scientists have realized thatprocesses that redistribute mass across Earth's surface mayinfluence deep earth processes (e.g., Davis et al., 1983; Beaumontet al., 1992; Willett et al., 1993; Willett, 1999). For example, globaland regional climate changes may enhance or subdue erosionwithin mountain belts, and these changes result in mass redis-tribution patterns that may affect the near-surface thermal state ofthe crust (e.g., Ehlers, 2005), the body forces within the crust (e.g.,Dahlen, 1984), and cause changes in mean elevations that result inisostatic adjustments (e.g., Molnar and England, 1990a). Many ofthe concepts linking surface processes to those deep within theearth arose from observations from Taiwan (e.g., Davis et al., 1983;Dahlen and Suppe, 1988), the Himalayas (e.g., Hodges, 2000;Thiede et al., 2005; Grujic et al., 2006), and geodynamic models(e.g., Beaumont et al., 2004; Stolar et al., 2007). In the centralAndes, tectonic deformation and erosional patterns also correlatewith one another, suggesting that here too they may be linked (e.g.,Horton, 1999; Montgomery et al., 2001).

[email protected] (I. Coutand).

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The South American convergent plate margin mostly consistsof the subduction of the Nazca Plate beneath South America alongan eastward-dipping subduction zone (Stauder, 1975; Barazangiand Isacks, 1976, 1979; Hasegawa and Sacks, 1981; Bevis and Isacks,1984; Cahill and Isacks, 1992; Allmendinger et al., 1997). Volcanicand tectonic processes have deformed and uplifted this margin,which hosts a variety of landscapes, including volcanoes in theWestern Cordillera, internally drained basins in the high-elevationAltiplano–Puna Plateau, thin- and thick-skinned fold-belts withinthe Eastern Cordillera and Subandean ranges, and basin-and-range-like topography created by basement-cored uplifts of theSierras Pampeanas (Fig. 1). In the central Andes, changes intopography and deformation style have been attributed to a varietyof tectonic factors, including the geometry of the down-goingsubducting plate (Isacks, 1988; Gephart, 1994), the thickness ofthe mantle lithosphere in the central segment of the mountain belt(e.g., Kay et al., 1994; Whitman et al., 1996; Garzione et al., 2006),interplate coupling (e.g., Gephart, 1994; Lamb and Davis, 2003),and pre-existing heterogeneities in the crust and mantle thatreflect the pre-Cenozoic geologic history of the area (e.g., Allmendingeret al., 1983; Coutand et al., 2001; Ramos et al., 2002). However, changesin erosion related to climate, and the exposure and erosion of differentrock types have been inferred to play a role in determining the patternsof deformation and topographyobservedwithin this orogen (e.g.,Maseket al., 1994). For example, the intensity of erosion that likely varies

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http://dx.doi.org/10.1016/j.tecto.2009.06.017

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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 reliefmap 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 showdepth to the downgoing Nazca plate, taken from Cahill and Isacks (1992), Bevis and Isacks (1984), and Barazangi and Isacks (1976).

79G.E. Hilley, I. Coutand / Tectonophysics 495 (2010) 78–92

because of climate patterns related to long-term global-scaleatmospheric circulation (e.g., Haselton et al., 2002; Parrish et al.,1982, Hartley et al., 1992; Masek et al., 1994; Strecker et al., 2007),may affect mountain-belt width (e.g., Willett, 1999; Montgomeryet al., 2001) or control the interplate frictional coupling bymodulating the sediment supply to the subduction trench (e.g.,Lamb and Davis, 2003). At a regional scale, studies have arguedthat erosion plays a key role in controlling the kinematics,geometry and pattern of deformation within the subandean fold-and-thrust belt (e.g., Masek et al., 1994; Horton, 1999), theArgentine Precordillera (Hilley et al., 2004), and the SierrasPampeanas (Sobel and Strecker, 2003; Hilley et al., 2005).

While erosion may play a role in shaping the central Andes, the first-order characteristics of the topography and deformation of this orogencorrelate strongly with changes in geodynamic features such as down-going slab geometry (e.g., Isacks, 1988; Cahill and Isacks, 1992; Gephart,1994; Allmendinger et al.,1997), and geologic heterogeneities in the crustthat include inherited structures and paleogeographic features (e.g.,Schmidt et al., 1995; Allmendinger et al., 1983, 1997; Ramos et al., 2002;Carlier et al., 2005). Based on expectations derived from the theoreticalframework that links erosion to tectonic processes, the strength of thelinkages between these factors should vary spatially depending on the

rheology of the crust and its heterogeneities, the lithologies exposed atthe surface, and regional to local climatic conditions. In this contribution,we review geological and geophysical observations, as well as modelingstudies to examine the strength and nature of coupling between tectonicand erosional processes in the central Andes. The spirit of this workfollows that of previous geological studies, inwhich correlations betweennatural phenomena are documented and interpreted in the context ofexperimental (e.g., Persson et al., 2004; Hoth et al., 2006) and theoretical(e.g., Dahlen, 1984) models that rationalize linkages between theobserved phenomena. Unfortunately, one cannot manipulate the under-lying tectonic processes or the geological history of natural orogens as isthe case in a controlled laboratoryenvironment (e.g., Persson et al., 2004;Hoth et al., 2006), and sowhile not wholly satisfying, we follow the pathof others inwhichwe seek correlations between various parameters thatcan be understood in terms of these simpler systems (e.g., Hoffmann-Rothe et al., 2006).

2. Conceptual framework linking tectonic and erosional processes

In the simplest conception, the state of stress within tectonic platesresults from the sum of tractions at the plate boundaries due to plate–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 andmantle, and stresses produced by the overburden of crustal material.The rheology of the crust relates deformation and/or deformation rateto these loading conditions (Fig. 2). The rheology of rocks depends onmany factors including composition, temperature, loading stresses,and grain size (e.g, Bürgmann and Dresen, 2008, and referencestherein), and so deformation may directly modify crustal rheologythrough at least three mechanisms: 1) heat redistribution in the crustdue to magmatism, exhumation, and frictional failure in the uppercrust, (e.g., Byerlee, 1968; Byerlee, 1978; Brace and Kohlstedt, 1980;Graham and England, 1976; Lachenbruch and Sass, 1980; Ryback andCermak, 1982; Edman and Surdam, 1984; Molnar and England, 1990b;Furlong et al., 1991; Royden et al., 1997; Beaumont et al., 2004;Byerlee,1978) Lachenbruch and Sass, 1980; Ryback and Cermak, 1982;2) grain size reduction by dynamic recrystallization (e.g., Rybacki andDresen, 2000; Bystricky andMackwell, 2001; Rutter and Brodie, 2004;Dimanov and Dresen, 2005); and 3) spatial redistribution ofrheologically distinct crustal units during deformation (e.g., Allmen-dinger et al., 1983).

Changes within the deep earth also influence the rheology of rockat shallower levels; for example, changes in mantle temperatures mayeventually be manifest by temperature changes within the crust (e.g.,Sleep, 2005, and references therein; Fig. 2). Convective removal ofmantle lithosphere or magmatic underplating may alter the crust andmantle lithosphere's thicknesses and consequently, the thermal stateof the crust (e.g., Kay et al., 1994; Conrad and Molnar, 1997). Finally,the rheological configuration of the crust prior to orogenisispotentially leaves a lasting imprint on the subsequent evolution ofthe deforming crust. Thus, in active convergent orogens, thedistribution of deformation can be spatially non-uniform and evolveover time as the state of stress and rheology of the crust change inspace and time.

Finally, erosion of Earth's surface can play a fundamental role indeformation by modifying the distribution of lithostatic stresses inthe crust (Fig. 2). For example, as surface elevations increase, so do thelithostatic stresses within the crust, and this may favor failure alongthe extremities of an orogen at the expense of internal shortening(e.g., Dahlen, 1984; Hilley et al., 2005) and may elicit an isostaticresponse where mean elevations change (Molnar and England,1990a). By implication, the state of stress and deformation withinEarth's crust is, at least in part, controlled by erosional processes. The

Fig. 2. Schematic representation of potential feedbacks between crust

strength of these erosional processes is likely moderated by factorsincluding climate (e.g., Knox, 1983; Bookhagen et al., 2005a,b;Strecker et al., 2007) and the lithology of the rocks exposed at thesurface (e.g., Gilbert, 1877; Hack, 1960; Stock and Montgomery, 1999;Sklar and Dietrich, 2004). In turn, local climate is strongly influencedby topographic relief that creates orographic gradients in precipitation(e.g., Barry, 1981; Bookhagen and Burbank, 2006), and global climateis perturbed by the uplift of large portions of Earth's surface such asthe Tibetan Plateau (e.g., Raymo and Ruddiman, 1992; Ruddiman,1997). Thus, surface uplift, climate, erosion, and deformation areinextricably linked in actively deforming orogens.

The lithology of rocks exposed at the surface partially controlserosional efficiency and topographic slopes. Consequently, as orogensevolve, changing distribution of lithologies likely changes the way inwhich deformation is accommodated. For example, as contractionaldeformation propagates toward the foreland, the exposure of poorlyconsolidated basin sediments may enhance erosion, which allowslarge quantities ofmaterial to be eroded evenwhere relief is subtle, anddelays the accumulation of lithostatic stresses and advancement ofdeformation to the orogen's external portions (e.g., Hilley et al., 2004).Similarly, orogens in which surface processes are veracious due toabundant precipitation may exhume deep crystalline rocks whoseresistance to erosion is high. The surface exposure of such competentrocks promotes the rapid steepening of orogenic slopes, increaseslithostatic loadswithin the orogen's interior, and forces deformation tomigrate towards the external portions of themountain belt (e.g., Hilleyet al., 2004; Hilley and Strecker, 2004). In addition, where erosion isinefficient during the waxing stages of orogenesis, the balancebetween lithostatic stresses at the base of the crust and the restoringstresses that buoyancy provides may be altered, causing the thickenedcrust to partially sink into themantle (e.g., Stüwe, 2002). As orogenesiswanes, lowering of the surface reduces the lithostatic stresses thatoppose buoyancy stresses, causing isostatic uplift of rocks (Molnar andEngland, 1990a). Thus, erosional removal of mass from Earth's surfacecan modify the local lithostatic stresses acting along individualstructures and within the frictional medium of the crust, as well asthe balance between lithostatic loading and the restoring stressesprovided by the relative buoyancy of the crust in mantle (Fig. 2).

In the simple scenarios outlined above, deep earth processesinteract with surface processes through the alteration of crustallithostatic loading and isostatic responses in complex ways that

al rheology, plate boundary forces, mantle processes, and erosion.


depend on the history of deformation, erosion, and heat transferwithin the crust. In addition, the pre-existing rheological state ofthe crust likely influences the deformation history during orogen-esis. Consequently, the initial distribution of physical properties ofthe crust when combined with global and local climatic effects,establishes the strength of coupling between erosion and deforma-tion throughout the history of orogenesis.

3. Physiographic, geologic, and climatic framework of thecentral Andes

3.1. Physiography

The central Andes are located between 70°W and 60°W, and15°S and 35°S (Fig. 1) along an active subduction margin in whichthe oceanic Nazca Plate is translated ENE relative to stable SouthAmerica and is underthrust beneath the South American Platealong an eastward dipping subduction interface (Stauder, 1975;Barazangi and Isacks, 1976; Hasegawa and Sacks, 1981; Bevis andIsacks, 1984; Cahill and Isacks, 1992; Allmendinger et al., 1997).This geodynamic configuration has led to the formation of a broadswath of high topography that roughly parallels the subductionzone (Fig. 1). This topography reaches peak elevations of 6962 mat Mt. Aconcagua (Argentina), the highest point in the westernhemisphere. The elevation and extent of the high topographychanges with latitude along the subduction zone: at ~18°S, themeridianal extent of the Andes reaches widths of ~750 km,whereas it becomes narrower both north and south of this location(Fig. 1).

A number of morphotectonic segments can be defined in thecentral Andes based on topographic attributes. The forearc domaincomprises, from west to east, the continental margin, the CoastalCordillera (up to 2000 m in elevation), the Central Depression, andthe North Chilean Precordillera reaching an average elevation of3000 m (e.g., Reutter et al., 2006). To the east, the WesternCordillera consists of the Miocene to Quaternary volcanic arc atopthe subduction zone (e.g., Isacks, 1988; Fig. 1). Between thelatitudes of 15°S and 26°S, the Western Cordillera is flanked onits eastern side by the high-altitude internally drained Altiplano–Puna Plateau (Fig. 1). Elevations within the Altiplano average~3700 m (Isacks, 1988), whereas the Puna is generally higher andhas average elevations of ~4400 m (Whitman et al., 1996). Acrossthe Altiplano, the Plateau surface is mostly flat; however, the Punais partitioned into a series of smaller basins by generally N–S-trending basement ranges that reach elevations of about 5000 m(Coutand et al., 2001; Fig. 1). South of the Puna (N26°S), thePrincipal, Frontal, and Pre-Cordilleras, which consist of a series ofrelatively narrow, N–S trending ranges that are externally drained(e.g., Jordan et al., 1993; Fig. 1). To the east at these latitudes, thecentral Andes consist of the Sierras Pampeanas, which is a zoneof discontinuous and widely spaced mountain ranges that arebroadly distributed across the foreland (Fig. 1; Stelzner, 1923;González-Bonorino, 1950; Jordan and Allmendinger, 1986). Bound-ing the eastern edge of the internally drained plateau north of~26°S is the Eastern Cordillera, which consists of a series of highranges that generally serve as the drainage divide between theinternally drained plateau and the externally drained foreland(Kennan et al., 1995; Kley, 1996). In a transitional area between~24°S and ~26°S, the Eastern Cordillera and the Sierras Pampeanasintermingle, with some Pampean-style ranges located to the east ofthe Eastern Cordillera. However, north of ~24°S, these vanish andgive way to a series of ranges that closely flank the EasternCordillera called the Santa Barbara system (Rolleri, 1976; Baldiset al., 1976). Finally, north of the Santa Barbara system, a series ofranges lying east of the Eastern Cordillera comprise the Inter-andean and Subandean ranges (Allmendinger et al., 1983; Fig. 1).

On average, elevations taper gently toward the foreland from6000–5000 to 200–100 m, although high local slopes result fromthe partitioning of the foreland into a series of meridianallyextensive, rhythmic and coherent ridges.

3.2. Relationship of morphology to current plate geometry

Thewidth and height of the orogenmimic changes in the geometryof the subducting Nazca plate. In the portion of the central Andesbetween 15° and 27°S, where the orogen reaches a maximum widthof 650–700 km, the downgoing Nazca plate dips steeply (~30°) downto 600 km beneath South America (Fig. 1B). By the time the Altiplano–Puna Plateau tapers both to the south and north, coinciding with asignificant narrowing of the orogenic system, the slab partially shoalsand dips shallowly (~5°) between depths of 100 and 125 km under-neath South America (e.g., Bevis and Isacks, 1984, Barazangi andIsacks, 1976; Cahill and Isacks, 1992). This phenomenon is reflected inthe symmetry of the Andean topography—the pole to the great circleabout which topography is most symmetric matches that about whichslab dip direction is also most symmetric (Gephart, 1994). In fact, thispole also corresponds to the modern rotation pole of the Nazca platerelative to a fixed South American plate, suggesting that the symmetryof the orogen is tightly linked to the subduction of the Nazca platebeneath South America (Gephart, 1994).

We show the topography of the central Andes projected aboutthe great circle defining its symmetry axis in Fig. 3A. When lookingat the average difference between elevations located north andsouth of this symmetry axis, the orogen is remarkably symmetric(Fig. 3B), although important excursions from this symmetric viewof the Andes exist. When considering the average differencebetween elevations (e.g., Gephart, 1994), differences within theexternal portions of the margin are naturally de-emphasized, sincethey lie at low elevations and are unlikely to have the correspond-ing differences in elevation expected in steep terrains. To adjust forthis, we normalized the difference between points north and southof this symmetry axis to their mean elevation. This allows smalldifferences in low relief areas to have similar weight as largedifferences at high elevation (Fig. 3C). Using this approach, we seethat the center portion of the mountain belt constituted by theAltiplano–Puna Plateau is highly symmetric; however, topographyalong the margins and within the foreland displays markedasymmetries (Fig. 3C). In particular, the enclave into the eastern-most portion of the Andes north of the symmetry axis that isabsent to the south constitutes a highly asymmetric feature thatwas obscured using the dimensional difference (compare Fig. 3Band C). In addition, the margins of the Eastern Cordillera, as well asthe whole of the Pampean and Santa Barbara ranges are highlyasymmetric with respect to their counterparts to the north. Thus,while the extent and location of the Plateau appears highlysymmetric, the external portions of the mountain belt, includingthe Eastern Cordillera, Principal, Frontal, and Pre-Cordilleras, theSierras Pampeanas and Santa Barbara system, and the Interandeanand Subandean ranges are far less so.

3.3. Pre-Cenozoic geologic history

The oldest rocks exposed within the back-arc area, south of 15°Sare located within the Sierras Pampeanas (Fig. 1) and includeMesoproterozoic high-grade metamorphic basement consisting ofgneisses and migmatites (Mon, 1979; Ramos et al., 1986) that may belocally intruded by several generations of late Neoproterozoic to earlyPaleozoic granitoids (Halpern and Latorre, 1973; Lork and Bahlburg,1993). Late Neoproterozoic to Early Cambrian greenschist-gradeturbidites and pelagic clays of the Puncoviscana Formation appearwithin the Puna and adjacent Eastern Cordillera (NWArgentina andsouthernmost Bolivia) and were deposited in a deep marine basin

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 circlewhose 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 asymmetrywas 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 becauseof 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 eachpoint 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 inwhich comparisons ofthe topographic symmetry were not possible due to the presence of the Pacific Ocean.


(thickness is ~2000 m; Ruiz Huidobro, 1960; Turner, 1960a,b;Aceñolaza,1973; Omarini, 1983). Importantly, the transition betweenthe crystalline basement and the Puncoviscana Formation is abrupt,and corresponds to the boundary between the Sierras Pampeanasand Eastern Cordillera at ~26°S (Fig.1).Whether this sharp transitionrepresents a buttress unconformity between a southerly crystallinebasement high (potentially the Pampia terranes, e.g., Ramos, 2008)within the Early Cambrian, or motion along a fault is unclear; how-ever, it certainly marks a persistent and profound change in thepaleogeography of the central Andes that occurred at least as early as

Fig. 4. Correlations between Central Andean (A) physiography, (B) mean annual precipSouth America, (D) horizontal displacements within each morphotectonic province relaand L, total displacement (see Table A1), (F) pre-Cenozoic foreland deposit thickness,style, and (J) precipitation distribution. Physiography modified after Isacks (1988) andtext. Mean annual precipitation taken from WMO (1975); thick black lines show locationprofiles reference specific sources given in Table A1). Calculations for total displacementforeland deposit thicknesses include the sediments that are undisrupted by Cretaceoudistance between slab depths of 100 and 150 km presented in Cahill and Isacks (1992

Cretaceous time, and perhaps during the Proterozoic to earlyPaleozoic (Allmendinger et al., 1983). These units are locallyunconformably covered by Cambrian quartzites (Hausen, 1925)deposited in an extensional back-arc basin (Salfity et al., 1976)associated with an extensional episode affecting the southwesternmargin of Gondwana (Ramos et al., 1986; deWit and Ransome, 1992;Bahlburg and Furlong, 1996). During the Ordovician, an up to 4000-m-thick turbiditic section is deposited across the Puna and EasternCordillera in Argentina (Bahlburg and Breitkreuz,1991; Bahlburg andFurlong,1996). In the late Ordovician, these units were deformed and

itation, (C) total horizontal displacements of the back arc region relative to fixedtive to fixed South America, (E) total shortening: (L/Lo) ×100 with Lo, initial length(G) slab angle between 100 and 150 km, (H) mountain belt width, (I) structuralCoutand (1999); numbered morphotectonic regions (1–3) are referenced to theof sections used to calculate displacements and shortening (Table A1; numbers on

, displacement by location, and total shortening are given in Table A1. Pre-Cenozoics rifting and are compiled in Table A2. Subduction angle calculated by measuring). Mountain-belt width measured perpendicular to structures shown on Fig. 1.


metamorphosed to greenschist grade (Turner and Méndez, 1979;Coira et al., 1982; Mon and Hongn, 1987) and subsequently intrudedby a granitoid and rhyodacitic belt known as the Faja Eruptiva de laPuna Oriental (FEPO; Mendez et al., 1972). The FEPO is interpreted asrepresenting a late Ordovician to Silurian magmatic arc that resultedfrom a subduction zone located to the west of the Puna, which likelydipped to the east (Coira et al., 1982). In northwestern Argentina,Silurian–Devonian sedimentation is mainly localized in the present-day Subandean ranges and eastern part of the Eastern Cordillera, andcomprises about 2000 m of marine sandstones and shales (Turner,1960a,b; Padula et al., 1967). To the west of the FEPO, Carboniferousand Permian plutons indicate that this margin was active during theLate Paleozoic.

The activity of the convergent margin during the Paleozoic ispartially represented within the sedimentary sequence lying east ofthe current Altiplano–Puna Plateau margin. These Paleozoic strataare present north of the crystalline basement/Puncoviscana transi-tion around ~26°S (Allmendinger and Gubbels, 1996, their Fig. 6). Inparts of the Sierras Pampeanas to the south, lower Paleozoic strataconsist of both marine and continental strata deposited in anextensional to trans-tensional tectonic environment, while Permianstrata record filling and amalgamation of these sub-basins (Fernan-dez-Seveso and Tankard, 1995; Fisher et al., 2002). In total, thesePaleozoic strata generally reach thicknesses of b2.5 km. As anexception, Paleozoic sediments currently exposed within thePrecordillera of Argentina (Fig. 1) reach thicknesses of up to 7 km(e.g., Zapata and Allmendinger, 1996). Between ~26°S and ~24°S,Paleozoic sedimentary rocks are often present, but generally oflimited thickness. However, north of ~24°S, a thick, eastwardlytapering section of Paleozoic clastic rocks reaches a thickness of~12 km in the Eastern Cordillera of Bolivia and thins to 4–9 km atthe edge of deformation in the Chaco Plain (Allmendinger et al.,1983; Allmendinger and Gubbels, 1996). These rocks were appar-ently deposited in the retro-arc basin of the Paleozoic volcanic arc(e.g., Ramos, 2008). Thus, as with the Precambrian and earliestCambrian rocks in the central Andes, the Paleozoic section starts tothicken around ~26°S, reaches maximum thicknesses in the currentBolivian foreland (e.g., Allmendinger and Gubbels, 1996; McQuarrieand DeCelles, 2001, and references therein) and finally tapers againnorthward to disappear around 17.5°S (Baby et al., 1994).

Mesozoic strata exposed within the central Andes generallyrecord an episode of intracontinental rifting starting in theNeocomian (Salfity, 1982). This rift basin system has a complexgeometry with N–S to E–W elongated depocenters (see Fig. 1 inMarquillas et al., 2005) filled with strata of the Salta Group (Turner,1959), which is Early Cretaceous to Late Eocene in age (e.g.,Marquillas et al., 2005). Deformed remnants of these basins arepreserved in the Altiplano–Puna Plateau, the Eastern Cordillera andthe foreland thrust belts. In the latter, Mesozoic strata are of limitedthickness or altogether absent south of ~27.5°S, reach a maximumthickeness of ~4–5 km between 25° and 26°S in the Santa Barbarasystem and disappear north of 23°S when the Lomas de Olmedobranch diverges eastward (see Fig. 1 in Marquillas et al., 2005; Salfityand Marquillas, 1994).

A plot of pre-Cenozoic stratal thickness in the modern forelandwith latitude shows the abrupt changes in thickness that accompanythe widespread exposure of Precambrian basement south of ~26°S(Fig. 4F). Importantly, changes in basin thickness correlate stronglywith the different morpho-tectonic provinces in the central Andes:the thick (up to 17 km) pre-Cenozoic foreland basin units areassociated with the fold-and-thrust belts in the Subandean and thePrecordillera fold-and-thrust belts, the irregular Cretaceous rift basinsare associated with thick-skinned deformation in the Santa Barbarasystem, and relatively thin (or absent) pre-Cenozoic basin strata areoften associated with the Sierras Pampeanas basement-cored upliftprovince (Fig. 4F).

3.4. Cenozoic geologic history and deformation

The style, location, and amount of Cenozoic deformation correlateswith zones defined by the current plate-boundary geometry, as well asby pre-Cenozoic paleogeography of the central Andes. While thetiming of Andean deformation within the Puna–Altiplano Plateau andbounding Eastern Cordillera are somewhat disputed and certainlyhighly diachronous, the commencement of modern deformation inthis area is broadly constrained to be between late-Eocene and mid-Miocene time (e.g., Allmendinger, 1986; Horton, 1997, 1998; Lamb andHoke, 1997, Marrett and Strecker, 2000; Coutand et al., 2001, 2006;Hongn et al., 2007).

3.4.1. Altiplano–Puna PlateauApatite fission-track thermochronology data from the Alti-

plano–Puna Plateau indicate that exhumation of this part of theorogen has been limited (Carrapa et al., 2005; Deeken et al., 2006;Ege et al., 2007; Letcher, 2007). The Bolivian Altiplano recordsexhumation in the Early Oligocene (33–27 Ma) at slow andconstant rate of 0.2 mm/yr until the late Miocene (11–7 Ma) (Egeet al., 2007). In the Puna Plateau, the onset of exhumation isrecorded by apparent AFT ages between 24 and 29 Ma (Deekenet al., 2006; Carrapa et al., 2005; Letcher et al., unpublished data). Ifexhumation serves as a crude proxy for total tectonic shorteningand rock uplift, these data suggest that shortening within theinterior of the Altiplano–Puna Plateau has been limited during thelast 20–25 Ma. The deformation that did occur within the Puna ismore distributed than the Altiplano Plateau to the north and hasuplifted N–S-trending bedrock blocks that partition the plateauinto a series of ranges and adjacent basins. These basins contain thedetritus that presumably resulted from uplift and exhumation ofthe ranges in the Plateau interior (Alonso et al., 1991; Vandervoortet al., 1995; Coutand et al., 2001). Basins that record thepartitioning of the Puna into disconnected depocenters may haveformed as early as the Oligocene, and certainly by the Miocene(Vandervoort et al., 1995). In fact, recent sedimentologic andstructural data from the Puna have been interpreted as evidence forEocene exhumation and depocenter development (Carrapa andDeCelles, 2008). Some of these basins are up to 5-km thick(Coutand et al., 2001), and those basins whose tops define the low-relief high elevation Plateau surface are internally drained (Isacks,1988; Vandervoort et al., 1995; Sobel et al., 2003). Shorteningestimates within the Puna are limited; however, those palinspaticreconstructions that exist suggest that less than 10–15% shortening(~37 km) has occurred within the Puna during the Cenozoic(Coutand et al., 2001).

The location and style of shortening within the interior of theAltiplano Plateau and along its margins differs from that of the Puna(e.g., Allmendinger, 1986; Lamb and Hoke, 1997; Lamb et al., 1997;Coutand et al., 2001;McQuarrie and DeCelles, 2001; Elger et al., 2005).The eastern margin of the Puna–Altiplano Plateau undergoes atransition around ~23°S in which thick-skinned, high angle reversefaults accommodate deformation to the south, while a west-vergingback-thrust belt accommodates deformation to the north (Sempereet al., 1990; McQuarrie and DeCelles, 2001; Elger et al., 2005).Within the Altiplano itself, ~65 km of shortening appears to beaccommodated between the Oligocene and mid-Miocene, withdeformation migrating to the orogen's margins after that time(Lamb et al., 1997; Lamb and Hoke, 1997; Elger et al., 2005). Defor-mation within the Plateau and along its margins has created up to a12-km-thick depocenter (Sempere et al., 1990; Kennan et al., 1995;Lamb and Hoke, 1997; Horton et al., 2002) that is currentlyinternally drained, and may have been so since Oligocene time(Horton et al., 2002; Sobel et al., 2003). Deformation within theAltiplano did not partition the depocenter into discrete, isolatedunits to the same degree as in the Puna (Allmendinger et al., 1997,


and references therein), creating the broad, largely uninterrupted,low relief Plateau surface (Fig. 4A).

Paleobotanical and stable-isotope data suggest that the broad,high-elevation portion of the central Andes, comprised mainly ofthe Puna–Altiplano Plateau abruptly rose sometimes after 10 Ma. Inparticular, paleobotanical evidence indicates that the Plateau hadobtained no more than a third of its modern elevation at 20 Ma, andhad risen to no more than half of its modern elevation by ~11 Ma(Gregory-Wodzicki, 2000). Likewise, fractionation of oxygen iso-topes recorded by paleosol carbonates, lacustrine carbonates,sandstone cements, and carbonates formed in paleo-marshes,indicate that the surface of the Bolivian Altiplano was uplifted~2.5–3.5 km between 10.3 and 6.8 Ma (Garzione et al., 2006). Theseresults are consistent with paleothermometric data from carbo-nates of the area, which have been interpreted to reflect ~3700±400 km of surface uplift between 10.3 and 6.7 Ma (Ghosh et al.,2006). Thus, at least within the Plateau, elevations appear to haveincreased during the late Cenozoic in what appears to be a rela-tively short-lived phase of rapid surface uplift (Garzione et al.,2006). This rapid uplift has been attributed to the convectiveremoval of eclogitized lower crust and lithosphere mantle beneaththe Altiplano–Puna Plateau (Molnar and Garzione, 2007; Hoke andGarzione, 2008).

3.4.2. Eastern CordilleraFission track data from the Eastern Cordillera record the earliest

onset of Cenozoic exhumation in the orogen during Late Eocene–EarlyOligocene in Bolivia (40 Ma) (Ege et al., 2007) and further south inArgentina (40–38 Ma) (Coutand et al., 2001, 2006). Interestingly, theend of exhumation in the Eastern Cordillera is diachronous in theNorth and the South. In Bolivia, it ceases in the Late Oligocene–EarlyMiocene, after which time exhumation resulting from tectonic rockuplift is transferred to the foreland area (Ege et al., 2007), while inArgentina it persists during the Late Miocene in the Angastaco area(Deeken et al., 2006; Coutand et al., 2006) until contractionaldeformation is transferred eastward to the Santa Barbara System inthe Pliocene (Kley and Monaldi, 2002).

The Eastern Cordillera is a basement-involved thrust system(Kley et al., 1996; Kley, 1996). In Argentina, Cenozoic deformationis taken up along moderate- to steeply-dipping reverse faults thataccommodate as much as 55–85 km of horizontal displacementrelative to stable South America (Grier et al., 1991; Baby et al., 1997;Coutand et al., 2001). Nested in this range, intramontane basinspreserve up to 6-km-thick Cenozoic sediments (Grier et al., 1991;Marrett and Strecker, 2000; Coutand et al., 2006) that have beenvariously cannibalized, reworked and deformed. In the southernEastern Cordillera, the current margins of the Cenozoic basinsclosely correspond with the paleo-borders of Cretaceous rift basins,emphasizing the impact of pre-existing structure in the modernarchitecture of the range (Grier et al., 1991; Deeken et al., 2006;Carrera et al., 2006; Carrera and Muñoz, 2008). Further north at~22°S, the thickening of the Paleozoic section triggers a change inthe style and amount of deformation accommodated within thisunit with the development of a back-thrust belt along the westernflank of the Eastern Cordillera and a deformation largely accom-modated along detachment horizons interlayered in the Paleozoicsection (e.g., Sempere et al., 1990; McQuarrie, 2002; Elger et al.,2005). Correspondingly, the Eastern Cordillera absorbs a largeramount of deformation, with total displacement accommodatedincreasing to ~100–140 km (Baby et al., 1997; McQuarrie andDeCelles, 2001; McQuarrie, 2002; Elger et al., 2005; McQuarrieet al., 2008a). Thus, while the Eastern Cordillera varies little in itsmeridianal location, the style and amount of deformation accommo-dated by these ranges changes abruptly along the strike. These changesspatially coincide with the distribution of pre-existing Late Paleozoicdetachmenthorizons (e.g., Allmendinger et al.,1983), againhighlighting

the deep influence of pre-existing paleogeographic features on themodern structure of the range.

3.4.3. Southernmost central AndesSouth of the termination of the Puna Plateau, the Aconcagua

fold-and-thrust belt, the Frontal Cordillera, and Precordilleradefine the eastern slope of the Andes (Fig. 4A). The emergence ofthe thin- and thick-skinned fold-and-thrust belts at this latitudecorresponds to the re-emergence of a thick pre-Tertiary sedimen-tary section in the modern foreland (Allmendinger et al., 1983;Ramos et al., 2002) that appears to have been deformed, and insome places continues to deform, from the mid-Miocene to present(e.g., Allmendinger et al., 1983; Ramos et al., 2002). At ~33°S,geologic relationships indicate that deformation propagated east-ward with time, and is broadly constrained to have occurredbetween 15 and 9 Ma in the Aconcagua fold-and-thrust belt, 9–6 Ma in the Frontal Cordillera, and 5–2 Ma in the Precordillera(Ramos et al., 2002). Further north at ~30.5°S, deformation withinthe Precordillera seems also to have migrated eastward with time,although the timing of deformation is slightly different than to thesouth (Allmendinger et al., 1990; Jordan et al., 1993). In particular,deformation within only the Precordillera appears to have begunaround ~20 Ma, and propagated eastward in a somewhat dis-continuous manner until the present (Jordan et al., 1993). In eithercase, deformation is broadly bracketed to have commenced in theearly- to mid-Miocene in this area, and currently continues todeform rocks at the boundary between the Precordillera andthe Sierras Pampeanas provinces (e.g., Zapata and Allmendinger,1996). Despite its relatively narrow width, total Cenozoic displace-ment accommodated within the Precordillera ranges from ~108to 135 km (relative to fixed South America) (Fig. 4C, D and E)(Allmendinger et al., 1990; Cristallini, 1996; Ramos et al., 2004,1996a,b).

3.4.4. Sierras PampeanasThe easternmost portions of the central Andes comprise, from

south to north, the Sierras Pampeanas, Santa Barbara system, andSubandean fold-and-thrust belt, respectively (Figs. 1 and 4A).Within the Sierras Pampeanas, deformation is accommodated byhigh-angle reverse faults whose locations often spatially corre-spond to terrain boundaries from the Paleozoic, or inverted normalfaults formed during the Cretaceous extension (Allmendinger etal., 1983; Schmidt et al., 1995; Ramos et al., 2002). Movementalong these thrust faults uplifts ranges of Late Proterozoic to EarlyPaleozoic metamorphic bedrock (Allmendinger et al., 1983; Jordanand Allmendinger, 1986). The utilization of pre-existing crustalweaknesses, whose along-strike extent is often limited, causesthese ranges to be discontinuous (e.g., Schmidt et al., 1995;Allmendinger and Gubbels, 1996). Intramontane depocenters asso-ciated with these ranges have been enduringly disconnected fromeach other and can preserve sediments up to 4000m in thickness(Sosic, 1972; de Urreiztieta et al., 1996; Fisher et al., 2002). AFTdata indicate that the frontal part of the Sierras Pampeanas (Sierrade Aconquija) was rapidly exhumed about 5 Ma ago (Sobel andStrecker, 2003) and warping of erosional surfaces in the NorthernSierras Pampeanas indicates that deformation remains active inthis area (e.g., Strecker et al., 1989). The Pampean Rangesaccommodate only about 10–20 km of horizontal displacementrelative to fixed South America (Fig. 4A, C and D) (Costa, 1992;Ramos et al., 2004).

3.4.5. Santa Barbara systemAt about 26°S, the northern boundary of the Sierras Pampeanas is

marked by an abrupt decrease in the wavelength of structures. There,the Santa Barbara system (Figs. 1 and 4A) accommodates deformationin the foreland largely by inversion of rift-related normal faults (Kley


and Monaldi, 2002). The thickness of Cretaceous strata varies sig-nificantly according to the spatial distribution of the rift depocenters(see Fig. 1 in Monaldi et al., 2008; Fig. 4F). The thickness of Cenozoicstrata exposed within the Santa Barbara ranges varies between 2.5 and5.5 km, the lower part being comprised of sediments reflecting thepost-rifting thermal subsidence stage and the upper part being forelandbasin strata related to the inversion of the rift (Kley and Monaldi,2002). While the utilization of high angle structures has createdsignificant topography, it has also allowed only modest amounts ofshortening to be accommodated: total displacement relative to stableSouth America is estimated to be between 25 and 30 km (Fig. 4B, C, andD) (Kley and Monaldi, 2002). The timing of contractional deformationis still controversial, but appears to have mostly taken place during thePliocene (Kley and Monaldi, 2002). The physiographic boundaries ofthe Santa Barbara system appear to be strongly controlled by the pre-Cenozoic paleogeography of the Andes. The southern extent of theSanta Barbara system lies at ~26°S, where the bedrock abruptlytransitions from the greenschist-grade Neoproterozoic to early Cam-brian Puncoviscana Formation to Meso-Neoproterozoic migmatites,gneisses, and granitoids (Allmendinger et al., 1983) and both thesouthern and northern boundaries of these ranges coincide with thetermination/bifurcation of the rift system (Kley and Monaldi, 2002).Specifically, the bifurcation of the Lomas de Olmedo depocenter to thenorth, causes the strike of old Cretaceous normal faults to becomeroughly orthogonal with respect to the Cenozoic shortening direction(Kley et al., 1999), rendering their reactivation mechanically difficult,and together with the tapering of Mesozoic deposits, apparently leadsto the termination of the Santa Barbara system.

3.4.6. Subandean rangesIn Bolivia, the eastern margin of the central Andes consists of

the Interandean Zone and the Subandean fold-and-thrust belt,which we consider together in this discussion for simplicity,although we acknowledge that important differences betweenthese two zones exist (e.g. Kley and Reinhardt, 1994). East of theBolivian Eastern Cordillera, a thin-skinned fold-and-thrust belt hasformed within the thick Paleozoic sedimentary section (Allmen-dinger et al., 1983). Deformation in this zone likely commencedduring the early Miocene and has propagated eastward with time(Kley, 1996). AFT data also document the migration of exhuma-tion from the EC around 30–20 Ma to the frontal part of thesubandes in the late Miocene–Pliocene (8–2 Ma) (Ege et al.,2007). This eastward propagation has contributed to the deposi-tion of late Cenozoic foredeep sediments of up to ~5 km thick infront of the advancing tip of the fold-and-thrust belt that was,and continues to be cannibalized as the belt widens (e.g., Kley,1996; Horton, 1997). Magnitude of relative displacement variesfrom north to south: in the north, it accommodates between 74and 135 km of displacement relative to stable South America(Roeder,1988; Babyet al.,1989,1997; Roeder and Chamberlain,1995), inits center ~163 km (McQuarrie, 2002), and in the south between 125and 150 km (Kley, 1996; Baby et al., 1997) (Fig. 4B, C, and D). Mecha-nically weak horizons in the Silurian, Devonian, and Carboniferouspart of the section accommodate the formation of many of thedecollement surfaces of the fold-and-thrust belt (e.g., Baby et al.,1990; Dunn et al., 1995). The geometry of these detachment horizonsalso changes from north to south—in the north, the basal decollementof the fold-and-thrust belt is ~4°, whereas it decreases to ~2° in thesouth—this southward shoaling corresponds to a widening of the fold-and-thrust belt (Roeder, 1988; Roeder and Chamberlain, 1995; Kley,1996; Horton, 1999). Likewise, some interpretations of GPS data(Norabuena et al., 1998) suggest that the distribution of active defor-mation within the fold-and-thrust belt changes along its length, without-of-sequence thrusting characterizing the northern part ofthe wedge (Horton, 1999), and deformation localized at its easterntoe in the south. Finally, the fold-and-thrust belt disappears to the south

due to the tapering of the Silurian and Devonian decollements in thePaleozoic basin (Belotti et al., 1995). This leaves a gap of ~60 km bet-ween the southern termination of the Subandean belt and theemergence of the Santa Barbara system to the south (Kley andMonaldi,2002). Thus, the Subandean belt is closely associated with the presenceof mechanically weak detachment horizons in pre-existing basindeposits (e.g., Allmendinger et al., 1983).

3.5. Current and Cenozoic climate

Today, the Andes constitute a major topographic barrier toatmospheric circulation in South America and separate the humidlowlands in the east from the arid Pacific margin in the west(see Fig. 1b in Bookhagen and Strecker, 2008). The semiarid tohyperarid climates observed on the western slope of the orogenand across the high Plateau (Fig. 1B; Strecker et al., 2007, andreferences therein), result from their proximity to upwelling ofcold water along the Pacific coast due to the cold Humboldtcurrent, which may have existed in some form since the earlyCenozoic (Zachos et al., 2001), and the fact that these regionsare located within a high-pressure sector with atmospheric sub-sidence (Houston and Hartley, 2003). Precipitation in the CentralAndes is largely derived from the westward penetration of Atlanticmoisture from the Chaco Plain and Argentine foreland (Rohmeder,1943; Halloy, 1982; Bianchi and Yañez, 1992; Garreaud et al., 2003)via the Andean low-level jet (Nogués-Paegele and Mo, 1997). Asmoisture encounters the topography of the eastern Andes,precipitation is orographically enhanced (e.g., Bianchi and Yañez,1992; Sobel et al., 2003; Strecker et al., 2007; Bookhagen andStrecker, 2008). Both the relief structure of the eastern flankof the central Andes and the source of incoming precipitation causelarge latitudinal changes in the amount of rainfall reaching theeastern slope and interior of the orogen (Fig. 1; Strecker et al.,2007; Bookhagen and Strecker, 2008). In northern Bolivia at ~15–17°S, orographic rainfall peaks above 3.5 m/yr at a mean elevationof 1.3 km once topographic relief of about 1 km is achieved(Bookhagen and Strecker, 2008). South of 17°S, orographicprecipitation decreases as the mountain belt widens (Montgomeryet al., 2001) but is still on the order of 1.5 m/yr at 26°S (Bookhagenand Strecker, 2008). As the long, linear ranges of the Subandeanfold-and-thrust belt give way to the largely discontinuous, steep-flanked and higher elevated ranges of the Santa Barbara system andSierras Pampeanas, precipitation is strongly focused on the wind-ward sides of these ranges (Strecker et al., 2007; Hilley andStrecker, 2005). Hence they shelter the eastern slope of the innerorogen from precipitation derived from the Atlantic (Fig. 1B). As aconsequence, most areas of the Sierras Pampeanas, and thePrincipal, Frontal, and Pre-cordilleras receive little moisture fromthe east (Sobel and Strecker, 2003).

Since ~18 Ma, the latitudinal location of the central Andes hasremained relatively stable (Scotese et al., 1988). In addition, theatmospheric and oceanic circulation that exists today appears tohave been established since at least the early Cenozoic (Parrishet al., 1982), and perhaps the Mesozoic (Hartley et al., 1992). In thebroadest sense, the persistent latitudinal position and stableatmospheric circulation has created a generally arid climate withinthe central Andes, with the western portions of the orogencharacterized by the most intense aridity, while the eastern flanksgenerally receive far more moisture (Fig. 1; Strecker et al., 2007).While this overall pattern has persisted since the Miocene,paleoclimate proxies record significant climate changes through-out this period (Strecker et al., 2007, and references therein).Furthermore, along the western Andean margin, late Miocene to earlyPliocene paleoclimate indicators are somewhat contradictory andvary latitudinally (Sáez et al., 1999; Gaupp et al., 1999; May et al.,2005); nonetheless, generally arid to semi-arid conditions with


periods of greatermoisture availability typify thewesternmargin of theorogen since the mid-Miocene, and perhaps as early as theEocene (Blanco et al., 2003). The termination of supergenealteration within the Atacama Desert and volcanic arc between 14and 8 Ma (Alpers and Brimhall, 1988; Sillitoe et al., 1991) suggest thatarid conditions commenced in these areas by mid-Miocene time.In addition, the deposition of thick halite and gypsum-bearingsedimentary units between 24 and 15 Ma within the Puna basins,indicate that by this time, the climate was arid (Alonso et al., 1991;Vandervoort et al., 1995).

In contrast, sediments found in areas east of the Puna andAltiplano show evidence for a more complex and variable climatehistory that appears strongly affected by the uplift of individualranges and the resulting downwind aridification (Strecker et al.,2007). For example, intramontane basins in the Eastern Cordil-lera of northwestern Argentina preserve sediments that record ashift from semi-arid/arid conditions to humid environment inthe late Miocene (Marshall and Patterson, 1981; Pascual et al.,1985; Pascual and Ortiz Jaureguizar, 1990; Nasif et al., 1997;Kleinert and Strecker, 2001; Starck and Anzótegui, 2001; Coutandet al., 2006; Strecker et al., 2007; Anzótegui, 2007; Hilley andStrecker, 2005). Further to the north in the Subandean ranges ofBolivia, climate also shifted towards more humid conditionsin the late Miocene (Uba et al., 2005, 2006), a time during whichmegafan were developed in this area (Horton and DeCelles,2001). In most of these marginal basins, the late Miocene changeto more humid conditions was abruptly terminated as the upliftof upwind ranges prevented moisture from penetrating intothese areas (Sobel et al., 2003; Sobel and Strecker, 2003;Coutand et al., 2006; Strecker et al., 1989, 2007, Hilley andStrecker, 2005). The timing of the onset of ensuing aridificationvaries from basin to basin, depending on the local tectonichistory that uplifted the upwind orographic barriers (e.g.,Strecker et al., 2007). This tectonic control on local precipitationpatterns is observed in the modern rainfall distribution on theeastern slope of the Andes (Figs. 1 and 4). In summary, both pastand current climate generally show an arid or-hyperarid westernslope, an arid interior, and an eastern slope that varies fromhumid to arid, depending on local topography and latitudinalposition.

4. Relationships between paleogeography, climate, and tectonicsin the central Andes

The location of paleogeographic provinces, total shortening,downgoing slab geometry, mountain belt width, and climatecorrespond to varying degrees within the central Andes (Fig. 4).In summarizing the above discussion, between ~15° and 22.5°–23°S total shortening is large (40%–70%) and is concentratedwithin the Eastern Cordillera and the Subandean ranges (Fig. 4 C,D, and E). The subducted Nazca plate dips steeply eastwardbeneath continental South America (Fig. 4G) and the location ofthe foreland fold-and-thrust belt coincides with pre-existing thickPaleozoic deposits containing multiple stratigraphic horizons thatserve as decollements (e.g., Allmendinger et al., 1983; Kley et al.,1999). However, from 17.5°S, the abrupt increase of the mountainbelt's width (Fig. 4A) spatially coincides with a decrease in meanannual precipitation (WMO, 1975; Montgomery et al., 2001)(Fig. 4B), the shallowing of the basal decollement angle of theorogenic wedge (Roeder, 1988; Roeder and Chamberlain, 1995;Kley, 1996; Horton, 1999) and the distribution of active deforma-tion, as documented by GPS data (Norabuena et al., 1998; Horton,1999). Between ~22.5°S and 26.5°S, the central Andes accommo-dates less shortening (30 to 40%) (Figs. 4C, D and E) and thiscorresponds spatially to the disappearance of the subandeanranges and Paleozoic foreland deposits with interlayered decolle-

ments. In the foreland, shortening becomes accommodated byinversion of Cretaceous normal faults in the thick-skinned SantaBarbara ranges (Jordan et al., 1983; Kley and Monaldi, 2002).Commensurate with these changes in total displacement, totalshortening, and structural style, the mountain-belt width alsodecreases (Fig. 4H). At ~26°S, the Santa Barbara system becomestransitionally replaced by the high-angle reverse faulted SierrasPampeanas to the south. Both these ranges focus Atlantic-derivedprecipitation on their eastern margin, resulting in large precipita-tion gradients between the humid flanks and arid interior of theorogen (Fig. 1B; Hoffmann, 1975; Sobel et al., 2003; Sobel andStrecker, 2003; Hilley and Strecker, 2005).

Finally, between 28°S and 34°S, pre-Cenozoic deposits areabsent in the foreland belt, and upper crustal shortening is takenup by reactivated Paleozoic structures across the Sierras Pampeanas(Allmendinger et al., 1983; Schmidt et al., 1995; Ramos, 2000;Ramos et al., 2002) that accommodate only 10–20 km of horizontaldisplacement (Fig. 4D). The distributed nature of the deformationwidens the mountain belt and consequently decreases the totalshortening (Fig. 4E). The subduction angle is low and thereforeactive volcanism is absent (Fig. 4G). Previous geologic studies havesuggested that tractions and/or thermal heating at the base of theSouth American lithosphere may have caused the eastwardmigration of deformation in this area (Jordan et al., 1983; Ramoset al., 2002) (Fig. 4H). In the arid westernmost portion of this zone(Figs. 1 and 4), the re-emergence of a thick pre-Cenozoicsedimentary section corresponds with the development of thePrincipal, Frontal, and the presently active Precordillera fold-and-thrust belt (Fig. 4A, D, and F), where a larger amount of dis-placement is accommodated than in the Sierras Pampeanas (Fig. 4Aand D).

In the Central Andes, the structural style (thin-skinned versusthick-skinned fold-and-thrust belt, basement-cored uplifts) of theexternal portion of the orogenic system mirrors the paleogeo-graphic zones that existed prior to the Andean orogenic event(Allmendinger et al., 1983). The fact that the amount of horizontaldisplacement absorbed by the various structural provinces ismarkedly different suggests that latitudinal changes in deforma-tion may be dominantly controlled by factors set in motion longbefore the construction of the modern mountain belt (Allmendin-ger et al., 1983). Thus, the paleogeographic heritage likely triggersmarked dissimilarities in the topography of the eastern Andeanmargin. In contrast, the topography of the high Plateau itself isremarkably symmetric (Fig. 3; Gephart, 1994). However, theamount and style of deformation accommodated along the easternboarder of the high plateau within the Eastern Cordillera appearsto show marked asymmetry along its length. In Bolivia, the Plateauboundary is seated within a thick section of Paleozoic metasedi-ments, and a large amount of shortening is accommodated withinthe Eastern Cordillera and its backthrust belt (McQuarrie andDeCelles, 2001; Elger et al., 2005). In contrast, along the ArgentinePuna margin, the Paleozoic deposits generally tend to becomethinner, high-angle reactivated structures accommodate deforma-tion and total shortening decreases (Baby et al., 1997; Kley et al.,2005). Thus, while the location of the eastern margin of thePlateau appears to reflect the geometry of the downgoing Nazcaplate, the nature of deformation and the amount of displacementaccommodated along the margin appears controlled by the paleo-geography of those areas.

The above observations allow us to draw some inferences aboutthe roles that slab geometry, paleogeography, and climate may playin shaping the central Andes. First, we observe that the extent ofthe Plateau coincides with major changes in the downgoing Nazcaplate geometry, suggesting a causal link between the two (Gephart,1994). As an alternative hypothesis, others have suggested that thepaleogeography of the South American craton may have influenced


the slab geometry itself: in the south where pre-existing high-angle structures are pervasive within the crust, the slab wasflattened as it was pushed underneath a continent that could notyield to the degree that it could in the north, where the Paleozoicbasin allowed large amounts of shortening to be accommodatedwithin the foreland (Allmendinger et al., 1983; Allmendinger andGubbels, 1996). Nonetheless, a combination of the paleogeographyand subduction zone geometry appears responsible for the large-scale architecture of the interior part of the orogen, rather thanclimate playing a primary role as some have previously suggested(e.g., Montgomery et al., 2001; Sobel et al., 2003). In addition, theamount and nature of deformation accommodated within theinterior and external parts of the orogen varies from north to south,leaving a strong imprint on the topography of the mountain range:(1) Total W–E shortening decreases stepwise to the south;(2) These variations are focused on boundaries of tectonicprovinces characterized by different styles of deformation; and(3) The total width of the mountain belt changes within each ofthese provinces. Changes in the precipitation distribution thatresult from both regional air-mass circulations and local orographiceffects broadly coincide with these tectonic and topographicboundaries. However, the spatial coincidence of these boundarieswith factors established prior to the Cenozoic makes it unlikelythat the current climate plays a dominant role in determining thelarge-scale architecture of the central Andes.

5. What role might climate play in the tectonics and topography ofthe central Andes?

If the paleogeography and plate interface geometry have left such astrong imprint on the morphology of the modern central Andes, thenwhat role, if any, might climate have played in controlling differenttectonic features? The coincidence of pre-Andean paleogeographywith the current Andean tectonic, climatic, and topographic attributesargues that most of the orogen-scale features of the central Andeshave been shaped in large part by this paleogeography. However, it ispossible that at a more local scale, climate, erosion, and tectonics maybe linked. To investigate this option, we further analyze three distinctmorphotectonic provinces in the central Andes: (1) the internallydrained Altiplano–Puna Plateau, (2) the foreland fold-and-thrust beltsof Bolivia and Argentina, and (3) the Sierras Pampeanas basement-cored uplift province (Fig. 4; labeled region 1, 2, and 3 respectively).The Altiplano–Puna plateau typically receives b400 mm of annualprecipitation (WMO,1975; Sobel et al., 2003; Bookhagen and Strecker,2008), and stratigraphic data indicate that the current western andeastern Altiplano margins were uplifted in the Eocene–Oligocene, andOligocene–Miocene, respectively, leading to the formation of internaldrainage in the area by ~25 Ma (Horton et al., 2002). Likewise, in thePuna, uplift of its present eastern margin and disparate basementuplifts within the plateau region may have led to the formation ofinternal drainage between Oligocene and Miocene time (e.g., Coutandet al., 2001, Vandervoort et al., 1995; Coutand et al., 2006). While theextent and symmetry of the Plateau appears to reflect the downgoingNazca plate geometry, erosion may help to reinforce the nature of thetopography and tectonics of this area. In particular, the establishmentof internal drainage decouples the local base-level of streams drainingthe hinterland from their regional base-level, allowing large-scaleaggradation to occur (Sobel et al., 2003; Hilley and Strecker, 2005).Integrated field studies and theoretical models indicate that thedisconnection and reintegration of individual basins base-levels mayoccur over timescales of ~105years (Allen, 2008; Hilley and Strecker,2005). The fact that regional disconnection of the foreland and Plateaubase-levels have persisted for much longer (e.g., Vandervoort et al.,1995) indicates that conditions that favor this disconnection havepersisted over these longer timescales. This has two important effectson erosional–tectonic coupling: (1) Regional aggradation reduces

relief in landscapes as channels aggrade and peaks degrade, and slopereduction decreases erosion in the orogen's interior; (2) Internaldrainage and uplifting upwind orographic barriers prevent massevacuation from this region, increase lithostatic loads, and reduce theisostatically driven rock uplift that would otherwise result from massremoval. Maintenance of lithostatic stresses by restriction of erosionalmass export may conspire with geodynamic processes to causedeformation tomigrate to the external parts of the orogen (Sobel et al.,2003). Together, geodynamic processes and an arid climate decouplesthe Plateau's base-level from the foreland, allowing a broad, low-relief, tectonically quiescent Plateau to persist over geologic time in amanner similar to the hypothetical case when erosion is absent (e.g.,Royden, 1996).

The foreland fold-and-thrust belts of Bolivia and Argentina (region2 in Fig. 4A) are consistently associated with thick pre-existingsedimentary units that contain mechanically weak layers that servedas decollement horizons. These regions are externally drained; thus,erosional mass efflux increases withmountain-belt width and/or withsteeper mean orogenic slopes parallel to material transport direction(e.g., Dahlen and Suppe, 1988; Willett et al., 1993). In contrast toplateau environments, field (Hilley et al., 2004), experimental (e.g.,Persson et al., 2004; Hoth et al., 2006) and theoretical (e.g., Dahlen,1984) studies suggest that the voracity of erosional processes likelyplay a role in controlling the long-term deformation and topography.Piggy-back basins in fold-and-thrust belts may act as significantstorage areas for sediment over 105year timescales (Allen, 2008;Hilley and Strecker, 2005). However, the resistant bedrock exposedwithin the ranges of the fold-and-thrust belt likely limit their rateof erosion (Whipple et al., 1999). Slope adjustments within the fold-and-thrust belt generally take ~106–107years to balance erosionalmass efflux with the flux of tectonically accreted material (e.g., Stockand Montgomery, 1999; Hilley et al., 2004). Both the Bolivian andArgentine fold-and-thrust belts have been active since the Miocene,suggesting that erosion and deformation in these areas have at leastpartially adjusted to one another. This adjustment may be reflected bychanges in the geometry of the fold-and-thrust belt in space and/ortime. For example, in northern Bolivia, a southward decrease in meanannual precipitation coincides with a widening of the fold-and-thrust belt (Fig. 4A and B) and perhaps a transition from diffuse tolocalized deformation along the topographic front of the Bolivianfold-and-thrust belt (Horton, 1999). Theoretical considerationssuggest that weaker erosion, presumably caused by decreasedprecipitation rates, results in widening of the orogen to evacuatetectonically accreted foreland sediments (e.g., Whipple and Meade,2004). Also, studies of the Argentine Principal, Frontal, and Pre-cordilleras (Fig. 4) indicate that changes in erosion rates, promptedby the exposure of different rock types, may result in changingfold-and-thrust belt geometry and topography over time (Hilley etal., 2004). Thus, theory predicts that externally drained areaspresenting the proper geologic conditions for the formation of fold-and-thrust belts, may show a strong imprint of climatically-moderated erosional processes on their morphology (Montgomeryet al., 2001). In the central Andes, segments of the mountain beltexperiencing less erosion, either due to lower precipitation rates(Montgomery et al., 2001) or to the exposure of more resistantlithologies (Hilley et al., 2004), tend to be wider than more erosiveareas. This is consistent with the view that the geometry of fold-and-thrust belts depends on the efficiency of erosion (e.g., Dahlenand Suppe, 1988). However, within the Bolivian fold-and-thrustbelt, recent area balancing of structural sections found littledifference in shortening within this area, which may suggest thaterosion may exert less of a control on orogenic structure here thanpreviously thought (McQuarrie et al., 2008b).

In contrast to fold-and-thrust belts, deformation along reacti-vated crustal-scale heterogeneities such as in the Sierras Pampeanas(region 3 in Fig. 4A) involve steep, pre-existing crustal-scale


weaknesses that may undergo frictional failure without necessitat-ing failure within the rest of the mountain range. An analysis of thestresses expected in this situation suggests that the development oferosionally moderated surface slopes should be decoupled from thereactivation of these structures (Hilley et al., 2005), except whenmean slopes across the range exceed a threshold value. In this case,the large lithostatic loads produced by these slopes opposehorizontal stresses and favor failure of frictionally stronger struc-tures without such loads. Because erosional processes moderatesurface slopes, vigorous erosion may prevent the build-up of steepslopes, and hence the migration of deformation to other structuresin low-lying areas. As in fold-and-thrust belts, bedrock erosion islikely the process that limits the timescale over which erosionresponds to tectonic uplift. However, the time it takes to buildtopography significant enough to cross this threshold slope dependson rates of uplift produced by fault motion, the erosional efficiency,the value of the threshold slope, and the geometry of the underlyingfault. This time-scale varies depending on local circumstances, butlikely occurs over ranges of 104–106years. These theoreticalconsiderations shed light on the development of two mountainranges within the northern Sierras Pampeanas that have undergonesimilar geologic histories but have vastly different kinematics. Thoseranges in wet areas tend to have deformation focused on a limitednumber of structures with large amounts of displacement, whilethose in drier sectors display more widely distributed deformationalong several faults with relatively small individual displacements(Sobel and Strecker, 2003). Unfortunately, the location of zones ofpre-existing weakness, their frictional properties, and their localloading conditions cannot be easily measured, and so a strict test ofthe idea that erosion may control a threshold in the development ofthese ranges remains elusive. Nonetheless, the details of thedevelopment of individual ranges in the Sierras Pampeanas areconsistent with qualitative theoretical expectations, and so it ispossible that the details of the development of individual ranges canbe influenced by climate moderated erosional processes.

While strict tests linking deformation to erosion remain elusive,detailed geologic observations are consistent with theory that links thetopographic and tectonic development of different components of thisorogen. In some cases, erosion may be strongly coupled to tectonicdeformation (e.g., fold-and-thrust belts), while it is less influential inothers (e.g., internally-drained plateaus). However, even in regionssuch as the internally drained plateau, the voracity of erosion may haveleft an important imprint on the landscape. For example, theestablishment of internal drainage within the Altiplano–Puna plateauin the Oligocene–Miocene restricted mass export from the hinterlandand may have allowed the low-relief plateau surface to form prior to itsuplift in late Miocene time (e.g., Garzione et al., 2006, and referencestherein). Had external drainage been maintained, erosional removal ofmaterial from the current plateau realm might have led toa topographically, and perhaps tectonically different orogen than thatobserved today.

6. Concluding remarks

Both the geodynamic and pre-Cenozoic geologic history of the centralAndes play an important, if not dominant role in the orogen-scalestructure of this mountain belt. However, at the scale of individualmountain ranges or morphotectonic provinces, it is also likely thaterosional processes playan important role in shaping the topographyandinfluence the kinematics of deformation. The construction of topographyin turn, modifies the local climate conditions, reinforcing or damping thestrength of erosional processes in the different provinces. The nature ofthe coupling between erosion and tectonics in the central Andes variesbetween each of these different morphotectonic zones. In some areas,erosionmay be strongly coupled to tectonic deformation, while in othersthese factors may be largely independent. In our view, the pre-orogenic

geologic history of the central Andes sets the stage for the nature andstrength of coupling between erosional processes, tectonic deformation,and mountain belt structure. This situation reflects the complexfeedbacks that we might expect to arise between the pre-existingrheological state of the crust, the geodynamic boundary conditions, andclimate-moderated erosional processes (Fig. 2).

Several important lessons may be gathered from interactionsobserved within the central Andes. Most importantly, the nature ofcoupling between erosion, climate, pre-orogenic geologic history, plate-boundary forces, andmantle processes is complex, and the impact of theinitial state of all of these variablesmaynoteasily be erasedover the life ofthe orogen. While our previous work has tried to view the orogenicstructure in terms of simple interactions between these factors (Sobel etal., 2003, Hilley et al., 2004; Hilley and Strecker, 2004; Hilley et al., 2005)and the work of others have attempted to understand overall orogenicdevelopment in terms of simple erosional and geodynamic feedbacks(e.g., Willett et al., 1993; Whipple and Meade, 2004), these approachesmaynotbealtogetherappropriate foranorogen suchas thecentralAndesthat has been subjected to a long and rich tectonic history. They may failto reveal the complex interactions between tectonic and erosionalprocesses that characterize these types of orogens. However, becausefactors affecting the Earth's mantle, crust, and surface at the onset of andduring mountain building profoundly influence the development of theorogen throughout its life, we argue that the path to understanding theseorogens requires an integrative approach that combines geodynamic,geologic, and geomorphologic observations and models.

Acknowledgments

This work has been supported by the France-Stanford Center forInterdisciplinary Studies and by the Fulbright Scholar Program.

Appendix A. Supplementary data

Supplementary data associated with this article can be found, inthe online version, at 10.1016/j.tecto.2009.06.017.

References

Aceñolaza, F.G., 1973. Sobre la presencia de Oldhamia sp. en la Formación Puncoviscanade Cuesta Muñano, Provincia de Salta, República Argentina. Revista de la AsociaciónGeológica Argentina, 28: 56.

Allen, P.A., 2008. Time scales of tectonic landscapes and their sediment routing systems.In: Gallagher, K., Jones, S.J., Wainwright, J. (Eds.), Landscape Evolution: Denudation,Climate and Tectonics Over Different Time and Space Scales. Special Publications,London. Geological Society, London, pp. 7–28.

Allmendinger, R.W., 1986. Tectonic development, southeastern boarder of the PunaPlateau, northwestern Argentine Andes. Geological Society of America Bulletin 97,1070–1082.

Allmendinger, R.W., Gubbels, T.L., 1996. Pure and simple shear plateau uplift, Altiplano–Puna, Argentina and Bolivia. Tectonophysics 259, 1–13.

Allmendinger, R.W., Ramos, V.A., Jordan, T.E., Palma, M., Isacks, B.L., 1983. Paleogeographyand Andean structural geometry, northwest Argentina. Tectonics 2 (1), 1–16.

Allmendinger, R.W., et al., 1990. Foreland shortening and crustal balancing in the Andesat 30 S latitude. Tectonics 9 (4), 789–809.

Allmendinger, R.W., Jordan, T.E., Kay, S.M., Isacks, B.L.,1997. The evolution of the Altiplano–Puna Plateau of the central Andes. Annual Reviews of Earth and Planetary Sciences 25,139–174.

Alonso, R., Jordan, T.E., Tuabbutt, K.T., Vandervoort, D.S., 1991. Giant evaporite belots ofthe Neogene central Andes. Geology 19, 401–404.

Alpers, C.N., Brimhall, G.H., 1988. Middle Miocene climatic change in the AtacamaDesert, northern Chile; evidence from supergene mineralization at La Escondida.Geological Society of America Bulletin 100, 1640–1656.

Anzótegui, L.M., 2007. Paleofloras del Mioceno en los Valles Calchaquíes, Noroeste deArgentina. Universidad Nacional Nordeste, Corrientes, Argentina.

Baby, P., et al., 1989. Structure de la zone Subandine de Bolivie: influence de la géométriedes séries sédimentaires antéorogéniques sur la propagation des chevauchements.Tectonique 309, 1717–1722.

Baby, P., et al., 1990. A late Oligocene–Miocene intermountain foreland basin in thesouthern Bolivian Altiplano. Comptes rendus de l'AcadeÂmie des sciences 11 (311),341–347.

Baby, P., Specht, M., Oller, J., Montemurro, D., Colletta, B., Letouzey, J., 1994. TheBoomerang-Chapare transfer zone (Recent Oil Discovery Trend in Bolivia):Structural interpretation and experimental approach. In: Roure, F., Ellouz, N.,

http://dx.doi.org/10.1016/j.tecto.2009.06.017


Shein, V.S., Skvortsov, I. (Eds.), Geodynamic Evolution of Sedimentary Basins,International Symposium, Moscow, 203–218.

Baby, P., Rochat, P., Mascle, G., Hérail, G., 1997. Neogene shortening contribution tocrustal thickening in the back-arc of the Central Andes. Geology 25, 883–886.

Bahlburg, H., Breitkreuz, C., 1991. Paleozoic evolution of active margin basins in thesouthern Central Andes (northwestern Argentina and northern Chile). Journal ofSouth American Earth Sciences 4, 171–188.

Bahlburg, H., Furlong, K., 1996. Lithospheric modeling of the Ordovician foreland basinin the Puna of northwestern Argentina: on the influence of arc loading on forelandbasin formation. Tectonophysics 259, 245–258.

Baldis, B.A.J., Gorroño, A., Ploszkiewicz, J.V., Sarudiansky, R.M., 1976. Geotectónica de laCordillera Oriental, Sierras Subandinas y comarcas adyacientes. Geol. Prov. BuenosAires, Relat. Congr. Geol. Argent. 6th., 1, pp. 4–22.

Barazangi, M., Isacks, B.L., 1976. Spatial distribution of earthquakes and subduction ofthe Nazca Plate beneath South America. Geology 4, 686–692.

Barazangi, M., Isacks, B.L., 1979. Subduction of the Nazca plate beneath Peru: evidencefrom spatial distribution of earthquakes. Geophysical Journal International 57 (3),537–555.

Barry, R.G., 1981. Mountain Weather and Climate. Methuen, New York.Beaumont, C., Fullsack, P., Hamilton, J.C., 1992. Erosional control of active compres-

sional orogens. In: McClay, K.R. (Ed.), Thrust Tectonics. Chapman and Hall, NewYork, pp. 19–31.

Beaumont, C., Jamieson, R.A., Nguyen,M.H., S.,M., 2004. Crustal channelflows: 1. Numericalmodels with applications to the tectonics of the Himalayan–Tibetan orogen. Journal ofGeophysical Research, B, Solid Earth and Planets 109, B06406.

Belotti, H.J., Saccavino, L.L., Schachner, G.A., 1995. Structural styles and petroleumoccurence in the sub-Anean fold and thrust belt of northern Argentina. In: Tankard,A.J., Suárez, R.S.,Welsink, H.J. (Eds.), Petroleum Basins of South America, pp. 545–555.

Bevis, M., Isacks, B.L., 1984. Hypocentral trend surface analysis: probing the geometry ofBenioff Zones. Journal of Geophysical Research 89, 6153–6170.

LasPrecipitacionesenelNoroesteArgentino. In:Bianchi,A.R., Yañez, C.E. (Eds.),AgropecuariaAgropecu. Estacíon Exp, Salta, Argent. 393 pp.

Blanco, N., Tomlinson, A.J.,Mpodozis, C., Pérez, d.A., C. andMatthews, C.Y., 2003. FormaciónCalama, Eoceno, II Region de Antofagasta (Chile): Estratigrafía e ImplicanciasTectónícas, Congresso Geologia Chileno, 10, Universidad Concepción.

Bookhagen, B., Burbank, D.W., 2006. Topography, relief, and TRMM-derived rainfallvariations along the Himalaya. Geophysical Research Letters 33, L08405.

Bookhagen, B., Strecker, M.R., 2008. Orographic barriers, high-resolution TRMM rainfall,and relief variations along the easternAndes.Geophysical ResearchLetters 35, L06403.

Bookhagen, B., Thiede, R.C., Strecker, M.R., 2005a. Abnormal monsoon years and theircontrol on erosion and sediment flux in the high, arid northwest Himalaya. Earthand Planetary Science Letters 231, 131–146.

Bookhagen, B., Thiede, R.C., Strecker, M.R., 2005b. Late Quaternary intensified monsoonphases control landscapeevolution in thenorthwestHimalaya.Geology33 (2),149–152.

Brace, W.F., Kohlstedt, D.L., 1980. Limits on lithospheric stress imposed by laboratoryexperiments. Journal of Geophysical Research 85, 6,248–6,252.

Bürgmann, R., Dresen, G., 2008. Rheology of the lower curst and upper mantle: evidencefrom rock mechanics, geodesy, and field observations. Annual Review of Earth andPlanetary Sciences 36, 531–567.

Byerlee, J.D., 1968. Brittle–ductile transition in rocks. Journal of Geophysical Research73, 4,741–4,750.

Byerlee, J.D., 1978. Friction of rocks. Pure and Applied Geophysics 116, 615–626.Bystricky, M., Mackwell, S., 2001. Creep of dry clinopyroxenite aggregates. Journal of

Geophysical Research 106, 13,443–13,454.Cahill, T., Isacks, B.L., 1992. Seismicity and shape of the subducted Nazca plate. Journal of

Geophysical Research, B, Solid Earth and Planets 97 (B12), 17,503–17,529.Carlier, G., et al., 2005. Potassic–ultrapotassic mafic rocks delineate two lithospheric

mantle blocks beneath the southern Peruvian Altiplano. Geology 33 (7), 601–604.Carrapa, B., DeCelles, P.G., 2008. Eocene exhumation and basin development in the Puna

of northwestern Argentina. Tectonics 27, TC1015.Carrapa, B., Adelmann, D., Hilley, G.E., Mortimer, E., Sobel, E.R., Strecker, M.R., 2005,

Oligocene range uplift and development of plateau morphology in the southerncentral Andes, Tectonics, Vol. 24, TC4011, doi:10.1029/2004TC001762.

Carrera, N., Muñoz, J.A., 2008. Thrusting evolution in the southern Cordillera Oriental(northern Argentine Andes): constraints from growth strata. Tectonophysics 459,107–122.

Carrera, N., Muñoz, J.A., Sabat, F., Mon, R., Roca, E., 2006. The role of inversion tectonicsin the structure of the Cordillera Oriental (NW Argentinean Andes). Journal ofStructural Geology 28, 1921–1932.

Coira, B., Davidson, J., Mpodozis, C., Ramos, V.A., 1982. Tectonic and magmatic evolutionof the Andes of northern Argentina and Chile, a symposium on the magmaticevolution of the Andes. Earth Science Reviews 18, 303–332.

Conrad, C.P., Molnar, P., 1997. The growth of Rayleigh–Taylor-type instabilities in thelithosphere for various rheological and density structures. Geophysical JournalInternational 129, 95–112.

Costa, C.H., 1992. Neotectónica del sur de la Sierra de San Luis. PhD, Unpublished Thesis,Universidad Nacional de San Luis, 390 pp.

Coutand, I., 1999. Tectonique cénozoïque du Haut Plateau de la Puna, Nord Ouestargentin, Andes Centrales. PhD thesis, Mémoires de Géosciences Rennes 92 381pp.(ISBN 2-905532-91-2).

Coutand, I., et al., 2001. Style and history of Andean deformation, Puna plateau,northwestern Argentina. Tectonics 20 (2), 210–234.

Coutand, I., et al., 2006. Propagation of orographic barriers along an active range front:insights from sandstone petrography and detrital apatite fission-track thermo-chronology in the intramontane Angastaco basin, NWArgentina. Basin Research 18,1–26.

Cristallini, E., 1996. La faja plegada y corrida de La Ramada, Geologìa de la región delAconcagua, provincias de San Juany Mendoza. Dirección Nacional del ServicioGeológico Anales, pp. 349–385.

Dahlen, F.A., 1984. Non-cohesive critical Coulomb wedges: an exact solution. Journal ofGeophysical Research 89, 10,125–10,133.

Dahlen, F.A., Suppe, J., 1988. Mechanics, growth, and erosion of mountain belts,Geological Society of America Special Paper 218. Geological Society of America,Denver, CO, pp. 161–178.

Davis, D., Suppe, J., Dahlen, F.A., 1983. Mechanics of fold-and-thrust belts andaccretionary wedges. Journal of Geophysical Research 88, 1153–1172.

de Urreiztieta, M., Gapais, D., Le Corre, C., Cobbold, P.R., Rossello, E.A., 1996. Cenozoicdextral transpression and basin development at the southern edge of the PunaPlateau, northwestern Argentina. Tectonophysics 254, 17–39.

deWit, M.J., Ransome, G.D., 1992. Regional inversion tectonics along the southern marginof Gondwana. In: De Wit, M.J., Ransome, G.D. (Eds.), Inversion Tectonics of the CapeFold Belt, Karoo and Cretaceous Basins of Southern Africa. Balkema, pp. 15–21.

Deeken, A., et al., 2006. Development of the southern Eastern Cordillera, NWArgentina,constrained by apatite fission track thermochronology: from early Cretaceousextension to middle Miocene shortening. Tectonics 25 TC6003-TC6003.

Dimanov, A., Dresen, G., 2005. Rheology of synthetic anorthite–diopside aggregates:implications for ductile shear zones. Journal of Geophysical Research 110, B07203.

Dunn, J.F., Hartshorn, K.G., Hartshorn, P.W., 1995. Structural styles and hydrocarbonpotential of the Subandean thrust belt of southern Bolivia. In: Tankard, A.J., S. R.,Welsink, H.J. (Eds.), Petroleum Basins of South America, pp. 523–543.

Edman, J.D., Surdam, R.C., 1984. Influence of ocerthrusting on maturation ofhydrocarbons in Phosphoria Formation, Wyoming–Idaho–Utah overthrust belt.AAPG Bulletin 68, 1803–1817.

Ege, H., Sobel, E.R., Scheuber, E., Jacobshagen, V., 2007. Exhumation history of thesouthern Altiplano plateau (southern Bolivia) constrained by apatite fission trackthermochronology. Tectonics 26, TC1004.

Ehlers, T.A., 2005. Crustal thermal processes and thermochronometer interpretation.Reviews in Mineralogy and Geochemistry 58, 315–330.

Elger, K., Oncken, O., Glodny, J., 2005. Plateau-style accumulation of deformation:southern Altiplano. Tectonics 24, TC4020.

Fernandez-Seveso, F., Tankard, A.J., 1995. Tectonics and stratigrpahy of the late PaleozoicPaganzo basin of western Argentina and its regional implications. In: Tankard, A.,Suárez Soruco, R., Adf, W.H.J. (Eds.), Petroleum Basins of South America. AAPG,pp. 285–304.

Fisher, N.D., Jordan, T.E., Brown, L., 2002. The structural and stratigraphic evolution ofthe La Rioja basin, Argentina. Journal of South American Earth Sciences 15,141–156.

Furlong, K., Hanson, R.B., Bowers, J.B., 1991. Modeling thermal regimes. Reviews inMineralogy and Geochemistry 26, 437–506.

Garreaud, R., Vuille, M., Clement, A.C., 2003. The climate of the Altiplano; observedcurrent conditions and mechanisms of past changes. Palaeogeography, Palaeocli-matology, Palaeoecology 194, 5–22.

Garzione, C.N., Molnar, P., Libarkin, J.C., MacFadden, B.J., 2006. Rapid late Miocene rise ofthe Bolivian Altiplano: evidence for removal of mantle lithosphere. Earth andPlanetary Science Letters 241 (3–4), 543–556.

Gaupp, R., Kött, A., Wörner, G., 1999. Paleoclimatic implications of Mio-Pliocenesedimentation in the high-altitude intra-arc Lauca Basin of northern Chile.Palaeogeography, Palaeoclimatology, Palaeoecology 151, 79–100.

Gephart, J.E., 1994. Topography and subduction geometry in the Central Andes: clues tothe mechanics of a non-collisional orogen. Journal of Geophysical Research 99,12,279–12,288.

Ghosh, P., Garzione, C.N., Eiler, J.M., 2006. Rapid uplift of the Altiplano revealed through13C–18O bonds in paleosol carbonates. Science 311, 511–515.

Gilbert, G.K., 1877. Report on the Geology of the HenryMountains. Government PrintingOffice, Washington, D.C. 170 pp.

González-Bonorino, F., 1950. Algunos problemas geológicos de las Sierras Pampeanas.Revista de la Asociacion Geologica Argentina 5, 81–110.

Graham, C.M., England, P.C.,1976. Thermal regimes and regionalmetamorphism in thevicinity of overthrust faults: an example of shear heating and invertedmetamorphic zonation from southern California. Earth and Planetary ScienceLetters 31, 142–152.

Gregory-Wodzicki, K.M., 2000. Uplift history of the central and northern Andes: areview. Geological Society of America Bulletin 112 (7), 1092–1105.

Grier, M.E., Salfity, J.A., Allmendinger, R.W., 1991. Andean reactivation of the CretaceousSalta rift, northwestern Argentina. Journal of South American Earth Sciences 4 (4),351–372.

Grujic, D., et al., 2006. Climatic forcing of erosion, landscape, and tectonics in the BhutanHimalayas. Geology 34 (10), 801–804.

Hack, J.T., 1960. Interpretation of erosional topography in humid temperate regions.American Journal of Science 258-A, 80–97.

Halloy, S., 1982. Contribucíon al estudio de la zona de Huaca huasi, Cumbres Calchaquíes(Tucumán Argentina). Univ. Nac., Tucumán, Tucumán, Argentina. 841 pp.

Halpern, M., Latorre, C., 1973. Estudio geocronológico inicial de las rocas del noroeste dela República Argentina. Revista Asociacion Geologica Argentina 28, 195–205.

Hartley, A.J., Flint, S., Turner, P., Jolley, E.J., 1992. Tectonic controls on the development ofsemi-arid, alluvial basins as reflected in the stratigraphyof thePurilactisGroup (UpperCretaceous–Eocene), northern Chile. Journal of South American Earth Sciences 5,275–296.

Hasegawa, A., Sacks, I.S., 1981. Subduction of the Nazca Plate beneath Peru as determinedfrom seismic observations. Journal of Geophysical Research 86, 4971–4980.

Haselton, K., Hilley, G.E., Strecker, M.R., 2002. Average Pleistocene climatic patterns inthe Southern Central Andes: controls on mountain glaciation and palaeoclimateimplications. Journal of Geology 110, 211–226.


Hausen, H., 1925. Sobre un perfil geológico del borde oriental de la Puna de Atacama.Acad. Nac. Cienc. Córdoba, Bol. XXVIII, 1–95.

Hilley, G.E., Strecker, M.R., 2004. Steady state erosion of critical Coulomb wedges withapplications toTaiwan and the Himalaya. Journal of Geophysical Research 109, B01411.

Hilley, G.E., Strecker, M.R., 2005. Processes of oscillatory basin filling and excavation in atectonically active orogen: Quebrada del Toro Basin, NW Argentina, GSA Bulletin117 (7–8), 887–901. doi:10.1130/B25602.1.

Hilley, G.E., Strecker, M.R., Ramos, V.A., 2004. Growth and erosion of fold-and-thrustbelts, with an application to the Aconcagua fold-and-thrust belt, Argentina. Journalof Geophysical Research 109.

Hilley, G.E., Blisniuk, P.M., Strecker, M.R., 2005. Mechanics and erosion of basement-cored uplift provinces. Journal of Geophysical Research 110, B12409.

Hodges, K.V., 2000. Tectonics of the HImalaya and southern Tibet from two perspectives.Geological Society of America Bulletin 112 (3), 324–350.

Hoffmann, J.A.J.,1975. Climatic atlas of South America.WorldMeteorological Organization,Geneva, pp. scale: 1:1,000,000, 28 pp.

Hoffmann-Rothe, A., et al., 2006. Oblique convergence along the Chilean margin:partitioning, margin-parallel faulting and force interaction at the plate interface. In:Oncken, O., et al. (Ed.), The Andes: Active Subduction Orogeny. Springer, Berlin,pp. 125–146.

Hoke, G.D., Garzione, C.N., 2008. Paleoelevation and geomorphic constraints on the lateMiocene rise of the Andes. Earth and Planetary Science Letters 271, 192–201.

Hongn, F., et al., 2007. Middle Eocene deformation and sedimentation in the Puna–EasternCordillera transition (23°–26°S): control by preexisting heterogeneities on the patternof initial Andean shortening. Geology 35 (3), 271–274.

Horton, B.K., 1997. The modern foreland basin system adjacent to the central Andes.Geology 25, 895–898.

Horton, B.K., 1998. Sediment accumulation on top of the Andean orogenic wedge:Oligocene to late Miocene basins of the Eastern Cordillera, southern Bolivia.Geological Society of America Bulletin 110, 1174–1192.

Horton, B.K., 1999. Erosional control on the geometry and kinematics of thrust beltdevelopment in the central Andes. Tectonics 18 (6), 1292–1304.

Horton, B.K., DeCelles, P.G., 2001. Modern and ancient fluvial megafans in the forelandbasin system of the central Andes, southern Bolivai: implications for drainagenetwork evolution in fold–thrust belts. Basin Research 13 (43–63).

Horton, B.K., Hampton, B.A., Lareau, B.N., Baldellon, E., 2002. Tertiary provenance history ofthe northern and central Altiplano (central Andes, Bolivia): a detrital record ofplateau-margin tectonics. Journal of Sedimenatary Research 72 (5), 711–726.

Hoth, S., Adam, J., Kukowski, N., Onken, O., 2006. Influence of erosion on the kinematicsof bivergent orogens; results from scaled sandbox simulations. In: Willett, S.D.,Hovius, N., Brandon, M., Fisher, D.M. (Eds.), Tectonics, Climate and LandscapeEvolution: Special Paper, Geological Society of America. Geological Socity ofAmerica, Denver, pp. 201–225.

Houston, J., Hartley, A.J., 2003. The central Andean West-Slope rainshadow and itspotential contribution to the origin of hyper-aridity in the Atacama Desert.International Journal of Climatology 23, 1453–1464.

Isacks, B.,1988. Uplift of the central Andeanplateau and bending of the Bolivian orocline.Journal of Geophysical Research, B, Solid Earth and Planets 93, 3211–3231.

Jordan, T.E., Allmendinger, R.W., 1986. The Sierras Pampeanas of Argentina: a modernanalogue of Rocky Mountain foreland deformation. American Journal of Science286, 737–764.

Jordan, T.E., et al., 1983. Andean tectonics related to the geometry of subducted Nazcaplate. Bulletin of the Geological Society of America 94, 341–361.

Jordan, T.E., Allmendinger, R.W., Damanti, J.F., Drake, R.E., 1993. Chronology of motion ina complete thrust belt: the Precordillera, 30–31 S, Andes Mountains. Journal ofGeology 101, 135–156.

Kay, S.M., Coira, B., Viramonte, J., 1994. Youngmafic back-arc volcanic rocks as indicatorsof continental lithospheric delamination beneath the Argentine Puna plateau,Central Andes. Journal of Geophysical Research, B, Solid Earth and Planets 99 (B12)242323-242339.

Kennan, L., Lamb, S., Rundle, C., 1995. K–Ar dates from the Altiplano and CordilleraOriental of Bolivia: implications for Cenzoic stratigraphy and tectonics. Journal ofSouth American Earth Sciences 8 (2), 163–186.

Kleinert, K., Strecker, M.R., 2001. Changes in moisture regime and ecology in response tolate Cenozoic orographic barriers: the Santa Maria Valley, Argentina. GeologicalSociety of America Bulletin 113, 728–742.

Kley, J., 1996. Transition from basement-involved to thin-skinned thrusting in theCordillera Oriental of Southern Bolivia. Tectonics 15 (4), 763–775.

Kley, J., Monaldi, C.R., 2002. Tectonic inversion in the Santa Barbara System of the centralAndean foreland thrust belt, northwestern Argentina. Tectonics 21. doi:10.1029/2002TC902003 NO.6, 1061.

Kley, J., Reinhardt, M., 1994. Geothermal and tectonic evolution of the Eastern Cordilleraand the Subandean Ranges of southern Bolivia. In: Reutter, K.-J., Scheuber, E.,Wigger, P.J. (Eds.), Tectonics of the Southern Central Andes. Springer-Verlag, NewYork, pp. 155–170.

Kley, J., Gangui, A.H., Krüger, D., 1996. Basement-involved blind thrusting in the easternCordillera Oriental, southern Bolivia: evidence from cross-sectional balancing,gravimentric and magnetotelluric data. Tectonophysics 259, 171–184.

Kley, J., Monaldi, C.R., Salfity, J.A., 1999. Along-strike segmentation of the Andeanforeland: causes and consequences. Tectonophysics 301, 75–94.

Kley, J., Rosello, E.A., Monaldi, C.R., Habighorst, B., 2005. Seismic and field evidence forselective inversion of Cretaceous normal faults, Salta rift, northwest Argentina.Tectonophysics 399, 155–172.

Knox, J.C., 1983. Responses of river systems to Holocene climates. In: WrightJr. Jr., H.E.(Ed.), Late-Quaternary Environments of the United States. University of MinnesotaPress, Minneapolis, pp. 26–41.

Lachenbruch, A.H., Sass, J.H., 1980. Heat flow and energetics of the San Andreas FaultZone. Journal of Geophysical Research 85, 6,185–6,222.

Lamb, S., Davis, P., 2003. Cenozoic climate change as a possible cause for the rise of theAndes. Nature 425, 792–797.

Lamb, S., Hoke, L., 1997. Origin of the high plateau in the central Andes, Bolivia, SouthAmerica. Tectonics 16, 623–649.

Lamb, S., Hoke, L., Kennan, L., Dewey, J., 1997. Cenozoic evolution of the Central Andes inBolivia and northern Chile. In: Burg, J.-P., Ford,M. (Eds.), Orogeny Through Time, Geol.Soc. Spec. Publ. Geological Society of America, Denver, CO, USA, pp. 237–264.

Letcher, A.J., 2007. Deformation history of the Susques basin (~23°S, 66°W), PunaPlateau, NW Argentina: new constraints by apatite (U–Th)/He thermochronologyand 40Ar–39Ar geochronology, and implications for plateau formation in the centralAndes. Stanford University, Stanford, CA. 111 pp.

Lork, A., Bahlburg, H., 1993. Precise U–Pb ages of monazites from the Faja Eruptiva de laPuna Oriental and the Cordillera Oriental, NW Argentina. XII Congreso GeológicoArgentino y II Congreso de Exploración de Hidrocarburos, pp. 1–6.

Marquillas, R.A., del Papa, C., Sabino, I.F., 2005. Sedimentary aspects and paleoenviron-mental evolution of a rift basin: Salta Group (Cretaceous–Paleogene), northwesternArgentina. International Journal of Earth Sciences 94, 94–113.

Marrett, R.A., Strecker, M.R., 2000. Response of intracontinental deformation in thecentral Andes to late Cenozoic reorganization of South American Plate motions.Tectonics 19, 452–467.

Marshall, L.G., Patterson, B., 1981. Geologic and geochronology of the mammal-bearingTertiary of the Valle de Santa Maria and Rio Corral Quemado, Catamarca Province,Argentina. Fieldiana Geology 9, 1–80.

Masek, J.G., Isacks, B.L., Gubbels, T.L., Fielding, E.J., 1994. Erosion and tectonics at themargins of continental plateaus. Journal of Geophysical Research 99, 13941–13956.

May, G., et al., 2005. Eocene to Pleistocene lithostratigraphy, chronostratigraphy andtectono-sedimentary evolution of the Calama Basin, northern Chile. RevistaGeologica de Chile 32, 33–58.

McQuarrie, N., 2002. The kinematic history of the central Andean fold–thrust belt,Bolivia: implications for building a high plateau. Geological Society of AmericaBulletin 114 (8), 950–963.

McQuarrie, N., DeCelles, P.G., 2001. Geometry and strutural evolution of the centralAndean backthrust belt, Bolivia. Tectonics 20 (5), 669–692.

McQuarrie, N., Barnes, J.B., Ehlers, T.A., 2008a. Geometric, kinematic, and erosionalhistory of the central Andean Plateau, Bolivia (15–17°S). Tectonics 27.

McQuarrie, N., Ehlers, T.A., Barnes, J., Meade, B., 2008b. Temporal variation in climateand tectonic coupling in the central Andes. Geology 36, 999–1002.

Mendez, V., Navarini, A., Plaza, D., Viera, V., 1972. Faja eruptiva de la Puna oriental. 5thCongresso Geol. Argent. Actas, 4, pp. 89–100.

Molnar, P., England, P.C., 1990a. Late Cenozoic uplift of mountain ranges and globalclimatic change: chicken or egg? Nature 346, 29–34.

Molnar, P., England, P.C., 1990b. Temperatures, heat flux, and frictional stress near majorthrust faults. Journal of Geophysical Research 95, 4833–4856.

Molnar, P., Garzione, C.N., 2007. Bounds on the viscosity coefficient of continentallithosphere from removal of mantle lithosphere beneath the Altiplano and EasternCordillera. Tectonics 26, TC2013.

Mon, R.,1979. Esquema tectónico de Los Andes del norte Argentino. Revisita AssociacionGeologica Argentina XXXIV, 53–60.

Mon, R., Hongn, F., 1987. Estructura del Ordovícico de la Puna. Revista de la AsociaciónGeológica Argentina 42, 31–38.

Monaldi, C.R., Salfity, J.A., Kley, J., 2008. Preserved extensional structures in an invertedCretaceous rift basin, northwestern Argentina: outcrop examples and implicationsfor fault reactivation. Tectonics 27, TC1011.

Montgomery, D.R., Balco, G., Willett, S.D., 2001. Climate, tectonics, and morphology ofthe Andes. Geology 29 (7), 579–582.

Nasif, N., Musalem, S., Esteban, G., Herbst, R., 1997. Primer registro de vertebrados para laFormación Las Arcas (mioceno tardio), Valle de Santa María, Provincia de Catamarca,Argentina. Ameghiniana 34, 538.

Nogués-Paegele, J., Mo, K.C., 1997. Alternating wet and dry conditions over South Americaduring summer. Monthly Weather Review 125, 279–291.

Norabuena, E., et al., 1998. Space geodetic observations of the Nazca–south Americanconvergence across the central Andes. Science 279, 358–362.

Omarini, R.H., 1983. Caracterización litológica, diferenciación y genesis de la FormaciónPuncoviscana entre el Valle de Lerma y la Faja Eruptiva de la Puna, UniversidadNacional de Salta, Argentina, Salta, Argentina, 220 pp.

Padula, E., et al., 1967. Devonian of Argentina. International Symposium on the DevonianSystem, pp. 165–199.

Parrish, J.T., Ziegler, A.M., Scotese, C.R., 1982. Rainfall patterns and the distribution ofcoals and evaporates in the Mesozoic and Cenozoic. Palaeogeography, Palaeocli-matology, Palaeoecology 40, 67–101.

Pascual, R., Ortiz Jaureguizar, E., 1990. Evolving climates and mammal faunas inCenozoic South America. Journal of Human Evolution (19), 23–60.

Pascual, R., Vucetich, M.G., Scillato-Yané, G.J., Bond, M., 1985. Main pathways ofmammalian diversification in South America. In: Stehli, F., Webb, S.D. (Eds.), TheGreat American Biotic Interchange. Plenum, New York, pp. 219–247.

Persson, K., Garcia-Castellanos, D., Sokoutis, D., 2004. River transport effects oncompressional belts: first results from an integrated analogue-numerical model.Journal of Geophysical Research 109, B01409.

Ramos, V.A., 2000. Evolución Tectónica de la Argentina. In: Caminos, R. (Ed.), GeologíaArgentina. InInstituto de Geología y Recursos Minerales, Buenos Aires, pp. 715–784.

Ramos, V.A., 2008. The basement of the central Andes: the Arequipa and relatedterranes. Annual Review of Earth and Planetary Sciences 36, 289–324.

Ramos, V.A., et al., 1986. Paleozoic terranes of the central Argentine–Chilean Andes.Tectonics 5, 855–880.

http://dx.doi.org/10.1029/2002TC902003

http://dx.doi.org/10.1029/2002TC902003


Ramos, V.A., et al., 1996a. Geología de la región del Aconcagua, Provincias de San Juan yMendoza. Direccíon Nacional del Servicio Geológico Anales 24, 1–510.

Ramos, V.A., Cegarra, M., Lo Forte, G., Comínguez, A., 1996b. Cenozoic tectonics of the HighAndes of west-central Argentina (30–36 S latitude). Tectonophysics 259, 185–200.

Ramos, V.A., Cristallini, E.O., Pérez, D.J., 2002. The Pampean flat-slab of the centralAndes. Journal of South American Earth Sciences 15, 59–78.

Ramos, V., Zapata, T., Cristallini, E., Intracaso, A., 2004. The Andean thrust system—

latitudinal variations in structural styles and orogenic shortening. In: McClay, K.R.(Ed.), Thrust Tectonics and Hydrocarbon Systems: AAPG Memoir, pp. 30–50.

Raymo, M.E., Ruddiman, W.F., 1992. Tectonic forcing of late Cenozoic climate. Nature359, 117–122.

Reutter, K.-J., et al., 2006. The Salar de Atacama Basin: a subsiding block within thewestern edge of the Altiplano–Puna Plateau. In: Oncken, O., et al. (Ed.), The AndesActive Subduction Orogeny. Springer Berlin Heidelberg, Berlin, pp. 303–325.

Roeder, D., 1988. Andean-age structure of Eastern Cordillera (province of La Paz,Bolivia). Tectonics 7, 23–39.

Roeder, D., Chamberlain, R.L.,1995. Structural geology of Subandean fold-and-thrust belt inNorthwestern Argentina. In: Tankard, A.J., Suárez, S.R., Welsink, H.J. (Eds.), PetroleumBasins of SouthAmerica. InAmericanAssociationofPetroleumGeologists, pp. 459–479.

Rohmeder, W., 1943. Observaciones meterológicas en la región encumbrada de lasSierras de Famatina y del Aconquija (republica Argentina). Anales de la SociedadCientifica Argentina 136, 97–124.

Rolleri, E.O., 1976. Sistema de Santa Bárbara: Una neuva provincia geológica Argentina,Sexto Congreso Geológico Argentino. Asoc. Geol. Argent, Bahía Blanca.

Royden, L., 1996. Coupling and decoupling of crust and mantle in covergent orogens:implications for strain partitioning in the crust. Journal of Geophysical Research,Solid Earth 101, 17679–17705.

Royden, L.H., et al., 1997. Surface deformation and lower crustal flow in eastern Tibet.Science 276, 788–790.

Ruddiman, W.F. (Ed.), 1997. Tectonic Uplift and Climate Change. Plenum Press, NewYork. 517 pp.

Ruiz Huidobro, O., 1960. Descripción Geológica de la Hoja 8e, Chicoana, Provincia deSalta, Buenos Aires.

Rutter, E.H., Brodie, K.H., 2004. Experimental grain size-sensitive flow of hot-pressedBraziliam quartz aggregates. Journal of Structural Geology 26, 2011–2023.

Ryback, L., Cermak, V., 1982. Radioactive heat generation in rocks. In: Angenheister, G.(Ed.), Landolt–Bornstein Numerical Data and Functional Relationships in Scienceand Technology, Group V. Springer-Verlag, Berlin, pp. 353–371.

Rybacki, E., Dresen, G., 2000. Dislocation and diffusion creep of synthetic anorthiteaggregates. Journal of Geophysical Research 105, 26,017–26,036.

Sáez, A., Barera, L., Jensen, A., Chong, G., 1999. Late Neogene lacustrine record andpaleogeography in the Quillangua–Llamara basin, Central Andean fore-arc (north-ern Chile). Palaeogeography, Palaeoclimatology, Palaeoecology 151, 5–37.

Salfity, J.A., 1982. Evolucion paleogeographica del grupo Salta (Cretacico–Eogenico),Argentina. 5th Cong. Latino Am. Geol. Actas, 1, pp. 11–26.

Salfity, J.A., Marquillas, R.A., 1994. Tectonic and sedimentary evolution of the Cretaceous–Eocene Salta Group Basin, Argentina. In: Salfity, J.A. (Ed.), Cretaceous Tectonics of theAndes. In: Earth Evolution Sciences. Vieweg and Sohn, Friedr, pp. 266–315.

Salfity, J.A., Omarini, R., Baldis, B., Gutierrez, W.J., 1976. Consideraciones sobre la evoluciongeologica del Precambrico y Paleozoico del norte Argentino. 2nd Congresso Ibero Am.Geol., 4, pp. 341–353.

Schmidt, C.J., Astini, R.A., Costa, C.H., Gardini, C.E., Kraemer, P.E., 1995. Cretaceous rifting,alluvial fan sedimentation, and Neogene inversion, southern Sierras Pampeanas,Argentina. In: Tankard, A.J., Suárez, R.S., Welsink, H.J. (Eds.), Petroleum Basins ofSouth America, pp. 341–358.

Scotese, C.R., Gahagan, L.M., Larson, R.L., 1988. Plate tectonic reconstructions of theCretaceous and Cenozoic ocean basins. Tectonophysics 155, 27–48.

Sempere, T., Herail, G., Oller, J., Bonhomme, M.G., 1990. Late Oligocene–early Miocenemajor tectonic crisis and related basins in Bolivia. Geology 18, 946–949.

Sillitoe, R., McKee, E.H., Vila, T., 1991. Reconnaissance K–Ar geochronology of theMaricunga gold–silver belt, Northern Chile. Economic Geology 86, 1261–1270.

Sklar, L.S., Dietrich, W.E., 2004. A mechanistic model for river incision into bedrock bysaltating bed load. Water Resources Research 40 21 pp.

Sleep, N.H., 2005. Evolution of the continental lithosphere. Annual Review of Earth andPlanetary Sciences 33, 369–393.

Sobel, E.R., Strecker, M.R., 2003. Uplift, exhumation, and precipitation: tectonics andclimatic control of Late Cenozoic landscape evolution in the northern SierrasPampeanas, Argentina. Basin Research 15.

Sobel, E.R., Hilley, G.E., Strecker, M.R., 2003. Formation of internally drained contractionalbasins by aridity-limited bedrock incision. Journal of Geophysical Research, B, SolidEarth and Planets 108 (B7), 2344.

Sosic, M.V.J.,1972. Descripción geológica de la hoja 14d, Tinogasta (Provincia de Catamarcay La Rioja). Dirección Nacional de Geología y Minería, Boletín, 129: 56 pp.

Starck, D., Anzótegui, L.M., 2001. The late Miocene climatic change: persistence of aclimatic signal through the orogenic stratigraphic record in northwesternArgentina. Journal of South American Earth Sciences 14, 763–774.

Stauder, W., 1975. Subduction of the Nazca plate under Peru as evidenced by focalmechanisms and by seismicity. Journal of Geophysical Research 80, 1053–1064.

Stelzner, A., 1923. Contribucion a la geología de la Republica Argentina, con la partelimitrofe de los Andes Chilenos entre los 32 y 33°S. Actas Acad. Nac. Cienc. Cordoba8, 1–28.

Stock, J.D., Montgomery, D.R., 1999. Geologic contraints on bedrock river incision usingthe stream power law. Journal of Geophysical Research, B, Solid Earth and Planets104 (B3), 4983–4993.

Stolar, D., Roe, G., Willett, S.D., 2007. Controls on the patterns of topography and erosionrate in a critical orogen. Journal of Geophysical Research, Earth Surface 112 F04002-F04002.

Strecker, M.R., Cerveny, P., Bloom, A.L., Malizia, D., 1989. Late Cenozoic tectonism andlandscape development in the foreland of the Andes: Northern Sierras Pampeanas,Argentina. Tectonics 8, 517–534.

Strecker, M.R., et al., 2007. Tectonics and climate of the southern Central Andes. AnnualReview of Earth and Planetary Sciences 35, 747–787.

Stüwe, K., 2002. Geodynamics of the Lithosphere, an Introduction. Springer Verlag, Berlin.Thiede, R.C., et al., 2005. From tectonically to erosionally controlled development of the

Himalayan orogen. Geology 33 (8), 689–692.Turner, J.C.M., 1959. Estratigrafía del cordón de Escaya y de la sierra de Rinconada

(Jujuy). Revista de la Asociación Geológica Argentina 13, 15–39.Turner, J.C.M., 1960a. Estratigrafía de la Sierra de Santa Victoria y adyaciencas. Bol. Acad.

Cienc. Córdoba 41, 163–196.Turner, J.C.M., 1960b. Estratigrafía del Nevado de Cachi y sector al oeste. Acta Geológica

Lilloana, Tucumán 3, 191–226.Turner, J.C.M., Méndez, V., 1979. Puna. II Simposio de geologia Regional Argentina,

pp. 13–56.Uba, C.E., Heubeck, C., Hulka, C., 2005. Facies analysis and basin architecture of the

Neogene Subandean synorogenic wedge, southern Bolivia. Sedimentary Geology180, 91–123.

Uba, C.E., Heubeck, C., Hulka, C., 2006. Evolution of the late Cenozoic Chaco forelandbasin, southern Bolivia. Basin Research 18, 145–170.

Vandervoort, D.S., Jordan, T.E., Zeitler, P.K., Alonso, R., 1995. Chronology of internaldrainage development and uplift, southern Puna plateau, Arentine central Andes.Geology 23, 145–148.

Whipple, K.X., Meade, B.J., 2004. Controls on the strength of coupling among climate,erosion, and deformation in two-sided, frictional orogenic wedges at steady-state.Journal of Geophysical Research, Earth Surface 109, F01011.

Whipple, K.X., Kirby, E., Brocklehurst, S.H., 1999. Geomorphic limits to climate-inducedincreases in topographic relief. Nature 401, 39–43.

Whitman, D., Isacks, B.L., Kay, S.M., 1996. Lithospheric structure and along-strikesegmentation of the central Andean Plateau: topography, tectonics, and timing.Tectonophysics 259, 29–40.

Willett, S.D., 1999. Orography and orogeny: the effects of erosion on the structure ofmountain belts. Journal of Geophysical Research 104 (B12), 28,957–29,981.

Willett, S.D., Beaumont, C., Fullsack, P., 1993. A mechanical model for the tectonics ofdoubly vergent compressional orogens. Geology 30, 175–179.

WMO, 1975. Climatic Atlas of South America. World Meteorological Organization,Geneva, pp. scale: 1:1,000,000, 28 pp.

Zachos, J., Pagani, M., Sloan, L., Thomas, E., Billups, K., 2001. Trends, rhythms, andaberrations in global climate 65 Ma to present. Science 292, 686–693.

Zapata, T.R., Allmendinger, R.W., 1996. Growth stratal records of instantaneous andprogressive limb rotation in the Precordillera thrust belt and Bermjo basin,Argentina. Tectonics 15 (5), 1065–1083.

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