SURFACE EXPOSURE DATING OF STREAM TERRACES IN THE CHINESE PAMIR: GLACIAL CHRONOLOGY AND PALEOCLIMATIC IMPLICATIONS

THESIS

Presented in Partial Fulfillment of the Requirements for the Degree Master of Science in

the Graduate School of The Ohio State University

By

Benjamin T. Kirby, B.S.

The Ohio State University 2008

Thesis Committee:

Professor Lindsay Schoenbohm, Advisor Approved By:

Professor Garry McKenzie ______

Professor Doug Pride Advisor Geological Sciences Graduate Program

ABSTRACT

The extreme spatial and temporal variability of climate systems in the Himalayas

has thus far limited our understanding of the evolution of the paleoclimate in

Central Asia and Tibet during the late Quaternary. Given the extreme pressures

currently being placed upon water resources in Asia by increasing populations, a

detailed understanding of the past and present climate has never been more sought-after by those seeking to mitigate climate change. In the Chinese Pamir, a region in which the South Asian monsoon and the mid-latitude westerlies vie for influence over the regional climate, surface exposure dating was performed on glacially-derived stream terraces to investigate the relative influences of these competing systems as well as possible climate teleconnections in the Northern

Hemisphere. Terrace ages of 53 ± 9 ka, 46 ± 12 ka, 36 ± 2 ka, ~17 ka, 15 ± 3 ka, and ~600 years reveal a periodicity similar to well-known climate events in the

Northern Hemisphere and may be indicative of the ability of the westerlies to convey climate signals globally. Considered in conjunction with glacial chronologies throughout the Western Himalayas, these data may also have implications for the evolution of the South Asian monsoon during the last glacial cycle.

ii ACKNOWLEDGMENTS

I am indebted to my advisor, Dr. Lindsay Schoenbohm, for her guidance in the field, her encouragement and motivation thereafter, and her desire to see me succeed.

My committee members, Drs. Garry McKenzie and Doug Pride, provided meaningful input and suggestions and have been more than helpful throughout the duration of the project.

Much of this work leans heavily upon the research of my good friend Dr. Lewis

Owen at the University of Cincinnati. The magnitude of his contribution to this project, through chemistry and conversation, cannot be overstated.

I must also thank Aron Buffen, Dr. Mary Davis, Natalie Kehrwald, and Shengxue

Lei for some informative discussions and, in the case of Mr. Lei, camaraderie and assistance during the search for the Kalagile and Qiaklak Faults, and other adventures in

Western China.

iii VITA

June 29, 1978 ...... Born, Lansing Michigan 2006…...... B.S. (Honors), University of Wisconsin-Madison 2006-Present...... Graduate Research/Teaching Associate The Ohio State University

FIELDS OF STUDY

Major Field: Geological Sciences

iv

TABLE OF CONTENTS

page Abstract...... ii Acknowledgements...... iii Vita...... iv List of Figures...... vi List of Tables ...... vii

Introduction...... 1 Geographic Setting...... 4 Climatic Setting ...... 6 Regional Glaciochronology ...... 10 Fluvial Terrace Formation ...... 12 Methods...... 14 Wulu Valley Results ...... 17 Zouba Valley Results...... 21 Glacial Chronology...... 23 Global Climate Records...... 33 Paleoclimatic Implications...... 40 Conclusions ...... 45

References Cited ...... 49 Appendix: Figures and Tables ...... 59

v

LIST OF FIGURES

page 1. The Western Himalayas, with Seasonal Air Currents ...... 59 2. The Chinese Pamir and Study Areas ...... 60 3. Compiled Regional Glacial Chronology…...... 61 4. Wulu Valley...... 62 5. Wulu Valley Transect ...... 63 6. Inheritance Curves- Wulu Valley ...... 64 7. Inheritance Curve- Zouba Valley...... 65 8. Inheritance Curves- 2005 Samples ...... 66 9. Zouba Valley...... 67 10. Zouba Valley Diagram...... 68 11. Zouba Valley Transect...... 69 12. Terrace Ages with Paleotemperature Curves...... 70 13. Compiled Chronology with Hypothetical Monsoon Regression...... 71

vi

LIST OF TABLES

Table 1- CRN Data and Results...... 72

vii

INTRODUCTION

The ability to accurately predict the effects of climate change on hydrologic systems has never been more crucial. Given the recent exponential growth of the human population, current and future water resources may soon be strained beyond their limits by unsustainable demand (Parry et al., 2007). Already, water shortages are threatening the lives and livelihoods of millions worldwide, and in a global warming scenario, it is likely that the extent of the problem will increase

(Kundzewicz et al., 2007). However, the exact impact of climate change on freshwater supplies in any one region remains uncertain due to the variability observed in local responses to climatic fluctuations (Cruz et al., 2007). To assess these localized responses and to design mitigation strategies, more data about the past and present climate is needed.

Seventy-five percent of the global population lives in the lower latitudes, areas that are subject to highly variable climate oscillations such as El Nino/Southern

Oscillation (ENSO) and the Asian Monsoon (Thompson, 2000). With respect to water supply, low-latitude glaciers serve as natural buffers against climatic variability by storing and releasing water year-round. This consistent water source is relied upon by billions of people for subsistence and survival. The future of low-latitude glaciers, therefore, has social and political ramifications, in addition to purely scientific importance. General Circulation Models (GCMs) have indicated that low-latitude

1 glaciers may be the first to experience significant mass loss in a global warming scenario (Bradley et al., 2006), and thus climate change increases the need for reliable data regarding these systems.

Few regions on earth are more reliant upon glacial runoff than Central Asia and

Tibet (Thompson et al., 1997). The Tarim Basin, north of the Tibetan Plateau, is one of the most arid places on earth, with some desert areas receiving only 10 mm of precipitation annually (Murzayev, 1971). Yet, glacial meltwater from the ranges that surround the Basin, the Tien Shan, Pamir, and Kunlun Shan, enable it to support a significant population (Qi and Cheng, 1998). Another example is Northern India, where one of the world’s most populous regions lies immediately adjacent to its highest mountain chain: the Himalayas. This region receives up to 80% of its annual precipitation during the summer monsoon period (Bradley et al., 1993), yet the seasonality of this phenomenon is dampened by glacial meltwater, which comprises up to 45% of annual river discharge (Chalise et al., 2003). Were glaciers in these regions to experience significant mass loss, it is likely that the downstream effects would be immediate and pronounced. It is only through study of past glaciations that the impacts of future climate change on Himalayan glaciers can be predicted (Benn and Owen, 2002).

Reconstruction of the extent and duration of past glacial advances is facilitated by measurement of glacial landforms, such as terraces and moraines (Benn and Owen, 2002). Mapping and dating these landforms allows compilation of a glacial history that can then be compared to global paleoclimatic records, and the relative influences of climatic changes on glaciers, as well as the effectiveness of

2 GCMs in simulating the paleoclimate, may then be studied. In Central Asia and

Tibet, physical and political boundaries have thus far limited the ability of scientists to compile a glacio-chronologic dataset of sufficient density to address many of the questions that remain regarding the regional paleoclimate (Benn and Owen, 2002).

One of the least accessible of the many ranges in and around the Western

Himalayas has been the Chinese Pamir. This range is immediately adjacent to the

Tarim Basin, and its glaciers serve as the primary source of fresh water for large cities such as Kashgar in the western Basin (Qi and Cheng, 1998). Furthermore, it is situated upwind of the Western Himalayas with respect to westerly air currents, and so may play an important role in both the past and present regional climate through orographic effects (e.g. Barry and Chorley, 1992). Preliminary studies have revealed extensive glacial landforms throughout the range which may record the history of regional glacial advances, and yet only recently has a detailed investigation of these landforms been undertaken (Schoenbohm et al., 2005; Robinson et al., 2004, 2007;

Seong et al., 2008).

For this study, stream terraces in the Chinese Pamir were mapped and dated using cosmogenic radionuclide (CRN) surface exposure dating. In conjunction with data from nearby ranges (Zech et al., 2005; e.g. Owen et al., 2005), these ages may reveal much about the relative influences of competing climatic regimes over Central

Asia and Tibet during the last glacial period. By adding to the growing paleoclimatic database in the region surrounding the Western Himalayas, we hope to improve the ability of the scientific community to assess the current and future welfare of glaciers in the Chinese Pamir, and throughout Central Asia and Tibet.

3 GEOGRAPHIC SETTING

The Pamir Mountains are situated at the far western end of the Tibetan Plateau in

the northern part of the western Himalayan Syntaxis (Fig. 1). The orogen overrides

the crust of the Tarim Basin, with as much as 300 km of north-south shortening accommodated on the Main Pamir Thrust fault (Burtman and Molnar, 1994; Thomas et al., 1994; Robinson et al., 2007). Crustal shortening in the Pamir likely began in the middle Cenozoic, and extensional systems within the range initiated in the Miocene, possibly concurrent with the uplift of the Tibetan Plateau (Meyerhoff et al, 1991;

Harrison et al., 1992; Robinson et al., 2004). Extension in the range primarily trends

east-west, and is focused along the Kongur Detachment Fault system in the east of the

range and the Karakul Rift system in the west (Arnaud et al., 1993; Brunel et al.,

1994; Robinson et al., 2004). The east-west trending Muji Range and north-south

trending Chinese Pamir separate the high topography of the syntaxis (> ~4000m

elevation) from the arid, low lying Tarim Basin (<~1200m elevation) (Fig. 1). The

topography of the Chinese Pamir is dominated by two mountain peaks, Kongur Shan

(7719m) and Muztagh Ata (7654m), which lie in the footwall of the west-dipping

Kongur Detachment (Robinson et al., 2004, 2007; Schoenbohm et al., 2005; Fig. 2).

The massifs are ~1500m higher than nearby peaks in the range, and are deeply incised by glaciers and glacially fed rivers (Schoenbohm et al., 2005).

Drainage in the Chinese Pamir is accommodated by two main rivers, both running parallel to the western range front in the hanging wall valley of the detachment: the

Muji River, flowing from northwest to southeast, and the Kengxuwar River, flowing

4 from southeast to northwest (Fig. 2). The rivers join near Kongur Shan, becoming the

Ghez River and cutting eastward across the range through the Ghez River gorge. The

entire western side of the Chinese Pamir drains though this narrow outlet and into the

Tarim Basin.

The landscape between the range front and the Muji River is a relatively homogeneous apron of glacial and alluvial sediments which, in places, is hundreds of meters thick. The uppermost abandoned fan surface stands 80-100 m above the river and can be correlated throughout the Muji Valley, though that correlation is often complicated by the presence of several glacial moraines which overlie the surface

close to the range (Schoenbohm et al., 2005; Seong et al., 2008). Overall, however, the uppermost surface is remarkably flat and extensive, covering up to 10 km from the Kongur Detachment to the Muji River. Evidence for soft-sediment deformation at the base of the apron suggests a single event of rapid emplacement. Glacial streams

have incised this alluvial surface, and the resulting deep channels are lined with a

succession of fill-cut terraces.

Terrace formation is likely to result from single episodes of surface

abandonment, which in turn could be related to climatic or tectonic events. Evidence

for this includes cobbles on the terrace surfaces, which often display a thick weathering patina, suggesting low to negligible erosion rates. Where the surfaces are vertically offset by one of the several minor normal faults in the valley trending sub- parallel to the main detachment, the scarps are very well preserved and are clearly visible in both aerial imagery and on the ground. Finally, the elevations of the terrace

5 surfaces above the active channel, usually tens of meters, precludes reoccupation of

the surfaces during even the most catastrophic flooding events.

CLIMATIC SETTING

The Chinese Pamir are situated within a ‘boundary zone’ between two major

circulation systems (Barry and Chorley, 1992; Fig. 1). Here, the mid-latitude westerlies and the Asian summer monsoon have overlapping and often competing

effects on temperature and precipitation (Benn and Owen, 1998). As a result, glacial

advances in the Pamir cannot immediately be linked to fluctuations in either system,

and must be considered in a broad, regional context wherein paleoclimatic records are

able to distinguish cyclicity and perturbations in both the westerlies and the Asian

Monsoon system.

In a simple, three-cell model of global air circulation, cold air descending

from the Subtropical Jet Stream forms a high pressure zone, sending air currents both

northward and southward from about 30º Latitude (Barry and Chorley, 1992). Winds

moving toward the poles are deflected to the east by Coriolis forces and form an

uninterrupted band of westerly winds between 30º and 60º Latitude (the ‘mid-latitude

westerlies’). At 60º, the winds have been heated sufficiently to rise and recirculate

southward, completing the loop. This circuit is known as the Ferrel Cell, and is

bounded in the south by the Hadley Cell (or Tradewinds) and in the north by the Polar

Cell (or Polar Easterlies) (Barry and Chorley, 1992).

6 Actual global air circulation is quite different from this idealized model.

Seasonality and variations in surface heat retention between land and water cause heterogeneity in latitudinal pressure zones (Barry and Chorley, 1992). Furthermore, elevated land surfaces can interrupt air flow and redirect or split major air currents.

Such is the case in Central Asia, where collision with the elevated Tibetan Plateau causes the westerlies to divide into a northern and southern branch each winter (e.g.

Benn and Owen, 1998; Fig. 1). The splitting of the westerlies is part of a much larger climate system over the Asian continent and is closely related to the annual pressure fluctuations which drive the summer monsoons. The northward shift and eventual abandonment of the southern branch of the Westerlies during late spring encourages the reversal of airflow throughout the region south of the Himalayas (Murakami,

1987; Barry and Chorley, 1992; Strahler and Strahler, 1996).

Multiple studies suggest that the westerlies have served to link the climate systems of the North Atlantic and Central Asia since at least the Pliocene (e.g. Porter

and An, 1995; Maslin et al., 1996). Cold climate events in the North Atlantic are recognized in countless region-wide paleoclimatic records, and the effects of sea surface temperature (SST) change on the regions immediately downwind are well known (Broecker, 1994; Cacho et al., 1999; Bard et al., 2000). ‘Heinrich Events’, periodic, large-scale releases of icebergs into the North Atlantic, are marked by an increase in the occurrence of ice rafted debris (IRD) and variations in the assemblages of marine fauna in marine sediment cores (e.g. Heinrich, 1988; Bond et al., 1992;

Broecker et al., 1992; Bond and Lotti, 1995). While the direct cause of these events is a subject of some debate, they are generally linked to a decrease in SST and so to a

7 colder North Atlantic climate. Porter and An (1995) documented that maxima in

mean grain size and magnetic susceptibility in the loess beds of Central Asia occurred

synchronously with Heinrich Events over the last ~70kyr. This long-distance

correlation was considered evidence of a link between the climates of the North

Atlantic and Central Asia via the westerlies. Even older links between the two

climates can be established by arguing that the initiation of modern NH glacial cycles

at ~2.6 Ma is evident in the loess plateau as well as north Pacific sediment cores, and

may be coeval with the onset of Asian monsoonal circulation (Maslin et al., 1996; An

et al, 2001).

The Asian Summer Monsoon is one of the most-studied climatic phenomena on the planet. The monsoon is defined as the annual response of air currents to a

pressure gradient between the Indian Ocean and the Asian continent (Ramage, 1971;

Barry and Chorley, 1992). A persistent low-pressure zone over the Tibetan Plateau

occurs in the spring, drawing moist air inland, and as the sea-breeze gains altitude

over land, its temperature drops and the moisture it carries condenses and falls as

torrential rains over much of the Indian subcontinent. This process occurs with great

regularity, beginning near the end of May every year, and is heavily relied upon for

agricultural production (Webster, 1998).

The effects of the monsoon are evident in a number of regional paleoclimate

proxies, and have been well documented (e.g. Bryson and Swain, 1981; van Campo et

al., 1982; Swain et al., 1983; Prell and van Campo, 1986; Prell and Kutzbach, 1987).

During the onset of the Indian monsoon, northerly winds blow across the Arabian

peninsula and out to sea, causing upwelling along the coast. These upwelling events

8 result in an increase in productivity and are identifiable as increased total organic carbon (TOC) in marine sediment cores from the Arabian Sea (Prell et al., 1992).

Likewise, powerful monsoonal winds transport coarser sediment, the deposition of which is evident as a periodic phenomenon within the cores. This high-resolution record indicates that the monsoon exhibits significantly more interannual variability than do the mid-latitude westerlies, a result of the fact that the monsoon is affected by more localized influences than is Ferrel Belt circulation (Webster, 1987).

Spectral analysis of monsoon paleointensity cycles reveals that peaks in monsoon intensity have varied with Milankovitch forcings, exhibiting 23 kyr and 41 kyr periodicity until 2.6 Ma, when the monsoon began responding more to the eccentricity (100 kyr) cycle (Clemens et al., 1996). This shift was coincident with the decoupling of monsoon intensity and NH ice extent (Clemens et al., 1996). The effect of this shift on glaciations in the Himalayas remains uncertain (Owen et al., 2005).

Therefore much may be learned by constraining the timing of glacial advance and retreat in regions affected by both NH climate shifts and monsoon cycles, such as the

Pamir.

The Pamir are situated upwind of the Tibetan Plateau with respect to the westerlies, near where the splitting of the air current takes place, and so experience a unique climate regime in comparison to surrounding ranges (Benn and Owen, 1998).

The range is at least partly affected by westerly air currents year round, whereas areas to the south of the Tibetan Plateau see the westerlies only in the winter and are more directly affected by the summer monsoon (Benn and Owen, 1998). Moreover, the

Pamir are more likely to experience a monsoonal influence than are ranges to the

9 north and east, placing the range within an important area of climatic overlap. Aizen et al. (2001) found that the Pamir receive two distinct precipitation maxima during the year: one in early spring, and the other in late autumn. This bimodal precipitation curve may be indicative of the dueling influences of the westerlies and the monsoon on the regional climate; the early precipitation maximum may correspond to the westerlies, and the latter to the monsoon. In comparison, the Northern and Central

Tien Shan, which lie immediately to the east of the Pamir and form the northern border of the Tarim basin, experience a single peak in precipitation in May and June, likely delivered via westerly air currents (Aizen et al., 2001).

Glaciers in the Pamir and surrounding ranges are likely to be more sensitive to changes in precipitation than in temperature (Derbyshire, 1981; Owen et al., 2005;

Seong et al., 2008) At these high altitudes (>3500 m), mean annual temperature is low, and drastic deviations are unlikely, even during the summer months; furthermore, given the distance of the Pamir from major moisture sources and the aridity of the local environment, precipitation is sporadic (Derbyshire, 1981; Aizen et al., 2001). Changes in the source direction and quantity of precipitation delivered to the range, therefore, are likely to have influenced past glaciations more than the minor temperature changes which may have occurred.

REGIONAL GLACIOCHRONOLOGY

A regional record of glacial advance and retreat may reveal much about the relative contributions of the westerlies and the monsoons to climate change. Benn and

10 Owen (1998) and Owen et al. (2005) compiled ages for glacial moraines across the

Indo-Asian collision zone and found that the timing of glacial advances, while in

many cases poorly constrained, showed regional asynchronicity on both millennial

and decadal/annual time scales (Fig. 3). This is indicative of the spatial and temporal

variability of the Himalayan paleoclimate as well as the importance of microclimate,

and highlights the need for improved spatial density of sample sites. Furthermore,

while glaciers in some regions (northern Tibet and the northern Karakoram) reached

maximum extent during the Northern Hemisphere (NH) Last Glacial Maximum

(LGM), glaciers in other regions (Lahul, Garwal, southern Karakoram) did not,

instead achieving maxima as early as 60 ka (Benn and Owen, 1998; Fig. 3). The

boundary between these two paleoclimatic regimes is as yet poorly defined.

Numerous studies have addressed glacial chronology in the Western

Himalayas (e.g. Owen et al., 1992, 2001, 2002, 2005, 2006; Phillips et al., 2000;

Sharma and Owen, 1996; Taylor and Mitchell, 2000), as well as in the western Pamir

(Zech et al., 2005) and the southern Chinese Pamir (Seong et al., 2008) through the dating of glacial moraines. Yet, there are a number of complicating factors to

consider when compiling a moraine-based glacial chronology, including post-

depositional instability of moraine surfaces, variations in glacial debris-mantling, and

the occasional difficulty in identifying moraines vs. mass wasting features (Benn and

Owen, 2002). These complications have been thoroughly addressed by the studies

cited here; however, more data is often considered necessary to verify these moraine-

based chronologies. Fluvial surfaces associated with glacial advances, while hardly

devoid of complications, are easily identified in the field and, in arid environments,

11 may experience little to no post-depositional modification. A terrace-based glacial

chronology, therefore, presents a valuable alternative to moraine studies and, when considered in conjunction with existing chronologies, may serve to reduce the uncertainty in paleoclimatic interpretation.

FLUVIAL TERRACE FORMATION

Fluvial terraces in glaciated regions provide measurable records of climatic

and tectonic change (Penck and Bruckner, 1909). In arid environments, especially,

these geomorphic features can be preserved for significant time periods (>100 ka) and

so may be used to investigate environmental fluctuations well into the past (Benn and

Owen, 2002). There are many possible external controls on fluvial dynamics: sea

level rise and fall, tectonic uplift or subsidence, and climate change all may influence

the behavior of a stream system. In an area such as the Pamir, where the Ghez River

empties into the internally drained Tarim Basin, sea level change can be ruled out as

an influence on stream behavior. Tectonic activity and climate change, however, have

both been influential in the region over the time scale in question, and so both may be addressed by studying fluvial surfaces.

Stream terraces (abandoned channel surfaces or floodplains) serve as a record

of the aggradation and degradation of a stream channel. The ability of a stream to

entrain sediment and incise its channel is sensitive to a number of climatic and

tectonic influences (Whipple and Tucker, 1999). The ‘balance model’ for aggradation

and degradation states that the behavior of the stream is the result of the relationship

12 between stream power and sediment supply (Blum and Tornqvist, 2000). That is, increasing sediment load or reducing the power (i.e., gradient or discharge) of the

stream will have the same effect: aggradation. The opposite is true for incision, where

a reduction of sediment load or an increase in stream power will cause the stream to

incise its channel, leaving behind the former stream bed as a terrace. All terraces

measured in this study are on the hanging wall of the Kongur Detachment rather than on the actively uplifting footwall. It can be assumed that fault activity has little impact on stream channel gradient, and so variations in channel morphology are likely to be due to variations in sediment supply and river discharge only.

In glaciated environments such as the Pamir, variations in sediment supply

and river discharge, and thus stream terrace formation, are thought to occur synchronously with climatic change. Molnar et al. (1994) used radiocarbon dating,

surface exposure dating, and relative dating methods to determine that nested terraces

in the northern Tien Shan corresponded to major global climate events which have

occurred since at least the mid-Pleistocene. They theorized that terrace surfaces were emplaced during the peak of glaciation, when they argued discharge was lowest and

sediment output from glaciers would have been at a maximum, and were abandoned

during deglaciation when the drop in sediment supply and an increase in discharge

promoted channel degradation. A different view is held by Formento-Trigilio et al.

(2003). In their study of stream terraces in the Alps of New Zealand, they determined

that terrace emplacement occurred immediately after peak glaciation. They proposed

that the increase in sediment supply occurred not during glacial maxima, but rather

during glacial retreat, when large swaths of unconsolidated till were exposed to

13 fluvial erosion. This increase in sediment load to glacial streams caused channel

aggradation downstream, and as time progressed and the easily-entrained sediment

was removed, sediment supply decreased and channel incision commenced, causing

terrace formation. Regardless of the exact mechanism which leads to fluvial surface

abandonment, the pacing of stream terrace emplacement should mimic that of glacial

fluctuations. Both models suggest that, as with glacial moraines, terrace abandonment

occurs during glacial retreat. The age of a terrace surface should therefore be

considered as marking the termination of a glacial stage.

METHODS

Terrace Mapping

Preliminary field investigation was performed using Advanced Spaceborne

Thermal Emission and Reflection Radiometer (ASTER) imagery with ~15 meter

resolution. Stream terraces, major faults, and drainage systems were identified in the

laboratory, with morphological interpretations later being verified in the field. Well

defined, laterally extensive terrace surfaces were evident within two valleys, Wulu

and Zouba (Fig. 2). These were selected as primary field study areas.

Terrace surface elevations were measured in the field using ‘eye heighting’, a

standard field mapping method in which, using a sighted level, elevations are determined in multiples of the height of the investigator. These rough elevation estimates were then verified using a DeLorme BlueLogger GPS receiver/recorder

(precision ~ ±1m), which was carried over and around terrace surfaces, recording 3

14 dimensional coordinates once per second. In the laboratory, the resulting datasets were uploaded into a GIS database and examined alongside the ASTER imagery

using ArcGIS software. Terrace extents and elevations were quantitatively defined

using this method.

CRN Sample Site Selection

Sample sites were selected based upon the following criteria: apparent lack of

surface erosion, presence of a cross-valley correlative surface, ease of sample

collection, and presence of a cross-fault correlative surface (the goal of a related study

is investigation the Kongur Detachment fault system). Natural wash-outs along

terrace surfaces, as well as a road-cut, were selected as sample sites due to the fact

that the recent removal of sediments from these sites minimized the excavation

necessary to retrieve CRN samples at depth. Sample sites were restricted to flat

surfaces strewn with cobbles displaying ‘desert varnish’, or patina. This reduced the

probability that the surfaces were significantly reworked following deposition. To

ensure that terraces were abandoned in response to climatic influences rather than

natural stream processes (Pan et al., 1999), cross-stream correlative surfaces were

identified for each sample site. At Wulu Valley, upper terrace surfaces were offset by a minor normal fault (Fig. 2). These surfaces were prioritized for sampling to allow the possibility of future fault slip rate assessment.

15 CRN Dating

CRN depth profile sampling was performed using the methods of Gosse and

Phillips (2000) to correct for inherited CRN concentrations. Sample pits were dug to a depth of ~2 meters, and bulk sediments were collected from 3-5cm thick horizons at

50cm intervals. Several kilograms were collected from each horizon, yielding a set of samples from each pit from depths of 2cm ± 2cm, 50cm ± 3cm, 100cm ± 3cm, 150cm

± 3cm, and 200cm ± 5cm. The latitude, longitude and elevation of each site were recorded, and topographic shielding was measured following Nishiizumi et al. (1989).

Quartz separation was performed by colleagues at Beijing University, China.

Samples were crushed and sieved, and the 180-500μm fraction was subjected to magnetic separation, heavy liquids, calcite dissolution, and frothing following the methods of Kohl and Nishiizumi (1992). Quartz separates were then shipped to the

Ohio State University. 10Be separation was performed at the University of Cincinnati, where two HF leaches were performed to eliminate meteoric 10Be. 9Be carrier was added, HF and HClO4 dissolutions were performed, beryllium was separated and purified using ion exchange chromatography (pH>7), and Be(OH)2 was precipitated and oxidized to BeO by ignition in quartz crucibles. After addition of Nb metal to the samples, 10Be/9Be ratios were measured using the accelerator mass spectrometer

(AMS) at the Purdue Rare Isotope Measurement Laboratory (PRIMELab) at Purdue

University, with the standards of Nishiizumi et al. (1999). 10Be/9Be ratios were converted to concentrations in atoms/g (quartz) using the methods of Gosse and

Phillips (2000). Inheritance values were calculated using a weighted polynomial linear regression in MATLAB (Appendix) wherein samples were weighted by inverse

16 error, and the resultant surface concentrations were converted to ages using the

CRONUS online age calculator (Balco et al., 2008) and the scaling factors of Stone

(2000).

A program which will utilize a Monte Carlo simulation to quantitatively define the uncertainty associated with depth-profile surface age models is currently under development at PRIMELab. This program is likely to greatly improve the precision of the age data reported here. For this paper, we utilize the standard errors of the regression coefficients (90% confidence interval) as a very conservative estimate of the age uncertainties.

WULU VALLEY RESULTS (38o 59’ N, 74o 32’ E)

60 km northwest of Kongur Shan, the alluvial apron which runs parallel to the

Muji Valley has been incised by a glacial stream, forming a nearly straight valley extending ~9 km from the range front to the Muji River with a width of 600-800 m.

This stream and its valley are referred to as ‘Wulu’ by this study, after the local name for the river. Based on analysis of aerial imagery, the river is fed by at least two and possibly five glacially carved alpine valleys. Analysis in the field and in the laboratory of GPS and eye-heighting data reveals 5 paired terraces at Wulu Valley and two unpaired terraces (Figs. 4, 5). Three terraces were sampled for surface exposure dating (Fig. 4).

17 T1. The uppermost surface, T1, is considered correlative with the upper alluvial surface(s) which was identified throughout the Muji river valley and dated by

Schoenbohm et al. (2005). The elevation of T1 is ~10m higher on the south side of the Wulu Valley than on the north (~76m above the stream channel in the south, vs.

~66m in the north); this is due to the presence of the normal Wulu fault which strikes northwest to southeast, with greater fault displacement in the south (Fig. 5). T1 was sampled for CRN dating at a road cut on the north side of the valley (Fig. 4).

The samples collected from T1 yielded low quartz percentages during the initial phase of processing in Beijing. Only three of the five samples were further processed for 10Be, and one of those, the surface sample, yielded trace amounts of 10Be which were below the detection limits of the AMS. This surface concentration (AMS error

> 100%) has been discarded, leaving only two data points, at 150 cm and 200 cm, to constrain the inheritance curve. An age estimation of 26 ka can be made using the curve-fitting method (Fig. 6a), but due to the lack of shallow data points to constrain the curve, this age is ascribed very little confidence. An age of 53 ± 9 ka for the upper surface in the Muji Valley (Schoenbohm et al., 2005; Fig. 8) will therefore be tentatively applied to T1. This uppermost terrace surface is likely correlative to that at

Wulu Valley (Fig. 2).

T2. Terrace 2 sits ~10 m below T1 (~54 m above the stream channel) on the north side of the valley only. Although some remnants of T2 may be present on the south side of the valley, these features are indistinct and were not mapped. T2 is narrow and heavily eroded on the west side of the Wulu fault and is more extensive to the

18 east. Due to the evidence of erosion and the lack of a cross-stream correlative

surface, T2 was not sampled for surface exposure data.

T3. Terrace 3 is ~50m above the stream channel, and is clearly manifest on both sides

of Wulu Valley (Fig. 4). As with T1, the southern T3 surface is slightly higher than

the northern surface, and this is again attributed to displacement on the Wulu Fault

(Fig. 5). The northern T3 surface is ~2.5 km long, while the southern surface is only

evident from the Wulu Fault to the valley terminus. The edge of T3 was sampled for

surface exposure dating on the northern side of the valley, east of the Wulu Fault

(Fig. 4).

Four of the T3 samples yielded appreciable quartz and provided well-constrained

10Be concentrations. The fifth sample, taken at the surface, yielded a low quartz

percentage and produced a measured 10Be concentration which was an order of magnitude greater than any other sample from the Wulu and Muji Valleys measured

by this study and Schoenbohm et al. (2005) (Table 1). This surface sample is

therefore disregarded due to the high likelihood of inherited 10Be and/or contamination. The resulting curve fit yields an age of 12 ± 8 ka (Fig. 6b), which is likely far younger than the actual age of the surface (see T4, below). As with T1, the lack of surface data to constrain the curve-fit has resulted in an erroneously young age

with low ascribed confidence. A method of Owen et al. (2006) allows extrapolation

of a surface concentration given concentrations at depth and the density of the alluvial

material. The density of 2.1g/cm3 for alluvium in the Muji Valley (after

Schoenbohm et al., 2005) yields a 10Be surface concentration of ~1.1 E+06 at/g,

19 which, in addition to the measured samples, produces an age of ~17 ka using the curve fitting method. We consider this to be a reasonable, if poorly constrained, estimate for the age of T3 (Fig. 12).

T4. Terrace 4 lies ~38m above the stream channel and is present on both the northern and southern sides of the valley, but only to the west of the Wulu Fault (Fig. 4). The southern surface was eroded and bore signs of man-made ditches or channels, perhaps for irrigation during an earlier, more humid period. The edge of the northern surface, which was relatively intact and bore no signs of anthropogenic disturbance, was sampled for CRN dating.

All five T4 samples yielded appreciable quartz percentages and provided well- constrained 10Be concentrations (Table 1). The weighted best-fit curve yields an age

of 15.4 ± 2.6 ka (Fig. 6c). Morphologically, T4 is younger than T3, and so T3 must

therefore be older than ~15 ka, lending some support to the re-calculated age for that

surface.

T5, T6, T7. The remaining Wulu terraces were mapped but were not sampled for age

control. Terraces 5, 6 and 7 were found to be 24, 20, and 8 meters above the stream

channel, respectively (Figs. 4, 5).

20 ZOUBA VALLEY RESULTS (39o 05’ N, 74o 35’ E)

Zouba Valley lies ~65km northwest of Kongur Shan along the Kongur

Detachment and is fed by several glaciated alpine valleys. The terraces here are

bisected by a small graben structure formed by the Kongur Detachment and a lesser

antithetic normal fault (Figs.1, 9). Analysis in the field and in the laboratory of GPS

and eye-heighting data reveals 4 terrace surfaces at Zouba Valley. One terrace

surface was sampled for surface exposure dating (Fig. 9).

Upper Surfaces. Given their location in the footwall of the active Kongur

Detachment Fault, the upper terrace surfaces at Zouba Valley (ZT1, ZT2, and ZT3)

are likely the result of tectonic activity rather than climate change. Episodes of fault

displacement may have created fault scarps crossing the stream channel, and the

resulting increase in channel gradient would have caused rapid headward incision and

fluvial surface abandonment. The lack of correlative surfaces on the hanging wall of

the fault may support this hypothesis (Fig. 10). The surfaces are ~43, ~38, and ~22 m

above the active stream channel, respectively (Fig. 11). A notable feature of these

upper terraces is that the inner edges, those closest to the active channel, have higher elevations than the outer edges. This may serve as evidence that the stream channel flowed more to the south during deposition of these surfaces, with the result that a

cross-section of the modern stream channel is not perpendicular to the paleo-channel.

In this oblique cross section, the inner terrace edges are upstream of the outer edges.

21 ZT4. Terrace 4 is morphologically distinct from the upper surfaces at Zouba

Valley. ~1.5km long and ~10m above the active channel, ZT4 is present on both the footwall and hanging wall of the Kongur Detachment, suggesting climatic rather than

tectonic events led to surface abandonment. ZT4 is primarily evident on the southern

side of Zouba Valley, although indistinct features on the northern side likely correlate

to this surface. The edge of the terrace was sampled for CRN dating on the hanging

wall of the Kongur Detachment.

All ZT4 samples yielded low quartz percentages (Table 1). One of the five

samples (200 cm) failed to yield appreciable quartz and was not processed further.

Of the remaining four, three produced meaningful 10Be ratios (the fourth, at 50cm, was an extremely small sample and was below the detection limits of the AMS). The three samples used for age calculation (2cm, 100cm and 150cm) had very low 10Be

concentrations. Consequently, the errors associated with these measurements, while

low in absolute terms, were large percentages of the results. The curve fitting method

yields an age of 637 years for ZT4 with a large range of error (Fig. 7).

CRN dating, and depth profile measurement in particular, is not intended to

resolve surface ages to the nearest decade or even century. Deposition and

abandonment of surfaces is a process which itself may take hundreds or thousands of

years, lending intrinsic uncertainty to any surface exposure age, no matter how

accurate the measurement. Given the uniformly low 10Be concentrations in the ZT4

samples and evidence of limited 10Be inheritance, we suggest that an age estimate of

600 years is reasonable for this surface, but allow that it may be as old as 1,000 years.

22 An age younger than 600 is also possible; however, younger ages become

increasingly unlikely due to the increasing incision rates implied.

GLACIAL CHRONOLOGY

The following glaciochronology studies have been performed in the regions

south and east of the Pamir (these data are summarized in Fig. 3):

-Owen et al. (2002) studied moraines in the Hunza Valley of the Northwestern

Karakoram using 10Be surface exposure dating. They discovered evidence of at least

four distinct glacial stages: 54.7–43.2 ka, 25.7–21.8 ka, 18.4–15.3 ka, and 10.8–9.0

ka. A possible fifth stage could not be dated, but was determined to be older than 60 ka.

- At Chitral in northern Pakistan, Owen et al. (2002) applied OSL dating

methods to moraines. Due to significant scattering of data, Owen et al. assign these

glaciations to marine isotope stages (MIS) rather than specific time intervals. An

early glaciation is thought to have occurred during MIS-3 (~60 - 30 ka). Three other

advances were tentatively identified as early Holocene, mid Holocene (~5 ka), and

Little Ice Age (~500 years).

-To the south, in the Swat Himalaya of Northern Pakistan, Owen et al. (1992)

measured loess beds overlying glacial moraines using thermoluminescence and

identified two glacial stages: ~22-18 ka and ~7-3 ka. However, Richards et al. (2000)

substantially revised these ages using optically stimulated luminescence dating on

glacial sediments. They propose ages of ~77 ka and ~38 ka for these glacial advances.

23 - At Nanga Parbat , morphologically complex exposures similarly make identification of older glacial stages difficult. Phillips et al. (2000) used OSL and

CRN dating methods to distinguish three glacial advances at 60-30 ka, 8.4-8.0 ka and

7.0-5.5 ka. Richards et al. (2000) identified a possible fourth stage at ca. 27 ka using

OSL dating, however, Phillips et al. (2000) dispute this result, citing uncertainty in the OSL methods used to obtain the ages.

-At Ladakh in northern Pakistan, Owen et al. (2006) obtained boulder ages on glacial moraines and identified 5 separate glacial stages. Two of these are older than

100 ka. The remaining three were poorly constrained, yet generally indicated glacial maxima at 50-60 ka, 40-50 ka, and 10-6 ka. The authors noted that, like many other study areas in the Himalayas, the extent of glaciations at Ladakh has progressively decreased since the local LGM at >50 ka.

-At Zanskar, southeast of Swat Himalaya, Taylor and Mitchell (2000) recognized two distinct glacial stages using OSL dating: ~78-40 ka and ~16-10 ka.

The first stage, they argue, corresponds to the local LGM, and the second could perhaps be older than the data suggest and may represent the NH LGM.

- At Lahul, Owen et al. (2001) revised previous interpretations of glacial moraine successions (Owen et al., 1995, 1996, 1997) and determined that only two glacial stages could be identified with confidence: 15.5-12 ka and 11.4-10 ka. An older, more extensive glaciation was hypothesized but could not be conclusively identified.

-In the Garhwal Himalaya, Sharma and Owen (1996) dated moraines using

OSL and lichenometry. They proposed a single, long-lived glacial stage which

24 achieved maximum extent at ~63 ka and persisted until the NH LGM. Younger moraines at Garhwal, with ages of ~17 ka and ~5 ka were attributed to brief periods of retreat within the overall glacial stage. Barnard et al. (2004) revised these data using CRN dating methods. They identified five glacial stages at Garhwal: 63-11 ka,

7 ka, 5 ka, 1 ka, and 200-300 years BP, the oldest stage being the most extensive.

These areas, considered as a whole, yield an array of glaciochronologic data points which traverses the Western Himalayas from foreland to hinterland and covers over 500 linear miles from northwest to southeast. It is within the boundaries of this area that the monsoon and the westerlies interact with one another. The maximum northern extent of monsoonal influence is a matter of some debate, and so the glacial chronology of the Pamir range, which is both the northernmost and westernmost range presented here, may yield end member constraints on the extent and timing of monsoonal influence on glaciation in the Himalaya. To date, however, only two glaciochronology studies have been conducted in the Pamir:

- In the western Pamir of Tajikistan, Zech et al. (2005) employed CRN techniques on glacial moraines. They identified five glacial stages: ~60 ka, ~47 ka,

40 ka, 27 ka, and 19 ka. It is important to note, however, that these ages were calculated using different scaling factors than those employed in the studies above, and may be significantly older than reported here (i.e., >10%). This will be further addressed in the discussion.

- In the southern Chinese Pamir at Muztagh Ata and Kongur Shan, Seong et al. (2008) identified three glacial stages using CRN dating. An older glacial stage, based upon boulders from a single moraine, was determined to be older than 150 ka.

25 A second stage ranging from >100 ka to 15 ka was assigned to several moraines

around the massif. The third stage was based upon ages obtained from terminal and

lateral moraines within glacial valleys, and consisted of 10 separate glacial advances

beginning at 17.1 ka and continuing through the Little Ice Age. These inferred glacial

stages are quite different from those measured in the Western Pamir and this study.

Pamir Glacial Chronology- Previous Studies

Zech et al. (2005) measured boulder ages from terminal and lateral moraines

in the Western Pamir of Tajikistan. They identified an extensive set of features older than ~60 ka, a second set of less hummocky moraines with ages between 47 and 11

ka, and lateral moraines with ages between 40 and 19 ka. These ages, however, were

calculated using the scaling methods of Desilets and Zreda (2003) rather than the

methods of Stone (2000), which are employed by the other surface exposure studies

cited here. Furthermore, uncertainty in the evolution of the local climate lends

uncertainty to the time-averaged erosion rate at any one location; as a result, many

glacial chronology studies, as well as this study, employ the convention of ‘zero-

erosion age’ reporting. A zero erosion age assumes that no nuclides have been

removed from the rock or sediment by erosion, and so the age reported is by default a

minimum age. Zech et al. (2005) have applied moderate erosion corrections to their

published ages for the glacial features in the Western Pamir. To facilitate comparison

with our data and other regional studies, we recalculate the ages of Zech et al. (2005).

Using the raw data of Zech et al. (2005) and the scaling factors and erosion

rates assumed for this study, the calculated ages of the features measured in the

26 Western Pamir increase significantly. The age of the oldest, most extensive surface becomes 61-75 ka, the second moraine ages become 14-61 ka, and the lateral moraine ages become 25-52 ka. The adjustment to these ages is significant in that it blurs the distinction between the first and second glacial stages, allowing for the possibility that both the extensive moraine and the smaller, hummocky moraine originated from a single, quasi-continuous glacial stage as has been suggested by Seong et al. (2008) for the southern Chinese Pamir. This glacial stage would have likely initiated before

75 ka, near which time it reached maximum extent, and then may have experienced a slow withdrawal until ~61 ka, where either another advance or a stillstand took place.

Glacial withdrawal continued after 61 ka, with a possible, brief advance at 52 ka which was extensive enough to override pre-existing lateral moraines, but did not advance far enough to obliterate the 61 ka moraine.

Seong et al. (2008) determined a glacial chronology for the Chinese Pamir based upon the lateral and terminal moraines within and around the glacial valleys of

Muztagh Ata and Kongur Shan. Seong et al. (2008) identified an extensive glacial advance with boulder ages between ~100 and 15 ka which was correlated to the last glacial cycle. Significantly, these boulder ages were verified with depth profile samples taken from glacial till near the termini of these moraines, and ages of 53 and

66 ka were the result. This may indicate that, in keeping with our interpretation of the re-calculated data from Zech et al. (2005), a large scale glacial advance took place throughout the Pamir before 75 ka and perhaps before 100 ka. The glaciers began a slow withdrawal and then entered a stillstand or perhaps re-advanced near 60 ka.

27 Citing clustered boulder ages from five glaciated valleys adjacent to the massifs, Seong et al. (2008) present evidence for no fewer than ten glacial advances since the global LGM. We interpret these glacial advances as being primarily of local significance; only one advance is evident in more than two of the valleys studied, and five are apparent only in single valleys (Seong et al., 2008). While these glacial advances could reflect global or regional climate signals, their limited spatial extent suggests these glaciers have reacted primarily to microclimatic and orographic changes taking place near the peaks throughout the Holocene. These changes may include such local phenomena as avalanches, variations in glacial debris cover, valley orientation and shading. Even if the asynchronicity of glacial advances from valley to valley is apparent rather than real because of irregular moraine preservation, as suggested by Seong et al. (2008), this in itself also points to significant local variability. It is possible that, as the peaks themselves are ~1500 m higher than any others in the surrounding range, glaciers on these particular mountains are not representative of those throughout the Pamir range as a whole.

Pamir Glacial Chronology- This Study

The stream terraces dated in this study are at the termini of fluvial systems which are each fed by multiple glaciers. This may serve to dampen the impact of localized glacial fluctuations on downstream fluvial morphology, as the advance or retreat of a single glacier among many would have limited influence on the total stream discharge. Furthermore, the glaciers which source the streams at Wulu and

Zouba Valleys are at significantly lower elevations than those on Muztagh Ata and

28 Kongur Shan and may be less sensitive than the glaciers on the massifs to fluctuations in high-altitude precipitation. Therefore, we propose that within our field study areas, stream terraces develop primarily in response to regional and global climatic influences, those which impact most or all of the glaciers upstream, and thus should correlate to only the most extensive glaciations identified by Zech et al. (2005) and

Seong et al. (2008).

An age estimate of 53 ka for the uppermost surface (T1) at Wulu Valley suggests that, in accordance with records from the Western Pamir, Muztagh Ata and Kongur

Shan, a significant climatic event took place throughout the Pamir at this time. This age matches exactly the depth profile age obtained by Seong et al. (2008) from alluvium in front of a moraine crest, and is well within the margin of error for a lateral moraine identified by Zech et al. (2005) in the Western Pamir. As both Zech et al. (2005) and Seong et al. (2008) discovered evidence for large-scale glaciation prior to this date, it is likely that the upper surface at Wulu was an active depositional surface during that stage, which likely began >75 ka. Active deposition continued during the stillstand which may have occurred between 66 ka and 61 ka, and the surface was not abandoned until after 53 ka when, after a brief advance, glaciers retreated rapidly. The surface likely aggraded during this time, but incision would have commenced shortly thereafter when the sediment supply and discharge decreased.

A lower terrace surface near Wulu Valley was dated 36 ± 2 ka by Schoenbohm et al. (2005), and a lower surface near Kongur Shan which was collected by

Schoenbohm et al. (2005) and dated by this study returned an age of 44 ± 12 ka (Fig.

29 8). Neither of these dates correspond to the moraine ages of Zech et al. (2005) and

Seong et al. (2008), and instead occur during the apparent period of gradual glacial withdrawal following the 53 ka event.

The age of Wulu Terrace 3 (T3) is poorly constrained; however, some conclusions may be drawn from our estimated age of ~17 ka. This age was also obtained by Seong et al. (2008) for two boulders on a single moraine crest (~17.1 ka), yet the lack of correlative features in any of the other valleys studied suggests that whatever triggered the creation of this moraine was local in extent or was not preserved in the Kongur/Muztagh moraines. The similarity in these ages may therefore be coincidental.

Wulu Terrace 4 (T4) yields a moderately well-constrained age of ~15 ka. None of the Kongur/Muztagh glacial advances correspond with this age, perhaps serving as further evidence of the local controls on these high-altitude glaciers. However, the termination of the long glacial stage in the Western Pamir identified by Zech et al.

(2005) from >100 ka to ~15 ka may be correlative. Were this the case, it would imply that the climatic episodes which led to the formation of terraces 2 and 3 at Wulu

Valley either did not occur in the Western Pamir or were minor in comparison to the

15 ka event. A slow glacial withdrawal beginning after the ~53 ka event, punctuated by brief stillstands or advances at ~17 ka and >17 ka (T2) would yield the observed pattern. The variations in sediment supply could lead to downstream terrace formation, but the end moraines associated with these events could be difficult to distinguish from the entire debris train deposited between 100 ka and 15 ka.

30 The age of the Zouba terrace (ZT4), and uncertainty in that age, makes correlation

problematic. Zech et al. (2005) did not sample any features younger than ~6 ka, although this does not eliminate the possibility that such features exist in the Western

Pamir. Seong et al. (2008) identified two sets of features which fall within the likely margin of error for ZT4- one at ~1.4 ka and the other at ~0.6 ka. These events were evident in two and three valleys, respectively, and so may be more representative of regional climate change than the earlier moraines. Given the brief time period in question, it is likely that the Zouba terrace corresponds to one or both of these events, or at least occurred simultaneously but due to unrelated forces.

In summary, the Chinese Pamir terraces correspond well with some glacial features measured in the Western Pamir and at Kongur Shan and Muztagh Ata. Other climatic events evident in those locations, however, either did not occur in our study areas or occurred yet did not result in terrace formation. The possible ~17 ka event marked by the formation of T3, as well as those which led to formation of the lower surfaces sampled by Schoenbohm et al. (2005), are apparently evident only in the

Pamir terraces; this may be due to lack of feature preservation in the Western Pamir and localized controls on the high-altitude glaciers of Muztagh Ata and Kongur Shan, or to the large uncertainties associated with some of these age estimates.

Comparison to Regional Glacial Chronology

The timing of glaciations in the Pamir shares some characteristics with that of nearby ranges. To the south of the Pamir, at Chitral (Hindu Kush), Owen et al.

(2002) identified an extensive glacial advance which took place between 60 ka and

31 the global LGM. Their re-calculation of some previously obtained data may extend

the age of this glacial stage to 70 ka. The valley occupied by this glacier had already

undergone significant erosion prior to this glaciation (Owen et al., 2002), raising the possibility that, like the Pamir, Chitral experienced a large-scale glacial advance prior to ~75 ka, which receded until ~60 ka and either advanced slightly or entered a stillstand. A similar pattern is evident at Hunza Valley in the Karakoram Range

(Owen et al. 2002), where an older, poorly defined glacial stage or stages occurred sometime prior to 60 ka. This glaciation was followed by a well-constrained glacial advance from 55 ka to 43 ka. The timing of this glacial stage raises the possibility that it is related to the ~53 ka event evident in the Pamir. Where the event in the

Pamir was likely a short lived advance followed by a rapid retreat, the advance in

Hunza Valley was slightly earlier and more persistent.

Swat Himalaya lies to the south of Chitral, and it is here that a pattern begins to emerge: as one moves south through study areas in the Western Himalayas, the duration of the earliest, most extensive glaciation becomes progressively longer from site to site, persisting for greater time periods as latitude decreases. This early glacial stage terminates at ~60 ka at Swat (Richards et al., 2000), ~56 ka at Nanga Parbat

(Phillips et al., 2000), 50-40 ka at Ladakh (Owen et al., 2001), 40-30 ka at Zanskar

(Taylor and Mitchell, 2000), >20 ka at Lahul (Owen et al., 2001), and 10-20 ka at

Garhwal (Sharma and Owen, 1996; Barnard et al., 2004) This pattern does not seem to apply to sites north of Swat. In the Pamir, the northernmost site in the dataset, one would expect the earliest termination of the glacial stage. This is not the case, however, as glaciation persisted or was in slow withdrawal until at least the 53 ka

32 event. The same applies to Chitral and Hunza, where glaciers were possibly

advancing at the same time that the extensive glacial stage at Swat was likely ending.

The contrast between areas north and south of Swat Himalaya will be discussed in

greater detail below.

One consistency between the glacial histories of all sites in the Western

Himalayas is that glaciations have been progressively less extensive since ~30 ka.

This pattern was noted by Owen et al. (2005) and was discovered to apply to the

Pamir by Seong et al. (2008) and this study. In all cases, the oldest measured glaciation is also the most extensive; however, this may be due to the fact that older,

less extensive glacial features were overridden by glaciers in later stages. The recent

(<30 ka) trend of decreasing glacial extent is attributed to increasing regional aridity

(Owen et al., 2005). Not only the extent, but also the duration of glacial stages has

decreased since ~30 ka, a phenomenon which may also be attributed to increasing

aridity.

GLOBAL CLIMATE RECORDS

The SPECMAP δ18O record provides an approximation of the global temperature

through time (Shackleton and Imbrie, 1990) and may be used to analyze the extent to

which glaciers in the Western Himalayas have responded to global temperature

fluctuations throughout the Quaternary. Discrepancies between the Himalayan

paleoclimate and the global temperature record can then be investigated using more

localized paleoclimate proxies, such as ice cores and individual marine sediment

33 cores, with the goal of determining the source of paleoclimatic fluctuation in the

Himalayas and Pamir (Fig. 12). Especially applicable to these ranges are the Guliya ice core (Thompson et al., 1997), the Greenland Ice Sheet Project Two (GISP2) ice core (Stuiver and Grootes, 2000), marine sediment cores in the Arabian Sea (Clemens and Prell, 2003; Schultz et al., 1998), lake levels near the Dead Sea (Bartov et al.,

2003), and the glacial record of Central Europe (e.g. Behr and van der Plicht, 1992).

The 53 ka Event (T1)

The earliest Pamir terrace (T1), possibly an indicator of a widespread climatic

event in the Pamir at ~53 ka, coincides with the early stages of Marine Isotope Stage

(MIS) 3 in the SPECMAP stack (Fig. 12a). MIS3 is characterized by a gradual

cooling trend between ~55 ka and ~25 ka, wherein fluctuations in temperature were

minor, indicating that the average global climate was relatively stable. Terrace

formation during MIS3, therefore, is unlikely to be related to global temperature

change, as no abrupt events which might influence glacial mass balance, such as rapid warming, are in evidence.

Where the SPECMAP curve shows slight warming following MIS4, the

Guliya core (Fig. 12c) displays a comparatively drastic spike in temperature. This temperature spike is followed by a commensurate drop in temperature at ~53 ka, a large amplitude climatic shift which again is not evident in the SPECMAP record.

This age corresponds almost exactly with the apparent climatic event in the Pamir which resulted in the formation of moraines in the Western and Eastern Pamir and abandonment of the upper alluvial surface in the Muji Valley. While this drop in

34 temperature may not have been explicitly responsible for the ~53 ka event in the

Pamir, it serves as an indicator that although the globally averaged climate may have been quiescent, extreme climate changes were occurring in Central Asia during this

time.

A survey of paleoclimate records from Greenland, Germany and the Arabian

Sea points to rapid Northern Hemisphere cooling immediately prior to 53 ka, and

rapid climate amelioration immediately thereafter. The GISP2 ice core, like the

Guliya core, records a dramatic cooling event at ~53 ka which is followed by a

requisite spike in temperature (Stuiver and Grootes, 2000; Fig. 12b). The magnitude and duration of this temperature swing are very similar to others which occur throughout the Pleistocene and are associated with well-documented Heinrich events in the North Atlantic; however, this particular event has not been correlated with a

Heinrich-style iceberg release (Heinrich, 1988). In Germany, soil stratigraphy provides a detailed record of glacial/interstadial cycles during the Pleistocene (Behr and van der Plicht, 1992). Here, a glacial advance persisting from 53.5 ka to 51.5 ka is marked by a sandy layer which is well-constrained by radiocarbon dating. The glacial period is followed immediately by the Glinde Interstadial (Behr and van der

Plicht, 1992).

Total Organic Carbon (TOC) records from marine sediment cores in the

Arabian Sea indicate upwelling and productivity associated with strong South Asian monsoonal winds (Schultz et al., 1998). At 53 ka, TOC records indicate that the

South Asian monsoon was undergoing a period of weakness which was followed at

~52 ka by a large amplitude increase in strength (Schultz et al., 1998). Given that

35 glacial mass balance in the Pamir is more likely to respond to fluctuations in

precipitation than in temperature (Derbyshire, 1981; Owen et al., 2005), it is possible that changes in moisture, rather than changes in temperature, are responsible for the observed moraine and terrace formation in the Pamir at this time. The summer monsoon was weak from ~55 ka to 52 ka, and so is unlikely to have caused precipitation changes in the Pamir at 53 ka. This leaves the westerlies as the most probable driver of the 53 ka event in the Pamir.

Lower Extensive Surfaces

An extensive terrace surface below T1 was determined by Schoenbohm et al.

(2005) to have a surface exposure age of ~44 ka. This age matches closely with

Heinrich Event 5 in the North Atlantic (Heinrich et al., 1988; Bond et al., 1992; Fig.

12), during which SST decreased and the westerlies likely became drier (Bartov et al.,

2003). The time period around 44 ka is also marked by cold temperatures in the

GISP2 core (Fig. 12b), a weak monsoon signal in the Arabian Sea cores, and a glacial

period in the German soil record. Unlike the 53 ka event, the 44 ka event is not

associated with glacial moraines in the Pamir, and so may simply represent a glacial

withdrawal brought about by the lack of moisture delivered by the westerlies. The

connection between glacial advance and withdrawal in the Pamir and SST in the

North Atlantic is reinforced with this finding.

The pattern begins to break down, however, with the second lower terrace

dated by Schoenbohm et al. (2005). This terrace was dated to 36 ka, a date which

falls between two Heinrich events and during a period of high amplitude, short

36 wavelength spikes in both the temperature record of the GISP2 core (Fig. 12) and the

monsoon record of the Arabian Sea cores. A temperature high occurs in the Guliya core at this time (Thompson et al., 1997; Fig. 12c), as does paleosol formation in eastern Tibetan loess sequences (Fang et al., 2003), suggesting that the entire Tibetan

Plateau was experiencing a warm climate. This may be an indication that local climatic controls, rather than the more regional influences transmitted by the westerlies, were dominant on the Plateau at ~36 ka. One explanation for this could be the fact that, at ~30 ka, global SSTs were approaching their minimum for MIS3

(Stuiver and Grootes, 2000), a trend which was reinforced by dropping SSTs in the

North Atlantic (Bond et al., 1992). It is possible that North Atlantic waters grew cold

enough during this period that, even in the absence of a Heinrich event, little moisture

was made available to the westerlies. Glaciers would have been static or retreating

during this dry period, and the ensuant decrease in albedo could have contributed to

higher temperatures. It is noteworthy that, as yet, no terraces in the Pamir have been

identified with ages between ~36 ka and ~17 ka. This may serve as further evidence

of quiescence in the Pamir during this period, as the westerlies, having crossed an

“evaporation threshold” after the 44 ka event, did not gain enough moisture to

significantly impact glaciations in the Pamir until after the global LGM (~18 ka).

Wulu Terraces 3 and 4

Our estimated age of ~17 ka for T3 suggests that it formed immediately after the

global LGM, and is therefore possibly related to the temperature change which was occurring at this time. This change is recorded globally by the SPECMAP record, as

37 well as locally by the Guliya ice core (Fig. 12). Sea surface temperatures in the North

Atlantic increased between 17.5 to 17.3 ka, and this warming period was followed

immediately by Heinrich event 1 (Lagerklint and Wright, 1999). The warming period

saw an increase in lake levels in the Dead Sea Basin indicating that evaporation and

moisture transport were increasing over the North Atlantic, and was followed by a

lake level drop coinciding with the Heinrich event (Bartov et al., 2003). As with

Heinrich event 4 at 44 ka, this drop in temperature was accompanied by a decrease in moisture availability and was possibly the driving force behind fluvial surface abandonment in the Pamir.

The next major introduction of meltwater into the Atlantic occurred at 14.6 ka

(Weaver et al., 2003), and is nearly concurrent with formation of Wulu terrace 4 in

the Pamir (~15 ka). This event, unlike Heinrich events, is not likely due to a massive

iceberg release, but instead is the result of large quantities of meltwater being released

into the world’s oceans during the initial stages of deglaciation (Weaver et al., 2003;

Clark et al., 2002). The source of this water, termed “meltwater pulse 1”, remains a

matter of some debate, but it has been well established that sea levels rose at

remarkably high rates during this period (>40mm a-1) and that global oceanic and air

circulation patterns were likely affected (Weaver et al., 2003; Clark et al., 2002).

Additionally, this age corresponds with initiation of modern, post-glacial monsoon

circulation as evidenced by Arabian Sea cores (Zonneveld et al., 1997). Following a

period of abnormally weak monsoons during the global LGM, monsoon intensity

rapidly increased to its highest level during the Holocene at ~14.6 ka.

38 Given the global nature of climate change at 14.6 ka, it is difficult to assign

terrace formation in the Pamir to any one phenomenon. It is possible that the

presence of meltwater in the North Atlantic led to the same drop in evaporation, and

hence drying of the westerlies, that occurred during Heinrich events 1 and 4.

However, meltwater pulse 1 does not appear to have caused a drop in North Atlantic

SSTs (Lagerklint and Wright, 1999), rendering this scenario unlikely. Another possibility is that renewed monsoon intensity caused reorganization of wind currents

over the Tibetan Plateau. A stronger monsoon could cause earlier and more effective

seasonal diversion of the westerlies to the north, possibly limiting the time during

which the Pamir were exposed to the moisture-bearing current (Ono and Irino, 2004).

Zouba Terrace

An age estimate of ~600 years for Zouba terrace 4 (ZT4) raises the possibility that

this terrace corresponds to a well known climatic oscillation in the late Holocene, the

Medieval Warm Period (800-1400AD) and the Little Ice Age (1400-1800AD). The

precise timing and duration of these events, and even their very existence, is a subject

of some debate (Keigwin et al., 1996; Hughes and Diaz, 1994; Grove and Switsur,

1994), yet evidence suggests that SSTs in the Atlantic varied significantly during this

period (e.g. Keigwin et al., 1996). The Little Ice Age (LIA) was likely triggered by a

decrease in insolation, and resulted in cooler temperatures, as well as glacial

advances, throughout much of Europe and North America. Though poorly

constrained, the age of 600 years for ZT4 suggests that glaciers in the Pamir retreated

at the very beginning of the LIA. As with other Pamir terraces, the age of ZT4 may

39 indicate that, in contrast with European and North American glaciers, Pamir glaciers

have tended to advance during periods of warmth in the North Atlantic and retreat

when SSTs cool. The age estimate of ZT4 is subject to large uncertainties, however,

and this terrace may have been created at any time during the Medieval Warm Period

or LIA.

PALEOCLIMATIC IMPLICATIONS

An examination of the data presented here, in addition to data compiled by other

studies (e.g. Owen et al., 2005), reveals several patterns suggested by the glacial chronology of the region which will be presented here as the following hypotheses:

(1) the timing of glacial advances in the Pamir, and to a lesser extent, at Hunza and

Chitral, was out of phase with Northern Hemisphere ice sheets during the late

Pleistocene; (2) the relative influence of the westerlies over glaciers in the Western

Himalayas has behaved as a function of North Atlantic SSTs; (3) as the monsoon weakened throughout the Pleistocene, the northernmost extent of monsoonal precipitation retreated steadily to the south, a regression which is apparent in the glacial record; and (4) a ‘dead-space’ in which there was little precipitation persisted

between the respective zones of influence of the monsoons and westerlies throughout

the Pleistocene.

Evaporation and Temperature: A Glacial ‘Phase Shift’

There is a strong correlation between inferred global temperature over the last

glacial cycle and the extent of Northern Hemisphere ice sheets (Miller and de Vernal,

40 1992; Andrews et al., 1988): colder weather resulted in ice sheet growth throughout

the Pleistocene, and as a result, the global temperature minimum at ~18 ka was

accompanied by a maximum in ice sheet extent throughout Europe and North

America. While this temperature minimum did occur on the Tibetan Plateau at ~18 ka

(Fig. 12c; Thompson et al., 1997), the maximum glacial extent in the Pamir and

throughout much of the Western Himalayas occurred earlier (~60 ka), as discussed

previously. Owen et al. (2005) note the widespread trend of increasingly limited

glaciations in the Himalayas during the Pleistocene, a finding supported by the ages

of the Pamir terraces. Glacial extent in the Western Himalayas, especially in areas

north of Swat Himalaya, decreased with global temperature during the late

Pleistocene, and thus was out of phase with NH ice sheets.

One phenomenon inversely related to temperature, and a possible explanation for

this phase shift, is evaporation. For example, in the mid-Holocene, an increase in NH insolation was accompanied by an increase in average SSTs, and thereby, an increase in evaporation (An et al., 2005). This period of increased marine evaporation coincides with maximum Holocene lake levels throughout India (An et al., 2005;

Enzel et al., 1999) and humidity across western China (An et al., 2005). Conversely,

the Southern Hemisphere was experiencing a summer insolation minimum during this

period, which coincides with maximum Holocene aridity in South America (An et al.,

2005; Baker et al., 2001). While the paucity of data prevents extrapolation of this trend into the Pleistocene for central Asia, the basic mechanisms affecting evaporation and precipitation were likely similar during this epoch. Because glaciers in the Pamir are precipitation rather than temperature sensitive, decreasing SSTs

41 during the Pleistocene may have led to increasing aridity and glacial retreat in the

Pamir and the surrounding ranges, yet simultaneously contributed to the advance of continental ice sheets.

Global Climate Teleconnections: SSTs and the Pamir

There is mounting evidence that SST fluctuations in the North Atlantic have had wide-ranging impacts on the global climate. Manifestations of Heinrich events can be found in ice and sediment cores on five continents, including Asia (Hemming,

2003), although the climatic fluctuations evident in the cores may not be directly caused by SST variations. Relationships between alpine glacial stages and Heinrich events have been established in regions quite distant from the North Atlantic. For example, Clark and Bartlein (1995) discovered that glacial advances in the western

United States coincided with Heinrich events during the late Pleistocene. It is not unreasonable, therefore, to suggest that there may be a link between Heinrich events

and glaciation in the Pamir, although to our knowledge no such link has yet been proposed. Additionally, there is a broad correlation between the Pamir terraces and warm periods in the Guliya ice core (Thompson et al., 1997; Fig. 12C). Temperatures on the Tibetan Plateau may therefore play a role in terrace formation in addition to moisture availability. This is reasonable given that, even in the absence of westerly precipitation, glaciers would be unlikely to withdraw significantly during periods of extreme cold.

That fact that SSTs in the North Atlantic are currently rising (Parry et al., 2007) while glaciers in the Pamir are apparently at their minimum Holocene extent (Seong

42 et al., 2008) suggests that the link between SSTs and glacial advances may be

complex, and that another phenomenon with close ties to SSTs, the monsoon, may be

a complicating factor. During the late Pleistocene, the northernmost extent of the monsoon was significantly farther south than present day (e.g. Shi, 2002). It has been

proposed that during periods of the last glaciation, the westerlies did not divert to the

north of the Tibetan Plateau in the summer, and instead flowed uninterrupted over the

Himalayas year-round (Ono and Irino, 2004). The present situation, in which the

monsoons and westerlies interact with one another over the western Himalayas, may

serve to disturb westerly air flow to the extent that precipitation is significantly

reduced over the Pamir.

Pleistocene Regression of the ‘Monsoon Regime Boundary’

The gradual withdrawal to the south of the maximum monsoon extent during the

late Pleistocene may be evident in the glacial record of the western Himalayas (Fig.

13). As noted previously, the character of glacial stages in the study areas south of

Swat Himalaya is different from that to the north. Specifically, the persistence of the

earliest, most extensive glacial stage apparently increased with decreasing latitude.

This pattern could be explained with a ‘monsoon regime boundary’, north of which

the precipitation contributed by the monsoon is minimal. Given that monsoon

intensity, and therefore its inland extent, is closely tied to SSTs (Tschuck et al.,

2004), and that global SSTs decreased throughout the last glacial cycle (Shackleton

and Imbrie, 1990), it follows that this regime boundary would gradually have receded

during the late Pleistocene.

43 We propose that as ranges emerged to the north of the monsoonal regime

boundary the abrupt decrease in precipitation led to the termination of the extensive glacial stage. These terminations progressed from Swat to the south during a period of 40 thousand years in the western Himalayas. There is evidence that glaciation continued in the wake of this withdrawal at Swat and Nanga Parbat (Richards et al.,

2000; Phillips et al., 2000); we suggest that these re-glaciations may mark the southernmost extent of westerly precipitation during the last glacial, as withdrawal of monsoonal influence would allow the westerlies to advance further to the south than usual (Ono and Irino, 2004).

‘Dead Space’

In each case of glacial termination discussed above, the glacial record suggests that at least 10 thousand years elapsed before initiation of the next glacial advance

(Owen et al., 2005; Fig. 13). The absence of glacial advances in the wake of the retreating monsoon translates to a ‘dead space’ extending approximately 300 km from the northern edge of the monsoon regime boundary (Fig. 13). The glacial chronology at Ladakh provides an example of this ‘dead space’: glaciers here began retreating at

~41 ka (Owen et al., 2001), and yet at Nanga Parbat to the north and at Zanskar to the south, glaciers were advancing (Phillips et al., 2000; Taylor and Mitchell, 2000).

Zanskar was still within the monsoon regime boundary at this time, and Nanga Parbat may have been undergoing re-glaciation fed by westerly precipitation.

The presence of this apparent ‘dead space’ suggests active diversion of precipitation from the area immediately to the north of the monsoon regime

44 boundary. As we suggest above, terrace ages may indicate that a strong monsoon interferes with westerly circulation in some way in and around the Pamir, reducing the likelihood of precipitation. This interaction between the monsoons and westerlies could explain the apparent ‘dead space’ between them, with monsoonal winds causing a northward diversion of westerly air currents and thus reducing the likelihood of precipitation. Alternatively, the upwelling air currents on the Tibetan

Plateau which drive the monsoons diverge in the upper atmosphere and fall back to earth; the ensuing cold, high pressure zone could impede precipitation in the areas around the Plateau and might also migrate as a function of monsoon intensity

(Kripalani et al., 1997).

CONCLUSIONS

Mapping and surface exposure dating of stream terraces at two field study areas in the Chinese Pamir, as well as data from previous studies (Schoenbohm et al.,

2005; Zech et al., 2005; Seong et al., 2008), have established the timing of glaciations throughout the late Quaternary in the Pamir. At Wulu Valley, 5 paired terraces and 2 unpaired terraces were mapped, and surface exposure dating was performed on terraces 1, 3 and 4. Due to samples with low quartz percentages, T1 and T3 failed to yield complete depth profiles. A very rough age estimate for T1 based upon two data points is 26 ka; however the surface is likely to be much older. Based upon previous studies, we suggest an age of 53 ka for T1. T3 did not yield a usable surface sample, and as a result, the age calculation in poorly constrained at ~12 ka. We suggest an

45 age of ~17 ka for this surface based upon an alternative calculation method and

superposition. The age of T4 is based upon a full depth profile, and is well

constrained at ~15 ka. At Zouba Valley, four terraces were mapped. The three

uppermost surfaces were likely to be tectonically controlled, and so our focus was on

the lowermost surface, terrace 4 (ZT4). Three usable samples for ZT4 provide a

depth profile age of ~600 years. Though poorly constrained, we believe an age of

‘several hundred years’ is quite reasonable for this terrace.

Each terraced valley was fed by multiple glaciers upstream, suggesting that the climate events which caused terrace formation were, at the least, regional in extent.

This is in contrast to the moraines measured by Seong et al. (2008) on the flanks of

Muztagh Ata and Kongur Shan, which likely correspond to more localized glacial

fluctuations near the peaks. Derbyshire et al. (1981) proposed that glaciers in the

Pamir were more likely to respond to changes in precipitation than in temperature.

Given that the northernmost extent of the South Asian monsoon was well to the south

of the Pamir during the last glacial cycle, it is therefore probable that terrace

formation in the Pamir corresponded to changes in precipitation delivered by the

westerlies.

In all, the Pamir terraces record six distinct climatic events with the following

ages: 53 ka, 44 ka, 36 ka, ~17 ka, ~15 ka, and a few hundred years ago. All but one

of these ages (36 ka) may correspond to well-known climate events in the North

Atlantic (primarily Heinrich events). While the data do not support inference of a

causal relationship between Heinrich events and stream terraces, the similar pacing of

climatic fluctuations in the North Atlantic and the Pamir is unlikely to be

46 coincidental. Some paleoclimatic records both upwind and downwind of the Pamir

display evidence of this same periodicity, suggesting that Northern Hemisphere

climate change during the last glacial cycle, transmitted via the westerlies, had wide-

ranging impacts.

Another factor influencing glacial timing in the western Himalayas during the last glacial cycle may have been a persistent ‘dead space’ between the areas receiving

moisture from the westerlies and the monsoons. Such a ‘dead space’ implies that the

monsoon affects the regional climate well to the north of the boundary beyond which

no monsoonal moisture is received. It may have been northward incursion of the

monsoonal regime and dead space which led to formation of T4 at ~15 ka.

The extreme spatial and temporal variability of climate systems in the Himalayas

has thus far limited our understanding of the evolution of the Tibetan paleoclimate

during the late Quaternary. The addition of the Pamir stream terraces to a growing

glacio-chronologic database in the Western Himalayas, while providing important

end-member constraints on the evolution of the monsoons and westerlies, has perhaps

raised more questions than it has answered. For instance, the 36 ka event, which may

be unique to the Chinese Pamir, remains mysterious, as does the general mechanism

by which glaciers in the Pamir may respond to the periodic climate fluctuations

marked by North Atlantic Heinrich events. The ages of several Pamir terraces are

poorly constrained, and at Wulu Valley, several lower terraces remain to be dated and

may provide important insight into the Holocene climate of the Pamir. Clearly,

further work is needed, in the Pamir and elsewhere, to compile a dataset of sufficient

density to address these questions.

47 If glacial extent in the Pamir is indeed related to sea surface temperatures, there may be cause for concern among populations who rely upon those glaciers for survival. Past SST increases have been accompanied by increases in monsoon intensity, and the results of this glacio-chronologic study suggest that intense monsoons have limited moisture availability in the Pamir. Global SSTs are currently increasing (Parry et al., 2007), a fact which should thus be taken into account by those managing water resources in and around the Pamir. Further data regarding the past and present climate, as well as models of the future climate, are needed to better meet this challenge.

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60 74° 30’ E 74° 45’ E 75° 00’ E 75° 15’ E 75° 30’ E

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r (Schoenbohm et al., 2005). Three us t surface ages were collected by Ghe z R. Schoenbohm (2005) and are summa-

38° 45’ N 39° 15’ N Kongur Shan rized further in Table X. Red and blue 7719m. asl squares enclose B) and C), respec- 5-13 Kengxuwar tively. (T1) 38° 30’ N ~53ka R. B) ASTER image of Wulu Valley, running for ~10km and incising the broad alluvial apron thought to

5-30 38° 15’ N correlate with the “Upper Surface” (Lower Surface) K throughout the Muji Valley. White ~44ka o Muztagh Ata n g u 7654m. asl square encloses Figure 4. r D et ac hm C) ASTER image of Zouba Valley,

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74° 30’ E 74° 45’ E 75° 00’ E 75° 15’ E 75° 30’ E 2km B 2km C

Kongur Detachment

Fig. 6 Zouba Valley

Fig. 4 Wulu Valley

Wulu Fault

61 3 x10yrs BP 70oE 75oE 80oE 85oE 90oE 100 90 80 70 60 50 40 30 20 10 0

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o o o o o o o 60 E 70 E 80 E 90 E 100 E 110 E 120 E Fig. 3 Compiled glacial chronology for the entire Western Himalayas, after Owen et al. (2005), presented with a precipitation map of the region (upper right) whose boundary is delineated by the red square on a DEM of Asia (bottom right). Data, in descending order, from Zech et al. (2005), Seong et al. (2007), this study, Owen et al. (2002), Owen et al. (1992), Phillips et al. (2000), Owen et al. (2006), Taylor and Mitchell (2000), Owen et al. (2001) and Sharma and Owen (1996). Precipitation map adapted from Owen et al. (2005). SPECMAP curve adapted from Shackleton and Imbrie (1990). The durations of glacial stages illustrated here are uncertain, and are depicted based primarily upon the morphological interpretations of the authors and of Owen et al. (2005). 78o 32` 00`` 78o 32` 15``

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Fig. 4 ASTER imagery of the terminus of Wulu Valley, with Wulu Fault (top) and terrace outlines (bottom). Terrace outlines are composed of three dimensional GPS data points recorded in the field and are displayed here in ArcGIS v. 9 as georectified XY datasets (see text). CRN sample sites are illustrated on lower diagram (see Table). Red line in upper diagram denotes transect location (Fig. 5). See Fig. 2 for location.

63 on the fault

meters above streambed based ions are 1, T3 and T4 1, were 0 80 70 60 50 40 30 20 10 1000 T1 T3 T4 T5 T6 N T7 400 600 800 T6 distance along transect (m) distance T4 = foreground T3 200 T2 View Upstream Wulu Valley Terraces Valley Wulu

T1

3500 3480 3460 3420 3440 0 meters asl meters Fig. 5 Fig. Transect of Wulu Valley terminus, perpendicular to river channel (Fig. 4). ~10x vertical exaggeration. Terrace surface elevat Terrace ~10x vertical exaggeration. 4). channel (Fig. river perpendicular to terminus, Valley Wulu of Transect upon GPS data and other field measurements (see text). The Wulu Fault runs parallel to the transect, in the background. Throw in the background. the transect, to runs parallel Fault Wulu The (see text). upon GPS data and other field measurements T surfaces upper terrace slightly higher in the south wall of valley. are and as a result, the south (right), to increases CRN dating in the northsampled for wall of the valley.

64 T1 Wulu T3 Wulu 6-004 depth profile 6-008 depth profile 0 0 A. B.

50 50

100 100 depth (cm) depth (cm) 150 150

200 200 10 Be Concentration: 1.74 E+06 atom/g 10 Be Concentration: 6.35 E+05 atom/g Age: 26.4 ka Age: 12.01 +/- 8.6 ka

250 250 0 0.5 1 1.5 2 2.5 012345678910 concentration (at/g) 6 5 x 10 concentration (at/g) x 10

65 T4 Wulu 6-006 depth profile 0 C.

50

100 depth (cm) 150

200 10 Be Concentration: 8.26 E+05 atom/g Age: 14.39 +/- 2.6 ka

250 012345678910 5 concentration (at/g) x 10 Fig. 6 Cosmogenic depth profiles for Wulu Valley terraces. Samples are depicted with depth (vertical) and AMS measurement (horizontal) error bars. See text for age calculation methods. 4 x 10 T4 Zouba concentration (at/g) concentration Be Concentration: 0.35 E+05 atom/g Be Concentration: 6-018 depth profile 10 Age: 637 +/- 2,486 yr Age: 345678910 012 0 Fig. 7 Fig. Cosmogenic depth profile for Zouba Valley terraces Samples are Samples are terraces Valley Zouba for depth profile Cosmogenic depicted error with depth (vertical) (horizontal) and AMS measurement age calculation methods. See text for bars.

50

100 150 200 250 depth (cm) depth

66 Lower Terrace T1 5-10 depth profile 5-13 depth profile 0 0

50 50

100 100 depth (cm) depth depth (cm) depth 150 150

10 200 Be Concentration: 2.01 E+06 atom/g 200 10Be Concentration: 2.85 E+06 atom/g Age: 35.7 +/- 2.3 ka Age: 52.5 +/- 8.7 ka

250 250 0 0.5 1 1.5 2 2.5 0 0.5 1 1.5 2 2.5 3 6 concentration (at/g) 6 concentration (at/g) x 10 x 10

67

Lower Terrace 5-30 depth profile 0 Fig. 8 50 Cosmogenic depth profiles for Muji Valley terraces. These profiles were collected in 2005 and profiles 5-10 and 5-13 were processed by Schoen- bohm et al. (2005). Profile 5-30 was processed by this study. See Figure 2 100 for profile locations. Samples are depicted with depth (vertical) and depth (cm) AMS measurement (horizontal) error bars. See text for age calculation 150 methods. 200 10 Be Concentration: 2.14 E+06 atom/g Age: 43.7 +/- 12 ka

250 0 0.5 1 1.5 2 2.5 3 6 concentration (at/g) x 10 74o35’15’’ 74o35’45’’

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Fig. 9 ASTER imagery of the terminus of Zouba Valley, with Kongur Detachment Fault and minor antithetic normal fault (top) and terrace outlines (bottom). Terrace outlines are composed of three dimensional GPS data points recorded in the field and are displayed here in ArcGIS v. 9 as georectified XY datasets (see text). Viewpoint for Fig. 10 and CRN sample site are displayed (bottom), as well as the transect for Fig. 11 (top).

68 ZT1 ZT2 N

ZT3 Zouba Valley Terraces ZT4 (View South) T3S

ZT3? 69

ZT4

Sample Site

Fig. 10 Simplified diagram of Zouba Valley terraces, based upon field sketches. Viewpoint shown in Fig. 9. Diagram employs slight vertical exaggeration, as well as some warping for perspective. ZT4 was sampled for CRN dating.

sed

meters above streambed 0 40 50 20 30 10 800 ZT1 = foreground 600 (Footwall) View Upstream Zouba Valley Terraces Valley Zouba ZT2 400 ZT3 Basement? ZT4 distance along transect (m) distance N 200

Lateral Moraine

3880 3910 3870 3890 3900 0 meters asl meters Fig. 11Fig. Kongur Detachment the runs through The as other field survey as well methods (see text). upon GPS data collected in the field, surfaces contradictory a apparent be explained by The of the upper terrace gradient may the transect. to parallel foreground, CRN dating in the hanging wall (foreground). ZT4 was sampled for of ZT3 (see text). emplacement since shift flow in stream Transect of Zouba Valley, perpendicular to modern Zouba River (Fig. 9). ~12x vertical exaggeration. Terrace elevations are ba are elevations Terrace ~12x vertical exaggeration. 9). (Fig. River modern Zouba perpendicular to Valley, of Zouba Transect

70 face 1 TerraceTerrace 4 3 Lower5-10 Sur Upper5-30 Surface Terrace5-13 MIS 1 1 e d 0 a b c 2

O (Normalized) 5 18

δ 3 4 -31 A SPECMAP -1

-33

-35 Heinrich Event

oo -37 /

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-45

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Fig. 12

Terrace ages (red bars) with 95% confidence bounds displayed with three δ18O records: SPEC- MAP (Shackleton and Imbrie, 1990), GISP2 (Stuiver and Grootes, 1992) and Guliya (Thompson et al., 1997). Heinrich events from Bond et al. (1992). Age estimate for Terrace 3 is constrained morphologically: T3 may not be younger than T4 (within error) and must not be older than the upper surfaces (within error). Although poorly constrained, the pacing of terrace formation in the Chinese Pamir may be similar to that of North Atlantic Heinrich-style events. The absence of a terrace between 5-10 and Terrace 3 may be the result of decreased moiture availability over the North Atlantic during late MIS3 and MIS2.

71 1 2,3 1 2,3 1 2,3 1 2,3 5 5 4 5 4 5 4 4 12 12 12 12 6 7 8 6 7 8 6 7 8 6 7 8 9 9 9 9 10 10 10 10 11 11 11 11

50-60ka 40-50ka 30-40ka 20-30ka

100 90 80 70 60 50 40 30 20 10 x10 3 yrs BP tern

Pamir Multiple 1 Wes 1 2,3 Multiple 2 5 Pamir Westerly Influence? 4 Chinese 12

6 7 8 3 THIS 9 STUDY 10 4 11 Hunza

72 5 Chitral 9 Uncertain 10-20ka Influence 6

Swat Monsoon Influence? 7

Westerly Influence 8 Parbat Nanga Dead Space 9 soon ce Mon Poorly Constrained 10

Influen Ladakh 11

Zanskar 12 = Guliya Ice Core Fig. 13 Precipitation map and glacial chronology for the Western Himalayas (after Owen et al. (2005), see Fig. 3 for other data sources). Precipitation sources are hypothetical, based upon likely northernmostLahul extent of the modern and LGM monsoon (Shi, 2002) and lack of glacial advances between the northern and southern climate regimes. Increasing aridity may have altered the system at ~30ka, leading to more abbreviated glaciations of uncertain precipitation source. Garhwal Sample Calculated Concentrations (atoms/g) Altitude Shielding Quartz 10Be* Sample ID Latitude Longitude Depth Be Carrier (g) 10Be/9Be (x10-15) Error % (m asl) Factor (g) (105 atoms/g) And Ages (yrs) with 1σ Uncertainties (cm)

5-30a 38.62089 74.99055 3244 0.999 0-5 31.06 0.4108 1865.04 ± 175.60 9 22.31 ± 2.01 5-30b 38.62089 74.99055 3244 0.999 17-23 25.15 0.4079 1516.87 ± 46.16 3 22.25 ± 0.67 In Situ 10Be: 2.14E+06 ± 5.75E+05 10

T2 5-30d 38.62089 74.99055 3244 0.999 82-88 5.40 0.401 125.22 ± 8.75 6 8.41 ± 0.50 Inherited Be: 3.78E+05 5-30e 38.62089 74.99055 3244 0.999 127-133 3.90 0.398 93.92 ± 10.03 10 8.67 ± 0.87 Age: 43,787 ± 12,032 Bulunko 5-30f 38.62089 74.99055 3244 0.999 192-198 28.29 0.4115 467.04 ± 20.94 4 6.15 ± 0.25

6-004a 38.98793 74.53998 3548 0.999 2 1.47 0.4013 152.74 ± 222.20 142 37.72 ± 53.56 In Situ 10Be: 1.74E+06 ± n/a 10 T1 6-004d 38.98793 74.53998 3548 0.999 150 3.60 0.4097 28.19 ± 2.97 8 2.90 ± 0.23 Inherited Be: 7.50E+04 Wulu 6-004e 38.98793 74.53998 3548 0.999 200 24.05 0.4012 125.84 ± 8.82 7 1.90 ± 0.13 Age: 26,500 ± n/a

6-008a 38.98170 74.52798 3515 0.999 2 4.05 0.3994 1036.74 ± 396.50 38 92.49 ± 35.1 6-008b 38.98170 74.52798 3515 0.999 50 10.42 0.4169 121.04 ± 7.06 5 4.38 ± 0.22 In Situ 10Be: 6.35E+05 ± 4.54E+05 10 73 T3 T3 6-008c 38.98170 74.52798 3515 0.999 100 13.14 0.4046 108.04 ± 5.55 5 3.01 ± 0.15 Inherited Be: 1.00E+05 Wulu 6-008d 38.98170 74.52798 3515 0.999 150 11.99 0.404 98.54 ± 6.81 7 3.00 ± 0.21 Age: 12,030 ± 8,640 6-008e 38.98170 74.52798 3515 0.999 200 13.93 0.3996 48.99 ± 4.25 7 1.27 ± 0.09

6-006a 38.98693 74.53902 3534 0.999 2 6.70 0.393 151.52 ± 10.42 6 8.04 ± 0.48 6-006b 38.98693 74.53902 3534 0.999 50 13.49 0.409 170.00 ± 8.34 5 4.66 ± 0.23 In Situ 10Be: 8.26E+05 ± 1.36E+05 10 T4 6-006c 38.98693 74.53902 3534 0.999 100 13.10 0.392 116.82 ± 6.62 5 3.16 ± 0.16 Inherited Be: 4.03E+04 Wulu 6-006d 38.98693 74.53902 3534 0.999 150 9.56 0.4019 33.16 ± 5.00 12 1.26 ± 0.15 Age: 15,391 ± 2,617 6-006e 38.98693 74.53902 3534 0.999 200 4.84 0.3977 13.64 ± 2.07 10 1.01 ± 0.10

6-018a 39.13698 74.53915 3740 0.999 2 3.75 0.4121 6.04 ± 1.90 18 0.60 ± 0.11 6-018b 39.13698 74.53915 3740 0.999 50 1.2 0.4147 -1.59 ± 0.94 35 -0.50 ± 0.17 In Situ 10Be: 3.78E+04 ± 1.48E+05 T4 6-018c 39.13698 74.53915 3740 0.999 100 4.04 0.4034 2.25 ± 1.33 15 0.20 ± 0.03 10 2.28E+04

Zouba Inherited Be: 6-018d 39.13698 74.53915 3740 0.999 150 10.37 0.419 9.11 ± 2.71 17 0.33 ± 0.06 Age: 637 ± 2,486

* Minus mean blank concentration Bold text: discarded from age calculation

Table 1: CRN sample data, including AMS results and age calculations with 1σ uncertainties.