oligocene range uplift and development of plateau...

Oligocene range uplift and development of plateau morphology

in the southern central Andes

B. Carrapa,1 D. Adelmann,2 G. E. Hilley,3 E. Mortimer,1 E. R. Sobel,1 and M. R. Strecker1

Received 26 October 2004; revised 14 February 2005; accepted 11 April 2005; published 13 August 2005.

[1] The Puna-Altiplano plateau in South America is ahigh-elevation, low internal relief landform that ischaracterized by internal drainage and hyperaridity.Thermochronologic and sedimentologic observationsfrom the Sierra de Calalaste region in the southwesternPuna plateau, Argentina, place new constraints onearly plateau evolution by resolving the timing ofuplift of mountain ranges that bound present-daybasins and the filling pattern of these basins duringlate Eocene-Miocene time. Paleocurrent indicators,sedimentary provenance analyses, and apatite fissiontrack thermochronology indicate that the originalpaleodrainage setting was disrupted by exhumationand uplift of the Sierra de Calalaste range between24 and 29 Ma. This event was responsible forbasin reorganization and the disruption of theregional fluvial system that has ultimately led to theformation of internal drainage conditions, which, inthe Salar de Antofalla, were established not later thanlate Miocene. Upper Eocene-Oligocene sedimentaryrocks flanking the range contain features that suggestan arid environment existed prior to and during itsuplift. Provenance data indicate a common similarsource located to the west for both the southern Punaand the Altiplano of Bolivia during the late Eocene-Oligocene with sporadic local sources. This suggeststhe existence of an extensive, longitudinally orientedforeland basin along the central Andes during thistime. Citation: Carrapa, B., D. Adelmann, G. E. Hilley,

E. Mortimer, E. R. Sobel, and M. R. Strecker (2005), Oligocene

range uplift and development of plateau morphology in the

southern central Andes, Tectonics, 24, TC4011, doi:10.1029/

2004TC001762.

1. Introduction

[2] The central Andean Altiplano-Puna plateau is ahyperarid, low internal relief, high-elevation region with

average and peak elevations greater than 3700 and 6000 m,respectively. Uplift of this high-elevation region has beenascribed to processes such as lithospheric thinning [Isacks,1988] following delamination [Kay et al., 1994], distributedcrustal shortening [Allmendinger et al., 1997], emplacementof regional basement thrust sheets [Kley et al., 1997;McQuarrie and DeCelles, 2001], and underthrusting ofthe Brazilian craton [Isacks, 1988]. Whereas these processesmay have created much of the surface uplift in the area, thelow-relief morphology, typical of continental orogenicplateaus, may result from simultaneous erosion of highpeaks and deposition within basins, as the incising powerof regional drainage systems is lost in such arid environ-ments [e.g., Sobel et al., 2003; Hilley and Strecker, 2005].Despite the importance of internal drainage conditions inmany of the world’s large continental plateaus and thepotential impact of sediment storage on their evolution aswell as adjacent foreland areas [Vandervoort et al., 1995;Metivier et al., 1998; Tapponnier et al., 2001; Horton et al.,2002; Sobel et al., 2003], the controls on their establishmentremain elusive.[3] The clastic fill preserved in intramontane basins

within the plateau contains the unique record of the timingand pattern of orogenic evolution and its relationship totectonics and climate. By constraining the sedimentaryevolution of such basins and the uplift of related ranges, abetter understanding of the processes leading to finalinternal drainage can be achieved.[4] Clastic sediments preserved in the Puna plateau

suggest that sedimentation within the plateau started in thelate Eocene in a broad foreland basin sourced from the west[Jordan and Alonso, 1987]. However, detailed sedimento-logical investigations which could assess this hypothesis arestill limited and therefore its validity must be further tested.During the Miocene, and possibly Oligocene, the appear-ance of evaporites in northern Argentina documents theonset of hyperarid conditions and has been directly relatedto the establishment of internal drainage in the area [e.g.,Vandervoort et al., 1995; Alonso et al., 1991]. Thermochro-nologic and provenance data from the eastern Puna marginsuggest that some topography may have formed as early asthe late Eocene–early Oligocene [Muruaga, 2001; Coutandet al., 2001; Deeken et al., 2004]. The eastern rangesconstituted a topographic high at least by middle Miocenetime [e.g., Strecker, 1987; Strecker et al., 1989; Grier andDallmeyer, 1990; Marrett and Strecker, 2000; Kleinert andStrecker, 2001; Hilley and Strecker, 2005]. However, therelationships between range exhumation and uplift, sedi-ment dispersal and final internal drainage developmentremain unclear.

TECTONICS, VOL. 24, TC4011, doi:10.1029/2004TC001762, 2005

1Institut fur Geowissenschaften, Universitat Potsdam, Potsdam,Germany.

2Institut fur Geowissenschaften, Friedrich-Schiller-Universitat Jena,Jena, Germany.

3Department of Earth and Planetary Science, University of California,Berkeley, California, USA.

Copyright 2005 by the American Geophysical Union.0278-7407/05/2004TC001762$12.00

TC4011 1 of 19

[5] In addition, constraints on the timing of initiation ofdeformation related to range uplift and basin compartmen-talization within the present plateau area in northwesternArgentina are limited. At present, a widely accepted modelproposes that deformation leading to crustal thickening andsubsequent uplift occurred during the middle-late Miocene[e.g., Allmendinger, 1986; Isacks, 1988; Allmendinger etal., 1997; Jordan et al., 1997, 2001]. However, subsequentstudies document pre-Miocene deformation in the presentplateau of Argentina [e.g., Coutand et al., 2001] andBolivia [McQuarrie and DeCelles, 2001; Horton et al.,2001, 2002; DeCelles and Horton, 2003; Ege, 2004; Elger,2004].[6] Extensive studies exist for the Altiplano and Eastern

Cordillera of Bolivia documenting an initial pattern offoreland basin development, followed by structural disrup-tion, drainage internalization and compartmentalization ofsediment basins during Eocene through early Miocenetime [e.g., Horton et al., 2001; Horton et al., 2002;Ege, 2004]. Comparable comprehensive investigations thatdocument the timing and pattern of such processes arescarce in the Puna region. In particular, it remains unclearif deformation driving marginal range uplift, basin com-partmentalization and subsequent infill of related sedimen-tary basins result from tectonic processes that form the

plateau or from climate conditions that may be largelyindependent of plateau formation [e.g., Hartley, 2003;Sobel et al., 2003].[7] Combined apatite fission track thermochronology

(AFTT) and sedimentologic investigations in the south-western Puna plateau (Figure 1) provide information onthe timing of exhumation and related uplift of basin-bounding ranges, the sedimentary dynamics within adja-cent basins, and climatic conditions during the tectonicevents that preceded plateau uplift. In the southern Puna,sedimentary basins that contain upper Eocene to Quater-nary sedimentary rocks are bounded by high-angle reversefaults whose hanging wall rocks form bedrock-coredmountain ranges (Figure 2). By constraining the timingof exhumation and related uplift of a bounding range andchanges in the sedimentary dynamics of the basins, wesuggest that exhumation and uplift of ranges might havebeen the trigger for basin compartmentalization whicheventually led to internal drainage development. Ourresults support the hypothesis that the deformation drivingrange uplift started at least by Oligocene time, contributingto the establishment of internal drainage and to thecharacteristic high elevation, low internal relief observedtoday within the Puna plateau.

2. Regional Setting

2.1. Tectonic Evolution of the Central Southern Andes

[8] The Puna-Altiplano plateau is part of the centralAndes and represents the second largest plateau on Earthafter Tibet. The southern Puna of northwestern Argentinaconstitutes the southern end of the intraorogenic plateau. Itis bounded to the west by a magmatic arc and to thenortheast by the Eastern Cordillera fold-and-thrust belt,while the eastern border is transitional to the high-anglereverse-fault bounded Sierra Pampeanas basement blocks(Figure 1). The �3700-m-high southern Puna plateau ischaracterized by Neogene contraction [Alonso, 1986;Allmendinger et al., 1997; Coutand et al., 2001; Ege,2004], meridionally trending mountain ranges (often inexcess of 6000 m elevation), and internally draining sedi-mentary basins.[9] During the late Eocene, contractional to transpres-

sional deformation (‘‘Incaic phase’’ [Steinmann, 1929])involved Upper Cretaceous to Paleogene rocks of thepresent Western Cordillera [Gunther et al., 1998]. Defor-mation and uplift of this belt is inferred to have triggereddeposition of the earliest clastic sequences in an extensiveforeland basin spanning both the present-day plateau andregions to the east [e.g., Jordan and Alonso, 1987; Hortonet al., 2001; DeCelles and Horton, 2003]. This model issupported by apatite fission track thermochronology in theChilean cordillera, indicating considerable exhumationbetween 50 and 30 Ma [Maksaev and Zentilli, 2000] andby sedimentological data indicating westerly sourcedlate Eocene-Oligocene sedimentary rocks in the Altiplano[Horton et al., 2002] and possibly in the Puna [Jordan andAlonso, 1987]. Sedimentation continued in isolated intra-montane basins that received sediments from more local

Figure 1. General map of the central Andes includingdifferent morphotectonic domains and area over 3 kmelevation (gray) (modified after Horton et al. [2001]).

TC4011 CARRAPA ET AL.: OLIGOCENE UPLIFT IN THE CENTRAL ANDES

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sources from at least Miocene until present time [e.g.,Jordan and Alonso, 1987; Vandervoort, 1993; Horton etal., 2001; DeCelles and Horton, 2003].[10] Basins within the plateau contain thick sequences of

continental evaporites and clastic deposits that yield funda-mental information as to the cooling/exhumation history ofhinterland sources, sediment dispersal, and provenance.These basins are bounded structurally by high-angle reversefaults [e.g., Jordan et al., 1997]. The timing of clasticsedimentation in basins within the plateau and along theeastern Puna border is relatively well known based onmagnetostratigraphy and 40Ar/39Ar dating on ash layers insynorogenic deposits [Coira et al., 1993; Kay et al., 1994;Marrett and Strecker, 2000; Coutand et al., 2001].

2.2. Geology of the Sierra de Calalaste Area

[11] The study area is located in the southern Punabetween the Salar de Antofalla and the Sierra de Calalaste(Figures 2 and 3). Tertiary E-W to WNW-ESE shorteningproduced a series of east and west vergent reverse andthrust faults striking parallel to the present Salar de

Antofalla (Figure 4) [Voss, 2000; Adelmann, 2001]. TheSierra de Calalaste constitutes low-grade metamorphicbasement rocks that were deformed during and after lateEocene time [Adelmann, 2001]. Within the Sierra deCalalaste, Paleozoic sedimentary rocks are thrust overTertiary sedimentary rocks along west and east vergingreverse faults (Figure 3). Crystalline basement rocks,including migmatitic gneisses, metabasites, granitoids andaplites, as well as Tertiary volcanics rocks crop out to thewest and southwest of the present-day Salar de Antofallaarea [e.g., Kraemer et al., 1999, and references therein](Figures 2 and 3). In contrast, Precambrian sedimentaryand low-grade metamorphic rocks are more widespread tothe east (Figures 2 and 3).[12] In the study area, sedimentation started in late

Eocene–early Oligocene time with the deposition of theQuinoas Formation during the Incaic deformation phase[e.g., Kraemer et al., 1999; Adelmann, 2001; Voss, 2002].During the late Oligocene, thick-skinned compressivedeformation [Adelmann and Gorler, 1998] triggered sedi-mentation of syntectonic coarse-grained alluvial fans con-stituting the Chacras Formation (Figure 4). This phase of

Figure 2. (a) Shaded relief map of the central Andes, based on GTOPO30 data (USGS). AB, ArizaroBasin; AD, Atacama Desert; CA, Campo Arenal; HM, Salar de Hombre Muerto; QT, Quebrada del Toro;H, Humahuaca; SA, Salar de Antofalla; SC, Sierra de Calalaste; Scu, Siete Curvas; Spg, Salar de PastosGrande; TC, Tres Cruces; VC, Valles Calchaquıes basin. Faults are modified after Reutter et al. [1994],Urreiztieta et al. [1996], and Coutand et al. [2001]. Dots indicate volcanoes. Star denotes 30.3 ± 3 Maapatite fission track (AFT) age, after Andriessen and Reutter [1994]; triangle marks AFT ages between30 ± 2 Ma and 38 ± 3 Ma from Coutand et al. [2001]. (b) Simplified geological map of the central Andesmodified after McQuarrie [2002a]. For a more detailed geological map of the study area we refer toFigure 3.


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Figure 3. Detailed geologic map of the Antofalla area [see Kraemer et al., 1999] with location of theanalyzed samples and sedimentological logs.


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early deformation is also marked by an angular unconfor-mity between the Quinoas and the Chacras formations.Immediately beneath the unconformity, a tuff layer yieldsan 40Ar/39Ar age of 28.9 ± 0.8 Ma (ID-51) [Adelmann andGorler, 1998]. The oldest strata of the overlying ChacrasFormation have been dated at 24.2 ± 0.9 Ma (ID-86) and22.5 ± 0.6 Ma (ID-18) [Kraemer et al., 1999].[13] During the early Miocene (�20–17 Ma), renewed

E-W to WNW-ESE shortening [Adelmann and Gorler,1998] reactivated the west vergent fault system that wasactive during the preceding deformation phase. Additionally,Miocene shortening produced new east and westwarddirected basement thrusts onto tilted alluvial fan sedimentsof the Potrero Grande Formation [Adelmann, 2001].In the middle Miocene, west vergent thrusts affectedLower Paleozoic, Permian and Tertiary rocks [Adelmann,1997; Voss, 2000; Adelmann, 2001]. This deformationtriggered deposition of the syntectonic Juncalito Formation

(middle Miocene-Pliocene) which is characterized by thickevaporites (dated as late Miocene [Kraemer et al., 1999])typical of a hyperarid internal drainage environment[Kraemer et al., 1999].

3. Methods

[14] Sediment logging, facies interpretation, paleocurrentand provenance analyses were carried out on upper Eocene–lower Miocene sedimentary rocks (Quinoas and Chacrasformations) to reveal changes in sediment source and sedi-mentary environments (Figures 5 and 6 and Table 1). Fifteenthin sections were analyzed using the Gazzi-Dickinsonmethod [Dickinson, 1970; Gazzi et al., 1973]. An averageof 400 framework grains per thin section was counted onunstained thin sections. Petrographic counting parametersand recalculated detrital modes are reported in Table 2.Paleocurrent direction was determined by measuring at least

Figure 4. Overview of the tectonostratigraphic and magmatic development of the Salar de Antofalla areain the southern Puna modified after Kraemer et al. [1999], Voss [2000], and Adelmann [2001]. Undulatedlines represent angular unconformities of regional significance separating main lithostratigraphic unitspresent in the Salar de Antofalla area.


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50 imbricated pebbles of clast-supported channelized stream-flow and proximal sheet flood units in alluvial fan deposits.[15] AFTT was conducted on six samples collected along

a vertical transect through the Sierra de Calalaste betweenelevations of 3729 and 4455 m to constrain the cooling andexhumation history of the central part of the range. Sampleswere separated and analyzed following the proceduredescribed by Sobel and Strecker [2003] (Table 3). Rawdata were reduced using the Trackkey program (I. Dunkl,Trackkey: Windows program for calculation and graphicalpresentation of EDM fission track data, version 4.2, 2002,available at http://www.sediment.uni-goettingen.de/staff/dunkl/software/trackkey.html). Measurements of fissiontrack etch pits were made to assess annealing kineticvariability [Donelick et al., 1999]. Length measurementswere attempted on all samples in order to gain information

on the degree of annealing [e.g., Wagner and Van derHaute, 1992]. Exhumation rate was calculated using theinverse slope of weighted least squares regression of theAFT elevation versus ages.

4. Sedimentology

4.1. Quinoas Formation

[16] The Quinoas Formation records the onset of sedi-mentation in the study area in the late Eocene as determinedby an 40Ar/39Ar age of 37.6 ± 0.3 obtained from an ashlayer in the basal part of the formation [Kraemer et al.,1999]. The Quinoas Formation is divided into two membersbased on facies associations and a change in the interpreteddepositional environments [Kraemer et al., 1999].

Figure 5. Correlation panel of the investigated sedimentological logs indicated in Figure 3. For moredetails, refer to Figure 6.


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Figure 6. Detailed sedimentological logs: (a) log R; (b) log U; (c) log G; (d) log H/I; (e) log S. Agesindicated in log U are from Kraemer et al. [1999]. For legend, refer to Figure 5. The most typical faciesare indicated in italics. For a description of the facies we refer to Table 1.


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Figure 6. (continued)


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4.1.1. Quinoas I[17] This member is deposited above an erosional uncon-

formity with the underlying Permo-Carboniferous rocks. Atits maximum recorded thickness, this member reaches�840 m. It is characterized by fine-grained, gypsiferoussiltstones and mudstones (Fsm/m, Fl) intercalated withmassive, horizontally stratified and imbricated conglomer-ates with undulating basal contacts (Gc); lenticular clast-supported conglomerates (Gt/p); and interbedded massiveand laminated sandstones (Sh/l, Sm) (Table 1). Combinedthicknesses of conglomerate units can exceed 250 m(Figure 5). In the farthest sections to the west (logs Uand R), these fine grained facies dominate the sections,while more proximal to the Sierra de Calalaste the section iscoarser and dominated by the conglomeratic facies, thoughstill with intercalated finer horizons (Figure 5).[18] We interpret the gypsiferous siltstones and mud-

stones (Fsm/m, Fl) as being deposited in a playa mudflat[e.g., Flint, 1985; Hartley et al., 1992]. The presence of finegrained, laminated sediments, often bioturbated and withrooting and soil remnants and gypsum, is indicative of anarid, or at least episodically dry environment [e.g., Hardieet al., 1978]. The coarse-grained lenticular conglomerates(Gt/p) are interpreted as having been deposited in fluvialchannels [e.g., Hartley et al., 1992; Miall, 1996]. Massive,bedded and imbricated conglomerates (Gc) are interpreted tohave been deposited under tractive flow [Rasmussen, 2000].Horizontally stratified conglomerates (Gh) might reflectdiscontinuous discharge and accretion during high-densityflows and sheet floods [Nemec and Steel, 1984]. Theinterbedded, laminated and massive sandstones (Sh/l, Sm)which occur both within the conglomerates and throughoutthe finer-grained section represent waning flow conditionsand are probably distal derivatives of debris flows on thealluvial fan [e.g., Lowe, 1979; Nemec and Steel, 1984]. This

facies association is typical of ephemeral discharge insemiarid to arid alluvial fan environment [e.g., Flint andTurner, 1988; Sohn, 1997; Rasmussen, 2000]. Paleocurrentdata from member I, on the west side of the present-daySierra de Calalaste, suggest a provenance mainly from theN-NE and S-SW (Figure 3).4.1.2. Quinoas II[19] This member reaches a thickness of about 500 m

adjacent to the Sierra de Calalaste (Figures 5 and 6; log G)and conformably overlies the underlying member. It com-prises fine to very coarse-grained, trough cross-beddedsandstones and pebbly sandstones (St). These occur in bedsbetween 0.5 and 5.0 m thickness. Vertically stacked beds ofthis facies can combine to form significant thicknesses ofmore than 100 m (e.g., log G, U). Alternating with theselarger coarse-grained units are finer-grained sandstones(decimeter-scale beds) with scoured bases and occasionalclimbing ripples occurring in many of the sandstone units(Sr) and mudstones (Fsm/m, Fl) (Table 1). These mudstonesoften contain desiccation cracks, and are bioturbated (e.g.,Figure 6b, log U).[20] We interpret the trough cross-bedded sandstones as

being derived from the migration of dune bed forms withinsand bed channels of a fluvial depositional environment[Miall, 1996]. Scour based sandstones with climbing ripplesrepresent a lower flow regime within a fluvial environment[e.g., Allen, 1963]. The mudstones represent the finestcomponent of the system, and are likely to have beendeposited onto a floodplain environment, possibly throughcrevasse splay deposition. The close vertical associationbetween channel and floodplain deposits suggests deposi-tion onto a fluvial plain [Miall, 1996]. Such a faciesassociation might be representative of either meanderingor braided river systems [Miall, 1978; Smith, 1987]. How-ever, because of the significant thicknesses of trough cross-

Table 1. Lithofacies Description and Interpretation Based on Work by Miall [1996]

Facies Codes Lithofacies Sedimentary Structures Interpretation

Gm conglomerates, matrixsupported

massive, faint gradation high-strength (cohesive) debris flow

Gc conglomerates, clastsupported

massive, faint horizontal laminations,imbrications

high-strength or low-strength (noncohesive)debris flow

Gh conglomerates, clastsupported

horizontal laminations, imbrications longitudinal gravel banks, lag deposits

Gt/p conglomerates, clastsupported

trough and planar cross beds minor channel fills, transverse bed forms

Sm sand, fine to coarse massive or faint lamination distal debris flowSt sand, fine to very coarse,

may be pebblytrough cross beds subaqueous 3-D dunes, Transition or upper

part of a flow regimeSp sand, fine to very coarse,

may be pebblyplanar cross beds subaqueous 2-D dunes

Sh/l sand, fine to very coarse,may be pebbly

horizontal laminations/low-angle (<15�)cross beds

scour fills

Sr sand, very fine to coarse ripple, cross lamination ripples (lower flow regime)Seod sand, fine to medium,

well sortedlarge-scale cross lamination (>25�) aeolian dunes

Fl sand, silt, mud fine laminations, very small ripples overbank, abandoned channel or waningflood deposits

Fsm/m silt, mud massive, desiccation cracks overbank, abandoned or drape depositsP paleosol carbonate pedogenic features: nodules filaments soil with chemical precipitation


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Table

2.SandstonePetrographyParam

etersBased

MethodsDescribed

byIngersollet

al.[1984]andDickinson[1985]a

Sam

ple

Stratigraphy

Locationb

Mean

Grain

Sizec

Qm

Qp

Plag

K-Feldspar

Ls

Lv

Lm

Lp

Chert

Fragments

Minor

Constituents

Mica

Cem

ent

Matrix

QF

LQm

FLt

Qp

Lvm

Lsm

A091

ChacrasFm

logG

mS

20.5

5.6

4.5

16.5

11.2

14.4

6.7

0.5

2.7

0.8

1.1

6.9

8.8

32

25

43

25

25

50

24

43

33

A092

ChacrasFm

logG

mS

15.6

11.5

4.9

29.2

4.1

17.5

7.7

01.9

1.9

0.8

4.4

0.6

29

37

34

17

37

46

38

50

12

Average

31

31

39

21

31

48

31

47

23

SD

28

66

83

10

515

A075

Quinoas

Fm

IIlogG

mS/cS

26.7

9.9

6.3

20.2

2.6

14.8

5.1

2.0

3.7

2.3

0.6

2.6

3.4

40

29

31

29

29

42

44

48

8A081

Quinoas

Fm

IIlogH/I

mS

21.0

13.6

3.2

13.6

1.3

10.6

13.3

3.7

2.1

0.8

0.5

7.2

9.4

42

20

38

25

20

54

57

38

5A082

Quinoas

Fm

IIlogH/I

mS

25.7

13.5

6.1

16.6

6.9

5.0

6.1

1.4

4.4

1.4

1.4

4.1

7.5

46

26

28

30

26

43

60

17

23

A443

Quinoas

Fm

IIlogS

mS

42.8

13.9

3.9

13.1

1.7

0.8

1.9

0.8

4.2

0.3

0.3

2.2

14.2

68

20

11

52

20

28

88

48

A444

Quinoas

Fm

IIlogS

mS

30.0

11.8

5.8

11.3

3.4

1.6

3.2

0.8

50.5

08.4

18.2

57

24

19

41

24

35

77

716

Average

51

24

25

35

24

40

65

23

12

SD

12

411

114

10

17

19

7A442

Quinoas

Fm

IlogS

mS

21.8

29.1

3.8

16.0

2.6

06.4

1.7

4.1

0.9

04.1

9.6

60

23

17

26

23

51

93

07

A441

Quinoas

Fm

IlogS

mS

39

14.0

5.8

11.6

3.6

1.1

1.9

0.0

5.0

1.1

0.8

3.6

12.7

65

21

14

47

21

32

80

515

A412

Quinoas

Fm

IlogR

mS

22.5

6.9

5.1

14.4

2.9

3.2

4.4

0.2

7.6

0.5

014.7

17.6

44

29

27

34

29

37

70

16

14

A411

Quinoas

Fm

IlogR

mS

19.1

6.9

9.1

17.7

1.7

2.4

6.7

0.0

8.4

1.0

0.5

16.7

10.0

36

37

27

27

37

36

79

12

9A090

Quinoas

Fm

IlogH/I

mS

18.4

6.1

5.9

23.5

2.7

6.7

5.1

0.8

4.3

1.9

09.3

15.5

33

40

27

25

40

35

53

34

13

A100

Quinoas

Fm

IlogH/I

mS

16.3

11.9

4.2

22.2

2.9

8.6

2.9

0.9

7.9

4.0

1.1

7.3

10.1

36

34

30

21

34

45

63

28

9A103

Quinoas

Fm

IlogH/I

mS

18.4

9.6

4.3

12.0

12.8

1.9

9.3

0.8

3.2

1.3

0.5

6.1

20.0

39

22

39

25

23

52

47

747

A108

Quinoas

Fm

IlogH/I

mS

28.3

19.5

1.9

9.3

11.3

1.7

11.0

0.6

1.7

0.8

1.9

5.2

6.9

56

13

31

33

13

54

62

533

Average

46

27

27

30

28

43

68

13

18

SD

12

98

89

915

12

14

aRecalculateddetritalmodes

arereported

forQFL,Qm,monocrystallinequartz;Qp,polycrystallinequartz;Plag,plagioclase;Ls,sedim

entary

rock

fragments;Lv,volcanicrock

fragments;Lm,metam

orphic

rock

fragments;Lp,plutonic

rock

fragments.Q,F,L,Lt,Lvm,andLsm

areparam

etersbased

ontheclassificationofDickinson[1970]andGraham

etal.[1976].SD,standarddeviation.

bLocationisgiven

inFigures5and6.

cHeremSismiddle

sandstoneandcS

iscoarse

sandstone.


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bedded sandstones, climbing ripples and minor interbeddedmud silt components we interpret this association as typicalof sheet flood and distal braided environments [Miall, 1978].

4.2. Chacras Formation

[21] Separated locally by an erosional unconformity, theQuinoas Formation is overlain by the 24.2 ± 0.9 to 22.5 ±0.6 Ma Chacras Formation [Kraemer et al., 1999] that isdeposited to the west and east of the Sierra de Calalaste(Figures 5 and 6). At its maximum recorded thickness, thismember reaches �650 m. At the southern end of Sierra deCalalaste, the Chacras Formation is missing.[22] The Chacras Formation is dominated by laterally

continuous, massive, stratified and imbricated conglomer-ates (Gc, Gh) and lenticular conglomerates (Gt/p). Clastswithin the conglomerates are angular to subangular, andthere are frequently boulders throughout. These conglom-erates are interbedded with medium to coarse-grained cross-bedded, and planar-bedded sandstones (St, Sp), and massivesandstones (Sm) (Table 1). In the upper �150 m of log U(Figure 6b) large-scale (5–10 m) cross sets of medium-grained sandstones are preserved with foreset dips of 15–25� that reach a combined thickness of up to 50 m.[23] We interpret the lenticular conglomerates (Gt/p) as

resulting from deposition in shallow, gravely, bed loadchannels, with planar stratification and imbrication (Gc,Gh) occurring due to tractional flow at the channel bases[e.g., Nemec and Steel, 1984]. Planar bedded conglomeratesalso probably result from high-density flows during high-discharge events [Smith, 1987; Flint and Turner, 1988;Adelmann, 2001] and variations in the amount of accumu-lation [Nemec and Steel, 1984]. The interbedded sandstonesare deposited from high-density flows during waning con-ditions [Rasmussen, 2000]. Such a close spatial relationshipbetween lenticular and planar conglomerates, and interbed-ded sandstones is typical of deposition through changingflow regimes on a shallow gravel braided streams onan alluvial fan [Miall, 1996]. The subangular clasts andpresence of boulders indicate a proximal source forthese deposits. The large-scale, cross-bedded sandstonesare interpreted as eolian dunes and exhibit geometries thatare typical of modern day examples [e.g., Hunter, 1977;Reading, 1996], suggesting that at this time depositionoccurred in an arid environment. Paleocurrent data fromthese sediments, on the east side of the present-day Sierra deCalalaste, suggest a provenance mainly from the W-NW(Figure 3).

4.3. Potrero Grande Formation

[24] The lower to middle Miocene Potrero Grande For-mation unconformably overlies the Chacras Formation.There are a series of interbedded tuffs that occur withinthis formation, and the oldest tuff has a 40Ar/39Ar age of18.0 ± 0.9 Ma [Kraemer et al., 1999]. A maximumthickness of 300 m was estimated west of the Salar deAntofalla [Adelmann, 2001]. Adjacent to the Sierra deCalalaste, the Potrero Grande Formation consists of 150 mof conglomerates, conglomeratic sandstones, and sand-stones that are interpreted as having been deposited in anT

able

3.AFTDataa

Sam

ple

Lithology

Elevation,

mLatitude,

decim

aldegrees

Longitude,

decim

aldegrees

Number

ofXlsb

Rho-S,c

�105

NSd

Rho-I,c

�105

NId

P(c

)2,e%

Rho-D

,f

�105

NDg

Age,

Ma

±1s

U,

ppm

Dpar,

mmSD

SCAD9

metasedim

ents

3729

�26.1436

�67.4215

19

5.176

327

46.711

2951

88

12.493

5166

25.8

1.6

47

2.2

0.2

SCAD5

metasedim

ents

3978

�26.1618

�67.4819

64.206

57

36.232

491

82

12.821

5166

27.7

3.9

34

2.0

0.1

SCAD6

metasedim

ents

4248

�26.1356

�67.4516

93.196

67

31.678

664

62

12.739

5166

23.9

3.1

30

2.0

0.2

SCAD2

metasedim

ents

4328

�26.1356

�67.4527

20

4.519

300

38.528

2558

82

13.076

5166

28.5

1.8

36

2.1

0.1

SCAD4

metasedim

ents

4335

�26.1526

�67.4606

20

3.227

252

26.699

2085

100

12.903

5166

29.0

2.0

26

1.9

0.1

SCAD3

metasedim

ents

4455

�26.1501

�67.4609

20

2.500

175

22.843

1599

46

12.985

5166

26.5

2.2

23

2.0

0.1

aSam

plesareanalyzedwithaLeica

DMRM

microscopewithdrawingtubelocatedaboveadigitizingtabletandaKinetek

computer-controlled

stagedriven

bytheFTStageprogram

[Dumitru,1993].

Analysisisperform

edwithreflectedandtransm

ittedlightat1250Xmagnification.Sam

pleswereirradiatedatOregonStateUniversity.Sam

pleswereetched

in5.5

molarnitricacid

at21�C

for20s.Following

irradiation,themicaexternaldetectorswereetched

with21�C

,40%

hydrofluoricacid

for45min.Thepooledageisreported

forallsamplesas

they

passthec2test.Errorisonesigma,calculatedusingthezeta

calibrationmethod[H

urford

andGreen,1983]withzeta

of373.2

±6.1

forapatite(B.Carrapa).Dpar,fissiontracketch

pitmeasurements;SD,therelatedstandarddeviation.

bNumber

ofXlsisthenumber

ofindividual

crystalsdated.

cRho-S

andRho-Iarethespontaneousandinducedtrackdensity

measured,respectively(tracks/cm

2).

dNSandNIarethenumber

ofspontaneousandinducedtrackscounted,respectively.

eP(c

)2isthechi-squareprobability[G

albraith,1981;Green,1981].Values

greater

than

5%

areconsidered

topassthistestandrepresentasingle

populationofages.

f Rho-D

istheinducedtrackdensity

inexternal

detectoradjacentto

CN5dosimetry

glass

(tracks/cm

2).

gND

isthenumber

oftrackscountedin

determiningRho-D

.


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TC4011

alluvial fan environment [Kraemer et al., 1999]. Paleoflowindicators within this formation show large variations indirection with a scattered pattern (Figure 5) [Kraemer et al.,1999]. Here, we focus on the early evolution of the Sierra deCalalaste region. The Potrero Formation documents the laterevolution, and as such we do not document it in detail herebut refer the reader to Kraemer et al. [1999] for furtherdiscussion.

5. Provenance

[25] Sandstone petrography for members I and II of theQuinoas Formation show a lithic-feldspatic to feldspatic-lithic composition (Figure 7a) with an increase in quartzfragments, shown in the QFL and QmFLt diagrams, an upsequence typical of an evolution toward a more crystallinesource. The Qp, Lvm, Lsm diagram shows a generallyquartz rich composition for these sediments (Figure 7c).[26] The only two samples available for the Chacras

Formation have a lithic-feldspatic composition suggestinga different composition compared to the Quinoas sediments(Figures 7a and 7b). Despite the limited number ofsamples available, the greater contribution from volcanicand sedimentary-metamorphic rocks typical of the Sierra deCalalaste range (Figure 7c) expressed by these two samples,is consistent with paleocurrent data measured east of Sierrade Calalaste, indicating an eastward direction and in turn asource located in this range (Figure 3).

6. AFTT Data

[27] All analyzed fission track samples show comparableresults within one standard deviation, with ages rangingbetween 24 ± 3 and 29 ± 1 Ma (Figure 8 and Table 3). Onlylimited lengths were available in sample SCAD3 (Table 3),giving a mean length of 12.98 ± 0.56 mm and suggestingthat partial annealing was not significant. Etch pit measure-ments indicate that the samples are monocompositional,suggesting a homogeneous closure temperature. The age-elevation pattern of the Sierra de Calalaste vertical profilepoints to an exhumation rate of �0.3mm/yr, which isconsistent with rates obtained in neighboring ranges to thenorth [Deeken et al., 2004].

7. Discussion and Conclusions

[28] Our multiple data sets show that sedimentationcommenced during the late Eocene–early Oligocene withthe deposition of the Quinoas sediments in a partiallysegmented foreland basin with sediments derived mainlyfrom the west, and with contribution from proximalsources. The general trend in the petrographical data forthe Quinoas Formation indicates an increase in Qz upsequence (Figure 7). This could be explained by anevolution of the source toward more crystalline inputs.Crystalline sources are generally typical of areas to thewest (Figure 7c). Also, the timing of the deposition of theQuinoas formation corresponds to the end of the exhuma-

tion episode in the Domeyko Cordillera in northern Chile,composed of granite and granodiorite, between 50 and30 Ma [Maksaev and Zentilli, 2000]. This clear associationbetween the onset of exhumation of the Chilean Cordilleraand the deposition of the Quinoas Formation, and thesandstone provenance data would indicate that the domi-nant input to sedimentation was from the west. However,the presence of coarse grained clastic sediments with morevariable paleocurrent directions in member I of the QuinoasFormation would suggest that the foreland basin systemseen during Quinoas time was probably already partiallysegmented by proximal highlands. These areas provided alocal, and minor sediment source, e.g., plutonic bodieslocated immediately to the NW of the study area or evenpart of the Sierra de Calalaste to the east (Figure 3)[Kraemer et al., 1999]. Also, late Eocene–early Oligoceneapatite fission track cooling ages have been reported to theeast-southeast of the study area [Coutand et al., 2001]introducing the possibility that there may have been sometopography to the east at this time that was responsible forthe compartmentalization of the foreland. In this respectit is interesting that Horton et al. [2002] proposed thatthe Eastern Cordillera of southern Bolivia was a source ofthe Altiplano sediments during the Paleocene and Oligo-Miocene. The western flank of the Eastern Cordillera wasstrongly deformed during the paleogene by west vergentbackthrusts [McQuarrie and DeCelles, 2001; McQuarrie,2002b]. The Eastern Cordillera continues southward alongstrike into the Puna plateau, and this range is a candidatesource for the investigated sediments (Figure 7).[29] However, to date, no clear sedimentological and

thermochronological evidence exists that the Eastern Cor-dillera provided detritus to the west into the Puna at this time.Furthermore, the metamorphic and volcanic lithic fragmentsthat would be characteristic of the Sierra de Calalaste and ingeneral of more eastern and southeastern sources [e.g.,Reutter et al., 1994] are exceedingly scarce in the Quinoassandstone, suggesting that such sources, if present, were notcontributing to sandstone detritus. Also, recent AFTT ongranitic rocks from eastern and northeastern areas in thePuna suggest that the nearest ranges were exhuming laterthan the deposition of the Quinoas sediments by �20 Ma[Deeken et al., 2004]. In light of the presently available datawe cannot rule out the possibility that the eastern boundaryof the Puna might have constituted some topographichigh already during the late Eocene–early Oligocene [e.g.,Coutand et al., 2001; Horton et al., 2002].[30] Interestingly, during the late Eocene, the local gyp-

sum and anhydrite layers within the lower Quinoas Forma-tion (member I), and their association with evaporitic playamud flats, suggest that at least a seasonal arid environmentexisted. This arid environment may have resulted fromdeformation and uplift of the Chilean Cordillera [Maksaevand Zentilli, 2000], which might have shielded the southernPuna from occasional moisture incursions at a time whenthe cold Humboldt current had not been fully establishedand thus provided sufficient moisture to generate precipita-tion [Alpers and Brimhall, 1988]. In addition to regionalclimatic effects, the prolonged aridity seen in the Calalaste


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Figure 7. Main sandstone compositions of the investigated sediments: (a) QFL diagram based on thetechnique described by Dickinson [1970]. Q, monocrystalline and polycrystalline quartz; F, feldspar;L, sedimentary, metasedimentary, and volcanic lithic fragments including chert. Gray area represents thecompositional field of the late Eocene-Oligocene Potoco sedimentary rocks from the Altiplano sourcedfrom the Western Cordillera (sections 1–2; after Horton et al. [2002]). (b) QmFLt diagram based onthe technique described by Graham et al. [1976]. Qm, monocrystalline quartz; F, feldspar;Lt, sedimentary, metasedimentary, and volcanic, lithic fragments including polycrystalline quartz andchert. (c) QpLvmLsm diagram based on the technique described by Graham et al. [1976]. Qp,polycrystalline quartz; Lvm, volcanic lithics; Lsm, metamorphic lithics.


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area may also have been further supported by the potentialtopographic culminations of the Eastern Cordillera at leastby early Miocene time [Deeken et al., 2004].[31] At a broader scale, time-equivalent clastic sedimen-

tary rocks (Potoco Formation) deposited in the BolivianAltiplano and sourced from the Western Cordillera showvery similar facies associations in turn suggesting a similarsedimentary environment [Horton et al., 2002]. It isimportant to acknowledge however, that the Altiplanoprovenance data were obtained from localities hundredsof kilometers to the north of the study area. In any case,similarities between these units may indicate that anextensive longitudinally oriented basin, with local high-lands, existed during the late Eocene–early Oligocene thatspanned the length of the central Andes. In such ascenario, the Quinoas Formation may represent a part ofa semicontinuous foreland basin located to the east of theEocene Incaic mountain belt.[32] The overlying upper Oligocene– lower Miocene

Chacras Formation shows a change in sediment dispersaland facies association, with a marked coarsening comparedto the Quinoas Formation, and paleocurrent and petrographicdata that suggest input from the Sierra de Calalaste. Despitethe fact that paleocurrent measurements within the ChacrasFormation are only available along the eastern margin of therange, these data unequivocally indicate a source from theSierra de Calalaste.[33] AFTT data show that exhumation of the range

occurred between 24 and 29 Ma, which is the time ofdeposition of the late Quinoas and early Chacras formations(28.9 ± 0.8 Ma; 24.2 ± 0.9 [Kraemer et al., 1999]).Therefore we propose that the observed change in sedimentdispersal is a direct response to the uplift and erosion of theSierra de Calalaste range, which must have had a profoundeffect on the fluvial systems within adjacent basins duringthe Oligocene. Finally, the facies association with thepresence of extensive eolian dunes within this formation

suggests that the arid climate established prior to or duringQuinoas time persisted during Chacras time.[34] In Sierra de Calalaste and adjacent basins, uplift

resulting from Oligocene deformation led to reorganizationof the depositional systems within the present-day Salar deAntofalla area. Paleocurrent and sandstone provenance datashow that the depositional system, originally mainly sourcedfrom western crystalline rocks, was reorganized during theOligocene. This reorganization was contemporaneous withdeformation within the Sierra de Calalaste, suggesting acausal linkage between uplift of the range and response ofthe adjacent basin. At least transient internally drainedconditions existed during depositions of the Quinoas sedi-ments though it is less clear if such conditions were presentduring the deposition of the Chacras sediments. However,the occurrence of thick evaporite units in the late Miocene(Juncalito Formation [Kraemer et al., 1999]) indicates thatthe Salar de Antofalla basin was internally drained by thattime. Therefore we suggest that the basin reorganizationseen between the Quinoas and Chacras formations heraldedthe beginning of the process of disruption of the regionalfluvial system that may have ultimately led to the formationof internal drainage.[35] Many workers have documented deformation possi-

bly driving range uplift and a transition from externalto internal drainage within basins of the Puna plateau inOligo-Miocene time [e.g., Alonso, 1986; Jordan andAlonso, 1987; Marrett, 1990; Alonso et al., 1991; Coira etal., 1993; Vandervoort, 1993]. Around 23�S latitude, theSalinas Grandes and Tres Cruces basins contain evidence ofdeformation beginning in the late Eocene to early Oligocene[Coutand et al., 2001]. At 24�S latitude, internal drainage atSiete Curvas may have formed as early as the late Oligoceneand certainly by late Miocene time [Vandervoort et al.,1995], whereas to the west internal drainage within theArizaro and Tolar Grande basins commenced no later thanearly Miocene time [Donato, 1987; Coutand et al., 2001],

Figure 8. AFT vertical profile and length histogram related to sample SCAD 3.


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and early to middle Miocene [Vandervoort et al., 1995],respectively. Within the Salar de Pastos Grandes to the westof Siete Curvas, thick evaporites show that internal drainageformed sometime between 11.2 Ma and perhaps as early asthe late Eocene–early Oligocene [Alonso et al., 1991].Within the Salar de Hombre Muerto area at 25�S latitude,evaporite deposition started by 15.0 ± 1.2 Ma, and internaldrainage may have been established as early as the Oligo-cene [Alonso et al., 1991; Vandervoort et al., 1995]. In thesame region the termination of supergene copper mineral-ization indicates that hyperaridity was established after14.7 Ma [Alpers and Brimhall, 1988]. Finally, in the CampoArenal area along the present Puna margin at 27�S latitude,AFTT data indicate that deformation had begun between 29and 38 Ma [Coutand et al., 2001].[36] In summary, these observations suggest that defor-

mation, exhumation of basement ranges, and the establish-ment of internal drainage within the Puna plateau werespatially diachronous [e.g., Vandervoort, 1993; Coutand etal., 2001]. In particular, the timing of the onset of internaldrainage in the area may be dependent on the details of theuplift of discrete mountain ranges that may be controlled bylocal structural or volcanic conditions [Segerstrom andTurner, 1972; Alonso, 1986]. Combined with our new data

set, this suggests that deformation driving uplift within thePuna may not have occurred progressively from west-to-east as previously suggested [e.g., Andriessen and Reutter,1994] and that the establishment of internal drainage maynot only have been dependent on local details of marginalrange uplift but also on the history of the uplift of rangesfarther upwind.[37] At a more regional scale, our new data show that

deformation and uplift of the southern Puna plateau startedalready in Oligocene time, if not earlier, and contributed tothe fragmentation and paleodrainage reorganization of anearlier semicontinuous foreland basin. This event occurredin an already arid climate environment, and created the idealmorphotectonic preconditions for the establishment of thesubsequent internal drainage environment that was in placeno later than middle Miocene time.

[38] Acknowledgments. Deutsche Forschungsgemeinschaft (DFG)and the Alexander von Humboldt Foundation are kindly acknowledgedfor financial support through grants to M. Strecker and B. Carrapa,respectively. Peter DeCelles, Teresa Jordan, Brian Horton, and an anony-mous reviewer are kindly thanked for constructive reviews of this manu-script. Also, we thank R. Marrett and R. Alonso for their help duringsample collection and logistics.

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TC4011

oligocene range uplift and development of plateau...

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