tectonics of sedimentary basins (recent advances) || cenozoic evolution of hinterland basins in the...

Chapter 21

Cenozoic evolution of hinterland basins in the Andes and Tibet

BRIAN K. HORTON

Department of Geological Sciences and Institute for Geophysics, Jackson School of Geosciences,University of Texas at Austin, Austin, USA

ABSTRACT

Sedimentary basins have been generated in hinterland regions of the Andean andHimalayan-Tibetan orogens during Cenozoic plate convergence. These hinterlandbasins record nonmarine sediment accumulation (commonly in high-elevation, low-relief, internally drained, arid/semiarid settings) during protracted deformation andsurface uplift of continental crust. In South America, Andean hinterland basins areproduced between the western magmatic arc and craton-directed fold-thrust belt to theeast. In the India-Asia collision zone, hinterland basins develop as elements of thegrowing Tibetan plateau between the Himalayan thrust belt and Asian plate interior tothe north. Hinterland sedimentary basins in theAndes andTibet are distinguished fromforeland basins adjacent to their respective fold-thrust belts by structural position,elevation, and stratigraphic evolution. Sediment accommodation in these hinterlandbasins is created by flexure, fault-induced subsidence, and topographic ponding, withcommon examples of remnant foreland basins succeeded by younger hinterland basins.

Keywords: Altiplano; Andes; Tibet; fold-thrust belts; foreland basins; hinterland;plateau

INTRODUCTION

The “hinterland” of a convergent orogen encom-passes the elevated interior region away from thelow-elevation foreland basin and external foothillsof the fold-thrust belt. Situated above regional baselevel, hinterland basins are susceptible to erosionand less likely to bepreserved in the geologic record.Nevertheless, hinterlands of large modern orogens,notably the Andes and Himalayan-Tibetan system,show long-term depositional records spanningtens of millions of years (Jordan and Alonso, 1987;Allmendinger et al., 1997; Yin and Harrison, 2000;Tapponnier et al., 2001). Stratigraphic historiesinclude nonmarine sedimentation at high eleva-tions, possibly up to 3–5 km above sea level(Garzione et al., 2000, 2006; Rowley and Currie,2006; DeCelles et al., 2007a; Saylor et al., 2009).

This chapter summarizes the tectonics of basinevolution in the Andean hinterland and Tibetanplateau, potential endmembers of noncollisionalretroarc orogens and collisional orogens, respec-

tively. Both systems have recorded crustal short-ening, thickening, and surface uplift of low-reliefhinterlands since early Cenozoic or Cretaceoustime. The Andes (Fig. 21.1) are the type exampleof an ocean-continent convergent plate boundary,providing insights into subduction zones,magmatic arcs, retroarc thrust belts, and retroarcforeland basins (Jordan et al., 1983; Isacks, 1988;Horton andDeCelles, 1997; James and Sacks, 1999;Kley et al., 1999; Beck and Zandt, 2002; Kay et al.,2005; McQuarrie et al., 2005; Uba et al., 2006). TheHimalayan-Tibetan orogen (Fig. 21.2) is a keyexample of continent-continent collision and acrucible of concepts on contractional tectonicsand peripheral foreland basins (Molnar andTapponnier, 1975, Tapponnier et al., 1982;Molnaret al., 1993; Royden et al., 1997; DeCelleset al., 1998; Chemenda et al., 2000; Hodges, 2000;Lave and Avouac, 2000; Beaumont et al., 2001).

The focus in this chapter is on the structuralsetting and generalized Cenozoic stratigraphy ofrepresentative hinterland basins in the Andes and

Tectonics of Sedimentary Basins: Recent Advances, First Edition. Edited by Cathy Busby and Antonio Azor.

� 2012 Blackwell Publishing Ltd. Published 2012 by Blackwell Publishing Ltd.

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Fig. 21.1. Topographicmap and satellite image of theAndes. Digital elevationmodel (fromNASA)depicts themajorAndeanhinterland basins (outlined and shaded) concentrated in the central andnorthernAndes. (a) Tres Cruces basin (Punaplateau,Argentina), (b) Altiplano basin (Altiplano plateau, Bolivia), (c) Callejon de Huaylas basin (Cordillera Blanca, Peru),(d) Cuenca basin (Interandean Valley, Ecuador), and (e) Magdalena Valley basin (Colombia). True-color satellite image(from NASA Terra/MODIS) shows saline lakes (SdU: Salar de Uyuni), freshwater lakes (LT: Lake Titicaca), and rivers (DR:Desaguadero river) of the central Andean hinterland.

428 Part 4: Convergent Margins

Tibet; details on the sedimentology, provenance,and dispersal patterns are accessible in cited refer-ences. In summarizing these two convergent sys-tems, it becomes clear that multiple modes of basindevelopment operate in hinterland regions,

including crustal flexure due to thrust loading,fault-induced subsidence along strike-slip or exten-sional faults, and topographic ponding in internallydrained regions. Although some hinterland basinsform as newly generated features upon a deformed

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Fig. 21.2. Topographic map and satellite image of the Himalayan-Tibetan orogen. Digital elevation model (fromMichael H.Taylor; after Yin, 2000) depicts the major Tibetan hinterland basins (outlined and shaded), with warmer colors representinghigher elevations. (a) Thakkhola basin (southern Tibet), (b) Lunpola basin (central Tibet), (c) Hoh Xil basin (north-centralTibet), (d) Qaidam basin (northern Tibet), and (e) Xining basin (northeastern Tibet). True-color satellite image (from NASATerra/MODIS) shows lakes of internally drained south-central Tibet (SC: Siling Co; NC: Nam Co), external drainage nearLhasa (L) in southeastern Tibet, and longitudinal drainage (TR: Tsangpo river) of southernmost Tibet.

Cenozoic Evolution of Hinterland Basins 429

orogenic substrate, a common pattern for both theAndean retroarc hinterland and Tibetan collisionalhinterland involves the conversion of former fore-land basin systems into elevated hinterland basinsas the deformation front advances cratonward awayfrom the orogenic interior, incorporating the fore-land basin into the growing orogen.

ANDEAN HINTERLAND BASINS

Tectonic setting

The Andes (Fig. 21.1) are the product of noncolli-sional deformation related to eastward subductionof the oceanic Nazca slab beneath South America.The Wadati-Benioff zone varies from moderateeastward dip (�30�) to subhorizontal orientationsalong the �7000-km-long convergent margin(James, 1971; Barazangi and Isacks, 1976). Subduc-tion and arc magmatism generally began in theMesozoic. Although overprinted by Cenozoicshortening, reconstructions suggest neutral to ten-sile stress during the Mesozoic (Dalziel, 1981;Mpodozis andRamos, 1989; Salfity andMarquillas,1994; Cooper et al., 1995). Absolute westwardmotion of SouthAmerica commencedwith riftingfrom Africa in the late Early Cretaceous (Coneyand Evenchick, 1994), but substantial compres-sion and mountain building was delayed untilCenozoic time.

The onset of Andean shortening, although spa-tially variable, is attributed to earlyCenozoicor LateCretaceous convergence (Steinmann, 1929;M�egard, 1984; Dalziel, 1986; Wilson, 1991; Dengoand Covey, 1993; Ramos andAleman, 2000; Hortonet al., 2001; McQuarrie et al., 2005, 2008;Mpodoziset al., 2005). Some workers consider shortening tobe limited to the past 25–30 Myr, with possibleevidence for accelerated shortening in lateMiocenetime (Isacks, 1988; Sempere et al., 1990; Gubbelset al., 1993;Hindle et al., 2002).Minimumestimatesof east-west shortening vary along strike and rangefrom �20 to �350km, with maximum shorteningand surface uplift observed in the central Andes(10–30�S) (Kley, 1996; Baby et al., 1997; Kley andMonaldi, 1998; Kley et al., 1999;McQuarrie, 2002a,2002b). In addition to crustal shortening and thick-ening, lithospheric thinning and/or removal bydelamination or other process is important inparts of the central Andean plateau (Kay andKay, 1993; Whitman et al., 1996; Beck andZandt, 2002; Schurr et al., 2006). Major shorteningappears to be linked to trenchward (westward)

acceleration of South America (Sobolev andBabeyko, 2005; Schellart, 2008), with additionalinfluences from erosional forces (Masek et al.,1994; Horton, 1999; Lamb and Davis, 2003), mag-matic processes (Babeyko et al., 2002; DeCelleset al., 2009), flat-slab dynamics (Gutscher et al.,2000; Ramos et al., 2002), and lower crustal flow(Lamb and Hoke, 1997; Kley and Monaldi, 1998).

Basin overview

Sedimentary basins occupy the retroarc hinter-land between the Andean magmatic arc andtrailing (western) flank of the craton-directed(east-vergent) thrust belt (Fig. 21.1). Cenozoicfill of these hinterland basins is exposed in out-crop belts typically parallel to regional tectonicstrike. Exposures are 10–100kmwide by 10–500kmlong and bounded by diverse faults, folds, and topo-graphic features (Jordan and Alonso, 1987; Kennanet al., 1995; Marocco et al., 1995; Horton, 1998).

The basins are characterized by structures devel-oped during and after filling, including fold-thruststructures, some strike-slip faults, and rare normalfaults. Andeanhinterland basins vary from�100 to50,000km2 and contain principally nonmarinesuccessions >2–12km thick with accumulationrates commonly >200m/Myr (Horton et al., 2001;Garzione et al., 2008). These thicknesses and ratesmatch those reported for Andean foreland, wedge-top, and piggyback basins (Reynolds et al., 1990;Jordan, 1995; Echavarria et al., 2003;Horton, 2005).

Modern accumulation in the Andean hinter-land involves extensive lacustrine, fluvial, andeolian sedimentation, commonly in arid/semiaridsettings. In internally drained regions such asthe Altiplano-Puna (central Andean) plateau(Fig. 21.1), freshwater lakes (e.g., Lake Titicaca)and saline lakes (e.g., Salar de Uyuni) coexist withaxial fluvial systems (e.g., Desaguadero river).Activedeposition also occurs in externallydrainedhinterland regions with axial rivers parallel totectonic strike, such as the Magdalena River inColombia and the Ucayali River in Peru.

Representative basins

Consideration of a limited set of Andean hinterlandbasins reveals depositional histories in five zones(Fig. 21.3): the Puna plateau (Argentina), Altiplanoplateau (Bolivia), Cordillera Blanca (Peru), Interan-dean Valley (Ecuador), and Magdalena Valley(Colombia). A compilation of the associated


deformational histories (Fig. 21.3) highlights theconsiderable uncertainties in estimating the onsetof shortening in different parts of the Andes.

The structural positions and stratigraphic histo-ries of Andean hinterland basins are distinguishedfrom the low-elevation forelandbasin systemoccu-pying the lowlands of the Amazon, Parana, andOrinoco drainage systems (Jordan, 1995; Hortonand DeCelles, 1997; de Berc et al., 2005). On thebasis of their structural boundaries, elevation,and protracted histories, the basins also can be

discerned from short-lived wedge-top or piggy-back basins situated between topographic boun-daries created by faults within the thrust belt(Beer et al., 1990; Marocco et al., 1995; Horton,1998, 2005; Mosolf et al., 2011).

(1) In the 4–5km-high Puna plateau (Fig. 21.1)of northern Argentina (21–26�S), compart-mentalized basins generally <100,000km2 and1–6km thick show similar depositional histo-ries (Fig. 21.3a). In most basins, Cretaceous

MAGDALENAVALLEY BASIN

(Colombia)

CUENCA BASIN(Interandean

Valley, Ecuador)

CALLEJON DE HUAYLAS BASIN

(Cord. Blanca, Peru)

ALTIPLANO BASIN(Altiplano, Bolivia)

TRES CRUCESBASIN

(Puna, Argentina)

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Cretaceous strataand pre-Cretaceousmetamorphic rocks

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Cretaceousstrata

Pan de Azucar Fm.(200 m)

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(200 m)

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(5000 m)

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(3000 m)

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(800 m)

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(2000 m)

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(2000 m)

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(4000 m)

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LisamaFm.

(1000 m)

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(1100 m)

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(800 m)

Mesa Fm.(600 m)

ColoradoFm.

(1200 m)

Turi Fm. (450 m)

Quingeo Fm.(1200 m)

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Mangan Fm.(1200 m)

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(>1000 m)

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RealFm.

(1800 m)

ConiriFm.

(2000 m)

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Biblian Fm.(300 m)

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Fig. 21.3. Cenozoic nonmarine stratigraphic records and inferred deformational histories for Andean hinterland basins(locations in Fig. 21.1). Stratigraphy: clastic deposition (stippled); limited marine deposition (shaded); syndepositionalvolcanism (V lettering); no preserved record (vertical ruling). Deformation: shortening (narrow black bars); extension (whitebars); strike-slip deformation (gray bars); uncertainties in timing constraints (dashed bars).


nonmarine tomarginalmarine sedimentation inthe Salta rift preceded early Cenozoic nonmar-ine deposition attributable to postrift thermalsubsidence (Jordan and Alonso, 1987; Salfityand Marquillas, 1994) or initial flexural subsi-dence in a foreland basin (Kraemer et al., 1999;Carrapa and DeCelles, 2008). The generation ofhinterland basins began with mid-Cenozoicfragmentation of the former foreland basin, theresult of cratonward (eastward) advance of thethrust belt (Schwab, 1985; Boll and Hern�an-dez, 1986; Coutand et al., 2001; Hongnet al., 2007; Carrapa et al., 2008). This led totopographic isolation and initiation of internaldrainage that persists today as sedimentationcontinues in various isolated basins, commonlyin evaporative saline lakes or salars (Alonsoet al., 1991; Vandervoort et al., 1995). Neogenegrowth of orographic barriers during uppercrustal shortening helped generate a uniquegeomorphic history of intermontane processesinvolving periodic topographic ponding andexcavation (Sobel et al., 2003; Hilley andStrecker, 2005; Coutand et al., 2006; Streckeret al., 2007, this volume; Hain et al., 2011).

(2) In Bolivia, the 4-km-high Altiplano plateau(Fig. 21.1) at 16–21�S contains a large modernbasin that has existed in a hinterland settingsince at least �25 Ma (Marshall et al., 1992;Allmendinger et al., 1997; Lamb and Hoke,1997; Horton et al., 2002a; Hampton and Hor-ton, 2007;Murray et al., 2010). In contrast to thePuna plateau, the Altiplano represents a singleintegratedbasin>500,000 km2with aCenozoicsuccession up to 12 km thick (Fig. 21.3b).Although a Cretaceous marginal marinesection is present locally, there is limited evi-dence for the Cretaceous extension thataffected the Puna region (Welsink et al.,1995). The Cenozoic history of the Altiplanoinvolved deposition of a thick nonmarine suc-cession attributed to initial foreland basin con-ditions, although strike-slip and normal faultsmay be locally important (Lamb andHoke, 1997; Elger et al., 2005). Topographicisolation of theAltiplano basin has been linkedto a major cratonward advance of the thrustfront (McQuarrie, 2002a; DeCelles and Hor-ton, 2003; McQuarrie et al., 2005). Uplift anderosion of the western arc and eastern thrustbelt are expressed in the provenance recordof the Altiplano (Horton et al., 2002a) andlow-temperature thermochronology of plateau

margins (Barnes et al., 2006; Gillis et al., 2006;Ege et al., 2007).

(3) The only well-developed extensional basinin the Andean hinterland is located adjacentto the Cordillera Blanca (Fig. 21.1), a >5km-high range in the PeruvianAndes. TheCallejondeHuaylas supradetachment basin (Fig. 21.3c)is situated in the hanging wall of the active,20–45� west-dipping Cordillera Blanca normalfault (Bonnot et al., 1988; Schwartz, 1988;Petford and Atherton, 1992; McNulty andFarber, 2002). The >1.3 km-thick successionwas deposited on folded Jurassic-Cretaceousstrata and lower Cenozoic arc volcanic rocks(Giovanni et al., 2010). Basin initiation roughlycoincided with late Miocene flattening of thesubductedNazca slab and cessation of arcmag-matism at 2–14�S (Hampel, 2002;McNulty andFarber, 2002). Normal faulting and basin evo-lution are the products of gravitational collapse(Dalmayrac and Molnar, 1981; S�ebrier et al.,1988) and strain partitioning related to obliqueplate kinematics (Petford and Atherton, 1992;McNultyetal., 1998). Incontrast to theCordilleraBlanca, other hinterland basins in Peru (Ayacu-cho, Cajabamba, San Marcos, and Namorabasins) are linked to localized transpressionaland transtensional structures (M�egard et al.,1984; Marocco et al., 1995; Wise et al., 2008).

(4) Small Neogene basins in the Interandean Val-ley of Ecuador (Fig. 21.1) are situated at �3 kmelevation between the Western Cordillera arcand Eastern Cordillera thrust belt. Individualbasins (Cuenca, Gir�on-Santa Isabel, Nab�on,Loja, and Malacatos-Vilcabamba basins) are<1000 km2 and contain <2–5 km of middleMiocene toQuaternary fill (Fig. 21.3d) (Lavenuet al., 1992, 1995). A lack of Paleogene fill inthe hinterland of Ecuador suggests erosion ornondeposition prior to Neogene activation ofdisconnected basins. Although these relation-ships are compatiblewith an elevated pre-Neo-gene orogen, marine fossils indicate that somehinterland areas remained at low elevationuntil �15–10 Ma (Hungerb€uhler et al., 1995,2002). Most basins are bounded by strike-slipand reverse faults near potential pre-Cenozoicsutures, suggestingsubsidence in transpressionalsettings involving possible fault reactivation(Hungerb€uhler et al., 2002; Winkler et al., 2005).

(5) The Magdalena Valley basin of Colombia(Figs. 21.1 and 21.3e) is situated between arelict Mesozoic-early Cenozoic arc (Central


Cordillera) and a Cenozoic fold-thrust belt(Eastern Cordillera). Although complicatedby terrane accretion in westernmost Colombia,Cretaceous and Cenozoic basin evolutionis attributed to a temporal transition fromregional extension to shortening (Dengo andCovey, 1993; Cooper et al., 1995). For the cen-tral and eastern parts of the Colombian Andes,marine sedimentation accompanying Creta-ceous synrift and postrift subsidence wasfollowed by accumulation of 2–5 km of non-marine sediment during early Cenozoic flex-ural subsidence. The switch to foreland basinconditions in the Magdalena Valley wasinduced by initial shortening and crustal load-ing in the Central Cordillera (G�omez et al.,2003, 2005a, 2005b; Horton et al., 2010).Subsequent partitioning of this early Cenozoicforeland basin by shortening-related uplift ofthe Eastern Cordillera, as expressed in thedetrital provenance record (Nie et al., 2010;Moreno et al., 2011), produced an isolatedhinterland basin that accommodated an addi-tional 2–5kmofNeogene sediment (VanHoutenand Travis, 1968; Van Houten, 1976).

TIBETAN HINTERLAND BASINS

Tectonic setting

The Himalayas and Tibetan plateau (Fig. 21.2)constitute Earth’s highest mountain belt and larg-est orogenic plateau. Cenozoic construction of theHimalayan-Tibetan orogen has been driven by col-lision and continued northward advance of theIndia plate relative toAsia. Estimates for the timingof India-Asia collision center on �55 Ma (Searleet al., 1987; Dewey et al., 1989; Rowley, 1996),withalternative suggestions of collision by �70–60 Ma(Jaeger et al., 1989; Yin and Harrison, 2000) or�34Ma (Aitchison et al., 2007). Following initialcollision, Indian lithosphere has been underthrustup to 700 km northward beneath southern andcentral Tibet (Le Pichon et al., 1992; Owens andZandt, 1997; DeCelles et al., 2002).

Initial north-south shortening during the colli-sion was focused in the presently elevated hinter-land region around the India-Asia (Indus-Yarlung-Tsangpo) suture zone and in the Tibetan plateauto the north. Estimating the onset of shorteningis complicated by pre-collisional deformationand uplift. Prior to collision, north-dipping sub-duction of the Indian oceanic slab accounted for the

Trans-Himalayan (Gangdese-Ladakh-Kohistan) mag-matic arc and associated retroarc shortening (Burget al., 1983; England and Searle, 1986; Kappet al., 2007; Leier et al., 2007a). Based on structuralrestorations, total shortening during the India-Asiacollision is estimated at 600–750km for the Hima-layas (DeCelles et al., 1998, 2002; Robinsonet al., 2006) and roughly 500–750km across theTibetan plateau (Murphy et al., 1997; Yin andHarrison, 2000; Kapp et al., 2005). Although paleo-magnetic data suggestmuch greater north-south con-vergence (up to 2500km) between Tibet and thestableAsian interior (Achache et al., 1984), structuralrestorations cannot account for such large values.Shallowing of magnetic inclination in Cretaceous-Cenozoic strata, rather than extreme shortening, mayexplain the paleomagnetic data (Tan et al., 2003).

Basin overview

Sedimentary basins in the interior of theHimalayan-Tibetan orogen (Fig. 21.2) formed during protractedIndia-Asia convergence. Hinterland basins of theTibetan plateau are distributed over a �1000km(north-south) by �2000 km (east-west) regionbetween the India-Asia suture zone in the southandthecratonic interiorofcentralAsia.Manybasinsare bounded by positive topographic features cre-ated by displacement along thrust and strike-slipfaults. Exposures of most Tibetan hinterland basinsoccur in restricted, fault-bounded regions (M�etivieret al., 1998; Liu et al., 2001; Horton et al., 2002b;DeCelles et al., 2007b; Saylor et al., 2009).

Faults associated with basin sedimentation sug-gest a broad northward progression of deformationtoward central Asia (Tapponnier et al., 2001;Royden et al., 2008). However, early Cenozoicactivation of some of the northernmost structuresof theTibetan plateau (Yin et al., 2002, 2007, 2008a,2008b; Horton et al., 2004; Dupont-Nivet et al.,2004) shows that much Cenozoic sedimentationtook place in a hinterland setting, rather than fore-land basin. Tibetan hinterland basins show largevariations in total thickness and accumulation rate.Whereas some basins exhibit <1km thicknessesand accumulation rates <50–100m/Myr (Hortonet al., 2002b), others contain successions up to6–10km thick with rates up to 1000m/Myr(M�etivier et al., 1998; Liu et al., 2001; Fanget al., 2003) comparable to the Himalayan forelandbasin (Burbank et al., 1996; DeCelles et al., 1998).

Active sedimentation in the modern plateau(Fig. 21.2) occurs dominantly under arid/semiarid


conditions in compartmentalized, internally orpoorly drained basins commonly with freshwaterand saline lakes. Modern fluvial processes aremore pronounced in eastern Tibet, where deepincision linked to large external drainages (e.g.,Yellow, Yangtze, Mekong, Salween, Red, andTsangpo rivers) has characterized the lateCenozoic to modern geomorphology. Neverthe-less, active accumulation occurs in select regionssuch as the modern Zoige basin of eastern Tibet

where wetlands fluviolacustrine depositionoccurs in part of the Yellow river drainage(Chen et al., 1999).

Representative basins

Collisional hinterland basins in the Tibetan pla-teau exhibit varied depositional anddeformationalhistories, as summarized here for five regions(Fig. 21.4): India-Asia suture zone (southernmost

ShangganchaigouFm.

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THAKKHOLA BASIN(southern Tibet)

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XINING BASIN(northeast Tibet)

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(1000 m)

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LuleheFm.

(1040 m)

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(1000 m)

XiayoushashanFm.

(1240 m)

Shizigou Fm.(> 1500 m)

QijiachuanFm.

(80 m)

HonggouFm.

(350 m)

MahalagouFm.

(370 m)

XiajiaFm.

(140 m)

ChetougouFm.

(240 m)

XianshuiheFm.

(60 m)

FenghoushanGp.

(4800 m)

Yaxicuo Fm.(670 m)

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(2800 m)

DingqingFm.

(1100 m)

Wudaoliang Fm.(370 m)

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Fig. 21.4. Cenozoic nonmarine stratigraphic records and inferred deformational histories for Tibetan hinterland basins(locations in Fig. 21.2). Stratigraphy: clastic deposition (stippled); and no preserved record (vertical ruling). Deformation:shortening (narrow black bars); extension (white bars); strike-slip deformation (gray bars); and uncertainties in timingconstraints (dashed bars).


Tibet), Lhasa-Qiangtang terrane boundary (centralTibet), Fenghuo Shan, Qaidam basin, and north-eastern Tibet. The onset of shortening in theseregions broadly corresponds to initial India-Asiacollision. High-elevation hinterland basins ofTibet (Fig. 21.2) developed independently of theHimalayan fold-thrust belt to the south, and aredistinguished by their structural settings andstratigraphic records from the low-elevationforeland basins and collisional successor basinsbordering the plateau, including the Himalayan,Sichuan, Hexi, Tarim, Junggar, Turpan, and vari-ous Mongolian basins (Hendrix et al., 1992;Graham et al., 1993, 2001; Burbank et al., 1996;DeCelles et al., 1998; Johnson, 2004; Johnson andRitts, this volume).

(1) In southernmost Tibet (Fig. 21.2), a series ofmiddle Miocene to Quaternary hinterlandbasins are the product of fault-induced subsi-dence along north-striking normal faults. Indi-vidual rift basins such as the Thakkhola graben(Fig. 21.4a) contain up to 5 km of nonmarinesediment over exposure areas of �20,000 km2

(Garzione et al., 2003). These north-trendingbasins are distinguished from east-trendingextensional basins associated with the SouthTibetan detachment system south of theIndia-Asia suture (Burchfiel et al., 1992;Hodges, 2000). East-west extension in Tibethas been attributed to gravitational collapse(Molnar et al., 1993), axial arc extension(McCaffrey and Nabalek, 1998), India collision(Kapp and Guynn, 2004), and subduction roll-back along the Pacific plate boundary(Northrup et al., 1995; Yin, 2000).

(2) In the Lhasa terrane of southern and centralTibet (Fig. 21.2), a Cretaceous retroarc forelandbasin evolvednorth of theGangdese arcprior toIndia-Asia collision (Kapp et al., 2007; Leieret al., 2007a). Others have suggested Mesozoicgrowth of backarc extensional basins(Zhang, 2000; Chen et al., 2003) and/or aperipheral foreland basin related to collisionof the Lhasa and Qiangtang terranes (Kappet al., 2005; Leier et al., 2007b). About 3–4kmof sediment was shed into these marine andnonmarine basins from shortening-relatedstructures to the south and possibly to thenorth in the Qiangtang terrane. Following ini-tial India-Asia collision, the former forelandand extensional basins were overprinted bydisconnected basins (including the Duba and

Lunpola basins; Fig. 21.4b) and their boundingthrust faults (Xu, 1984; Leeder et al., 1988).These Cenozoic basins, represented by 4–5km-thick successions exposed over areas up to20,000km2, were generated by flexure and/orfootwall tilting and filled by proximal fan andlacustrine sedimentation. Isotopic studies ofupper Eocene to Miocene carbonates suggestbasin filling at 3–5 km elevation (Rowleyand Currie, 2006). Additional Neogene basinsinvolved subsidence induced by extension(Blisniuk et al., 2001; Kapp et al., 2008) andlateral extrusion of wedge-shaped blocksbetween kinematically linked, conjugatestrike-slip faults (Taylor et al., 2003, thisvolume).

(3) In north-central Tibet, the Fenghuo Shanregion (Fig. 21.2) consists of the Hoh Xilbasin complex (Fig. 21.4c) partially separatedby thrust-related topography. Three principalsubbasins (Fenghuoshan-Hantaishan, Wudao-liang, and Zhuolai Lake subbasins) wereformed in proximity to fold-thrust structuresactive during Eocene-Oligocene convergence(Liu and Wang, 2001; Liu et al., 2001).Nonmarine fill reaches a maximum thicknessof �6 km, with exposures distributed over100,000 km2. In contrast to southern Tibet,isotopic studies imply theHohXil basin devel-oped at elevations near �2km and was lateruplifted to �5km (Cyr et al., 2005; Wanget al., 2008). During the early Miocene, contin-ued filling overtopped a topographic barrier,leading to integration of adjacent subbasins. Inother places, sedimentation had ceased by thelatest Oligocene with extensive erosion gener-ating mature geomorphic (peneplain) surfaces(Wang et al., 2002). A series of small, elongatenonmarine basins in east-central and south-eastern Tibet (Fig. 21.2) are partial equivalentsof the larger Hoh Xil basin complex in theFenghuo Shan region. The narrow Nangqian-Yushu and Gonjo basin systems are 2–4 km-thick features parallel to tectonic strike andbounded by thrusts and limited strike-slipfaults (Horton et al., 2002b; Spurlin et al.,2005; Studnicki-Gizbert et al., 2008).

(4) The Qaidam basin of northern Tibet (Figs. 21.2and 21.4d) is an internally drained region bor-dered by topographic barriers produced bymotion along major thrust and strike-slipfaults. Although the �100,000 km2 basin hasregions of active accumulation, deformation


along the margins has reduced the originaldepositional area (Bally et al., 1986; Yinet al., 2002, 2007, 2008a). The Qaidam basincontains a nearly complete record of Cenozoicsedimentation, with up to 10 km of principallynonmarine fill (Song and Wang, 1993; Yinet al., 2008b), although early Cenozoic marinefossils have been found along the northernedge of Tibet (Ritts et al., 2008). During theearly Cenozoic, subsidence was presumablydriven by flexural subsidence due to crustalthickening along the southern margin (Ballyet al., 1986; Yin et al., 2007, 2008a, 2008b).Later accumulation is considered the productof topographic isolation by basin closure pro-moted by slip along the Altyn Tagh strike-slipfault bounding the northern margin of Tibet(M�etivier et al., 1998; Tapponnier et al., 2001;Yin et al., 2002). The Qaidam basin is distin-guished from the larger collisional successorbasins of central Asia (Hendrix et al., 1992;Graham et al., 1993) on the basis of its struc-tural position within the Tibetan plateau,higher elevation, and genetic links to basin-bounding structures.

(5) In the northeastern corner of Tibet (Fig. 21.2),compartmentalization (partitioning) of formerextensive basins typifies the Cenozoic record.The region contains a series of thrust systems(Qilian Shan, Nan Shan, and Liupan Shan)kinematically linked to major strike-slip faults(Altyn Tagh and Haiyuan faults) (Tapponnieret al., 1990; Burchfiel et al., 1991). Althoughthese structures are generally attributed to lat-est Cenozoic deformation, paleomagnetic andstratigraphic evidence indicate pre-Neogenedeformation (Dupont-Nivet et al., 2004; Hortonet al., 2004). Prior to India-Asia collision, accu-mulation of 1–3 km-thick Jurassic-Cretaceoussuccessions is partially linked to extension ortranstension (Vincent and Allen, 1999). Theselarger basins were compartmentalized duringCenozoic uplift that promoted accumulation inlocalized basins, such as the Xining basin(Fig. 21.4e), controlled by variable thrust andstrike-slip structures (Fang et al., 2003; Hortonet al., 2004).

DISCUSSION AND CONCLUSIONS

Despite fundamental differences in the tectonicsettings of noncollisional retroarc orogens and

collisional systems, some similarities are identi-fied between the hinterland regions of the Andeanand Himalayan-Tibetan orogens. Both systemscontain high elevation plateaus, regions of internaldrainage, zones of active hinterland sedimenta-tion, and exposures of older basin fill that accu-mulated during Cenozoic mountain building.Although many Andean and Tibetan hinterlandbasins have distinct depositional histories, severalprocesses appear to govern basin evolution in theorogenic interiors of retroarc and collisional sys-tems. Two particular modes of hinterland basinevolution are evident in the Andes and Tibet.

Thefirstcategoryconsistsofhinterlandbasinsthatdeveloped as newly activated features in regionspreviously experiencing erosion and/or deforma-tion. In this case, basin evolution can be attributedto the initiation of slip on particular structures asso-ciated with shortening, extension, and strike-slipdeformation within the orogenic interior. Reverseslip on steep, basement-involved faults generatedisolated ranges bounded by individual flexuralbasins in the Puna plateau (Schwab, 1985; Hongnetal.,2007)andDuba-LunpolaregionofcentralTibet(Leederetal., 1988).Hinterlandextensionproducedriftandsupradetachmentbasins inlocationssuchasthe Peruvian Andes (McNulty and Farber, 2002;Giovanni et al., 2010) and southernmost Tibet(Yin, 2000; Garzione et al., 2003). Additional exam-ples of fault-induced and flexural subsidence helpexplain basin evolution in transtensional and trans-pressional hinterland settings, notably in centralTibet (Taylor et al., 2003, this volume) and Ecuador(Hungerb€uhler et al., 2002; Winkler et al., 2005).

A second common mode of basin evolutioninvolves the evolution of former foreland basins,including wedge-top depozones (DeCelles andGiles, 1996;DeCelles, this volume), intohinterlandbasins as the deformation front propagates craton-ward and incorporates the foreland basin into theadvancing orogen (Figs. 21.5 and 21.6). Key exam-ples of hinterland basins overprinting or succeed-ing foreland basins include the Altiplano-Puna(central Andean) plateau and parts of southernandnorthernTibet. However, each system is some-what unique. First, in the central Andean plateau(Fig. 21.1), the orogenic front did not necessarilymigrate as a steadily advancing wave of deforma-tion. The Puna plateau experienced a nonsystem-atic pattern of deformation, with isolated basinscreated at different times during Cenozoic short-ening (Coutand et al., 2001, 2006; Carrapa andDeCelles, 2008). The Altiplano hinterland basin


was created when the deformation front abruptlyjumped�200 km cratonward (eastward), inducinga pronounced shift in sediment dispersal patterns(Fig. 21.5) (Horton et al., 2002a; DeCelles andHorton, 2003; McQuarrie et al., 2005). A similarjump in the orogenic front may have promoted thehinterland growth phase of the Magdalena Valleybasin in the northern Andes (G�omez et al., 2003,2005a, 2005b; Nie et al., 2010; Moreno et al., 2011).Second, in the southern and central Tibetan pla-teau (Fig. 21.2), a Cretaceous forelandbasinwas thecomplex product of crustal loading on both flanksof the basin. Flexure was the result of peripheralloading in the north related to the Qiangtang-Lhasaterrane collision and retroarc deformation in thesouth associated with the ocean-continent subduc-tion boundary (and Gangdese arc) that precededIndia-Asia collision (Leier et al., 2007a).Hinterland

basin conditions were only established later afterthe early Cenozoic India-Asia collision promotedshortening across much of the Tibetan plateau.Third, the Hoh Xil basin of northern Tibet(Fig. 21.2) represents a potential case where oro-gen-directed backthrusting during northwardgrowth of the plateau induced a shift from forelandto hinterland accumulation (Fig. 21.6) (Liuet al., 2001; Yin et al., 2007). Farther north, theQaidam basin was trapped by a combination ofstrike-slip deformation and shortening along thepresent edge of the plateau (Metivier et al., 1998;Yin et al., 2002, 2008a, 2008b). Although theseexamples from the Andes and Tibet are variationson a theme, a potentially unifying element is thatnewly activated uplift of topographic barriershelped to shield the former foreland basin andsucceeding hinterland basin from large-scale

Western Cordillera

Altiplano EAST

back-bulgeforebulge

Eastern Cordillera

(a)

(c)

(b)

craton

craton

craton

WEST

Early Miocene

Late Eocene – Oligocene

hinterlandbasin

foredeep

forebulgethrust belt jumps

toward craton foredeep

Subandean Zone

Early – MiddleEocene

Fig. 21.5. Schematic reconstruction of Cenozoic foreland and hinterland basin evolution in the Altiplano plateau andneighboring regions of the central Andes, showing sediment accumulation (stippled), forebulge zone of erosion or severelylimited accumulation (thick dashed lines), sediment disperal patterns (large arrows), and generalized thrust structures atdepth. Source: After DeCelles and Horton (2003).

FenghuoShan

EasternKunlun Shan

Qaidam basinHoh Xilbasin

foredeep

hinterland basin deformation front jumps toward craton

Early Miocene

Eocene

Oligocene

NORTH

(a)

(c)

(b)

SOUTH

backthrust

backthrust

Fig. 21.6. Schematic reconstruction of possible Cenozoic transition from foreland to hinterland basin evolution in the HohXil basin and neighboring regions of northern Tibet, showing sediment accumulation (stippled), sediment disperal patterns(large arrows), and generalized thrust structures at depth. Source: Modified from Yin et al. (2007).


erosional processes. These new barriers help createorographic rain shadows (Sobel et al., 2003;Strecker et al., 2007), further promoting arid/semi-arid conditions, internal drainage, and continuedinfilling of elevated hinterland basins that rival thethickness and lifespan of low-elevation forelandbasins.

ACKNOWLEDGMENTS

Constructive reviews by Andrew Leier, CariJohnson, and Antonio Azor improved the manu-script. I thank the following individuals for bene-ficial discussions on various regions of (1) theAndes and (2) Tibet: (1a) Ricardo Alonso, BarbaraCarrapa, Peter DeCelles, Fernando Hongn, EdwardSobel, Daniel Starck, and Manfred Strecker; (1b)Robert Gillis, Brian Hampton, N�estor Jim�enez,Ramiro Matos, Nadine McQuarrie, and BryanMurray; (1c) Melissa Giovanni and BrendanMcNulty; (1d)DominikHungerb€uhler; (1e)GermanBayona, Andr�es Mora, and Mauricio Parra; (2a)Michael Murphy, Joel Saylor, and Daniel Stockli;(2b) Peter Blisniuk, PaulKapp, andMichael Taylor;(2c) Brian Currie, Peter Lippert, Zhifei Liu,Matthew Spurlin, and Xixi Zhao; (2d) RaymondIngersoll,MichaelMcRivette, BradleyRitts, andAnYin; and (2e) Guillaume Dupont-Nivet, CarmalaGarzione, Junsheng Nie, and Jiangyu Zhou.

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tectonics of sedimentary basins (recent advances) || cenozoic evolution of hinterland basins in the...

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