36Cl Chronologies and ELA reconstructions from the northern boundary of the South American Arid Diagonal

A thesis submitted to the Graduate School of the University of Cincinnati in partial fulfillment of the requirements of the degree of

Master of Science

in the Department of Geology of the McMicken College of Arts and Sciences by

Rachel Thornton

B.S. (Geology), Kent State University, 2011

Advisory Committee:

Dylan Ward, Ph.D. (Committee Chair) Thomas Lowell, Ph.D. Aaron Diefendorf, Ph.D.

i ABSTRACT

This study focused on glaciation in the South American Arid Diagonal (AD) climatic feature which intersects in the Andes mountain range at approximately 24°S. This NE-SW- trending zone of hyper-aridity separates the Atlantic-sourced Easterly moisture belt from the

Pacific-sourced Westerly precipitation in the central Andes. We present new geomorphological maps of the landforms, particularly of glacial moraines across two mountain ranges within the

Western Cordillera of the central Andes. Additionally, new 36Cl chronologies using surface exposure dating were produced from two sites near the northern boundary of the modern AD.

One site is located in the Cordillera del Tatio at ~21°S (Tatio field site), the other is the southernmost valley in a stratovolcano chain called the Cordón de Puntas Negras at ~24°S (SPN field site). Paleo-ELAs were reconstructed utilizing a two-dimensional numerical model to simulate glaciers over a 90m digital elevation model of the upstream catchment.

We find that four glacial stages were present at both field sites sampled. Glacial stabilizations from the El Tatio field site, once assumed Lateglacial in age are on the order of 10-

20 thousand years older than previously published. Tatio ages ranged from ~25 ka to as early as

MIS 5. We also found that dated glacial stages from El Tatio and the Puntas Negras are not synchronous with lake highstands on the Altiplano, but projected ages of the younger glacial stages indicate both sites were glaciated during the Tauca phase. Regional comparison indicated the possibility of glacial occupation of the Tatio site into the Lateglacial during the Coipasa phase (12 ka) after rapid deglaciation of the SPN field site. Regional ELA comparisons supported the trend of moisture loss from the NE to SW portions of the climatic AD feature.

ii

iii ACKNOWLEDGEMENTS

First and foremost, I thank my advisor Dr. Dylan Ward for this opportunity, his patience, and for pushing me to accomplish things I once considered out of reach. I’d also like to thank my supportive committee Dr. Thomas Lowell and Dr. Aaron Diefendorf for their feedback, guidance, and encouragement. I would not have made it through this project without the help of my lab group, particularly Chris Sheehan and Jason Cesta for their help collecting and processing my samples, for which I also owe thanks to Sarah Hammer. I’d like to acknowledge Dr. Esteban Sagredo for his advice and help during our field season in Chile. For help, support, much-needed laughter, and friendship I thank my friends inside and outside the Department of Geology. Finally, I’d like to thank my family, particularly my husband Robert Naylor, for supporting me and encouraging me the whole way. This project was funded by a National Science Foundation award from Geomorphology & Land-use Dynamics (GLD) (EAR-1226611) to Dr. Dylan Ward. Additional funding was granted by the Geological Society of America and Association for Women Geoscientists.

iv TABLE OF CONTENTS

Abstract…………………………………………………………………………………………….i

Acknowledgements………………………………………………………………...... iv

List of Figures...... ………………………...... ix

List of Tables...... ………………………...... x

List of Appendices...... ……………...... xi

1. Introduction...... 1

2. Geologic Background...... 4

3. Regional Climate...... 6

4. Field Sites...... 10

4.1. Site selection...... 10

4.2. SPN site...... 10

4.3. Tatio site...... 12

5. Literature Review...... 14

5.1. Overview...... 14

5.2. Lake and salar basins...... 14

5.3. Hydrologic budget...... 16

5.4. Glacial history...... 17

6. Methods...... 21

6.1. Cosmogenic surface exposure dating...... 21

6.1.1. Production & scaling...... 22

6.1.2. Limitations of surface exposure dating...... 24

v 6.2. This study...... 26

6.2.1. Sample locations...... 28

6.2.2. Sampling methodology...... 28

6.3. Bulk rock 36Cl sample preparation...... 29

6.3.1. Physical preparation...... 30

6.3.2. Chemical preparation...... 30

6.3.3. Carrier/Isotope dilution...... 31

6.3.4. Dissolution & procedural blank preparation...... 31

6.3.5. Precipitation...... 32

6.3.6. Anion exchange chromatography...... 33

6.3.7. Final sample preparation...... 34

6.3.8. Target preparation for AMS measurement...... 34

6.4. TCN age calculation...... 35

6.5. Geomorphological mapping methods...... 37

6.6. Numerical modeling methods...... 42

7. Results

7.1. El Tatio region mapping...... 43

7.1.1. Tatio site mapping...... 45

7.2. Puntas Negras region...... 48

7.2.1. SPN site mapping...... 50

7.3. Mapping interpretations...... 53

7.4. Cosmogenic results...... 55

7.4.1. SPN ratios...... 55

7.4.2. Tatio ratios...... 57

vi 7.4.3. SPN apparent ages...... 57

7.4.4. Tatio apparent ages...... 58

7.5. Apparent age interpretations: SPN & Tatio...... 60

7.5.1. Scenario 1...... 63

7.5.1.a. SPN...... 63

7.5.1.b. Tatio...... 64

7.5.2. Scenario 2...... 65

7.5.2.a. SPN...... 65

7.5.2.b. Tatio...... 65

7.5.3. Scenario 3...... 66

7.5.3.a. SPN...... 66

7.5.3.b. Tatio...... 67

7.6. Numerical modeling results...... 68

7.7. ELA interpretations...... 70

8. Discussion...... 72

8.1. Regional glacial trends: Northern Altiplano...... 72

8.2. Tatio site chronology: Cosmogenic ages...... 73

8.3. SPN site chronology: Cosmogenic ages...... 75

8.4. Paleo-ELA plots...... 77

9. Conclusions...... 79

References...... 81

Appendix...... 96

vii

viii LIST OF FIGURES

Figure 1: Global atmospheric circulation cells and climatic features affecting South America....2

Figure 2: Moisture loss to the southwest AD between Tauca Phase and the Lateglacial...... 3

Figure 3: Regional DEM and morpho-tectonics of study sites in the Central Andes...... 4

Figure 4: South Puntas Negras field site...... 7

Figure 5: AD modern climatic conditions and glacially-delineated AD boundaries...... 11

Figure 6: How surface processes affect cosmogenic apparent ages...... 25

Figure 7: Field images of boulder sampling...... 29

Figure 8: Examples of landforms in the field...... 41

Figure 9: El Tatio regional glacial map and elevation profile...... 45

Figure 10: Map of glacial stages and sample locations from El Tatio field site ...... 46

Figure 11: SPN regional glacial map and elevation profile...... 49

Figure 12: Map of glacial stages and sample locations from SPN field site...... 50

Figure 13: Map of glacial stages and sample ages from SPN field site...... 58

Figure 14: Map of glacial stages and sample ages from El Tatio field site...... 58

Figure 15: TCN apparent ages and errors plotted with lake curve from Laguna Miscanti...... 61

Figure 16: Probability density functions or camel plots for Tatio and SPN glacial stages...... 62

Figure 17: Locations of each formerly-glaciated site used to calculate paleo-ELAs...... 69

Figure 18: Numerical modeling paleo-ELA distance plots sites...... 70

Figure 19: Regional synthesis of paleo-ELA and projected moraine ages...... 71

ix LIST OF TABLES

Table 1. Boulder sample properties...... 56

Table 2. Sample apparent ages...... 59

x APPENDICES

Appendix A.

Table 3. Raw AMS data SPN site...... 103

Table 4. Raw AMS data Tatio site...... 104

xi 1. Introduction

Studying paleoclimatic transitions is crucial to understanding the impact and extent of modern climate components during periods of environmental change. Glaciers are excellent tracers of climatic conditions (Oerlemans, 1991; Lowell, 2000); thus, preserved palaeoglacial deposits provide an integral archive as a proxy for conditions able to sustain patterns of prior glacial occupation. A more thorough understanding of patterns, extent, and timing of glaciation in Chile are pivotal to understanding climatic events from the Southern

Hemisphere.

The South American Arid Diagonal (AD) is a NW-SE-trending climatic feature in the arid Andes between 18-27° S. This zone of hyper-aridity physically separates the northern

Andes, which receive summer Atlantic moisture, from the southern Andes, which receive winter westerly moisture from the Pacific. Within the AD climatic feature is a zone with no evidence of prior glaciation from 24-27° S, which represents the modern glacially-delineated AD. It is clear from the glacial deposits that the AD has not always been this dry, which begs the question: when was this region last glaciated?

For the last 30 years, the majority of glacial chronologies from South America have been produced in the tropical northern Andes and southern Andes near the Chilean lake district and

Patagonia. In the modern climate, these regions are affected by the Northern and Southern

Trades and Westerlies, respectively. Work in this area has resulted in understanding primary glacial sensitivity (Sagredo and Lowell, 2012) (Figure 1B) as a lens to interpret levels of temperature and precipitation changes supporting observed glacial patterns. It is, however, unclear whether glacial deposits adjacent to the Arid Diagonal (AD) spatially or temporally

1 correlate across the hyper-arid Central Andes, where there is a marked paucity of hydrologic proxies in the literature.

There was little to no age constraint in the

southern Puna plateau portion of the AD until a study

from the Chajnantor Plateau (23°S) added a valuable

new data set which, at the time, was the southernmost

data set from the northern edge of the AD (Ward et al.,

2015). Results from Ward et al., 2015 indicated

widespread glaciation during Tauca phase (15-20 ka),

named after the largest Altiplanic lake highstand from

the Lake Titicaca basin. During the Lateglacial (13-10

ka), however, there are glacial deposits in the

subtropics, but they disappear southwest of ~21° S

latitude (Figure 2). Adjacent field sites exhibited

evidence of glaciation both north and eastward of the

Chajnantor Plateau (Zech et al., 2009; Blard et al.,

2014) during the Lateglacial. Figure 1A. Global atmospheric circulation cells. 1B atmospheric and oceanic circulation cells affecting field Given the primary moisture-limiting sensitivity sites and seasonal climatic features responsible for seasonal gradient in precipitation patterns. Colored and numbered polygons represent primary glacial in this region, further examination of an apparent sensitivities. Number 3 in pink is affected by Atlantic precipitation and is both temperature- and precipitation- sensitive. Group 4 & 5 in green and blue are subtropical spatiotemporal trend of moisture loss southwest of and mid-latitude glacial groups. These groups are precipitation-sensitive. Yellow markers are cosmogenic glacial records from El Tatio (north) and Nevado de 21°S is the motivation for this study. The goal of this Chañi (east). The red marker is the south Puntas Negras site (SPN)

2 project was to develop a

new cosmogenic glacial

chronology near the

northern boundary of the

AD climatic feature and

south of the Chajnantor

plateau where there are no

current glacial records. Figure 2: Apparent loss of moisture delivery to support glaciation between Tauca Phase (left) through the Lateglacial (right) SW of ~20°S. Orange and blue dashed lines indicate modern AD climatic feature. The red polygon represents the boundaries of the glacially-delineated AD with no This new data set, along preserved evidence of prior glaciation. Yellow markers are glacial chronologies published in other studies and red indicates glacial chronology from this study. The Chajnantor plateau chronology is represented with a star. Figure modified from Ward et al., 2015. with modeling efforts, will shed light on the hypothesis that the glacially-delineated AD expanded and/or migrated to the northeast, encompassing the Western Cordillera of the central Andes during the last glacial occupation of the AD.

3

2. Geologic background

The Andes, along the western margin of South America, are composed of semi- continuous mountain ranges and volcanic complexes spanning 7000 km. The Central Andes

(21°- 25°S) are characterized by distinct structurally-controlled physiographic units resulting from subduction of oceanic crust from the Pacific plate under the continental South American plate since the Jurassic (Figure 3).

Figure 3: Regional digital elevation map of study sites in the Central Andes Along the Pacific coast, these morphotectonic units include the Coastal Cordillera and

Central Depression which comprise the Andean forearc. Arc magmatism originated in the

Coastal Cordillera during the Jurassic and migrated inland and eastward to the Western

Cordillera throughout the Miocene-Quaternary from continuous subduction.

4 The Altiplano-Puna plateau is bounded by the Eastern and Western Cordillera of the

Andes. The surface of the plateau is characterized by ignimbrite flows, dacite domes, dacite- andesite composite stratovolcano chains, and basalt-andesite scoria cones. This central portion of the Andes from 14°-28° S, the Central Volcanic Zone, was built from volcanism beginning as early as the Upper Oligocene to recent (de Silva 1989; Isacks 1988; Matteini et al., 2002).

Average elevation is ~4000 m across the arid Altiplano-Puna with volcanic peaks as high as

6000 m.

Crustal shortening trending WNW-WSE and extension trending ENE-WSW created

Basin-and-Range-like surface relief. The intersection of extension and NW-SE transversal lineaments led to a transtensional regime opening pull-apart basins along the Central Puna. The topographic lows associated with both types of deformation are the modern N-S aligned salares, which are characteristic features of the modern landscape. Uplift of the Eastern flank of the

Andes from 14 to 9 Ma is thought to be the mechanism for drainage separation from the forearc and establishment of Altiplano-Puna internal drainage ~15 Ma (Vandervoort et al., 1995;

Hartley, 2003).

5 3. Regional Climate

The Andes range intersects regional and global atmospheric circulation cells (Figure 2A).

The central Andes are located at the sub-tropical zone of high pressure between the south Hadley circulation cell and the south Ferrel Cell; as such, they fall between two different precipitation regimes which exhibit steep east-west and seasonal gradients. Mean annual precipitation (MAP) is highly variable with respect to latitude due to these differing precipitation sources and the topographic modification of moisture distribution (Vuille 1999; Garreaud 1999). As such, the range is separated into broad climatic sections: the tropical Andes, the arid central Andes region containing the Atacama Desert, and the southern Andes.

The central Andes remain dry due to the upwelling Humboldt current off the west coast of Chile, the descending arm of the S Hadley circulation cell, and the orographic effect across the

Andes range. Glaciers cannot be sustained in the arid Andes despite peaks exceeding 6000 m due to precipitation levels under 200 mm/yr (Betancourt et al., 2000; Clapperton 2000; Ammann et al., 2001; Vuille et al., 2012). The climate of the AD is also affected inter-annually by the South

American Seasonal Monsoon and El Niño/Southern Oscillation (ENSO) circulation (Garreaud,

R.D. et al., 1999; Kanner et al., 2012; Baker and Fritz, 2015).

North of the AD, easterly precipitation reaches a maximum of 400 mm/y at 18°S and decreases linearly from 18° to 20°S with a lapse rate of 100 mm/y per degree of latitude. At approximately 23°S, precipitation rapidly plunges to a minimum of 100-200 mm/yr from 23-

26°S (Ammann et al., 2001). Average elevations within the AD are ~4000 m and mean annual temperatures ~4°C (Kull and Grosjean 2000); therefore, the high peaks intersect the 0°C isotherm, a condition previously described as “thermal readiness” (Messerli, 1973). Despite continuous permafrost (Grenon, 2007; Kull and Grosjean, 2000; Amman et al., 2001), however,

6 there is no modern glaciation from 18°S to 27°S (De Martonne, 1934; Houston and Hartley,

Zech et al., 2008) with the exception of small snowfields and rock glaciers (Azócar & Brenning

2010).

Figure 4: Region of the AD affected by the easterly trades during austral summer is shaded in orange beginning at the equator and tapering off around 18°S latitude. Region affected in winter by the westerlies shaded blue starting at ~28°S latitude. The 0° isotherm is plotted on the figure and is continuous across the arid diagonal as it intercepts the topography throughout the region. Precipitation data is plotted in the center across the AD, illustrating the average precipitation around 200 mm/year. This moisture-limitation is why snowline disappears across the AD and boundaries of the AD are marked by orange and blue vertical lines. Field sites are plotted at their latitude within the AD. The yellow markers are El Tatio field site and Nevado de Chañi, which were not visited during the 2016 field season. The SPN site is marked in red as this field site was visited during austral summer in 2016. Figure modified from Ammann et al., 2001. The tropical Andes receive Atlantic-sourced moisture from the easterly trade winds of the

Intertropical convergence zone (ITCZ), which are displaced southward during austral summer

(DJF). During this time, components of the South American Seasonal Monsoon (SASM) (Zhou and Lau, 1998; Vera et al, 2006) direct easterly moisture to the eastern Andes and southern

Altiplano (Garreaud et al., 2009; Lenters and Cook, 1997). A continental low over the Gran

Chaco region in Argentina (~25°S) forms due to transport of Atlantic moisture onto the continent. Convective moisture from the eastern lowlands and Amazon basin is diverted southward via a low-level jet (Seluchi et al., 2003). Upper-level (300 hPa) anticyclonic circulation resulting from latent condensational heat forms the “Bolivian High” (BH), (Lenters

7 and Cook, 1997) which transports monsoonal precipitation southward to the northern Puna and portions of the arid Andes. In summary, strengthened and southward migration of the BH increases moisture flux to the northern AD. The opposite (northward migration and decreased moisture flux) occurs when the BH is weakened. The BH system is modulated by the SASM and continental convection, which is closely tied to both Atlantic and Pacific convergence zones, or

ITCZ.

In addition to the annual anticyclonic system described above, the central Andes are affected by phases of El Niño/Southern Oscillation (ENSO). ENSO drives precipitation and atmospheric circulation heterogeneity on inter-annual timescales (Vuille, 1999; Garreaud and

Aceituno, 2001). Using daily precipitation and reanalysis data, Vuille (1999) demonstrated that

ENSO significantly decreases summer precipitation over the Altiplano.

The Southern Andes receive advective extra-tropical moisture sourced from the Pacific and transported by westerly circulation such as the South American Westerly Wind Belt. The

Altiplano receives very little precipitation from May to October due to the maximum northwardly-displaced position (27°S) of the subtropical westerly jet stream, typically focused at

~50°S (Zech et al., 2007). Upwelling along the western margin of South America associated with the Humboldt current limits evaporation and transport of Pacific moisture onto the continent.

Prevailing dry, cool westerly air flow collides with easterly winds in the upper troposphere, limiting convective precipitation on the eastern flank of the Andes during austral winter months

(JJA). Rare winter and inter-seasonal precipitation in the central Andes have been attributed to extra-tropical polar cold fronts or cut-off low pressure systems from the westerly wind zone.

These produce short-lived but significant precipitation events (Vuille 1999).

8 In summary, hemispheric (meridional) and tropical-subtropical pressure gradients are key regulatory components of the easterly (wet) – westerly (dry) seasonal climate dynamic in this region (Garreaud et al., 2003).

9 4. Field sites

4.1. Site selection

Two sampling sites were chosen to analyze the timing and extent of glaciation across the northern edge of the AD. The northernmost site El Tatio and a valley from the southern Cordón de Puntas Negras (SPN). The Tatio site, near El Tatio geyser field, is located at 23°44’ S, 67°46’

W, approximately 3° of latitude further north from SPN. Although the Tatio moraines had not been dated, they were reported to be late glacial in age by earlier studies which made the assumption that they were concurrent with Altiplanic lake phases (Jenny et al., 1996; Kull and

Grosjean, 2000; Glasser et al., 2009).

4.2. SPN site

The SPN valley is the southernmost formerly-glaciated valley of the Puntas Negras stratovolcano chain (23.8°S, 67.5°W). The Cordon de Puntas Negras, approximately 40 km south of Lascar volcano, are the northern-most chain of stratovolcanoes associated with the Upper

Cenozoic, NW-SE-trending Calama-Olacapato-El Toro (COT) lineament. Here, Quaternary eruptions produced basaltic-andesitic lava flows and scoria cones which are characteristic of this portion of the Puna Plateau.

The location of the SPN site was chosen foremost for the preservation of multiple stages of glaciation identifiable by aerial photography. SPN is also adjacent to El Laco mining facility with infrastructure such as access roads and off-road parking to ease sampling in this remote location. The SPN site is a relatively narrow valley with an impressive expression of nested lateral, frontal, and ground moraines (Figure 5). It is located 30 km from the Quebrada

Nacimiento fault which bounds three endorheic lake basins: Laguna Miñiques, Miscanti, and

10 A B

Figure 5: A. Overview of South Puntas Negras (SPN) field site. B: Lower left photo from Laguna Miscanti on the western end of the SPN site. C: Lower right photo from the formerly-glaciated valley on the southeast end of the SPN site. Preliminary glacial stages are marked on this photo. The red is the relatively older glacial stage I. The blue is the relatively younger glacial stage II.

Pampa Varela (Ramirez and Gardeweg, 1982; Grosjean et al., 2001). The valley is 12 km from

Laguna Miscanti (23°44’ S, 67°46’ W), a saline, alkaline (pH 8 - 8.8) lake 10 m in depth that has been cored and dated using radiocarbon (Grosjean et al., 2001). The presence of Miscanti allowed lake stages and glacial stabilizations to be temporally compared to test the hypothesis that these moraines and lake highstands occurred synchronously.

11 4.3. Tatio site

The Cordillera del Tatio is a stratovolcano chain built predominantly from interbedded

Late Cenozoic-to-modern volcanic peaks, lava domes, and expansive dacitic ignimbrite sheets

(Davidson et al., 1995). Topography of this portion of the arid Andes is affected by the Altiplano

Puna Volcanic Complex (APCV) (de Silva, 1989b; Pritchard and Simons, 2002; Lucchi et al.,

2009a,b). The APCV is highly uplifted and bordered by the thick-skinned fold-and-thrust Eastern

Cordillera to the east and an additional thrust system, the West-Vergent Thrust System to the west (Lucchi et al., 2009). Today there is active magmatism here, responsible for the Tatio geothermal field and relatively young dacitic domes such as Co. La Torta (34 ka). This volcanic arc region is bounded to the west by the Preandean Depression with several N-S striking tectonic basins such as the Salar de Atacama from extensional tectonic activity.

The Tatio site, near El Tatio geyser field, is located at 23°44’ S, 67°46’ W, approximately

3° of latitude further north from SPN. Although the Tatio moraines had not been dated, they were reported to be late glacial in age by earlier studies which made the assumption that they were concurrent with Altiplanic lake phases (Jenny et al., 1996; Kull and Grosjean, 2000;

Glasser et al., 2009).

A recent 10Be chronology from Nevado de Chañi in Argentina at ~ 24-25°S by Martini et al., 2017 was compared with the southern Puntas Negras and Tatio records from this study. The record from Chañi is approximately the same latitude and altitude as the SPN site but located in the Eastern Cordillera. The addition of the Martini 2017 study enabled the evaluation of regional trends in glaciation from the eastern and western flanks of the central Andes. The Tatio, SPN, and Chañi sites envelop the northern boundary of the AD, which is where an apparent threshold of moisture delivery may have existed during the last glacial occupation of the region (Ward et al., 2015).

12

13 5. Literature Review

5.1. Overview

Moisture variability and changes in precipitation gradient near and within the AD have been studied using numerous hydrologic budget proxies. Sediment cores and shoreline deposits from lake and salar basins (Minchin, 1882; Servant and Fontes, 1978; Servant et al., 1995;

Sylvestre et al., 1999; Plackzek et al., 2006a,b; 2009a,b, 2011) are often correlated with glacial records and hydrologic modeling (Hastenrath and Kutzbach, 1985; Blodgett et al., 1997) to identify periods of increased precipitation or aridity. Glacial chronologies are paired with these proxy data to study trends in relative effective moisture variability, as glacial geometry is affected by precipitation and temperature (Ohmura et al., 1992). Lacustrine records from this region yield continuous paleohydrologic data exceeding 100 ka and are used to construct time series analyses to correlate the southern hemisphere expressions of north Atlantic climatic events such as the Younger Dryas and Heinrich events. Reconstructions from the Altiplano reveal that the AD may not have been as arid throughout the Pleistocene as it is today. The AD may have migrated or contracted/expanded through the Pleistocene/Holocene transition to include portions of the Western Andes around the time of deglaciation. Below, I review the key literature and summarize existing knowledge of paleoclimate variability in the AD and adjacent Altiplano.

5.2. Lake and Salar Basins

Between 14° and 22°S on the Altiplano, there are four expansive internally-drained basins: Titicaca (16°S, 69°W), which is separated from the Poópo (18.5°S, 67°W) basin by the

Río Desaguadero, Coipasa (19.5°S, 68°W), separated from the Poópo by the Laka sill, and Uyuni

(20°S, 67°W). Fresh-water Lake Titicaca occupies 8560 km2 of the northern Altiplano with a depth of 285 m (Argollo and Mourguiart, 2000; Placzek et al., 2013). On the southern Altiplano,

14 Lake Poópo is 2530 km2. The Coipasa and Uyuni salar basins on the southern Altiplano are separated by a mere 1m topographic high in the modern climate; therefore, during past lake transgressions, they have overflowed to connect with a total area of ~12,100 km2 (Argollo and

Mourguiart, 2000; Placzek et al., 2013). Two cores were collected from the Salar de Uyuni

(Risacher and Fritz, 2000; Fornari et al., 2001; Baker et al., 2001a; Fritz et al., 2004), both containing alternating salt and mud layers, salt indicating desiccation, mud indicating expansion, respectively. This evidence along with shoreline deposits exposed the current lake levels and salt flats indicate that moisture delivery to the Altiplano/Puna region has fluctuated in geologic time

(Minchin, 1882; Servant and Fontes, 1978; Sylvestre et al., 1995; Sylvestre et al., 1999; Placzek et al., 2006a,b; 2009a,b, 2011). The amalgamation of shoreline deposits, ages from Uranium- series geochronological methods used on sediment cores, natural gamma data, and crosscutting landform observations were used to re-construct lake levels.

Pioneering studies of the Altiplanic lakes focused on shoreline deposits as evidence for the largest lake highstand, the Tauca highstand from 17-14 ka and utilized hydrologic budget modeling to estimate lake surface area and volume (Servant and Fontes, 1978; Hastenrath and

Kutzbach, 1985). Re-evaluation of U-Th and Ar/Ar dates from two cores collected from the

Uyuni-Coipasa basin (Fornari et al., 2001; Risacher and Fritz, 2000; Baker et al., 2001a; Fritz et al., 2004) and coupling with shoreline tufa deposits (Placzek et al., 2006a,b) allowed identification of six lake oscillations from a unified chronology (Placzek et al., 2013). Lake highstands include: Ouki (125-95 ka), Salinas (95-80 ka) which the authors state could have been a separate phase or perhaps was part of the Ouki phase, Inca Huasi (50-40 ka), Sajsi (25-20 ka),

Tauca (17-14), and Coipasa (13-11 ka). The lake stages were corroborated by both natural gamma and salt content data from the core analyzed by Fornari et al., 2001; moreover, among the Late Pleistocene lake phases, only Ouki and Tauca were deep lakes. Lake Tauca is estimated

15 to have reached a maximum depth of 140m during highstand (Bills et al., 1994; Baker et al.,

2001a). Baker et al. (2001a) attributed the wet periods from the Salar de Uyuni to maxima in summer insolation. This group suggested that land-sea temperature gradients were increased compared to the modern climate, enhancing the SASM and bringing more easterly-sourced precipitation to the tropics in South America.

Further south on the Altiplano of northern Chile (23°S, 68°W), a salt core was drilled from the Salar de Atacama from the western end of the central Andes. Bobst et al., 2001 used sedimentary structures and evaporite textures to study net hydrologic balance and identify wet/dry periods. They determined that from 75.7-60.7 ka, 53.4-15.3 ka, and 26.7-16.5 ka were wet periods. Moreover, the largest, wettest lake stage was from 26.7 to 16.5 ka. Brief phases of expansive mudflats rather than lake stages also occurred in the Holocene from 11.4 to 10.2 ka and from 6.2 to 3.5 ka. These wet periods match hydrologic budget studies from the Bolivian

Altiplano with the wettest cycle (26.7 to 16.5) overlapping the Tauca phase. The early Holocene stage (11.4 to 10.2) is synchronous with the Coipasa phase.

5.3. Hydrologic budget

Hydrologic budget and balance modeling have provided estimates of mean annual precipitation (MAP) and temperature changes necessary to support field observations. These models follow the basic premise that in this closed basin, evaporation must not outpace precipitation during a lake transgression.

Hastenrath and Kutzbach (1985) concluded, by using deposits estimating the former surface of paleolake Tauca, that a 300 mm/yr increase in precipitation must have occurred to support the Tauca highstand. Similarly, in his 2001 paper, Grosjean estimated precipitation rates of >500 mm (an increase of ~300 mm relative to modern) occurred at 23°S. These estimates are

16 in agreement with glacier modeling done in the valley sampled at the Tatio site by Kull and

Grosjean (2000) suggesting precipitation rates must have been 515±45mm/yr in the northern

AD. These modeling studies utilized the paleolake deposits to determine P/E relationships that match the paleo-shorelines, estimated depth, and estimated heat balance. These studies, along with the “thermal readiness” (Messerli 1973) concept, illustrated in Figure 4, provided support to accept assumptions of moisture sensitivity to glaciation during the late Pleistocene, similar to modern AD conditions. It has since been an assumption from field- and model-based support that glacial moraines from the northern AD are approximately Lateglacial in age.

Some of the first researchers to use absolute dating methods on deposits from Lake Tauca once thought meltwater filled Lake Tauca with highstands occurring post-deglaciation (Servant and Fontes, 1978). This was later considered implausible based on calculations of the LGM glacial ice volume falling far short of the estimates for total volume of the maximum extent of

Lake Tauca (Hastenrath and Kutzbach, 1985; Blodgett et al., 1997). This exemplifies the value of combining field and hydrologic modeling studies to evaluate existing hypotheses about lake and glacial stages assumed to be contemporaneous.

5.4. Glacial History

As earlier qualitative work gave way to technological advancements and additional fieldwork in this remote area, evidence suggested glaciers may have re-advanced or reached maximum extent at the same time as the Tauca phase (~14-17 ka BP) (Seltzer, 1992; Clayton and Clapperton, 1997; Ammann et al., 2001; Placzek et al., 2013). Early modeling by Hastenrath and Kutzbach, 1985 suggested an increase in summer precipitation may have been the forcing responsible for glacial maxima and lake transgressions occurring synchronously. This hypothesis is supported by correlations of terminal moraine complexes in similar configurations from the

17 Andean Eastern Cordillera in Southern Peru (Seltzer, 1990), Bolivia, as well as the northern boundary of the AD between 18 and 25°S (Jenny et al., 1996). Additionally, Salar de Uyuni and the Salar de Atacama Tauca and Coipasa phases also align with high accumulation and low d18O values from the Sajama ice core (Thompson et al, 1998) from southwest Bolivia (18°S).

A data set describing the chronology of the Tauca phase from erosional and depositional shoreline deposits surrounding the Uyuni-Coipasa basin (Servant and Fontes, 1978) was produced by Sylvestre et al., 1995. This study also noted a terminal moraine complex of three stages at roughly 15, 9, and 6 km, respectively, from modern valley glaciers of the Andean

Eastern Cordillera. Moraines in a similar configuration were also observed in Southern Peru

(Seltzer, 1990) as well as the northern boundary of the AD (18-25°S) (Jenny et al., 1996). These studies qualitatively provided support to accept the assumption of humidity-limitation to glaciation during the late Pleistocene, similar to modern AD conditions.

Additional support for the tentative Lateglacial age assignment (Jenny et al., 1996) was found after stages were further described, correlated via geomorphological descriptions and ELA reconstructions (Clayton and Clapperton, 1997), and paired with minimum radiocarbon ages

(Amman et al., 2001). The oldest and most degraded stage, “N-I” from Amman et al., 2001 relatively pre-dated an intermediate stage “N-II”, characterized by high-relief, sharp-crested lateral moraines and observed as far south as 25°S, linked most probably to “Advance 3” from

Clayton and Clapperton, 1997.

Clayton and Clapperton inferred that their “Advance 3” was synchronous with Tauca phase based initially on fan-deltas of glacifluvial outwash associated with Advance 3 building into the paleolake Tauca basin. The fans lie below the altitude of the highest paleoshoreline, implying a glacial advance followed by a lake level rise. Advance 3 glacifluvial outwash fan- deltas were built into the basin of paleolake Tauca, which imply they respectively advanced and

18 transgressed synchronously. The “N-II”/Advance 3 correlation assigns a radiocarbon minimum age of 13,300 14C yr BP from peat at the base of a terminal moraine, placing the age of “N-

II”/Advance 3, and local LGM at this latitude in the Late Glacial (Clayton and Clapperton, 1997;

Sylvestre et al., 1999; Blard et al., 2009; Placzek et al., 2013).

Clayton and Clapperton re-constructed equilibrium line altitudes (ELA) from the N-

II/Advance 3 moraine using maximum altitude of lateral moraines (Meierding, 1982) from valley glaciers where lateral moraines were well-preserved with clearly-delineated limits. ELA for advance 3 or N-II, or the Tauca phase moraine was estimated at 4435±50 m for Co. Anzanaques and 4525±50m for Co. Tunupa. The authors cite the 90m difference as evidence suggesting

Advance 3 regional ELA exhibited a gradient which parallels the maximum snowline altitudes from Gosse et al., 1995; implicating a NE or Atlantic source of humidity. The modeling efforts of Clayton and Clapperton, 1997 estimated 3 possible precipitation and temperature configurations (relative to modern values) necessary to support the geometry and ELA of

Advance 3. They include: (1) a 4.5-7°C temperature depression with no change in precipitation

(2) a 300 mm/yr increase in precipitation, suggested initially by Hastenrath and Kutzbach, with a temperature depression of 3-5°C, or (3) an increase by 600 mm/yr in precipitation, as suggested by Grosjean (1994). The 600mm/yr with a temperature depression of 1.8-3.6°C match an interpreted 15-25 m rise in paleolake Lejía (23°S, 67°W, 4325m), bringing the surface to 9-11 km2 as opposed to just 2 km2 today. Laguna Lejía is considered representative of the Altiplano area from 21-24°S.

The Altiplano lake phases from the Salar de Uyuni not only overlap wet phases from the

Salar de Atacama sediment core, they also match glacial moraine stabilizations north of ~20°S and south of ~40°S. There is a bias in available data sets, though, as Bobst et al., 2001 suggest.

Most paleoclimate records from the Altiplano are less than 100 ka and the majority of records are

19 from the northern tropical or southern Andes. There is little data from the arid, central Andes where there is no modern glaciation and the majority of lakes are evaporite-encrusted pans. One of the most recent studies from this region was from the Chajnantor plateau, which is located between the Tatio and SPN field sites at ~22.5°S. Samples were collected and dated using cosmogenic 36Cl from the peaks of the Chajnantor Plateau surrounding the Atacama Large millimeter/submillimeter Array (ALMA) observatory. New 36Cl ages from Ward et al. (2015) were reviewed in context with proximal previously published glacial records (Figure 2). Results from this study indicated glacial stabilizations at 25-40 ka, 15-17 ka, and 12-14 ka. The youngest of these moraines, however, were absent south of ~20°S, which implicates a loss of enough moisture delivery to support glaciation south of 20°S within the climatic AD.

Pairing glacial moraine ages with the age constraints on the salar and lake basins is where the research evolved in the tropical portion of the range, but as glacial records become scarce toward the AD, this coupling becomes more challenging.

6. Methods

6.1. Cosmogenic Surface Exposure Dating

Advances in techniques such as terrestrial cosmogenic nuclide (TCN) exposure dating have improved the quantitative study of Earth-surface processes and rates of landscape evolution. We are now able to study a range of spatial and temporal scales which were not possible just decades ago (National Research Council, 2010). Directly dating exposure times of young volcanic rocks, geomorphic surfaces, and soils is a powerful tool for investigating the timing and rates of surface processes. In situ TCN concentration studies have contributed vital chronological constraints for regional and global paleoclimatic reconstructions (Gosse et al.,

20 1995; Ivy-Ochs et al., 1999; Owen et al., 2003; Liccardi et al., 2004; Douglass et al., 2006). This dating method is based on the influx of cosmic radiation, or particles (predominantly protons and alpha particles; 83 and 13 percent, respectively), at energy levels sufficient (~1 GeV < E < ~1010

GeV) to induce reactions with Earth’s upper atmosphere, producing secondary particle showers.

Resultant secondary cosmic rays interact with target atoms in minerals at and within the first ~3-

5 meters of the Earth’s surface. Past this depth, cosmic rays are attenuated with depth due to successive energy loss, described by Beer’s law. The product of spallation reactions of target minerals rare isotopes, in situ cosmogenic nuclides, which are diagnostic of surface exposure.

Six terrestrial cosmogenic nuclides (3He, 10Be, 14C, 21Ne, 26Al, and 36Cl) are commonly used for geologic applications due to: (1) lack of natural occurrence from other processes, making radiogenic origin clear (2) reasonable decay rates relative to durations of surface Earth processes being investigated (3) abundance of target mineral (4) production rate is high enough for them to be measured within detection limits of current analytical methods. In principle, the time since exposure of a sample can be determined by measuring the net concentration of cosmogenic nuclides in a sample (accounting for variation in cosmic radiation flux, decay of the nuclide, and erosion) and dividing it by the rate of production. (Gosse and Phillips, 2001).

6.1.1. Production & Scaling

Cosmogenic isotopes are produced by a few different types of reactions including: spallation, secondary thermal and epithermal neutron capture, and muon-induced reactions (Lal

& Peters 1967, Lal 1988a). Spallation reactions involve collision of secondary particles possessing sufficient energy levels to shatter nuclei of target minerals, which produce rare daughter nuclides with smaller mass relative to the parent element. Muon capture reactions involve negatively charged muon particles occupying the electron shells of target atoms, where

21 they are subsequently captured by the nucleus to produce daughter cosmogenic nuclides.

Relatively less apt to interact with atoms than other particles, muons maintain energy levels suitable for nuclide production (attenuation length approximately 1500 g cm-2; Brown et al.,

1995a); thus, are a more prominent production source at depth (Heisinger et al., 1997; Stone et al., 1998b). Thermal neutron absorption reactions occur due to energy discrepancy with other particles in the radiation flux; therefore, low energy thermal neutrons are effectively suspended between paths of particles with higher energy and follow paths described by Brownian motion until target nuclei uptake them to produce cosmogenic nuclides. Contrary to muonic production, thermal neutron absorption occurs more readily at and proximal to the Earth’s surface (Kurz,

1986b; Fabryka-Martin, 1988). One or more of these production mechanisms may contribute to the production rate of an individual nuclide.

Accuracy of TCN dating is dependent, in part, upon appropriate estimations of regional production rates. Rates of nuclide production are affected by temporal variability in the strength of the geomagnetic field as well as solar modulation of cosmic-ray flux (Lal & Peters, 1967;

Kurz et al., 1990; Lifton et al., 2008). A weaker magnetic field increases the production of cosmogenic nuclides due to greater penetration of cosmic rays into the atmosphere. Increased solar activity decreases the cosmic ray flux to Earth’s atmosphere, particularly at high latitude and altitude. Production rates are additionally affected by site-specific variables including: latitude, atmospheric pressure, and compositional properties of target material (Lal, 1988a; Lal

1991). To calculate an age constraint on an exposure event as a function of time in years (a), an estimated local production rate must be used (reported in units of atoms g-1 a-1) which accounts for the processes and variability described above.

22 Numerous scaling schemes have been developed to incorporate time and spatial variability of TCN production (Lal, 1991; Stone, 2000; Dunai, 2000; Dunai, 2001a; Desilets and

Zreda, 2003; Pigati and Lifton, 2004; Farber et al., 2005; Lifton et al., 2005; Lifton et al., 2014).

The scaling scheme based on observations of Lal & Stone, abbreviated “St” in the review by

Balco, 2008 is most prevalent in the literature as this was the first and only scaling scheme for over a decade. This scheme scaled as a function of latitude and atmospheric pressure (Stone,

2000). Production rates of nuclides were assumed constant over time; thus, temporal variability of the geomagnetic field was not incorporated into this scaling scheme. The scaling schemes of

Dunai 2000 “Du” and Desilets et al., 2006 “De” incorporate magnetic field modulation of production. Due to additional data availability of cosmic ray flux near the surface from additional calibration locations, the elevation-dependence of production was modified to fit modern observation data more closely. The Lifton et al., 2005 “Li” model added temporal solar variability. The “Lm” model was produced to retain altitude scaling from Lal 1991 with the addition of a paleomagnetic correction by Nishiizumi et al., 1989. Fully described in Lifton et al.,

2014, the Lifton-Sato-Dunai scaling model incorporated analytical parameterizations, allowing for calculation of cosmic ray spectra based on first principals of physics rather than empirical fit models. Temporal variability in production rates from solar and geomagnetic modulation is also retained in this new scaling framework. Two versions of the Lifton-Sato-Dunai scaling model were used as part of the CRONUS calculator. The “LSD” denoted “SF” in the calculator is flux- based and the “LSDn” denoted “SA” model contains nuclide-dependent and energy-dependent terms resulting in a nuclide-dependent scaling factor.

Accuracy of empirical measurements used to calibrate scaling schemes as well as increased geographical distribution of calibration data sets will inevitably lead to improvements in scaling methods. Future discrepancies in exposure ages calculated from the same nuclide

23 concentrations are expected due to ongoing advancements of age calculators and scaling schemes; therefore, detailed observations and measurements necessary to recalculate results are reported as legacy data to ensure future applicability of this data set.

6.1.2. Limitations of Surface Exposure Dating

Cosmogenic exposure dating methods are routinely used to constrain minimum or maximum ages of deglaciation by dating landforms deposited during or after ice retreat.

Exposure ages from glacial-geomorphic landforms such as moraine boulders and striated bedrock are assumed to equate to the duration of exposure; therefore, age of moraine or time of deglaciation, respectively (Putkonen and Swanson, 2003; Fabel et al., 2012; Stone and

Ballantyne, 2005; Small et al., 2016). TCN exposure ages are affected by many sources of uncertainty including 2-5% from analytical errors and accelerator mass spectrometry (AMS) and

10-20% from production rate uncertainty associated with the number and distribution of calibration sites and equipment used to measure production variation with latitude and altitude

(Gosse and Phillips, 2001). Finally, there is geological uncertainty, which is difficult if not impossible to quantify. Geological uncertainty of chronological methods was described by Small et al., 2016 as falling into two categories: (1) factors affecting the net concentration before sampling and (2) assumptions involving equivalence, or strength of contemporaneous association between events of interest and the material used for dating. Both types of geological sources of uncertainty are discussed in this section, action to minimize both types of uncertainty are explained in sample methodology section (7.2.2.).

Erosion and weathering processes result in nuclide loss, yielding a younger exposure age than predicted by rates of production. Nuclide loss from partial exposure due to topographic, dust, ash, or snow cover are examples of shielding, an example of the first type of geological

24 uncertainty. Partial exposure of samples due to boulder exhumation or post-depositional toppling are additional ways to calculate erroneously young ages of a target event. The latter examples of partial exposure have the additional component of the second type of geological uncertainty.

Samples from glacially-transported boulders are collected under the assumption that the material had been continuously irradiated since deglaciation with no periods of prior exposure.

Previous events of exposure accumulate TCN concentration that, if not eliminated by erosion or decay before a subsequent exposure event, yields an older exposure age than expected for a duration of irradiation. This excess nuclide inventory from a previous process with respect to the target exposure event is referred to as inheritance; thus, inheritance results in an erroneously old age for the landform or event being dated. Deposition of rockfall- or landslide- sourced boulders from valley walls or topography upstream of the glacial catchment are examples of processes that contribute to inheritance. Inheritance is most common in

allochthonous deposits

with sediment that may

have more complex

exposure histories from

reworking and in

landscapes lacking

sufficient erosion to reset

the sample.

Surface exposure ages

Figure 6: From J Heyman et al., 2011, this figure illustrates how cosmogenic results can vary based of boulders on moraines on surface processes. A: Ideal case with complete shielding, no prior exposure, and no post- depositional modifications which would affect ages. B: prior exposure affecting the exposure ages of boulders making them appear older than age of deglaciation. C: Shielding and erosional and striated bedrock processes make ages appear younger than the age of deglaciation. calculated from a site-

25 scaled production rate, sample properties, and nuclide concentrations theoretically represent a minimum age estimate of deposition/abandonment. For a sample set yielding scatter that is not attributed to inheritance, the maximum apparent exposure age indicates a minimum deglaciation age (Heyman et al., 2011). Geologic uncertainty is incredibly difficult to quantify as the processes described above as issues of equivalence and exposure histories prior to sampling are impossible to confirm; therefore, geologic context must be incorporated into the interpretation of calculated apparent exposure ages (Heyman et al., 2011; Douglass et al., 2006). Figure 6, modified from Heyman et al., 2011, illustrates this concept of how processes including inheritance and partial exposure affect apparent exposure ages.

6.2. This Study

Mafic igneous lithologies from the Cordon de Puntas Negras including andesite, basaltic andesite, and dacite, render 36Cl the most suitable nuclide for this study. Concentrations of in situ cosmogenic 36Cl were measured from moraine boulders to calculate surface exposure ages, or ages of deposition/ice abandonment of glacial landforms. The use of 36Cl accumulation for exposure dating was first suggested over 60 years ago (Davis and Schaeffer, 1955) as the half- life (308,000 years) makes this nuclide of particular interest for studying the Pleistocene epoch.

As such, 36Cl or the stable cosmogenic isotope 3He are often utilized to analyze glacial landforms lacking sufficient crystalline quartz for 10Be dating including mafic volcanics and carbonate-rich lithologies. After advancement in analytical techniques of accelerator mass spectrometry (AMS) measurements, the method was further characterized and reviewed by Phillips et al., 1986 who stressed the hydrophilic nature of chlorine and the use for this mobility in separating in situ from meteoric components.

26 At the surface, 36Cl primarily forms by spallation of 39K and 40Ca and epithermal and thermal-neutron absorption of 35Cl. Spallation production from Ti and Fe is considered insignificant at the surface due to atmospheric attenuation of secondary particles (Zreda et al.,

1991). In the subsurface, muon capture by 40Ca is an important production mechanism due to muon penetration beyond depths of spallation-production attenuation (1-2 meters) (Phillips et al.,

1986; Zreda et al., 1991; Gosse and Phillips, 2001). Whole-rock 36Cl analysis was used for this study, although plagioclase mineral separate analyses may be performed on a subset of samples in the future.

Due to the variety of production pathways, rates of 36Cl production are the most difficult to constrain of commonly measured cosmogenic nuclides (Marrero et al., 2016; Phillips et al.,

2016). One of the most prominent sources of computational error in 36Cl exposure age calculations is the accuracy of the production rate used to convert net concentration measurements into years of irradiation, which is inherently tied to availability of calibration data.

Data from calibration localities exhibit evidence to support the following criteria: minimal erosion, minimal to no shielding or partial exposure, continuous exposure to cosmic rays, and an independent age constraint such as radiocarbon ages. The 36Cl production rates used by the

CRONUS-Earth calculator were calibrated using three sites to constrain Ca and K spallation.

Together the sites yielded a rate of 56.0 ± 4.1 atoms of 36Cl (g Ca)-1 yr-1 and 155 ± 11 atoms of

36Cl (g K)-1 yr-1, scaled using the Lifton-Sato-Dunai “SA” model (Marrero et al., 2016; Lifton et al., 2014). Ages and erosion rates acquired from moraine samples surrounding Quelccaya ice cap in Huancané, Peru (13° S latitude, 4850m elevation) were among the three spallation production calibration sites (Mercer and Palacior, 1977; Kelly et al., 2015, 2012; Phillips et al., 2016;

Marrero et al., 2016). As such, the CRONUS-Earth calculator and the “SA” scaling model were the most appropriate to use for exposure age calculation for this study.

27

6.2.1. Sample locations

Eight glacially-transported andesite boulders were sampled from lateral moraines along the walls of the formerly-glaciated valley at the Tatio site in 2014. During our field season in

2016, twenty-one samples were collected from large glacial erratic boulders at the SPN site. Of the 21 samples from SPN, 2 were collected from the hummocky, degraded terminal moraine, glacial stage 1, GS1). The remaining samples were collected from two additional glacial stages

(GS2b and GS3) up the valley with higher relief relative to GS1.

6.2.2. Sampling Methodology

Several samples were collected from the same moraine ridge, where possible, due to presence of stable boulders acceptable for sampling. Samples were taken via hammer and chisel from the largest boulders in a stable position at the top of moraine ridges, where possible, to minimize the risk of interrupted exposure from till, dust, snow shielding, or toppling. Boulders were chosen for sampling if there was a flat surface with evidence of varnish and ventifaction, indicating little erosion since deglaciation and a stable position since deposition. Field notes were taken for each sample including the dimensions of the boulder, the precise lat/long location, lithology, and sample thickness.

28 Figure 7: Images of our lab group collecting samples from the upper centimeters of Andesite boulders.

6.3. Bulk Rock 36Cl Sample Preparation

Chemical extraction of 36Cl from bulk-rock dissolutions was performed at the University of Cincinnati (UC) Cosmogenic Chlorine Laboratory and purified as AgCl using the procedure given in Stone et al., 1996 with revision per preparatory recommendations from PRIME lab facility at Purdue University (Radler, personal communication). 36Cl was measured with accelerator mass spectrometry (AMS) at Purdue Rare Isotope Measurement (PRIME) Lab,

Purdue University with the addition of two chemical blanks for each unique combination of preparatory reagents and a direct precipitation blank for each carrier solution to correct for reagent contamination. Geochemical analysis of bulk and trace elemental composition of samples was performed at Acme Analytical Laboratories Ltd., now operating as Bureau Veritas

29 Mineral Laboratories in Vancouver, BC, Canada. Exposure ages were calculated using the

Cosmic Ray-Produced Nuclide Systematics on Earth Project (CRONUS-Earth Project) calculator.

6.3.1. Physical Preparation

To acquire the recommended 250 grams of fine grains (<250 μm), samples were first crushed using a Sturtevant 2x6 fixed-plate jaw crusher. Second, each sample was pulverized in a

Bico UD direct drive disc mill, sieved, and sorted by grain sizes (>500 μm, 250-500 μm, and

<250 μm). The jaw crusher, disc mill, and sieves were carefully and thoroughly cleaned of grains and dust between each use. The jaw crusher and disc mill were vacuumed and cleaned with a wire brush and compressed air. Sieves were cleaned using brushes and grain picks. They were also wiped free of dust between each sample. A small hand sample was reserved to measure density of each sample using mass and volumetric water displacement and to archive for future reference.

6.3.2. Chemical Preparation

Two sets of approximately 10 g of <250 μm grains were reserved for whole-rock geochemical analysis. One aliquot was set aside before chemical leaching (described below) and one after. Geochemical analysis of bulk and trace elemental sample composition was performed on the post-leach aliquot at Acme Analytical Laboratories Ltd., now operating as Bureau Veritas

Mineral Laboratories in Vancouver, BC, Canada.

Approximately 100 g of the fine fraction of grains (<250 μm) from each crushed sample were initially leached for over 12 hours using 70mL of 18 MΩ water and 7-8 mL of TraceMetal

Grade (TMG) nitric acid (HNO3) solution in 250 mL nalgene bottles. This process removed

30 potential contamination from atmospheric, or meteoric 36Cl. Leaching solution was decanted, and samples were rinsed with 18 MΩ water to neutralize residual HNO3 and to separate and discard fine particles, then dried overnight.

6.3.3. Carrier/Isotope dilution

After measuring 10 g of post-leached sample and reserving it for geochemical analysis, as mentioned above, ~30g of each sample (exact mass ±0.0001g recorded on lab bench sheets and in sample spreadsheet for archival and working use) was transferred to a 500mL bottle for dissolution and the addition of ~1g of a 35Cl-enriched carrier solution. Adding a spike enriched in isotopically-stable Cl, or the isotope dilution method, has been widely adopted due to the following benefits: (1) ability to measure 36Cl/Cl and total Cl on the same AMS target (2) increase measurement precision for samples with low Cl concentration (3) reduce minimum sample size (4) reduce sensitivity of Cl concentrations to contamination from reagents (Desilets et al., 2006; Sharma et al., 2000; Elmore et al., 1997). Carrier solutions were prepared at the

University of Cincinnati with ICON Isotope Industries 35Cl standard using the approximate ratio of 1.0000 (mg/g) 35Cl. To calculate the amount (mg) of carrier present in each sample, exact concentration of the carrier solution and weight of carrier added to each sample was recorded to ±0.0001 g on bench sheets.

6.3.4. Dissolution & procedural blank preparation

Samples were chemically dissolved in a solution containing 150g of 18 MΩ water,

~150g, or 500% of sample mass, of low chloride 40% hydrofluoric acid (HF) solution, and ~45g, or 150% of sample mass, concentrated TMG 70% HNO3 solution. At this stage, each stock solution and reagent were given a unique identifier. This name was recorded alongside other measurements on the sample bench sheets, which was used to determine the minimum number of

31 procedural blanks to prepare. Each unique combination of HF, HNO3, and carrier spike used on the aliquot of samples was mixed in duplicate pairs and as a procedural blank. Procedural blanks were processed using the same procedure as the bulk-rock samples. Eight procedural blanks were prepared beginning with the addition of ~1 g of carrier spike solution and the solutions for digestion of the rock samples. The blanks serve the purpose of developing a correction factor for chlorine contamination in the reagents used to prepare the sample for AMS measurement.

Samples reacted with the digestion solution to form a gaseous product and fluoride gel; therefore, bottles were vented in a fume hood for 12 hours and occasionally perturbed to ensure the entire sample was exposed to the solution. Samples were subsequently capped tightly and transferred to a heated rolling apparatus for constant agitation at 60°C. The heat ensured complete dissolution and prevented the grains from being coated in fluoride. Time required for dissolution ranged from 2-4 days.

Dissolved samples were transferred from their 500 mL bottle to two 250 mL Nalgene bottles for compatibility with Beckman Coulter Allegra X-14 centrifuge. They were then centrifuged for 30 minutes at 2600 rpm. This step condensed and separated the fluoride solid from solution and enabled sample solution to be decanted into a clean and leached 500 mL bottle in preparation for initial precipitation of silver chloride (AgCl).

6.3.5. Precipitation

AgCl was precipitated from solution by adding 20 drops of 0.15M HNO3 followed by 24-

48 hours of refrigeration while the crystalline solid settled. A direct precipitation blank for the carrier solution used to spike samples was prepared in a 50 mL centrifuge tube. At this stage in the procedure, samples and both varieties of blanks went through the same procedural steps.

Supernatant fluid was carefully decanted into Ag/HF/HNO3 waste without disturbing the solid at the bottom of the bottle. The solid AgCl product was transferred to clean and leached 50

32 mL centrifuge tubes along with a rinse of 18 MΩ water to transfer any residual solids. Samples were centrifuged at 3600 rpm for 10 minutes to compact the solid for rinsing with 10 mL of 18

MΩ using a VWR Analog Vortex Mixer. The sample was then compacted once more via centrifuge and rinse water was decanted into Ag/HF/HNO3 waste.

6.3.6. Anion exchange chromatography

Crude samples were purified for AMS measurement using anion exchange chromatography. One positively charged, stationary phase resin column per sample was suspended over a waste container and conditioned with 10 mL of 4.0 M HNO3. After the HNO3 drained to the resin bed, the pH was neutralized to ~7 using 18 MΩ water (60-80 mL).

To prepare samples for ion exchange columns, 20 drops of ~30% TraceMetal Grade ammonium hydroxide (NH4OH) and enough 18 MΩ water to bring the total volume to 10 mL was added to each centrifuge tube to dissolve the AgCl solid. Samples were re-capped, the pellet of crude AgCl was dislodged from the bottom of the sample and was fully dissolved by mixing with the VWR Analog Vortex. The silver chloride samples, now in solution, were transferred to their corresponding pre-conditioned anion exchange columns. Each centrifuge tube was rinsed with an additional 10 mL of 0.1 M NH4OH solution, vortexed, and decanted into the respective column to transfer residual sample. The columns were drained to the surface of the resin bed.

Next, 10 mL of 0.05 M HNO3 was eluted through each column.

Clean pre-leached 50 mL centrifuge tubes were labeled and prepped with 2 mL of silver nitrate solution in each. After the columns were drained to the resin bed, the new batch of centrifuge tubes replaced the waste bottles below the suspended columns. Chlorine from each column was eluted into the centrifuge tubes by adding 20 mL of 1.5 M HNO3. Isolated AgCl began to precipitate in a white turbidity in the centrifuge tubes. Finally, 20 drops of TraceMetal grade HNO3 was added to each sample. Samples were capped, vortexed for 10 seconds, and

33 placed in a refrigerator for 12-18 hours to allow the AgCl to precipitate from solution and accumulate.

6.3.7. Final Sample Preparation

After precipitation of AgCl was complete, the samples were centrifuged for 10 minutes at

3600 rpm to discard supernatant fluid. Samples were rinsed using 10 mL of 18 MΩ water and vortex. AgCl was re-centrifuged following the rinse in preparation for transfer to 1.7 mL microcentrifuge tubes for final consolidation of the solid sample pellet. Pre-leached microcentrifuge tubes were labeled and weighed, mass recorded in grams. Each microcentrifuge tube received 1 mL of 18 MΩ water as a medium for sample transfer from the 50 mL centrifuge tube to the 1.7 mL microcentrifuge tube using a separate 2 mL disposable pipette per sample.

The 50 mL centrifuge tubes were rinsed, and the rinse water was transferred to the corresponding microcentrifuge tube. The solid sample, now in the microcentrifuge tubes, was compacted into a pellet by centrifuging for 10 minutes at 3600 rpm. Supernatant liquid was drawn up with a new pipette, one for each sample to prevent contamination, and samples were dried in an oven for 12-

24 hours. After the samples were completely dry and had cooled to room temperature, the final weight of the microcentrifuge tube plus sample was weighed and recorded on bench sheets.

6.3.8. Target Preparation for AMS Measurement

Samples were loaded into copper cathodes with a lens of silver bromide (AgBr) for AMS measurement at Purdue Rare Isotope Measurement (PRIME) lab. A loading workspace was cleaned by kimwipe and covered with an 11x7 sheet of cardstock. A 6x6 cm metal plate with a drilled center for holding copper cathodes in place was lined with kimwipes (set on top of one and an additional covered the plate itself) to be replaced between each sample. Hand tools such

34 as microspatulas, forceps, and one 11/64” diameter high speed steel drill blank for each sample were ultrasonically cleaned, sanded, and sterilized with acetone to prevent contamination. The

Cu cathodes were heated in glass vials via hot plate for 15 minutes. Heating the cathodes before and after they have been loaded removes any water moisture, makes the AgBr less brittle, and reduces 36S. After the cathode was initially heated, it was placed into the slot of the metal plate standing freely with the AgBr lens facing upward. A drill blank pin was used to press a small dimple in the center of the AgCl target to focus the sample in the center of the cathode. Each sample microcentrifuge tube was opened and passed through an antistatic machine to prevent the sample pellet from clinging to the sides. The sample was carefully placed in the center of the cathode and pressed into the AgBr using the drill blank pin and a ball hammer to secure the loose material. The work surface was cleaned between each sample to reduce the risk of cross- contamination. After the cathodes were loaded, they were placed back onto a hot plate for an additional 15 minutes. They were then capped, and sample information was recorded. Sample

ID, Cu cathode holder number, and PRIME lab sample ID were recorded on a sample loading spreadsheet.

6.4. TCN Age Calculation

PRIME Lab facility directly measures 36Cl/37Cl and 35Cl/37Cl, which are used to calculate the total atoms of rock chlorine using equation (1):

[][] [��] = 1 + (1)

35 37 where [37��] and [35��] are atoms of Cl and Cl carrier, is the measured stable isotope ratio, and is 3.127, the natural ratio.

35 The 36Cl/Cl ratio from AMS measurement can be represented using the following relationship:

[ ] = (2) [][][][]

The 36Cl/Cl ratio in the rock, , can be calculated using equation 3:

[ ] = ∗ 1 + (3) [] where [��] represents total atoms of carrier spike added to a sample, [��] represents the total atoms of rock Cl, and is the measured ratio described in equation (2). The net concentration of cosmogenic nuclides in a rock as a function of time in years (a) of surface irradiation is fundamentally described by the following equation:

= 1 − � (4a) () where P represents a site-specific production rate in atoms g-1 a-1 and � is the decay constant of the nuclide in a-1; therefore, the resultant AMS concentration of accumulated rare nuclides is used to compute an apparent exposure age.

As discussed in Limitations of Surface Exposure Dating section, there are many factors and processes that affect the build-up and net concentration of cosmogenic nuclides such as geographic location, sample depth, and exposure history. A considerable fraction of the uncertainty due to surface processes has been studied and computationally expressed by the

Cosmic-Ray Produced Nuclide Systematics on Earth Project (CRONUS-Earth Project) calculator. For example, equation (4b) has been modified from (4a):

= 1 − � + � � (4b) ()

36 -1 and incorporates the term �� for instances of a constrained value for inheritance, � in g cm for density of sample material, � in cm a-1 for erosion rate, and attenuation length, Λ in g cm-2, for samples at depth. Similarly, latitude, altitude, and a factor for shielding, among other modifications, are incorporated into the production term, P. Examples above are just a few of the constantly evolving computational improvements based on observational studies and advancements in measurement precision that improve accuracy of exposure age calculations.

The CRONUS model was used to calculate apparent exposure ages as it is among the most rigorously detailed in process characterization and computationally consistent calculators to date.

Additionally, the production rate used in the CRONUS calculator was obtained using a calibration site proximal to the sample site.

6.5. Geomorphologic mapping methods

Mapping of landforms was done first by referencing previous mapping studies in the literature (Jenny and Kammer 1996, Amman et al., 2001) to cross-reference and recognize features using current and historical Google Earth satellite imagery and aerial photography.

Working from this calibration and map drafts, ground-truthing was performed with field observations the summer of 2016 in the Cordón de Puntas Negras. In the central arid Andes there are many landforms which appear similar to one another and must be distinguished in the context of geomorphological observations or by direct observation in the field. The following is a list of commonly found topographic features and their diagnostic features in this portion of the Andes.

• Rock glacier- landform elongated down valley axis, composed of unconsolidated

debris with frozen interstitial water. They are typically larger on north-facing

peaks and cirques. The surface of rock glaciers contain pockmarks of melt and

flow patterns (ridges and furrows) due to creep downward from high peaks. A 35-

37 45% slope front scarp is a defining feature of rock glaciers as well. They are

indicators of permafrost as they are packages of frozen soil and debris deforming

downslope into large tongue-shaped lobes. Rock glaciers form in both periglacial

and glaciogenic landscapes and are typically found on the highest peaks of the

Andes western cordillera.

• Hummocky moraines- landforms with sinuous ridges showing signs of

denudation or melt of ice-cored moraines to produce a hummocky surface

expression. Hummocky moraines may also be achieved due to meltwater creating

rills which divide a once single moraine into eroded segments. Positive

identification of these landforms is based on the direction of their long axis,

perpendicular to the direction of glacial flow. Hummocky moraines are typically

found in valleys nearest terminal moraine complexes where valley drainage

interacts with branches of tributary channels to erode moraines arranged

perpendicular to the drainage. This creates much smaller segmented moraine

fragments near the terminus.

• Lateral moraines- landforms parallel to the glacial flow direction built from

supraglacial till. At both SPN and El Tatio field sites, lateral moraines are often

high-relief (4-5 m) and sharp-crested. When multiple stages of glaciation are

present in a valley, lateral moraines are often present in sets, with inset lateral

moraines stratigraphically younger as the altitude of their crests decrease along

the valley walls. They are typically present on both sides of the valley walls and

it is sometimes possible to trace them to stratigraphically-correspondent frontal

moraines.

38 • Polygonal patches/patterned ground- landforms creating a field of hummocky

polygons, often elongated in the direction of glacial flow. Sizes of these features

range from 1-4 meters in length and width. Throughout the Central Andes these

polygon fields are located near terminal moraine complexes and in the upper

reaches of the valleys toward the headwall. These polygon patches may be formed

by a number of processes, making their origin somewhat unclear. Surface

processes that may be responsible for creating such features include stagnant rock

glaciers, sublimation and melt of ice carrying supraglacial debris, or post-

depositional cryoturbation. The latter of these processes may be a proxy for

permafrost.

• Irregular drift- unconsolidated glacial debris with defined furrows and ridges

perpendicular to the direction of flow down the long axis of a glacial valley and

into the piedmont where glaciers have flattened and flowed over flat terrain.

• Roche moutonnée- landform elongated in the direction of glacial flow constructed

of exposed bedrock scoured by glacial sliding. These landforms are short, parallel

bedrock ridges in the direction of glacial flow, which is opposite the direction of

lava flow, rock glacier, and flow till ridges and furrows. Positive identification of

these landforms is based on the direction of their long axis parallel to drainage

direction in the formerly-glaciated valley.

• Gelifluction lobes- landforms constructed from ground ice loss in a zone of

permafrost. Ice loss lubricates soil and/or till to the point that it flows and forms

distinctive front scarps. Gelifluction lobes appear similar to rock glaciers in the

field but are much smaller. Gelifluction lobes may be present in glacial, para- and

periglacial landscapes. They are often the result of repeated freeze and thaw;

39 moreover, they flow down valley due to creep. In satellite imagery, it can be

difficult to tell rock glaciers and gelifluction lobes apart. Positive identification

typically comes down to rock glaciers being most prevalent in cirques and near

the highest peaks; however, gelifluction lobes form along the valley walls at both

similar and lower altitudes than rock glaciers. This makes it easier to distinguish

them from similarly tongue-shaped lobate rock glaciers.

• Lava flows- at the field sites for this study, lava flows are ridge and furrowed

tongues of mafic lithologic composition. Young lava flows have the sharpest

ridges and in certain spots, override glacially eroded portions of the landscape.

Older lava flows are sometimes overridden by younger lava flows or have been

glacially eroded to the point that they no longer have distinctive ridges and

furrows but a hummocky appearance. The front scarp of these flows are at a

higher angle of repose than rock glaciers. In the field it is easier to approach them

and determine whether the landform is solid or made of unconsolidated material.

At a distance, e.g. across an entire valley or in satellite imagery, identification

comes down to ridge-and furrow surface texture, cross-cutting relationships, and

proximity to volcanic peaks.

Examples of these landforms can be found in Figure 8 below.

40

Figure 8: Images of landforms in the field. A: Gelifluction lobes (23.8°S, 67.5°W) B: Lava flow (23.8°S, 67.7°W) C: Roche moutonnée (23.8°S, 67.5°W) D: Flow till/irregular drift (22.5°S, 67.9°W) E: Rock glacier (22.2°S, 67.9°S) F: Sinuous moraines (23.8°S, 67.5°W). Images from Google Earth.

The surrounding regions were mapped using the landform criteria above prior to interpretation of glacial history. Landform mapping was finalized by incorporating field observations and preliminary landform delineation from remotely sensed data. Google Earth Pro

KML/KMZ files were exported and overlain onto 1-arcsecond (30 m) resolution ASTER Global

41 DEM data. Hillshade relief, shapefiles, and contours were generated in QGIS. Map feature labels and age annotations were created with Adobe Illustrator CC 2015.3.

6.6. Numerical Modeling Methods

Modeling efforts for this study were adapted from a two-dimensional finite difference model from Kessler et al., 2006. The model simulates glacial formation and growth over a clipped 90 m digital elevation model DEM of field sites using ice flux and mass conservation over a time series. Equations from this mass balance model have been adapted from those used by other workers (Oerleman, 1986; MacGregor et al., 2000; Plummer and Philips, 2003). Ice thickness is calculated at each grid cell using the following equation for continuity:

�ℎ �� �� = � − − �� �� �� where hi is the ice thickness, bz is the accumulation or ablation rate of the ice, and q is the volumetric specific discharge of ice in or out from neighboring cells. The ice is transported between cells via ice deformation, basal sliding, and avalanching.

The glacier models were run for each field site separately from clipped DEMs which incorporated as much of the upstream catchment as possible which could feed the glacier. Using

Matlab software, the model was run until the glacier reached steady-state at maximum extent according to moraine maps. These model runs helped determine response times of the glaciers, which was an input parameter for the stepped ELA model simulations. For these calculations, an initial ELA was defined as well as a step interval. A spin-up time was also input for each glacier to ensure that the glacier stabilized at maximum extent before an ELA step function ran until the

DEM was fully deglaciated.

After the model runs were completed, a path was defined along the long axis of the valley from which moraines were sampled. For each site, a flow path down the thalweg of the valley

42 was created. This flow path enabled the calculation of terminal position and a cumulative distance down the valley so that ELA could first be plotted against distance from the headwall of the valley. Finally, each glacial record was plotted on a chart with markers for each glacial stage.

This chart was anchored by the most reliable ages and the other glacial stages projected stratigraphically into age bins of common regional moraine stabilizations. The ELAs were read from the previous distance chart and plotted in these age bins.

7. Results

7.1. El Tatio region mapping

The western portion of the mountainous region surrounding El Tatio geyser field was once glaciated, leaving behind some of the most spectacularly-preserved valley glacier deposits in the AD. The portion of the Andes range near El Tatio is split into two sub-ranges, both part of the western cordillera. The western peaks reach elevations of 5550 m. The western end of the range is steep, gaining approximately 5300-5550 m of elevation in 3.5-4.5 km and a total length of 23 km. Eastward, there is a ~3500 m-wide, NW-SE-trending valley between the sub-ranges with elevations ranging from 4900-5300 m. The valley walls are significantly steeper on the eastern side, but the eastern sub-range itself does not exhibit the same asymmetry as the western sub-range. The eastern sub-range is 14.2 km in total length with the highest peaks reaching 5400 m and averaging ~5350 m. Sub-ranges, along with the valley, all strike in the NW-SE direction.

The majority of moraines are on the western sub-range, extending ~6000-7000 m from peaks on the western flank of the Western Cordillera, as seen in Figure 6. The moraines are present at elevations between 4125-5300 m. On the eastern portion of the Western Cordillera, glacial moraines extend 4.5-5 km from peaks and are confined to elevations between 4400-5000 m. The eastern flank is where the vast majority of patterned ground is located. Glacial features

43 on the eastern end of the eastern sub-range from the western cordillera are particularly difficult to identify or differentiate from one another. There are expansive lobes (greater than 500 km2 with frontal and lateral moraine-like deposits; however, the low-relief ridges and furrows are no longer whole or connected but broken into segmented polygonal patches. These patches of what appears to be patterned ground are longest perpendicular to the direction of flow. They range from ~90-200 m in length and resemble the outermost hummocky terminal moraines from the

SPN site further south as they too are dissected from drainage. These polygonal patches are superimposed on top of expansive lobate till deposits which may either be flow till or moraines.

The eastern polygonal patches are markedly different based on aspect near the El Tatio site. The western portion of the western sub-range contains large patches of patterned ground in the northern portion of the range. These polygons range from ~60-120 m in both length and width with no particular orientation, just semi-circular clusters of polygonal patches.

The easternmost glacial deposits are on the eastern flank of the eastern row of volcanic peaks which abut the relatively smaller patches of patterned ground at the center of the range.

Most previously-described large patches of flow till with polygonal features are concentrated at the SE end of the range and surrounding a flat, pancake-shaped dacite dome intrusion referred to as “La Torta” between El Tatio and Sairecabur. If the zone between the two chains of high peaks were formerly occupied by a SE-flowing ice field, this would have produced and transported the volume of sediment required to form these features (Figure 9). These features are common across the eastern cordillera, particularly at the SE end of the range as the valleys open up and flatten, enabling piedmont lobes to spread south near La Torta.

In the valley between the two subranges, this area may have been formerly occupied by ice fields feeding valley glaciers to the west and/or east. There is evidence on the northern side of the range just south of a large active rock glacier at approximately 5000 m elevation where

44 valley glacier deposits on the west and east sides of the range are fed by the same upstream source.

7.1.1. Tatio Site

Mapping N El Tatio

field site (22.3°S,

67.98°W) is about

3 km north of the

El Tatio geyser

field. The site is

on the western

flank of the

western subrange

(described above)

in the arid central

Andes.

Preservation and

delineation of

moraines in this

formerly- Figure 9: Overview of El Tatio field site. Moraines pictured in white, flow till in blue polygons, rock glaciers in magenta polygons. Elevation profile across this portion of the range from A to A’ pictured in bottom panel. Location of Figure 10 marked with black rectangle. Images from Google Earth. glaciated valley stands out relative to neighboring glacial deposits. Surrounding glacial features are low-relief polygonal patches, which may be indicative of cold-based thin ice (Ward et al., 2017). It is difficult to interpret the

45 number of glacial stages from the surrounding moraines due to hummocky patterned ground deposits. The sampled valley contains 4 stages of glaciation throughout the valley floor and walls in the form of well-preserved, sharp-crested, high-relief lateral moraines, frontal moraines, and ground moraine.

As pictured in Figure 10, Stage 1 (GS1)-The outermost and stratigraphically oldest glacial stage is a terminal moraine complex which extends 6.1 km from the head of the valley at

4150 m

elevation. The

left- lateral

moraine has a

relief of 6 m.

There are two

right-lateral

moraines as 22.3°S, 68°W well. The

inner right-

lateral Figure 10: Valley sampled at El Tatio with glacial stages indicated (GS1 in red, GS2 in blue). Moraines with no age constraints are purple. Green polygons indicate patterned ground. Magenta polygons are rock glaciers. moraine has the same relief of ~6 m as the left-lateral. An incised river valley exists between the lateral moraines which eroded away nearly all frontal moraines from the complex. The valley is 50 m deep. An additional right-lateral moraine has about 10 m of relief. All moraines from this terminal complex exhibit sharp crests and are symmetrical. Approximately 675 m up the valley is a portion of a frontal moraine that is preserved on the north valley wall which could be the

46 remnants of an advancement or recession. Additionally, a portion of the left- lateral outermost moraine may have been overridden by GS2 lateral moraines on the north valley wall.

Stage 2 (GS2)-Located 1 km up the valley from GS1 moraines, are the Stage II moraines.

This set has lateral moraines on both sides and some have been overridden on the valley walls.

There is also a distinct set of frontal moraines, but it is difficult to match the frontal moraines to laterals since this portion of the valley has steep walls and the lateral moraines appear to be deposited on top of one another on both the north and south valley walls. There are two left- lateral moraines on the northern end of the valley, both with more rounded crests relative to the pristinely preserved sharp-crested right-lateral moraine on the south valley wall. The outermost left-lateral moraine has a relief of 11 m at 4300 m elevation. Moving down the valley from north to south, another left-lateral moraine crest is located 85 m inset from the first at 4290 m elevation. This slightly inset left-lateral moraine is joined to a corresponding frontal moraine with 13 m relief. The drainage valley between the right- and left-lateral is 50 m deep. In this portion of the valley, there are 3 different frontal moraines nested less than 100 m apart from one another with approximately the same relief as the 13 m frontal moraine described earlier. The right-lateral moraine on the southern valley wall has a relief of 10 m. This moraine is asymmetrical with a steep northern slope contributing to the sharp crest. GS2 moraines are located at a bend in the glacial valley where the long axis switches from NE-SW-trending to an orientation of NW-SE moving up in elevation from the terminal moraine complex, GS1.

Stage 3 (GS3)- Located 1525 m up the valley from GS2, GS3 is characterized by a series of lateral moraines plastered on top of one another on both valley walls. One frontal moraine connects two lateral moraines at ~4400 m elevation. The relief of the GS3 moraines range from just 1-3m. The stratigraphically older lateral moraines on the north and south sides of the glacial valley are much sharper-crested than the lateral moraines that connect to a frontal moraine

47 further down the valley walls. The north side of the connecting lateral moraine is hummocky and disconnected with respect to the stratigraphically older lateral moraines from GS3. This stage is located at a bend in the valley where the long axis switches back from NW-SE orientation to NE-

SW.

Stage 4 (GS4)- Located 2155 m up the valley from GS3 at ~4600 m are three separate frontal moraine lobes at the upper recesses of the valley. These moraines range from 4-10 m in relief and are located roughly 1 km south of gelifluction lobes and rock glaciers that occupy the headwall of the valley. This moraine stage is the best preserved and least degraded with respect to the other glacial stages from the valley.

Three samples were collected from the outermost lateral moraines on the southern valley wall. These moraines connect continuously to a terminal moraine complex at the foot of the valley. On the northern side of the valley, five samples were collected near a lateral moraine of slightly lower altitude associated with GS2 on the valley wall that represents a stabilization separate from the outermost and stratigraphically oldest GS1 moraine deposit.

7.2. Puntas Negras Region

Glacial deposits from the western side of the Cordon de Puntas Negras were more spatially dense than glacial deposits from the eastern flank of the range., as seen in Figure 11.

This now de-glaciated landscape has been affected by pre- and post-glacial volcanic activity.

Lava flows, volcanic cones/domes and vents, landslide deposits, and swaths of hydrothermal alteration of the exposed regolith cover all 500 km2 of the range. Examination of the landscape in the field and from aerial photography revealed glacial deposits radially emanating from the center of the range. Evidence of glaciation on the eastern and southern side of the range closely resembles steep valley glacier deposits. Moraines on the NW portion of the range, were at

48 similar altitudes (4350-5000 m) N in a series of roughly NE-SW-

trending lobes which may have

been sourced from an ice field

or cap. The glacial deposits on

the NW side of the range overlie

degraded and eroded lava flows

(relative to young, high-relief

lava flows with relict flow

patterns) on the SE side of the Figure 11: Overview of the Puntas Negras range. Moraines are represented with white lines, rock glaciers are in magenta. The elevation profile from A-A’ is the bottom panel. Laguna Miscanti is located on the west side of the image. Terrain image from Google range. Earth. It is possible to work out a few key cross-cutting relationships based on morphology and lithology of the flows, but there is little detailed geochronological information for lava flows and ignimbrite deposits in this region.

Ignimbrite volcanism began in the late Miocene ~10.4 Ma and continued into recent times (de

Silva 1989a,b). It is, therefore, not implausible that evidence from prior glaciations has not been well-preserved, particularly on the eastern flank of the range, due to post-glacial eruption of lava flows. Additionally, due to the presence of landslide scars, rock glaciers, gelifluction lobes, lava flows, glacial moraines, and glaciated lava flows, it can be difficult to nearly impossible to distinguish between landforms via remote sensing techniques alone in the absence of ground- truthing field observations; yet, the spatial distributions, glacial extents, and similar elevation of glacial deposits suggest it was once capped by ice, making the east and western lobes fed from the same upstream source. Due to challenges to landform identification, mapping was done prior to visiting the site, during the field season, and also revised after confirming identification of

49 remotely sensed landforms at the field site. Sample locations were carefully selected from glacially deposited landforms with little to no signs of re-working.

7.2.1. SPN site mapping

The SPN site (23.8°S, 67.52° S) is a formerly-glaciated valley between the peaks of Co.

Tuyajto to the west and a peak to the east referred to henceforth as Co. Laco at the southern end

23.8°S, 67.4°W

Figure 12: Aerial view of the SPN field site. Glacial stages are marked in red (GS1), blue (GS2b), and orange (GS3). Moraines outlined in purple have no age constraint. Close-up panels on the left side of the figure identify the sample markers. Channels are in blue and till is in green. The brown polygons are bedrock and tan are alluvial fans. of the Cordon de Puntas Negras. The valley is located ~4 km west of the El Laco mining facility.

From remote analysis of satellite imagery, aerial photography, and digital elevation models of the region, two preliminary stages of glaciation were identified based on the distinctively different morphology of moraine deposits. After visiting the site, four stages of glaciation were identified.

Glacial stage 1 (GS1)- Located ~10.5 km from the two peaks that feed into this particular valley, at 4500 m elevation are sinuous, hummocky degraded frontal moraines. Moraine ridges

50 also dissected from valley drainage, resembling short (200-250 m), rolling crests with low relief

(1-2 m). The GS1 terminal moraine complex is covered with a dense concentration of andesitic coarse gravel, with cobbles of andesite, dacite, magnetite, and tufts of grasses. Few boulders with sufficient height and/or stability were available to sample; as such, just two boulders were sampled from this moraine. Patches of polygonal patterned ground features were identified from satellite imagery and mapped at the front of this terminal moraine complex. Relief of these features is <1 m with dimensions of 150-250 m in length and width. A section of resistant bedrock ridges separated this terminal moraine complex from lateral moraine ridges along the eastern valley wall. It remains unclear which if any of the individual lateral moraine ridges may have met the sinuous frontal moraines. One km up the valley from GS1, directly eastward of the bedrock ridge separating GS1 from lateral moraines along the eastern valley wall, two lateral moraine crests extend down-slope toward the valley center. The bend in these moraines toward the western side of the valley lose relief moving westward and abut till packages resembling outwash terraces or ground moraine (relief <1 m) depicted in figure # as green polygons. These features produce the smoothed knobby surface texture between GS1 and GS2.

At the same approximate altitude and latitude as the till deposits on the eastern portion of the valley is a basin slightly over-deepened relative to other portions of the lower valley. The basin contains relatively parallel deposits of outwash and re-worked glacial deposits following the hydrologic drainage pattern toward the basin. This location was not visited in the field, but some of the parallel ridges have been identified by other researchers as roche moutonnée (Jenny et al., 1996). Most features here with positive relief are bedrock and some may be mounds of outwash and unconsolidated material. From the perspective of the center of the valley looking

NW-W toward the bedrock ridges surrounding the smaller inset basin, this topographic high may have once been a drift limit for one of the glacial stages.

51 Glacial stage 2a (GS2a)- A cluster of moraine deposits adjacent to the eastern edge of outwash features which closely follow valley drainage into the over-deepened basin are located approximately 1.5 km up the valley. The relief on these deposits are <1m. From the mapping, this cluster of moraines appears as if it was once a frontal moraine loop, which was dissected by the main channel running down the entire valley. They are at the same latitude (23.84°S) as ground moraine/till on the east portion of the valley. Since they were located “upstream” from the basin along the valley drainage, they may have been affected by stored and/or transient meltwater from the channel that runs through the moraine; thus, this moraine was not sampled.

Lateral moraines are present on the western valley wall here west of the basin as well as the eastern wall in this portion of the valley. These lateral moraines are delineated on the map in

Figure 10 as purple polygons meaning there are moraines present but with no age constraint. Due to the proximity of this heavily dissected, low relief moraine cluster, it was labeled as part of

GS2.

Glacial stage 2b (GS2b)-The next set of sampled moraines (GS2b), ~1km up-slope from the oldest set (GS1) and a mere 240 m from GS2a. GS2b moraines range from 4-8 m in relief and are sharp-crested. A distinguishing feature of this set of moraines are the abundance of lateral moraines along the valley walls. It is unclear from mapping whether this intermediate set of moraines represents one or two distinct stages of glaciation, particularly as there are drainage channels dividing what may be either one wide moraine eroded by channels or two moraine ridges. Relative to GS1 moraines, there is a lower density of cobbles, and less vegetation was supported. There were significantly more boulders with sufficient height and apparent stability, rendering this set of moraines more suitable for sampling.

Glacial stage 3 (GS3) moraines are located ~1.3 km up-valley from the larger GS2 cluster. Stratigraphically, they are the youngest set of moraines with accessibility for sampling.

52 This stage consisted of lower relief (1-3 m) frontal moraines relative to the older stages of glaciation and was primarily characterized by high-relief moraine ridges toward the center of the valley. Similar to GS2b, many lateral moraines are present in the area delineated as GS3.

Glacial stage 4 (GS4) moraines are located ~1.3-1.6 km toward the valley headwall with respect to the GS3 moraines. GS4 moraine altitudes ranged from 4575-4600 m. The GS4 stage was not accessible for sampling. GS4 lateral moraines were on the western valley wall and the rest of the moraines here were frontal moraines with a small patch of roche moutonnée. The lateral moraines were relatively sharp-crested compared to the hummocky sinuous moraines toward the valley center. The sinuous ridges ranged from 160-240 m. These moraines appear to have been heavily dissected from valley channels. These moraines are very low relief (<1 m) and the majority of them are polygonal patches. The fourth stage of glaciation, or recessional position, is located in the upper recesses of the valley near rock glaciers. The GS4 moraines have not been dated and were analyzed in the context of the ages from GS1-3 in the lower portions of the valley floor and from mapping utilizing remotely sensed terrain data from Google Earth and a digital elevation model (DEM).

7.3. Mapping Interpretations

The El Tatio sample site has four stages of glaciation. Down the long axis of the valley, GS1 extended 7.2 km, GS2 5.7 km, GS3 3.7 km, and GS4 1.6 km from the headwall of the valley, occupied by rock glaciers. GS1 and GS2 are the most prominent stages of glacial stabilizations.

Both GS2 and the GS1 terminal complex have frontal and lateral moraines which are wide enough to map as polygons on Figure 14 with high relief (6-13 m). Patches of patterned ground are mapped north and south of the deglaciated valley at ~4500 m elevation. This is the same

53 altitude as GS3, which was discontinuous, particularly on the northern left-lateral moraine. This

GS3 moraine is broken up into polygons, which is most likely a result of cryoturbation processes at this altitude. It also appears that when GS3 was the glacier terminus, there were a few advances and recessions, creating the series of nested lateral moraines along the valley walls which are so prominent for this glacial stage. GS4 is a set of three frontal moraines which meet short lateral moraines, and these three lobes are not nested inside one another as they are with the rest of the glacial stages. All three occur as one stage next to one another at ~4600 m, 1 km down the valley from rock glaciers and gelifluction lobes.

The SPN sample site has 4 glacial stages. Down the long axis of the valley, GS1 extended

7.3 km, GS2a 6 km, GS2b 5.3 km, GS3 3.7 km, GS4 1.3-1.6 km. The extent of glacial stages at the SPN site are nearly identical to the el Tatio site. Similar to Tatio, GS1 and GS2(a&b) are the most prominent stages of glacial stabilization, the latter being the most prominent in the SPN valley. Rather than separating GS2a&b into two stages, the consistency of landform relief across channels led me to conclude these were most probably a moraine complex representing the most prominent glacial stage with multiple advances and recessions over time when this was the terminus. SPN moraines are much more hummocky (relative to Tatio) with polygonal, segmented moraines and patterned ground. This is to be expected, though, as all moraines mapped at SPN are above 4600 m and 4500 m was the altitude at which moraines appeared broken up into polygon patches at the el Tatio site. Lateral moraines are prominent features of

GS1-3 at the SPN site, much like GS1-3 from Tatio; although, Tatio’s GS3 exhibited the most lateral moraines and SPN’s GS2a&b are the stages with the most lateral moraines. GS4 moraines at the SPN site are low-relief, hummocky polygonal patches in the center of the valley with very few lateral moraines. This is a stark contrast to the Tatio GS4 moraines which have higher relief and are intact frontal moraines which are easily identified from satellite imagery. SPN GS4

54 moraines, however, look very similar to the patterned ground patches found at both sites. The commonality between the GS4 stages at both sites is their proximity to rock glaciers near the headwall of the valley. Rock glaciers were identified at SPN beginning at 4670 m elevation, whereas at the Tatio site they were identified from 4775-4940 m.

7.4. Cosmogenic results

7.4.1. SPN ratios

The data tables below detail the sample properties (Table 1), raw AMS data (appendix

A), and calculated ages (Table 2) from both SPN and Tatio samples. The SPN sample 36Cl ratios

(x10-15) ranged from 375.9 ± 7.7 to 6058.8 ± 83.1. It should be noted that sample SPN-16-1 was considered an outlier due to poor current during the AMS measuring procedure.

Cl ratio uncertainties range from 7.7 from sample SPN 16-6 with the lowest 36Cl ratio to 83.1 from SPN-16-2 with the highest ratio. These uncertainty results are expected given that uncertainty increases as Cl ratios increase. Blanks from the SPN site ranged from 5.1 ± 0.6 to

14.2 ± 2. The lowest blank ratio from SPN, sample BLK-1B, was excluded in the calculations for ages as the weight of Cl contamination was negative, indicating a measurement error. As such, the lowest ratio from viable blanks was 6.7 ± 0.9. Cl ratio uncertainties from viable blanks range from 0.8 to 1.3. The weight of Cl contamination (mg) from viable SPN blanks ranged from 0.043

± 0.001 to 0.124 ± 0.002. AMS results from blanks very little Cl contamination. AMS results establish the minimum amount of uncertainty for each sample before concentration uncertainties are later propagated as a results of apparent age calculation using CRONUScalc (Marrero et al.,

2016; Heyman et al., 2011). The wide range of Cl ratios from these samples can be attributed to geologic uncertainty rather than analytical error (Applegate et al., 2010; Fabel and Harbor, 1999;

Hallet and Putkonen, 1994; Putkonen and Swanson, 2003).

55

Table 1: Boulder sample properties

Field Boulder Sample Sample ID Latitude Longitude Elevation Lithology site height thickness °S °W m cm cm SPN-16-1 -23.8532 -67.5170 4510 SPN 120 8.0 Andesite SPN-16-2 -23.8522 -67.5140 4519 SPN 100 4.0 Andesite SPN-16-3 -23.8233 -67.5205 4563 SPN 70 2.0 Andesite SPN-16-4 -23.8235 -67.5257 4578 SPN 60 3.0 Andesite SPN-16-5 -23.8235 -67.5249 4554 SPN 130 2.0 Andesite SPN-16-6 -23.8241 -67.5251 4566 SPN 100 3.0 Andesite SPN-16-7 -23.8269 -67.5238 4560 SPN 40 4.0 Andesite SPN-16-8 -23.8269 -67.5241 4554 SPN 30 2.0 Andesite SPN-16-9 -23.8269 -67.5241 4554 SPN 100 3.0 Andesite SPN-16-10 -23.8263 -67.5210 4560 SPN 170 3.0 Andesite SPN-16-11 -23.8263 -67.5209 4560 SPN 70 2.0 Andesite SPN-16-12 -23.8346 -67.5278 4579 SPN 40 2.0 Andesite SPN-16-13 -23.8345 -67.5273 4582 SPN 70 3.0 Andesite SPN-16-14 -23.8351 -67.5271 4576 SPN 50 4.0 Andesite SPN-16-15 -23.8353 -67.5270 4573 SPN 90 3.0 Andesite SPN-16-16 -23.8361 -67.5269 4578 SPN 90 1.0 Andesite SPN-16-17 -23.8361 -67.5266 4570 SPN 90 3.0 Andesite SPN-16-18 -23.8370 -67.5263 4564 SPN 40 3.0 Andesite SPN-16-19 -23.8383 -67.5235 4544 SPN 40 3.0 Andesite SPN-16-20 -23.8383 -67.5234 4549 SPN 100 4.0 Andesite SPN-16-21 -23.8391 -67.5231 4546 SPN 40 3.0 Andesite Tat-14-1 -22.2942 -68.0012 4387 Tatio 100 2.0 Andesite Tat-14-2 -22.2943 -68.0003 4395 Tatio 60 3.0 Andesite Tat-14-3 -22.2945 -67.9996 4405 Tatio 70 3.0 Andesite Tat-14-4 -22.2945 -67.9992 4417 Tatio 120 0.5 Andesite Tat-14-5 -22.2954 -67.9972 4423 Tatio 150 5.0 Andesite Tat-14-6 -22.3004 -68.0056 4325 Tatio 120 1.5 Andesite Tat-14-7 -23.3005 -68.0056 4329 Tatio 70 1.0 Andesite Tat-14-8 -22.3006 -68.0067 4320 Tatio 100 6.0 Andesite

56 7.4.2. Tatio ratios

Tatio sample 36Cl ratios (x10-15) ranged from 805.6 ± 23.2 to 2268.3 ± 58.3. Cl ratio uncertainty ranged from 20.4 from sample Tat-14-8 with the lowest ratio to 58.3 from sample

Tat-14-6 with the highest ratio. Again, these uncertainty values are expected as they scale relatively linearly with 36Cl ratios. Blanks from the Tatio site ranged from 6.4 ± 0.8 to 8.7 ± 1.5.

Chlorine ratio uncertainties from Tatio blanks ranged from 0.820 from sample Tat B to 1.281 from Tat A. The weight (mg) of Cl contamination ranged from 0.15 to 0.19, which are ~35% higher than blanks from the SPN aliquots yet low enough that contamination does not appear to be an issue. As such, much of the range in 36Cl ratios may also be attributed to geologic uncertainty.

7.4.3. SPN apparent ages

Two boulders were sampled from the terminal moraine complex from the SPN valley.

One of these dates was measured with low current at PRIME lab and was considered an outlier due to these circumstances. The single remaining age for the GS1 stage dated to 255 ± 81 ka.

The GS2b moraine stabilization is a relatively scattered data set with age ranging from 13 ± 6 to

101 ± 43 and majority of the dates clustering around 40-50 ka. The full list of GS2b ages are located in Table 2. Ages from GS3 ranged from 18.5 ± 8.8 ka to 96 ± 14 ka. Again, there is a considerable amount of scatter from this glacial stage. The full list of GS3 ages are listed in

Table 2 and can be seen on the aerial map of the SPN field site in Figure 13. GS3 moraines cluster at approximately 25 ka as well as 40-50 ka.

57 23.8°S, 67.4°W

Figure 13: mapping symbols and colors are the same as Figure 13 with the addition of the cosmogenic ages and error, both in thousands of years (ka).

7.4.4. Tatio apparent ages

The outermost lateral moraines sampled in the El Tatio valley dated to 82.0 ± 15.0, 27.0 ± 2.0, and 19.8 ± 0.9 ka. These are the moraines that can be followed directly to the terminal moraine complex. Lateral moraines inset from these outermost moraines on the northern side of the valley dated to 40.9 ± 4.6, 34.6 ± 3.3, 57.0 ± 11.0, 24.9 ± 2.3, and 26.0 ± 3.0 ka. The full list of ages is located in Figure 14 below, as well as Table 2.

58 Table 2: Sample apparent ages

Sample Age Glacial (ka) Stage (GS) SPN-16-1 14.4 ± 3.15 1 SPN-16-2 255 ± 35.36 1 SPN-16-3 65 ± 21.47 2b SPN-16-4 70 ± 37.49 2b

SPN-16-5 101 ± 39.37 2b SPN-16-6 13 ± 5.47 2b SPN-16-7 28 ± 3.10 2b SPN-16-8 49 ± 15.97 2b SPN-16-9 44.7 ± 2.42 2b SPN-16-10 69 ± 8.06 2b SPN-16-11 88 ± 54.75 2b SPN-16-12 18.5 ± 8.34 3 SPN-16-13 63 ± 21.22 3

SPN-16-14 65 ± 5.14 3 SPN-16-15 84 ± 13.35 3 SPN-16-16 27.4 ± 8.81 3 SPN-16-17 71 ± 15.14 3 SPN-16-18 46 ± 14.37 3 SPN-16-19 42.1 ± 3.60 3 SPN-16-20 96 ± 5.52 3 SPN-16-21 49.9 ± 3.97 3 Tat-14-1 40.9 ± 4.6 2 Tat-14-2 34.6 ± 3.3 2 Tat-14-3 57.0 ± 11.0 2 Tat-14-4 24.9 ± 2.3 2 Tat-14-5 26.0 ± 3.0 2 Tat-14-6 82.0 ± 15.0 1 Tat-14-7 27.0 ± 2.0 1 Tat-14-8 19.8 ± 0.9 1

Table 2: Samples listed with sample names, apparent ages in ka and glacial stage from the SPN and Tatio field sites. Probable outliers are listed in red text.

59 22.3°S, 68°W

Figure 14: Glacial map of Tatio field site. Glacial stage colors are the same as the SPN site (GS1 red & GS2 blue). Cosmogenic ages and error (in ka) are listed for each yellow sample marker. Patterned ground patches are represented in green, moraines with no age constraint are in purple, and rock glaciers are marked in magenta.. both in thousands of years (ka).

7.5. Apparent age interpretation: SPN & Tatio

Due to significant scatter in the 36Cl ages, illustrated in Figures 15 and 16, each glacial stage was interpreted in the context of three separate scenarios: maximum, minimum, and arithmetic mean/median for the Tatio and SPN data sets. Without additional data from this region, a straightforward interpretation of glacial history is difficult to conclude due to scatter and overlap.

60 As such, three different scenarios which researchers have used to interpret this type of data were considered. Surface processes responsible for each of the scenarios are discussed in further detail below one-by-one.

Figure 15: TCN apparent ages plotted with error bars in ka. The colored boxes correspond to the glacial stage colors from the maps (red GS1, blue GS2, orange GS3). Laguna Miscanti curve from Grosjean et al., 2001 plotted for comparison.

61

Figure 16: Probability density functions, or camel plots, of ages from entire Tatio data set, entire SPN data set, and individual glacial stages from both sites. Median ages are marked with a blue line and mean ages in green. Individual ages are shown by thin red lines. The thick black line represents the probability distribution of the representative age population.

62 7.5.1. Scenario 1

Maximum ages from each moraine stage are considered correct under the assumption that the younger ages are due to partial exposure and/or surface processes including erosion, exhumation, or toppled boulders (Putkonen and Swanson 2003; Zreda et al., 1995; Phillips et al.,

1997; Blard et al., 2014).

7.5.1.a. SPN

Out of two samples from SPN GS1, only the oldest date (SPN-16-2) may be considered for interpretation, as the Cl concentration from sample SPN-16-1 was measured with poor current and not considered a reliable measurement. This means that the oldest stage of glaciation from the SPN valley may date back to 255 ± 35.36 ka. This age makes sense stratigraphically, as it is expected that the oldest age from the valley will be located on the furthest extent of an ice-distal terminal moraine. The age is, however, significantly older than ages from all other glacial stages in the cosmogenic data set and may represent an outlier. This means that age constraint on SPN

GS1 is poor.

The oldest ages from the remaining glacial stages date to 101 ± 39.37 ka from GS2b and 84 ±

13.35 from GS3. Again, the ages from these nested moraines correspond correctly to the stratigraphic positions of the glacial stages. Additionally, moraines from all stages may have experienced post-depositional reworking, as evident from GS1 sinuous, hummocky moraine crests which have been dissected from channels. Moraines from GS2b exhibit some of the polygonal patterns which may indicate cryoturbation processes, particularly at elevations exceeding 4500 m, which may have led to boulder exhumation and toppling. Rock glaciers from the high peaks of the SPN valley indicate permafrost conditions and the possibility of ice-cored moraines, especially for moraines located up-valley near the cirque headwall. Partial exposure

63 due to snow cover is less of a factor in this location with strong wind, as evident from desert varnish; however, supraglacial sediment and/or ash shielding may also impact cosmogenic ages here, as this range of stratovolcanoes have been active since the Miocene.

7.5.1.b. Tatio

The oldest age from Tatio GS1 outermost left lateral moraines date to 82.0 ± 15.0 ka

(sample Tat-14-6). Remaining ages may have been eroded as GS1 terminal moraine complex has experienced dissection from channels that cut through the frontal moraines. The oldest age from

GS2 dates to 57.0 ± 11.0 ka. The maximum ages from these nested moraines make stratigraphic sense with older ages in the outermost position and younger ages from GS2 lateral moraines further up valley. The sampled moraine stages from Tatio (GS1 and GS2) are below the altitude at which polygonal patterns become prominent, so cryoturbation may be less of a factor on these moraines. Moreover, the samples were taken from sharp moraine crests at the top of the ridge to minimize sampling reworked boulders from further up the valley. It is, however, more likely toward the crest of the ridge that boulder exhumation and/or toppling had occurred and contributed to scatter.

These two oldest ages were considered anomalously old outliers exhibiting inheritance by previous authors (Ward et al., 2017). If these ages are considered outliers, that makes the oldest

Tatio moraines 27.0 ± 2.0 (GS1) and 24.9 ± 2.3 ka (GS2). These ages fit the stratigraphy and imply that these glacial deposits are much younger than the outlier values.

64

7.5.2. Scenario 2

Minimum limiting ages from each moraine are considered correct to assign deglaciation ages under the assumption that older ages are due to inheritance from a lack of resetting by erosion or decay of a previous irradiation period (Benson et al., 2005; Heyman et al., 2011).

7.5.2.a. SPN

The youngest apparent age from SPN GS1 is the same as the maximum given the single viable measurement, 255 ± 35.36 ka. The youngest age from remaining sampled moraines GS2b and GS3 are 13 ± 5.47 and 18.5 ± 8.34 ka, respectively. The date from GS2b violates the morpho-stratigraphy of the nested glacial stages from the SPN valley and there is substantial overlap between stages GS2b and GS3. Neighboring moraine records from the Chajnantor

Plateau, which experienced similar climatic conditions and erosion rates, contained age disparity that was attributed to cosmogenic inheritance (Ward et al., 2015, 2017). In the SPN valley, jagged lava flows along the western edge of the valley were smoothed by glacial erosion, similar to the roche moutonnée bedrock ridges but neither of them appeared to be affected by deep abrasion or plucking. Given the evidence for extremely low erosion rates and the hyper-aridity of this field site, it is likely that the samples contain an unconstrained fraction of inherited TCN concentration. Moreover, considering the overlap from stages GS2b and GS3 in addition to the scatter, glaciers may have re-advanced after recession to override previously deposited moraines.

7.5.2.b. Tatio

The youngest apparent ages from glacial stages at the Tatio field site are 19.8 ± 0.9 from

Tatio GS1 and 24.9 ± 2.3 from Tatio GS2. These dates contradict the stratigraphic position of

65 these moraines. It was, however, difficult to discern particular lateral moraines from GS1 and

GS2 along the northern valley wall at the Tatio field site. Samples were collected from the high- relief, sharp crests of lateral moraines on both sides of the valley. This eliminates much of the likelihood that these boulders were re-worked by upstream rock fall, slope diffusion, or transportation via the channels that run along the long axis of the valley. Moreover, glacial stages

GS1 and GS2 are located below the 4500 m altitude, at which ground-ice loss and cryoturbation processes become more of a factor (Martini et al., 2017). This is evident from the lack of polygonal patches at the far ice-distal reaches of this formerly glaciated valley along with the estimated altitude of the permafrost layer in the region.

7.5.3. Scenario 3

Arithmetic median ages from each glacial stage best represent the true age of the landform.

7.5.3.a. SPN

The median age from GS2b from the SPN valley is 65 ± 21.45 ka. The median age from

GS3 is 56.5 ± 13.4 ka. Median ages from these two moraines also violate the morpho- stratigraphy of the valley and there is considerable overlap between these two glacial stages.

Camel plots from GS2b cluster around ~25 ka, ~45-50 ka (which is the highest peak from the probability density function, ~65 ka and ~96 ka, Figure 17. This is the glacial stage that most closely resembles a normal distribution. Using the mean age as a minimum from this group and the maximum as a conservative age range (Briner et al., 2005a,b; Smith et al., 2009) would mean

GS2b from the SPN valley was emplaced ~ 58.6±20.9 – 101±39.5 ka as early as marine isotope stage (MIS) 5 or as late as to the boundary between MIS 3 and 4.

66 Using the mean age as a maximum and the youngest age as a minimum for a conservative range (Applegate et al., 2012; Heyman et al., 2011; Smith et al., 2009), would mean that GS2b was emplaced between 13±5.4 - 58.6±20.9 ka. This could place the GS2b stabilization anywhere from MIS 3 to Lateglacial. The younger age range may indicate synchronous glaciation and

Altiplanic lake transgressions for this moraine-building period. Given the presence of the undated GS2a in close proximity to GS2b, it is possible that GS2 does not represent one moraine-building phase, but that the terminus fluctuated at this position in the SPN valley. There are many lateral moraines on the eastern valley wall that could be attributed to these fluctuations in mass balance and moraine-building. Ages from GS3 cluster around ~13 ka, ~25 ka, ~45-47 ka

(age of highest peak from the camel plot), and ~60 ka.

Using the same age range method as described for GS2b, GS3 could have been emplaced from 18.5±8.3 - 56.3±13.36 ka using the mean as a maximum age. Using the mean as a minimum limiting age, it could have been emplaced 56.3±13.36 – 96±5.5 ka, meaning as late as

MIS 5 to as early 56.3 ka near the boundary between MIS 3 to 4. This boundary is contemporaneous with a period of desiccation in the Salar de Atacama (Bobst et al., 2001). The minimum limiting age for the younger age range places this glacial stage as young as 18.5±8.3 ka, which is just prior to the Altiplanic lake transgression, Tauca phase. The 18.5 age relative to the other ages on the GS3 appears anomalously young as an outlier. In that case, the age range would be shifted from 18.5±8.33 - 56.45±13.3 ka to 27.4±8.8 - 56.45±13.3 ka, during MIS 3.

Regionally, ages from the latter age range are present at both the Chajnantor Plateau and the

Tatio field site.

7.5.3.b. Tatio

After removing the outliers from this data set as done in Ward et al., 2017 (Tat-14-3:

57±11, Tat -14-6: 82±15, and Tat-14-8: 19.8±0.9 ka), the age range (using median age as this is

67 the only age not eliminated as an outlier) for GS1 at the Tatio field site is 27±2 ka. Ages from approximately the same time, pre-LGM, were found on the Chajnantor Plateau as well (Ward et al., 2015). The minimum range using median and youngest viable age for GS2 is 24.9±2.3 -

30.3±3.15 ka, making this moraine pre-LGM as well. Maximum age range using median again as a minimum limiting age is 30.3±3.15 – 40.9±4.6 ka. This means that the Tatio GS2 moraine could have been emplaced just before the global LGM during what appears to be the local LGM to MIS 3. Alternatively, the maximum range could mean this moraine was emplaced near the boundary of MIS 2 & MIS 3.

7.6. Numerical Modeling results

Calculated paleo-ELAs were plotted for each site from this study and a neighboring site,

Co. Uturuncu. Locations of the five sites are plotted on Figure 16. The purpose of this comparison was to be able to determine if different stages could have been emplaced or deposited concurrently from site to site. This will narrow the large age ranges from both SPN and Tatio field sites and allow investigation of ELA gradients from site-to-site.

ELA trendlines were plotted (Figure 18) according to flowline length, or distance, on a normalized axis, eliminating the variable of valley hypsometry in order to compare between sites. Markers correspond to the cosmogenic ages from each moraine. For those from the SPN and Tatio valleys, median ages were used due to scatter from the cosmogenic results.

Glacial parameters such as ice thickness, viewpoint azimuth and elevation, depth to water table, and frequency of calving and avalanching were treated the same across all field sites. Thus, if there is a bias in the data set due to these parameters, it is a systematic bias across all sites used in the simulation.

68

Figure 17: Field sites of glacial chronologies are represented with black markers. The yellow numbers represent field sites that were not visited for this study but used for analysis. These include (1) El Tatio and (5) East & (4) West Chañi. The red marker represents (3) SPN, and the grey site (2) is from neighboring Co. Uturuncu ELAs ranged from 5200-5500 m from the Uturuncu site, which reasonably compares to the 65 ka moraine ELA of 5280 m and 16-14 ka moraine ELA of 5350 m from Blard et al.,

2014). Three moraine stages were present, with ages in stratigraphic order. The west Chañi site

ELAs ranged from 5050-5500 m with two moraine stages in stratigraphic order but without an

LGM moraine. ELA estimates were within ±40 m of estimated ELAs from Nevado de Chañi

(Martini et al., 2017). ELAs at the SPN field site ranged from 5000-5400 m with four glacial stages, 3 dated and one without age constraint. East Chañi ELAs range from 5000-5400 m with four glacial stages, three were dated and one was not. El Tatio ELAs range from 4950-5450 m

(compared to 4830m from Kull & Grosjean, 2000) with four glacial stages. Two of these glacial stages are dated and two of them have no age constraint. The two markers, one at 5400 m and the

69 other at 5450 m are a single stage (moraines are at the same altitude and are not nested, as they are located where tributaries join the main valley).

7.7. ELA Interpretations

ELA-terminus position trendlines from each site were calculated from the Kessler glacial model (Figure 18). Position data was then normalized using the distance of the outermost Regional ELAs SPN_Puntas Tatio W_Chañi E_Chañi Uturuncu 12 ka 14-17 ka LGM Pre-LGM-MIS 3 MIS 4 MIS 6 No ages 5800 5600 5400 5200 ELA (m) 5000 4800 4600 0 0.2 0.4 0.6 0.8 1 Normalized Distance

Figure 18: Results of paleo-ELA data generated by the numerical model from Kessler et al., 2006. Each field site is represented by a colored trendline. Blue and yellow from Martini et al., 2017, Green from Blard et al., 2014. Orange and grey trendlines are from this study. Markers are the positions of moraine stabilizations along the distance axis with distance normalized. Markers with age constraint are filled and markers with no fill color are those with no age constraint. terminal moraine from each site. Reliable ages (colored markers on Figure 19) included the

Uturuncu record, east Chañi, west Chañi, the oldest glacial stage from SPN and the two moraines from Tatio. These were plotted on Figure 19 according to age and calculated ELA. The reliable ages were used to anchor a regional ELA analysis plot in Figure 19. The solid lines represent the reliable ages. Please note that Tatio and SPN moraines were plotted according to mean ages at

Tatio and the date from the terminal moraine was used from SPN. ELA values were plotted on

Figure 19 for each of the field sites included in the analysis. The dashed lines are the projected

70 ages in reference to reliable ages. Glacial stages were plotted in stratigraphic order in regionally recurring age bins similar to a histogram. The projected ages (dashed lines) were treated as if there was no data for age constraint and were plotted according to stratigraphic order from their respective valleys.

Figure 19: This plot is anchored by regional glacial records with relatively better age constraint than the cosmogenic data set from this study. Reliable ages are plotted with filled markers. Undated moraines are plotted with unfilled markers. The SPN site is in orange, Tatio in grey, E Chañi in blue, W Chañi in yellow, and Uturuncu in green. Please note the break in the age scale between 30-40 ka and MIS 6. Undated glacial stages were projected in stratigraphic order into age bins from regional glacial records. This figure illustrates trends in glacial ELAs geographically and through time.

Both of the dated glacial stages (GS1 and GS2) from Tatio fell into the age range of Pre-LGM-

MIS3. Additionally, the terminal moraines from Uturuncu, W Chañi, and Tatio all dated within this age range (Pre-LGM-MIS3). Figure 19 indicates widespread glaciation at 30-40 ka and during the Tauca phase (14-17 ka).

According to model data synthesized in Figure 19, the undated moraine stages from the

SPN were projected at 30-40 ka, 20-25 ka, and 14-17 ka. The Tatio ages from the upper portions of the valley were projected to be 14-17 ka and 12 ka.

71 8. Discussion

New cosmogenic data sets from this study are the first chronologies this far south in the subtropics and mid-latitudes of northern Chile. The ages from these sites are notably scattered; therefore, it is difficult to determine probable outliers. Analytical errors, as listed in the appendix tables, are low enough to suggest that geological error e.g.: pre-depositional (inheritance from prior exposure) and/or post-depositional geologic surface processes (erosion, shielding, moraine degradation, boulder toppling, and exhumation) are the cause for the age scatter rather than analytical error. These factors make it nearly impossible to determine a straightforward glacial chronology targeting the true ages of deglaciation from the landforms with confidence. As such, scenarios were put forth in the Results chapter for both cosmogenic data and ELA comparisons.

8.1. Regional glacial trends: Northern Altiplano

Moraines from the northern portion of the Altiplano-Puna plateau at Co. Tunupa, located at the northern end of the Salar de Uyuni, have two preserved stages of glaciation (Blard et al.,

2013a) The oldest stage at Tunupa and local last glacial maximum persisted until 16-15 ka; therefore, synchronous (within error) with the Lake Tauca highstand. Bedrock samples indicated rapid retreat during the lake regression at 14.5 ka. The youngest glacial stage at Co. Tunupa, their M2 recessional moraine, was emplaced between 15.5 and 14.5 ka (Blard et al., 2013a).

At Co. Uturuncu, on the southern end of the Lake Titicaca basin, five stages of glaciation were identified, two dated. The terminal moraine stage and maximum extent of glaciation (M5) was emplaced 65-37 ka, during early MIS 3. As this moraine stage spanned many thousands of years, the authors made a point to state that the mode from this moraine was 40 ka and ELA was

5250 m (Blard et al., 2014). These old ages, at the time of publication, were the first to suggest that the local last glacial maximum on the Altiplano occurred prior to global last glacial

72 maximum (21 ka). At Uturuncu, glaciers persisted at their maximum extent until as late as 18 ka.

The youngest glacial stage at Co. Uturuncu, their M2 moraine, was emplaced by 34-17 ka, recessed 1 km at ~17 ka, followed rapid deglaciation. Their striated bedrock ages 15-14 ka also indicated, similar to Tunupa, that deglaciation occurred during the Tauca phase lake regression

(Blard et al., 2011; Placzek et al., 2006b; Sylvestre et al., 1999).

8.2. Tatio Site Chronology: Cosmogenic Ages

Four glacial stages were identified from the El Tatio field site at ~21°S, just 1° south of the Uturuncu site. The outermost two stages were dated. Oldest ages from both glacial stages and the youngest age from GS2 were considered outliers from the data set. The justification for identifying oldest ages as outliers was the aridity and little evidence for erosion, ideal conditions for inheritance. The youngest age was eliminated due to signs of moraine degradation. Heyman et al., 2011 suggest that the latter is the dominant process leading to scatter in data sets of glacial boulders older than a few thousand years. At this field site, though, there is evidence that pre- depositional processes may contribute to the scatter as well. The removal of outliers which may eliminate both over- and under-estimating true ages (Tat-14-3; 57.0±11.0, Tat-14-6; 82.0±15.0, and Tat-14-8; 19.8±0.9) make age estimates 27.0 ± 2.0 ka for GS1 and a range from 24.9-40.9 ka for GS2.

Median age for GS2 is 30.8 ka as the majority of ages from Tatio GS2 date from as early as MIS 3 to the boundary of MIS 2 and MIS 3. The youngest ages from each of these moraines remain in stratigraphic order, but the age range from GS2 violates the stratigraphy of the landforms. It is still permissible, however, that local last glacial maximum may have indeed occurred prior to the global LGM at this site. Additionally, if GS2 is Pre-LGM-MIS 3 in age, the morphostratigraphy allows the undated GS3 could be the LGM moraine. As discussed in the

73 mapping results, there is a well-preserved moraine (GS4) in a recessed position proximal to the headwall of the Tatio valley that was not dated due to inaccessibility in the field. This moraine stage contains 2 well-defined arcuate moraine loops with 4-10 m relief next to one another rather than nested. This description matches that of the pre-LGM to Tauca moraine (M2) from the Co.

Uturuncu record (Blard et al., 2014); therefore, it is possible that the undated GS4 moraine could date to the Tauca highstand. Even if this were the case, the potential Tauca moraine was not a significant glacial standstill further than 22° S latitude. All three of these northern sites are proximal to the Lake Titicaca basin, which housed Altiplanic paleolakes.

Tauca phase may have created a positive precipitation feedback, supporting glaciation in areas affected by a disturbance in spatial distribution of precipitation delivery. Moreover, positions of the Tauca moraine are further recessed moving southwest from Co. Tunupa across the Altiplano plateau, as would be expected from amplified rainfall of the Lake Tauca episode.

This indicates that even the southern Altiplano/northern Puna plateau region may have been affected by the Tauca Highstand, synchronous with Heinrich 1 event in the North Atlantic (Lea et al., 2003; Andersen et al., 2004; Blard et al., 2007). Blard et al., 2013 attribute the contemporaneous Co. Tunupa deglaciation and Lake Tauca regression to warm and dry conditions following a northward migration of the SASM after 14.5 ka. The onset of warm, dry climatic regime on the Altiplano occurred synchronously (within error) with the abrupt onset of the Bolling-Allerod warming (Chiang et al., 2003; Blard et al., 2009; Blard et al., 2014). Median ages from Tatio GS2 moraines are similar to moraines from the neighboring Chajnantor to the south of El Tatio as well. This further supports Pre-LGM-MIS 3 glaciation rather than LGM- and

Tauca-aged moraines. The regional cosmogenic ages point to a most plausible conclusion that

Tatio cosmogenic true ages are much older than previously-assumed Lateglacial ages. If one of

74 the undated glacial stages is Tauca in age, it is in the upper recessed position of the valley and does not indicate a major advance at this time.

8.3. SPN Site Chronology: Cosmogenic Ages

Four glacial stages were identified at the SPN site, three of which were dated. Similar to the El Tatio field site, the distribution of ages is skewed toward older exposure ages. The following discussion is a set of interpretations considering the cosmogenic ages from the glacial stages alone without consideration of the paleo-ELAs calculated for each of the moraine stages.

Anomalously old and young samples were eliminated as outliers, for the same reasons from

Tatio (SPN-16-1; 14.4 ± 3.15 ka, SPN-16-6; 13 ± 5.47 ka, SPN-16-12; 18.5 ± 8.34 ka, and SPN-

16-5; 101 ± 39.37 ka). This left no age constraints on the outermost moraine, SPN GS1. SPN

GS2b ages ranged from 27.4-96 ka, with a median of 63 ka. SPN GS3 ages, after outlier removal, ranged from 28-88 ka with a median of 57 ka. The median ages fit the morpho- stratigraphic position of the glacial stages in the SPN valley, but there is significant overlap between GS2b and GS3. This may be attributed to GS3 either overriding the previously deposited GS2b stage or boulder re-working from GS3 down to GS2b.

Two of the samples from SPN were collected ~28 m apart from each other from the same moraine in GS2b, SPN-16-14 and SPN-16-15, which returned ages of 65 ± 5.14 and 84 ± 13.35, respectively. These ages overlap one another within error. An additional pair from the same moraine, SPN-16-19 and SPN-16-20, was collected a mere 12 m apart, and yielded ages of 42.1

± 3.60 and 96 ± 5.52. These ages are significantly different from one another and clearly highlight the possibility of pre-depositional, post-depositional, or both types of geomorphological processes contributing to scatter. Samples from GS3 (SPN-16-8 and SPN-16-

9), however, were collected just under 3 m apart and yielded almost identical ages (49 ± 15.97

75 and 44.7 ± 2.42 ka, respectively), which are remarkably close to the 42.1 ± 3.60 ka sample as well. Therefore, boulders from the GS3 stage may have been reworked and deposited alongside boulders from GS2b, located 1 km further down the valley. This also supports the possibility that

GS3 may have overridden GS2b. Support for the 35-40ka ages being true ages of glacial stabilizations are the presence of samples which date to the same time period from the

Chajnantor Plateau (~65 km north of the SPN field site and ~85 km south of Tatio). At the

Chajnantor site, there is little (ages eliminated as outliers) to no evidence to support the presence of a Tauca moraine that far south. This, however, could be due to a similar situation from El

Tatio and SPN where the highest moraines in the valley may not have been sampled due to inaccessibility.

The cosmogenic data sets from Tatio and SPN both have a distribution of ages clustering toward the older ages rather than younger, which according to Applegate et al., 2010 is characteristic of a data set affected by inheritance. The aridity of the field area and low erosion rates also support the conclusion that these samples may not have experienced enough erosion to be reset from a pre-depositional exposure (Owen et al., 2003; Heyman et al., 2011). This is in direct contrast to the youngest chosen outliers for SPN as discussed above. If inheritance was a large fraction of the geologic error from this field site, it would be the younger ages from the set that would more accurately describe true ages for these moraines. Support for this hypothesis emerges from the paleo-ELA patterns of this region, which make a case for young ages of glacial stabilizations than the cosmogenic ages suggest.

76 8.4. Paleo-ELA plots

Paleo-ELA trend plots derived from the numerical modeling exercise were plotted on a distance axis and normalized to the terminal moraine from each site. This enabled comparing different field sites without influence of valley hypsometry. Figure 18 from the Results chapter plots ELA trend lines from modeling along the normalized distance axis. Moraines were potted at their appropriate distance from the terminus and marked according to their cosmogenic median ages for the SPN and Tatio sites. Moraines from Uturuncu and Nevado de Chañi were plotted based on their published ages. Figure 19 utilizes ELA values read from Figure 18 to plot the ELA from each glacial stage on a time series separated into bins of recurring ages from regional cosmogenic records. The two dated glacial stages from the Tatio field site were originally marked on Figure 18 with the same symbol of Pre-LGM-MIS 3 since the only viable

GS1 age (27 ± 2 ka) overlapped significantly with the ages from GS2b, as discussed earlier.

Beginning from north to south along the Western Cordillera of the Andes, ELAs rise and are on average 100 m higher at the SPN field site compared to the Tatio site in the north. As these are precipitation-limited glaciers in the arid Andes, a higher ELA represents a site that received less moisture than a site with a lower ELA. This indicates a loss of moisture from north to south across the climatic AD feature. It also appears from Figure 19 that the SPN site deglaciated rapidly after the Tauca phase but the Tatio site remained glaciated just prior to the

Lateglacial. This is in contrast to the assumption from the mapping section that the highest moraines in the upper recesses of the Tatio valley are Tauca highstand in age. If this were the case, the moraines from Tatio long assumed to be Lateglacial in age, are still older than once thought with the possibility of two moraine stabilizations synchronous with Altiplanic lake highstands (Tauca 14-17 ka and Coipasa 12 ka) (Placzek et al., 2013). This could mean that glacial stages were concurrent with lake highstands much further south than suggested from

77 previous work in this region. This idea entertains the possibility that positive precipitation feedbacks from the large surface area of Altiplanic lakes cast a larger radius than suggested by previous workers.

From east to west, the ELAS from Nevado de Chañi sites are the lowest values from the chart followed by the Tatio field site, this also implies a moisture loss westward from the Eastern

Cordillera. Finally, records 2 and 3 (Uturuncu and SPN) have the highest ELA values. This, intuitively, makes sense as these sites are on the central plateau rather than from the edges. Sites from Chañi and Tatio are more proximal to the easterly and westerly moisture belts, respectively.

The Uturuncu and SPN sites are located centrally on the Altiplano-Puna plateau and distal to the southern boundary of the Altiplanic paleolake shorelines.

If this were the case, one of the core hypotheses of this project (that preserved moraines from Tatio will date to as late as Tauca with no moraines that date to the Lateglacial) may be permissible or challenged and more age data from the undated moraines is needed to further assess the possibility. As the two younger, undated stages may correspond to the Tauca highstand and the Coipasa highstand present in the E Chañi record, this does not yet refute the argument that moraine stages in the AD will be concurrent with Altiplanic lake highstands.

78 9. Conclusions

In summary, extensive geomorphological mapping was done throughout the Cordillera del Tatio and Cordón de Puntas Negras. Andesite boulders were sampled at each site to create cosmogenic 36Cl glacial chronologies. Additionally, paleo-ELAs were produced using a glacier model with clipped DEM inputs from regional field sites for a trend analysis. Due to significant scatter in the age distributions for each glacial stage, this study did not yield straightforward 36Cl glacial chronologies from the Tatio and SPN field sites. A synthesis from available data implies that the ages of these glacial deposits are either MIS 3 and older according to exposure ages. The modeling, however, points toward the possibility that moraines nearest the headwalls of the valleys could have been synchronous with Altiplanic lake phases, particularly the Tatio field site.

Tauca moraines at both Tatio and SPN were not dated due to inaccessibility in the field, but the modeling exercise projected that both the SPN and Tatio sites may have been glaciated during the Tauca phase. Modeling data also indicates that the Tatio site was glaciated through the transition into the Lateglacial. In contrast, the SPN site appears to have been glaciated up to

Tauca phase but then rapidly deglaciated before 12 ka. Additionally, the final synthesis from the modeling data confirmed that in this field area with precipitation-sensitive glaciers, there is merit to the apparent trend of moisture loss moving progressively SW across the AD climatic feature in the central Andes. Evidence in support of this trend included the ELA rise along the Western

Cordillera between the Tatio and SPN field sites as well as ELA rise from the easternmost sites at Nevado de Chañi to the sites on the plateau and in the Western Cordillera.

Geological uncertainty is the largest fraction of error from this data. Factors such as pre- depositional exposure (inheritance) and/or post-depositional boulder exhumation, toppling, and erosion may be responsible for the age distributions. Additionally, these dates may represent glacial margins advancing and retreating multiple times to create some overlap from additional

79 age populations on one moraine. Without more certainty about depositional and degradational surface processes, assumptions of equivalence come into question. Equivalence is fundamental to surface exposure dating, particularly in a field area where landforms are difficult to identify from remotely sensed data.

The most robust findings from this study are the contributions to mapping in the AD and additional questions raised regarding whether or not the glacial stages from these sites were synchronous with Altiplano lake phases. Ages from the uppermost glacial moraines may be useful to further constrain the valley chronologies. Advancement in cosmogenic dating techniques such as the growing utilization of 3He may also be useful in this field area; although, additional work on surface processes to better understand geologic error affecting existing records is vital precursory work. More time spent in the field to bolster landform identification would be useful here as well, especially considering how difficult it is to identify landforms from satellite imagery alone. Finally, cosmogenic dating of striated bedrock, if present, would also be useful as a means to further interpret durations and rates of glacial response from SPN and Tatio.

80 References Cited

Allmendinger, Richard W., Sharp, James W., Von Tish, Douglas, Serpa, Laura, Brown, Larry, Kaufman, Sidney, Oliver, Jack, Smith, Robert B., 1983. Cenozoic and Mesozoic structure of the eastern Basin and Range province Utah, from COCORP seismic-reflection data. Geology. v. 11(9). p. 532-536. doi: https://doi.org/10.1130/0091-7613(1983)11<532:CAMSOT>2.0.CO;2

Ammann, C., Jenny, B., Kammer, K., Messerli, B., 2001. Late Quaternary glacier response to humidity changes in the arid Andes of Chile (18-29S). Palaeogeography, Palaeoclimatology, Palaeoecology. v. 172(3-4). p. 313-326. doi: 10.1016/S0031-0182(01)00306-6.

Andersen, K.K., Azuma, N., Barnola, J.M., Bigler, M., Biscaye, P., Caillon, N., Chap- pellaz, J., Clausen, H.B., DahlJensen, D., Fischer, H., Fluckiger, J., Fritzsche, D., Fujii, Y., Goto-Azuma, K., Gronvold, K., Gundestrup, N.S., Hansson, M., Huber, C., Hvidberg, C.S., Johnsen, S.J., Jonsell, U., Jouzel, J., Kipfstuhl, S., Landais, A., Leuen- berger, M., Lorrain, R., Masson- Delmotte, V., Miller, H., Motoyama, H., Narita, H., Popp, T., Rasmussen, S.O., Raynaud, D., Rothlisberger, R., Ruth, U., Samyn, D., Schwander, J., Shoji, H., Siggard-Andersen, M.L., Steffensen, J.P., Stocker, T., Sveinbjornsdottir, A.E., Svensson, A., Takata, M., Tison, J.L., Thorsteinsson, T., Watanabe, O., Wilhelms, F., White, J.W.C., 2004. High-resolution record of Northern Hemisphere climate extending into the last interglacial period. Nature. v. 431. p. 147– 151.

Applegate, P. J., Urban, N.M., Laabs, B.J.C., Keller, K., Alley, R.B. 2010. Modeling the statistical distributions of cosmogenic exposure dates from moraines. Geoscientific Model Development. v. 3 p. 293-307. doi: https://doi.org/10.5194/gmd-3-293-2010.

Applegate, Patrick J., Urban, Nathan M., Keller, Klaus, Lowell Thomas V., Laabs, Benjamin J.C., Kelly, Meredith A., Alley, Richard B. 2012. Improved moraine age interpretations through explicit matching of geomorphic process models to cosmogenic nuclide measurements from single landforms. Quaternary Research. v. 77(2). p. 293-304. doi: https://doi.org/10.1016/j.yqres.2011.12.002.

Argollo, Jaime, Mourgruiart, Phillippe. 2000. Late Quaternary climate history of the Bolivian Altiplano. Quaternary International. v. 72(1). p. 37-51.

Azócar, G. F., Brenning, A. 2010. Hydrological and Geomorphological Significance of Rock Glaciers in the Dry Andes, Chile (27°-33°S). Permafrost and Periglacial Processes. v. 21. p. 42-53. doi: 10.1002/ppp.669

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. v. 409(6821). p. 698-701. doi: 10.1038/35055524.

Baker, Paul A., Fritz, Sherilyn C. 2015. Nature and causes of Quaternary climate variation of tropical South America. Quaternary Science Reviews. v. 124. p. 31-47.

81 Balco, Greg, Stone, John O., Lifton, Nathaniel A., Dunai, Tibor J., 2008. A complete and easily accessible mean of calculating surface exposure ages or erosion rates from 10Be and 26Al measurements. Quaternary Geology. v. 3(3). p. 174-195.

Benson, Larry, Madole, Richard, Landis, Gary, Gosse, John. 2005. New data for Late Pleistocene Pinedale alpine glaciation from southwestern Colorado. Quaternary Science Reviews. v. 24(1-2). p. 49-65. doi: https://doi.org/10.1016/j.quascirev.2004.07.018.

Betancourt, J. L., Latorre, C., Rech, J. A., Quade, J., Rylander, K. A. 2000. A 22,000-Year Record of Monsoonal Precipitation from Northern Chile’s Atacama Desert. Science. v. 289(5484). p. 1542- 1546. doi: 0.1126/science.289.5484.1542.

Bills, Bruce G., de Silva, Shanaka L., Currey, Donald R., Emenger, Robert S., Lillquist, Karl D., Donnellan, Andrea, Worden, Bruce. 1994. Hydro-isostatic deflection and tectonic tilting in the central Andes: Initial results of a GPS survey of Lake Minchin shorelines. Geophysical Research Letters. v. 21(4). doi: https://doi.org/10.1029/93GL03544.

Blard, P.-H. Lavé, J., Pik, R., Wagnon, P., Bourlès, D. 2007. Persistence of full glacial conditions in the central Pacific until 15,000 years ago. Nature. v. 449. p. 591-594.

Blard, P., Lavé, J., Farley, K., Fornari, M., Jiménez, N., Ramirez, V. 2009. 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). Quaternary Science Reviews. v. 28. p. 27-28. doi: 10.1016/j.quascirev.2009.09.025.

Blard, P.-H., Sylvestre, F., Tripati, A.K., Claude, C., Causse, C., Coudrain, A., Condom, T., Seidel, J.-L. Vimeux, F., Moreau, C., Dumoulin, J.-P., Lavé. 2011. Lake highstands on the Altiplano (Tropical Andes) contemporaneous with Heinrich 1 and the Younger Dryas: new insights from 14C, U-Th dating and 18O of carbonates. Quaternary Science Reviews. v. 30(27-28) p. 3973- 3989. doi: https://doi.org/10.1016/j.quascirev.2011.11.001.

Blard, P.-H., Lavé, J., Sylvestre, F., Placzek, C.J., Claude, C., Galy, V., Condom, T., Tibari, B. 2013. Cosmogenic 3He production rate in the high tropical Andes (3800 m, 20°S): Implications for the local last glacial maximum. Earth and Planetary Science Letters. v. 377-378. p. 260-275. doi: https://doi.org/10.1016/j.epsl.2013.07.006.

Blard, P.H., Lave, J., Farley, K.A., Ramirez, V. Jimenez, N., Martin, L.C.P., Charreau, J., Tibari, B., Fornari, M., 2014. Progressive glacial retreat in the Southern Altiplano (Uturuncu volcano, 22°S) between 65 and 14ka constrained by cosmogenic 3 He dating. Quaternary Research. v. 82(1). p. 209-221.

Blodgett, T.A., Lenters, J.D., Isacks, B.L., 1997. Constraints on the origin of paleolake expansions in the Central Andes. Earth Interactions. v. 1(1). p. 1-28. doi: 10.1175/1087- 3562(1997)001<0001:COTOOP>2.3.CO;2.

Bobst, Andrew L., Lowenstein, Tim K., Jordan, Teresa E., Godfrey, Linda V., Ku, The-Lung, Luo, Shangde. 2001. A 106 ka paleoclimate record from drill core of the Salar de Atacama, northern

82 Chile. Palaeogeography, Palaeoclimatology, Palaeoecology. v. 173(1-2). p. 21-42. doi: https://doi.org/10.1016/S0031-0182(01)00308-X.

Borchers, Brian, Marrero, Shasta, Balco, Greg, Caffee, Marc, Goehring, Brent, Lifton, Nathaniel, Nishiizumi, Kunihiko, Phillips, Fred, Schaefer, Joerg, Stone, John. 2016. Geological calibration of spallation production rates in the CRONUS-Earth project. Quaternary Geochronology. v. 31. p. 188-198.

Brenning, A. 2008. The impact of mining on rock glaciers and glaciers: examples from Central Chile. In: B. S. Orlove, E. Wiengandt & B. Luckman (eds.), Darkening peaks: glacier retreat, science, and society. University of California Press, Berkeley. Chapter 14, p. 196-205.

Briner, J.P., Kaufman, D.S., Manley, W.F., Finkel, R.C., Caffee, M.W., 2005a. Cosmogenic exposure dating of late Pleistocene moraine stabilization in Alaska. Geololgical Society of America Bulletin. v. 117 p.1108–1120.

Briner, J.P., Miller, G.H., Davis, P.T., Finkel, R.C., 2005b. Cosmogenic exposure dating in arctic glacial landscapes: implications for the glacial history of northeastern Baffin Island, arctic Canada. Canadian Journal of Earth Science. v. 42 p. 67–84.

Brown, E.T., Bourles, D.L., Colin, F., Raisbeck, G.M., Yiou, F., Desgar- ceaux, S., 1995a. Evidence for muon-induced in situ production of Be in near-surface rocks from the Congo. Geophysical Research Letters v. 22, p. 703-706.

Chiang, John C.H., Biasutti, Michela, Battisti, David. 2003. Sensitivity of the Atlantic Intertropical Convergence Zone to Last Glacial Maximum boundary conditions. Paleoceanography. v. 18(4). doi: https://doi.org/10.1029/2003PA000916.

Clapperton, Chalmers. 2000. Interhemispheric synchroneity of Marine Oxygen Isotope Stage 2 glacier fluctuations along the American cordilleras transect. Journal of Quaternary Science. v. 15(4). p. 435-468. doi: https://doi.org/10.1002/1099-1417(200005)15:4<435::AID-JQS552>3.0.CO;2-R.

Clark, Peter U., Dyke, Arthur S., Shakun, Jeremy D., Carlson, Anders E., Clark, Jorie, Wohlfarth, Barbara, Mitrovica, Jerry X., Hostetler, Steven W., McCabe, Marshall A. 2009. The Last Glacial Maximum. Science. v. 325(5941). p. 710-714.

Clayton, J.D., Clapperton, C.M., 1997. Broad synchrony of a Late-glacial glacier advance and the highstand of palaeolake Tauca in the Bolivian Altiplano. Journal of Quaternary Science. v. 12(3). p. 169-182. doi: 10.1002/(SICI)1099-1417(199705/06)12:3<169::AID-JQS304>3.0.CO;2-S.

Coira, B., Mahlburg, Kay S., Viramonte, J., 2010. Upper Cenozoic magmatic evolution of the Puna-A model for changing subduction. International Geology Review. v. 35(8). p. 677-720.

Davidson, J.P., de Silva, S.L.1995. Late Cenozoic magmatism of the Bolivian Altiplano. Contributions to Mineralogy and Petrology. v. 119(4). p. 387-408.

De Martonne, E. 1934. The Andes of the North-West Argentine. The Geographical Journal, 84 (1), p. 1- 14. doi: 10.2307/1786827.

83

Denton, G.H., Anderson, R.F., Toggweiler, J.R., Edwards, R.L., Schaefer, J. M., Putnam, A.E. 2010. The Last Glacial Termination. Science. v. 328 (5986). p. 1652-1656. doi: 10.1126/science.1184119.

Denton, G.H., Heusser, C. J., Lowell, T.V., Moreno, P. I., Anderson, B. G., Heusser, Linda E., Schlühter, Marchant, D.R. 2003. Interhemispheric linkage of paleoclimate during the Last Glaciation. Geografiska Annaler. v. 81(2). p. 107-153.

Denton, G.H., Heusser, C. J., Lowell, T.V., Moreno, P. I., Anderson, B. G., Heusser, Linda E., Schlühter, Marchant, D.R. 2003. Interhemispheric linkage of paleoclimate during the Last Glaciation. Geografiska Annaler. v. 81(2). p. 107-153.

Denton, G.H., Broecker, W.S., Alley, R.B., 2006. The mystery interval 17.5 to 14.5 kyrs ago. PAGES News. v. 14. p. 14-16.

Desilets, D., Zreda, M., 2003. Spatial and temporal distribution of secondary cosmic-ray nucleon intensities and applications to in situ cosmogenic dating. Earth and Planetary Science Letters. v. 206(1-2). p. 21-42. doi: https://doi.org/10.1016/S0012-821X(02)01088-9.

Desilets, Darin, Zreda, Marek, Almansi, Peter F., Elmore, David, 2006. Determination of cosmogenic 36Cl in rocks by isotope dilution: innovations, validation and error propagation. Chemical Geology. v. 233(3-4). p. 185-195.

De Silva, S.L., 1989a. Geochronology and stratigraphy of the ignimbrites from 21°30’S to 23°30’S portion of the central Andes of northern Chile. Journal of Volcanology and Geothermal Research, v. 37(2). p. 93-131. doi:10.1016/0377-0273(89)90065-6.

De Silva S.L., Francis, P.W. 1989b. Correlation of large ignimbrites-Two case studies from the Central Andes of northern Chile. Journal of Volcanology and Geothermal Research. v. 37(2). p. 133-149. Doi: https://doi.org/10.1016/0377-0273(89)90066-8.

Douglass, D.C., Singer, B.S., Kaplan, M.R., Mickleson, D.M., Caffee, M.W., 2006. Cosmogenic nuclide surface exposure dating of boulders on last- glacial and late-glacial moraines, Lago Buenos Aires, Argentina: interpretative strategies and paleoclimate implications. Quaternary Geochronology. v. 1. p. 43–58.

Dunai, T., 2001. Influence of secular variation of the magnetic field on production rates of in situ produced cosmogenic nuclides. Earth and Planetary Science Letters v. 193. p. 197-212.

Dunai, T.J., 2000. Scaling factors for production rates of in situ produced cosmogenic nuclides: a critical reevaluation. Earth and Planetary Science Letters. v. 176(1). P. 157-169.

Fabel, D., Harbor, J. 1999. The use of in-situ produced cosmogenic radiocublides in glaciology and glacial geomorphology. Annals of Glaciology. v. 28. p. 103-110. doi: https://doi.org/10.3189/172756499781821968.

84 Fabel, Derek, Ballantyne, Colin K., Xu, Sheng. 2012. Trimlines, blockfields, mountain-top erratics and the vertical dimension of the last British-Irish Ice Sheet in NW Scotland. Quaternary Science Reviews. v. 55. p. 91-102.

Fabryka-Martin, J.T., 1988. Production of radionuclides in the earth and their hydrogeologic significance, with emphasis on chlorine-36 and iodine-129. Ph.D. Thesis, University of Arizona, Tucson, 400.

Farber, Daniel L., Hancock, Gregory S., Finkel, Robert C., Rodbell, Donald T. 2005. The age and extent of tropical alpine glaciation in the Cordillera Blanca, Peru. Journal of Quaternary Science. v. 20(7-8). p. 759-776. doi: 10.1002/jqs.994.

Fornari, M., Risacher, F., Féraud, G., 2001. Dating of paleolakes in the central Altiplano of Bolivia. Palaeogeography, Palaeoclimatology, Palaeoecology. v. 172(3-4). p. 269-282. doi: https://doi.org/10.1016/S0031-0182(01)00301-7.

Fritz, S.C., Baker, P.A., Lowenstein, T.K., Seltzer, G., Rigsby, C., Dwyer, G., S., Tapia, P., M., Arnold, K., K., Ku, T.-L., Luo, S., 2004. Hydrologic variation during the last 170,000 years in the southern hemisphere tropics of South America. Quaternary Research. v. 61(1). p. 95-104. doi: 10.1016/j.yqres.2003.08.007.

Gardeweg, P.M. and Ramirez, C.F., 1987. The La Pacana Caldera and the Atana ignimbrite. A major ashflow and resurgent caldera complex in the Andes of northern Chile: Bull. Volcanol., v. 49, p. 547-566. doi: 10.1007/BF01080449.

Garreaud, René D. and Battisti, David S. 1999. Interannual (ENSO) and Interdecadal (ENSO-like) Variability in the Southern Hemisphere Tropospheric Circulation. American Meteorological Society. v. 12. p. 2113-2123. doi: https://doi.org/10.1175/1520- 0442(2003)016<0281:SOTERR>2.0.CO;2.

Garreaud, R.D., Aceituno, P., 2001. Interannual Rainfall Variability over the South American Altiplano. Journal of Climate. v. 14(12). P. 2779-2789. doi: 10.1175/1520- 0442(2001)014<2779:IRVOTS>2.0.CO;2.

Garreaud, R., Vuille, M., Clement, A., 2003. The climate of the Altiplano: observed current conditions and mechanism of past changes. Palaeogeography, Palaeoclimatology, Palaeoecology. v. 194 (3054). p. 1–18.

Garreaud, R.D., Vuille, M., Compagnucci, R., Marengo, J., 2009. Present-day South American climate. Palaeogeography, Palaeoclimatology, Palaeoecology. v. 281(3-4). p. 180-195. doi: 10.1016/J.PALAEO.2007.10.032.

Glasser, Neil F., Clemmens, Samuel, Schnabel, Christoph, Fenton, Cassandra R., McHargue, Lanny. 2009. Tropical glacier fluctuations in the Cordillera Blanca, Peru between 12.5 and 7.6 ka from cosmogenic 10 Be dating. Quaternary Science Reviews. v. 28. p. 3448-3458.

Gosse, J.C., Klein, J., Evenson, E.B., Lawn, B., Middleton, R., 1995. 10Be dating of the duration and retreat of the Last Pinedale glacial sequence. Science. v. 268, p. 1329-1333.

85

Gosse, J.C., Phillips, F.M., 2001. Terrestrial in situ cosmogenic nuclides: theory and application. Quaternary Science Reviews. v. 20. p. 1475-1560.

Graf, K. 1991. Ein Modell zur eiszeitlichen und heutigen Vergletscherung in der bolivianschen Westkordillere. Bamberger Geogr. Schr. v. 11. p. 139-154.

Grenon, M., 2007. Nature around the ALMA Site-part 1. ESO: the Messenger, p. 59-63

Grosjean, M. 1994. Paleohydrology of the Laguna Lejía (north Chilean Altiplano) and climatic implications for late-glacial times. Paleogeography, Palaeoclimatology, Paleoecology. v. 109(1). p. 89-100. Doi: 10.1016/0031-0182(94)90119-8.

Grosjean, M., Núñez, L., Cartajena, I., & Messerli, B., 1997. Mid-Holocene Climate and Culture Change in the Atacama Desert, Northern Chile. Quaternary Research, 48(2), 239-246. doi:10.1006/qres.1997.1917

Grosjean, M., van Leeuwen, J.F.N., van der Knaap, W.O., Gehy, M.A., Ammann, B., Tanner, W., Messerli, B., Núñez, L.A., Valero-Garcés, B.L., Veit, H., 2001. A 22,000 14C year BP sediment and pollen record of climate change from Laguna Miscanti (23°S), northern Chile. Global and Planetary Change. v. 28(1-4). p. 35-51. doi: 10.1016/S0921-8181(00)00063-1

Hall, S.R., Farber, D.L., Ramage, J.M., Rodbell, D.T., Finkel, R.C., Smith, J.A., Mark, B.G., Kassel, C., 2009. Geochronology of Quaternary glaciations from the tropical Cordillera Huayhuash, Peru. Quaternary Science Reviews. v. 28(25-26). p. 2991-3009. doi: 10.1016/j.quascirev.2009.08.004.

Hallet, Bernard, Putkonen, Jaakko. 1994. Surface dating of dynamic landforms: Young boulders on aging moraines. Science. v. 265(5174). p. 937-940. doi: 10.1126/science.265.5174.937. Hartley, A., 2003. Andean uplift and climate change. Journal of the Geological Society, London. v. 160(1). p. 7-10. doi: 10.1144/0016-764902-083.

Hastenrath, S., Kutzbach, J., 1985. Late Pleistocene climate and water budget of the south American Altiplano. Quaternary Research. v. 24(3). p. 249-256. doi: 10.1016/0033-5894(85)90048-1.

Heisinger, B., Niedermayer, M., Hartmann, J.F., Korschinek, G., Nolte, E., Morteani, G., Neumaier, S., Petitjean, C., Kubik, P., Synal, A., Ivy-Ochs, S., 1997. In-situ production of radionuclides at great depths. Nuclear Instruments and Methods in Physics Research v. B(123). p. 341-346.

Heyman, Jakob, Stroeven, Arjen P., Harbor, Jonathan M., Caffee, Marc W. 2011. Too young or too old: Evaluating cosmogenic exposure dating based on an analysis of compiled boulder exposure ages. Earth and Planetary Science Letters. v. 302(1-2). p. 71-80.

Houston, J., Hartley, A.J., 2003. The central Andean west-slope rainshadow and its potential contribution to the origin of the hyper-aridity in the Atacama Desert. International Journal of Climatology. v. 23(12). p. 1453-1464. doi: 10.1002/joc.938

Isacks, Bryan L., 1988. Uplift of the Central Andean Plateau and bending of the Bolivian Orocline. Journal of Geophysical Research, v. 93(B4). p. 3211-3231. doi:10.1029/JB093iB04p03211.

86

Ivy-Ochs, S., Schluchter, C., Kubik, P., Denton, G.H., 1999. Moraine exposure dates imply synchronous Younger Dryas glacier advances in the European Alps and in the Southern Alps of New Zealand. Geografiska Annaler. v. 81, p. 313–323.

Jenny, C., Kammer, B., Ammann, C., 1996. Climate change in den trockenen Anden. Geographic Bernensia. Verlag des Geographischen Institutes der Universität Bern (IV).

Jomelli V., Favier, V., Vuille, M., Braucher, R., Martin, L., Blard, P.-H., Colose, C., Brunstein, D., He, F., Khodri, M., Bourlés, D. L., Leanni, L., Rinterknecht, V., Grancher, D., Francou, B., Ceballow, J. L., Fonseca, H., Liu, Z. & Otto-Bliesner, B. L. 2014. A major advance of tropical Andean glaciers during the Antarctic cold reversal. Nature. v. 513. p. 224-228.

Kanner, Lisa C., Burns, Stephen J., Cheng, Hai, Edwards, Lawrence R. 2012. High-latitude forcing of the South American Summer Monsoon during the last glacial. Science. v. 335. p. 570-573. doi: 10.1126/science.1213397.

Kelly, Meredith A., Lowell, Thomas V., Applegate, Patrick J., Smith, Colby, A., Phillips, Fred M., Hudson, Adam M. 2012. Late glacial fluctuations of Quelccaya Ice Cap, southeastern Peru. Geology. v. 40(11). p. 991-994. doi: https://doi.org/10.1130/G33430.1.

Kelly, Meredith A., Lowell, Thomas V., Applegate, Patrick J., Phillips, Fred M., Schaefer, Joerg M., Smith, Colby A., Kim, Hanul, Leonard, Katherine C., Hudson, Adam M. 2015. A locally calibrated, late glacial Be-10 production rate from a low-latitude, high-altitude site in the Peruvian Andes. Quaternary Geochronology. v. 26. p. 70-85.

Kessler, A., 1985: Zur Rekonstruktion von spatglazialem Klimaund Wasserhaushalt auf dem perusanisch-bolivianischen Altiplano. Zeitschrift fur Gletscherkunde und Glazialgeologie, 21: 107-114.

Kessler, Mark A., Anderson, Robert S., Stock, Greg M. 2006. Modeling topographic and climatic control of east-west asymmetry in Sierra Nevada glacier length during the Last Glacial Maximum. Journal of Geophysical Research. v. 111. F02002. doi: 10.1029/2005JF000365.

Kull, C., Grosjean, M., 2000. Late Pleistocene climate conditions in the north Chilean Andes drawn from a climate–glacier model. Journal of Glaciology. v. 46(155). p. 622-632. doi: https://doi.org/10.3189/172756500781832611.

Kull, C., Grosjean, M. & Veit, H. 2002. Modeling Modern and Late Pleistocene Glacio-climatological conditions in the Northern Chilean Andes (29-30°). Climatic Change. v. 52(3). p. 359-381. doi: https://doi.org/10.1023/A:1013746917257.

Kurz, M.D., 1986b. In situ production of terrestrial cosmogenic helium and some applications to geochronology. Geochimica et Cosmochimica. Acta 50. p. 2855-2862.

Kurz, M.D., Colodner, D., Trull, T.W., Moore, R.B., O'Brien, K., 1990. Cosmic ray exposure dating within situ produced cosmogenic He : Results from young Hawaiian lava flows. Earth and Planetary Science Letters. v. 97. p. 177-189.

87

Lal, D., 1988a. In situ-produced cosmogenic isotopes in terrestrial rocks. Annual Review of Earth and Planetary Sciences. v. 16. p. 355-388.

Lal, D., 1991. Cosmic-ray labeling of erosion surfaces: in situ nuclide production rates and erosion models. Earth and Planetary Science Letters. v. 104(2-4). p. 424-439.

Lal, D., Peters, B., 1967. Cosmic ray produced radioactivity on the earth. In: Sitte, K. (Ed.), Handbuch der Physik. Springer, Berlin, pp. 551-612.

Lea D. W., Pak D. K., Peterson L. C., Hughen K. A. 2003: Synchroneity of Tropical and High-Latitude Atlantic Temperatures over the Last Glacial Termination. Science. v. 301. p. 1361- 1364.

Lenters, J.D., Cook, K.H., 1997. On the origin of the Bolivian high and related circulation features of the South American climate. Journal of Atmospheric Sciences. v. 54(5). p. 656-678. doi: https://doi.org/10.1175/1520-0469(1997)054<0656:OTOOTB>2.0.CO;2.

Licciardi, J.M., Clark, P.U., Brook, E.J., Elmore, D., Sharma, P., 2004a. Variable response of western U.S. glaciers during the last deglaciation. Geology. v. 32(1). p. 81-84. doi: https://doi.org/10.1130/G19868.1.

Licciardi, Joseph, M., Clark, Peter U., Brook, Edward J., Elmore, David, Sharma, Pankaj. 2004b. Variable responses of western U.S. glaciers during the last deglaciation. Geology. v. 32(1). p. 81- 84. doi: https://doi.org/10.1130/G19868.1.

Lifton, Nathaniel A., Bieber, John W., Clem, John M., Duldig, Marc L., Evenson, Paul, Humble, John E., Pyle, Roger. 2005. Addressing solar modulation and long-term uncertainties in scaling secondary cosmic rays for in situ cosmogenic nuclide applications. Earth and Planetary Science Letters. v. 239(1-2). p. 140-161.

Lifton, N., Smart, D.F., Shea, M.A., 2008. Scaling time-integrated in situ cosmogenic nuclide production rates using a continuous geomagnetic model. Earth and Planetary Science Letters. v. 268(1-2). p. 190-201.

Lifton, Nathaniel, Sato, Tatsuhiko, Dunai, Tibor J., 2014. Scaling in situ cosmogenic nuclide production rates using analytical approximations to atmospheric cosmic-ray fluxes. Earth and Planetary Science Letters. v. 386. p. 149-160.

Lifton, Nathaniel, Marrero, Shasta M., Nishiizumi, Kuni, Reedy, Robert C., Stone, John O.H. 2016. Where now? Reflections on future direction for cosmogenic nuclide research from the CRONUS Projects. Quaternary Geochronology. v. 31. p. 155-159. doi: https://doi.org/10.1016/j.quageo.2015.04.010.

Lowell, Thomas V. 2000. As Climate changes, so do glaciers. Proceedings of the National Academy of Sciences of the United States of America. v. 97(4). p. 1351-1354. doi: https://doi.org/10.1073/pnas.97.4.1351.

88 Lucchi, Frederico, Tranne, Claudio Antonio, Rossi, Piermaria Luigi, Gallardo, Claudio, De Astis, Gianfilippo, Pini, Gian Andrea. 2009a. Volcanic and tectonic history of the El Tatio area (central Andes, northern Chile): explanatory notes to the 1:50,000 scale geologic map. GeoActa. v. 9045(2). p. 1-30.

Lucchi, F., Tranne C.A., Gallardo C., Rossi, P.L., Pini G.A., De Astis, G. Geological map of the el Tatio volcanic area (central Andes- northern Chile). 2009b. Dip. Scienze della Terra e Geologico- Ambientali, Università di Bologna.

MacGregor, K.R., Anderson, R.S., Anderson, S.P., Waddington, E.D. 2000. Numerical simulations of glacial-valley longitudinal profile evolution. Geology. v. 28(11). p. 1031-1034.doi: https://doi.org/10.1130/0091-7613(2000)28<1031:NSOGLP>2.0.CO;2.

Maldonado, A., Betancourt, J. L., Latorre, C. and Villagran, C. 2005. Pollen analyses from a 50 000-yr rodent midden series in the southern Atacama Desert (25° 30′ S). Journal of Quaternary Sci., v. 20(5). p. 493–507. doi: 10.1002/jqs.936.

Malone, Andrew G. O., Pierrehumbert, Raymond T., Lowell, Thomas V., Kelly, Meredith A., Stroup, Justin S. 2015. Constraints on southern hemisphere tropical climate change during the Little Ice Age and Younger Dryas based on glacier modeling of the Quelccaya Ice Cap, Peru. Quaternary Science Reviews. v. 125. p. 106-116. doi: https://doi.org/10.1016/j.quascirev.2015.08.001.

Marrero, Shasta M., Phillips, Fred M., Caffee, Marc W., Gosse, John C. 2016. CRONUS-Earth cosmogenic 36Cl calibration. Quaternary Geochronology. v. 31. p. 199-219.

Marrett, R. A., Allmendinger, R.W., Streckeisen, A., Zanettin, B., 1986. A chemical classification of volcanic rocks based on the total alkali-silica diagram. Journal of Petrology. V. 27. P. 745-207.

Marrett, R. A., Allmendinger, R. W., Alonso, R. N., Drake, R. E., 1994. Late Cenozoic tectonic evolution of the Puna Plateau and adjacent foreland, northwestern Argentine Andes. Journal of South American Earth Sciences. v. 7(2). p. 179-207. doi: https://doi.org/10.1016/0895- 9811(94)90007-8.

Martini, Mateo A., Kaplan, Michael R., Strelin, Jorge A., Astini, Ricardo A., Schaefer, Joerg M., Caffee, Marc W., Schwartz, Roseanne. 2017. Late Pleistocene glacial fluctuations in Cordillera Oriental, subtropical Andes. Quaternary Science Reviews. v. 171. p. 245-259.

Matteini, M., Mazzuoli, R., Omarini, R., 1999. Geochemical, petrographic and isotopic characterisation of Tultul, Del Medio and Pocitos volcanoes, Central Andes. XIV Congreso Argentino-Salta Acta II, 228 – 231.

Matteini, M., Mazzuoli, R., Omarini, R., Cas, R., Maas, R. 2000. The geochemical variations of the upper Cenozoic volcanism along the Calama-Olacapato-El Toro transversal fault system in central Andes (~24°S): petrogenic and geodynamic implications. Tectonophysics. v. 345. p. 211- 227.

Matteini, M., Mazzuoli, R., Omarini, R., Cas, R., Maas, R., 2002. The geochemical variations of the upper Cenozoic volcanism along the Calama-Olacapato-El Toro transverse fault system in

89 central Andes (~24°S): petrogenetic and geodynamic implications. Tectonophysics, v. 345(1-4). p. 211-227. doi:10.1016/S0040-1951(01)00214-1.

Matteini, M., Mazzouli, R., Omarini, R., Maas, R., 2002. Geodynamical evolution of Central Andes at 24°S as inferred by magma composition along the Calama-Olacapato-El Toro transversal volcanic belt. Journal of Volcanology and Geothermal Research. v. 118(1-2) p. 205-228. doi: https://doi.org/10.1016/S0377-0273(02)00257-3. Mayes, C. L., Lawver, L. A., and Sandwell, D. T., 1990. Tectonic history and new isochron chart of the south Pacific. Journal of Geophysical Research: Solid Earth, v. 95(B6). p. 8543-8567. doi:10.1029/JB095iB06p08543.

Meirding, Thomas C., 1982. Late Pleistocene glacial Equilibrium-Line Altitudes in the Colorado Front Range: a comparison of methods. Quaternary Research. v. 18(3). p. 289-310. doi: 10.1016/0033- 5894(82)90076-X.

Mercer, J. H., Palacios, Oscar M., 1977. Radiocarbon dating of the last glaciation in Peru. Geology. v. 5. p. 600-604.

Messerli, B., 1973. Problems of vertical and horizontal arrangement in the high mountains of the extreme arid zone (central Sahara). Arctic and Alpine Research. V. 5(3). p. A139-A147. doi: 10.2307/1550163.

Minchin J., 1882. Notes of a journey through part of the Andean table- land of Bolivia in 1882. Proceedings of the Royal Geographic Society and Monthly Record of Geography, London. v. 4(11). p. 671–676. doi: 10.2307/1800303.

National Research Council. 2010. Landscapes on the Edge: New Horizons for Research on Earth's Surface. Washington, DC: The National Academies Press. https://doi.org/10.17226/12700.

Nishiizumi, K., Winterer, E .L., Kohl, C. P., Klein, J., Middleton, R., Lal, D. J., Arnold, J. R. 1989. Cosmic ray production rates of 10Be and 26Al in quartz from glacially polished rocks. Journal of Geophysical Research: Solid Earth. v. 94(B12).

Oerlemans, J. 1986. An attempt to simulate historic front of Nigardsbreen, Norway. Theoretical and Applied Climatology. v. 37(3). p. 126-135. doi: https://doi.org/10.1007/BF00867846.

Oerlemans, J. 1991. A Model for the surface balance of ice masses: Part I. Alpine glaciers. Zeitschrift Für Gletscherkunde Und Glazialgeologie. Band 27/28. S. p. 63-83.

Ohmura, A., Kasser, P., Funk, M., 1992. Climate at the Equilibrium Line of Glaciers. Journal of Glaciology. v. 38(130). P. 397-411. doi: 10.3189/S0022143000002276.

Owen, L.A., Finkel, R.C., Haizhou, M., Spencer, J.Q., Derbyshire, E., Barnard, P.L., Caffee, M.W., 2003. Timing and style of Late Quaternary glaciation in northeastern Tibet. Geological Society of America Bulletin v. 115, p. 1356-1364.

90 Phillips, Fred M., Leavy, Brian D., Jannik, Nancy, O., Elmore, David, Kubik, Peter W. 1986. The accuculation of cosmogenic chlorine-36 in rocks: a method for surface exposure dating. Science. v. 231(4733). p. 41-43.

Phillips, Fred M., Zreda, Marek G., Gosse, John C., Klein, Jeffrey, Evenson, Edward B., Hall, Robert D., Chadwick, Oliver A., Sharma, Pankaj. 1997. Cosmogenic 36Cl and 10Be ages of Quaternary glacial and fluvial deposits of the Wind River Range, Wyoming. GSA Bulletin. v. 109(11). P. 1453-1463. doi: https://doi.org/10.1130/0016-7606(1997)109<1453:CCABAO>2.3.CO;2.

Phillips, Fred M., Argento, David C., Bourlès, Didier L., Caffee, Marc W., Dunai, Tibor J., Goehring, Brent, Gosse, John C., Hudson, Adam M., Jull, Timothy, A.J., Kelly, Meredith, Lifton, Nathaniel, Marrero, Shasta M. ,Nishiizumi, Kuni, Reedy, Robert C., Stone, John O.H. 2016. Where now? Reflections on future directions for cosmogenic nuclide research from the CRONUS Projects. Quaternary Geochronology. v. 31. p. 155-159.

Pigati, J.S., Lifton, N.A. 2004. Geomagnetic effects on time-integrated cosmogenic nuclide production with emphasis on in situ 14C and 10Be. Earth and Planetary Science Letters. v. 226. Issues 1-2. p. 193-205.

Placzek, C., Quade, J., Patchett, P.J., 2006a. Geochronology and stratigraphy of late Pleistocene lake cycles on the southern Bolivian Altiplano: implications for causes of tropical climate change. Geological Society of America Bulletin. v. 118(5-6). p. 515–532. doi: https://doi.org/10.1130/B25770.1.

Placzek, C., Patchett, P.J., Quade, J., Wagner, J.D.M., 2006b. Strategies for successful U–Th dating of paleolake carbonates: an example from the Bolivian Altiplano. Geochemistry, Geophysics, Geosystems. v. 7(5). p.1525-2027. doi: 10.1029/2005GC001157.

Placzek, C., Betancourt, J.A., Quade, J., Patchett, P.J., Latorre, C., Rech, J., Holmgren, C., English, N.B., Matmon, A., 2009a. Climate in the dry, central Andes over geologic, millennial, and interannual timescales. Annals of the Missouri Botanical Garden. V. 96(3). p. 386–397. doi: 10.3417/2008019.

Placzek, C., Quade, J., Rech, J., Patchett, P.J., Pérez de Arce, C., 2009b. Geochemistry, chronology and stratigraphy of Neogene tuffs of the Central Andean region. Quaternary Geochronology. v. 4(1). p. 22–36. doi: 10.1016/j.quageo.2008.06.002.

Placzek, C.J., Quade, J., Patchett, P.J., 2011. Isotopic tracers of paleohydrologic change in large lakes of the Bolivian Altiplano. Quaternary Research. v. 75(1). p. 231–244. doi: 10.1016/J.YQRES.2010.08.004.

Placzek, C., Quade, J., Patchett, P. 2013. A 130ka reconstruction of rainfall on the Bolivian Altiplano. Earth and Planetary Science Letters. v. 363. p. 97-108. doi: 10.1016/j.epsl.2012.12.017.

Plummer, Mitchell A., Phillips, Fred M. 2003. A 2-D numerical model of snow/ice energy balance and ice flow for paleoclimatic interpretation of glacial geomorphic features. Quaternary Science Reviews. v. 22(14). p. 1389-1406. doi: https://doi.org/10.1016/S0277-3791(03)00081-7.

91 Pritchard, Matthew E., Simons, Mark. 2002. A satellite geodedic survey of large-scale deformation of volcanic centres in the central Andes. Nature. v. 418(1). p. 167-171.

Putkonen, Jaakko, Swanson, Terry. 2003. Accuracy of cosmogenic ages for moraines. Quaternary Research. v. 59(2). p. 255-261. doi: https://doi.org/10.1016/S0033-5894(03)00006-1.

Ramirez, C. F., Gardeweg, M., 1982. Hoja Toconao, region de Antofogasta, Srevicio Nacional de Geologia y Mineria. Carta Geol. Chile. V. 54. P. 122.

Raymond, Davis Jr., Schaeffer, Oliver A. 1955. Chlorine-36 in Nature. Annals of the New York Academy of Sciences. v. 62(5). p. 107-121.

Risacher, F., Fritz, B., 2000. Bromine geochemistry of salar de Uyuni and deeper salt crusts, Central Altiplano, Bolivia. Chemical Geology. v. 167(3-4). p. 373-392. doi: 10.1016/S0009- 2541(99)00251-X.

Sagredo, E. A., Lowell, T. V. 2012. Climatology of Andean glaciers: A framework to understand glacier response to climate change. Global and Planetary Change. v. 86-87. p. 101-109.

Salfity, J.A., 1985, Lineamientos transversales al rumbo andino en el noroeste argentino: Congreso Geológico Chileno, 4th, Antofagasta, 1985, Actas, v. 1, p. 119−137.

Seltzer, G.O., 1990. Recent glacial history and paleoclimate of the Peruvian-Bolivian Andes. Quaternary Science Reviews. v. 9(2-3). p. 137-152. doi: 10.1016/0277-3791(90)90015-3.

Seltzer, G.O., 1992. Late Quaternary glaciation of the Cordillera Real, Bolivia. Journal of Quaternary Science. v. 7(2). p. 87-98. doi: 10.1002/jqs.3390070202.

Seluchi, M.E., Norte, F., Satyamurty, P., Chou, S., 2003. Analysis of Three Situations of the Foehn Effect over the Andes (Zonda Wind) Using the Eta–CPTEC Regional Model. Weather and Forecasting. v. 18(3). p. 481-501. doi: 10.1175/1520-0434(2003)18<481:AOTSOT>2.0.CO;2.

Sharma, P., Bourgeois, M., Elmore, D., Granger, D., Lipschutz, M. E., Ma, X., Miller, T., Mueller, K., Rickey, F., Simms, P., Vogt, S. 2000. PRIME lab AMS performance, upgrades and research applications. Nuclear Instruments and Methods in Physics Research Section B: Beam Interactions with Materials and Atoms. v. 172(1-4). p. 112-123. doi: https://doi.org/10.1016/S0168-583X(00)00132-4.

Small, David, Fabel, Derek. 2015. A lateglacial 10Be production rate from glacial lake shorelines in Scotland. Journal of Quaternary Science. v. 30(6). p. 509-513.

Small, David, Fabel, Derek. 2016. Was Scotland deglaciated during the younger Dryas? Quaternary Science Reviews. v. 145. p. 259-263.

Smith, Colby A., Lowell, Thomas V., Caffee, Marc W. 2009. Lateglacial and Holocene cosmogenic surface exposure age glacial chronology and geomorphological evidence for the presence of cold‐based glaciers at Nevado Sajama, Bolivia. Journal Quaternary Science. v. 24 p. 360–372. ISSN 0267‐8179.

92

Stone, J.O., Ballantyne, C.K., Field, L.K., 1998a. Exposure dating and validation of periglacial weathering limits, northwest Scotland. Geology. v. 26. p. 587-590.

Stone, J.O.H., Evans, J.M. Fifield, L.K., Allan, G.L., Cresswell, R.G. 1998b. Cosmogenic chlorine-36 production in calcite by muons. Geochimica et Cosmochimica Acta. v. 62(3). p. 433-454. doi: https://doi.org/10.1016/S0016-7037(97)00369-4.

Stone, J.O., 2000. Air pressure and cosmogenic isotope production. Journal of Geophysical Research. v. 105. p. 23753-23759. doi: 10.1029/2000JB900181.

Stone, J. O. and Ballantyne, C. K. 2005. Dimensions and deglacial chronology of the Outer Hebrides Ice Cap, northwest Scotland: implications of cosmic ray exposure dating. J. Quaternary Sci., v.. 21 p. 75–84.

Servant M, Fontes J.C., 1978. Les lacs quaternaires des hauts plateaux des Andes Boliviennes: premiéres interpretations paléoclimatiques (in French). Cahiers de l’ORSTOM, Série Géologique. v. 10(1). p. 9–23. ISSN 0029-7232.

Sylvestre, F., Servant-Vildary, S., Fournier, M. et al., 1995. Lake levels in the southern Bolivian Altiplano (19°-21°S) during the Late Glacial based on diatom studies. International Journal of Salt Lake Research. v. 4. p. 281-300. doi: https://doi.org/10.1007/BF01999113.

Sylvestre, F., Servant, M., Servant-Vildary, S., Causse, C., Ybert, J.-P., 1999. Lake-Level Chronology on the Southern Bolivian Altiplano (18°–23°S) During Late-Glacial Time and the Early Holocene. Quaternary Research. v. 51(1). p. 54-66. doi: 10.1006/qres.1998.2017.

Tassara, Andrés. 2005. Interaction between the Nazca and South American plates and formation of the Altiplano-Puna plateau: Review of a flexural analysis along the Andean margin (15°-34° S). Tectonophysics. v. 399(1-4). p. 39-57. Thompson, L.G., Davis, M.E., Mosley-Thompson, E., Sowers, T.A., Henderson, K.A., Zagorodnov, V.S., Lin, P.-N., Mikhalenko, V.N., Bolzan, J.F., Cole-Dai, J., Francou, B. 1998. A 25,000-Year tropical climate history from Bolivian ice cores. Science. v. 282(5395). p. 1858-1864. doi: 10.1126/science.282.5395.1858.

Thompson, Lonnie G. 2000. Ice core evidence for climate change in the tropics: implications for our future. Quaternary Science Reviews. v. 19(1-5). p. 19-35. doi: https://doi.org/10.1016/S0277- 3791(99)00052-9.

Thompson, Lonnie G., Mosley-Thompson Ellen, Brecher, Henry, Davis, Mary, León, Blanca, Les, Don, Lin, Ping-Nan, Mashiotta, Tracy, Mountain, Keith. 2006. Abrupt tropical climate change: Past and present. Proceedings of the National Academy of Sciences of the United States of America. v. 103(28). p. 10536-10543. doi: https://doi.org/10.1073/pnas.0603900103.

Valero-Garcés, B.L., Grosjean, M., Schwalb, A., Kelts, A.K., Schreier, H., Messerli, B. 1996. Limnogeology of Laguna Miscanti: evidence for mid to late Holocene moisture changes in the Atacama Altiplano (Northern Chile). Journal of Paleolimnology.

93 Vandervoort, D.S., Jordan, T.E., Zeitler, P.K. & Alonso, R.N., 1995. Chronology of internal drainage development and uplift, southern Puna plateau. Argentine Central Andes. Geology. v. 23(2). p. 145-148. doi: https://doi.org/10.1130/0091-7613(1995)023<0145:COIDDA>2.3.CO;2.

Vera, C., Silvestri, G., Liebmann, B., González, P., 2006. Climate change scenarios for seasonal precipitation in South America from IPCC-AR4 models. Geophysical Research Letters. v. 33(13). p. L13707. doi: 10.1029/2006GL025759. Vuille, M., 1999. Atmospheric circulation over the Bolivian Altiplano during dry and wet periods and extreme phases of the Southern Oscillation. International Journal of Climatology. v. 19(14). p. 1579-1600. doi: 10.1002/(SICI)1097-0088(19991130)19:14<1579::AID-JOC441>3.0.CO;2-N.

Vuille, M., Bradley, R.S., Keimig, F., 2000a. Climate variability in the Andes of Ecuador and its relation to tropical Pacific and Atlantic sea surface temperature anomalies. Journal of Climate. v. 13(14). p. 2520–2535. doi: 10.1175/1520-0442(2000)013<2520:CVITAO>2.0.CO;2.

Vuille, M., Burns, S. J., Taylor, B. L., Cruz, F. W., Bird, B. W., Abbott, M. B., Kanner, L. C., Cheng, H., and Novello, V. F. 2012. A review of the South American monsoon history as recorded in stable isotopic proxies over the past two millennia, Climate of the Past, 8, p.1309-1321. doi:https://doi.org/10.5194/cp-8-1309-2012, 2012.

Ward, D.J., Cesta, J.M., Galewsky, J., Sagredo, E. 2015. Late Pleistocene glaciations of the arid subtropical Andes and new results from the Chajnantor Plateau, northern Chile. Quaternary Science Reviews. v. 128. p. 98-116. doi: 10.1016/j.quascirev.2015.09.022.

Ward, D. Thornton, R., Cesta, J. 2017. Across the Arid Diagonal: Deglaciation of the western Andean cordillera in southwestern Bolivia and northern Chile. Quadernos de Investigación Geográphica. v. 43(2). p. 667-696. doi: http://doi.org/10.18172/cig.3209.

Zech, Roland. Kull, Christopher, Veit, Heinz. 2006. Late Quaternary glacial history in the Encierro Valley, northern Chile (29°S), deduced from 10Be surface exposure. Paleogeography, Paleoclimatology, Plaeoecology. v. 234(2-4). p. 277-286. doi: https://doi.org/10.1016/j.palaeo.2005.10.011.

Zech R., Kull, Ch., Kubik, P.W., Veit., H. 2007. LGM and Late Glacial glacier advances in the Cordillera Real and Cochabamba (Bolivia) deduced from 10Be surface exposure dating. Climate of the Past Discussions, European Geosciences Union (EGU). v. 3(3), p.839-869. doi: .

Zech, R., May, J., Kull, C, Ilgner, J., Kubik, P., Veit, H., 2008. Timing of the late Quaternary glaciation in the Andes from ~15 to 40° S. Journal of Quaternary Science. v. 23(6-7). p. 635-647. doi: 10.1002/jqs.1200.

Zech, J., Zech, R., Kubik, P. W., Veit, H., 2009. Glacier and climate reconstruction at Tres Lagunas, NW Argentina, based on 10 Be surface exposure dating and lake sediment analyses. Paleogeography, Paleoclimatology, Paleoecology. v. 284. p. 180-190.

94 Zhou, J., Lau, K.M., 1998. Does a Monsoon Climate exist over South America? Journal of Climate. v. 11(5). p. 1020-1040. doi: https://doi.org/10.1175/1520- 0442(1998)011<1020:DAMCEO>2.0.CO;2

Zreda, Marek G., Phillips, Fred M., Elmore, David, Kubik, Peter W., Sharma, Pankaj, Dorn, Ronald I. 1991. Cosmogenic chlorine-36 production rates in terrestrial rocks. Earth and Planetary Science Letters. v. 105(1-3). p. 94-109.

Zreda, Marek, G., Phillips, Fred M., 1995. Insights into alpine moraine development from cosmogenic 36Cl buildup dating. Geomorphology. v. 14(2). p. 149-156. doi: https://doi.org/10.1016/0169- 555X(95)00055-9.

95 Appendix A.: Raw AMS data from SPN and Tatio

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