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Quaternary Science Reviews xxx (2009) 1–14

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Quaternary Science Reviews

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Late local glacial maximum in the Central Altiplano triggered by cold andlocally-wet conditions during the paleolake Tauca episode (17–15 ka, Heinrich 1)

P.-H. Blard a,b,*, J. Lave b, K.A. Farley a, M. Fornari c,d, N. Jimenez e, V. Ramirez e

a Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA, USAb Centre de Recherches Petrographiques et Geochimiques, UPR 2300, CNRS, Nancy-Universite, Vandœuvre-les-Nancy, Francec Institut de Recherche pour le Developpement, Franced Geoazur, UMR 6526, CNRS, Universite de Nice, Nice, Francee Universidad Mayor de San Andres, La Paz, Bolivia

a r t i c l e i n f o

Article history:Received 8 July 2009Received in revised form18 September 2009Accepted 28 September 2009

* Corresponding author. Centre de Recherches PetroUPR 2300, CNRS, Nancy-Universite, Vandœuvre-les-Na42 23.

E-mail address: [email protected] (P.-H. Bla

0277-3791/$ – see front matter � 2009 Elsevier Ltd.doi:10.1016/j.quascirev.2009.09.025

Please cite this article in press as: Blard, P.-Quaternary Science Reviews (2009), doi:10.1

a b s t r a c t

The timing and causes of the last deglaciation in the southern tropical Andes is poorly known. In theCentral Altiplano, recent studies have focused on whether this tropical highland was deglaciated before,synchronously or after the global last glacial maximum (w21 ka BP). In this study we present a newchronology based on cosmogenic 3He (3Hec) dating of moraines on Cerro Tunupa, a volcano that islocated in the centre of the now vanished Lake Tauca (19.9�S, 67.6�W). These new 3Hec ages suggest thatthe Tunupa glaciers remained close to their maximum extent until 15 ka BP, synchronous with the LakeTauca highstand (17–15 ka BP). Glacial retreat and the demise of Lake Tauca seem to have occurredrapidly and synchronously, within dating uncertainties, at w15 ka BP. We took advantage of thesynchronism of these events to combine a glacier model with a lake model in order to reconstructprecipitation and temperature during the Lake Tauca highstand. This new approach indicates that, duringthe Tauca highstand (17–15 ka BP), the centre of the Altiplano was characterized by temperature w6.5 �Ccooler and average precipitation higher by a factor ranging between �1.6 and �3 compared to thepresent. Cold and wet conditions thus persisted in a significant part of the southern tropical Andesduring the Heinrich 1 event (17–15 ka BP). This study also demonstrates the extent to which thesnowline of glaciers can be affected by local climatic conditions and emphasizes that efforts to drawglobal climate inferences from glacial extents must also consider local moisture conditions.

� 2009 Elsevier Ltd. All rights reserved.

1. Introduction

The Tropics are recognized to play a major role in paleoclimaticcycles (Stocker et al., 2001) because this area may potentiallygenerate forcing (i) at the global scale, e.g. through the tight linkbetween tropical sea-surface temperature (SST) and atmosphericCO2 and CH4 concentrations (Lea, 2004), and (ii) at the regionalscale, e.g. through the feedbacks between moisture transport andthe thermohaline circulation (Leduc et al., 2007). However, ourunderstanding of the exact mechanisms involved in these inter-actions is limited by the lack of precise chronologies in several keytropical areas (Hastenrath, 2009). This is particularly true for thehighlands of the continental realm, where the past extents of

graphiques et Geochimiques,ncy, France. Tel.: þ33 3 83 59

rd).

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H., et al., Late local glacial m016/j.quascirev.2009.09.025

mountain glaciers can be used to reconstruct with precisionchanges in temperature and precipitation (e.g. Hostetler and Clark,2000). The Altiplano is a key area within this tropical puzzle: thishigh plateau is located at the interface between tropical andmid-latitude atmospheric circulation (Kull et al., 2008) and so islikely to have experienced substantial changes in precipitation andtemperature during the last deglaciation (20–10 ka), and particu-larly during the 17.5–14.5 ka period termed ‘‘Mystery Interval’’ by(Denton et al., 2006).

Although the presence of ancient mountain glaciers is reportedin several places on the Altiplano (Smith et al., 2008; Zech et al.,2008), the spatial pattern and the timing of paleosnowline changesremain challenging issues (Rodbell et al., 2009). A discussionbetween Clark (2002) and Seltzer et al. (2002) focused on therelative timing between the global last glacial maximum (LGM,21 ka) and the late Pleistocene glaciation in the tropical Andes.Cosmogenic 10Be (10Bec) ages from moraines located in thenorthern part of the Altiplano suggest that the last glacialmaximum occurred as early as w25 ka BP in the Tropical Andes

aximum in the Central Altiplano triggered by cold and locally-wet...,

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P.-H. Blard et al. / Quaternary Science Reviews xxx (2009) 1–142

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(Smith et al., 2005b, 2008). However, other Altiplano locationsyielded younger local LGM ages of w15 ka (Seltzer, 1992; Claytonand Clapperton, 1997; Zech et al., 2007). This contrasting behaviorsuggests that local effects, such as asynchronous changes in theprecipitation field, were superimposed on the regional temperaturetrend that affected the tropical Andes. Indeed glacier fluctuationscan result either from changes in precipitation or temperature(Ohmura et al., 1992). This ambiguity makes it difficult to interpretlocal glacier advances as a global temperature signal in the absenceof an independent estimate of local precipitation conditions.

Paleolakes may be very useful proxies in that, when combinedwith glacier-based paleoclimatic reconstructions, they can be usedto estimate the relative effects of precipitation and temperature.Although the Altiplano is now dry and covered by large saltydeserts, large palaeolakes were developed on this plateau at varioustimes in the late Pleistocene period (Minchin, 1882; Servant andFontes, 1978; Seltzer, 1994; Sylvestre et al., 1999). The chronology ofthe last and highest lake level stage, known as the Tauca phase(w3770 m) is now relatively well-established between w17 and15 ka thanks to 14C and 230Th/234U dating of lacustrine formations(Sylvestre et al., 1999; Placzek et al., 2006). Past occurrence ofglaciers in areas that are now ice-free is documented by morainesdeposited on the slopes surrounding the palaeolake Tauca. Therelationship between the local last glacial maximum and the Taucahas been debated for a long time, and at least two scenarios havebeen proposed:

i) The glaciers receded from their maximum position before therise of the Lake Tauca (Servant and Fontes, 1978), and,according to these authors, the lake was filled by the glaciersmeltwater. However this idea was later refuted by budgetcalculations showing that the total volume of LGM mountainglacier ice was insufficient to supply the deep Tauca paleolake(Hastenrath and Kutzbach, 1985; Blodgett et al., 1997).

ii) The glaciers reached their maximum extent - or significantlyreadvanced – synchronously with the Tauca episode between17 and 15 ka BP (Seltzer, 1992; Clayton and Clapperton, 1997;Clapperton et al., 1997; Zech et al., 2007). In this case, bothevents could result from the same climate forcing.

In order to define the relative chronology of ice advance and lakehighstand, and to improve our understanding of the paleoclimaticconditions on the Altiplano at the end of the Pleistocene, wepresent here new 3Hec exposure ages measured on pyroxenesphenocrysts sampled on glacial landforms on an andesitic volcanolocated at the centre of palaeolake Tauca. After having confirmedsynchronism between the Central Altiplano local LGM and theTauca highstand (17 and 15 ka BP), we investigate a joint inversionof the paleoclimatic conditions (precipitation, temperature) pre-vailing during those glacial times by using numerical models ofpaleoglacier extent and paleolake level.

2. Geomorphological and (paleo)-climatic settings of thecentral Altiplano

The Altiplano plateau (14–22�S) is a large (w500,000 km2), flat,high-elevation area (mean elevation range between 3600 and3800 m) located in the tropical Central Andes. Precipitation mainlyoriginates from humid air masses brought by the tropical easterliesfrom the Tropical Atlantic. Indeed, although the Altiplano is alsounder the influence of westerly winds from the Central Pacific,clouds coming from the northeast are currently responsible forw80% of the small annual rainfall on the plateau (Vuille, 1999). Itpresents a highly seasonal pattern, with 80% falling from October toApril (Austral summer), as a consequence of the southward

Please cite this article in press as: Blard, P.-H., et al., Late local glacial mQuaternary Science Reviews (2009), doi:10.1016/j.quascirev.2009.09.025

deflection of the Inter Tropical Convergence Zone (ITCZ) during theso-called ‘‘South American Monsoon season’’ (Fig. 1) (Garreaudet al., 2009). As a consequence of orographic effects, there is alsoa strong northeast to southwest gradient in annual precipitation:rainfall decreases from 800 mm a�1 in the northern part of theBolivian Altiplano to less than 100 mm a�1 in the Lipez area insouthern Bolivia (Fig. 1). The endorheic watershed of the centralAltiplano includes two subcatchments: the Titicaca watershed inthe north and the Salar de Uyuni catchment in the south. Howeverthese catchments are hydrologically connected through the RioDesaguadero, the outlet of Lake Titicaca (Fig. 1).

Due to low precipitation (<1000 mm a�1) and mean annualpotential evaporation ranging between 1000 and 2000 mm a�1

(Montes de Oca, 1989), the present moisture budget of the Altiplanois in deficit. This results in shallow oligosaline lakes (e.g. LakePoopo) and large saline basins (Salar de Uyuni and Salar de Coipasa)(Fig. 1). However, evidence for several episodes of deep LatePleistocene lakes (Seltzer, 1994; Sylvestre et al., 1999; Placzek et al.,2006) attest to the fact that Altiplano hydrographic conditions weresignificantly different in the past (Argollo and Mourguiart, 2000;Placzek et al., 2006). These features document a high sensitivity toregional moisture and temperature conditions that have variedthrough time. Previous studies proposed reconstructions of thepaleoclimatic conditions that triggered and maintained the paleo-lake Tauca highstand (Hastenrath and Kutzbach, 1985; Blodgettet al., 1997). Using evaporation models at steady-state, theseauthors concluded that Lake Tauca resulted from rainfall rangingbetween a few percent and 80% greater than present mean condi-tions. However these reconstructions suffered from the fact that airtemperature was not independently constrained during the Taucaphase (w15 ka). Moreover part of the variability also arises fromthe calibration of the model parameters (Blodgett et al., 1997).Therefore, the present study takes advantage (see Section 4.2) ofthe hydrological model of (Condom et al., 2004), which offers theadvantage of having been parameterized using the present-dayconditions of the Titicaca watershed.

Under present-day conditions, the glacier equilibrium-linealtitude (ELA) is at w5200 m in the northern part of the Altiplano(Ribstein et al., 1995), and rises south-westward up to w6000 m inthe Lipez area, the extreme southern part of the plateau (Klein et al.,1995). This regional ELA trend parallels the present negativeprecipitation gradient from northeast to southwest, suggestinga strong influence of the regional snowfall pattern (Ammann et al.,2001). The paleo-ELA reconstructed from several morainessuggests that during the last glacial maximum snowlines rangedbetween 800 and 1200 m lower than today (Clayton and Clapper-ton, 1997; Dornbusch, 2000). However, due to limited dating, itremains unclear whether this snowline depression was synchro-nous at the regional scale of the Altiplano. Importantly, the influ-ence of temperature reduction and precipitation increase on thespatial variations in ELA depression are not yet well-established(Clayton and Clapperton, 1997; Kull et al., 2008).

3. Cosmogenic 3He glacial chronology of Cerro Tunupa

3.1. Glacial features of Cerro Tunupa

Cerro Tunupa (19.86�S, 67.61�W) is an inactive volcano locatedin the centre of the now vanished paleolake Tauca (Fig. 1). It ismainly built of Quaternary andesite lavas that erupted less than5 Ma (Villeneuve et al., 2002). The flanks of the Tunupa volcanoexhibit several landforms and glaciogenic deposits testifying to thepresence of ancient glaciers (Fig. 1). These glacial features areremarkably well-preserved, suggesting that erosion rates are notsignificant, thanks to the low precipitation in this area (w200 mm


Fig. 1. Geomorphological and climate setting of the Central Altiplano. A – Present climate setting of the Tropical Andes. B – Altiplano map Draped on the Landsat image of theAltiplano: Lake Tauca paleoshorelines (3770 m) (dashed blue line), the limit of the endoreic Tauca watershed (orange dashed line), and the isohyets curves (white lines) of presentannual precipitation (New et al., 2002). ZM: Zongo–Milluni area where moraines have been dated by 10Be in Smith et al. (2005b). HL: Huara Loma – Rio Suturi area dated by 10Be inZech et al. (2007). CA: Cerro Azanaques dated by 14C in Clayton and Clapperton (1997). Sajama, 6542 m, is the highest peak in Bolivia.

P.-H. Blard et al. / Quaternary Science Reviews xxx (2009) 1–14 3

ARTICLE IN PRESS

per year) (New et al., 2002). One of the largest glacial landforms isdeveloped on the south-east flank of Tunupa, in the Chalchalavalley (Fig. 2) (Clapperton et al., 1997). Detailed mapping of thisvalley was achieved through field observations and analyses ofaerial photos and satellite images (Fig. 2).

This glacial valley was initially connected to an old moraineoutlet (M0), which has the shape of crab pincers and is orientedsouthward. A prominent sharp-crested moraine, M1 crosscut thisold moraine M0, and presently delimits the former glaciated U-shaped Chalchala valley with w100 m-high-shoulders. This mainmoraine M1 appears to be genetically connected to a distal


glacifluvial fan delta Mf-1 (Fig. 2). Two distinct lobes can be iden-tified on this fan: the older lobe, Mf-1a, is overlapped by themoraine M1, and has two major erosional benches whose altitudes(3760 and 3770 m) suggest that they result from the Tauca paleo-lake highstands that occurred between 17 and 15 ka (Sylvestreet al., 1999; Placzek et al., 2006). These strandlines are conspicu-ously developed on both sides of the fan Mf-1. The younger fan,Mf-1b, lies in the axis of the former glacier tongue that formed M1moraines. In contrast with Mf-1a, Mf-1b does not record a trace ofthe highest shoreline (3770 m), but only a depositional bench incontinuity with the lower erosional strandline (3660 m) (Fig. 2).


Fig. 2. A – Mapping of the Cerro Tunupa glacial deposits, with the sample locations and 3Hec ages. B – Picture of boulder TU1B sampled on the M1 moraine. C – Picture of boulderTU5A sampled on the Mf-1b fluvio-glacial fan delta.


ARTICLE IN PRESS

These stratigraphic relationships therefore suggest that: 1. the earlyM1 stage of the glacier advance, which led to the frontal depositionof the glacifluvial material of Mf-1a, is antecedent to the highestlake stand; 2. the maximum glacial advance and M1 construction,accompanied by frontal deposition of the glacifluvial material ofMf-1b, post-dates the highest lake stand but is coeval with thefollowing high stand (10 m below). This scenario is also consistentwith sedimentologic observations on these fan deltas in the Chal-chala valley outlet and in the nearby Pocolli valley, indicating thatMf-1 is cross-stratified with the 3760 m and 3770 m shorelinedeposits of the Tauca paleolake (Clayton and Clapperton, 1997;Clapperton et al., 1997) (Fig. 2).

If this interpretation is correct, the deposition age of Mf-1b fancan thus be bracketed between 16.5 and 15 ka and considered asa reference site for testing the accuracy of the used 3Hec productionrates.

Moreover, if the Mf-1 fan is the distal equivalent of the laststages of the moraine M1 building (or M2), all these observationsprovide strong support for a tight temporal coincidence betweenthe termination of moraine M1 (or M2) and the Tauca highstand,dated between 17 and 15 ka (Placzek et al., 2006).

Just after the building episode of the M1 moraine, the initialstage of glacier retreat built the lateral moraine, M2, which isinterlocked within the inner part of the lateral moraine M1. Thesubsequent glacier retreat left the bottom of the Chalchala valleyfree of debris deposits and exposed numerous well-preservedstriated rocks. Two elongated moraines M3 delineate a later andshort duration re-advance of a very narrow glacier tongue thatdeveloped above these striated rocks and results from an over-spilling originating from the upper glacier cirque. The smallmoraine M3 is thus stratigraphically younger than the highestroches moutonnees of the valley at w4500 m.

Here we use cosmogenic 3He dating to complement theseobservations with an absolute chronology of these glacial features.For this purpose, we sampled several meter-sized boulders on themoraines M0, M1, Mf-1 and M3 (Fig. 2 and Table 1), paying specialattention to select the top of the most prominent boulders located


at the top of moraine crests (Fig. 2). We also collected severalstriated rocks (TU2 and TU4 samples) at the flat bottom of theChalchala valley. The well-preserved striations suggest that thesesurfaces have suffered limited erosion. Moreover, the sampling wasdone in the upper part of prominent knobs or whale-backs tominimize the probability of temporary shielding by till or sedimentcover. Finally, the deep downcutting of the glacier between thephases M0 and M1 ensures that glacial erosion has been sufficientlyintense to fully reset the striated rocks, as demonstrated in regionsthat experienced extensive glacial erosion (e.g. Gayer et al., 2006).

3.2. Cosmogenic 3He dating of the Tunupa glacial landforms

3.2.1. MethodsCosmogenic 3He data and calculated exposure ages are given in

Table 1.These data were obtained from samples bearing millimeter-size

pyroxenes phenocrysts. This mineral is well-suited for cosmogenic3He dating because it has high helium retentivity (Trull et al., 1991;Trull and Kurz, 1993; Blard et al., 2008). Empirical calibration fromindependently dated surfaces established its production rate of 3Hebetween 128 and 142 at g�1 a�1 at sea-level high latitude,depending on the used scaling factor (Ackert et al., 2003; Blardet al., 2006; Amidon et al., 2009). Both helium isotopes weremeasured at Caltech with a MAP 215-50 mass spectrometer. Toestimate the magmatic helium component, several aliquots werefirst crushed in vacuum and the extracted gas analyzed (Supp. Table1). All pyroxene samples were then fused in a vacuum furnace tocompletely release and analyze the cosmogenic 3He (Table 1) andthe radiogenic 4He* components (Supp. Table 2). Measurement,correction for magmatic and nucleogenic 3He and calculations ofexposure ages are detailed in Appendix A. The total correction fornon-cosmogenic 3He is less than 5% for all samples and does notrepresent a major source of uncertainty (see Appendix A), as sup-ported by the excellent agreement between different replicates,including those which vary in Li content (the source of nucleogenic3He and cosmogenic thermal-neutron production; see Table 1).


Table 1Description of the Tunupa moraine samples and cosmogenic 3He ages.

Sample Object Boulderheight(cm)

Altitude(m)

Latitude(�S)

Longitude(�W)

Mineral Size(mm)

Mass(mg)

4Het

(1012 at g�1)

3Het

(107 at g�1)Li(ppm)

3Hen

(105 at g�1)aCTN P3

(at g�1 a�1)b

3Hec

(107 at g�1)cDepthcorrectiond

Paleomagneticcorrectione

(applied tothe Stone factor)

Exposure age (ka) ELA fromnumericalmodeling

Lifton et al.,2005

Stone,2000;Dunai,2001

Tunupa Moraine M0TU7A Boulder 200 3828 19.8816 67.5971 Green

diopside0.2–0.5 11.9 2.95� 0.1 17.4� 0.4 18 1.5� 0.2 5 17.3� 0.4 0.98 1.25 115� 4

(121� 4)142� 6(150� 6)

4200

TU7B Boulder 100 3829 19.8813 67.5973 Greendiopside

0.2–0.5 18.7 2.73� 0.05 18.7� 0.5 10 0.7� 0.1 3 18.6� 0.5 0.95 1.25 127� 4(135� 5)

160� 6(170� 6)

4200

TU7C Boulder 150 3829 19.8813 67.5973 Greendiopside

0.3–0.5 35.1 3.00� 0.07 20.2� 0.5 31 3.7� 0.4 8 20.1� 0.5 0.97 1.25 129� 5(138� 5)

163� 7(175� 7)

4200

Fluvio-glacial Mf1a.TU5A Boulder 100 3760 19.8742 67.5880 Green

diopside0.2–0.5 257.4 3.26� 0.03 1.59� 0.06 19 2.3� 0.2 5 1.55� 0.05 0.98 0.98 Average Average

TU5A Blackaugite

0.3–0.5 41.4 9.55� 0.06 1.50� 0.05 24 2.9� 0.3 6 1.45� 0.05 0.98 .97 13.8� 0.5 16.2� 0.6 4380


0.2–0.5 258.9 4.03� 0.04 1.51� 0.05 36 4.3� 0.4 9 1.45� 0.06 0.96 0.97 13.3� 0.5 15.6� 0.6 4380

Mf1bTU6 Boulder 250 3810 19.8789 67.5955 Green

diopside0.2–0.5 183.9 4.50� 0.04 2.94� 0.09 15 1.7� 0.2 4 2.9� 0.09 0.98 1.12 24� 0.9 27� 1.1 4380

TU8 Boulder 200 3778 19.9632 67.5941 Greendiopside

0.2–0.5 318.8 4.68� 0.04 1.90� 0.05 60 6.4� 0.6 1.82� 0.05

TU8fa Greendiopside

0.2–0.3 178.2 5.16� 0.05 1.91� 0.08 60 6.4� 0.6 1.83� 0.08

TU8fb Greendiopside

0.2–0.3 189.3 5.26� 0.05 1.90� 0.07 60 6.4� 0.6 1.82� 0.08

TU8 Greendiopside

0.2–0.3 79.0 5.32� 0.1 1.93� 0.06 60 6.4� 0.6 1.85� 0.06

TU8G Greendiopside

0.3–0.5 183.2 3.90� 0.04 1.92� 0.08 60 6.4� 0.6 1.84� 0.08 Average Average

15 0.97 1.01 13.4� 0.5 17.8� 0.8 4380Moraine M1TU1A Boulder 250 4260 19.8643 67.6128 Green

diopside0.2–0.5 62.2 3.52� 0.08 2.00� 0.06 15 1.8� 0.2 4 1.96� 0.06 0.97 0.98 14.1� 0.5 17.2� 0.7 4380


0.2–0.5 191.5 3.89� 0.03 2.21� 0.07 23 2.8� 0.3 6 2.16� 0.07 0.98 1.00 14.8� 0.5 17.9� 0.7 4380

TU1C Boulder 120 4206 19.8652 67.6099 Greendiopside

0.2–0.5 377.0 3.67� 0.03 2.84� 0.06 20 1.9� 0.2 2.8� 0.06

TU1Cf Greendiopside

0.2–0.3 159.1 4.29� 0.04 3.05� 0.1 20 1.9� 0.2 3.02� 0.1

TU1CG Greendiopside

0.3–0.5 196.9 3.42� 0.03 2.96� 0.08 20 1.9� 0.2 2.93� 0.08 Average Average

5 0.99 1.06 19.4� 0.7 23.2� 0.9 4380TU1D Boulder 300 4083 19.8691 67.6054 Green

diopside0.2–0.5 42.9 2.67� 0.09 1.76� 0.05 9 1.1� 0.1 2 1.74� 0.05 0.99 0.97 13.7.0� 0.5 16.5� 0.7 4380

TU4 Striatedrock/rochemoutonnee

4260 19.8570 67.6180 Greendiopside

0.2–0.5 236.6 3.09� 0.03 1.85� 0.06 20 2.4� 0.2 5 1.81� 0.06 0.99 0.96 12.8� 0.5 15.7� 0.6 4550

TU2 Striatedrock/rochemoutonnee

4450 19.8455 67.6272 Greendiopside

0.3–0.5 36.0 2.53� 0.1 2.01� 0.06 16 2� 0.2 4 1.98� 0.06 0.99 0.98 12.7� 0.5 15.6� 0.6 4720

(continued on next page)

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Please cite this article in press as: Blard, P.-H., et al., Late local glaciQuaternary Science Reviews (2009), doi:10.1016/j.quascirev.2009.09

al m.025

Because the Altiplano is located in the Tropics, the main source ofuncertainty in the moraine ages arises from the scaling factor usedto calculate the local production rate. We thus computed 3Hec agesusing two different scaling procedures: the one of (Lifton et al.,2005) and the one of (Stone, 2000) implemented with the time-integrated correction of (Dunai, 2001). Correction for topographicshielding and snow cover are negligible. Zero-erosion assumptionis discussed in Appendix A.

3.2.2. 3Hec dating results and choice of the scaling factorThe cosmogenic 3He ages are presented in Table 1. They range

between 115� 4 ka and 163�7 ka for moraine M0, between12.7� 0.5 and 27�1.1 ka for moraine M1 and striated rocks, andbetween 9.8� 0.4 and 13.2� 0.6 ka for moraine M3, depending onthe scaling factor used (Table 1). The error ages, given at the 1slevel, include analytical uncertainties only.

The scaling of 3Hec production rate is potentially an importantsource of uncertainty for cosmogenic dating at low latitudes andhigh elevation (Balco et al., 2008). In the present study, we canhowever rely on the stratigraphic relationships existing betweenthe distal glacial deposits (lobe Mf-1) and the lake shorelines.Indeed the boulders TU5A and TU5B sit on the fan Mf-1b whosedeposition is probably synchronous (Fig. 2 and Section 3.1 fordetails) with the lake highstand (>3760 m) dated between w17and 15 ka BP (Sylvestre et al., 1999; Placzek et al., 2006). The scalingfactors of (Lifton et al., 2005) yield 3Hec ages that are younger by atleast w1.5 ka compared to the lake highstand episode. On thisbasis, we conclude that this correction factor is probably notadequate for this period and this region (Supp. Fig. 1). We conse-quently decided to consider only the ages calculated with the(Stone, 2000; Dunai, 2001) polynomial to discuss the Tunupaglacier chronology and its climatic implications.

3.2.3. The glacial chronology of Cerro Tunupa – synchronismbetween the paleolake Tauca and the local last glacial maximum

The 3Hec concentrations measured in the three boulders TU7A, Band C (M0 moraine) indicate that the deposition of the M0 moraineended between 142� 6 and 163�7 ka under the assumption of noerosion, and between 150� 6 and 175�7 ka if a maximum erosionrate of 0.4� 0.1 m Ma�1 (Smith et al., 2005a) is considered. Therelatively good cluster of these three ages suggests that pre- andpost-depositional processes have not significantly disturbed theexposure history of this old moraine (Smith et al., 2005a). This oldglacial stage recorded on Cerro Tunupa is consistent with theglacier re-advance observed by (Smith et al., 2005a) between 170and 125 ka in Peru. At a broader scale, there is also a good temporalcoincidence between this glacier stillstand and the penultimateglaciation of Oxygen Isotope Stage 6, between 140 and 170 ka (Petitet al., 1999).

The exposure ages measured on Mf-1, M1 and M3 rangebetween 27�1.1 and 12.3� 0.5 ka by using the (Stone, 2000;Dunai, 2001) scaling. These 3Hec ages along with our field obser-vations allow the following chronology to be proposed:

� The 3Hec date of 27�1.1 ka from the boulder TU6 providesa maximum age for the first pulse of the glacier tongue that ledto the deposition of the Mf-1 fluvio-glacial fan. TU6 is locatedon the fluvio-glacial deposit Mf-1, clearly above the highestshoreline, so there is little probability that the exposure age ofTU6 has been affected by post-depositional processes, such asexhumation due to wave erosion at the lake shoreline. If this isthe correct interpretation of this date, such an early glacieradvance would be in agreement with the 10Be ages of w25 kaobtained by (Smith et al., 2005b) in the northern part of theAltiplano. However, inheritance of 3Hec cannot be excluded for


5 10 15 20 250 303640

3680

3740

3780

3720

3760

3700

3660

Elev

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lake

sur

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Lake chronology (ka BP)

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4400

4500

4600

4700

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3Hec chronology of the Tunupa glacial deposits (ka BP)

Mod

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A (m

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YDBA

Striated rockBoulder

Data by Placzek et al., 2006

This study

U-Th dating14C dating

Lake Tauca

Lake Sajsi

Lake Coipasa

Dry periodDry period

H1

A

B

Fig. 3. A – 3Hec glacial chronology of Cerro Tunupa. Scaling factors are computed usinga combination of the scaling corrections of (Stone, 2000) and (Dunai, 2001) and a SLHL3Hec production rate of 128 at g�1 a�1 (Blard et al., 2006). The ELA is modeled by usingthe approach described in Section 4.1. H1 refers to the Heinrich 1 event, BA to theBølling–Allerød and YD to the Younger Dryas. B – Lake level chronology from (Placzeket al., 2006).


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TU6, and w27 ka should thus be considered as a probablemaximum age for the deposition of the Mf-1a formation.� All the other boulders belonging to the main moraine M1 or the

fluvio-glacial deposit Mf-1 are characterized by 3Hec agesranging between 15.6� 0.6 and 17.9� 0.7 ka, with the excep-tion of TU1C, from moraine M1, whose high age (23.2� 0.9 ka)suggests this boulder is an outlier affected by inheritance.Despite the fact that these absolute ages are dependent on thechoice of scaling factor (see discussion in Appendix A), theirrange suggest that the peak of the local last glacial maximumoccurred between 18 and 15 ka. This timing is not significantlydifferent from a recent LLGM dating in southern Perou(Bromley et al., 2009), while the LLGM in the Cordillera Realoccurred several thousand years before (w25 ka) (Smith et al.,2005b). Importantly, these new 3He ages from Cerro Tunupastrongly suggest that the glacier remained close to itsmaximum position until w15 ka. Such scenario is consistentwith the conclusions of (Clayton and Clapperton, 1997; Clap-perton et al., 1997) and our morphostratigraphic observationssuggesting that the glacial stage that produced the M1 morainepersisted during the w2 ka duration of the paleolake Taucahighstand between 17 and 15 ka (Fig. 3) (Sylvestre et al., 1999;Placzek et al., 2006).� Striated rocks TU2 (4450 m) and TU4 (4260 m) from roches

moutonnees yielded 3Hec dates of 15.6� 0.6 and 15.7� 0.6 ka,


respectively. These ages are additional support for thehypothesis of a late (w15 ka) retreat of the Cerro Tunupaglaciers from their maximum position. Moreover, this glacialretreat exhibits a striking similitude with the drop of the Taucawater level at w15 ka. Although uncertainty in the cosmogenicdates does not permit us to conclude which event occurred first(ice retreat or lake level drop), the timing and amplitude of thetwo abrupt events are very similar. This suggests either a causallink between these two retreats, or that both lake and moun-tain glaciers responded almost simultaneously to an externalforcing due to an abrupt regional climatic change affecting theCentral Altiplano.� 3Hec ages of the boulders belonging to the small moraine M3

(TU3A and TU3B) indicate a glacial re-advance between12.3� 0.5 and 13.2� 0.6 ka. Although limited in amplitude,this stillstand seems to be synchronous with the Coipasa wetphase corresponding to a new rise of the lake level at 3700 mbetween 13 and 12 ka (Fig. 3). Moreover, such timing is alsocompatible, within uncertainty, with the Younger Dryas (YD)cold event that occurred between 12.9 and 11.6 ka in theNorthern Hemisphere (Andersen et al., 2004). Our data are thusin agreement with previous 10Be ages in the Cordillera Blancaby (Farber et al., 2005) supporting the hypothesis of a YDsignature in the South tropical Andes. However, because of thedating uncertainty and the limited number of samples, it is notpossible to exclude the possibility that this re-advance wassynchronous with the Antarctic cold reversal event, dated atw14 ka (Jouzel et al., 1995). Additional ages and calibrationswill thus be useful to address the question of a YD occurrencein the Central Altiplano.

The climatic implications of this glacial chronology are exam-ined in the next section by using numerical modeling of paleo-climatic conditions, with a special emphasis on the synchronismbetween the highstand of paleolake Tauca and the local last glacialmaximum between 17 and 15 ka.

4. Numerical modeling of paleoclimatic conditions(precipitation, temperature) during the paleolake Taucaepisode (17–15 ka)

4.1. Modeling of the ELA depression

4.1.1. Model descriptionIn order to interpret the observed Tunupa glacier extents in

terms of paleoclimatic conditions (temperature and precipitation),we used a numerical approach that combines an ice mass-balancemodel (Blard et al., 2007) with a simplified 2-D ice flow model(Harper and Humphrey, 2003). The computation of mass-balancerelies on a positive-degree-month model in which ablation isproportional to temperature and shortwave solar radiation (Hock,1999). Different ablation parameters were used for snow and ice inorder to include the influence of albedo on melting. This empiricalmodel does not directly include the effects of wind and relativehumidity. However, the adopted melting factors have been cali-brated using the mass-balance and climatic dataset monitoredmonthly on the Zongo glacier (16�S) over the 2003–2005 period(Blard et al., 2007). Given the proximity of these locations, and thefact that the modeled altitude range for these glaciers overlaps, it isreasonable to conclude that they are characterized by comparablemelting processes, so that the reconstructed melting rates are notsignificantly biased by their calibration on the Zongo glacier.Moreover, in order to be conservative with the uncertaintiesattached to our reconstruction, we considered the maximum range(from �0.7 to �1.3) reported for these melting parameters in the


ELA = 4380 m

N

0

50

100

150

Ice th

ickn

ess (m

)

Glacier mass balance (mwater.a-1)

Elev

atio

n (m

)

-10 -8 -6 -4 -2 0 2

4000

4400

4800

5200

5600

6000

3600

Present-day ELA = 5680 m

Palaeo ELA (~17-15 ka) = 4380 m

Tunupa summit

Elev

atio

n (m

)

Eastern distance (km) Southern distance (km)

M1

~17-15 ka glaciers

M1: Local LGM moraine

Mf-1b: Fluvio-glacial deposit

Mf-1b

Tauca Lake 3770 m

~17-15 ka

ΔT = -6.5°C

xP = × 3 = 600 mm.a-1

5200

4400

4800

4000

3600

42

0

68

10 8 6 4 20

BA

Fig. 4. A – Numerical modeling of the local last glacial maximum during the 17–15 ka period. B – Modeled mass balance vs elevation for the present (orange curve, ELA¼ 5680 m)and the LLGM (blue curve, ELA¼ 4380 m). See Section 4.1 and (Blard et al., 2007) for details.


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tropical area (Hock, 2003). The input data used in the model aremonthly temperature and precipitation (see Supp. Table 3), directsolar radiation during the 17–15 ka period (Paillard et al., 1996;Laskar et al., 2004), and a digital elevation model of the South-Eastflank of the Tunupa volcano. A local lapse-rate ofw6.5� 0.2 �C km�1 was assumed (Klein et al., 1999). The presentEquilibrium-Line Altitude (ELA) modeled with the present-dayconditions stands at 5680�100 m. Cerro Tunupa, 5400 m high, isnot covered by permanent ice now and cannot be used to check themodel, but this calculated ELA is consistent with the snowlinecurrently observed on the rare glaciated Altiplano summits, such asSajama (18�060S–68�520W) (Fig. 1) and Parinacota volcano(18�100S–69�080W) (Ammann et al., 2001; Hastenrath, 2009), thatare under precipitation and temperature conditions similar to thoseof Tunupa.

The code was run iteratively for various annual precipitation andtemperature conditions, until the modeled ice tongue fit themapped glacial deposits of the local LGM (LLGM) at 17–15 ka(moraine M1) (Fig. 4).

4.1.2. Results of glacier modelingSeveral paleoprecipitation and paleotemperature couples (P, T)

are able to reproduce the glacial extent corresponding to themoraine M1. The solutions corresponding to the 17–15 ka episode(LLGM) are thus shown (Figs. 4 and 5) as a curve in (P, T) space.According to our model, the ice extension observed during the17–15 ka period requires a paleo-ELA at 4380�100 m, whichcorresponds to a depression of w1300 m from present conditions.The modeled glacial conditions may have been maintained bya cooling between 6 and 7.5 �C, and with local paleoprecipitationsranging between 800 and 200 mm a�1, respectively (Fig. 5).

4.2. Modeling of the paleolake Tauca highstand (3770 m, 17–15 ka)

4.2.1. Model descriptionIn order to reconstruct the (P, T) scenarios that can reproduce

the highstand of the Tauca paleolake during the 17–15 ka intervalwe used a model derived from the one developed by (Condom et al.,2004). A complete description of the model is available in AppendixB. In this hydrological model, the equation (B2) used to calculateevaporation is derived from the generalized equation of (Xu andSingh, 2000) that relies on an energy balance budget. In the modelof (Condom et al., 2004) the runoff is computed by consideringtemporal and spatial variations the soil evapotranspiration(Makhlouf and Michel, 1994) (see Appendix B for details).


Input data are regional grids of precipitation and temperature(from New et al., 2002), a regional grid describing the geographicextent of the lake area at 3770 m, and watershed domains anddaylight length for the 17–15 ka period (computed from Laskar et al.,2004 and Paillard et al., 1996). The code was run to find all the (P, T)couples solving the general equation describing the steady-statehydrological balance (Precipitation–Evaporation) * lake surfaceþRunoff¼ 0).

4.2.2. Results of modeling of the Tauca paleolake highstand(3770 m)

The precipitation field is not homogenous at the scale of theAltiplano (Fig. 1), and the spatial field of paleoprecipitation duringthe Tauca episode is not known a priori. The area of the Tunupawatershed is small (few km2) compared to the whole Taucacatchment (w200,000 km2 including lakes surfaces); consequentlythe precipitation amount inferred from paleoglacier modelingalone cannot be used. Furthermore, local evaporation may haveproduced a local positive anomaly in precipitation centered overthe lake. This convective precipitation is due to local recycling ofwater, a phenomenon which is observed today on large intertrop-ical lakes such as Titicaca and Victoria (see Supp. Fig. 2) or duringthe Late Pleistocene at Lake Bonneville (e.g. Hostetler et al., 1994).

We thus considered two endmember scenarios to establisha large range for the (P, T) solution curves corresponding to thepaleoclimatic conditions over the Tunupa area during the 17–15 kaepisode (Fig. 5): 1) Scenario 1: The ‘‘minimum precipitation’’solution corresponds to an homogeneous increase of the regionalprecipitation field (with no lake-induced local anomaly), 2)Scenario 2: The ‘‘maximum precipitation’’ solution is obtained fromthe assumption that paleolake Tauca induced a local increase ofprecipitation (the centre of the anomaly is 80% above thesurrounding rainfall value) due to local recycling of the lake Taucaevaporation (see Supp. Fig. 2).

The geometry of the hypothetic precipitation anomaly due tolocal water recycling over Lake Tauca was generated based on threehypothesis (see Supp. Fig. 2): i) The shape of the anomaly roughlyfollows the shape of the paleolake shoreline; ii) The maximumrainfall increase is located at the centre of the lake, as roughlyobserved above Lake Titicaca; iii) An intensity of 80% is consideredas an upper limit for the peak anomaly amplitude, similar to themaximum anomaly currently observed on Lake Victoria, a lake thatis superficially similar to paleolake Tauca (w51,000 km2).

Given that Cerro Tunupa is located at the very centre of paleo-lake Tauca, the potential effect of the recycling anomaly is maximal


Fig. 5. A – Modeled PT solutions able to maintain the Tunupa glacier (ELA at 4380 m, dark blue curve) and the lake highstand (3770 m) during the Tauca interval (17–15 ka).Scenario 2 curve assumes an 80% increase of the regional precipitation field due to local recycling anomaly over Lake Tauca. Scenario 1 curve assumes a uniform increase over theTauca watershed. Surfaces show the 1s confidence interval. B – Precipitation field over the Tauca watershed for the 2 modeled scenarios.


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in the studied area (up to 80%). The scenario 2 curve can thus beconsidered as an upper limit of the precipitation estimate overTunupa during the Tauca episode (Fig. 5).

Our approach relies on the finding that the last glacial maximumof Tunupa and the Tauca highstand were synchronous (within thelimit of the 3Hec dating uncertainty) during the 17–15 ka Taucainterval (Fig. 3). Because lake levels and glacier mass balance aresensitive to precipitation and temperature in distinctly differentways, the intersection of the two curves defines the most probablepaleotemperature and paleoprecipitation conditions over Tunupaduring the Tauca episode. This method indicates that the atmo-spheric cooling was between 7 and 6 �C, with a corresponding localannual rainfall between 320 and 600 mm a�1 (i.e. between �1.6and �3 the present level).

The main uncertainty of our reconstruction arises from theestimate of the amplitude and localization of the precipitationanomaly due to the lake effect, but the precision of this paleo-thermometry approach remains better than those of othermethods, such as geochemical thermomethers (Ghosh et al., 2006).


5. Discussion – paleoclimatic implications

5.1. Comparison with other paleo-glacier and paleo-lakereconstructions

The range of paleoclimatic conditions that we obtain (Fig. 5A)for the 17–15 ka period (DT¼�6 to �7 �C and DP¼þ120 (�1.6) toþ400 (�3) mm a�1) is not significantly different from the estimateof (Clayton and Clapperton, 1997), although these authors useda much simpler model for reconstructing paleo-ELA.

In contrast, (Blodgett et al., 1997) used an energy balance modelwith a higher level of complexity to reconstruct the (P, T) condi-tions required to maintain the Lake Tauca highstand (3770 m).However, the model of (Blodgett et al., 1997) did not take intoaccount the non-linear effect of soil infiltration to calculate therunoff. For comparison the (P, T) lake solution curve reconstructedby (Blodgett et al., 1997) is plotted on Fig. 5. Although this alter-native model is not based on the same assumptions, its (P, T)solution is not significantly different from our reconstructions,



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which provides independent support of accuracy of our model(Fig. 5A).

Regarding the climatic reconstruction based on the 3Hec-datedglacial features, our model implies that the ELA depression wasw1300 m over Tunupa until 15 ka, for a cooling of w6.5 �C (forP� 1.6 to �3 over Tunupa) (Fig. 5A). This temperature drop iscompatible with the w6 �C cooling inferred by (Kull et al., 2008)using a different energy balance based model to reconstruct paleo-ELA at w4300 m in different areas of the Central Andes. Such ELAdepression lies, however, within the largest snowline dropsreported by (Porter, 2001) for the whole Tropics.

The reconstructed shift is also much higher than the ELA drops ofabout 800 m inferred from for the contemporaneous paleoglaciersdated by (Smith et al., 2005b) in the Cordillera Real, 400 km north ofTunupa. If this discrepancy in ELA-shift was only due to temperature,it would require a w3.5 �C difference between the Cordillera Real(w16�S) and the centre of paleolake Tauca (w20�S). However, sucha large difference seems unrealistic because the present-daytemperature difference between these two locations at similarelevations is only 0.5 �C (Ammann et al., 2001). Consequently, it ismore probable that this heterogeneous ELA depression was mainlydriven by spatial variations in precipitation. This interpretation iscompatible with the hypothesis that glacier advances in the CentralAltiplano (w20�S) are mainly moisture sensitive (Zech et al., 2007;Zech et al., 2008) and also with the hypothesis that the rainfalldistribution was different during the Tauca wet episode, because ofa lake-induced precipitation anomaly centered on Lake Tauca.

If such ‘‘maximum precipitation solution’’ (intersection of the‘‘glacier’’ curve and the ‘‘Scenario 2’’ curve, Fig. 5) can be consideredthe most plausible scenario to bridge the gap in ELA depressionwith the Cordillera Real (Smith et al., 2005b), it implies that theTunupa glacier was sustained by cold and wet paleoclimaticconditions during the Tauca period with DT¼w�6 �C, and Pw600 mm a�1.

5.2. Implications for regional and global climate

The main result of this study is that the 17–15 ka period (Taucapaleolake episode) was characterized both by wetter DP¼þ120(�1.6) toþ400 (�3) mm a�1 and colder DT¼�6 to�7 �C conditionscompared to present. This reconstruction thus constrains therespective changes in precipitation and temperature and may thusbring new insights to the debate regarding the interpretation of thestable isotopes data from tropical ice cores of Illimani and Sajama(Thompson et al., 2000; Vuille et al., 2003; Vimeux et al., 2005). Thesignature of the wet Tauca episode in the central Altiplano is alsoattested by records of magnetic susceptibility and clay contents fromTiticaca drill-cores (Baker et al., 2001a,b) and Salar de Uyuni sedi-ments (Baker et al., 2001a), which are qualitative proxies for moisture.The present study thus adds a quantitative constraint on temperatureduring this period: it suggests that both cool and wet conditionscharacterized the Tauca episode, between 17 and 15 ka.

Unfortunately, the glacier features preserved on Tunupa do notallow the amplitude of the ELA depression to be determined beforew20 ka. Hence this record does not permit us to infer a precisetemperature estimate during the global LGM (21�1 ka). However,the snowline depression reported in the Eastern Cordillera duringthe global LGM indicate a cooling between 9 and 6 �C (Kull et al.,2008). Our result imply that, in the Central Andes, such a significanttemperature depression may have persisted until the rapid glacierretreat observed on Tunupa, at w15 ka. The low lake level duringthe 21–17 ka period (Placzek et al., 2006) however suggests that theCentral Altiplano remained dry during this period, and thata significant precipitation increase occurred after w17 ka. Theabrupt character of these rainfall variations is comparable with the


pattern of sea-surface temperatures recorded in the Caribbean Sea(Lea et al., 2003), as well as with the Greenland temperaturepattern (Andersen et al., 2004). In particular, the two wet and coldevents inferred from our lake and glacier analysis, at 17–15 ka andat w12.5 ka, appear to be coincident with the Heinrich 1 and theYounger Dryas events, respectively (Fig. 3). Alternatively, the dryand warm period observed on the Central Altiplano between w14.5and 13.5 ka corresponds to the warm Bølling–Allerød period. Thus,since w17 ka, cold and dry events in the Caribbean and Amazonia(Lea et al., 2003; Cruz et al., 2005) seem to be in phase with cold andwet episodes on the Central Altiplano.

Several studies have suggested that the mechanism of this tightlink might lie in the modulation of the ITCZ position by the AtlanticSST gradient (e.g. Peterson et al., 2000; Haug et al., 2001; Bakeret al., 2001a; Kull et al., 2008), with possibly a significant influenceof the sea ice cover (e.g. Denton et al., 1999; Chiang et al., 2003).However, if this mechanism occurred during Heinrich 1 event, itremains unclear why the global LGM boundary conditions did notinduce a similar rainfall increase over the Altiplano during the21–17 ka period. Further investigations and paleoclimatic recon-structions will be useful to understand this apparent paradox.

6. Conclusions

Our new 3Hec exposure ages obtained from the glacial land-forms of Cerro Tunupa indicate that the oldest glaciation in theCentral Altiplano occurred as early as w160 ka. After a probableperiod of retreat followed by glacier re-advance, the glaciers per-sisted on Tunupa in their maximum position until w15 ka. Theseresults confirm that local LGM in this area was synchronous withthe paleolake Tauca highstand between 17 and 15 ka, as alreadysuggested by (Clayton and Clapperton, 1997). The data also suggestthat the glacier readvanced during the Younger Dryas, before thecomplete deglaciation of Tunupa after w12 ka.

The synchronism between the local LGM and the Tauca high-stand is a remarkable result because it permitted a new approach toreconstruct paleoclimatic conditions during the Tauca episode(17–15 ka). Indeed, combining numerical modeling of paleoglacierextent and of paleolake level, we were able to propose tightquantitative constraints on the temperature drop and precipitationchange during this period. This reconstruction indicates that theTauca episode is the result of both colder (DT¼�6 to �7 �C) andwetter conditions (DP from þ120 (�1.6) to þ400 (�3) mm a�1)compared to current conditions. The main uncertainty in thisreconstruction arises from the fact that the spatial pattern ofpaleoprecipitation is unknown. Further research is required toimprove the reconstruction of this paleoprecipitation field.However, the reconstructed temperature and precipitation condi-tions are sufficiently precise to propose that the climatic changeobserved on the Altiplano are concordant in timing and amplitudeboth with the tropical Atlantic SST (Lea et al., 2003; Cruz et al.,2005) and with the ITCZ position (Cruz et al., 2005; Peterson et al.,2000) during the 17–12 ka period. Our results support thehypothesis that there is a tight link between the paleoclimate of thecentral Altiplano and the dynamics of the Atlantic Ocean duringTermination 1.

It should be noted that such a late local LGM (w15 ka) in thecentre of the Altiplano is apparently in disagreement with the earlyglacial retreat (w25 ka) determined by (Smith et al., 2005b) for thenorthern part of the Altiplano. This finding stresses how sensitiveglacier extents are to both local temperature and precipitation (Kullet al., 2008). Consequently, it can be inappropriate to infer regionalor global climate conclusions from the behavior of individualglaciers. Our approach based on the combined modeling of ancientglaciers and lake level is one strategy by which to separate the



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respective influence of temperature and precipitation. Synchro-nism is however necessary and, so, dating is key.

Acknowledgments

Constructive comments by Wallace S. Broecker, Joerg M.Schaefer and Ano N. Ymous improved an earlier version of thismanuscript. This work was mainly funded by the French INSUprogram ‘‘Relief de la Terre’’ and by the Caltech Tectonics Obser-vatory. It is part of the Post Doc of P.-H. Blard. The authors thank theIRD (Institut de Recherche pour le Developpement) of La Paz,Bolivia, which provided a precious technical and logistical assis-tance in the field. This is CRPG contribution #2017.

Appendix A. Cosmogenic 3He dating of Tunupa glacialdeposits

Methods

The andesite samples were crushed, sieved and the 0.2–0.3 mmand 0.2–0.5 mm fractions were processed to isolate pure pyroxenephenocrysts, ranging in size between 0.2 and 0.6 mm, throughsuccessive magnetic and heavy liquid separation techniques.Although it is difficult to ensure that all the analyzed grains are notfragments from larger minerals, theireuhedral shape suggests that themajority of them are unbroken pyroxenes. Additionally, all sampleswere checked under a binocular microscope to remove grains withadhering lava groundmass or minerals which are not 3He-retentive.

Several aliquots (TU1C, TU6, and TU8) were first step-crushed(for 0.5, 3 and 10 min) under vacuum to estimate both theconcentrations and the isotopic composition of any magmatichelium (Supp. Table 1). Uncrushed samples were fused in a highvacuum furnace to extract all the matrix-sited helium, which in thiscase is composed of cosmogenic 3He, radiogenic 4He, and anyremaining magmatic helium. Typical sample size was between fewtens and hundreds of mg (Table 1). Because the finest crushedgrains may potentially be affected by 3Hec depletion (Blard et al.,2006), only the fractions larger than 150 mm were fused. Smallminerals generally bear lower amounts of magmatic helium (Wil-liams et al., 2005); we thus analyzed aliquots of the sameuncrushed samples (TU1C and TU8) with different grain sizes, 0.2–0.3 mm and 0.3–0.5 mm (Table 1). This experiment was an inde-pendent check of the accuracy of the magmatic correction.

The extracted gas was purified, cryofocused and separated fromneon before being inlet in a MAP 215-50 mass spectrometer. 3Heand 4He were measured by peak-jumping according to the standardprocedure used at Caltech (Patterson and Farley, 1998). The abso-lute sensitivity was determined measuring gas standards of knowncomposition and pressure. The size of the standard was adjusted sothat the 4He pressure in the mass spectrometer was similar forsamples and standards (Burnard and Farley, 2000). Sensitivitieswere w1.7�10�5 cps at�1 and w3�10�7 mV at�1 for 3He and 4He,respectively. Total analytical uncertainties attached to themeasured 3He and 4He concentrations ranged between 2 and 4%(given as 1s), with blank correction <1%.

The major element composition of several samples wasmeasured by electron microprobe. Despite little compositiondifferences, most of the analyzed green pyroxenes have diposidecomposition. However, when another pyroxene specie (blackaugite, with higher Li content) was present in the same rock, bothminerals were analyzed separately to assess possible productionrates differences due to composition. Additionally, the U, Th and Liconcentrations of host rocks and minerals were analyzed by isotopedilution and Inductively Coupled Plasma Mass Spectrometry (Far-ley, 2002) (Table 1 and Supp. Table 2). Those data are crucial to


calculate the amount of 3He produced by nucleogenic process(Andrews and Kay, 1982; Andrews, 1985) and by thermal neutroncapture (Dunai et al., 2007).

Cosmogenic 3He results

Helium results are summarized in Table 1. Total 3He released byfusion ranges between 1.50� 0.05�107 at g�1 and20.2� 0.5�107 at g�1 and 4He between 2.53� 0.1�1012 at g�1 and21.2� 0.1�1012 at g�1. The determination of accurate 3He agesrequires estimating properly the non-cosmogenic components(nucleogenic 3He, magmatic 3He). This can be done througha complete budget of the 3He components contained within thephenocrysts:

3Hec ¼ 3Hetotal � 3Hemag �ZTc

0

Pn$dt (A.1)

where 3Hetotal is the total 3He extracted by fusing the phenocrysts,3Hemag (at g�1) is the inherited (i.e. magmatic) component, Pn

(at g�1 a�1) the nucleogenic production rate of 3He and Tc (a) is theclosure age of the sample.

Correction for magmatic 3He

3Hemag is mainly contained within melt and fluid inclusions,which implies that this component is preferentially released bycrushing (Kurz, 1986a). Consequently, the prolonged vacuumcrushing performed on TU1C, TU6 and TU8 phenocrysts (Supp.Table 1) allowed us to establish the order of magnitude of theinherited 3He: 3Hemag ranges between <0.2�105 and3.6�105 at g�1, yielding an average of 1.5�105 at g�1 anda 3He/4He ratio <5 Ra. We chose to use this average value and thusto apply a 3Hemag correction of 1.5.105�1.5.105 at g�1 to all of thesamples. Such correction could be criticized arguing that cosmo-genic or nucleogenic 3He may be extracted by prolonged crushing(Scarsi, 2000), leading to overestimation of the true 3Hemag

concentration. However, the correction we applied represents lessthan 1% of the total 3He extracted by fusion, even for the pheno-crysts with the lowest 3Hetotal concentration (Table 1). Additionally,although minerals with different sizes are supposed to have con-trasting magmatic helium contents (Williams et al., 2005), thefusion of unbroken aliquots of different grain size (0.2–0.3 and0.3–0.5 mm) yielded indistinguishable 3Hetotal concentrations(Table 1). This result indicates that the magmatic 3He correction isnot a significant source of uncertainty.

Correction for nucleogenic 3He

To obtain a precise constraint on the helium closure age Tc,which is relevant to estimate the nucleogenic build-up

R Tc0 Pn$dt,

we used the (U–Th)/4He* chronometer (Supp. Table 2):

Tc [ 4He � =P4 (A.2)

4He* is the concentration of radiogenic 4He and P4 the produc-tion rate of 4He* within the pyroxenes. The 4He* concentration wascorrected for the magmatic 4He (4Hemag) extracted by melting:

4He � [ 4Hetotal—4Hemag (A.3)

4Hemag was estimated from the crushing experiments (Supp.Table 1). This correction is less than 2% of the total 4He extractedfusing the samples. P4 is calculated from the U and Th concentra-tions measured both in the lava and in the phenocrysts, by applying



ARTICLE IN PRESS

corrections for ejection and implantation of 4He* (Blard and Farley,2008).

The three (U–Th)/4He* ages obtained from samples TU1C, TU3B,TU6, TU7A and TU8 range between 0.5� 0.1 and 1.9� 0.2 Ma, witha weighted average of 1.1�0.5 Ma (Supp. Table 2). The uncertaintymainly arises from the estimate of the 4He* production rate (due tothe implantation/ejection corrections). These closure ages areconsistent with the eruption age of< 5 Ma estimated by (Ville-neuve et al., 2002) for Tunupa volcano.

The flux of radiogenic neutron and the production rate ofnucleogenic 3He, Pn, are calculated for each sample from (Andrewsand Kay, 1982; Andrews, 1985) taking into account the meancomposition of these andesitic rocks and the Li concentrationmeasured in the phenocrysts and the surrounding lavas. This esti-mate takes into account the implantation (and ejection) of 3Hen

from (out of) these 0.2–0.5 mm grains (Farley et al., 2006). Giventhe closure ages of w1 Ma obtained by (U–Th)/4He* dating, theamount of nucleogenic 3He range between 0 and 5% of the total 3Hereleased by fusing the phenocrysts (Table 1).

Cosmogenic 3He production rates and choice of scaling factors

We used two different scaling factors to convert the measured3Hec concentrations into ages. One of them combines the scaling of(Stone, 2000) with the time-dependent one of (Dunai, 2001) withthe (Carcaillet et al., 2004) geomagnetic database. The other usedscaling model was designed by (Lifton et al., 2005). Both calcula-tions include geographic and time-dependent geomagneticcorrections. The factor of (Lifton et al., 2005) also includes theeffects of solar fluctuations on the primary cosmic ray flux. Forconsistency, we used sea-level high latitude (SLHL) productionrates which were calculated using the relevant correcting poly-nomial: 128� 10 (2s) at g�1 a�1 for the (Stone, 2000; Dunai, 2001)factor and 142� 8 (2s) at g�1 a�1 for the (Lifton et al., 2005) factor.These SLHL spallogenic production rates result from the scaling ofthe data from the empirical calibration studies by (Cerling andCraig, 1994; Licciardi et al., 1999; Dunai and Wijbrans, 2000; Ackertet al., 2003; Blard et al., 2006; Licciardi et al., 2006) after correctionfor radiogenic 4He (Blard and Farley, 2008). It is important to notethat these revised production rates are in agreement with the 3Hec

production rates determined in olivines and pyroxenes by cross-calibration against 10Be in quartz (Amidon et al., 2009).

In addition to spallation, 3He is also produced by capture ofcosmogenic thermal neutrons (CTNs) by 6Li (Dunai et al., 2007).This mechanism was considered here, by measuring for eachsample i) the Li concentration of the mineral and its surroundings,ii) the size of the mineral, to correct for ejection and implantation,and iii) the major and trace elements of the whole rock hosts.Although it is potentially important, we ignored the effect of snowon the thermal neutron profile. However, the applied correction forCTN 3He yielded similar ages, within uncertainties, for pyroxenesbelonging to the same rocks but having contrasting Li content(Table 1). For example, in the case of TU3A, the two pyroxenesspecies, with 13 and 75 ppm of Li, yielded 3Hec concentrations of1.51�0.06 and 1.57� 0.06 ka BP, respectively, after correction fornucleogenic 3He. This suggests that the production from CTNs is nota major source of uncertainty for these rocks.

Geometric correction to the production rate took into accountsampling depth (Kurz, 1986b) and the surrounding topography(Dunne et al., 1999). Those corrections are <5% for all the samples(Table 1). The possible influence of erosion was considered bycalculating the exposure ages for no erosion and for denudationrates of 0.4� 0.1 m Ma�1, which is the maximum rate estimated by(Smith et al., 2005a) for comparable objects in this area. However, itmust be noted that, even with this maximum erosion rate, this age-


dependent correction is <1% for the samples of the last glaciation(<20 ka) and w7% for the oldest moraine M0 (>100 ka) (Table 1).

Appendix B. Description of the hydrological model used toreproduce the P–T conditions necessary to maintain the LakeTauca highstand (3770 m, 17–15 ka)

The model used in this study is derived from the one developedby (Condom et al., 2004). Because the goal of the exercise is toreconstruct the precipitation–temperature (P–T) conditions of theLake Tauca episode (constant lake level at w3770 m), the calcula-tion aimed to determine all P–T couples solving the generalhydrological balance equation at steady-state (dVlake/dt¼ 0), for thewhole Tauca whatershed:

Plake L Elake D RWS [ 0 (B1)

where Plake (m3 a�1) is the rainfall flux falling over the Taucalake, Elake (m3 a�1) the evaporation flux over the Tauca lake freesurface and RWS is the runoff (m3 a�1) of the endorheic Taucawatershed. This watershed includes the Titicaca subcatchment,which is in hydrological connection with the southern Tauca lakesubcatchment (Condom et al., 2004). Given the Tauca watershed isendorheic, no external sink or source is considered.

The model was run using a Matlab� code. This code computes thehydrological budget for each 15�15 km cell of a grid describing thelandscape of the whole Tauca lake watershed. The input grid wasgenerated using a 15 km resolution SRTM DEM. The contour of Titi-caca and Tauca lakes were determined from their respective waterlevel during the 17–15 ka period (3810 m for the Titicaca and 3770 mfor the Tauca). The resulting areas of lakes and sub-watershed were:

- Tauca surface: 50,900 km2

- Titicaca surface: 8750 km2

- Tauca sub-watershed: 87,975 km2 (138,875 km2 with LakeTauca)

- Titicaca sub-watershed: 49,500 km2 (58,250 km2 with LakeTiticaca)

- Total watershed: 137,475 km2

- Total (watershedsþ lakes): 197,125 km2

In order to take into account the non-linear effect of infiltrationon evapotranspiration, the model uses two different evaporationlaws for the soil and the lake surface. Both evaporation laws arebased on the generalized evaporation equation of (Xu and Singh,2000). For a 3-month period (n) the evaporation from the lakesurface, Elake (in mm/quarter�1) is calculated from:

ElakeðnÞ ¼3� 365

12��

0:1þ 0:7� Dtrue

DCS

�� Re

� ðT þ 17:8Þ � 0:0145595� 0:51� T

(B2)

Where Re¼ 11794 J cm�2 day�1 is the extraterrestrial radiation, T(�C) is the mean trimestrial air temperature, DCS (h day�1) is theclear sky trimestrial daylength, Dtrue (h day�1) the actual trimestrialdaylength. Because of cloudiness, Dtrue is shorter than DCS. Climaticobservations show that the actual daylight (Dtrue) is closely linkedto precipitation (Condom et al., 2004). In the model, Dtrue wascalculated from the empirical relationship proposed by (Condomet al., 2004) (R2¼ 0.8):

Dtrue [ 9:18—0:008 3 P (B3)

P (mm/quarter�1) being the trimestrial rainfall.The calculation ofthe runoff (RWS) takes into account the quantity of water available



ARTICLE IN PRESS

in the soil. Indeed, the evapotranspiration is closely linked to thewater content of the soil. The model used here includes a moduledescribing the saturation state of the soil. This module is largelybased on the one developed by (Makhlouf and Michel, 1994). Ateach 3 month time-step (n), the model takes into account the watercontent of the soil of the previous time step, H(n� 1) (in mm/quarter�1), to compute the water evaporated from the soil (Esoil (n))(in mm/quarter�1) (Condom et al., 2004):

EsoilðnÞ ¼�

Hðn� 1ÞK

��

2� Hðn� 1ÞK

�� E (B4)

with:

E ¼ 3� 36512

��

0:1þ 0:7� Dtrue

DCS

�� Re �

ðT þ 17:8Þ � 0:0145595� 0:51� T

(B5)

K (in mm/quarter�1) is a parameter defining the soil capacity. IfH(n)> K, then runoff occurs, if H(n)< K then there is water infil-tration into the soil. For each time step, the new saturation state ofthe soil H(n) is calculated through a complete hydrological budget,taking into account the previous saturation state of the soilH(n� 1), the local precipitation P(n), the local evaporation Esoil(n).

The model also takes into account the fact that infiltration isimpossible in areas of land covered by permanent ice: abovea certain elevation (4400 m for the local LGM), we indeed consid-ered that all rainfall contributes to the hydrological budget as runoff(and assumed that ice and snowmelt is dominant over sublima-tion). The code was run for a 24 month duration to ensure the initialconditions of soil saturation reached steady-state. The K-parameterwas calibrated by using current meteorological observations (Newet al., 2002) and the present-day hydrological budget of the Titicacawatershed:

PTiticaca L ETiticaca D RWSTiticaca [ QDes (B6)

where QDes¼ 1.1�109 m3 a�1 is the mean annual flow of the RıoDesaguadero, which is the Titicaca lake overflow. This calibrationyielded a K-parameter value of 200 mm quarter�1. This is notsignificantly different from the value of 240 mm quarter�1 obtainedby (Condom et al., 2004) by using a 0D approach and a differentmeteorological dataset.

However, in order to evaluate the sensitivity of our results to theK-parameter, the precipitation–temperature curve solving theTauca lake was also computed by running the model with K value of100 and 300 mm quarter�1. This allowed an uncertainty area to beplotted (Fig. 5) for both scenarios (uniform precipitation increaseand Tauca lake anomaly).

Appendix. Supplementary data

Supplementary data associated with this article can be found inthe online version, at doi:10.1016/j.quascirev.2009.09.025.

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