Late Local Glacial Maximum in the Central Altiplano Triggered by Cold and Locally-Wet Conditions During the Paleolake Tauca Episode (17–15 Ka, Heinrich 1)

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Quaternary Science Reviews xxx (2009) 1–14

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Late local glacial maximum in the Central Altiplano triggered by cold and locally-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. Jimeńez e, V. Ramirez e a Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA, USA b Centre de Recherches Pe´trographiques et Geóchimiques, UPR 2300, CNRS, Nancy-Universite´, Vandœuvre-le`s-Nancy, France c Institut de Recherche pour le De´veloppement, France d Geóazur, UMR 6526, CNRS, Universite´ de Nice, Nice, France e Universidad Mayor de San Andre´s, La Paz, Bolivia article info abstract

Article history: The timing and causes of the last deglaciation in the southern tropical Andes is poorly known. In the Received 8 July 2009 Central Altiplano, recent studies have focused on whether this tropical highland was deglaciated before, Received in revised form synchronously or after the global last glacial maximum (w21 ka BP). In this study we present a new 18 September 2009 3 3 chronology based on cosmogenic He ( Hec) dating of moraines on Cerro Tunupa, a volcano that is Accepted 28 September 2009 3 located in the centre of the now vanished Lake Tauca (19.9 S, 67.6 W). These new Hec ages suggest that the Tunupa glaciers remained close to their maximum extent until 15 ka BP, synchronous with the Lake Tauca highstand (17–15 ka BP). Glacial retreat and the demise of Lake Tauca seem to have occurred rapidly and synchronously, within dating uncertainties, at w15 ka BP. We took advantage of the synchronism of these events to combine a glacier model with a lake model in order to reconstruct precipitation and temperature during the Lake Tauca highstand. This new approach indicates that, during the Tauca highstand (17–15 ka BP), the centre of the Altiplano was characterized by temperature w6.5 C cooler and average precipitation higher by a factor ranging between 1.6 and 3 compared to the present. Cold and wet conditions thus persisted in a signiﬁcant part of the southern tropical Andes during the Heinrich 1 event (17–15 ka BP). This study also demonstrates the extent to which the snowline of glaciers can be affected by local climatic conditions and emphasizes that efforts to draw global climate inferences from glacial extents must also consider local moisture conditions. Ó 2009 Elsevier Ltd. All rights reserved.

1. Introduction mountain glaciers can be used to reconstruct with precision changes in temperature and precipitation (e.g. Hostetler and Clark, The Tropics are recognized to play a major role in paleoclimatic 2000). The Altiplano is a key area within this tropical puzzle: this cycles (Stocker et al., 2001) because this area may potentially high plateau is located at the interface between tropical and generate forcing (i) at the global scale, e.g. through the tight link mid-latitude atmospheric circulation (Kull et al., 2008) and so is between tropical sea-surface temperature (SST) and atmospheric likely to have experienced substantial changes in precipitation and CO2 and CH4 concentrations (Lea, 2004), and (ii) at the regional temperature during the last deglaciation (20–10 ka), and particu- scale, e.g. through the feedbacks between moisture transport and larly during the 17.5–14.5 ka period termed ‘‘Mystery Interval’’ by the thermohaline circulation (Leduc et al., 2007). However, our (Denton et al., 2006). understanding of the exact mechanisms involved in these inter- Although the presence of ancient mountain glaciers is reported actions is limited by the lack of precise chronologies in several key in several places on the Altiplano (Smith et al., 2008; Zech et al., tropical areas (Hastenrath, 2009). This is particularly true for the 2008), the spatial pattern and the timing of paleosnowline changes highlands of the continental realm, where the past extents of remain challenging issues (Rodbell et al., 2009). A discussion between Clark (2002) and Seltzer et al. (2002) focused on the relative timing between the global last glacial maximum (LGM, 21 ka) and the late Pleistocene glaciation in the tropical Andes. * Corresponding author. Centre de Recherches Pe´trographiques et Ge´ochimiques, Cosmogenic 10Be (10Be ) ages from moraines located in the UPR 2300, CNRS, Nancy-Universite´, Vandœuvre-le`s-Nancy, France. Tel.: þ33 3 83 59 c 42 23. northern part of the Altiplano suggest that the last glacial E-mail address: [email protected] (P.-H. Blard). maximum occurred as early as w25 ka BP in the Tropical Andes

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(Smith et al., 2005b, 2008). However, other Altiplano locations deflection of the Inter Tropical Convergence Zone (ITCZ) during the yielded younger local LGM ages of w15 ka (Seltzer, 1992; Clayton so-called ‘‘South American Monsoon season’’ (Fig. 1)(Garreaud and Clapperton, 1997; Zech et al., 2007). This contrasting behavior et al., 2009). As a consequence of orographic effects, there is also suggests that local effects, such as asynchronous changes in the a strong northeast to southwest gradient in annual precipitation: precipitation field, were superimposed on the regional temperature rainfall decreases from 800 mm a1 in the northern part of the trend that affected the tropical Andes. Indeed glacier fluctuations Bolivian Altiplano to less than 100 mm a1 in the Lipez area in can result either from changes in precipitation or temperature southern Bolivia (Fig. 1). The endorheic watershed of the central (Ohmura et al., 1992). This ambiguity makes it difficult to interpret Altiplano includes two subcatchments: the Titicaca watershed in local glacier advances as a global temperature signal in the absence the north and the Salar de Uyuni catchment in the south. However of an independent estimate of local precipitation conditions. these catchments are hydrologically connected through the Rio Paleolakes may be very useful proxies in that, when combined Desaguadero, the outlet of Lake Titicaca (Fig. 1). with glacier-based paleoclimatic reconstructions, they can be used Due to low precipitation (<1000 mm a1) and mean annual to estimate the relative effects of precipitation and temperature. potential evaporation ranging between 1000 and 2000 mm a1 Although the Altiplano is now dry and covered by large salty (Montes de Oca, 1989), the present moisture budget of the Altiplano deserts, large palaeolakes were developed on this plateau at various is in deficit. This results in shallow oligosaline lakes (e.g. Lake times in the late Pleistocene period (Minchin, 1882; Servant and Poopo) and large saline basins (Salar de Uyuni and Salar de Coipasa) Fontes, 1978; Seltzer, 1994; Sylvestre et al., 1999). The chronology of (Fig. 1). However, evidence for several episodes of deep Late the last and highest lake level stage, known as the Tauca phase Pleistocene lakes (Seltzer, 1994; Sylvestre et al., 1999; Placzek et al., (w3770 m) is now relatively well-established between w17 and 2006) attest to the fact that Altiplano hydrographic conditions were 15 ka thanks to 14C and 230Th/234U dating of lacustrine formations significantly different in the past (Argollo and Mourguiart, 2000; (Sylvestre et al., 1999; Placzek et al., 2006). Past occurrence of Placzek et al., 2006). These features document a high sensitivity to glaciers in areas that are now ice-free is documented by moraines regional moisture and temperature conditions that have varied deposited on the slopes surrounding the palaeolake Tauca. The through time. Previous studies proposed reconstructions of the relationship between the local last glacial maximum and the Tauca paleoclimatic conditions that triggered and maintained the paleo- has been debated for a long time, and at least two scenarios have lake Tauca highstand (Hastenrath and Kutzbach, 1985; Blodgett been proposed: et al., 1997). Using evaporation models at steady-state, these authors concluded that Lake Tauca resulted from rainfall ranging i) The glaciers receded from their maximum position before the between a few percent and 80% greater than present mean condi- rise of the Lake Tauca (Servant and Fontes, 1978), and, tions. However these reconstructions suffered from the fact that air according to these authors, the lake was filled by the glaciers temperature was not independently constrained during the Tauca meltwater. However this idea was later refuted by budget phase (w15 ka). Moreover part of the variability also arises from calculations showing that the total volume of LGM mountain the calibration of the model parameters (Blodgett et al., 1997). glacier ice was insufficient to supply the deep Tauca paleolake Therefore, the present study takes advantage (see Section 4.2)of (Hastenrath and Kutzbach, 1985; Blodgett et al., 1997). the hydrological model of (Condom et al., 2004), which offers the ii) The glaciers reached their maximum extent - or significantly advantage of having been parameterized using the present-day readvanced – synchronously with the Tauca episode between conditions of the Titicaca watershed. 17 and 15 ka BP (Seltzer, 1992; Clayton and Clapperton, 1997; Under present-day conditions, the glacier equilibrium-line Clapperton et al., 1997; Zech et al., 2007). In this case, both altitude (ELA) is at w5200 m in the northern part of the Altiplano events could result from the same climate forcing. (Ribstein et al., 1995), and rises south-westward up to w6000 m in the Lipez area, the extreme southern part of the plateau (Klein et al., In order to define the relative chronology of ice advance and lake 1995). This regional ELA trend parallels the present negative highstand, and to improve our understanding of the paleoclimatic precipitation gradient from northeast to southwest, suggesting conditions on the Altiplano at the end of the Pleistocene, we a strong influence of the regional snowfall pattern (Ammann et al., 3 present here new Hec exposure ages measured on pyroxenes 2001). The paleo-ELA reconstructed from several moraines phenocrysts sampled on glacial landforms on an andesitic volcano suggests that during the last glacial maximum snowlines ranged located at the centre of palaeolake Tauca. After having confirmed between 800 and 1200 m lower than today (Clayton and Clapper- synchronism between the Central Altiplano local LGM and the ton, 1997; Dornbusch, 2000). However, due to limited dating, it Tauca highstand (17 and 15 ka BP), we investigate a joint inversion remains unclear whether this snowline depression was synchro- of the paleoclimatic conditions (precipitation, temperature) pre- nous at the regional scale of the Altiplano. Importantly, the influ- vailing during those glacial times by using numerical models of ence of temperature reduction and precipitation increase on the paleoglacier extent and paleolake level. spatial variations in ELA depression are not yet well-established (Clayton and Clapperton, 1997; Kull et al., 2008). 2. Geomorphological and (paleo)-climatic settings of the central Altiplano 3. Cosmogenic 3He glacial chronology of Cerro Tunupa

The Altiplano plateau (14–22S) is a large (w500,000 km2), flat, 3.1. Glacial features of Cerro Tunupa high-elevation area (mean elevation range between 3600 and 3800 m) located in the tropical Central Andes. Precipitation mainly Cerro Tunupa (19.86S, 67.61W) is an inactive volcano located originates from humid air masses brought by the tropical easterlies in the centre of the now vanished paleolake Tauca (Fig. 1). It is from the Tropical Atlantic. Indeed, although the Altiplano is also mainly built of Quaternary andesite lavas that erupted less than under the influence of westerly winds from the Central Pacific, 5Ma(Villeneuve et al., 2002). The flanks of the Tunupa volcano clouds coming from the northeast are currently responsible for exhibit several landforms and glaciogenic deposits testifying to the w80% of the small annual rainfall on the plateau (Vuille, 1999). It presence of ancient glaciers (Fig. 1). These glacial features are presents a highly seasonal pattern, with 80% falling from October to remarkably well-preserved, suggesting that erosion rates are not April (Austral summer), as a consequence of the southward significant, thanks to the low precipitation in this area (w200 mm

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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 the Altiplano: 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 present annual 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 in Zech et al. (2007). CA: Cerro Azanaques dated by 14CinClayton and Clapperton (1997). Sajama, 6542 m, is the highest peak in Bolivia.

per year) (New et al., 2002). One of the largest glacial landforms is glacifluvial fan delta Mf-1 (Fig. 2). Two distinct lobes can be iden- developed on the south-east flank of Tunupa, in the Chalchala tified on this fan: the older lobe, Mf-1a, is overlapped by the valley (Fig. 2)(Clapperton et al., 1997). Detailed mapping of this moraine M1, and has two major erosional benches whose altitudes valley was achieved through field observations and analyses of (3760 and 3770 m) suggest that they result from the Tauca paleo- aerial photos and satellite images (Fig. 2). lake highstands that occurred between 17 and 15 ka (Sylvestre This glacial valley was initially connected to an old moraine et al., 1999; Placzek et al., 2006). These strandlines are conspicu- outlet (M0), which has the shape of crab pincers and is oriented ously developed on both sides of the fan Mf-1. The younger fan, southward. A prominent sharp-crested moraine, M1 crosscut this Mf-1b, lies in the axis of the former glacier tongue that formed M1 old moraine M0, and presently delimits the former glaciated U- moraines. In contrast with Mf-1a, Mf-1b does not record a trace of shaped Chalchala valley with w100 m-high-shoulders. This main the highest shoreline (3770 m), but only a depositional bench in moraine M1 appears to be genetically connected to a distal continuity with the lower erosional strandline (3660 m) (Fig. 2).

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3 Fig. 2. A – Mapping of the Cerro Tunupa glacial deposits, with the sample locations and Hec ages. B – Picture of boulder TU1B sampled on the M1 moraine. C – Picture of boulder TU5A sampled on the Mf-1b ﬂuvio-glacial fan delta.

These stratigraphic relationships therefore suggest that: 1. the early at the top of moraine crests (Fig. 2). We also collected several M1 stage of the glacier advance, which led to the frontal deposition striated rocks (TU2 and TU4 samples) at the flat bottom of the of the glacifluvial material of Mf-1a, is antecedent to the highest Chalchala valley. The well-preserved striations suggest that these lake stand; 2. the maximum glacial advance and M1 construction, surfaces have suffered limited erosion. Moreover, the sampling was accompanied by frontal deposition of the glacifluvial material of done in the upper part of prominent knobs or whale-backs to Mf-1b, post-dates the highest lake stand but is coeval with the minimize the probability of temporary shielding by till or sediment following high stand (10 m below). This scenario is also consistent cover. Finally, the deep downcutting of the glacier between the with sedimentologic observations on these fan deltas in the Chal- phases M0 and M1 ensures that glacial erosion has been sufficiently chala valley outlet and in the nearby Pocolli valley, indicating that intense to fully reset the striated rocks, as demonstrated in regions Mf-1 is cross-stratified with the 3760 m and 3770 m shoreline that experienced extensive glacial erosion (e.g. Gayer et al., 2006). deposits of the Tauca paleolake (Clayton and Clapperton, 1997; Clapperton et al., 1997)(Fig. 2). 3.2. Cosmogenic 3He dating of the Tunupa glacial landforms If this interpretation is correct, the deposition age of Mf-1b fan can thus be bracketed between 16.5 and 15 ka and considered as 3.2.1. Methods 3 3 a reference site for testing the accuracy of the used Hec production Cosmogenic He data and calculated exposure ages are given in rates. Table 1. Moreover, if the Mf-1 fan is the distal equivalent of the last These data were obtained from samples bearing millimeter-size stages of the moraine M1 building (or M2), all these observations pyroxenes phenocrysts. This mineral is well-suited for cosmogenic provide strong support for a tight temporal coincidence between 3He dating because it has high helium retentivity (Trull et al., 1991; the termination of moraine M1 (or M2) and the Tauca highstand, Trull and Kurz, 1993; Blard et al., 2008). Empirical calibration from dated between 17 and 15 ka (Placzek et al., 2006). independently dated surfaces established its production rate of 3He Just after the building episode of the M1 moraine, the initial between 128 and 142 at g1 a1 at sea-level high latitude, stage of glacier retreat built the lateral moraine, M2, which is depending on the used scaling factor (Ackert et al., 2003; Blard interlocked within the inner part of the lateral moraine M1. The et al., 2006; Amidon et al., 2009). Both helium isotopes were subsequent glacier retreat left the bottom of the Chalchala valley measured at Caltech with a MAP 215-50 mass spectrometer. To free of debris deposits and exposed numerous well-preserved estimate the magmatic helium component, several aliquots were striated rocks. Two elongated moraines M3 delineate a later and first crushed in vacuum and the extracted gas analyzed (Supp. Table short duration re-advance of a very narrow glacier tongue that 1). All pyroxene samples were then fused in a vacuum furnace to developed above these striated rocks and results from an over- completely release and analyze the cosmogenic 3He (Table 1) and spilling originating from the upper glacier cirque. The small the radiogenic 4He* components (Supp. Table 2). Measurement, moraine M3 is thus stratigraphically younger than the highest correction for magmatic and nucleogenic 3He and calculations of roches moutonneés of the valley at w4500 m. exposure ages are detailed in Appendix A. The total correction for Here we use cosmogenic 3He dating to complement these non-cosmogenic 3He is less than 5% for all samples and does not observations with an absolute chronology of these glacial features. represent a major source of uncertainty (see Appendix A), as sup- For this purpose, we sampled several meter-sized boulders on the ported by the excellent agreement between different replicates, moraines M0, M1, Mf-1 and M3 (Fig. 2 and Table 1), paying special including those which vary in Li content (the source of nucleogenic attention to select the top of the most prominent boulders located 3He and cosmogenic thermal-neutron production; see Table 1).

Please cite this article in press as: Blard, P.-H., et al., Late local glacial maximum in the Central Altiplano triggered by cold and locally-wet..., Quaternary Science Reviews (2009), doi:10.1016/j.quascirev.2009.09.025 laect hsatcei rs s lr,P-. ta. aelclgailmxmmi h eta liln rgee ycl n locally-wet..., and cold by triggered Altiplano Central the in maximum glacial local Late al., et P.-H., doi:10.1016/j.quascirev.2009.09.025 Blard, (2009), as: Reviews press Science in Quaternary article this cite Please Table 1 Description of the Tunupa moraine samples and cosmogenic 3He ages.

4 3 3 3 Sample Object Boulder Altitude Latitude Longitude Mineral Size Mass Het Het Li Hen CTN P 3 Hec Depth Paleomagnetic Exposure age (ka) ELA from 12 1 7 1 5 1 a 1 1 b 7 1 c d e height (m) ( S) ( W) (mm) (mg) (10 at g ) (10 at g ) (ppm) (10 at g ) (at g a ) (10 at g ) correction correction numerical (cm) (applied to modeling the Stone factor) Lifton et al., Stone, 2005 2000; Dunai, 2001 Tunupa Moraine M0 TU7A Boulder 200 3828 19.8816 67.5971 Green 0.2–0.5 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 4 142 6 4200 diopside (121 4) (150 6) TU7B Boulder 100 3829 19.8813 67.5973 Green 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 4 160 6 4200 diopside (135 5) (170 6) TU7C Boulder 150 3829 19.8813 67.5973 Green 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 5 163 7 4200 diopside (138 5) (175 7) Fluvio-glacial Mf1a.

TU5A Boulder 100 3760 19.8742 67.5880 Green 0.2–0.5 257.4 3.26 0.03 1.59 0.06 19 2.3 0.2 5 1.55 0.05 0.98 0.98 Average Average 1–14 (2009) xxx Reviews Science Quaternary / al. et Blard P.-H. diopside TU5A Black 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 0.97 13.8 0.5 16.2 0.6 4380 PRESS IN ARTICLE augite TU5B Boulder 350 3760 19.8742 67.5880 Green 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 diopside Mf1b TU6 Boulder 250 3810 19.8789 67.5955 Green 0.2–0.5 183.9 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 diopside TU8 Boulder 200 3778 19.9632 67.5941 Green 0.2–0.5 318.8 4.68 0.04 1.90 0.05 60 6.4 0.6 1.82 0.05 diopside TU8fa Green 0.2–0.3 178.2 5.16 0.05 1.91 0.08 60 6.4 0.6 1.83 0.08 diopside TU8fb Green 0.2–0.3 189.3 5.26 0.05 1.90 0.07 60 6.4 0.6 1.82 0.08 diopside TU8 Green 0.2–0.3 79.0 5.32 0.1 1.93 0.06 60 6.4 0.6 1.85 0.06 diopside TU8G Green 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 diopside 15 0.97 1.01 13.4 0.5 17.8 0.8 4380 Moraine M1 TU1A Boulder 250 4260 19.8643 67.6128 Green 0.2–0.5 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 diopside TU1B Boulder 300 4290 19.8638 67.6144 Green 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 diopside TU1C Boulder 120 4206 19.8652 67.6099 Green 0.2–0.5 377.0 3.67 0.03 2.84 0.06 20 1.9 0.2 2.8 0.06 diopside TU1Cf Green 0.2–0.3 159.1 4.29 0.04 3.05 0.1 20 1.9 0.2 3.02 0.1 diopside TU1CG Green 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 diopside 5 0.99 1.06 19.4 0.7 23.2 0.9 4380 TU1D Boulder 300 4083 19.8691 67.6054 Green 0.2–0.5 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 diopside TU4 Striated 4260 19.8570 67.6180 Green 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 rock/roche diopside moutonne´ e TU2 Striated 4450 19.8455 67.6272 Green 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 rock/roche diopside moutonne´ e (continued on next page ) 5 ARTICLE IN PRESS

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Because the Altiplano is located in the Tropics, the main source of ). uncertainty in the moraine ages arises from the scaling factor used 3 to calculate the local production rate. We thus computed Hec ages

numerical modeling using two different scaling procedures: the one of (Lifton et al., 0.5 4690 0.6 4690 2005) and the one of (Stone, 2000) implemented with the time- Supp. Table 1 integrated correction of (Dunai, 2001). Correction for topographic shielding and snow cover are negligible. Zero-erosion assumption is discussed in Appendix A. 0.4 13.2 0.4 12.3

3 He* dating (See 3.2.2. He dating results and choice of the scaling factor

4 c Average Average Exposure age (ka) ELA from 9.8 10.7 The cosmogenic 3He ages are presented in Table 1. They range between 115 4 ka and 163 7 ka for moraine M0, between

e 12.7 0.5 and 27 1.1 ka for moraine M1 and striated rocks, and between 9.8 0.4 and 13.2 0.6 ka for moraine M3, depending on the scaling factor used (Table 1). The error ages, given at the 1s

Paleomagnetic correction (applied to the Stone factor) level, include analytical uncertainties only.

d 3 The scaling of Hec production rate is potentially an important source of uncertainty for cosmogenic dating at low latitudes and ). high elevation (Balco et al., 2008). In the present study, we can Depth correction

c however rely on the stratigraphic relationships existing between ) 1

the distal glacial deposits (lobe Mf-1) and the lake shorelines. 0.06 0.99 0.95 0.06 0.99 0.96 0.05 0.99 0.96 at g

c Indeed the boulders TU5A and TU5B sit on the fan Mf-1b whose 7 ). Closure age is determined from (U–Th)/ 1.7 He deposition is probably synchronous (Fig. 2 and Section 3.1 for 3 (10

b w ) details) with the lake highstand (>3760 m) dated between 17 1 Carcaillet et al., 2004

a and 15 ka BP (Sylvestre et al., 1999; Placzek et al., 2006). The scaling 3 1

3 factors of (Lifton et al., 2005) yield Hec ages that are younger by at Andrews, 1985

CTN P (at g least w1.5 ka compared to the lake highstand episode. On this a )

1 basis, we conclude that this correction factor is probably not adequate for this period and this region (Supp. Fig. 1). We conse- 0.2 3 1.51 0.9 19 1.57 0.6 11 at g n 5 quently decided to consider only the ages calculated with the He 3 (10 (Stone, 2000; Dunai, 2001) polynomial to discuss the Tunupa

. glacier chronology and its climatic implications. 2 Li (ppm) ) 1

3.2.3. The glacial chronology of Cerro Tunupa – synchronism 0.06 13 1.6 0.05 75 8.5 0.05 42 6.2 . at g between the paleolake Tauca and the local last glacial maximum 1 t 7 ). The 3He concentrations measured in the three boulders TU7A, B He c 3 (10 at g ) and by using the magnetic database of ( )

5 and C (M0 moraine) indicate that the deposition of the M0 moraine 1 ended between 142 6 and 163 7 ka under the assumption of no 0.1 1.54 0.06 1.67 0.04 1.78 at g 1.5.10 erosion, and between 150 6 and 175 7 ka if a maximum erosion t 12 1 5 He rate of 0.4 0.1 m Ma (Smith et al., 2005a) is considered. The Dunai, 2001 4 (10

Dunai et al., 2007 relatively good cluster of these three ages suggests that pre- and post-depositional processes have not signiﬁcantly disturbed the Mass (mg) exposure history of this old moraine (Smith et al., 2005a). This old He is 1.5.10 3 and an attenuation length of 160 g cm glacial stage recorded on Cerro Tunupa is consistent with the 3 0.3–0.5 28.0 21.2 0.3–0.5 38.9 1.03 0.3–0.5 67.0 4.85 (mm) glacier re-advance observed by (Smith et al., 2005a) between 170 and 125 ka in Peru. At a broader scale, there is also a good temporal coincidence between this glacier stillstand and the penultimate diopside augite diopside Mineral Size glaciation of Oxygen Isotope Stage 6, between 140 and 170 ka (Petit et al., 1999).

W) The exposure ages measured on Mf-1, M1 and M3 range Longitude ( between 27 1.1 and 12.3 0.5 ka by using the (Stone, 2000; 3 Dunai, 2001) scaling. These Hec ages along with our ﬁeld obser- S)

vations allow the following chronology to be proposed: Latitude ( He is performed assuming magmatic

3 3 The Hec date of 27 1.1 ka from the boulder TU6 provides

Altitude (m) a maximum age for the first pulse of the glacier tongue that led to the deposition of the Mf-1 fluvio-glacial fan. TU6 is located ) is calculated by taking into account the whole rock composition and the Li content of the phenocrysts ( n on the fluvio-glacial deposit Mf-1, clearly above the highest He height (cm) 3

) shoreline, so there is little probability that the exposure age of He (

3 TU6 has been affected by post-depositional processes, such as exhumation due to wave erosion at the lake shoreline. If this is continued

( the correct interpretation of this date, such an early glacier advance would be in agreement with the 10Be ages of w25 ka

Nucleogenic The production rate from cosmogenicCorrection thermal for neutron non is cosmogenic computedDepth following correction ( is calculated assumingCorrection a for rock time density variations of of 2.7 the g geomagnetic cm ﬁeld are performed following ( obtained by (Smith et al., 2005b) in the northern part of the c Moraine M3 TU3A Boulder 50 4420 19.8455 67.6263 Green TU3A Boulder 50 4420 19.8455 67.6263 Black TU3B Boulder 50 4420 19.8455 67.6263 Green Sample Object Boulder Table 1 a e b d 3 Altiplano. However, inheritance of Hec cannot be excluded for

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respectively. These ages are additional support for the 3Hec chronology of the Tunupa glacial deposits (ka BP) A hypothesis of a late (w15 ka) retreat of the Cerro Tunupa 0 5101520 25 30 4800 glaciers from their maximum position. Moreover, this glacial This study Striated rock retreat exhibits a striking similitude with the drop of the Tauca water level at w15 ka. Although uncertainty in the cosmogenic Boulder 4700 dates does not permit us to conclude which event occurred ﬁrst (ice retreat or lake level drop), the timing and amplitude of the two abrupt events are very similar. This suggests either a causal 4600 link between these two retreats, or that both lake and mountain glaciers responded almost simultaneously to an external 4500 forcing due to an abrupt regional climatic change affecting the

Modelled ELA (m) ? ? Central Altiplano. 3 Hec ages of the boulders belonging to the small moraine M3 4400 (TU3A and TU3B) indicate a glacial re-advance between 12.3 0.5 and 13.2 0.6 ka. Although limited in amplitude, 4300 this stillstand seems to be synchronous with the Coipasa wet YD BA H1 phase corresponding to a new rise of the lake level at 3700 m B 3780 Data by Lake Tauca between 13 and 12 ka (Fig. 3). Moreover, such timing is also 3760 Placzek et al., 2006 compatible, within uncertainty, with the Younger Dryas (YD) cold event that occurred between 12.9 and 11.6 ka in the 14C dating 3740 Northern Hemisphere (Andersen et al., 2004). Our data are thus U-Th dating in agreement with previous 10Be ages in the Cordillera Blanca 3720 by (Farber et al., 2005) supporting the hypothesis of a YD Lake Coipasa signature in the South tropical Andes. However, because of the 3700 dating uncertainty and the limited number of samples, it is not possible to exclude the possibility that this re-advance was 3680 synchronous with the Antarctic cold reversal event, dated at Lake Sajsi w14 ka (Jouzel et al., 1995). Additional ages and calibrations Elevation of lake surface (m) 3660 Dry period Dry period will thus be useful to address the question of a YD occurrence in the Central Altiplano. 3640 0 5101520 25 30 The climatic implications of this glacial chronology are exam- Lake chronology (ka BP) ined in the next section by using numerical modeling of paleo-

3 climatic conditions, with a special emphasis on the synchronism Fig. 3. A– Hec glacial chronology of Cerro Tunupa. Scaling factors are computed using a combination of the scaling corrections of (Stone, 2000) and (Dunai, 2001) and a SLHL between the highstand of paleolake Tauca and the local last glacial 3 1 1 Hec production rate of 128 at g a (Blard et al., 2006). The ELA is modeled by using maximum between 17 and 15 ka. the approach described in Section 4.1. H1 refers to the Heinrich 1 event, BA to the Bølling–Allerød and YD to the Younger Dryas. B – Lake level chronology from (Placzek 4. Numerical modeling of paleoclimatic conditions et al., 2006). (precipitation, temperature) during the paleolake Tauca episode (17–15 ka) TU6, and w27 ka should thus be considered as a probable maximum age for the deposition of the Mf-1a formation. 4.1. Modeling of the ELA depression All the other boulders belonging to the main moraine M1 or the 3 fluvio-glacial deposit Mf-1 are characterized by Hec ages 4.1.1. Model description ranging between 15.6 0.6 and 17.9 0.7 ka, with the excep- In order to interpret the observed Tunupa glacier extents in tion of TU1C, from moraine M1, whose high age (23.2 0.9 ka) terms of paleoclimatic conditions (temperature and precipitation), suggests this boulder is an outlier affected by inheritance. we used a numerical approach that combines an ice mass-balance Despite the fact that these absolute ages are dependent on the model (Blard et al., 2007) with a simplified 2-D ice flow model choice of scaling factor (see discussion in Appendix A), their (Harper and Humphrey, 2003). The computation of mass-balance range suggest that the peak of the local last glacial maximum relies on a positive-degree-month model in which ablation is occurred between 18 and 15 ka. This timing is not significantly proportional to temperature and shortwave solar radiation (Hock, different from a recent LLGM dating in southern Perou 1999). Different ablation parameters were used for snow and ice in (Bromley et al., 2009), while the LLGM in the Cordillera Real order to include the influence of albedo on melting. This empirical occurred several thousand years before (w25 ka) (Smith et al., model does not directly include the effects of wind and relative 2005b). Importantly, these new 3He ages from Cerro Tunupa humidity. However, the adopted melting factors have been cali- strongly suggest that the glacier remained close to its brated using the mass-balance and climatic dataset monitored maximum position until w15 ka. Such scenario is consistent monthly on the Zongo glacier (16S) over the 2003–2005 period with the conclusions of (Clayton and Clapperton, 1997; Clap- (Blard et al., 2007). Given the proximity of these locations, and the perton et al., 1997) and our morphostratigraphic observations fact that the modeled altitude range for these glaciers overlaps, it is suggesting that the glacial stage that produced the M1 moraine reasonable to conclude that they are characterized by comparable persisted during the w2 ka duration of the paleolake Tauca melting processes, so that the reconstructed melting rates are not highstand between 17 and 15 ka (Fig. 3)(Sylvestre et al., 1999; significantly biased by their calibration on the Zongo glacier. Placzek et al., 2006). Moreover, in order to be conservative with the uncertainties Striated rocks TU2 (4450 m) and TU4 (4260 m) from roches attached to our reconstruction, we considered the maximum range 3 moutonneés yielded Hec dates of 15.6 0.6 and 15.7 0.6 ka, (from 0.7 to 1.3) reported for these melting parameters in the

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

A B 6000 ~17-15 ka glaciers Present-day ELA = 5680 m 5200 5600 M1: Local LGM moraine 150 Tunupa summit Mf-1b: Fluvio-glacial deposit 4800 5200 100

4400 50 4800 M1 ELA = 4380 m Ice thickness (m) 0 Elevation (m) 4000 Elevation (m) 4400 Palaeo ELA (~17-15 ka) = 4380 m

3600 4000 -1 0 ΔT = -6.5°C

2 Mf-1b N xP = × 3 = 600 mm.a Eastern dis4 tance (km) Tauca Lake 3770 m 3600 ~17-15 ka 0 -10 -8 -6 -4 -2 0 2 6 2 8 4 -1 6 Glacier mass balance (mwater.a ) 10 8 Southern distance (km)

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, ELA ¼ 5680 m) and the LLGM (blue curve, ELA ¼ 4380 m). See Section 4.1 and (Blard et al., 2007) for details. tropical area (Hock, 2003). The input data used in the model are Input data are regional grids of precipitation and temperature monthly temperature and precipitation (see Supp. Table 3), direct (from New et al., 2002), a regional grid describing the geographic solar radiation during the 17–15 ka period (Paillard et al., 1996; extent of the lake area at 3770 m, and watershed domains and Laskar et al., 2004), and a digital elevation model of the South-East daylight length for the 17–15 ka period (computed from Laskar et al., flank of the Tunupa volcano. A local lapse-rate of 2004 and Paillard et al., 1996). The code was run to find all the (P, T) w6.5 0.2 Ckm1 was assumed (Klein et al., 1999). The present couples solving the general equation describing the steady-state Equilibrium-Line Altitude (ELA) modeled with the present-day hydrological balance (Precipitation–Evaporation) * lake surface þ conditions stands at 5680 100 m. Cerro Tunupa, 5400 m high, is Runoff ¼ 0). not covered by permanent ice now and cannot be used to check the model, but this calculated ELA is consistent with the snowline 4.2.2. Results of modeling of the Tauca paleolake highstand currently observed on the rare glaciated Altiplano summits, such as (3770 m) Sajama (18060S–68520W) (Fig. 1) and Parinacota volcano The precipitation field is not homogenous at the scale of the (18100S–69080W) (Ammann et al., 2001; Hastenrath, 2009), that Altiplano (Fig. 1), and the spatial field of paleoprecipitation during are under precipitation and temperature conditions similar to those the Tauca episode is not known a priori. The area of the Tunupa of Tunupa. watershed is small (few km2) compared to the whole Tauca The code was run iteratively for various annual precipitation and catchment (w200,000 km2 including lakes surfaces); consequently temperature conditions, until the modeled ice tongue fit the the precipitation amount inferred from paleoglacier modeling mapped glacial deposits of the local LGM (LLGM) at 17–15 ka alone cannot be used. Furthermore, local evaporation may have (moraine M1) (Fig. 4). produced a local positive anomaly in precipitation centered over the lake. This convective precipitation is due to local recycling of 4.1.2. Results of glacier modeling water, a phenomenon which is observed today on large intertrop- Several paleoprecipitation and paleotemperature couples (P, T) ical lakes such as Titicaca and Victoria (see Supp. Fig. 2) or during are able to reproduce the glacial extent corresponding to the the Late Pleistocene at Lake Bonneville (e.g. Hostetler et al., 1994). moraine M1. The solutions corresponding to the 17–15 ka episode We thus considered two endmember scenarios to establish (LLGM) are thus shown (Figs. 4 and 5) as a curve in (P, T) space. a large range for the (P, T) solution curves corresponding to the According to our model, the ice extension observed during the paleoclimatic conditions over the Tunupa area during the 17–15 ka 17–15 ka period requires a paleo-ELA at 4380 100 m, which episode (Fig. 5): 1) Scenario 1: The ‘‘minimum precipitation’’ corresponds to a depression of w1300 m from present conditions. solution corresponds to an homogeneous increase of the regional The modeled glacial conditions may have been maintained by precipitation field (with no lake-induced local anomaly), 2) a cooling between 6 and 7.5 C, and with local paleoprecipitations Scenario 2: The ‘‘maximum precipitation’’ solution is obtained from ranging between 800 and 200 mm a1, respectively (Fig. 5). the assumption that paleolake Tauca induced a local increase of precipitation (the centre of the anomaly is 80% above the 4.2. Modeling of the paleolake Tauca highstand (3770 m, 17–15 ka) surrounding rainfall value) due to local recycling of the lake Tauca evaporation (see Supp. Fig. 2). 4.2.1. Model description The geometry of the hypothetic precipitation anomaly due to In order to reconstruct the (P, T) scenarios that can reproduce local water recycling over Lake Tauca was generated based on three the highstand of the Tauca paleolake during the 17–15 ka interval hypothesis (see Supp. Fig. 2): i) The shape of the anomaly roughly we used a model derived from the one developed by (Condom et al., follows the shape of the paleolake shoreline; ii) The maximum 2004). A complete description of the model is available in Appendix rainfall increase is located at the centre of the lake, as roughly B. In this hydrological model, the equation (B2) used to calculate observed above Lake Titicaca; iii) An intensity of 80% is considered evaporation is derived from the generalized equation of (Xu and as an upper limit for the peak anomaly amplitude, similar to the Singh, 2000) that relies on an energy balance budget. In the model maximum anomaly currently observed on Lake Victoria, a lake that of (Condom et al., 2004) the runoff is computed by considering is superficially similar to paleolake Tauca (w51,000 km2). temporal and spatial variations the soil evapotranspiration Given that Cerro Tunupa is located at the very centre of paleo- (Makhlouf and Michel, 1994) (see Appendix B for details). lake Tauca, the potential effect of the recycling anomaly is maximal

P.-H. Blard et al. / Quaternary Science Reviews xxx (2009) 1–14 9

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 the Tauca watershed. Surfaces show the 1s confidence interval. B – Precipitation field over the Tauca watershed for the 2 modeled scenarios. in the studied area (up to 80%). The scenario 2 curve can thus be 5. Discussion – paleoclimatic implications considered as an upper limit of the precipitation estimate over Tunupa during the Tauca episode (Fig. 5). 5.1. Comparison with other paleo-glacier and paleo-lake Our approach relies on the finding that the last glacial maximum reconstructions of Tunupa and the Tauca highstand were synchronous (within the 3 limit of the Hec dating uncertainty) during the 17–15 ka Tauca The range of paleoclimatic conditions that we obtain (Fig. 5A) interval (Fig. 3). Because lake levels and glacier mass balance are for the 17–15 ka period (DT ¼6to7 C and DP ¼þ120 (1.6) to sensitive to precipitation and temperature in distinctly different þ400 (3) mm a1) is not significantly different from the estimate ways, the intersection of the two curves defines the most probable of (Clayton and Clapperton, 1997), although these authors used paleotemperature and paleoprecipitation conditions over Tunupa a much simpler model for reconstructing paleo-ELA. during the Tauca episode. This method indicates that the atmo- In contrast, (Blodgett et al., 1997) used an energy balance model spheric cooling was between 7 and 6 C, with a corresponding local with a higher level of complexity to reconstruct the (P, T) condi- annual rainfall between 320 and 600 mm a1 (i.e. between 1.6 tions required to maintain the Lake Tauca highstand (3770 m). and 3 the present level). However, the model of (Blodgett et al., 1997) did not take into The main uncertainty of our reconstruction arises from the account the non-linear effect of soil infiltration to calculate the estimate of the amplitude and localization of the precipitation runoff. For comparison the (P, T) lake solution curve reconstructed anomaly due to the lake effect, but the precision of this paleo- by (Blodgett et al., 1997) is plotted on Fig. 5. Although this alter- thermometry approach remains better than those of other native model is not based on the same assumptions, its (P, T) methods, such as geochemical thermomethers (Ghosh et al., 2006). solution is not significantly different from our reconstructions,

10 P.-H. Blard et al. / Quaternary Science Reviews xxx (2009) 1–14 which provides independent support of accuracy of our model pattern of sea-surface temperatures recorded in the Caribbean Sea (Fig. 5A). (Lea et al., 2003), as well as with the Greenland temperature 3 Regarding the climatic reconstruction based on the Hec-dated pattern (Andersen et al., 2004). In particular, the two wet and cold glacial features, our model implies that the ELA depression was events inferred from our lake and glacier analysis, at 17–15 ka and w1300 m over Tunupa until 15 ka, for a cooling of w6.5 C (for at w12.5 ka, appear to be coincident with the Heinrich 1 and the P 1.6 to 3 over Tunupa) (Fig. 5A). This temperature drop is Younger Dryas events, respectively (Fig. 3). Alternatively, the dry compatible with the w6 C cooling inferred by (Kull et al., 2008) and warm period observed on the Central Altiplano between w14.5 using a different energy balance based model to reconstruct paleo- and 13.5 ka corresponds to the warm Bølling–Allerød period. Thus, ELA at w4300 m in different areas of the Central Andes. Such ELA since w17 ka, cold and dry events in the Caribbean and Amazonia depression lies, however, within the largest snowline drops (Lea et al., 2003; Cruz et al., 2005) seem to be in phase with cold and reported by (Porter, 2001) for the whole Tropics. wet episodes on the Central Altiplano. The reconstructed shift is also much higher than the ELA drops of Several studies have suggested that the mechanism of this tight about 800 m inferred from for the contemporaneous paleoglaciers link might lie in the modulation of the ITCZ position by the Atlantic dated by (Smith et al., 2005b) in the Cordillera Real, 400 km north of SST gradient (e.g. Peterson et al., 2000; Haug et al., 2001; Baker Tunupa. If this discrepancy in ELA-shift was only due to temperature, et al., 2001a; Kull et al., 2008), with possibly a significant influence it would require a w3.5 C difference between the Cordillera Real of the sea ice cover (e.g. Denton et al., 1999; Chiang et al., 2003). (w16S) and the centre of paleolake Tauca (w20S). However, such However, if this mechanism occurred during Heinrich 1 event, it a large difference seems unrealistic because the present-day remains unclear why the global LGM boundary conditions did not temperature difference between these two locations at similar induce a similar rainfall increase over the Altiplano during the elevations is only 0.5 C(Ammann et al., 2001). Consequently, it is 21–17 ka period. Further investigations and paleoclimatic recon- more probable that this heterogeneous ELA depression was mainly structions will be useful to understand this apparent paradox. driven by spatial variations in precipitation. This interpretation is compatible with the hypothesis that glacier advances in the Central 6. Conclusions Altiplano (w20S) are mainly moisture sensitive (Zech et al., 2007; 3 Zech et al., 2008) and also with the hypothesis that the rainfall Our new Hec exposure ages obtained from the glacial land- distribution was different during the Tauca wet episode, because of forms of Cerro Tunupa indicate that the oldest glaciation in the a lake-induced precipitation anomaly centered on Lake Tauca. Central Altiplano occurred as early as w160 ka. After a probable If such ‘‘maximum precipitation solution’’ (intersection of the period of retreat followed by glacier re-advance, the glaciers per- ‘‘glacier’’ curve and the ‘‘Scenario 2’’ curve, Fig. 5) can be considered sisted on Tunupa in their maximum position until w15 ka. These the most plausible scenario to bridge the gap in ELA depression results confirm that local LGM in this area was synchronous with with the Cordillera Real (Smith et al., 2005b), it implies that the the paleolake Tauca highstand between 17 and 15 ka, as already Tunupa glacier was sustained by cold and wet paleoclimatic suggested by (Clayton and Clapperton, 1997). The data also suggest conditions during the Tauca period with DT ¼ w6 C, and P that the glacier readvanced during the Younger Dryas, before the w600 mm a1. complete deglaciation of Tunupa after w12 ka. The synchronism between the local LGM and the Tauca high- 5.2. Implications for regional and global climate stand is a remarkable result because it permitted a new approach to reconstruct paleoclimatic conditions during the Tauca episode The main result of this study is that the 17–15 ka period (Tauca (17–15 ka). Indeed, combining numerical modeling of paleoglacier paleolake episode) was characterized both by wetter DP ¼þ120 extent and of paleolake level, we were able to propose tight (1.6) to þ400 (3) mm a1 and colder DT ¼6to7 C conditions quantitative constraints on the temperature drop and precipitation compared to present. This reconstruction thus constrains the change during this period. This reconstruction indicates that the respective changes in precipitation and temperature and may thus Tauca episode is the result of both colder (DT ¼6to7 C) and bring new insights to the debate regarding the interpretation of the wetter conditions (DP from þ120 (1.6) to þ400 (3) mm a1) stable isotopes data from tropical ice cores of Illimani and Sajama compared to current conditions. The main uncertainty in this (Thompson et al., 2000; Vuille et al., 2003; Vimeux et al., 2005). The reconstruction arises from the fact that the spatial pattern of signature of the wet Tauca episode in the central Altiplano is also paleoprecipitation is unknown. Further research is required to attested by records of magnetic susceptibility and clay contents from improve the reconstruction of this paleoprecipitation field. Titicaca drill-cores (Baker et al., 2001a,b) and Salar de Uyuni sedi- However, the reconstructed temperature and precipitation condi- ments (Baker et al., 2001a), which are qualitative proxies for moisture. tions are sufficiently precise to propose that the climatic change The present study thus adds a quantitative constraint on temperature observed on the Altiplano are concordant in timing and amplitude during this period: it suggests that both cool and wet conditions both with the tropical Atlantic SST (Lea et al., 2003; Cruz et al., characterized the Tauca episode, between 17 and 15 ka. 2005) and with the ITCZ position (Cruz et al., 2005; Peterson et al., Unfortunately, the glacier features preserved on Tunupa do not 2000) during the 17–12 ka period. Our results support the allow the amplitude of the ELA depression to be determined before hypothesis that there is a tight link between the paleoclimate of the w20 ka. Hence this record does not permit us to infer a precise central Altiplano and the dynamics of the Atlantic Ocean during temperature estimate during the global LGM (21 1 ka). However, Termination 1. the snowline depression reported in the Eastern Cordillera during It should be noted that such a late local LGM (w15 ka) in the the global LGM indicate a cooling between 9 and 6 C(Kull et al., centre of the Altiplano is apparently in disagreement with the early 2008). Our result imply that, in the Central Andes, such a significant glacial retreat (w25 ka) determined by (Smith et al., 2005b) for the temperature depression may have persisted until the rapid glacier northern part of the Altiplano. This finding stresses how sensitive retreat observed on Tunupa, at w15 ka. The low lake level during glacier extents are to both local temperature and precipitation (Kull the 21–17 ka period (Placzek et al., 2006) however suggests that the et al., 2008). Consequently, it can be inappropriate to infer regional Central Altiplano remained dry during this period, and that or global climate conclusions from the behavior of individual a significant precipitation increase occurred after w17 ka. The glaciers. Our approach based on the combined modeling of ancient abrupt character of these rainfall variations is comparable with the glaciers and lake level is one strategy by which to separate the

P.-H. Blard et al. / Quaternary Science Reviews xxx (2009) 1–14 11 respective influence of temperature and precipitation. Synchro- calculate the amount of 3He produced by nucleogenic process nism is however necessary and, so, dating is key. (Andrews and Kay, 1982; Andrews, 1985) and by thermal neutron capture (Dunai et al., 2007). Acknowledgments Cosmogenic 3He results Constructive comments by Wallace S. Broecker, Joerg M. Schaefer and Ano N. Ymous improved an earlier version of this Helium results are summarized in Table 1. Total 3He released by manuscript. This work was mainly funded by the French INSU fusion ranges between 1.50 0.05 107 at g1 and program ‘‘Relief de la Terre’’ and by the Caltech Tectonics Obser- 20.2 0.5 107 at g1 and 4He between 2.53 0.1 1012 at g1 and vatory. It is part of the Post Doc of P.-H. Blard. The authors thank the 21.2 0.1 1012 at g1. The determination of accurate 3He ages IRD (Institut de Recherche pour le De´veloppement) of La Paz, requires estimating properly the non-cosmogenic components Bolivia, which provided a precious technical and logistical assis- (nucleogenic 3He, magmatic 3He). This can be done through tance in the field. This is CRPG contribution #2017. a complete budget of the 3He components contained within the phenocrysts: Appendix A. Cosmogenic 3He dating of Tunupa glacial deposits ZTc 3 3 3 $ Hec ¼ Hetotal Hemag Pn dt (A.1) Methods 0 3 3 The andesite samples were crushed, sieved and the 0.2–0.3 mm where Hetotal is the total He extracted by fusing the phenocrysts, 3 1 and 0.2–0.5 mm fractions were processed to isolate pure pyroxene Hemag (at g ) is the inherited (i.e. magmatic) component, Pn 1 1 3 phenocrysts, ranging in size between 0.2 and 0.6 mm, through (at g a ) the nucleogenic production rate of He and Tc (a) is the successive magnetic and heavy liquid separation techniques. closure age of the sample. Although it is difficult to ensure that all the analyzed grains are not 3 fragmentsfrom larger minerals,theireuhedral shape suggests that the Correction for magmatic He majority of them are unbroken pyroxenes. Additionally, all samples 3 were checked under a binocular microscope to remove grains with Hemag is mainly contained within melt and fluid inclusions, adhering lava groundmass or minerals which are not 3He-retentive. which implies that this component is preferentially released by Several aliquots (TU1C, TU6, and TU8) were first step-crushed crushing (Kurz, 1986a). Consequently, the prolonged vacuum (for 0.5, 3 and 10 min) under vacuum to estimate both the crushing performed on TU1C, TU6 and TU8 phenocrysts (Supp. concentrations and the isotopic composition of any magmatic Table 1) allowed us to establish the order of magnitude of the 3 3 5 helium (Supp. Table 1). Uncrushed samples were fused in a high inherited He: Hemag ranges between <0.2 10 and 5 1 5 1 vacuum furnace to extract all the matrix-sited helium, which in this 3.6 10 at g , yielding an average of 1.5 10 at g and 3 4 case is composed of cosmogenic 3He, radiogenic 4He, and any a He/ He ratio <5 Ra. We chose to use this average value and thus 3 5 5 1 remaining magmatic helium. Typical sample size was between few to apply a Hemag correction of 1.5.10 1.5.10 at g to all of the tens and hundreds of mg (Table 1). Because the finest crushed samples. Such correction could be criticized arguing that cosmo- 3 3 grains may potentially be affected by Hec depletion (Blard et al., genic or nucleogenic He may be extracted by prolonged crushing 3 2006), only the fractions larger than 150 mm were fused. Small (Scarsi, 2000), leading to overestimation of the true Hemag minerals generally bear lower amounts of magmatic helium (Wil- concentration. However, the correction we applied represents less 3 liams et al., 2005); we thus analyzed aliquots of the same than 1% of the total He extracted by fusion, even for the pheno- 3 uncrushed samples (TU1C and TU8) with different grain sizes, 0.2– crysts with the lowest Hetotal concentration (Table 1). Additionally, 0.3 mm and 0.3–0.5 mm (Table 1). This experiment was an inde- although minerals with different sizes are supposed to have con- pendent check of the accuracy of the magmatic correction. trasting magmatic helium contents (Williams et al., 2005), the The extracted gas was purified, cryofocused and separated from fusion of unbroken aliquots of different grain size (0.2–0.3 and 3 3 neon before being inlet in a MAP 215-50 mass spectrometer. He 0.3–0.5 mm) yielded indistinguishable Hetotal concentrations 3 and 4He were measured by peak-jumping according to the standard (Table 1). This result indicates that the magmatic He correction is procedure used at Caltech (Patterson and Farley, 1998). The abso- not a significant source of uncertainty. lute sensitivity was determined measuring gas standards of known composition and pressure. The size of the standard was adjusted so Correction for nucleogenic 3He that the 4He pressure in the mass spectrometer was similar for samples and standards (Burnard and Farley, 2000). Sensitivities To obtain a precise constraint on the helium closureR age Tc, w 5 1 w 7 1 3 4 Tc $ were 1.7 10 cps at and 3 10 mV at for He and He, which is relevant to estimate the nucleogenic build-up 0 Pn dt, respectively. Total analytical uncertainties attached to the we used the (U–Th)/4He* chronometer (Supp. Table 2): measured 3He and 4He concentrations ranged between 2 and 4% [ 4 (given as 1s), with blank correction <1%. Tc He =P4 (A.2) The major element composition of several samples was 4He* is the concentration of radiogenic 4He and P the produc- measured by electron microprobe. Despite little composition 4 tion rate of 4He* within the pyroxenes. The 4He* concentration was differences, most of the analyzed green pyroxenes have diposide corrected for the magmatic 4He (4He ) extracted by melting: composition. However, when another pyroxene specie (black mag augite, with higher Li content) was present in the same rock, both 4 [ 4 4 He Hetotal— Hemag (A.3) minerals were analyzed separately to assess possible production 4 rates differences due to composition. Additionally, the U, Th and Li Hemag was estimated from the crushing experiments (Supp. concentrations of host rocks and minerals were analyzed by isotope Table 1). This correction is less than 2% of the total 4He extracted dilution and Inductively Coupled Plasma Mass Spectrometry (Far- fusing the samples. P4 is calculated from the U and Th concentra- ley, 2002)(Table 1 and Supp. Table 2). Those data are crucial to tions measured both in the lava and in the phenocrysts, by applying

12 P.-H. Blard et al. / Quaternary Science Reviews xxx (2009) 1–14 corrections for ejection and implantation of 4He* (Blard and Farley, dependent correction is <1% for the samples of the last glaciation 2008). (<20 ka) and w7% for the oldest moraine M0 (>100 ka) (Table 1). 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, with a weighted average of 1.1 0.5 Ma (Supp. Table 2). The uncertainty Appendix B. Description of the hydrological model used to mainly arises from the estimate of the 4He* production rate (due to reproduce the P–T conditions necessary to maintain the Lake the implantation/ejection corrections). These closure ages are Tauca highstand (3770 m, 17–15 ka) consistent with the eruption age of < 5 Ma estimated by (Ville- neuve et al., 2002) for Tunupa volcano. The model used in this study is derived from the one developed The flux of radiogenic neutron and the production rate of by (Condom et al., 2004). Because the goal of the exercise is to nucleogenic 3He, P , are calculated for each sample from (Andrews reconstruct the precipitation–temperature (P–T) conditions of the n w and Kay, 1982; Andrews, 1985) taking into account the mean Lake Tauca episode (constant lake level at 3770 m), the calcula- composition of these andesitic rocks and the Li concentration tion aimed to determine all P–T couples solving the general measured in the phenocrysts and the surrounding lavas. This esti- hydrological balance equation at steady-state (dVlake/dt ¼ 0), for the 3 whole Tauca whatershed: mate takes into account the implantation (and ejection) of Hen from (out of) these 0.2–0.5 mm grains (Farley et al., 2006). Given P L E D R [ 0 (B1) the closure ages of w1 Ma obtained by (U–Th)/4He* dating, the lake lake WS 3 3 3 1 amount of nucleogenic He range between 0 and 5% of the total He where Plake (m a ) is the rainfall flux falling over the Tauca 3 1 released by fusing the phenocrysts (Table 1). lake, Elake (m a ) the evaporation flux over the Tauca lake free 3 1 surface and RWS is the runoff (m a ) of the endorheic Tauca Cosmogenic 3He production rates and choice of scaling factors watershed. This watershed includes the Titicaca subcatchment, which is in hydrological connection with the southern Tauca lake We used two different scaling factors to convert the measured subcatchment (Condom et al., 2004). Given the Tauca watershed is 3 Hec concentrations into ages. One of them combines the scaling of endorheic, no external sink or source is considered. (Stone, 2000) with the time-dependent one of (Dunai, 2001) with The model was run using a MatlabÒ code. This code computes the the (Carcaillet et al., 2004) geomagnetic database. The other used hydrological budget for each 15 15 km cell of a grid describing the scaling model was designed by (Lifton et al., 2005). Both calcula- landscape of the whole Tauca lake watershed. The input grid was tions include geographic and time-dependent geomagnetic generated using a 15 km resolution SRTM DEM. The contour of Titi- corrections. The factor of (Lifton et al., 2005) also includes the caca and Tauca lakes were determined from their respective water effects of solar fluctuations on the primary cosmic ray flux. For level during the 17–15 ka period (3810 m for the Titicaca and 3770 m consistency, we used sea-level high latitude (SLHL) production for the Tauca). The resulting areas of lakes and sub-watershed were: rates which were calculated using the relevant correcting polynomial: 128 10 (2s)atg1 a1 for the (Stone, 2000; Dunai, 2001) - Tauca surface: 50,900 km2 factor and 142 8(2s)atg1 a1 for the (Lifton et al., 2005) factor. - Titicaca surface: 8750 km2 These SLHL spallogenic production rates result from the scaling of - Tauca sub-watershed: 87,975 km2 (138,875 km2 with Lake the data from the empirical calibration studies by (Cerling and Tauca) Craig, 1994; Licciardi et al., 1999; Dunai and Wijbrans, 2000; Ackert - Titicaca sub-watershed: 49,500 km2 (58,250 km2 with Lake et al., 2003; Blard et al., 2006; Licciardi et al., 2006) after correction Titicaca) for radiogenic 4He (Blard and Farley, 2008). It is important to note - Total watershed: 137,475 km2 3 2 that these revised production rates are in agreement with the Hec - Total (watersheds þ lakes): 197,125 km production rates determined in olivines and pyroxenes by cross- calibration against 10Be in quartz (Amidon et al., 2009). In order to take into account the non-linear effect of infiltration In addition to spallation, 3He is also produced by capture of on evapotranspiration, the model uses two different evaporation cosmogenic thermal neutrons (CTNs) by 6Li (Dunai et al., 2007). laws for the soil and the lake surface. Both evaporation laws are This mechanism was considered here, by measuring for each based on the generalized evaporation equation of (Xu and Singh, sample i) the Li concentration of the mineral and its surroundings, 2000). For a 3-month period (n) the evaporation from the lake 1 ii) the size of the mineral, to correct for ejection and implantation, surface, Elake (in mm/quarter ) is calculated from: and iii) the major and trace elements of the whole rock hosts. Although it is potentially important, we ignored the effect of snow 3 365 Dtrue ElakeðnÞ¼ 0:1 þ 0:7 Re on the thermal neutron profile. However, the applied correction for 12 DCS 3 CTN He yielded similar ages, within uncertainties, for pyroxenes ðT þ 17:8Þ 0:0145 (B2) belonging to the same rocks but having contrasting Li content 595 0:51 T (Table 1). For example, in the case of TU3A, the two pyroxenes 2 1 species, with 13 and 75 ppm of Li, yielded 3He concentrations of Where Re ¼ 11794 J cm day is the extraterrestrial radiation, T c 1 ( C) is the mean trimestrial air temperature, DCS (h day ) is the 1.51 0.06 and 1.57 0.06 ka BP, respectively, after correction for 1 nucleogenic 3He. This suggests that the production from CTNs is not clear sky trimestrial daylength, Dtrue (h day ) the actual trimestrial a major source of uncertainty for these rocks. daylength. Because of cloudiness, Dtrue is shorter than DCS. Climatic Geometric correction to the production rate took into account observations show that the actual daylight (Dtrue) is closely linked sampling depth (Kurz, 1986b) and the surrounding topography to precipitation (Condom et al., 2004). In the model, Dtrue was calculated from the empirical relationship proposed by (Condom (Dunne et al., 1999). Those corrections are <5% for all the samples 2 (Table 1). The possible influence of erosion was considered by et al., 2004)(R ¼ 0.8): calculating the exposure ages for no erosion and for denudation Dtrue [ 9:18—0:008 3 P (B3) rates of 0.4 0.1 m Ma1, which is the maximum rate estimated by (Smith et al., 2005a) for comparable objects in this area. However, it P (mm/quarter1) being the trimestrial rainfall.The calculation of must be noted that, even with this maximum erosion rate, this age- the runoff (RWS) takes into account the quantity of water available

P.-H. Blard et al. / Quaternary Science Reviews xxx (2009) 1–14 13 in the soil. Indeed, the evapotranspiration is closely linked to the Fujii, Y., Goto-Azuma, K., Gronvold, K., Gundestrup, N.S., Hansson, M., Huber, C., water content of the soil. The model used here includes a module Hvidberg, C.S., Johnsen, S.J., Jonsell, U., Jouzel, J., Kipfstuhl, S., Landais, A., Leuenberger, M., Lorrain, R., Masson-Delmotte, V., Miller, H., Motoyama, H., describing the saturation state of the soil. This module is largely Narita, H., Popp, T., Rasmussen, S.O., Raynaud, D., Rothlisberger, R., Ruth, U., based on the one developed by (Makhlouf and Michel, 1994). At Samyn, D., Schwander, J., Shoji, H., Siggard-Andersen, M.L., Steffensen, J.P., each 3 month time-step (n), the model takes into account the water Stocker, T., Sveinbjornsdottir, A.E., Svensson, A., Takata, M., Tison, J.L., Thorsteinsson, T., Watanabe, O., Wilhelms, F., White, J.W.C., 2004. High-reso- content of the soil of the previous time step, H(n 1) (in mm/ lution record of Northern Hemisphere climate extending into the Last Inter- 1 quarter ), to compute the water evaporated from the soil (Esoil (n)) glacial period. Nature 431, 147–151. (in mm/quarter1)(Condom et al., 2004): Andrews, J.N.,1985. The isotopic composition of radiogenic helium and its use to study groundwater movement in confined aquifers. Chemical Geology 49, 339–351. Andrews, J.N., Kay, R.L.F., 1982. Natural production of tritium in permeable rocks. Hðn 1Þ Hðn 1Þ E ðnÞ¼ 2 E (B4) Nature 298, 361–363. soil K K Argollo, J., Mourguiart, P., 2000. Late Quaternary climate history of the Bolivian Altiplano. Quaternary International 72, 37–51. with: Baker, P.A., Rigsby, C.A., Seltzer, G.O., Fritz, S.C., Lowenstein, T.K., Bacher, N.P., Veliz, C., 2001a. Tropical climate changes at millennial and orbital timescales on the Bolivian Altiplano. Nature 409, 698–701. 3 365 Dtrue ðT þ 17:8Þ0:0145 E ¼ 0:1 þ 0:7 Re Baker, P.A., Seltzer, G.O., Fritz, S.C., Dunbar, R.B., Grove, M.J., Tapia, P.M., Cross, S.L., 12 DCS 595 0:51 T Rowe, H.D., Broda, J.P., 2001b. The history of South American tropical precipitation for the past 25,000 years. Science 291, 640–643. (B5) Balco, G., Stone, J.O., Lifton, N.A., Dunai, T.J., 2008. A complete and easily accessible 10 26 1 means of calculating surface exposure ages or erosion rates from Be and Al K (in mm/quarter ) is a parameter defining the soil capacity. If measurements. Quaternary Geochronology 3, 174–195. H(n) > K, then runoff occurs, if H(n) < K then there is water infil- Blard, P.-H., Farley, K.A., 2008. The influence of radiogenic 4He on cosmogenic 3He tration into the soil. For each time step, the new saturation state of determinations in volcanic olivine and pyroxene. Earth and Planetary Science the soil H(n) is calculated through a complete hydrological budget, Letters 276, 20–29. 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