Research Collection

Doctoral Thesis

Reconstructing paleoclimate and landscape history in Antarctica and Tibet with cosmogenic nuclides

Author(s): Oberholzer, Peter

Publication Date: 2004

Permanent Link: https://doi.org/10.3929/ethz-a-004830139

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ETH Library DISS ETH NO. 15472

RECONSTRUCTING PALEOCLIMATE

AND LANDSCAPE HISTORY

IN ANTARCTICA AND TIBET

WITH COSMOGENIC NUCLIDES

A dissertation submitted to the

SWISS FEDERAL INSTITUTE OF TECHNOLOGY ZURICH

For the degree of

Doctor of Sciences

presented by

PETER OBERHOLZER

dipl. Natw. ETH

born 14. 08.1969

citizen of

Goldingen SG and Meilen ZH

accepted on the recommendation of

Prof. Dr. Rainer Wieler, examiner

Prof. Dr. Christian Schliichter, co-examiner

Prof. Dr. Carlo Baroni, co-examiner

Dr. Jörg M. Schäfer, co-examiner

Prof. Dr. Philip Allen, co-examiner 11 Contents

1 Introduction 1 1.1 Climate and landscape evolution 2 1.2 Archives of the paleoclimate 2

1.3 Antarctica and Tibet - key areas of global climate 3

1.4 Surface exposure dating 3

2 Method 7 2.1 A brief history 8 2.2 Principles of surface exposure dating 8 2.3 Sampling 14 2.4 Sample preparation 17

2.5 Noble gas extraction and measurements 18

2.6 Non-cosmogenic noble gas fractions 20

3 Antarctica 27 3.1 Introduction 28 3.2 Relict ice in Beacon Valley 32 3.3 Vernier Valley moraines 42 3.4 Mount Keinath 51 3.5 Erosional surfaces in Northern Victoria Land 71

4 Tibet 91 4.1 Introduction 92 4.2 Tibet: The MIS 2 advance 98

4.3 Tibet: The MIS 3 denudation event 108

5 Europe 127 5.1 Lower Saxony 128 5.2 The landslide of Kofels 134

6 Conclusions 137 6.1 Conclusions 138 6.2 Outlook 139

iii Contents

A Appendix 155 A.l Additional samples 156 A.2 Sample weights 158 A.3 Sample codes of Mt. Keinath and Black Ridge samples 160

IV Summary

The main goal of this thesis was to reconstruct parts of the glacial history in two regions that are outstanding on the Earth regarding their climate, relief and position: Antarctica and Tibet. The approach chosen to achieve this was to date distinct landscape elements of glacial origin and make conclusions about the climate events related to their formation. With surface exposure dating using cosmogenic nuclides a technique was applied that has proven to be highly suitable for dating glacial features in cold and arid regions.

Antarctica

In Antarctica, glacial features were studied in the Dry Valleys of McMurdo and in the Terra Nova Bay area of Northern Victoria Land.

The ice of Ferrar Glacier, one of the large outlet glaciers of the East Antarctic Ice Sheet (EAIS), reached its maximum position of 125 m above the present level in Vernier Valley 3.4 Ma BP. It retreated continuously, with stagnation phases or re-advances at 1 Ma BP and 440 ka BP. These minimal local variations of the ice surface are extrapo¬ lated towards nearby Taylor Dome, indicating minimal volume fluctuation of the EAIS since the Pliocene.

The floor of Beacon Valley is covered by a till layer that covers remnants of ancient ice. Erratics from the surface yielded an exposure age of 2.6 Myr. Since then, the valley bottom was not reworked. This is a minimum age for the buried ice. Samples excavated from the ice shed light on the evolution of the ice-level in Beacon Valley. The compar¬ ison of the measured nuclide concentration with the shielding depth of buried samples indicates that the ice must have been substantially thicker at the time of deposition. Its thinning occurred continuously by sublimation and not as a collapse. In order to test the validity of paleoclimatic findings from the Dry Valleys for other regions of Antarctica, glaciogenic surfaces in Northern Victoria Land were investigated. Exposure ages from near the coast of Terra Nova Bay region imply that ice receded from the mountains near the coast more than 5 Ma ago. Subsequent fluctuations of the ice volume in the valleys have occurred, but were of limited extent.

A very similar picture emerges from datings of ancient dolerite surfaces from nunataks on the inner side of the Transantarctic Mountains. These prove that the volume varia¬ tions of the EAIS were very small since 3.3 Myr.

The results from the Dry Valleys and Northern Victoria Land are consistent with each other. They strongly indicates that the Antarctic climate that was constantly cold since at least the late Pliocene. This stability was characteristic not only for Southern Victoria Land, but also further north.

v Tibet

The second main focus of this work is the Tibetan Himalaya. It is not clear what the forcing climate mechanism of Tibetan glaciers advances is. In order to assess the relative influence of the Indian Monsoon and the North Atlantic climate on the glaciation of Tibet, the monsoonal area of Nyalam is investigated. A sequence of large and old- looking moraines was deposited more than 180 ka ago. On the same moraines, a cluster of apparently younger erratics indicates that strong denudation and moraine degradation took place between 48 and 65 ka BP, a phase of intensified Indian Monsoon.

A second Nyalam moraine set testifies for an influence of North Atlantic climate signals on Tibetan glaciations. During the global glacial maximum in MIS 2, glaciers in Nyalam advanced to a limited extent. A second advance was dated to 10.6 ± 0.4 ka BP, which is consistent with the Younger Dryas cold interval. The limited amplitude of these glaciations is explained with three arguments: a reduced monsoon activity during cold phases, the large distance to the moisture source of the oceans and with high-altitude effects.

Both moraine sequences hence indicate that the monsoon is of minor importance for triggering glacial advances in the Tibetan Himalaya. Instead, it seems very likely that Tibetan glaciers advanced in phase with the ice sheets of the Northern hemisphere during Late Pleistocene.

Surface exposure dating

With a large set of samples from Antarctica it is demonstrated that the degassing pattern of noble gases from Pyroxene in stepwise heating is strongly dependent on the elemental composition of the mineral. It is shown that - unlike previously assumed - cosmogenic neon can be released in substantial amounts below 850 °C. The percentage is inversely proportional to the Mg content. Furthermore, it is likely that cosmogenic helium is not necessarily retained quantitatively in pyroxene. In order to separate cosmogenic from non-cosmogenic noble gas fractions and obtain a reliable helium measurement, the type of pyroxene that is analysed has to be determined and accurately separated from others occurring in the sample.

In a pilot study using shielded samples from erratic boulders, the utility of such control samples to assess pre-exposure irradiation was tested. Together with a simple irradiation model, shielded samples allow to detect exposition prior to the deposition of a boulder and to roughly quantify it.

Zusammenfassung

Das Hauptziel der vorliegenden Arbeit war es, einen Teil der Gletschergeschichte der Antarktis und Tibets zu rekonstruieren. Diese beiden Gegenden sind einzigartig, was

vi ihr Klima, ihr Relief und ihre Position auf der Erde betrifft. Um dieses Ziel zu er¬ reichen, wurden Landschaftsformen datiert, welche deutlich glazialer Herkunft sind. Aus ihrem Alter wird auf das klimatische Ereignis geschlossen, das zum entsprechen¬ den Gletschervorstoss geführt hat. Die verwendete Datierungstechnik, Oberflächen- Altersbestimmung mit kosmogenen Nukliden, hat sich als zuverlässiges Werkzeug zur Datierung von Gletscherrückzügen in kalten und ariden Gebieten bewährt.

Antarktis

In der Antarktis wurden glazigene Oberflächen in den Trockentälern von McMurdo und in der Region der Terra Nova-Bucht in Nord-Viktorialand untersucht.

Das Eis des Ferrar-Gletschers, einem der grossen Ausfluss-Gletscher des ostantark¬ tischen Eisschildes (EAIS), erreichte seinen Höchststand im Vernier Valley 3.4Ma vor heute. Es reichte 125 m höher als heute. Sein Rückzug war kontinuierlich und nur von zwei Stagnationsphasen oder erneuten Vorstössen vor 1 Ma und vor 440'000 Jahren unterbrochen. Diese geringen Schwankungen der Eisoberfläche lassen sich extrapolieren auf den Taylor-Dom, was minimale Volumenänderungen des ostantarktischen Eisschildes anzeigt. Der Talboden von Beacon Valley wird gebildet von einem Moränenschleier, unter dem Überreste von altem Eis begraben sind. Blöcke aus dieser Moränenschicht ergaben ein Mindestalter von 2.6 Myr. Seither wurde der Talboden nicht mehr umgestaltet. Gesteinsproben aus dem begrabenen Eis erlauben Rückschlüsse auf die Geschichte des Eisstandes im Beacon Valley. Vergleicht man die gemessenen Nuklid-Konzentrationen mit der momentanen Abschirmung durch die Moräne, so zeigt sich, dass das Eis in der Vergangenheit deutlich dicker gewesen sein muss. Es is kontinuierlich dünner geworden, also nicht durch ein kollaps-ähnliches Ereignis, sondern durch stetige Sublimation. Um die Gültigkeit von paläoklimatischen Erkenntnissen aus den Dry Valleys für andere Regionen der Antarktis zu testen, wurden gletscherüberprägte Oberflächen in Nord- Viktorialand untersucht. Expositionsalter aus küstennahen Gebieten im Bereich der Terra Nova-Bucht legen nahe, dass die Berge dort seit mehr als 5 Myr eisfrei sind. Es gab zwar spätere Schwankungen der Gletscher, aber sie waren von beschränktem Ausmass. Ein ganz ähnliches Bild ergibt sich aus Datierungen von Dolerit-Oberflächen auf der Innenseite des transantarktischen Gebirges. Diese belegen, dass die Volumenänderungen des ostantarktischen Eisschildes seit mindestens 3.3 Myr sehr klein waren. Die Ergebnisse aus diesen zwei Gebieten sind konsistent miteinander. Sie belegen, dass da Klima der Antarktis seit dem oberen Pliozän permanent kalt war. Dies war nicht nur in Süd-Viktorialand der Fall, sondern auch nördlich davon.

Tibet

Das zweite Hauptaugenmerk dieser Arbeit lag auf dem tibetischen Himalaya. Es ist nicht ganz klar, welcher Klima-Mechanismus die Gletscher in Tibet wachsen lässt. Um die

vu Anteile des indischen Monsuns einerseits und des nordatlantischen Klimas andererseits an der Bestimmung des Klimaverlaufs abzuschätzen, wurde das vom Monsun geprägte Gebiet um Nyalam untersucht. Eine Sequenz von grossen, verwittert erscheinenden Moränen wurde vor mehr als 180'000 Jahren abgelagert. Die scheinbaren Alter von vier Blöcken derselben Moränen liegen eng beieinander zwischen 48'000 und 65'000 Jahren. Dies wird als Hinweis interpretiert, dass der in jener Zeit verstärkte Monsun durch vermehrte Niederschläge die Moränen stark abgetragen hat. Weitere Moränen aus derselben Gegend belegen, dass Klimasignale aus der Region des Nordatlantiks auch in Tibet wirksam sind. Während dem globalen Gletscher-Maximum der Sauerstoff-Isotopenstufe 2 stiessen die Gletscher in Nyalam vor, wenn auch in begren¬ ztem Mass. Ein zweiter kleinerer Vorstoss vor 10'600±400 Jahren stimmt zeitlich mit der Kaltzeit der Jüngeren Dryas überein. Das begrenzte Ausmass dieser beiden Vorstösse scheint drei Gründe zu haben: die reduzierte Stärke des Indischen Monsuns während dieser beiden Kaltphasen, die grossen Distanz zur Feuchtigkeitsquelle der Ozeane und Hochgebirgs-Effekte. Die Alter beider Moränensequenzen weisen übereinstimmend darauf hin, dass der Monsun keine bestimmende Kraft in der Auslösung von Vergletscherungen im tibetischen Himalaya ist. Stattdessen scheinen sich die Gletscher im späten Pleistozän zeitgleich mit den Eisschilden der Nordhemisphäre ausgedehnt zu haben.

Oberflächen-Datierung

Mit einem umfangreichen Datensatz aus der Antarktis wird gezeigt, dass das Entga¬ sungsmuster von Edelgasen aus Pyroxenen während des schrittweisen Heizens stark von der Zusammensetzung des Minerals abhängt. Anders als bisher angenommen können beträchtliche Mengen an kosmogenem Neon schon unterhalb von 850 °C freigesetzt wer¬ den. Der Anteil ist proportional zum Mg-Gehalt. Weiter ist es wahrscheinlich, dass kosmogenes Helium nicht immer vollständig in Pyroxenen zurückgehalten wird. Um kosmogene und nicht-kosmogene Neon-Anteile zu trennen und eine verlässliche Helium- Menge zu erhalten, sollte das Pyroxen-Mineral, das analysiert wird, vor der Messung genau bestimmt und separiert werden. In einer Pilotstudie mit abgeschirmten Proben von Erratikern wurde die Nützlichkeit solcher Kontroll-Proben zur Identifizierung von Vorbestrahlung getestet. Werden sie mit Modellrechnungen der kosmischen Einstrahlung an der genauen Probenstelle kombiniert, so kann Vorbestrahlung festgestellt und ihre Dauer abgeschätzt werden.

vni 1 Introduction

î 1 Introduction

1.1 Climate and landscape evolution

The surface of our planet with the landscapes we are cultivating and shaping at high rates today has evolved over long periods, and is still evolving, under the influence of strong external forces: rivers, winds, glaciers and gravity, mostly in an interplay of two ore more of these. Most of these forces and the extent to which they shape landscapes, as well as the rates at which they do this, are crucially depending on the climate. This is the reason why the evolution of climate and landscape are so closely interlinked. To learn about the history of one is often achieved best by looking closely at the history of the other.

Understanding the history of a landscape not only involves identifying the main actors and the most important events - to a considerable part, it consist of the time factor: find out when things happened, at what rates and over which length of time. Therefore, the crucial task in reconstructing the history of climate is to assign an age to the elements of a landscape and/or the processes which created them.

1.2 Archives of the paleoclimate

Our present knowledge about the characteristics of the climate in the past and their variations is mainly based on the archives of the deep oceans (e. g. Imbrie et al. (1984)) and ice core records from northern (e.g. Dansgaard et al. (1993)) and southern (e.g. Petit et al. (1999)) polar ice sheets on one side, and lacustrine sediments (e.g. Björck et al. (1996)), pollen profiles and tree rings from the continents on the other. As the youngest periods of the Earth's history, the ice ages, were characterised by large ice sheets and ice caps on the continents, the terrestrial archive would be excellently suited for the reconstruction of how the climate varied in these last few million years. Glaciers are extremely sensitive climate indicators responding to changes in both temperature and precipitation. Therefore they record changes in ocean and atmospheric circulation patterns, as well as temperature variations. This information is stored in the sedimentary and erosional features that the glaciers and ice sheets have left behind: Valley troughs and fillings, and moraines in mountain belts, fluvioglacial gravel bodies, erratics, tors and till sheets in their forelands and along the former margins of ice sheets and ice caps. By mapping their distribution, the amplitude of past cold climate excursions can be reconstructed. If their age can be determined, the time frame of these climate variations may be established, too. Comparing timing and amplitude of past glacial advances in different latitudes sheds light on which areas have a driving influence on the Earth's climate system. However, the terrestrial record is much less complete than marine archives are, although climate variations certainly did have an immediate impact on the continents through advancing or retreating glaciers and ice sheets. This lack of evidence ("more gaps than records" on the continents) is due to the constant activity of weathering as well as episodic erosion,

2 1.3 Antarctica and Tibet - key areas of global climate

which wipe out most of the traces of glacial advances. Glacial landscape features are best preserved, therefore, where the atmospheric processes that lead to weathering are least intense - in the cold and arid regions of the Earth, i. e. in high mountains and near the poles. This is why Antarctica and Tibet were chosen as the main areas of research for this dissertation.

1.3 Antarctica and Tibet - key areas of global climate

Both Tibet and Antarctica offer a wealth of well-conserved witnesses for the glacial activ¬ ity from large time ranges, because both were severely affected by ice coverage whenever the climate of our planet deteriorated in the past. This makes them invaluable archives for the study of glacial landscape evolution and the reconstruction of paleoclimate. Both areas also are regarded as potential agents in causing climate variations. This is due to their size and their potential to store large ice masses. For example, Antarctica, by locking large volumes of ice, defines the global sea level and with it the size of the landmasses and continental shelves; changes of the sea level therefore alter circulation patterns of both the oceans and the atmosphere. The Tibetan Plateau is located at the latitude of the westerly jet stream. Because it acts as a huge heat source in summer, a snow or ice cover of the area, reducing the heat supply, would influence the circulation of the stratosphere and be able to shift the jet stream to a more northern latitude.

1.4 Surface exposure dating

A problem inherent to paleoclimate studies in cold and dry regions is the scarcity of organic material that can be used for radiocarbon dating. The resulting lack of absolute dating has hindered the reconstruction of climate history through establishing of glacial chronologies in regions like Tibet and Antarctica. The advent of the method of surface exposure dating (SED) with cosmogenic nuclides provided a means to directly and ab¬ solutely date land forms of ages between a few thousand up to 10 million years without the requirement of organic material. Particularly in Antarctica, this new technique soon led to the first dating studies (Ivy-Ochs et al., 1995; Brook et al., 1995b,a; Bruno et al., 1997). It has become an indispensable tool as well for the paleoglaciological problems on the Tibetan Plateau (Schäfer et al., 2002; Tschudi et al., 2003) and in the surrounding mountains (Owen et al., 2001; Zech et al., 2003). Another advantage of SED is that it directly dates an event, while other techniques determine the age of a related object or process, e. g. the onset of vegetation in the forefield of a vanished glacier or the damming of a lake by a moraine (radiocarbon, pollen stratigraphy, dendrochronology). For its suitability for cold regions and for its ability to provide direct absolute dat¬ ing, SED was the technique of choice for the problems addressed in this work. The

3 1 Introduction

application of SED has seen something like an inflation in recent years. Known only to physicists and a few glacial geomorphologists a decade ago, the method is used in countless contributions at every Quaternary geology or geomorphology congress. Surface exposure dating - when applied properly - is a successful alliance of physics and Earth sciences. A meaningful and accurate exposure age is the result of interdisciplinary work and requires both field expertise to collect an appropriate sample, and physical and tech¬ nical knowhow to carry out a reliable measurement. For these reasons, it was essential for this thesis project to involve experienced geologists into the collection of samples, as well as physicists into the analysis and interpretation of their isotopic content.

Organisation of this thesis

The indispensable partnership of physics and Earth sciences for exposure dating is re¬ flected in the collaboration of universities, departments and institutes that led to this thesis project and its predecessors since 1994. The technical and physical base of knowl¬ edge for the noble gas isotope analysis in this thesis were provided by Rainer Wieler and Heiri Baur from the institute of Isotope Geology and Mineral Resources at ETH Zürich. For those of our projects that involved radionuclides, I was strongly supported by Susan Ivy-Ochs and Peter Kubik from the Institute of Particle Physics at ETH Zürich. The geological problems and approaches to their solution as well as many rock samples and guidance on sampling excursions came from Christian Schlüchter from the Institute of Geology at the University of Bern, and from Carlo Baroni from the University of Pisa, Italy. Jörg Schäfer from the Lamont-Doherty Earth Observatory in Palisades, USA, acted as the linking connection between the field and laboratory sides. Finally, Philip Allen from the Institute of Geology at ETH Zürich supplied the viewpoint of the surface dynamics.

This work was substantially supported by the Swiss National Science Foundation (SNF) Grants No. 2000-053942.98/1 and 2000-065153.01/1.

Outline

This thesis is structered as follows:

I will give a short introduction to the basic principles of exposure dating for those who are not familiar with it in chapter 2. The details of the analyses carried out for this thesis will be described, and the results of methodological experiments are discussed.

Chapter 3.1.3 presents four studies from the Antarctic continent. Two of them come from the Dry Valleys of McMurdo, the "classic" spot of glacial geological research in

Antarctica. The other two extend the Antarctic data base of exposure ages further north to the Terra Nova Bay region and compare local glacial records with those from the south.

4 1.4 Surface exposure dating

In chapter 4, I report the data from our 2001 expedition to the Tibetan Himalaya. Evidence was collected that north Atlantic climate signals of the latest Pleistocene are recorded in Tibet. Furthermore, an alternative explanation for the "local Tibetan Last Glacial Maximum (LGM)" during MIS 3 is suggested. Finally, two pilot studies are reported in chapter 5. The first is about an attempt to date Pre-LGM glaciations in Northern Germany. The second is the attempt to calibrate the 21Ne production rate at a site of intermediate latitude and altitude. The final part consists of general conclusions and an outlook to possible future research activities.

5 1 Introduction

6 2 Surface exposure dating with cosmogenic nuclides

7 2 Method

2.1 A brief history

Surface exposure dating (SED) with in-situ produced terrestrial cosmogenic nuclides (TCN) is in broad application for less than two decades, but the fact that nuclides of some isotopes are produced by the interaction of cosmic-ray particles with rocks has been known much longer. Cosmogenic nuclides were first discovered in 1952 in iron meteorites (Paneth et al., 1952). Davis and Schaeffer (1955) were the first to measure cosmogenic nuclides in terrestrial material. Craig et al. (1979) discovered excess 3He and 21Ne in terrestrial metals. Besides that, not much research was done in the field of terrestrial cosmogenic nuclides for thirty years as mass spectrometry was not yet ready to reliably measure the low concentrations occurring in terrestrial samples. A large number of pioneer studies were published in 1986. In-situ produced cosmo¬ genic radionuclides 10Be, 26A1 in quartz (Nishiizumi et al. (1986); Klein et al. (1986)) and 36C1 in volcanic rocks (Phillips et al., 1986) were reported. Cosmogenic noble gases 3He and 21Ne were detected in volcanic olivine and pyroxene (Kurz, 1986a,b; Craig and Poreda, 1986; Marti and Craig, 1987)). This initial surge of studies was followed by in¬ tensive research on theoretical aspects, along with numerous applications and substantial improvements in the field of mass spectrometry. Today, attempts are made to use other cosmogenic nuclides for dating and other minerals for SED, e.g. garnet (Dunai and Roselieb, 1996), sanidine (Kober et al., 2004) or fluorine apatite (Farley et al., 2000). Nuclide production rates and their scaling to latitude and altitude are still under discussion. Production rate issues will be further discussed on paragraphs 2.2.2 and 2.2.3 of this chapter. Today, SED with TCN is a valuable dating tool that is widely accepted by the Earth science community, and the range of problems that can be addressed with it has consid¬ erably widened. The method's significance is expressed by the large and still increasing numbers of SED papers published in journals or presented at geochemistry, geology or geomorphology conferences.

2.2 Principles of surface exposure dating

Only a brief summary of the physical principles of SED is given here, as these issues are excellently covered in a large number of application studies and review papers, e. g.Niedermann (2002); Gosse and Phillips (2001); Cerling and Craig (1994a); Bierman (1994); or Kurz and Brook (1994). The following explanations address the production mechanisms of cosmogenic noble gases, but some references to radionuclides are made. The lengthy expression "Surface exposure dating using in-situ produced terrestrial cosmogenic noble gases" for once is a quite helpful description of the most important aspects of this method: surface SED samples are collected from the very surface of a rock. The production rate decreases exponentially with the rock depth below the surface. Therefore, only the

8 2.2 Principles of surface exposure dating

uppermost 2 cm are usually used for analysis. exposure What is determined is the duration for which the rock surface was exposed to cosmic radiation, ideally without cover of any kind (snow, ice, sediment, soil, rock or thick vegetation). dating Assigning an age to a surface is not the only goal that can be achieved by mea¬ suring the concentration of cosmogenic nuclides. Erosion rates can be constrained too (see below). in-situ produced The nuclides that are used for SED are produced in the sampled rock surface itself. This in contrast to other techniques such as radiocarbon dating which uses the cosmogenic nuclide 14C produced in the upper atmosphere and incorporated into organisms. terrestrial SED is a valuable tool for geomorphological problems on our planet. How¬

ever, the method was first used to determine exposure ages of meteorites and still is an indispensable tool in cosmochemistry. cosmogenic The radiation that leads to the production of TCN originate in the cosmos. The cosmic radiation has a galactic and a solar component. The first originates outside our solar system and its particles are of much higher energies than those of the second. For SED, the solar component is of minimal importance. nuclides The nuclides used for SED are rare isotopes of various elements. The most

commonly used ones are the noble gases 3He and 21Ne and the radionuclides 10Be, 26Al and 36C1. The use of cosmogenic 14C (from in-situ production), 15N, 38Ar and 53Mn is tested.

Some of these points are described in more detail in the following.

2.2.1 Nuclide production and interfering processes

Protons of the cosmic radiation collide with the particles of the upper atmosphere. The main product of these spallation reactions are secondary neutrons. As the secondary neu¬ trons travel towards the Earth, they induce more spallation reactions with the nuclides of the atmosphere and finally with the matter on the Earth's surface. The cosmic-ray flux is strongly diminished by these interactions. Among the products of the spallation reactions are the two helium and all three neon isotopes. Table 2.1 gives an overview of the most important nuclides used for exposure dating. A very useful comparison of these nuclides is given in Table 1 of Gosse and Phillips (2001). Cosmogenic nuclides are not only produced by spallation, but also by the capture of stopped negative muons. At the surface, production by spallation is dominant, but at a depth of 3 to 4 metres, muon production becomes more important (Cerling and

9 2 Method

Table 2.1: The cosmogenic isotopes used for dating and their main target elements in the lithosphère (from Cerling and Craig (1994a)).

Isotope half-life target element [yrs]

3He stable 0, Si, AI, Mg etc. 10Be 1.5xl06 0, Si, Al 21Ne stable Mg, Al, Si, Na 26 Al OJlxlO6 Al, Si 36C1 0.3 xlO6 Cl, K, Ca

Craig, 1994a). As muon production is negligible for noble gases at the surface (Brown et ah, 1995), it is not discussed any further here. It is an important issue in radionuclide dating, though. The produced cosmogenic nuclides accumulate over time in a rock surface. The nuclide concentration at the very surface is a direct measure of the duration of exposure:

A^(l-e-«-) (2,)

In the case of radionuclides, the accumulation is opposed by radioactive decay. The corresponding equation is

P f. „.„mwA „ „ " G -T iVQO (2_2) A + ^± j A

N [atg"1] nuclide concentration P [atg_1yr~-1] nuclide production rate T H duration of exposure P [g cm-3] rock density t [cmyr-1] steady state erosion rate A [g cm-2] cosmic ray attenuation length A [a"1] decay constant of the radionuclide

Cosmogenic isotope production takes place in all minerals, but not all minerals retain the produced nuclides. By comparing 3He and 21Ne concentrations found in quartz grains, it was shown that 3He is lost by diffusion (Graf et ah, 1991), while 21Ne is retained quantitatively (Niedermann et ah, 1993). In the mafic minerals pyroxene and olivine, both isotopes are thought to be retained (Bruno et al. (1997); Schäfer et al. (1999)), although recent studies suggest that this may not be true for all pyroxenes (see section 3.5). The radionuclide 10Be can be used only in quartz because in pyroxene, the

10 2.2 Principles of surface exposure dating

meteoric component can not be quantitatively removed (Ivy-Ochs et ah, 1998). This makes cosmogenic noble gases a very valuable dating tool for quartz-free lithologies, such as those of mafic volcanic rocks. For exposure times of several millions of years, they are even the only isotopes available for exposure dating, as the time range of 36C1 is limited by radioactive decay. If a mineral is retentive for a nuclide, the exposure age assuming constant production rate and zero erosion can be calculated simply as:

N T=- (2.3)

Assuming zero erosion and calculating T is one limiting case of eq. 2.1. The other limiting case, assuming infinite exposure and calculating e, is discussed below.

The assumption of a constant production rate in eq. 2.3 is never entirely fulfilled. Several factors can reduce the production rate temporarily or permanently. While some can be corrected for, others can be partially controlled by appropriate sampling (see section 2.3):

Temporary shielding by snow, sediment, or vegetation. Corrections can be made using estimates of the nature and duration of a shielding period. A good example is Ku- bik et al. (1998). Shielding for a substantial period can be detected if two nuclides

of different half-lives or a stable and a radioactive nuclide are measured. Because

the nuclides that were built up before the shielding period decay at different rates (or do not decay at all as the stable nuclides) during the shielding, they yield different ages.

Shielding by surrounding topography. The equations above are only valid for an ob¬ ject irradiated on a flat surface. An elevated horizon reduces the incoming cosmic- ray flux. This can be corrected for if the horizon angle is measured at the sampling site.

Overturning of a boulder. If a moraine boulder is moved after its initial deposition

on the moraine, e. g. it overturns, a fresh surface can become exposed to cosmic

radiation. Then, the duration of exposure determined for the "new" top surface

does not represent the age of the landform. For this reason particularly, it is essential to take more than one sample from a landform. Putkonen and Swanson (2003) precisely suggest the number of boulders to be taken from a moraine as a function of the expected age and its height. However, the number of samples is usually limited by laboratory capacity and financial boundary conditions as well

as the time available.

Erosion. At the same time as cosmogenic nuclides accumulate in a rock surface, they are permanently removed by erosion. This can happen at a steady-state rate, or episodic. Mathematical models exist for both cases (e.g. Lai (1991); Nishiizumi

11 2 Method

et al. (1986); Gosse and Phillips (2001)). An ideal case is the presence of glacial striae on the surface to be dated, because it proves that erosion must have been

very limited, else, the striae would not have been preserved. If the exposure time T is either known from an independent dating or can be reasonably estimated, a steady-state erosion rate can be calculated from eq. 2.1. A maximum erosion rate is obtained by assuming the exposure time T as infinite. This is the second limiting case of eq. 2.1.

Alternatively, a long-term erosion rate can often be estimated graphically from the so-called erosion island plot (Lai, 1991) if two nuclides with different half-lives are measured in the same sample. Such a graph shows the ratio of two nuclides versus the concentration of one of them and contains curves representing different erosion rates. This approach is used for the beryllium and neon data from Mount Keinath in section 3.4).

All these factors potentially inhibit nuclide accumulation in the sampled surface. If they are not corrected for, the production rate is overestimated and leads to too young exposure ages. As we can never be sure which of them affected the surface, and we can mostly make assumptions only, exposure ages generally have to be regarded as minimum ages. This fact always has to be kept in mind when interpreting surface exposure ages in a geomorphological context. The sensitivity of an exposure age to erosion is illustrated in section 3.4 for samples from Antarctica.

So far, only processes that underestimate the exposure age were discussed. Three processes make the measured nuclide concentration higher than what corresponds to continuous exposure at the present-day location:

Pre-exposure A surface can be irradiated prior to its deposition at the sampling site. For a moraine boulder this could take place in the rock wall above the glacier, on an older, reworked moraine, or during transportation on the glacier surface. Pre-exposure can be revealed by measuring two cosmogenic nuclides of different half life. If the period of pre-exposure was followed by a substantial period of burial, two isotopes yield differing ages due to different decay rates. Furthermore, a partially shielded sample collected along with an exposed sample can be used to identify pre-exposure (see section 2.6.3).

Uplift. When a surface is uplifted during its irradiation through tectonic processes, the production rate steadily increases. To use a constant production rate scaled to the present-day elevation of the sampling site would therefore lead to too young expo¬ sure ages. As uplift commonly occurs at low rates, its effect is mostly neglected. If uplift must be considered, i. e. in active mountain belts, an uplift rate U must

be independently determined or taken from literature. Using its exposure age as¬ suming no uplift, the elevation of the sample at the beginning of irradiation and the respective production rate are calculated. Through integration, an averaged

12 2.2 Principles of surface exposure dating

production rate is obtained. This process is repeated iteratively, and the exposure

age corrected for uplift is calculated.

However, uplift is a concern mainly for old samples, unless extraordinarily high uplift rates are involved.

Nucleogenic production A nucleogenic neon component (see section 2.6) that has a composition similar to atmospheric or cosmogenic may not be recognised as such.

In this case the real age is lower than the apparent. The importance of this problem

decreases with increasing exposure age of the sample.

2.2.2 Production rates I: Values

If nothing else is remarked, all production rate values given in this chapter refer to their sea level, high latitude value.

3He production rate (P3)

In this work, P3 in pyroxenes is regarded as independent from the composition of the target mineral (Eugster, 1988). For both pyroxene and olivine samples, I have used the value of 115 ± 4atg_1 yr-1. It was determined by Cerling and Craig (1994b) from olivine and pyroxene-bearing, 17'500 14C yrs old volcanic rocks and is widely used today. Licciardi et al. (1999) (116 ± 3atg-V_1) and Dunai and Wijbrans (2000) (118 ± 11 atg-1 yr-1, calculated with their own scaling fromalism) suggest similar values.

21Ne production rate (P21)

Cosmogenic 21Ne is produced by several reactions and target minerals. In pyroxene and Olivine, the three elements Al, Mg and Si are responsible for more than 99% of the nuclide production. Here, the the elemental production rates from Niedermann (2000) and Schäfer et al. (1999, 2000) were applied:

- P2i(Si) = 42.5at g"1 a"1

- P2i(Mg) = 185at g"1 a"1

- P2i(Al) = 51at g"1 a"1

The concentrations of the main target elements were determined for the Vernier Valley and Beacon Valley samples by ICP MS for Si and laser ablation ICP-OES for Mg and Al at the Institute of Inorganic Chemistry at ETH Zürich; for the Northern Victoria Land samples by atomic adsorption spectroscopy at the Centre de Recherches Pétrographiques et Géochimiques in Vandoeuvre-lès-Nancy, France. Quartz has a simple and invariable composition, and accordingly, P2i is independent of the mineral chemistry. It has been determined at an independently dated calibration

13 2 Method

site. Niedermann et al. (1994) analysed glacially polished granite of the Tioga period in the Sierra Nevada (USA) and calculated a P2i value of 21 at g_1 yr-1. After the age of the glacial stage was revised, this value was adapted to 19.0 ± 3.7atg_1yr_1 (Niedermann, 2000). I adopted the value of 20 at g_1 yr-1 obtained from the data of Niedermann et al. (1994) with the revised age of the Tioga stage, but scaled to sea level and high latitude with the scaling formalism of Lai (1991) instead of that of Dunai (2000) (see tablet in Niedermann (2000)).

2.2.3 Production rates II: Spatial scaling

The nuclide production rate at any location on the Earth's surface is a function of its geomagnetic latitude and its elevation. This is because two main agents modulate the incoming neutron flux: The Earth's magnetic field and its atmosphere. The magnetic field is easiest to penetrate for the primary protons at the poles, where the field lines are at a right angle with the Earth's surface . Production rates are therefore highest at the poles and lowest at the equator. Because the atmospheric shielding decreases with elevation, the neutron flux and thus the nuclide production rate increase. The sensitivity of production rates to elevation is much stronger, as illustrated in Fig. 2.1.

As a consequence of these relationships, production rates have to be adapted to the elevation and geomagnetic latitude of a sampling site. Various scaling algorithms are in use. For years, that by Lai (1991) was standard. Dunai (2000), Stone (2000), Desilets and Zreda (2001) and others introduced new formalisms. They use partially new ap¬ proaches, but also base them on new neutron flux data. Differences in production rate scaling make it difficult to directly compare exposure ages from the same region but different studies. However, the deviations are known, and if authors clearly declare how they scaled their production rates, exposure ages can be transformed from one system into an other.

The production rate is one of the major sources of uncertainty in SED. This is not only due to the scaling, but also to the sea-level, high-latitude value that is used. Efforts of research groups from all over the world to ameliorate the situation are concerted in the recently initiated, US and Europe-based CRONUS-Earth (Cosmic Ray-Originated NUClide Systematics on Earth's) project. Planned for at least the period from 2004 to 2007, the main goals of this project are to systematically measure neutron fluxes, expose artificial targets in latitude and altitude transects, calibrate production rates at independently dated sites (see section 5.2) and model geomagnetic and paleoclimatic effects upon nuclide production.

2.3 Sampling

As mentioned above, several problems inherent to SED can be circumvented by appro¬ priate sample strategies in the field. This applies for the large to intermediate scale

14 2.3 Sampling

1 40

vel * • n a.s.l - • I /

4 30 / /

é - - t / I ! / 20<3 g>2 • - — —

-

10

-

0 10 20 30 40 50 60 70 80 90 0 1000 2000 3000 4000 5000 6000 Latitude n altitude [m a.s. I.]

ure 2.1: Variation of the scaling factor for nuclide production rates with geo¬ graphic latitude at sea level and at 1500 m a. s.l. (left), and altitude at 45° (right). The variation with altitude is much stronger. Note the

different scales on the ordinates. Scaled with the formalism of Stone (2000).

15 2 Method

(from the climate zone to the geomorphological environment) as well as to the small scale (e. g. on the moraine and on the selected boulder). The landform to be dated must be selected according to the questions addressed and the nuclide to be measured. This implies that a rough age estimate should be made. For example, a Late Glacial moraine in a precipitation-rich climate is difficult to date with cosmogenic noble gases, as the duration of exposure would be rather short given the high erosion rate and the possible interferences of atmospheric and nucleogenic con¬ tribution. Radionuclides with their low non-cosmogenic background and low detection limit, and 3He with a high production rate are more promising in this situation. On the other hand, attempting to date a very old-looking rock in Antarctica, where the oldest surfaces of the Earth are found, with radionuclides can be disappointing because the concentration most probably will be in saturation. The pitfalls on the small scale are less obvious. The requirements to the "perfect" sample surface are so numerous that it simply does not exist. Nevertheless, some char¬ acteristics of a promising sample for dating glacial features shall be listed here:

- The rock contains a mineral that retains cosmogenic nuclides, and the latter can be can be quantitatively extracted: commonly quartz for neon and the radionuclides,

pyroxene or olivine for both helium and neon. 36C1 can be measured in whole rock samples.

- The surface looks old compared to others occurring on the same block or moraine,

e. g. it displays desert varnish or a weathering rind, both indicators of long expo¬

sure.

- At the same time the surface does not look heavily eroded. Glacial striae or tectonic harnish are a good indicator. If traces of strong erosion like pit holes or tafoni are visible, the sample should be taken at a spot between the holes, at a protruding surface. There are no signs that recently, chips or slabs of rock have spalled, or that the entire block is falling apart.

- An erratic is in a stable position: It rests on flat terrain. On a sloping surface, it is stably embedded into the soil. On a moraine, the block sits on the very top of the crest. This strongly indicates that it is in its initial position; if a block starts moving from the crest of a moraine, it most certainly goes all the way down to the valley.

- The sample is taken from the top of a large block or a highly elevated part of bedrock. This reduces the probability that the sample is shielded by snow or vegetation. Note that if there is no vegetation today, there might still have been

some earlier.

- The surface is horizontal. Else, the neutron flux is reduced. If the slope is mea¬

sured, this can be corrected for.

16 2.4 Sample preparation

- The surrounding country is open, i. e. free of large obstacles that shield the cosmic radiation like huge blocks, hill slopes or high trees. Shielding by distant mountains can be corrected for if the horizon angle is measured.

- The landform offers the opportunity to collect other samples. Three or more samples from one moraine are desirable (Putkonen and Swanson, 2003).

- The sample is taken from the center of a large block, not from its edge; else, the production rate is reduced because of neutrons escaping through the sides of the block (Masarik and Wieler, 2003)

Samples for this thesis were collected with a hammer and chisel. Where a block showed a projecting part at the top, it was attempted to detach it with a sledgehammer. In some rare cases in Antarctica, a fuel-powered hand driller was available that was used to loosen parts of the rock which were then hammered off. The amount of sample required depends on the abundance of the target mineral in the rock and the number of nuclides that can be expected. This in turn is a function of the exposure time and the production rate, i. e. mainly the elevation of the sampling site (see Fig. 2.1). For noble gas analysis, a few grams of pure mineral grains is sufficient. This is can be extracted from a few hundred grams of rock. For radionuclides, though, a few hundred grams of quartz may be required. For an old sample or one from high latitudes, less material has to be collected than for a young sample in mid-latitudes. A spring-balance is helpful to optimise the weight of the sample in the field. Gosse and Phillips (2001) give a useful instruction estimate the minimum sample weight as a function of the elevation and the expected age.

2.4 Sample preparation

The following is a description of the standard procedures of preparing mineral separates, as they were applied to most of the samples in this thesis. However, adaptations were made for some sample of difficult nature or for experiments. Rock samples were photographed in order to document the characteristics of their sur¬ face such as weathering features, the details of the geometry and the general appearance. A thin section was made of each sample.

2.4.1 Pyroxene

Rock slabs of the uppermost two to three centimetres were sawed off, crushed in a steel mortar and sieved in a Pb free sieve. The grain size fraction of 100 to 315 //m was used for further processing. The grains that came from the very surface of dolerite samples often showed a heavy reddish-brown stain coat which hindered the magnetic separation of pyroxene. In those

17 2 Method

cases, the coat was removed by washing the grains in 30% hydrochloric acid and deion- ized water at about 95 °C for 10 minutes before the separation. In this process, the grains usually lost their individual greenish-brown color along with the stain coat and all appeared uniformly light grey.

By magnetic separation in up to four steps with a Frantz LB-1 device, many, for some samples most, of the mineral phases other than pyroxene were removed. Pyroxene grains with impurities or a deep staining have magnetic properties different than pure pyroxene and were removed as well.

Pyroxenes were further enriched in the sample with heavy liquid separation in methy¬ lene iodine. Usually clinopyroxenes sink at a density of 3.20 g cm-3, orthopyroxenes at 3.28 to 3.4gcm~3 (Balco et ah, 2002). If necessary, pyroxenes or the disturbing mineral phases were handpicked in a final step to obtain a higher purity of the separates.

In order to release a part of the nucleogenic noble gases contained in fluid inclusions, the pyroxene grains were crushed to less than 100 //m in an agate mortar, washed in water, cleaned from very fine grained material with an ultrasonic bath and dried.

Finally, 20 to 50 mg of each sample were wrapped in small pieces of household alu¬ minium foil and loaded to the furnace of the extraction line.

2.4.2 Quartz

For the analysis of noble gases from quartz samples, aliquots of the separates prepared for radionuclide analysis were usually used. They were prepared following the procedure of Kohl and Nishiizumi (1992), modified after Ivy-Ochs (1996) at the Lamont-Doherty Earth Observatory's cosmogenic laboratory or at the Institute of Particle Physics at ETH. If radionuclides were not measured in a sample, quartz for noble gas analysis was separated with basically the same procedure as pyroxene. The Frantz apparatus was used in the diamagnetic mode, and the density fraction from 2.65 to 2.69 g cm-3 was used (Niedermann et al. (1993); Kohl and Nishiizumi (1992)) and crushed to 100 ßm in an agate mortar in order to reduce the nucleogenic neon contribution from fluid inclusions (see section 2.6.2). Wrapping and loading was done in the same manner as for pyroxene, see above.

2.5 Noble gas extraction and measurements

2.5.1 Extraction line

In the common procedure, noble gases were extracted from the mineral separates by stepwise heating in an electron impact furnace (compare paragraph 2.6.1). From some samples, an additional aliquot was crushed with an on-line crusher device (see paragraph 2.6.2).

18 2.5 Noble gas extraction and measurements

The extracted gas was then cleaned in an ultra high vacuum (UHV) extraction line. This line included three Zr-V-Fe getters, three non-evaporable SAES getters with Zr and Al as getter materials, and three charcoal cold traps. He and Ne were separated by freezing the Ne component to a cryogenic cold trap with a 0.5 //m high porosity filter of sintered stainless steel at 14 K before helium was admitted into the spectrometer.

2.5.2 Mass spectrometer

Noble gas concentrations were measured with a custom-made all-metal, statically op¬

° erated 90 sector field noble gas mass spectrometer ("Tom Dooley") at the noble gas laboratory of the Institute of Isotope Geology and Mineral Resources (IGMR) at ETH Zürich. Its sensitivity is essentially constant over the pressure range relevant for this work.

A particular feature of this spectrometer is its ion source. It is a small version of a Baur-Signer source (Baur, 1980), equipped with a compressor (Baur, 1999). The compressor works like an inverse molecular-drag pump. Magnetic fields levitate the rotor and drive it at 86'000 revolutions per minute . By concentrating the neutral sample gas in the ionization chamber, the compressor reduces the effective volume of the spectrometer. This enhances the sensitivity of the spectrometer by a factor of 120 for 3He, 150 for 4He and 200 for 21Ne compared to the same ion source without the compressor running (Baur, 1999). This increase in sensitivity allows much smaller gas amounts to be measured. As a direct consequence, the sample weight could be reduced from 150 to 300 mg to 20 to 40 mg.

Neon isotopes, the interfering ions and 3He were counted with an electron multiplier. 4He was detected on a Faraday cup. The measured count rates and currents, respectively, were calculated into gas volumes by measuring a helium-neon standard (produced at IGMR) with the precisely same procedure as a sample. The short-term variations in sensitivity were recorded with "fast" calibrations three to four times a day, for which a helium or a neon standard gas was let directly into the spectrometer and measured. Depending on the exact settings of the ion source and the filament condition, the sensi¬ tivities were in the order of 3 to 4xl015 Hz ccSTP"1 21Ne and 6 to 9xl014 Hz ccSTP"1 3He.

The measurements on mass 20 were corrected for interferences of H2180 (0.2 % of the total H20) and the doubly-charged ion 40Ar++ (0.045% of the total 40Ar), respectively. Typical Blank values are 105 at 3He and 107 at 21Ne.

Details about the extraction line and the spectrometer are reported in Beyerle et al. (2000)(referred to as spectrometer "MSI").

19 2 Method

2.6 Non-cosmogenic noble gas fractions

Noble gas concentrations in minerals are always a mixture of several gas components from different sources. As they differ in their isotopic composition, they can often be separated. For different elements, the various components have different importance.

These sources are:

Nucleogenic Produced by nuclear reactions that are induced by neutrons or a particles from the decay of U and Th (Wetherill, 1954); mainly 3He and 20_22Ne.

Radiogenic Direct products of the decay of 235>238U, 232Th and 147Sm; 4He.

Mantle Noble gases incorporated during rock formation. Helium of mantle origin has an isotopic ratio of up to 30xRa (the atmospheric 3He/4He ratio, see below).

Trapped Originating mainly from the atmosphere, but also from the mantle or the crust. Atmospheric ratios are (Eugster, 1988)

- 3He/4He=JRA=1.386xl0-6

- 21Ne/2ONe=2.959xl0-3

- 22Ne/20Ne=0.102

As nucleogenic and radiogenic contributions are produced related to nuclear decay, they are a function of the rock formation age and the concentrations of the respective mother elements.

In the case of neon, different contributions in a sample can be distinguished by plotting the isotope ratios in a three-isotope plot (Fig. 5.3). Data that plot on the atmospheric- cosmogenic mixing line and are not of atmospheric composition are suitable for age calculation. A nucleogenic contribution is indicated if the data plot away from that line.

If a sample contains a mixture of atmospheric and cosmogenic neon, the cosmogenic excess in 21Ne (21Necos) over air is calculated under the assumption that all 20Ne is atmospheric:

= - x 21Necos 21Netot [ ^- ] 20Netot

atm 21 20 at„ n nrr.n -., in-3 w = 2VNetot - 2.959 x 10"3 x MNetot (2.4)

In the following, different approaches to separating cosmogenic and non-cosmogenic noble gases will be described: stepwise heating, crushing and shielded samples. The best results are often acheived by combining two of them.

20 2.6 Non-cosmogenic noble gas fractions

i

c g

o 3 "O > . ./ o —

z ^

AIR? Nucleogenic 21Ne production (1sO(a,nfNe)

>• 21Ne/20Ne

Figure 2.2: Neon three isotope plot. Arrows indicate cosmogenic and nucleogenic production processes. After Niedermann et al. (1993).

21 2 Method

2.6.1 Stepwise heating

This approach is well-suited to separate cosmogenic from nucleogenic neon. It takes advantage of the fact that these contributions occupy different locations in the mineral grain. While cosmogenic neon sits in the crystal lattice, nucleogenic neon is contained in fluid inclusions (Niedermann et ah, 1993). When the mineral is heated, these two fractions leave their place at different temperatures and can be measured separately. An accurate degassing pattern was first established for quartz by Niedermann et al. (1993), who found that 21Necos is extracted below 600 °C, with a maximum at 320 °C. Schäfer et al. (2000) investigated the behaviour of cosmogenic noble gases in heated pyroxene. According to that study, 3Hecos is almost completely released between 400 and 900 °C, along with 50 % of the trapped neon. 21Necos is released below 1400 °C.

It seems that the degassing pattern from pyroxene is not of general validity, and it seems that not all pyroxenes are entirely retentive for cosmogenic helium. In section 3.5, the degassing pattern of pyroxene from the same geologic unit that Schäfer et al. (2000) inferred the pyroxene degassing pattern from, the Ferrar Dolerite Group in Antarctica is described. In some of those samples, up to 95 % of the 21Necos are released at 850 °C, which is not consistent with Schäfer et al. (2000). 3Hecos, in contrast, behaved as pre¬ dicted. A dependency of the neon degassing pattern on the Mg content of the sample is suggested.

2.6.2 Crushing

Non-cosmogenic 3He and 21Ne components that are trapped in fluid inclusions can effec¬ tively be released by cracking the inclusions through crushing the mineral grains. The two ways to do this are briefly described below.

Off-line crushing

The simplest way to crack fluid inclusions is to crush pyroxene separates in an agate mortar before loading them to the spectrometer. A loss of sample weight has to be expected because of the very fine grain size fraction that remains in suspension when the sample is cleaned and when it transferred to the aluminium foil. The serious problem that one faces is that large amounts of atmospheric noble gases are adsorbed to the the mineral grains as their surface is multiplied by the crushing. If the cosmogenic ecxess in the respective sample is small due to short exposure or low production rates, the only effect of off-line crushing is that 21Necos is masked by atmospheric instead of nucleogenic noble gases. This is illustrated by a test that was carried out with the quartz sample Lit-4b from Tibet (Schäfer et ah, 2002). One aliquot of the grain size 100 to 315 (im and one of the grain size <100/xm were analysed for cosmogenic neon by stepwise heating at 400, 600 and 1750 °C. The nucleogenic imprint that was recognisable in the neon three-isotope

22 2.6 Non-cosmogenic noble gas fractions

plot of the coarse aliquot was not visible in the fine fraction. Instead, the latter was not distinguishable of atmospheric composition, although in all temperature steps, the gas amounts were less than half as large as in the coarse fraction. It seems that a lot of neon was released from the fluid inclusions, but the effect of isolating 21Necos was lost to the adsorbed atmospheric neon.

In another experiment, it was attempted to crack fluid inclusions by heating mineral grains in a household microwave oven. The idea was to increase the pressure inside the inclusions so much that they burst. Theoretically, the pressure reaches nearly 1 kbar if the liquid is heated above 100 °C. Two aliquots of a quartz sample were analysed for their neon composition, one with microwave treatment and one without. The effect was almost nil. Both gas amounts and isotopic ratios of the two aliquots are identical within less than 8 %. This failure may have two reasons: First, the volume of the inclusions is very small, and only a very limited amount of energy is absorbed by the small amount of water in it, and boiling temperature is not reached. This little energy most likely dissipates completely to the surrounding mineral lattice. Second, the inclusions have thick walls if they are not located at the surface of a crystal. The total force exerted on the wall of a spherical inclusion of 0.1 mm diameter by a pressure of 1 kbar is a few N. This was obviously not enough to crack the inclusions.

On-line crushing

Crushing a sample in vacuo has two crucial advantages compared to off-line crushing: First, no atmospheric neon is absorbed to the newly created mineral surface; and second, the composition of the gas released from the cracked inclusions can be analysed. The on-line crushing device of the Institute of Isotope Geology at ETH Zürich con¬ sists of of a sample tray and a crushing chamber, both directly connected to the UHV extraction line. The sample tray can be moved inside the open crushing chamber and turned upside down with a magnet outside the system. In this manner, the sample is poured into the crushing chamber. To crush the sample, the crushing chamber is closed and a stainless steel ball is driven across the sample grains again by moving a strong magnet outside the UHV system. In order to prevent possible friction heating and local degassing of the sample in the crushing step, the crushing is carried out in 30-second intervals separated by 30-second breaks. After 20 intervals, the released gas is analysed. After repeated crushing and analysis of the released gas, the crushing chamber is heated with a hot air gun in order to release 21Necos from the sample. The crushed sample is later removed from the crusher, wrapped in aluminium foil and analysed in the furnace at 600 °C and 1750 °C, like the off-line crushed and uncrushed samples. This is a control if all 21Necos was released. The resulting grain size after on-line crushing is estimated as

<10 /an. The ETH on-line crushing device is described in more detail in Schäfer (2000), and its application to a quartz sample is reported in paragraph 5.2. The principle of on-line crushing is promising for the treatment of samples rich in

23 2 Method

nucleogenic neon. However, it has the disadvantage that it is time consuming. Only one sample at a time can be loaded into the ETH crushing device. Reloading means admitting atmosphere into a considerable part of the extraction line, which then requires pumping and external heating of the system for many hours. A crushing device that allows inserting several samples at a time and heating to temperatures above 800 °C would strongly increase the applicability of the technique. Another problem is that it is difficult to recover the complete crushed sample from the crushing chamber. The isotopic ratios can be measured, but not the original gas amounts. To summarize this paragraph, it must be said that for quartz samples which display a strong nucleogenic noble gas component after stepwise heating, on-line crushing is the technique to use. For all other samples, off-line crushing to 300 /iin and subsequent stepwise heating is the appropriate procedure. If a strong nucleogenic component is present, the analysis can be repeated with the <100//m fraction. For pyroxenes, finally, a crushing device is required that can be heated to 1500 °C in order to liberate 21Necos. However, as fluid inclusions are not a serious problem in pyroxenes, this issue is of lesser importance.

2.6.3 Shielded samples

A promising and in theory very simple approach to quantify the non-cosmogenic back¬ ground in a sample is to take a "zero sample" from the object that is to be dated with no irradiation. The basic assumption behind this strategy is the following: Any noble gas fraction that is contained in the top surface of an erratic is also contained in a piece of rock from underneath the block - except the nuclides produced in-situ in the top sample through cosmic ray bombardment. If two such samples are analysed from the same boulder, the difference in total 21Ne between the shielded sample and the exposed one equals the cosmogenic 21Ne. There are some limitations to this idea, though. First, a completely shielded sample is extremely difficult to obtain without drilling or using explosives, particularly if solid bedrock is involved. In a very favourable case, a tunnel, cave or cavern can be sampled. Second, an important precondition is that the lithology of the two samples is identical. Else, the trapped components might differ considerably. A thinsection helps to verify if the precondition of similar lithology is sufficiently fulfilled. In the following, it is suggested how a bottom sample can be used even if it is not completely shielded. 21 Ne excess over air found in the bottom sample can have three sources: Irradiation at the present location of the boulder, pre-exposure, or nucleogenic neon. If the shielding geometry of the bottom sample has been documented accurately in the field, the neutron fluxes / at both the bottom and top sites can be calculated. The result is compared to the concentrations C of 21Neea;c (excess over air) measured in both aliquots. If the calculated flux ratio and the measured concentration ratio agree, i. e.

Jbottom ^bottom /<-. r\

Jtop ^top

24 2.6 Non-cosmogenic noble gas fractions

Figure 2.3: Sketch of a boulder with the positions of an exposed and a shielded sample. Incoming neutrons are indicated by arrows.

25 2 Method

the excess of the bottom sample is exclusively cosmogenic and originates entirely from in-situ irradiation. Thus, no pre-exposure is involved and the non-cosmogenic 21Ne of the shielded sample can be used to calculate the cosmogenic excess in the exposed top sample:

21Necos(top) =21 Neexc(top) - (21Neexc(bottom) -21 Necos(bottom)) (2.6)

A calculated flux ratio that is smaller than the measured concentration ratio indicates pre-exposure or nucleogenic imprint. The amount of this background can be quantified because we can calculate the in-situ cosmogenic neon analogous to Eq. 2.5:

- =21 2lNecos{boUom} in situ) Necos{top) x ^^ (2.7) Jtop

The neon component in the bottom sample consisting of either nucleogenic neon or cosmogenic neon from pre-exposure or both is the given as

21Newe,nuc{bottorn) =21 Neexc(bottom) —21 Necos(bottom, in — situ) (2.8)

where 21Neexc(bottom) is the total 21Ne excess over air. 21Nepre>nuc(bottom) can then be regarded as a maximum estimate of pre-exposure under the assumption that no nucleogenic neon is involved. Of course pre-exposure at one spot of a boulder does not automatically imply pre¬ exposure at other locations on the boulder. But one would expect that pre-exposure occurring during transport on the surface of a glacier affects the entire surface of a block because it is repeatedly rotated. We have applied the approach of collecting shielded samples during the 2001 Tibet field campaign. In Nyalam, we encountered favourable conditions for taking bottom samples, as the erratics often featured large holes and openings at the bottom. This allowed us to take bottom samples for 14 out of 17 blocks sampled. The results are presented in section 4.3. That study shows that shielded samples are an efficient means of identifying the maximum pre-exposure time of glacial boulders.

26 3 Limited ice volume variations in Victoria Land, Antarctica, since middle Pliocene

27 3 Antarctica

3.1 Introduction

3.1.1 Antarctica in the global context

Antarctica is the coldest continent of the Earth and has the highest average elevation.

80% the global sweet water are locked up in the Antarctic ice sheet and glaciers. This is one of the most important aspects that make this continent to a key player in global ocean and atmosphere circulation. If a small part of the Antarctic ice melts, the global sea level will rise perceptibly in response, as it did in the past (e.g. Bentley (1999)). In the light of the ongoing climate change towards warmer conditions, the amplitude and the timing of the response of the Antarctic ice is a crucial parameter in predictions of future climate scenarios, and to understand the dynamics of the Antarctic ice sheets therefore is of more than academic interest. The cryosphere of the Antarctic is a complex system, though, and its behaviour difficult to predict with computer models. We can learn more about the nature of Antarctica's response to climate warming by looking at how the vast ice masses responded to warm climate phases of the past. This will help to create a more reliable prediction for the future (Denton et ah, 1991). Furthermore, computer models can bee improved by using descriptions of the Antarctic cryosphere at one point in the past as input parameters and then make the model calculate the ice geometry of the present. As much as "The present is the key to the past" (James Hutton, 1726-1797), the past must be regarded as the key to the future. I. e., to improve our understanding of the dynamics of Antarctic ice masses on intermediate and long timescales and under changing climatic conditions, it is crucially necessary to deepen our knowledge about the amplitude and the rates of ice volume variations in the past. A comprehensive description of the traces that the ice left during the Cenozoic is given in Denton et al. (1991). The Antarctic cryosphere consists of three main systems:

- The continental ice sheets of the East and the West

- The ice shelves fringing the land mass

- The valley glaciers and small ice caps of the Transantarctic mountains

The East and West Antarctic Ice Sheets (EAIS and WAIS) are the major parts to consider when discussing the role of Antarctica in sea-level variations. The ice shelf are already afloat in the oceans and would not contribute to a rise in sea level if they melted. The largest part of the continental ice, the EAIS, is mainly based above sea level and is expected to response with a large time lag to climate warming. Along the Transantarctic mountains (TAM), EAIS ice is dammed and spills over to the Ross Sea through wide trough valleys in the so-called outlet glaciers. These are joined by local alpine-style glaciers and small ice caps. The WAIS, in contrast, is to a large part based on the sea floor. This makes it very vulnerable to climate change. If the sea level rises, the grounding lines of the West

28 3.1 Introduction

Atlantic Ocean

270

Indian Ocean

Pacific Ocean 1000 km 180° j

Figure 3.1: Map of the Antarctic continent with the main parts of the continental

ice sheet. Shaded areas are ice shelves.

29 3 Antarctica

Antarctic ice shelves will recede, and large parts of the WAIS are predicted to become afloat. This would have dramatic impact, as the entire ice sheet could become unstable and start flowing and sliding to the sea, which in turn would drastically increase the sea level rise (DeAngelis and Skvarca, 2003; Stone et ah, 2003). It is expected that by such events also the lower marginal parts of the EAIS could float and become unstable.

3.1.2 The "stabilist - dynamist" controversy

On the long run, and regarding large-scale climate fluctuations, the greater contribution to sea-level variations can be expected from the EAIS as it is more than seven times larger than the WAIS. The history of the EAIS volume and its stability is controversial. Two main groups of scientists propagate their hypotheses and supply rich evidence in support of them. The "dynamist" hypothesis describes the EAIS of the past as a dynamic feature that underwent massive volume fluctuations in the Cenozoic (e.g. Webb et al. (1984), Webb and Harwood (1991)). As recently as 2.5MaBP, the EAIS had shrunken to about one third of its present size. Temporary marine transgression of the continent's interior lead to the deposition of marine diatoms, subsequent glaciation and rapid uplift of the TAM brought them to high mountain sites where they are found today in the Sirius Formation. The age of the Sirius Formation and the origin of its diatoms is crucial to the understanding of the Cenozoic history of the ice sheet, as the deposition of these sediments mark the last large-scale déglaciation of the continent. The other

that the ice sheet was stable since late Miocene and hypothesis , however, proposes only experienced minimal volume fluctuations (e.g. Denton et al. (1991), Denton et al. (1993), Marchant et al. (1993b), Sugden et al. (1995a) and Sugden et al. (1995b)). This "stabilist" hypothesis is supported by geomorphological and glacial geological evidence. Datings of volcanic ash and surface exposure ages of glacial sediments in the Dry Valleys of McMurdo testify for a climate that was as dry and cold as today for ar least 10 Myr (e.g., Brook et al. (1995b), Bruno et al. (1997), Ivy-Ochs et al. (1995), Marchant et al. (1996), Nishiizumi et al. (1991), Schäfer et al. (1999) and Schäfer et al. (2000), and references therein). Even though the Sirius Formation remains a puzzle (e.g. Ashworth and Kuschel (2003)), the stabilist point of view is more and more widely accepted.

3.1.3 .. .and beyond the Dry Valleys?

Most of the arguments for and against a stable ice sheet are based on evidence from the Dry Valleys, the classic spot of glacial geological research in Antarctica. This is because this region contains the largest ice free areas of the continent and is comparably well- accessible. As the Dry Valleys' tectonic block may have experienced an uplift history different from that in other blocks along the TAM, it was argued that the arguments for the stable ice sheet theory may not be valid for the entire mountain chain. Van Der Wateren et al. (1999) suggested that the climatic stability was confined to the Dry Valleys region. The logical next step on the search for an answer to this question is

30 3.1 Introduction

to investigate ice free areas beyond the Dry Valleys. Indeed, such areas are abundant

- although to a smaller extent - further north along the TAM, in the Prince Albert Mountains, in the mountains ranges inland from Terra Nova Bay and around Rennick Glacier. Many of the glacial landscape elements of Northern Victoria Land are described in Orombelli et al. (1991), Denton et al. (1991) and Denton et al. (1986). Absolute ages for glacial deposits and glaciogenic rock surfaces are required to extend our knowledge of landscape evolution in the TAM. For these regions, local chronologies of glacial advance and retreat have to be established, linked to regional ones and then correlated with the glacial record that is found in the South. This chapter includes four studies that represent first steps in the approach described above. The first, section 3.2, enlarges the number of numerical datings for the puzzling buried ice of Beacon Valley, testifying for a landscape that is deep-frozen since millions if years. Moraines from the margin of the polar plateau are dated in section 3.3. They allow an estimation for past surface elevation changes of Taylor Dome on the EAIS. Some of the first exposure ages from the Terra Nova Bay region are reported in section 3.4. They are among the oldest found outside the Dry Valleys so far. Finally, a correlation of ancient glaciogenic surfaces across Northern Victoria Land is attempted in the last of these four studies in section 3.5.

The whole of the data gathered in this chapter are consistent with a climate that was persistently cold and hyperarid since at least mid-Pliocene, throughout the entire Ross Sea section of the Transantarctic Mountains. The analysis of Pleistocene sediments for the reconstruction of the more recent glacial history is under way, too.

31 3 Antarctica

3.2 Relict ice in Beacon Valley, Dry Valleys: Extending the data set

3.2.1 Introduction

The cold desert landscape of the Dry Valleys of McMurdo in Antarctica has often been taken as a model for the surface of Mars. This is because of the similarity in appearance of the Red Planet to the barren, rocky plains full of dark, stained and desert-varnished stones and blocks in the Dry Valleys. Then this parallel also refers to the occurrence of ice in the subsurface that was is common to both places, as recently proved (Head et ah, 2003). A place where this resemblance to a remote and alien planet is particularly striking is Beacon Valley, a tributary valley to Taylor Valley near McMurdo Sound. A lobe of Taylor Glacier today enters the mouth of Beacon Valley in upvalley direction. It is assumed that, if Taylor Glacier grows as a consequence of an increasing East Antarctic Ice Sheet (EAIS), that lobe advances into Beacon Valley.

Less evident is that the debris that covers the floor of the entire valley in its flat central part overlies remnants of relict ice. According to 40Ar/39Ar dating of in-situ ash layers in the overlying debris, this ice must be older than 8.1 Myr (Sugden et ah, 1995b). In that paper, the ice is interpreted as remnants of a glacier, and the debris accordingly is a sublimation till. That ice could be preserved for such a long time is interpreted as strong evidence that the climate in the Dry Valleys has been as cold and arid as it is today since the Miocene. The remnant ice hypothesis is supported by Schäfer et al. (2000), who dated the ice with cosmogenic nuclides. An erratic from the till surface in central Beacon Valley yielded a minimum exposure age of 2.3 Myr. This age is even higher if erosion is taken into account. The exposure age of a second erratic recovered from the ice underneath the till is twenty times lower, whereas the present shielding situation of the boulder corresponds to a production rate that is only three times smaller than at the surface (Schäfer et ah, 2000). This indicates that the ice has continually lowered since at least late Pliocene and has never been thinner than today. Finally, in the steeper uppermost part of Beacon Valley, ice is present in the subsurface as rock glaciers that enter the valley from the valley slopes. They replace ice that is lost from the valley through sublimation. Rignot et al. (2002) estimate the residence time of ice in the flat central part of the valley to 0.3 to 12 Myr, linearly increasing with the ice thickness assumed.

For its special glacial environment with a large glacier front at its mouth and ice remnants underneath its rocky surface, Beacon Valley is a unique archive for climate variations in the Dry Valleys. As it acts as a sort of overflow basin for Taylor Glacier, the valley documents cold events in the region by recording advances of Taylor Glacier. The buried ice is a second climate indicator, one that is sensitive to climate warming.

In this section, a set of exposure ages for erratics from the till surface and such from

32 3.2 Relict ice in Beacon Valley

underneath the till in central Beacon Valley are presented. These extend and corroborate the findings from a pilot study of Schäfer et al. (2000) which was based on two samples. The questions that are addressed with these measurements are:

1. Was the boulder field in central Beacon Valley reworked by an advance of the Taylor Glacier lobe or does it represent an undisturbed sublimation till, as suggested by Sugden et al. (1995b)?

2. How old is the dead ice?

3. How did it evolve over time?

3.2.2 Sample description

Eight dolerite rocks from the surface of the debris field were analysed for cosmogenic 3He and 21Ne, along with five rocks that were excavated from different depths in the till and ice. They were collected in upvalley direction from those presented in Schäfer et al. (2000) (Fig. 3.2). AU surface rocks are covered by brownish desert varnish, on all sides. They are affected by oxidation to a depth of up to 5 mm. Some show distinct pithole weathering, mostly one one side only, indicating that they have remained in the same position for a long time. In contrast, the rocks excavated from the till and/or ice look fresh, nearly unweathered (Fig. 3.3). They are slightly rounded and of a light grey color. An exception is ALX-5B which has a reddish color.

ALX-5A was buried under 55 cm of till and oxidised sand. ALX-2A stuck 18 cm deep in the ice which was covered by 50 cm of sand and till. ALX-5B was recovered from 57 cm depth in the ice.

3.2.3 Methods

Pyroxene separates were produced and analysed by stepwise heating, as described in sections 2.4 and 2.5, with the difference that the compressor of the ion source was turned off.

For the calculation of the neon production rates, the pyroxenes were analysed for their Si, Mg and Al content by ICP-OES and laser ablation ICP-MS at the Institute of Inorganic Chemistry at ETH Hönggerberg. Element concentrations are given in Tab. 3.1. See caption of Tab. 3.4 for more details about the production rates.

3.2.4 Results and Discussion

The 3He and 21Ne dataand the resulting exposure ages are listed in tables3.2, 3.3 and 3.4. All samples released most of their 3He at 850 °C and most of their 21Necos at

33 3 Antarctica

^W - %:

Figure 3.2: Map of Beacon Valley with sampling locations. Black dots indicate sur¬ face samples, white dots are buried samples. Ages in the boxes are

minimum ages of surface samples. Numbers in white boxes are the high¬ est surface sample age in kyr for each location (compare Tab. 3.4). The white dotted line marks the approximate limit of Granite Drift (Tschudi et al., sub).

34 3.2 Relict ice in Beacon Valley

Figure 3.3: Samples ALX-2B (above) and ALX-2A (below), taken at the same sam¬ pling site, the first from the surface, the second from within the relict ice. Note the striking difference in weathering degree.

35 3 Antarctica

Table 3.1: Relative concentrations of the elements that produce Necos and the 21Ne production rate at sea level and high latitude (slhl),

calculated with P2i(Si) = 42.5 at g"1 yr"1, P2i(Mg) = 185atg_1yr_1,

P2i(Al) = 51atg-1yr"1 (Schäfer et al., 2000).

Sample Si Mg Al P21 slhl (weight %) (weight %) (weight %) (atg--lyr"1)

ALX11-1A 25.2 9.2 1.4 28.5

ALX-11-2B 22.2 0.3 12.5 16.4

ALX-2A 21.8 6.0 2.0 21.4

ALX-2B 23.5 6.8 5.2 25.1

ALX-2C 24.0 8.6 1.0 26.6

ALX-5A 23.4 8.2 1.0 25.7

ALX-5B 27.0 7.8 1.1 26.5

DLE-14S 26.7 6.0 4.1 24.6

DLG-14X 26.1 8.0 2.6 27.3

DLX-06Y 22.6 6.5 2.6 23.0

DLX-H 26.8 6.5 2.2 24.6

Table 3.2: Helium isotope data from the 850 °C steps, used for age calculation. Un¬ certainties are 2

Sample 4He 3He 3He/4He (1012 at g"1) (106 at g"1) R/Ra

ALX-11-1A 43.52 ± 0.10 797.58 ± 41.23 13.24 ± 0.69

ALX-11-2B 8.30 ± 0.03 8.52 ± 0.75 0.74 ± 0.07

ALX-2B 49.41 ± 0.14 376.88 ± 19.66 5.51 ± 0.29

ALX-2C 122.98 ± 0.78 165.85 ± 6.17 0.97 ± 0.04

DLE-14S 182.48 ± 0.50 84.91 ± 4.74 0.34 ± 0.02 DLG-14X (TE) 111.18 ± 0.30 96.10 ± 5.17 0.62 ± 0.03 DLX-06Y 360.84 ± 3.42 184.15 ± 10.12 0.37 ± 0.02 DLX-H (TE) 447.84 ± 1.69 82.54 ± 4.55 0.13 ± 0.01 ALX-2A base (TE) 274.19 ± 0.67 22.54 ± 1.43 0.06 ± 0.00 ALX-2A top 251.86 ± 1.04 18.83 ± 1.35 0.05 ± 0.00 ALX-5A base 172.12 ± 1.07 16.62 ± 0.88 0.07 ± 0.00 ALX-5A top 185.50 ± 0.68 16.29 ± 1.22 0.06 ± 0.00 ALX-5B 223.54 ± 1.49 179.60 ± 6.67 0.58 ± 0.02 > > > > > > > > > > > a a a a a a a a y^—N rr> t-l t-l f f f f t-1 t-1 t-1 t-1 t-1 t-1 t-1 t-1 t-1 t-l t-1 t-1 t-1 M M M M M M M M M t-l t-1 M O B î^s î^s î^s î^s î^s î^s î^s î^s n n r^ X X X X X X X X X P*s P*s X s? I I I I I I I I I I I I I I I I I I I-1 Cn en Cn en en en en en en to to to 0 o o h^ tG tG tG tG tG tG M M M M M M B CD" td tö td 01 Ol Ol td tel td > > > > > > > > > ce co O O O 1 1 1 1 1 1 p K| !" > > 00 œ W tî t) M M 00 Ol Ol -4 Ol Ol ~4 I—1 U! O O S Ol Ol Ol c—1 c—1 '/) c—1 oo Ol o o t—i O i/j M M 1/1 O O CO M M 00 •—s M M 00 o o oo UJ oo Ol Ol -a en en 93 BS P oo 5' O S Ol Ol S Ol Ol P P oo CO M M o M O o o o 00 o c—> g "—' 00 o o -a en en 00 O O M M O o Ol Ol Ol o o O CD en en oo p p "oo P P ~oo P P P 00 o O M M O o o CO M M O CL O en en -—' CD P O

OO M CD P&3COP03HffiCOW OO to ^ o go' to OO ^ en oo to ai to SOOlOltOOlOPKJ oo o en «D M h^ oo «D M O oo en to ai 00 to ai en oo ai oo to oo ai to 00 o to oo en en oo o to O 00 OO ai 00 00 00 -a 00 o to 00 en en 00 oo ai en CO 00 m H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- aio^i^ioo^icr>entoenoo^oooh^ootOh^cr>ooooenooooooo^iooh^ ^ ^—^ CD H-H-H-H-H-H-H-H-H-H-H-H-H-H-H-H-H-H-H-H-H-_H-H-H-H-H-H-H-H-H-H-H- ~^^ B M to p M M M o o o o o o o o o o o O o o o o o o ai o o o o o o o o M o c O o P f£T 1 ai 00 to en ai to to en ai to o oo o o to to o o tO h 0Û ^ CO ^-ï OO en en en en to en en en 00 en en 00 to oo ai 00 to 00 h CD ft) ai O -a -a o oo - — - CO CO~ ^ o o o o o o o o o o o O o O o o o o o o o o o o o o o o o o o o c O CD Cad o o o o O O O o o o O a m M o O o M o O o O o O O O o o o o o o to o to 00 en ai 00 to 00 -a en oo to 00 oo to oo -a to to en en en to oo en -a 00 oo -a en to ai ai en -a ai 00 00 ai ai ai o o en to p d o to to p H- H- H- oH- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- 1 o o o o o o o o o o o o o o o o o o o o o o o o o o o o o o o o o o CD CD o* to o o o o o o o o o o o o o o o o o o o o o o o o o o o o CD P o to ai oo oo 00 to M 00 to M M to oo oo ai M ooo -a ai M to 1 o* o o o o 00o o o oen o ooo o o oto oai o-a oto o to to o en oen o00 o00 oai o-a o00 o 00o ooo CD o o o o o o o o o o to oo oo oo 00 to o oo to oo ta oo oo en to to oo en to to 00 o «D o to to o 00 M ai M oo ai en to en oo oo 00 ai OO O So o O^ to O en O ai to 00 00 00 ai to en ai ai to en to o H- H- 03 to Ö H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- H- oH- H- H- H-

1500°C. ALX-11-2B is an exception. Its 21Necos was released at 850°C, and very little 3He was measured. This sample contains at least 20 times less Mg than the other samples (Tab. 3.1), the element that has the highest production rate of 21Necos. In section 3.5 it is shown that in dolerite samples, the Mg content is inversely proportional to the percentage of 21Necos that is released below 850 °C. This sample is a striking confirmation of that hypothesis. The almost complete absence of 3He indicates that Mg poor pyroxenes might not be retentive for cosmogenic 3He.

In Tab. 3.4, the exposure ages of the two samples from Schäfer et al. (2000) are in¬ cluded. They are based on the concentrations given in Schäfer et al. (2000), but calcu¬ lated using the scaling by Stone (2000), as for the samples from this study. The original ages are 32 % higher on average.

The helium and neon ages agree within errors for eight of the fifteen samples, whereas for the others, the helium age is lower. As exposure ages generally are minimum ages, the higher age obtained for each sample is used for the discussion. The same holds for ALX-2A and ALX-5A, two buried samples for which an aliquot from the top of the rock and one from its bottom was measured. The goal of this experiment was to test if the effect of self-shielding of the rock is measurable. This was not the case in either of these samples, as the ages from top and bottom agree within errors, but the top ages are higher (Tab. 3.4). It is assumed that the shielding by the boulders of 11 and 15 cm thickness was too weak to be detected in the isotope concentrations.

Surface samples

The oldest surface sample has been exposed for at least 2.6 Myr. This number is the minimum age of the last reworking of the boulder field in Beacon Valley. Most probably, the Beacon Valley lobe of Taylor glacier has not reached the central part of the valley for a much longer time, though. At 2.6 Myr ago, this boulder was possibly unburied from the sublimating ice, whereas the latter had been deposited much earlier. This idea explains why the ash datings of Sugden et al. (1995b) (8.1 Myr) from Beacon Valley yield considerably higher ages for the remnant ice.

The spatial distribution of the ages of the surface samples shows a trend: the ages become generally higher in downvalley direction. The uppermost part of the valley in the SSE (Fig. 3.2) is slightly inclined and features rock glaciers with the typical pressure ridges indicating deformation in flow direction as well as transverse to it. On this active surface, rocks are repeatedly overturned, buried and uncovered again. Therefore, surfaces are comparatively young in this area. The central part (sampling locations 1 to 5), in contrast, is flat and without flow pattern. Surface motion is only caused by frost cracking (patterned ground). The samples discussed here, however, were collected in the centres of frost polygons, where such were present, and it is not probable that they have fallen into a frost crack or were heaved and overturned by frost action. The lowest ages are therefore found in the uppermost sampling sites. Thus, according to the distribution of surface samples, surface activity seems to decrease in downvalley

38 3.2 Relict ice in Beacon Valley direction. Rignot et al. (2002) have determined the surface velocity field in Beacon Valley with Interferometric Synthetic-Aperture Radar measurements. Average surface velocities on the rock glaciers are 10 to 20 mm yr-1, whereas outside the rock glaciers, surface motion on the valley bottom is absent. The age distribution described above is consistent with those results.

Buried samples

With the exception of ALX-5B, all buried samples have lower exposure ages than the surface samples. ALX-5B, however, shows a stained surface, which indicates that the rock might have been exposed to weathering and - as its 21Necos content shows - cosmic radiation before it was entrained in the ice. Therefore, this sample is excluded from the discussion of buried samples.

As the buried sample ALX-2A and the surface samples ALX-2B and ALX-2C were taken at the same spot, their 21Necos concentrations can be compared in order to find out more about the ice thickness variation. ALX-2A contains less than 18 % of the 21Necos found in ALX-2B at the surface. The present shielding by 50 cm of sand and 18 cm of ice, however, corresponds to a production of 46 % of that at the surface (calculated using densities of 2.0 g cm-3 for sand and 0.9 g cm-3 for ice, and an attenuation length of 150 g cm-2). This difference must be caused by a thickness reduction of the shielding ice. The thinning has been continuously and not by collapsing, because boulder ages in Beacon Valley are rather evenly distributed between the lowest and the highest age, which means that boulders are permanently, although slowly, transported to the surface. In the case of a rapid downwasting of the ice, surface boulders would all have similar ages. A continuous lowering of the ice surface, most likely, has occurred by sublimation, as no traces of meltwater can be found in Beacon Valley.

3.2.5 Implications for the paleoclimate

1. The floor of central Beacon Valley was not reworked since at least 2.6 Myr. Thus, Taylor Glacier and the EAIS have not undergone significant volume changes since

then. This age is a minimum for the age of the buried relict ice. As it might represent boulder exhumation and not ice deposition, it is consistent with the

Miocene age of ash layers found in the till that covers the ice.

2. The ice under the till has never been thicker than today. Ice loss has occurred continuously since 2.6 Myr through sublimation and not through collapsing.

3. The preservation of ice over such a long time and its continuous thinning require a constantly cold and hyperarid climate. This rules out any significant warming of Antarctic climate since Mid-Pliocene.

39 3 Antarctica

The results presented in this section are entirely consistent with those from Schäfer et al. (2000). The implications drawn from them thus confirm the findings made by that study.

40 3.2 Relict ice in Beacon Valley

Table 3.4: Helium and neon exposure ages calculated from the 850 °C (He) and 1500 °C (Ne) steps, resp. Production rates: 3He, 115atg_1yr_1 (Cer¬ ling and Craig, 1994b); 21Ne, see Tab. 3.1. Scaling according to Stone (2000). "Location" refers to Fig. 3.2.

21 Sample Location 3He minimum Ne minimum age (kyr) age (kyr)

Surface samples ALX-11-1A 2 1546±80 1931±43

ALX-11-2B+ 2 17±1 2592±135

ALX-2B 3 730±38 1118±48

ALX-2C 3 321±12 276±49

DLE-14S 5 165±9 220±28

DLG-14X 5 186±10 188±28 DLX-06Y 4 357±20 424±104

DLX-H 6 160±9 278±42

DME3* 1 1729 ± 54 1737 ± 131 Buried samples ALX-2A base 3 44±3 200±53 ALX-2A top 3 36±3 106±52 ALX-5A base 3 32±2 51±26 ALX-5A top 3 32±2 28±13 ALX-5B 3 348±13 322±32

DME1* 1 80 ± 7 91±17

+ Age based on 850 °C neon

* Calculated with isotope data from Schafer et al (2000), using different production rates and scaling formalism

41 3 Antarctica

3.3 Relative and absolute chronologies of glacial drifts in Vernier Valley: Indications for a stable climate and ice sheet in Antarctica since Mid-Pliocene

In this chapter, surface exposure ages of moraines on Pivot Peak nunatak in Vernier Valley, Antarctica are presented. These ages are part of a larger study that incorporates investigation of weathering characteristics, volcanic ash properties and climate data from the area. That study is authored by Jane Staiger, David R. Marchant, Peter Oberholzer, Joerg M. Schaefer and Adam R. Lewis (Staiger et al., unp) and, at the time of this writing (January 2004), is almost ready for submission to Geomorphology. As soon as more data on the composition of volcanic ash samples are available, the manuscript will be submitted. This chapter reports exposure ages from Vernier Valley and draws some conclusions from them. I have prepared and analysed these samples at the noble gas laboratory at ETH Zürich together with Jörg Schäfer. For a more general interpretation, however, data from the manuscript by Staiger et al. (unp) is incorporated. All information on the local geomorphology and the properties of the glacial drifts is taken from that manuscript.

Abstract

Vernier Valley on Pivot Peak nunatak, Western Dry Valleys, opens onto upper Ferrar Glacier, an outlet glacier of the East Antarctic Ice Sheet. A lateral lobe of Ferrar Glacier enters the valley. Glacial drifts in Vernier Valley deposited by this lobe record ice surface

variations of upper Ferrar Glacier and thus of nearby Paylor Dome. Phese drifts were dated

with in-situ produced cosmogenic noble gases. A relative stratigraphy was established by examining the weathering characteristics of glacial boulders. Phe minimum exposure ages show that the ice in Vernier Valley was 125 m above the present ice surface in Pliocene, and had lowered to 50 m above it by mid-Quaternary. Phese minimal variations can be modelled into surface elevation changes on Paylor Dome, which are roughly half as large as in Vernier Valley. Pogether with strong evidence that these glacial drifts were deposited under cold basal conditions, these small volume changes since the Pliocene imply that

the Dry Valleys region has experienced cold-desert climate conditions - much like those

prevailing today - since at least Mid-Pliocene.

3.3.1 Introduction

Large glaciers that drain the Antarctic inland ice by cutting through the Transantarctic Mountains are called outlet glaciers. The largest of them are Taylor, Ferrar, David and Campbell Glaciers. Their deposits alongside the valley walls are of great interest for paleoclimate research because the variations in the outflow of ice through the outlet glaciers reflects ice volume variations of the feeding inland ice. Thus, if the record of the outlet glacier deposits in the today ice free areas of the continent can be deciphered, a history of ice volume change can be established. It is indeed the deposits of the outlet glaciers in the Dry Valleys of McMurdo that contain much of the evidence for the stable ice sheet hypothesis (e. g. Schäfer et al.

42 3.3 Vernier Valley moraines

(1999); Sugden et al. (1995a); Marchant et al. (1993b)), as well as against it (e.g. Webb and Harwood (1991), Webb et al. (1984)). This debate is summarized in section 3.1. The aim of this study is to expand the available database of absolute ages for glacial deposits in the Transantarctic Mountains. By examining evidence of past glacial activity that comes from close to the inland ice, a quantitative constraint on the magnitude of the volume fluctuations should be possible. Changes in ice surface elevation are recorded immediately on the nunataks closest to the ice sheet, without the influence of valley topography or grounding ice on the sea. We investigate glacial deposits in lower Vernier Valley, western Dry Valleys, where it opens onto the upper part of Ferrar Glacier. When the Ferrar Glacier surface rises, as expected in response to a ice-surface rise of the nearby Taylor Dome, the ice enters Vernier Valley. This has happened several times in the past as indicated by a set of moraines in the lower part of the Valley. In order to find out when such highstands of the inland ice have occurred, these moraines were dated with the method of surface exposure dating using in-situ produced cosmogenic noble gases 3He and 21Ne. The weathering characteristics of the boulders and pebbles composing these deposits were examined. Furthermore, the chemical composition of volcanic ashes that have been found between the pebbles and boulders were analysed and their micromorphology studied in order to identify their geographical origin and the mechanism of their entrainment into the till.

3.3.2 Geomorphological setting

Ferrar Glacier is an outlet glacier that, unlike nearby Taylor Glacier, does not terminate on land, but in the Ross Sea. It drains part of Taylor Dome on the East Antarctic Ice Sheet (EAIS), passes between the Royal Society Range and the Kukri Hills in the Dry Valleys region and ends in the sea ice of McMurdo Sound. The lowest ten kilometers of Ferrar Glacier are floating. Vernier Valley is a 3.5 km long valley on Pivot Peak nunatak, halfway between Mount Feather and Table Mountain. The valley joins Ferrar Glacier in its uppermost part. The glacier enters Vernier Valley and presently deposits a lateral moraine in its lowest part. On the righthand side of the valley, three different glacial drifts are found. These drifts are informally called Ferrar 2, Ferrar 3, and Ferrar 4 by Staiger et al. (unp). Each drift is limited uphill by a small, but well-developed moraine ridge. These ridges are sub-parallel to each other and to the modern ice surface. At present, an ice-cored lateral moraine is being developed by the glacier ice. This is called Ferrar 1 drift. All three ancient drifts consist of sandstone and dolerite clasts of gravel to boulder size, with a sparse matrix of grus and sand. Glacial striation is absent, but perched boulders occur in Ferrar 1 and 2 drifts. The following is a brief description of the sampled drifts:

Ferrar 1 The lateral moraine presently developing on the margin of the glacier. Clasts

are little weathered, some are perched. Where the thickness exceeds 30 cm, ice sub-

43 3 Antarctica

mfr^mm

^r HE/ I î 1

10 km (USBS, 19

Figure 3.4: Map of the southern part of the Dry Valleys with Beacon Valley and Pivot Peak. Taylor and Ferrar Glaciers are visible in the upper right part, flowing in eastern direction.

44 3.3 Vernier Valley moraines

limation is reduced and an ice core has developed. Maximum elevation: 1450 m a. s .1., at present glacier margin.

Ferrar 2 Occurring from the present ice margin up to 1500 m a. s .1.. A continuous sheet of dolerite and sandstone clasts displaying desert varnish. Between angular clasts and within the sandy matrix, silt-sized volcanic ash is found.

Ferrar 3 Extending up to 1550 m a. s .1., a layer of scattered boulders, in places piled up

to build a moraine. Contains more than 5 % very well-preserved volcanic ash in pores between openwork gravels. The drift rests directly on bedrock, undisturbed colluvium and Sessrumnir till of pre-Mid-Miocene age (Marchant et ah, 1993a).

Ferrar 4 Maximum elevation of 1575 ma. s .1., i. e. 125 m above the modern glacier sur¬ face. Clasts are similar to those of Ferrar 3, resting on the same units, but the clasts are more heavily affected by desert varnish and pithole weathering. Ashes occur in lenses and layers.

From their position and geometry, it can be inferred that these drifts were deposited by the same ice tongue that occupies the bottom of the lower valley today. The ice level of the Ferrar Glacier in its upper part, hence also of its lobe entering Vernier Valley, reflects the combined effects of

1. volume changes of the East Antartic Ice Sheet (EAIS) on nearby Taylor Dome,

2. ice volume fluctuations of the alpine glaciers covering the surrounding mountains, and

3. grounded sea ice in McMurdo Sound, potentially damming the discharge of ice into the sea by Ferrar Glacier.

3.3.3 Methods

Moraines were absolutely dated by determining the exposure age of boulders collected on their top. All boulders were of Ferrar Group Dolerite. High-purity pyroxene sep¬ arates were obtained by physical separation techniques described in section 2.4 of this thesis. The concentrations of cosmogenic noble gases were measured in a stepwise heat¬ ing procedure on a non-commercial, ultra-high sensitivity noble gas spectrometer at the Institute for Isotope Geology and Mineral Resources at ETH Zürich. Cosmogenic 3He was measured at 850 °C, cosmogenic 21Ne at 850 °Cand 1500 or 1750 °C. Details about noble gas extraction and purification are given in section 2.5. In order to determine accu¬ rate neon production rates, the relative concentrations of Si, Al and Mg in the pyroxene samples were measured with ICP-OES. The obtained sea-level, high-latitude production rates were then scaled to the elevation of the sampling site with the procedure described

45 3 Antarctica

by Stone (2000). This procedure takes the air-pressure anomaly in Antarctica into ac¬ count. For helium, a sea-level, high-latitude production rate of 115 at g_1 (Cerling and Craig, 1994b) was used. As a check for a possible irradiation of the boulders during transport on the glacier surface, samples were collected from Ferrar 1 drift, the presently deposited debris at the glacier margin. Any 21Ne excess over air that is found in Ferrar 1 samples is likely to be the result of such pre-exposure, as they have not been irradiated in their present position for a substantial period. If a considerable pre-exposure is found, it can be used as a tentative correction on the samples used for dating of Ferrar 2 to 4 drifts.

3.3.4 Results

The concentrations of cosmogenic noble gases are given in table 3.5 on page 48. Neon ratios are plotted in figure 3.5. Neon ages were calculated from both the 850 °C and the 1500 °C steps for samples DXP-99-37 to 40. Only the 1500 °Cdata were used for the other samples because their low-T data plot below the cosmogenic/atmospheric mixing line. The helium ages were calculated from the 850 °C step alone. For DXP-99-48, which was analysed in a single step at 1750 °C, no neon excess over air was detected.

The agreement of helium and neon ages on the ancient moraines is excellent, with the exception of the control sample from the modern moraine. All samples have identical neon and helium ages within errors. The age difference in the control sample DXP-

99-42 from Ferrar 1 can not be caused by pre-exposure, as any cosmogenic irradiation produces identical helium and neon ages. Nucleogenic neon is improbable, because the data plot on the atmospheric-cosmogenic mixing line. The only reasonable explanation would be a nucleogenic fraction of coincidentially cosmogenic composition.

3.3.5 Discussion

Exposure ages

Because cobbles at the surface of moraines may have rotated after their deposition, and because we neglected the effects of nuclide loss through erosion in our age calculations, results reported in table 3.5 are strict minimum drift ages. Although denudation rates for the western Dry Valleys region are extremely low (Summerfield et al. (1999); less than 0.27m Ma-1 in the nearby Quartermain Mountains, see Brook et al. (1993), field observations indicate that some boulders on Ferrar moraines in Vernier Valley may have experienced splitting or spallation. This is, together with possible overturning of a boulder, an explanation for the considerable age difference between samples DXP-99- 38 and DXP-99-39, which were collected on the same moraine.

In order to estimate the duration of possibly occurring pre-exposure, two control sam¬ ples were taken from the recently developing lateral moraine at the actual ice margin {Ferrar 1 drift). Any cosmogenic neon in these samples (DXP-99-42and DXP-99-48) is

46 3.3 Vernier Valley moraines

0.120 442 40 41

' 0.100 37

0.080 0.060 0.0000 0.0800 0.1600 0.0020 0.0040 0.0060 B 0.140

0.130- 0) Panel C 0.120 r

2> 0.110- AIRm]®

850 "COT 1500 "CflS 1750 "CflS

0.0100 0.0200 0.0300 21Ne/20Ne

Figure 3.5: Neon three-isotope plot of Vernier Valley samples. Filled symbols are

data points used for age calculation, and the open symbols stand for data points that are not regarded as cosmogenic. Sample labels are given with¬ out the prefix DXP-99-. The equation of the atmospheric-cosmogenic mixing line is: y=(1.069±0.035)x+0.099 (Schäfer et al., 1999).

47 3 Antarctica

Table 3.5: Isotope data for the Vernier Valley boulders.1

3 21 Sample Deposit Elevation 3He 21Ne He age Ne age (ma. s .1.) (108 at g"1) (107 at g"1) (kyr) (kyr)

DXP-99-42 Ferrar 1 1420 0.11±0.01 0.53±0.01 20±1 48±9 DXP-99-48 Ferrar 1 1400 0.16±0.01 0.90±1.50 31±2 (77±128) DXP-99-41 Ferrar 2 1480 2.45±0.15 5.30±0.36 438±26 409±28

DXP-99-46 Ferrar 2 1510 n. a. 2.55±0.14 — 193±11

DXP-99-40 Ferrar 3 1510 n. a. 14.61i0.76 — 1021±53

DXP-99-37 Ferrar 4 1550 8.89±0.25 18.99il.25 1505±42 1339±88

DXP-99-38 Ferrar 4 1550 9.34±0.32 19.19±2.06 1581±54 1555±167 DXP-99-39 Ferrar 4 1480 18.84±0.06 38.64±2.04 3370±169 3134±166

1 Production rates were calculated and scaled to geomagnetic latitude and elevation according to Stone

(2000), using a sea-level, high latitude value that was corrected the Si, Al and Mg contents of the sample. Errors are 2

21Ne calculated as excess over air. All 3He is assumed to be cosmogenic. n. a.: not analysed.

thought to represent irradiation during transport on the glacier surface or at a different site before the transport. It is not possible to tell what part of it occurred during trans¬ port or before. However, the age difference between these two samples underlines the following: Neither have all samples undergone the same journey on the glacier surface, nor is it probable that the duration of exposure at a different site was identical. The measured 21Ne excess over air in DXP-99-42 corresponds to an exposure period of 48 kyr, and the helium ages are 20 and 31 kyr (see table 3.5. In order to obtain a maximum esti¬ mation of the pre-exposure, the highest age, i. e. the neon age of DXP-99-42, is used. It corresponds to 2 to 4 % of the Ferrar 4 samples, 5 % of the Ferrar 3 sample, and 12 % of the Ferrar 2 sample, respectively. Thus, substantial pre-exposure can be excluded for all samples. If all boulders had experienced a similar exposure history during their transport to Vernier Valley, then the neon ages presented in table 3.5 would be reduced by 48 ka. Thus, if the isotope data are interpreted conservatively, Ferrar 2 drift is Qua¬ ternary in age, and Ferrar 3 and 4 drifts are Pliocene in age. This would not alter any of the conclusions drawn in this study.

The exposure ages of the three glacial drifts analysed are entirely consistent with the relative position of the deposits in the field. The highest and most distal moraine is the oldest, and the lowest one is the youngest. This implies a simple progressive deposition of the Ferrar 1 to 4 drifts by a lobe of Ferrar Glacier retreating from Vernier Valley without significant readvances. Accordingly, 3.4MaBP, the ice level was roughly 125m higher than today. By about 1 MaBP, the ice reached up to 100 m, by 440 ka ago, up to 50 m above the present-day ice surface.

48 3.3 Vernier Valley moraines

Weathering characteristics

This numerical chronology is also consistent with the relative chronology that is inferred by the weathering characteristics of the four drifts. While the frequency of pitholes and the thickness of desert varnish do not vary much among the clasts of a moraine, they show sharp increases from one moraine to the next older one. This pattern also strongly suggests a progressive deposition of the four drifts by a ice body that retreated with phases of stagnation which led to the deposition of moraines.

Ice dynamics

From these data, it can not be excluded that the ice has re-advanced between the reces¬ sional phases documented by the sampled drifts, leaving no deposit behind. However, given the strong katabatic winds that occur in Vernier Valley, it is most likely that any substantial glacial advance would result in a distinct sediment. Therefore, we regard the probability of a substantial ice advance into Vernier Valley after 3.4 Myr BP as very small.

It is very probable that the drifts Ferrar 2 to 4 were deposited under cold based conditions like the recent moraine Ferrar 1. This is based on the absence of glaciofluvial features in Vernier Valley, on the fact that unconsolidated till and colluvium was pre¬ served under the glacial drifts, and on the existence of ice tongues thinner than 125 m for a considerable period.

As mentioned above, the record of ice-level variations on upper Ferrar Glacier is the result of three factors: volume fluctuations on Taylor Dome, variations in glaciation on the surrounding mountains and grounded ice in McMurdo Sound. Hence, the limited ice-level variations in Vernier Valley documented by the dated moraines does not quan¬ titatively reflect fluctuations of the EAIS volume alone. They rather represent an upper limit to them. I. e., assuming that the variations of the ice level on upper Ferrar Glacier are caused entirely by ice-volume changes at Taylor Dome, the ice-level variations de¬ termined in Vernier Valley can be used to model changes in surface elevation at Taylor Dome accordingly. Simple 1-D models assuming plastic ice flow (e. g. Marchant et al. (1994)) assign to the 125m Pliocene ice level maximum in Vernier Valley an ice surface rise at Taylor Dome of 63 m. The Quaternary glacier highstand of 50 m in Vernier Valley would then reflect an elevation increase at Taylor Dome of 25 m above the present sur¬ face. These numbers are maximum values, though, as the ice-level variations in Vernier Valley are not necessarily caused by inland ice volume fluctuations alone.

3.3.6 Conclusions

The ages of three glacial moraines in Vernier Valley, upper Ferrar Glacier, were deter¬ mined with surface exposure dating using cosmogenic noble gases 3He and 21Ne. Sub¬ stantial pre-exposure has been ruled out by a control sample from the present-day ice

49 3 Antarctica

margin. The results show that these sub-parallel moraines were deposited by an ice lobe of Ferrar Glacier continuously retreating from the Valley. The ice was highest in Pliocene, at least 3.4 Myr ago, when it reached 125 m above the present-day ice-surface (deposition of Ferrar 4 drift). At IMaBP, the ice stood 100m higher than today (deposition of Ferrar 3 drift). Until 440kaBP, the ice had lowered by 75m (deposition of Ferrar 2 drift). This age sequence is confirmed by a relative stratigraphy set up by the distribution pattern of weathering parameters (abundance of weathering pits and thickness of desert varnish) on surface pebbles from the drifts. According to 1-D models, these surface elevation changes of upper Ferrar Glacier correspond to ice-level variations on nearby Taylor Dome of less than 63 m above the modern surface in Pliocene and less than 25 m in Quaternary. These minimal ice volume fluctuations of Taylor Dome since 3.4 Myr, together with the fact that in this time, ice existed most likely under cold basal conditions, imply that the climate in the western

Dry Valleys - and probably in the entire region - was persistently cold as it is today. This finding supports the hypothesis of long-term stability of the EAIS.

50 3.4 Mount Keinath

3.4 Limited Pliocene/Pleistocene glaciation in Deep Freeze Range, northern Victoria Land, Antarctica, derived from in-situ cosmogenic nuclides1

PETER OBERHOLZER1*, CARLO BARONI2, JOERG M. SCHAEFER3, GIUSEPPE OROMBELLI4, SUSAN IVY OCHS5, PETER W. KUBIK6, HEINRICH BAUR1 AND RAINER WIELER1

Institute of Isotope Geology and Mineral Resources, ETH Zürich, Sonneggstrasse 5, 8092 Zürich, Switzerland

Dipartimento di Scienze della Terra, Universita di Pisa, e Istituto di Geoscienze e Georisorse CNR, 3 Via S. Maria 53, 56126 Lamont-Doherty Earth Observatory, Columbia University, Palisades, NY 10964, USA 4 Dipartimento di Scienze dell'Ambiente e del Temtorio, Universita di Milano-Bicocca, Piazza della Scienza 1, 20126 Milano, Italy Institute of Particle Physics, ETH Zürich, ETH Hönggerberg, 8093 Zürich, Switzerland 6 Paul Scherrer-Institut, c/o Institute of Particle Physics, ETH Zürich, ETH Hönggerberg, 8093 Zürich, Switzerland * corresponding author: [email protected]

Abstract Phe question of how stable the climate in Antarctica has been during the last few million

years compared to the rest of the planet is still controversial. Phis study attempts to add new information to the dicussion by reconstructing the timing and spatial extent of glacial

advances in northern Victoria Pand over tens of thousands to millions of years. In Perra

Nova Bay region, surface exposure ages and erosion rates of glacially rounded bedrock and glacial erratics have been determined using the cosmogenic nuclides 3He, 10Be and 21Ne.

Phree morphological units have been analysed. Phey yield minimum ages of 11 to 34 ka, 309 ka, and 2.6 Ma, respectively. Erosion rates were as low as 20cmMa~1 since middle Pliocene time. Paking erosion into account, the oldest surface is 5.3 Ma old. Pleistocene

glacier advances had considerable extent, reaching up to 780 metres above modern ice levels, but have been restricted to the valleys since at least mid-Pliocene. Phe existence

of landscapes of mid-Pliocene age in northern Victoria Pand implies that the climatic stability of the Dry Valleys is not unique within the Pransantarctic Mountains, but rather the expression of a constantly cold and hyperarid climate regime in entire Victoria Pand.

Keywords: Surface exposure dating, glacial drifts, erosion rates, East Antarctic Ice Sheet, Pleistocene, Terra Nova Bay.

This article was published in Antarctic Science 15 (4), 2003, pp 493 - 502

51 3 Antarctica

3.4.1 Introduction

The glacial history of the Antarctic continent during late Tertiary and Quaternary times is still not fully understood. Was Antarctica decoupled from global climate variations, or has it been subject to similar fluctuations as took place elsewhere on the planet? This study addresses this question by reconstructing glacial advances using surface exposure dating with in-situ produced cosmogenic nuclides together with detailed geomorpholog- ical and glacial geological mapping.

The Terra Nova Bay region is situated in the Transantarctic Mountains on the Ross Sea coast at 74°30'S and contains extensive ice-free areas. Well-preserved glacial features such as rounded mountain tops, striated bedrock and erratic blocks are widespread. This makes the area particularly interesting for the study of landscape evolution, since ice- free areas are very rare in Antarctica. On the south ridge of Mt. Keinath, a sequence of glacial landforms and deposits is preserved that yields a relative chronology. The glacial record also provides valuable information on ice mass changes of the East Antarctic Ice Sheet (EAIS), because glacial drifts along its outlet glaciers record the limits of inland ice surface elevation (Marchant et ah, 1994). To establish a chronology of Antarctic ice volume expansion and reduction and landscape history, dating of the imprint left by the ice is essential. In the extremely cold and arid environment of the Antarctic, surface exposure dating has proven to be a valuable tool to date glacial deposits of wide age range (e. g.Nishiizumi et al. (1991), Brook et al. (1993), Schäfer et al. (1999), Ackert et al. (1999)). The history of the EAIS and its stability is controversial and framed by two opposing hypotheses. One advocates a dynamic ice sheet and temporary marine transgression of the continent's interior (e.g.Webb et al. (1984), Webb and Harwood (1991)). The other hypothesis proposes a stable ice sheet with minimal volume fluctuations in the last 14 Ma (e.g.Denton et al. (1991), Denton et al. (1993), Marchant et al. (1993b), Sugden et al. (1995a) and Sugden et al. (1995b)). The latter point of view is supported by geomorphological and glacial geological evidence dated on volcanic ash, and surface exposure ages on glacial sediments in the Dry Valleys, 300 km south of Terra Nova Bay (e.g., Brook et al. (1995b), Bruno et al. (1997), Ivy-Ochs et al. (1995), Marchant et al. (1996), Nishiizumi et al. (1991), Schäfer et al. (1999) and Schäfer et al. (2000), and references therein). As the Dry Valleys' tectonic block may have experienced an uplift history different from that in other blocks along the Transantarctic Mountains, the arguments for the stable ice sheet theory may not be valid for the entire mountain chain (Van Der Wateren et ah, 1999). Thus, absolute ages from old surfaces outside the Dry Valleys are required to extend knowledge of landscape evolution in the Transantarctic Mountains and put existing data in a wider context. The mountain ridges west of Terra Nova Bay are highly appropriate terrain for obtaining such data.

In this study, the radionuclide 10Be was used in concert with the stable noble gas nuclides 3He and 21Ne This approach helps to reveal complex exposure histories and allows erosion rates to be determined (Graf et ah, 1991).

52 3.4 Mount Keinath

3.4.2 Geomorphological setting

Landscape elements

The two sites sampled for this study, Mt. Keinath (1090 m a. s .1.) and Black Ridge (200 to 1500 m a. s .1.) are part of the NW-SE striking Deep Freeze Range of the Transantarctic Mountains in northern Victoria Land (Fig. 3.6). The Deep Freeze Range is flanked by the Campbell and Priestley Glaciers and is partially covered by local mountain glaciers. Both the Campbell and Priestley Glaciers discharge into Terra Nova Bay. Campbell Glacier is exclusively fed by local mountain glaciers and ice fields, whereas Priestley Glacier drains the East Antarctic Ice Sheet (EAIS). Mt. Keinath is located at the confluence of two tributaries of Priestley Glacier, Boomerang and Browning Pass glaciers. Black Ridge is situated where Corner Glacier flows into Priestley Glacier (Fig. 3.6). Blue ice zones indicating net ablation are abundant on these glaciers and testify to the strong katabatic winds in the area. These winds are responsible for the widespread ice-free zones in Terra Nova Bay region. In an inventory of the geomorphic elements of the area, Orombelli et al. (1991) dis¬ tinguished four major landscapes:

1. deep glacial troughs cutting backwards inland from the coast;

2. surfaces of alpine topography consisting of well developed ridges, spires, spurs and

cirques, in many places above a clearly visible trimline; this topography extends from the deep troughs into the intervening mountain blocks;

3. relict summit mesas, often carrying thin ice caps, scalloped by the fringing alpine topography;

4. coastal piedmonts, also relict features of an earlier landscape.

Well-defined trimlines are visible on the highest peaks of the area and on the valley walls above Priestley and Campbell glaciers and their tributaries. Trimlines mark the upper limit of ice cover. Their elevations imply that ice thickening was minimal in the upper reaches of outlet glaciers, but considerable (more than 1000 m) in the coastal areas (Orombelli et ah, 1991). Mt. Keinath and Black Ridge are situated on the edge of deep glacial troughs, below the alpine topography.

Quaternary deposits

Baroni and Orombelli (1989) divide the Cenozoic glacial sediments in northern Victoria Land into three units: a young drift, an older drift and an oldest drift. Later, Orombelli et al. (1991) did not further differentiate the older and oldest drift as the geomorpho¬ logical evidence is ambiguous. The young drift is a massive, matrix-supported diamict in thin sheets. Moraines exist only in very few places, one of which is Black Ridge.

53 3 Antarctica

Figure 3.6: Map of a part of northern Victoria Land, showing the sampling sites Black Ridge and Mount Keinath. Priestley Glacier drains Priestley Nevé, a part of the East Antarctic Ice Sheet, while Campbell Glacier originates in the Southern Cross Mountains. Grey shading indicates ice-free areas.

54 3.4 Mount Keinath

Subangular clasts and perched boulders are common features of this drift and indicate that it is at least partly composed by supraglacial and englacial debris. Its weathering degree is moderate. Reworked marine fauna is widespread in the matrix of the young drift deposits along the coast of the Northern Foothills (Fig. 3.6). The young drift was correlated to Ross Sea I glaciation (Baroni and Orombelli, 1989) and Late Wisconsin (Orombelli et ah, 1991). It was called younger drift by Stuiver et al. (1981) and Terra Nova Drift by Orombelli et al. (1991).

The older drift is a thin and discontinuous sheet of strongly weathered glacial sed¬ iments, lacking constructional forms. In places it occurs as scattered large boulders resting directly on bedrock or regolith. It is present at higher elevations than the young drift.

Above all observed glacial drifts and below a distinct erosional trimline, glacially rounded bedrock topography is present, showing deep weathering on the scale of metres. Here, glacial sediment or erratics are absent.

Available dating

More than hundred radiocarbon dates have been determined for marine shells, penguin and seal remains on raised beaches and marine fossils in till deposits along the coast (Baroni and Orombelli (1989), Baroni and Orombelli (1991), Baroni and Orombelli (1994), Orombelli et al. (1991)). These dates provide age brackets for the young drift in the Northern Foothills (between 25 and 8 ka) and assign a Holocene age to the raised beaches. Apart from that, there are few absolute ages from northern Victoria Land. Van Der Wateren et al. (1999) determined surface exposure ages on glacial erosion surfaces in northern and southern Victoria Land of up to 11.2 Ma. K-Ar and Rb-Sr dating of volcanic rocks from northern Victoria Land suggest that the temperate glacial conditions under which the serrated ridges of the alpine topography were cut into a summit mesa landscape ended about 7.5 to 8 MaBP and cold polar glacier conditions have prevailed ever since (Armienti and Baroni, 1999).

3.4.3 Sample description

A detailed geomorphological and glacial geological survey has been carried out in the Terra Nova Bay region (Baroni, 1996). Aerial photographs were analysed, and the geo¬ morphology was mapped during numerous field traverses. Thus, the spatial relationship of different units of glacial sediments as well as the distribution of glacial erosional features are well defined in the investigated area.

Rock samples for surface exposure dating were collected at two sites, namely Mt. Keinath (Fig. 3.7a; coordinates from USGS map sheet Mt. Melbourne: 74°32'10"S, 163°57'12"E) and Black Ridge (Fig. 3.7b; 74°24'11"S, 163°36'39"E), in the Southern part of the Deep Freeze Range during the 1997/1998 Italian Antarctic expedition. Wherever

55 3 Antarctica

possible, samples were chosen so as to exclude post-depositional overturning of the block

- in the case of erratics - and to minimize temporary snow cover or erosion.

Mt. Keinath samples

The Mt. Keinath samples originate from three different geomorphic features (I to III in Fig. 3.7a):

1. Rounded bedrock (I in Fig. 3.7a) from below the trimline, in the summit area from 1090 m a. s.l. down to about 800 m a. s.l.: The coarse granite bedrock on top of Mt. Keinath is well rounded and completely lacks glacigenic sediments. In places, it shows polished and hardened surfaces that are strongly stained (Munsell Colour 5YR4/6), a weathering rind of several centimetres thickness and a discontinuous patina. In contrast, the slopes beneath the summit are deeply sculpted by ex¬ tensive cavernous weathering, pseudokarren features and weathering pits. Surface

exfoliation is common and in places concentrated. Samples were taken from pro¬ truding parts of polished, hardened and stained flat bedrock surfaces. The sampled surfaces were not directly affected by cavernous weathering. Exfoliation is seen in the pits and between the protruding cores. From the exposed topography of the summit and the prevailing strong winds, it can be assumed that seasonal snow cover has been negligible.

2. Erratics field (II in Fig. 3.7a): Erratic blocks occur scattered between 850 and 580 m a. s.l. on the southeast ridge. The abundance of blocks continuously in¬ creases with decreasing elevation. The subangular to subrounded boulders are up to 1.5 m in diameter and show strong weathering with yellowish red or red staining (Munsell Colour 5YR4/6 to 2.5YR4/6). Forms of cavernous weathering with ta- foni, pseudo-karren and pitholes are widespread, particularly on the coarse-grained granites, both on the bedrock and on the overlying erratics.

3. Continuous drift (III in Fig. 3.7a), occurring as a continuous thin deposit on a nearly flat bench at 580 m a. s .1.: This is a diamicton that is moderately weathered and in places matrix-supported. Granitic boulders are subangular to subrounded and show a brown to reddish staining (Munsell Colour 7.5YR), but are not affected by deep cavernous weathering. The topography is locally hummocky and ice-cored with diffused frost cracks and evidence of ice wedge polygons.

The modern surface of Boomerang Glacier at the foot of Mt. Keinath is approximately 330 m below the continuous drift, at an elevation of 250 m a. s .1. (Fig. 3.7a).

Black Ridge samples

On Black Ridge, samples were collected from two surfaces (II and III in Fig. 3.7b):

56 3.4 Mount Keinath

Figure 3.7: a. The sampling locations at Mt. Keinath with Boomerang Glacier in the foreground and on the left, seen from Southeast. Black crosses rep¬ resent sampling sites. The upper limits of the sampled landscape units

are indicated by dashed lines. In the lower part of the slope, two more,

presumably younger moraines featuring extensive frost polygons are vis¬ ible, b. The moraine (arrows) on Black Ridge, formed by the continuous drift (III). It displays a nearly horizontal crest and rests directly on the erratics field (II) which covers the area above the moraine. White crosses indicate sampling locations.

57 3 Antarctica

1. A discontinuous blocky glacial drift of prevalently granitic boulders (II). Fine¬ grained boulders are deeply weathered and oxidized, generally showing the same

characteristics as those from the erratics field at Mt. Keinath. A few boulders are

as much as 3 m across, but most are smaller than 1.5 m. While the lower boundary of this drift is clearly marked by an onlapping moraine (see below), its upper limit is unclear as the boulders thin out uphill.

2. A moraine ridge at 950 m a. s .1. with prevalent granite and metamorphic boulders (III). It rests directly on the drift described above. The degree of weathering is much lower than in the erratics field (II) and similar to that observed on the continuous drift at Mt. Keinath. Sandy matrix, striated pebbles and cobbles and perched boulders are present. The moraine can be followed for several hundred metres.

The present surface of Priestley Glacier beneath the sampling sites is 750 m lower at an elevation of 200 m a. s .1. The erratics field (II) is generally thin and discontinuous at both sampling sites, and the upper limit is poorly defined. Thus, the erratics field cannot be clearly assigned to a single glaciation, but could represent several ice advances, each leaving a small number of erratics or a thin sheet of glacial drift. It was referred to as older drift in Orombelli et al. (1991). In the nearby Northern Foothills (Fig. 3.6), soil has developed on this drift. It shows an incipient differentiation with a thin surficial, dark yellowish-brown (10YR3/3 to 4/4) horizon, slightly more oxidized than the underlying sediment. In the same area - and in particular on Mt. Browning, 5 km southeast from Mt. Keinath - this drift covers residual patches of a strongly developed paleosol with well defined horizons, abundant salt concentration and clay skins (B2t, B3sa, 10R4/6 in colour). This paleosol is comparable to the oldest Antarctic soils (weathering stages 5 and 6) as described by Campbell and Claridge (1987), who attributed them to pre-Pleistocene times. In terms of the degree of weathering, lithology and geomorphological position, the moraine on Black Ridge (III in Fig. 3.7b) is very similar to the continuous drift on Mt. Keinath (III in Fig. 3.7a), although it is found at a much higher elevation than the latter. The moderate degree of weathering, the presence of matrix and perched boulders are also characteristics of the young drift on the Northern Foothills (Baroni and Orombelli, 1989). There, its age is known from 14C dates on marine shells and penguin remains to lie between 8 and 25 14C kaBP (Orombelli et al. (1991), Baroni and Orombelli (1991), Baroni and Orombelli (1994), and unpublished data), a time span which incorporates Marine Isotope Stage 2.

3.4.4 Experimental

The surface exposure ages of nine erratic boulders were determined using in-situ pro¬ duced cosmogenic 3He, 10Be and 21Ne. In samples K3, B2, and B3 (Table 3.6), both

58 3.4 Mount Keinath

10Be and 21Ne were measured. Quartz was separated from eight coarse-grained magmatic rocks from the Ordovician Harbour Granite Intrusions complex (Ghezzo et ah, 1987). High-purity quartz separates for the noble gas analysis were obtained by magnetic and density separation, followed by hand-picking. For 10Be analysis, the crushed rocks were leached in 5 % hydrofluoric acid following the procedure given by Kohl and Nishiizumi (1992). Cosmogenic 3He was measured in olivine from a basalt boulder, separated by magnetic and density separation and hand-picking. Noble gas analyses were carried out at the Department of Earth Sciences at ETH Zürich, using a non-commercial, all-metal magnetic sector mass-spectrometer (90°, 210 mm radius) equipped with a modified Baur-Signer ion source whose sensitivity is es¬ sentially constant over the pressure range relevant for this work (Baur, 1999). This ion source features a compressor device that increases the sensititvity compared to a con¬ ventional source by a factor of 120 for 3He and 200 for 21Ne (Baur, 1999). A stepwise heating procedure was applied in order to separate the cosmogenic gas fraction from nucleogenic or trapped gases (Niedermann et al. (1993), Schäfer et al. (1999)). It was assumed that all 3He is cosmogenic. To calculate the concentrations of cosmogenic 21Ne, 20Ne was assumed to be entirely atmospheric, and an atmospheric 21Ne/20Ne ratio of 0.002959 was used (Eberhardt et ah, 1965). 10Be was measured on the PSI/ETH Zürich Tandem AMS facility. First analyses of the olivine sample (K6) were carried out on the grain size fraction of <315 //mm. Neon data from three temperature steps plotted on the atmospheric-MORB mixture line with 22Ne/20Ne ratios below 0.095. Offline crushing to <25 /iin removed most of the MORB neon. However, the cosmogenic excess over air was still too small and did not allow to calculate a neon age. The helium age for this sample was calculated from the data of the <25 /iin fraction.

3.4.5 Results

Nuclide concentrations are given in Table I, neon isotopic data in Fig. 3.8, and the cal¬ culated minimum exposure ages in Table II and Fig. 3.9.

59 CA (Sample Table 3.6: Isotope data for all samples 1. names are short codes. The originalsample names are listed in the appendix.)

Sample Grain size T 3He 3He/4He 20Ne 21Ne 21Ne/20Ne 22Ne/20Ne 10Be (mm) (°C) (106at g~l) (io-6) (109at g"1) (106at g"1) (IO"3) (105at g"1) Kl <315 TE 6±10 3.4±5.7 11.84±0.23 180±16 18.21±0.90 0.1175±0.0033 - K2 <100 600 6±8 0.7±0.1 8.30±0.45 148±10 20.75±0.86 0.1148±0.0025 - 1750 - - 11.03±0.60 16±5 4.44±0.40 0.1027±0.0033 - K3 <100 600 11±1 16.2±1.4 19.61il.04 167±10 11.48±0.24 0.1090±0.0013 225.5±12.2 1750 - - 5.06±0.27 1±2 3.08±0.45 0.1022±0.0021 K4 <100 600 5±1 1.8±0.4 28.31il.50 13±1 3.41±0.05 0.1020±0.0010 - 1750 - - 22.88il.22 9±1 3.34±0.08 0.1024±0.0014 - Bl-A <100 600 - - 7.63±0.41 8±3 3.99±0.33 0.1027±0.0015 - 1750 - - 3.94±0.21 20±2 8.09±0.41 0.1027±0.0013 - Bl-B <100 600 1.5±0.3 0.1Ü0.02 14.56±0.77 11±2 3.71±0.16 0.1019±0.0010 - 1750 - - 46.18±2.45 12±3 3.23±0.06 0.0982±0.0006 - B2 <100 600 - - 2.92±0.16 20±2 9.75±0.62 0.1079±0.0021 48.7±2.6 1750 - - 49.42±2.62 44±6 3.85±0.11 0.0978±0.0007 K5 <100 ------4.0±0.3 K6 <25 800 7.3±0.4 226.0±150.6 19.85±19.38 3±8 3.13±0.24 0.1013±0.0024 - 1400 0.5±0.1 61.8±174.4 18.30±17.87 5±8 3.22±0.12 0.1014±0.0017 - 1800 - - 11.91±11.64 2±5 3.13±0.22 0.0952±0.0135 - K6 <315 400 0.3±1.2 0.4±1.3 2.4±0.2 0.1±1.6 3.02±0.5 0.0962±0.0026 - 800 6.2±1.0 1.3±0.2 1.6±0.1 0.8±1.9 3.5±1.1 0.0947±0.0070 - 1400 2.1±0.6 1.3±0.4 1.0±0.1 0.2±1.6 3.2±1.5 0.0902±0.0052 - 1800 0.9±0.7 6.5±5.2 1.7±0.1 2.8±2.0 4.6±1.1 0.0904±0.0044 - B3 <100 600 - - 8.70±0.47 49±6 8.63±0.62 0.1033±0.0023 1.9±0.2 1750 - - 9.80±0.53 3926±210 403.64±5.88 0.2275±0.0056

1 Noble gas data errors are 2

amounts are not included,but are thought to be <3%. 10Be data errors are Is,includingthe statistical error due to the normalisation to standards and blanks, and a sample processing reproducibilityerrors of 5%. Degassing times were: 600°Cand 800 °C, 45 minutes; 1400 °C,

1750°Cand 1800°C, 15 minutes (TE = total extraction in a singlestep at 1750 °C). Phe low temperature steps (600°Cforquartz, 800°Cfor olivine) were used for age calculation. Table 3.7: Exposure ages of all samples.1

Type of Site Sample Sample Lithology Mineral Elevation 3 He iuBe 2iNe surface type (ma s 1 ) minimum minimum minimum age (ka) age (ka) age (ka)

I Kei Kl b granite quartz 1090 - - 2560±230

rounded Kei K2 b granite quartz 1090 - - (2090±150)

bedrock Kei K3 b granite quartz 1090 - 1750Ü30 2370±140

II Kei K4 e monzogranite quartz 660 - - 266±26

erratics BR Bl2 e granite quartz 980 - - 147±33

field BR B2 e granite quartz 980 - 305±16 309±32

III Kei K5 e granite quartz 560 - 33 9±2 8 -

continuous Kei K63 e basalt olivine 580 28 4±1 3 - -

drift BR B3 e granite quartz 950 - 11 2±0 9 (790±96)

1Sea level, high latitude production rates of P3 = 115at g_1 ar1 (Cerling and Craig, 1994b), Pio = 5 44atg_1 yr-1 (Kubik et al (1998), re-evaluated after Heismger et al (2002b) and Heismger et al (2002a)) and P2i=20at g_1 ar1 (Niedermann, 2000) were scaled to latitude and elevation according to the procedure given in Stone (2000) For Pio, nucléons, stopped negative muons and fast muons were scaled separately according to Heismger et al (2002b) and Heismger et al (2002a) Ages in brackets were not taken into account for interpretation Age calculations are based on attenuation

P = 150 cm~2 et al of = 2 7 cm~3 sites are length g (Brown , 1992), density quartz p g Sampling Mt Keinath (Kei) and Black Ridge (BR) Sample types are solid bedrock (b) and erratic block (e) 2Average of two aliquots Bl-A and Bl-B

3Gram size <25mm

Some data points for quartz samples plot below the atmospheric-cosmogenic mixing line in the neon three-isotope plot (Fig. 3.8). This may indicate a trapped neon compo¬ nent enriched in nucleogenic 21Ne (Hetzel et ah, 2002). In order to test the reliability of our neon data, the critical samples B2, B3and K3 were cross-checked by 10Be analysis. The beryllium age of B3 differs considerably from its neon age (Table II), and a sig¬ nificant non-cosmogenic neon contribution must be assumed. Assuming the production rates are accurate, the 10Be concentrations of sample B2are consistent with the neon isotope data. They do not indicate a substantial nucleogenic neon component, neither in-situ produced nor as the result of a non-atmospheric composition of trapped neon. The neon data plot within error on the mixing line (Fig. 3.8). For sample K3, the beryl¬ lium age is 26 % lower, and a minor nucleogenic contribution appears possible. This sample is slightly offset from the mixing line.

We may estimate the influence of such a non-atmospheric trapped neon component on the age calculations. The three data points B2, K3and Kl appear to be consistent with a trapped neon ratio (21Ne/20Ne)tr of 5T0~3 (compared to 2.959T0-3 for atmospheric neon). Though somewhat arbitrary, this value can be deduced in Fig. 3.8 from a fit line through these three data points parallel to the atmospheric-cosmogenic mixing line. Adopting this (21Ne/20Ne)tr value leads to 21Ne ages of 2.2 Ma, 1.8 Ma and 216 ka for the samples Kl, K3and B2, respectively. These values are about 25% lower than those calculated above with trapped neon of atmospheric composition, a difference that would not substantially alter any conclusion of this paper. However, this estimation of a non- atmospheric trapped component is not quantitative, and the interpretation must be

61 0.600

0.500

0.400 - y -

- 0.300 -

' 0.200 - Fig. 3b)/

0.100 - * t , , ,-

0.000 0.100 0.200 0.300 0.400 0.130

0.005 0.010 0.015 0.020

2iNe/20Ne

3.8: Three-isotope plots of all neon data from quartz (all temperature steps). The dashed line in all panels is the atmospheric-cosmogenic mixing line

(y = (1.120 ± 0.021)a; + 0.098; Niedermann et al. (1993)). Data that plot on this line display a two-component mixture of atmospheric and cosmogenic neon. A position significantly below this line points towards an inherited component, a. all neon data points; b. enlargement of the

section that contains the neon data points used for age calculation. 3.4 Mount Keinath

based on the assumption of atmospheric trapped neon. In the case of K3, the putative non-atmospheric trapped neon ratio would explain the divergence between the neon and beryllium minimum ages as the calculated neon age of 1.8 Ma matches the beryllium age.

Identical 10Be and 21Ne minimum ages of an old sample, on the other hand, would imply zero erosion, which is unlikely even in the hyperarid Antarctic climate. For all other samples apart from Kl, K3and B2, we consider such a non-atmospheric trapped neon composition unlikely as they have the same or very similar lithologies and plot on the atmospheric-cosmogenic mixing line (Fig. 3.8), and the ages calculated with atmospheric trapped neon composition are adopted. The sample K2 and the neon age of B3 are not considered in the following discussion because of the possible nucleogenic imprint. Additional age constraint for the continuous drift (III) comes from the helium analysis of the olivine sample K6 (Table II).

All 3He is assumed to be cosmogenic for the calculation of the helium age of sample K6. This assumption is supported by three points. First, the offline crushing step removed the MORB signal in neon (see data in Table I). It therefore can be assumed that any MORB helium fraction was also released from the sample by crushing. Second, the 3He/4He ratio of the presented helium data is very high. A correction for an assumed mantle component of 9 times the atmospheric ratio (Lupton, 1983) would reduce the 3He concentration by only 5 %. Third, another sample from the continuous drift dated with 10Be (K5) even has a slightly higher age than the helium dated K6, which would not be expected in the case of a significant non-cosmogenic helium contribution in K6 as 10Be concentrations are not affected by MORB contribution. Thus, a substantial mantle- derived helium fraction seems unlikely.

The difference between the 10Be and 21Ne age of sample K3 seems to point toward pre¬ exposure of that sample. This is very unlikely, however, as the samples from the summit of Mt. Keinath were taken from solid bedrock and not from erratics. Furthermore, if very old samples have experienced erosion, their minimum ages without erosion correction, are not expected to be the same. Indeed, the 10Be and 21Ne ages of K3 agree well when corrected for erosion (see below). For younger sample B2 from the erratics field, pre¬ exposure can be ruled out as the radionuclide and noble gas ages are almost identical.

Note that all surface exposure ages were calculated neglecting any uplift of the sam¬ pling sites during the period of exposure as uplift rates are very difficult to assess. However, if a sample is uplifted during its exposure, its measured age is too low due to the lower production rates at low elevations. Thus, correction for uplift would increase the exposure ages.

3.4.6 Discussion

Minimum exposure ages

The minimum exposure ages cluster in three clearly distinct groups (Fig. 3.9): Two samples are older than 1.8 Ma, three have an intermediate age between 147 and 309 ka and

63 3 Antarctica

three samples have low ages between Hand 34ka. This age pattern is consistent with the relative glacial stratigraphy: The old ages represent the strongly weathered summit surface of Mt. Keinath (I in Fig. 3.7a), the intermediate ages come from the erratics fields referred to as older drift (II in Figs. 3.7aand 3.7b), and the lowest ages were determined on the lower elevation continuous drift representing the young drift of Baroni and Orombelli (1989) (III in Figs3.7aand 3.7b). These three units represent at least three, but probably more, glacial events:

1. A round-topped and strongly weathered bedrock landscape in the higher parts

of the mountain range below the erosional trimline has been shaped prior to

2.6 Ma ago and has remained ice-free ever since. Because surface exposure ages

are minimum ages, the oldest age measured on a single surface is assumed to be least biased by erosion or temporal shielding and is therefore adopted as its

minimum age.

2. A glacial event occurred before 309 ka BP and deposited the boulders of the erratics fields (older drift) in the lower reaches of Priestley Glacier and below 850 m a. s .1. around Mt. Keinath. These deposits are situated up to 780 and 600 m above the modern glacier surface, respectively. The depositing glaciers have occupied the present-day network of glacial valleys, but did not cover the mountain ranges. A

mid-Pleistocene age of the erratics field is consistent with the occurrence of the presumably pre-Quaternary paleosol under the erratics field on the slope of Mt. Browning (see above).

An increase in weathering degree and changes in the clast lithology with increasing elevation as observed in the older drift by Baroni and Orombelli (1989) suggest that the erratics fields represent multiple glaciations. With the data obtained in this study, this notion seems to be confirmed as the three samples yielded three different ages. However, the low age of sample Bl has to be regarded with caution. It was collected close to sample B2, which is of the same lithology, and shows a weathering stage so similar that deposition by the same ice body is highly probable. The age difference to B2 is most probably due to post-depositional overturning or temporary shielding. The same considerations apply to sample K4. Thus, there is no clear-cut evidence from our data that the boulders sampled from the erratics field were deposited by more than one glacial event in mid-Pleistocene.

3. The continuous drift at Mt. Keinath developed before 33.9 kaBP, 380 m above the present ice level. This is the first piece of evidence of a late Pleistocene ice advance older than the Last Glacial Maximum in this area. The chronological constraint for the younger drift in the coastal area supplied by shells and penguin remains yields an age of 8 to 25ka (Orombelli et ah, 1991). It is inferred that the young drift, as the older drift, was possibly formed by more than one glacial phase in the

Terra Nova Bay area. But this can not be resolved by the exposure dates presented

64 3.4 Mount Keinath

O Beryllium ages A Neon ages O Helium ages 5000 r- -, 5000 2 S g. ' s 1 Rounded bedrock 2000 2K 2000

1000 1000

CM m 500 500 2

' H S II Erratics field m 0 200 200 ï 100 100

50 2 50

IE

III Continuous drift. 20 CO 20 m

10 0- 10 Mount Keinath Black R dge samples samples

ure 3.9: Age distribution of all samples. The three groups of exposure ages are in correspondence with the three landscape units. Note that the ordinate scale is logarithmic. K2 is not included in the discussion because of a possible nucleogenic imprint (see text).

65 3 Antarctica

here, and differentiation of this drift based on geomorphological evidence and the weathering degree was not possible.

The exposure age of the continuous drift is much lower at Black Ridge than it is at Mt. Keinath. The geomorphological and sedimentological characteristics and the weathering degree are very similar at both sites, though, and exposure ages

are minimum ages. Thus, it seems that the continuous drifts at Mt. Keinath and Black Ridge can be correlated to the Northern Foothills glacial drift which is

attributed to the Last Glacial Maximum based on Radiocarbon dates. The low

apparent age of 11.2 ka at Black Ridge must be explained by a temporal shielding or post-depositional overturning of the sampled block.

Erosion rates and erosion-rate corrected exposure ages

Fig. 3.10 displays the evolution of 21Ne/10Be with 10Be for different erosion rates. The positions of the samples K3 from the summit of Mt. Keinath and B2 from the erratics field at Black Ridge indicate that both have experienced a continuous history of exposure without burial or temporal shielding (e.g. by seasonal snow-cover) (Graf et ah, 1991). The resulting erosion rate for Mt. Keinath is 20±3cmMa_1. Sample B2 plots on the left-hand side of the erosion island where no erosion rates can be inferred.

The erosion rate of 20 cm Ma-1 for the bedrock on the summit of Mt. Keinath is some¬ what higher than those reported previously for sandstone bedrock or dolerite clasts from the Dry Valleys (e.g.Ivy-Ochs et al. (1995), Bruno et al. (1997), Schäfer et al. (1999)), but well within the range of those determined by Ivy-Ochs et al. (1995) for granite clasts from the Dry Valleys and Nishiizumi et al. (1991) for quartz grains in sandstone from the S0r Rondane Mountains. Considering the lower error limit of 17 cm Ma-1 shifts the neon exposure age for sample K3 from 2.4 to 4.2 Ma (Fig. 3.11). Assuming that the same sam¬ ple was eroded by the more probable rate of 20 cm Ma-1, the corrected exposure time is 5.3 Ma, and the upper limit of 23 cmMa"1 yields an age of 9.7Ma (Fig. 3.11). Correcting the 10Be age for erosion at 17and 20cmMa_1 leads to ages of 3.3and 5.2Ma, respec¬ tively, i.e., the beryllium and the neon age agree well. At 23 cm Ma-1, beryllium is in saturation. These erosion-corrected ages are in agreement with the finding of Armienti and Baroni (1999) that no wet-based glacier has covered this area since 7.5MaBP.

3.4.7 Implications

10Be and 21Ne surface exposure ages obtained from glacially rounded bedrock on the summit of Mt. Keinath show that the mountains in Deep Freeze Range, northern Victoria Land have been free of ice cover since 5.3 Ma. This implies that the Transantarctic Mountains have not been overridden by erosive ice after mid-Pliocene. Since then, erosion has been very low, with denudation rates of 17 to 23 cm Ma-1, allowing the preservation of relict surfaces.

66 3.4 Mount Keinath

Q I I I I I I I I 4x106 8x106 12x10e 16x106 10Be[atg-1]

Figure 3.10: The effect of erosion on the relative concentrations of a stable and a radioactive nuclide for different erosion rates (solid lines). This kind of plot allows the determination of long-term erosion rates of the sampled surfaces (Graf et al., 1991). The erosion rate curves end where neon reaches erosion saturation, i.e. cosmogenic 21Ne does not accumulate

any more due to erosion. The dashed line connects these end points. Samples plotting outside the area between the curve for zero erosion

and the dashed line are likely to have a complex exposure history. In

the case of exposure prior to the present exposure period, data would plot above this area because of the decay of lOBe and the increase of the 21Ne/10Be ratio. Shown are K3 (I, rounded bedrock at Mt. Keinath) and B2 (III, erratics field at Black Ridge). The resulting erosion rate for K3 is 20±3cmMa_1. For B2, no erosion rate can be read from

the plot as the data point falls into the range where all erosion curves

converge. For parameters used for calculation see caption of Table II.

67 3 Antarctica

16 1 I ' I ' I

CO

CO o - - Q. 8 T3

tS

&_ (3

i.i.i

0 5 10 15 20 25 Erosion rate (cm Ma-1)

Figure 3.11: Variation of erosion rate-corrected exposure time with erosion rate for the neon data of sample K3. The continuous straight lines indicate the erosion rates as determined from Fig. 3.10 with their errors. The dashed lines show the resulting exposure times: 5.3+3.4/-l.lMa. For parameters used for calculation see caption of Table II.

68 3.4 Mount Keinath

Erratics on valley walls yield much younger 3He, 10Be and 21Ne exposure ages than the summit bedrock on Mt. Keinath. Although the data discussed here do not permit to create a detailed glacial chronology, apparently glacial advances were of limited extent throughout the Pleistocene. The ice did not reach the mountain crests in the local glacier system or along the EAIS outlet streams. Glaciers were more extensive than today in mid-Pleistocene at least once before 309kaBP. Another ice advance occurred before 34kaBP. During the earlier of these phases, the ice reached a level of 780 m and more above the present-day ice surface around Black Ridge and about 600 m above the present-day ice surface at Mt. Keinath.

3.4.8 Conclusions

Surface exposure dating and detailed field mapping were applied to construct a prelimi¬ nary glacial history of Terra Nova bay area and contribute to reconstructing Pleistocene climate fluctuations in Antarctica. This study presents one of the very first combined applications of noble gas and radionuclide isotopes for the calculation of erosion rates in Antarctica. The results demonstrate how critical even small erosion rates are when calculating high exposure ages. The preservation of mid-Pliocene granite bedrock surfaces in the mountains of Terra Nova Bay area and the very low denudation rates strongly point towards a permanently cold and hyperarid climate in this part of the Transantarctic Mountains since the forma¬ tion of these surfaces. This finding is consistent with earlier studies on the late Tertiary climate from the Dry Valleys and from the northern Victoria Land glacial system. It does not support the hypothesis that the landscape stability of the Dry Valleys' crustal block throughout the Pliocene is a unique feature due to a particular tectonic history. Independent from the history of uplift, there is strong evidence that the summit area of Mt. Keinath has been sculpted by wet-based ice before 5.3MaBP. This surface has not been glacially modified any more after that time.

According to the exposure age data presented here, the ice surface lowered in the Terra Nova Bay area mountains before the mid-Pliocene. This is a minimum age for the erosional trimline and consistent with K-Ar and Rb-Sr dating of volcanic rocks presented by Armienti and Baroni (1999) who suggested the trimline was older than 7.5 Ma. Subsequent glacial events reached elevations hundreds of metres above modern ice levels, but were restricted to the valleys. Such advances took place at least twice throughout the Pleistocene.

The evidence presented in this study is among the first to indicate Pleistocene glacial fluctuations and thus for the amplitude of climate oscillations throughout the Quater¬ nary outside the Dry Valleys. However, to improve our knowledge on the timing and amplitude of the Last Glacial Maximum in Antarctica, more data from selected sites are required. The analysis of samples from various places across Northern Victoria Land is under way.

69 3 Antarctica

Acknowledgements

We thank Walter Wittwer for separating the minerals. This work was supported by the Italian Program on Antarctic Research (PNRA) and the Swiss National Science Foundation. The manuscript was greatly improved thanks to the reviews by Robert Ackert, Mike Bentley, Kevin Hall and David Sugden.

70 3.5 Erosional surfaces in Northern Victoria Land

3.5 Cosmogenic noble gases in Pyroxenes used for dating Late-Cenozoic erosional surfaces in Victoria Land, Antarctica 2

Peter Oberholzer "*, Carlo Baroni b, Maria Cristina Salvatore c, Heinrich Baur a and Rainer Wieler a

a Institute of Isotope Geology and Mineral Resources, ETH Zürich, Sonneggstrasse 5, 8092 Zürich, Switzerland b Dipartimento di Scienze della Terra, Universita di Pisa, e Istituto di Geoscienze e Georisorse CNR, Via S. Maria 53, 56126 Pisa, Italy : Dipartimento di Scienze della Terra, Universita dette studi Roma "La Sapienza", P. le A. Moro 5, 00185 Roma, Italy * corresponding author: [email protected]

Abstract

We present 3He and 21Ne exposure ages of erosional glaciogenic rock surfaces on nunataks

in the Prince Albert Mountains and near Mesa Range, Northern Victoria Land, Antarc¬

tica. These nunataks are located at the margin of the polar plateau and therefore directly record ice volume changes of the East Antarctic Ice Sheet. These surfaces that overlook

the present ice surface by less than 200 m were last covered by ice at 3.5 and 2.7 Ma BP, respectively. This clearly indicates that the ice sheet has not been substantially larger

than today since the lower Pliocene. Erosion rates on these surface are lower than 16 cmMyr-1.

First absolute ages are reported for the Alpine Topography above the trimline that marks

the upper limit of glacial activity in Northern Victoria Land. With 1.6 Myr, they are

apparently younger than the erosional surfaces below the trimline. This is a result of the

high relief that strongly enhances erosional processes. The age of 3.5 Ma therefore is a

minimum age for the trimline itself, too.

With a large sample set we show that the degassing pattern of cosmogenic neon from py¬

roxene is a function of its Mg content. It is suggested that the Mg content is determined prior to isotope analysis.

3.5.1 Introduction

For a long time, the discussion about the history of the East Antarctic Ice Sheet (EAIS) has been dominated by the debate between "dynamists" and "stabilists". The dynamic ice sheet hypothesis proposes that the EAIS after its formation around 14 Ma ago was reduced several times to one-third to the half of its present size due to warmer climatic

! To be submitted to Journal of Quaternary Science

71 3 Antarctica

conditions, the last time in late Pliocene (Webb et ah, 1984; Webb and Harwood, 1991). The stabilists, on the other hand, advocate an ice sheet that did not vary much in size as the Antarctic climate is supposed to have been constantly cold and arid since at least 15Myr (Denton et ah, 1991; Marchant et ah, 1993b; Schäfer et ah, 2000). The controversy is not settled, and many questions remain open. Many key arguments of both opponents in this debate are based on evidence from the Dry Valleys of McMurdo, as they contain the largest ice-free areas of the continent with well-conserved glacial land forms. Van Der Wateren et al. (1999) argue, based on this fact, that landscape stability as inferred from evidence from the Dry Valleys can not be generalised for the entire Transantarctic Mountains (TAM). Therefore, to improve our knowledge of ice volume fluctuations of the EAIS - in both a temporal and a spatial sense - areas outside the Dry Valleys need to be investigated. Ice-free terrain indeed is abundant also north of the Dry Valleys along the Transantarc¬ tic Mountains, occurring in mountain ranges and as nunataks at the edge of the vast inland ice. Some of these nunataks protrude from the present-day ice surface by less than 200 m. Their vicinity to the EAIS makes the nunataks at the western, or poleward, side of the TAM excellent archives for the history of the ice volume. Glacial land forms on these nunataks are a direct record of ice elevation variations, not influenced by the damming effect of narrow valley sections or a grounded Ross Ice Shelf, as is the case for the large outlet glaciers that drain the ice sheet through the TAM. Furthermore, reconstructions of the inland ice level based on evidence from the periphery of the ice sheet do not depend upon backward extrapolations of former glacier profiles (Marchant et ah, 1994; Denton and Hughes, 2000).

In an attempt to obtain absolute ages for Pliocene and Pleistocene variations of the

EAIS ice level from the internal TAM ranges, we investigated glacial erosional surfaces in two areas of NVL: in the Prince Albert Mountains south of David Glacier and in the Mesa Range south of Rennick Glacier (Fig. 3.12). Only very few authors so far have methodically addressed the use of the stepwise heating technique for separating cosmogenic and non-cosmogenic noble gas fractions in pyroxene, providing data from less than a dozen samples. As a second focus of this study, new information about the degassing pattern of cosmogenic neon from pyroxene in dependence of the chemical composition of the mineral is supplied.

3.5.2 Glacial geological and geomorphological setting

The morphological units of Northern Victoria Land are described in detail in Orombelli et al. (1991); Baroni and Orombelli (1989); Orombelli (1989); Denton et al. (1986). Particular features are glacial trimlines that mark the upper limit of the erosional activity of ice, and the so-called Alpine Topography above the uppermost trimlines, untouched by the ice. It is assumed that the development of the Alpine Topography began with the onset of glaciation, through undercutting of rock walls. We have dated the most extensive glaciation in Northern Victoria Land by dating the

72 3.5 Erosional surfaces in Northern Victoria Land

Chisholm Hills

w- -s* ïm,2\i *? /> ^ JS-

_-^

200 km

Figure 3.12: Northern Victoria Land with the sampling areas Chisholm Hills, Mount Bowen and Ford Peak. Note that North is toward the bottom of the figure.

73 3 Antarctica

uppermost trimline. This was attempted in two different ways: First by dating surfaces of the Alpine Topography at Ford Peak and Chisholm Hills; second, by dating erosional surfaces below the trimline at Chisholm Hills and Mount Bowen, which are expected to yield a minimum age for the trimline. Samples were taken in locations where a trimline is directly related to the EAIS, but trimlines are also created by local ice caps and mountain glaciers.

Alpine Topography

The landform that Orombelli et al. (1991) referred to as Alpine Topography is charac¬ terised by steep ridges and spurs, serrated crests, horns and cirques. It occurs both in granite and dolerite surfaces and is widespread in NVL. The rock surface displays an oxidation crust or weathering rind and is often affected by cavernous weathering. Its lower limit is in places obvious as a distinct trimline that marks the upper limit of direct glacial action. In other places, the Alpine Topography reaches down to the present-day ice level. Orombelli et al. (1991) give a detailed description of the distribution of the trimline between the Mesa Range and David Glacier.

Erosional surfaces below the trimline

Beneath the trimline, sedimentary and erosional landscape elements of glacial origin are abundant in Northern Victoria Land. Seven samples were taken from erosional surfaces, four of them from a particular erosional feature that commonly are referred to as "Antarctic Tors" or "Gargoyles": rock columns that are sculpted into Ferrar Dolerite bedrock (Fig. 3.13). The columns are up to 1.2m high, but mostly from 60 to 80 cm, and have diameters from 10 to 40 cm, some becoming narrower at the base. They stand close to each other, at most a few metres apart, and the ground in-between is covered by sand to pebble-sized debris, toppled columns, or is bare bedrock. A common feature is excessive pithole-weathering on the top surface (diameters of weathering pits are from

1 to 5 cm) and a strong weathering rind with a - often shiny - oxidation crust on the top and lateral surfaces, resembling the surface properties of the Alpine Topography. Gargoyles occur in groups on flat terraces that are thought to be structurally controlled, i.e. to coincide with the original surface of the dolerite intrusion. It is assumed that the Gargoyles developed between intersecting cooling cracks, similar to basalt columns. The Gargoyles investigated in this study are all found at lower elevations than the highest trimline. Hence, from their mutual topographic position, it appears possible that the ice that has sculpted the trimline has covered the Gargoyle terraces, but there is no direct evidence for this. For example, at Mount Bowen, where the local ice flow patterns are complicated (see below), it is quite likely that the terrace with the Gargoyles is not covered by ice when the ice sheet surface is 50m higher than today, i.e. it reaches its maximum past elevation as marked by the trimline. The Gargoyles in their present-day condition are so delicate that it seems very unlikely that they could have survived ice

74 3.5 Erosional surfaces in Northern Victoria Land

Figure 3.13: Gargoyles or Antarctic Tors on Mount Bowen. The column in the picture is approximately 1 m high.

75 3 Antarctica

cover.

3.5.3 Sampling sites

David Glacier Area

Mount Bowen (75°45' S, 161°00' E) Mount Bowen is a nunatak 50 km inland from the Ross Sea coast, very near to the inland ice plateau. It is 6 km long and 3 km wide. As the ice flows around Mount Bowen in a generally northeasterly direction, its surface drops from 1400 to 800 m a. s.l. This abrupt topographic step locally increases the prevailing strong katabatic winds, which leads to high ablation and blue-ice surfaces in the lee of the nunatak and a locally reversed ice flow direction. The sampled Gargoyles are found in a field of several hundred square metres at 1525 m a. s.l., clearly below the oldest erosional trimline that appears a few kilometers away at 1750 m a. s .1.

In addition to the Gargoyles, we sampled a small plateau 50 m below the oldest trim- line, at 1700 m a. s.l. Flat dolerite tables of more than one square metre are covered with the same stain coat and desert varnish, but glacial striae are preserved as well. These striae indicate ice flow in northeasterly direction towards David Glacier. This is consistent with the flow pattern observed on nearby Ricker Hills by Van Der Wateren et al. (1999).

Ford Peak Ford Peak lies about 18 km northwest of Mount Bowen and has about the same size. The crests of the mountain display Alpine Topography with indented ridges and numerous spires. The narrow plains at the foot of Ford Peak are partially debris-covered and in places contain Gargoyles. The trimline does not crop out on Ford Peak.

We sampled the top of single dolerite spires on a high and exposed crest at 1800 m a. s .1.. These spires are rubified, oxidized and affected by pithole weathering.

Chisholm Hills (Mesa Range)

Chisholm Hills (73°30' S, 163°14' E) is a group of small nunataks at the southern end of Mesa Range, 90 km from the Ross Sea coast and over 260 km north of the two sampling sites described above. Like those, the Chisholm Hills are located at the inner side of the Transantarctic Mountains, roughly at the elevation of the local inland ice surface. On Chisholm Hills, samples were collected at two locations:

- At the foot of a steep, sharp and serrated ridge in deeply weathered dolerite rock, sculpted by Alpine Topography (2790 ma. s .1.). This ridge is entirely surrounded by ice. The trimline is not visible, but may be covered by accumulated snow at the foot of the ridge.

76 3.5 Erosional surfaces in Northern Victoria Land

- On a small flat dolerite bench entirely covered with Gargoyles (2700 m a. s .1.). This terrace is at a lower elevation than the serrated ridge, thus, the situation is comparable to that at Mount Bowen and Ford Peak, where the Gargoyles are found below the trimline and thus below the Alpine Topography. The Gargoyles at Chisholm Hills are in general shorter and narrower than those at Mount Bowen, reaching heights of 50 cm.

3.5.4 Methods

Nuclide measurements

Rock samples were crushed and sieved to <315/xm. Pyroxene separates were produced by magnetic and heavy liquid separation and by hand-picking. 20 to 40 mg of each

° sample were measured at ETH Zürich with a non-commercial 90 sector field noble gas mass spectrometer equipped with a Baur-Signer ion source (Beyerle et ah, 2000) and a compressor that increases the sensitivity by two orders of magnitude (Baur, 1999). Cosmogenic and non-cosmogenic noble gases were separated by stepwise heating. The temperature steps were chosen at 850 °C to extract cosmogenic 3He and nucleogenic 21Ne, at 1500 °C for cosmogenic 21Ne and at 1750 °C for total extraction, following Schäfer et al. (2000). The sample gas was cleaned with Zr-Ti and SAES non-evaporable getters and nitrogen-cooled cold traps. Neon was frozen to a cryogenic cold trap at 14 K to separate it from helium. Interferences of H2180 and 40Ar++ on mass 20 in the order of 0.5 % were corrected for.

Age calculation

All 3He was assumed to be cosmogenic. The cosmogenic neon excess (21Necos) was cal¬ culated as the excess over air. For 3He, a sea level, high latitude production rate of 115 at g_1 yr-1 (Cerling and Craig, 1994b) was adopted. For 21Ne, we used elemental production rates of 42.5atg_1 yr-1 (Si), 185atg_1 yr-1 (Mg), and 51atg_1yr_1 (Al) (Niedermann, 2000; Schäfer et ah, 2000). The relative concentrations of these elements were determined by ICP-OES and atomic absorption spectroscopy (Tab. 3.10). Produc¬ tion rates were scaled to the geomagnetic latitude and elevation of the sampling sites with the formalism suggested by Stone (2000) which compensates for the Antarctic air pressure anomaly. If a sample is not taken from a large horizontal plain, but e. g. at the periphery of the top surface of a cube-shaped boulder, or atop a hemisphere or pyramid, some of the incoming cosmic-ray neutrons escape from the rock back to the atmosphere (Masarik and Wieler, 2003), reducing the production rate. The distinct columnar shape of the Gargoyles is particularly affected by this. A correction in the range of 4 to 7% was applied according to Masarik and Wieler (2003).

77 3 Antarctica

Table 3.8: Helium isotope data for all samples. Errors are 2 a and include uncertain¬ ties in mass discrimination and sensitivity; uncertainties due to calibration

gas concentrations are not included, but expected to be smaller than 3%.

Sample T (°C), 3He 4He 3He/4He time (min) (106 at g"1) (1012 at g"1) (1CT6)

BOW1 Px 850 45 213.01 ± 5.50 241.07 ± 0.24 0.88 ± 0.02

1500 15 1.18 ± 0.41 0.51 ± 0.01 2.28 ± 0.80

BOW2 850 45 123.61 ± 3.77 247.05 ± 5.31 0.50 ± 0.02

1500 15 0.69 ± 0.15 0.60 ± 0.01 1.16 ± 0.25

BOW3 850 45 209.46 ± 5.70 203.24 ± 4.17 1.03 ± 0.04

1500 15 9.76 ± 0.55 5.08 ± 0.10 1.92 ± 0.11

BOW3-2 850 45 219.68 ± 5.21 198.90 ± 0.10 1.10 ± 0.03

1500 15 7.39 ± 1.18 4.26 ± 0.01 1.74 ± 0.28

BOW4 850 45 1646.95 ± 82.17 133.29 ± 0.07 12.36 ± 0.62

1500 15 20.66 ± 2.38 2.87 ± 0.00 7.19 ± 0.83

BOW6 850 45 1081.76 ± 56.38 190.74 ± 0.10 5.67 ± 0.30

1500 15 8.59 ± 1.02 3.34 ± 0.00 2.58 ± 0.31

CH2 850 45 87.71 ± 3.97 154.70 ± 0.12 0.57 ± 0.03

1500 15 0.60 ± 0.35 0.35 ± 0.01 1.69 ± 1.00

CH3 850 45 2.00 ± 0.68 10.57 ± 0.01 0.19 ± 0.06

1500 15 0.90 ± 0.39 1.88 ± 0.01 0.48 ± 0.21

CH4 850 45 3433.19 ± 317.61 164.63 ± 0.30 20.85 ± 1.93

1400 15 21.25 ± 0.97 0.82 ± 0.00 25.80 ± 1.18

1750 15 1.02 ± 0.18 0.28 ± 0.00 3.61 ± 0.62

CH5 850 45 2214.57 ± 198.01 145.00 ± 0.18 15.27 ± 1.37

1400 15 35.57 ± 1.70 2.33 ± 0.00 15.28 ± 0.73

FOR2 850 45 1159.17 ± 84.46 24.65 ± 0.02 47.02 ± 3.43

1400 15 28.01 ± 2.21 1.15 ± 0.00 24.44 ± 1.93

FOR4 850 45 819.45 ± 57.21 32.02 ± 0.02 25.59 ± 1.79

1500 15 28.78 ± 1.28 2.23 ± 0.00 12.88 ± 0.57

RHP1 850 45 115.67 ± 3.37 312.45 ± 6.61 0.37 ± 0.01

1500 15 3.07 ± 0.35 9.55 ± 0.20 0.32 ± 0.04

RHP2 850 45 40.48 ± 1.86 63.92 ± 1.76 0.63 ± 0.03

1500 15 1.32 ± 0.25 1.37 ± 0.04 0.96 ± 0.18

RHP9 850 45 88.04 ± 2.02 132.06 ± 1.58 0.67 ± 0.02

1500 15 11.64 ± 0.57 26.28 ± 0.31 0.44 ± 0.02 Ci 1- ce

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3.5.5 Results

Isotope data are shown in tables 3.8 and 3.9 and figures 3.14, 3.16 and 3.15. In BOW1, quartz could be separated along with pyroxene and as a cross-check was analyzed for cosmogenic 21Ne. The quartz neon age agrees with the pyroxene total neon age (see Fig. 3.16). However, the quartz data show a slight nucleogenic contribution in the rele¬ vant T steps, and the quartz neon age therefore is a maximum estimate.

Data for three moraine boulders from Ricker Hills nunatak west of Ford Peak are added for the investigation of the pyroxene neon degassing pattern (RHP samples). These samples will not be considered in the discussion of the exposure ages here, though (Oberholzer et ah, 2003a).

In the samples presented here 21Neea;c is released from pyroxene in significant amounts at 850 °C. This is in contrast to all previous studies that have applied stepwise heating to pyroxene (Schäfer et ah, 2000; Marti and Craig, 1987; Staudacher and Allègre, 1991, 1993), where cosmogenic neon was shown to be released above 900 °C. With the exception of FOR 2 and 4, all samples released more than 25, one sample even 93 % of 21Necos at 850 °C (Fig. 3.15). The 850 °C and 1500 °C data points plot, with two exceptions, within 2(7 errors on the atmospheric-cosmogenic mixing line in the neon three-isotope plot, so that a nucleogenic contribution can be excluded. The 1750 °C steps were entirely atmospheric, as expected. Therefore, we use both the 850 and 1500 °C steps for the calculation of the neon exposure ages (see discussion below). 3He was, as previously observed, uniformly released at 850 °C.

3.5.6 Discussion

Three samples have to be discarded from discussion. For BOW4 and BOW6, taken a few metres apart, the neon ages agree well, whereas the helium age of BOW6 is much smaller than for BOW4. Substantial helium loss from BOW6 must have occurred. The helium age of BOW6 will therefore be excluded from the discussion of the neon degassing pattern. From their geomorphological position on the Alpine Topography, CH2 and CH3 are expected to be contemporary to or older than CH4 and CH5. However, they show nearly no 21Ne excess and correspondingly little 3He. Substantial nuclide loss due to very high rock denudation rates must have occurred, as thick slabs of rock are falling off at these sampling sites. Therefore, CH2 and CH3 are not considered for the discussion of the exposure ages. However, these samples are still useful for the investigation of noble gas systematics in pyroxene.

Neon degassing pattern

The small number of studies that so far have addressed the subject of stepwise heating to release cosmogenic neon from pyroxene (Schäfer et ah, 2000; Marti and Craig, 1987; Staudacher and Allègre, 1991, 1993) give results for at most three samples each. Thus,

80 3.5 Erosional surfaces in Northern Victoria Land

0.400 , , , , r p 1 1 1 1

0.000 LJ—V l 0.0000X0.0500 0.1000 0.1500 0.2000

0.140 i * 1 . 1 .

0.090 I—X1 ' ' L \ 0.0100 0.0200 0.0300

0.116 , r*—, . . . .

Figure 3.14: Three - isotope plot of the neon data from pyroxene. Open diamonds are 850 °C steps, and filled diamonds are 1500 °C steps. 1750 °C steps are omitted for better readability of the plot as they are of atmospheric composition within errors. The star symbol in each panel represents atmospheric composition (2.959xlO-3 / 0.102, Eberhardt et al. (1965)).

The atmospheric / cosmogenic mixing line is drawn as y = 1.069a; + 0.099 (Schäfer et al., 1999).

81 3 Antarctica

100%

80%

60% o CD W CO D1750°C 0 40% 1500°C 850°C

20%

5 Ö 2 | S m ^r t- (n cd O O O O O I T I I OOITT CÛOQDQDÛCQU ÜÜÜU-LLDCüiQ:

Figure 3.15: Relative amounts of Necos released at each temperature step.

82 3.5 Erosional surfaces in Northern Victoria Land

Helium I D NeonHT O Necniot 3000 T i ï

2000 t ï ï ï £ 5 t 1000 IP

3E

-XΠr- ^ 0 "

illlloinSniSlïû-tt.t 00000xx±X004;xx

Figure 3.16: 3He and 21Ne ages of all Northern Victoria Land Samples. Neon HT: calculated from the 1500 °C step only; Neon total: calculated from 850 °C and 1500 °C steps.

83 3 Antarctica

limited data is available for comparison with the results in this study.

To find out if the use of both the 850 °C and the 1500 °C neon data (21Netot) for age calculation is justified, the two different neon ages are compared to the helium age in

3.16. Whereas this the use of for six it the Fig. plot suggests 21Netot samples , supports use of 21Ne#y for five samples. The helium age is identical to the Ne#y for two samples (BOW2 and RHP9) and to the Netot age for one (BOW4). In the following we consider the possibility that the degassing temperature of cosmo¬ genic neon depends on the chemical composition of the mineral. From Fig. 3.17, it is clear that the percentage of cosmogenic neon that is released at 1500°C is positively correlated to the content in Mg (correlation coefficient r=0.91), to a lesser degree to Si (r=-0.64) and Al (r=-0.74). I.e., the more Mg a sample contains, the more of its 21Necos is released above 850 °C. This behaviour corresponds well with the observation that diffusion is slower in Mg-rich magmatic systems due to the higher closure temperature (Balco et ah, 2002). The activation energy of particles in these systems is higher and therefore it takes higher temperatures to release them from the crystal lattice.

The same observation is made in a set of 3He and 21Ne measurements in pyroxenes from dolerite boulders from Beacon Valley, Dry Valleys (open symbols in Fig. 3.17; article in preparation). The correlation coefficient of the percentage of 21Necos released at 1500 °C with the Mg content for all samples here is even r=0.96.

As a consequence of these findings, it is suggested that the Mg concentration is mea¬ sured before the neon analysis, so the extraction procedure can be optimised. For py¬ roxenes very poor in Mg, it may not be possible to separate the cosmogenic from the non-cosmogenic neon fraction by stepwise heating. For these, a total extraction might be sufficient. As in all samples-with the exception of the Mg-poorest sample CH3-only minor amounts of the 3He are released at 1500 °C, total extraction would be sufficient for helium analysis, too. In samples of intermediate to high Mg content (>8 %), however, cosmogenic and non-cosmogenic neon phases can be efficiently separated by stepwise heating to 850, 1500 and 1750 °C, as suggested by Schäfer et al. (2000).

Retentivity of pyroxene for 3He

Figure 3.16 shows that for six out of eleven samples the helium age is lower than the the total neon age, for some considerably. This raises the question if 3He is retained completely. Again, a correlation to the elemental composition is observed. In figure 3.17, the ratio of cosmogenic nuclides as a measure of helium retentivity is plotted against the concentrations of Si, Mg and Al. If both isotopes are completely retained and exclusively cosmogenic, the ratio 3He/21Necos equals one. The correlation coefficient r is 0.89 for Mg, -0.75 for Al and -0.65 for Si. Thus, the degree to which cosmogenic 3He is retained in pyroxene seems to be strongly positively correlated to the Mg content, and, less strongly, negatively to the contents in Al and Si. The data from Beacon Valley show the same characteristics (r=0.86 for Mg) and are given in figure 3.17 for comparison.

84 3.5 Erosional surfaces in Northern Victoria Land

• 1 1 1 1 1 1 ' 1 1 - o^ o o - 100

. o . . o .

_ - _ 40 *. \

- - 20

'

>> . . I t' l0 . 0 0 24 28 32 36 4 0 4 8 1 D 4 8 12 1

I - o „ < * o .o *.

-oto* - - -°o** - o«vï o *o 1 '

.

1 ' » ' 1 s—— 1 — i tft— 24 28 32 36 40 0 4 8 Si (wt %) Mg (wt %) Al (wt %)

ure 3.17: Upper three panels: The concentrations of Si, Mg and Al plotted against (upper three panels) the percentage of 21Necos released at 1500°C. Lower three panels: The concentrations of Si, Mg and Al plotted against the concentration ratio of the cosmogenic isotopes. In order to eliminate the effect of the elemental production rates on the neon concentration, nuclide concentrations are divided by the respective production rate scaled for latitude and elevation. Filled symbols: data from this paper;

open symbols: data from Beacon Valley.

85 3 Antarctica

Table 3.10: Relative concentrations of the elements that produce 21Necos and the 21Ne production rate at sea level and high latitude (slhl),

calculated with P2i(Si) = 42.5atg_1 yr"1, P2i(Mg) = 185atg-1yr_1>

?2i(Al) = 51atg_1 yr-1 (Schäfer et al., 2000). Uncertainties of he con¬ centrations are below 1 % for Mg and Si and below 4% for Al. Samples

between horizontal lines were taken at the same location.

Sample Si Mg Al P21 slhl (weight %) (weight %) (weight %) (atg_1yr_1)

BOW1 28 8 26 68 20 5 BOW2 27 6 26 3 6 18 3

BOW3 27 4 11 2 4 2 34 6 BOW4 24 9 66 30 24 2

BOW6 25 7 43 2 9 20 3

CH2 26 7 1 9 76 18 7 CH3 38 3 0 1 3 9 18 3

CH4 24 9 10 4 1 1 30 4 CH5 24 6 10 1 1 7 30 1

FOR2 24 7 110 1 0 31 5 FOR4 25 1 11 7 0 9 32 7

RHPl 24 7 54 44 22 7 RHP2 26 1 52 48 23 2

RHP9 25 0 8 1 28 27 0

This dependence is not surprising. As helium is even more diffusive than the larger neon atoms, it is more affected by the higher diffusion constants in Mg-rich minerals (Brady, 1995). To avoid this problem, it is advisable to separate Mg-rich pyroxenes, whenever possible.

Exposure ages

We interpret our data conservatively, i. e. we calculate exposure ages from the 21Neexc of all temperature steps that indicate a cosmogenic contribution in the neon three isotope plot (Fig. 3.14). These are both 850 °C and 1500 °C data, except for two samples (850 °C of FOR2 and 1500 °C of CH3 are not used). This is somewhat tentative as the degassing behaviour of cosmogenic noble gases from pyroxene is not entirely clear from the data discussed above. However, the neon isotope ratios show that all the steps used for age calculation are strongly, although probably not entirely, cosmogenic. The resulting ages are listed in Tab. 3.11.

Surface exposure ages have to be regarded as strict minimum ages as long as they are not corrected for the effects of erosion and possible uplift. Even when corrected, the apparent age can still be too low due to possible temporary shielding. Therefore, we consider the highest minimum age on a landform for discussion.

The Gargoyles on Mount Bowen have a helium minimum age of 3.0 Myr and a neon minimum age of 3.5 Myr; those on Chisholm Hills are 2.7 and 2.1 Myr old, respectively.

86 3.5 Erosional surfaces in Northern Victoria Land

Table 3.11: Neon and helium surface exposure ages for all samples from erosional surfaces. Production rates: see paragraph 3.5.4 and Tab. 3.10. Scaling of production rates according to Stone (2000). Gargoyle ages are corrected for geometry effects (see text). Ages in brackets: see text. Surface types:

sb = striated bedrock; G = Gargoyle; at = Alpine Topography.

Sample Surface Helium age Total neon age type (kyr) (kyr)

BOW1 Qz sb — 806±74 BOW1 Px sb 321±8 733±39 BOW2 sb 186±6 541±49 BOW3 sb 316±9 396±19 BOW4 G 3041±152 3331±136 BOW6 G (1998±104) 3485±126 (CH2 at 62±3 ( 210±16 ) (CH3 at 1±1 (111±21 ) CH4 G 2685±248 2098±85 CH5 G 1749±156 1412±60

FOR2 at 1620±118 1299±63

FOR4 at 1145±80 971±42

The Alpine Topography at Chisholm Hills could not be dated due to spalling rock surfaces; at Ford Peak, it yielded ages of 1.6 (helium) and 1.3Myr (neon). This "age reversal" reflects higher erosion rates on the Alpine Topography. The very steep relief above the trimline enhances the continuous processes of salt and frost weathering and leads to higher denudation rates. Episodic erosion by block detachment is facilitated, too, as it is well visible at the sampling sites of samples CH2 and CH3. Maximum erosion rates calculated under the assumption of infinite exposure time (Kurz and Brook, 1994) are given in table 3.12. Maximum erosion rates of the Alpine Topography range from 34 to 57cmMyr_1, whereas they range from 16 to 39cmMyr_1 for the Gargoyles. These values are similar to the 20 cm Myr-1 determined on granite bedrock near Terra Nova Bay (Oberholzer et ah, 2003b). However, the true erosion rates are assumed to be considerably lower. Thus, the time at which erosive ice sculpted the Gargoyle flats, engraved the trimline and initiated the development of the Alpine Topography is best approximated by the highest exposure age obtained, which comes from the Gargoyles. The most stringent limit to the age of the Gargoyles in the David Glacier area is 3.5 Myr, determined on

Mount Bowen. The outermost domes of the EAIS have never reached the same elevations as the Gargoyle terraces again after that time. This implies minimal variations of the ice surface level since Mid-Pliocene. On Mount Bowen, the Gargoyles are located at a lower elevation than the adjacent ice sheet surface. Present-day ice flow patterns suggest that the inland ice is likely to overflow the terrace with Gargoyles if its elevation in upstream direction is only 200 m higher. A constraint for the maximum ice level variation near Mesa Range is obtained from the Gargoyles on Chisholm Hills which today overlook the

87 3 Antarctica

Table 3.12: Maximum erosion rates Emax calculated from 3He and 21Ne data under

the assumption of infinite exposure (Kurz and Brook, 1994). Surface

types: sb = striated bedrock; G = gargoyle; at = Alpine Topography.

Sample Surface Emm ( He) Emax (21Ne) mean type (cmMa-1) (cmMa-1) (cmMa-1)

BOW1 Qz sb — 69±6 — BOW1 Px sb 173±4 76±4 124 BOW2 sb 298±9 103±9 200 BOW3 sb 176±5 140±7 158

BOW4 G 20±1 18±1 19 BOW6 G 30±2 17±1 23 (CH2 at 891±40 264±20 578) (CH3 at 39139±13258 500±93 19820) CH4 G 22±2 28±1 25 CH5 G 33±3 41±2 37

For2 at 34±2 43±2 39

For4 at 49±3 57±2 53

ice by no more than 250 m.

Pyroxene neon ages for the higher platform on Mount Bowen (samples BOW1 to BOW3) range from 396 to 733 kyr, and helium ages from 186 to 321 kyr. The BOW1 quartz age of 806 kyr is an upper limit to their age, because it might be affected by a small nucleogenic component, according to its position in the three isotope plot. Thus, the flat surfaces are considerably younger than the Gargoyles, although they are located 175 m higher than those. Erosion is not likely to be the reason for this difference, as it is not expected to be stronger on this flat platform than on the Gargoyles. It is more probable that local ice flow patterns, ruled by the strong katabatic winds, created an ice surface with a high relief, leading to an ice cover that at times of a slightly raised ice sheet surface reached the flat rock surfaces at 1700m a. s.l., but not the Gargoyles on the terrace at 1525 m a. s .1. a few kilometres away. The small variations in surface elevation of the EAIS in Northern Victoria Land testify for persistent stability of climate and landscape in this area. This indicates that the respective findings from the Dry Valleys do not reflect a particular uplift history of that area as suggested by Van Der Wateren et al. (1999). Instead, it seems reasonable to assume that the landscape of large parts of the Transantarctic Mountains has remained unchanged since Mid-Pliocene.

3.5.7 Conclusions

From the data presented in this study, conclusions can be drawn for two different sub¬ jects: First, the glacial history of Northern Victoria Land, and second, the noble gas systematics in pyroxene. Ancient erosional rock surfaces from the northern and the southern part of Northern

88 3.5 Erosional surfaces in Northern Victoria Land

Victoria Land yield minimum exposure ages of 2.1 to 3.5 Myr. These ages were deter¬ mined on dolerite rock columns referred to as Gargoyles or Antarctic Tors. An exposure age of 1.6 Myr is the first direct age constraint for the Alpine Topography, a type of morphology with serrated ridges and steep rock walls, widespread in Northern Victoria Land. This is a strict minimum age, though, as the Alpine Topography is strongly af¬ fected by erosion due to the steep relief and the associated enhancement of gravitational processes. From the position of the sampling locations relative to the present-day ice surface it is inferred that the EAIS surface elevation in Northern Victoria Land has not changed more than 250 m around Chisholm Hills for at least 2.7 Myr, and not more than 200 m around Mount Bowen for at least 3.5 Myr. Correction for erosion would increase these time ranges, but maximum erosion rates were as low as 16 cm Ma-1 on the Gargoyle terraces, and two to three times higher on the steep Alpine Topography. Thus, our data set contributes to the large body of evidence in support of a stable behaviour of the EAIS since early Pliocene. Particularly, it demonstrates that the Dry Valleys are not a special case regarding its landscape stability. So far, only scarce data existed on the separation of cosmogenic and non-cosmogenic noble gases in pyroxene by stepwise heating. This study supplies new information with an extensive data set, although the results are not easy to interpret. It is shown that sub¬ stantial amounts of the total cosmogenic neon can be released from pyroxene at 850 °C. The temperature at which cosmogenic neon starts to be released is strongly dependent on geochemical parameters. We have identified the Mg content as one important agent.

It is strongly positively correlated (r = 0.915) to the fraction of 21Necos that is released above 850 °C. This finding is confirmed by a second data set from the Dry Valleys. It is recommended to measure the Mg concentration in a sample before noble gas analysis in order to decide if stepwise heating is indicated. More extensive analyses of cosmogenic nuclides in pyroxenes along with careful examination of their composition will provide deeper insight into the geochemical parameters that determine the degassing pattern of the mineral.

Acknowledgements

This work was supported by the Italian Program on Antarctic Research (PNRA) and the Swiss National Science Foundation. We thank the pilots of Helicopter New Zealand and the Mountain Guides at Terra Nova Bay Station very much for making it possible to sample Ford Peak and other sites difficult to access. To Francesco Fasano we are grateful for his precious help with the collection of the larger samples.

89 3 Antarctica

90 4 Late Pleistocene glaciations in the monsoonal Tibetan Himalayas

91 4 Tibet

4.1 Introduction

The Tibetan Plateau represents a unique geological and climatological feature. The climatic feedbacks on the plateau impact directly the hydrological system feeding one of the most densely populated areas on earth. The understanding of Tibet's climate, fundamental to predict climate future scenarios ruling Tibet's hydrography, is therefore of outstanding importance, particularly in view of the growing anthropogenic impact on our climate system. It remains unclear how Tibet is climatically coupled to that of the rest of the planet. Did prominent climate changes such as the Last Glacial Maxi¬ mum impact Tibet in a similar way as mid-latitude, low-altitude sites on the northern hemispheres? How significant is the monsoon system as a driver of glaciations in Tibet?

A very effective strategy to improve our understanding of the nature of climate change and the sensitivity of the climate system to anthropogenic forcing is the reconstruction of past climate using paleorecords from various regions on our planet. We present here reconstructions of glacial advances, witnesses of significant reorganization of Tibet's hydrography in the past. We focus on an area in the monsoonal Himalayas which is particularly sensitive to the Indian Monsoon due to an orographic low in the Himalaya chain. Our goal is to evaluate the influence of moisture transport by the monsoon system for glaciations in Tibet. This work represents an extension of earlier studies on the eastern margin of the plateau, influenced by the East Asian monsoon system (Schäfer et ah, 2002; Tschudi et ah, 2003). We present our results in two coordinated manuscripts:

- "Late Pleistocene glaciations in the monsoonal Himalayas - Part 1: The MIS 2 glacial advances" by Schäfer et al.

- "Late Pleistocene glaciations in the monsoonal Himalayas - Part 2: Glaciation or denudation event in MIS 3?" by Oberholzer et al.

Both manuscripts contribute to the recent controversy whether or not the structure of late Pleistocene glaciations in Tibet are similar to other areas in the northern hemisphere.

Tibet and the Indian monsoon

The major moisture source of the Tibetan Plateau is the Indian Ocean monsoon system. It is therefore suggested to be the driving force of glaciations on the plateau. The moisture from the North Atlantic Ocean transported by the westerlies is precipitated mainly over the Eurasian continent, and the winds have a strongly continental character when they arrive in southern Asia.

In the northern summer, beginning in March or April, the Indian sub-continent and the Tibetan Plateau are heated up. The resulting land/sea temperature difference leads to a stable low-pressure zone over the area. As a consequence, the Inner Tropical Convergence Zone (ITCZ) is shifted from its winter position south of the equator to north of the

92 4.1 Introduction

Himalaya. The SE Passat winds that flow into the ITCZ change their direction due to the Coriolis force as they cross the equator and become SW winds (Fig. 4.1). The warm and humid air is carried across the Arabian Sea and the Himalaya mountains to the Tibetan Plateau as the Indian Ocean summer monsoon winds. They carry strong seasonal precipitation and are responsible for the heavy summer rains from June to September. Inversely, the ITCZ returns back to the south of the equator when the radiative heating of the continent is reduced in winter time due to lower insolation and increased albedo. The winter monsoon then drives cold and dry air southwards, from the Mongolian deserts across the Tibetan Plateau and the Himalaya onto the Arabian Sea (Fig. 4.1). Thus, both precipitation and temperature patterns in Tibet are tightly linked to the intensity of the monsoon circulation. While an enhanced summer monsoon results in increased precipitation, a strong winter monsoon leads to cold temperatures.

Driving mechanisms of Tibetan glaciations

Due to its mountainous topography combined with a high elevation, conditions in Tibet are suitable for extensive glaciations under favourable climate conditions. Therefore, Tibet potentially is an active player in global climate change and may have been a major factor in initiating Quaternary glaciations in the northern hemisphere (Raymo and Ruddiman, 1992). A persistently snow or ice covered plateau has a high albedo and in this way which weakens the entire Indian summer monsoon circulation (Vernekar et ah, 1995; Anderson, 1993). The glacial conditions on the Tibetan Plateau therefore influence the climate on a supraregional scale. However, the role of the Tibetan Plateau and the Himalaya in the global climate system of the past is still not understood. In particular, the temporal and causal rela¬ tionship between monsoon intensity and cold events in the North Atlantic region, such as the Wisconsian/Weichselian ice sheet growth, the Heinrich events or the Younger Dryas remains unclear. To answer this question, we have to tackle the directly related problem of the driving mechanisms of Tibetan glaciations, and therefore we have to know more about the spatial and temporal structure of glacial advances in Tibet. The pattern of glaciation in the Himalayas and on the Plateau as emerged so far is complex and marked by a strong local variability. This may be partially caused by small- scale variability of the factors that determine the mass balance of a glacier. A dominating parameter is the temperature during the ablation season. The glacier dynamics, however, critically depend on several other internal and external factors. Internal ones are the aspect and slope of the glacier, its elevation range, the thickness of its sediment bed, or the bedrock topography. External variables are, besides temperature, precipitation and local wind patterns. This local variability is expressed in the dating of the Last Glacial Maximum (LGM) by different authors. A number of articles reported a "local LGM" in early marine isotope stage (MIS) 3 or late MIS 4, 40 to 60 kyr ago, out of phase with the global LGM in MIS 2. This was found in the Himalayas, the Hindukush, the Karakorum

93 4 Tibet

Figure 4.1: The monsoon systems of the Earth in the Northern Hemisphere sum¬ mer and winter. The Tibetan Himalaya is under the influence of the Indian Ocean Monsoon, in the figure labelled as Indian Southwest Mon¬ soon, ISWM. In summer time, the monsoon flows towards the Himalaya from a clearly southwestern direction. Monsoon according to a strict climatological definition is active in the areas indicated by the dashed rectangle. Cross-hatched areas have the highest surface temperatures. From Webster (1987).

94 4.1 Introduction

and the mountains north of the plateau (Finkel et ah, 2003; Richards et ah, 2000;

Benn and Owen, 1998; Owen et ah, 2003a) . Early MIS 3 was a period of intensified Indian summer monsoon (Prell and Kutzbach, 1987), but an interglacial on the Northern hemisphere. This raises the question whether a strong monsoon alone can cause ice advance in the Himalaya and on the plateau, while the rest of the northern hemisphere undergoes interglacial conditions. In other words, can sufficient moisture supply alone trigger a glaciation and overrule the effect of a higher-than-average temperature? On the other hand, North Atlantic climate events have been recorded in Tibet. (Tschudi et ah, 2003) have found moraines deposited during the Younger Dryas cold reversal. The north Atlantic LGM in MIS 2 has made Tibetan glaciers advance, too (Schäfer et ah, 2002), although to a smaller extent than in other parts of the northern hemisphere. In the Tanggula Shan in the centre of the plateau, the most extensive ice advance occurred in early MIS 6, or probably MIS 7 (Schäfer et ah, 2002). Finally, older glaciations are recognized at various times throughout MIS 5 to 8 (Derbyshire et ah, 1991). Thus, it is not clear if climate signals from the north Atlantic region are transferred to central and south Asia, or if regional climate mechanisms determine the pattern of ice ages in Tibet through precipitation effects. The interpretations of the available evidence are contradictory in this respect.

Glacial geological setting

The southern slope of the region around Mt. Qomolangma (Mt. Everest) and Mt. Shisha- bangma is characterized by deeply incised valleys between high mountains. Many of the valleys are occupied by valley glaciers, and glacial deposits of past glaciations are abundant. According to Zheng (1988), the southern slope of Mt. Shishabangma (8012 m a. s .1., Fig. 4.2) has experienced at least four major glaciations since the Pliocene. Deposits of these are excellently conserved in the Fuqu Valley upriver from the village of Nyalam (3750 ma. s .1.). Zheng (1988) names them (from younger to older) Little Ice Age, Neoglacial, Puluo (or Qomolangma II), Fuqu (or Qomolangma I), and Nyanyax¬ iongla (or Nyalam) glaciations. The latter three of these stages were assigned middle and late Pleistocene ages based on their morphology, paleobiology and soil cover. On the south slope of Mount Shishabangma, traces of the Shishabangma glaciation are found only at high elevations. Harper and Humphrey (2003) questioned if the Nyanyaxiongla moraines could be correlated with the Nyanyaxiongla glaciation that has its type site 45 km north on Yaruxiongla Pass. It seems dubious that the deposits at the type site are of glacial origin. Therefore, this glaciation is referred to as Nyalam in this paper. West and south of Nyalam, Zheng (1988) found lateral moraine terraces along the modern Fuqu river which he assigned to Fuqu glaciation. In the field, the glacial origin of these deposits seemed questionable to us. They may well be fluvial terraces. Therefore, we concentrated on the deposits of unequivocally glacial character.

95 4 Tibet

85° 50' E 86° 00' E

10 km

28°10'N-

Lhasa / *ç ^V study area

Figure 4.2: Schematic map of the southern slope of Mount Shishabangma. A: Nyalam moraines discussed in section 4.3; B: Puluo moraines discussed in section 4.2. Naisa Valley is located 15 km north of Nyalam along the road to Tingri/Lhasa.

96 4.1 Introduction

Figure 4.3: Sketch map of the moraine walls of lower Fuqu valley. The moraine ridge the samples were collected from are labelled with triangles (Nyalam) and diamonds (open diamond: Puluo 1; filled diamond: Puluo 2). From Academia-Sinica (1980).

97 4 Tibet

4.2 Pleistocene glaciations in the monsoonal Himalayas Part 1: The MIS 2 glacial advances1

Joerg M. Schaefer"'*, Peter Oberholzer6, Christian Schluechterc, Zhizhong Zhaod, Susan Ivy-Ochs6, Peter W. Kubik6'^, Heinrich Baur6, Rainer Wieler6

a : Lamont-Doherty Earth Observatory, Columbia University, Palisades NY 10964, USA b: Isotope Geology, ETH Zürich, Switzerland c: Institute of Geology, University of Bern, Switzerland d: Chinese Academy of Geological Sciences, Institute of Geomechanics, Beijing, PR China

e : Particle Physics, ETH Zürich, 8093 Zürich, Switzerland f : Paul Scherrer Institute, Vilhngen, Switzerland *: corresponding author: Joerg Schaefer, Lamont-Doherty Earth Observatory, Geochemistry (Room 61), Route 9W, Palisades, NY 10964, USA Tel: (001) 845 365 8703, Fax: (001) 845 365 8155 email: schaeferQldeo. Columbia, edu

Abstract

We present here evidence for two distinct glacial advances during the late stage of the last glacial cycle in the monsoonal Himalayas. Glaciers were in advanced position dur¬ ing Marine Isotope Stage (MIS) 2, as well as prior to 10.5 kyr. The corresponding snow line depression was 650 ± 100 m for the older and between 50 and 150 m for the younger advance, respectively. This is consistent with earlier studies from the eastern margin of the Tibetan Plateau. The snow line lowering is about half to two third of the classical northern and southern hemisphere snow line depressions during this time, which is most likely explained by decreased monsoon intensity during cold periods.

In view of conclusions of earlier studies, we infer that the Tibetan LGM and Late Glacial

events were of plateau-wide extent. It seems very probable that the late Pleistocene glaciations in Tibet have been synchronous to the classic Northern Hemisphere events.

This holds true even in monsoonal areas, where moisture supply is low during cold periods. Arguments presented here as well as in a companion paper (Oberholzer et al., this issue) suggest that Tibetan glaciations did coincide with periods of cooler temperatures and decreased monsoon intensity. This assigns a minor role to the monsoon as drivers of

Tibetan In turn this that even for the mass-balance glaciations. , implies tropical glaciers

on time scales of 100 yrs and longer is much more sensitive to supraregional changes in temperature during the ablation season than to changes in precipitation patterns.

4.2.1 Introduction

Excursions of Central Asia's climate trigger catastrophic natural hazards in the most populated area of our planet, such as disastrous flooding or storm events. The reliability

1 To be submitted to Quaternary Science Reviews, coordinated with section 4 3.

98 4.2 Tibet: The MIS 2 advance

of predicting future climate excursions critically depends on the knowledge of the pro¬ cesses that govern the climate system. An effective strategy to improve this knowledge is the reconstruction of the climate system in the past as it is archived in various ma¬ rine and terrestrial records. Glacial moraines on the Tibetan Plateau mirror sequences of glacial cycles and represent a unique archive of past fluctuations of the atmospheric regime and the precipitation patterns of Central Asia in general, and past excursions of the monsoon in particular.

Recently, the progress of modern dating techniques, particularly of Surface Exposure Dating (SED) using cosmogenic nuclides, has allowed to gain new insight into the past glaciations on the Tibetan Plateau. Various recent studies suggest that glaciations in Tibet have been asynchronous to the classic Northern Hemisphere glaciations, but rather occurred during periods of intensified monsoon (e.g. Finkel et al. (2003); Owen et al. (2002a, 2001); Phillips et al. (2000)). In particular, several of these studies report that the local Last Glacial Maximum (LGM) in Tibet did not occur during the cold period during MIS 2 (12-24 kaBP) but rather during the relatively warm MIS 3 (24- 60kaBP). Further, these studies do not find much evidence for glacial advances during the Late Glacial period, while the "classic" mid-latitude northern records testify of glacial advances during the Younger Dryas period (e. g. Finkel et al. (2003); Gosse et al. (1995); Ivy-Ochs et al. (1999); Owen et al. (2002b, 2001); Phillips et al. (2000)) between 13 and 11 ka BP. However, these Tibetan studies do report early Holocene glaciations consistent with a reported monsoon maximum. These results are in disagreement with the timing of mid-latitude glaciations and imply that Tibet's glaciers rather react to changes in the moisture transport, i.e. in the monsoon intensity, than to temperature changes during the ablation season. On the other hand, studies from the monsoonal eastern margin of the plateau report glacial events during MIS 2 (Schäfer et ah, 2002) and discuss a potential glacial advance during the Late Glacial period consistent with the Younger Dryas cold reversal (Tschudi et ah, 2003). These latter studies suggest that glaciations in monsoonal Tibet react to the same driving mechanisms as mid-latitude glaciers and assign a rather limited role to the monsoon system in driving glaciations on the Plateau. To clarify this fundamental conflict how the climate of Tibet is coupled to the north¬ ern hemisphere climate and the monsoon system, respectively, we investigate timing and amplitude of past glacial advances in the highly monsoonal Nyalam County in the Himalaya of southern Tibet. Glaciers in this area are particularly exposed to the Indian monsoon winds, due to an orographic low in the Himalayan mountain chain and due to relatively low-altitude. We calculate the underlying snow line lowering and use SED to date the glacial advances. The results are compared to data from classic mid-latitude moraine records. We focus on two key questions:

1. Is the timing of Late Pleistocene glaciations in the monsoonal Himalayas consistent

with the classic northern hemisphere LGM - Late Glacial pattern?

2. Are the snow line depressions comparable to those elsewhere, which were some 900

99 4 Tibet

to 1000 m during the LGM and some 200 to 400 m for the Late Glacial event?

4.2.2 Methods and geological setting

Sampling

Sampling sites are shown in Fig. 4.2 and 4.3. Nyalam County is situated in the mon¬ soonal Himalayas, at the southern margin of the Plateau. We investigated two valleys. The glaciations in Fuqu Valley, of which the Puluo Valley is a tributary, are discussed in the companion paper. Here we investigate the timing and amplitude of the Puluo glaciation (Fig. 4.3), and the Naisa glaciation in Naisa valley (a tributary valley to Poiqu Valley, 15km north of Nyalam village, see Fig. 4.2). We only sampled large boulders (larger than 1.5m from the ground) to minimize the probability of occasional snow cover. In addition, we concentrated on boulders whose weathered appearance indicates long-term exposure to avoid recently exhumed rocks. From the weathered boulders, we sampled the best-preserved surface section. We sampled two erratics on the Puluo 1 moraine, three erratics on Puluo 2, and four on the Naisa moraines. The rocks are of biotite-gneissic lithology. Due to the high biotite content, this lithology is less resistant to steady grain-by-grain erosion than e. g. quartz-rich granites. This becomes obvious from looking at the pit-holes in the rock's surface.

Isotope analysis

Quartz separation and chemical processing for 10Be analysis was performed at the Surface Exposure Dating laboratory at the Lamont-Doherty Earth Observatory following an updated protocol based on the procedure given in Kohl and Nishiizumi (1992). Samples were analyzed for in-situ 10Be at the AMS facility at ETH Zürich. Cosmogenic neon isotopes were analyzed at the laboratory for Isotope Geology at ETH Zürich using a split of the quartz separates prepared for the 10Be analysis.

Puluo glaciations

The Puluo glaciations have deposited two distinct moraines sets, representing different glacial events. Puluo 1 is more extensive than Puluo 2. The Puluo 2 terminal moraine is inside Puluo valley. It appears likely that the Puluo Glacier has merged with the Fuqu glacier during Puluo 1 glaciation, but since the terminal moraine is has been eroded, this is not certain. We sampled the left lateral moraines of both sets. The left lateral moraine of the more extensive Puluo 1 event is about 50 m higher than the Puluo 2 left lateral moraine.

100 4.2 Tibet: The MIS 2 advance

Naisa glaciations

The glacial deposits in Naisa Valley consist of a multi-ridged moraine set. At least three sub-advances are documented by a moraine-belt. The distance of the terminal moraine ridges of the innermost to the outermost moraine is less than 200 m. Both lateral and terminal moraines are preserved. We sampled three boulders on the innermost left lateral/terminal moraine and two of the outermost right lateral moraine with the aim to date the oldest (Naisa-a) and youngest (Naisa-b) glacial event of this set. In contrast to the Puluo events, the two sampled moraine ridges belong to the same major glacial event, we therefore label them Naisa "a" and "b" and not "1" and "2". Both glaciers, Puluo and Naisa, still exist, which allows us to estimate snow line depressions related to the glacial advances. The accumulation areas are at an altitude of about 5500 m, which avoids the high-altitude desert effect as reported by Harper and Humphrey (2003).

Production rates

For the 10Be exposure ages we used the sealevel and high-latitude value of 5.1 at g_1 yr-1 as well as the scaling procedure given in Stone (2000). For the 21Ne production rates we used the sealevel and high-latitude value of 20atg_1yr_1 given in Niedermann (2000). The use of any other production rates and scaling procedures, respectively, does not affect any of the conclusions here.

Snowline depression

To estimate the snow line lowering underlying the respective glacial advances, we applied the THAR (Toe-to-Headwall Altitude Ratio) method, using a THAR of 0.4 (Meierding, 1982). This method is the most robust one when no high-resolution maps are available. We regard this method as sufficiently accurate for the semi-quantitative comparison of Tibetan snow lines relative to northern mid- and high-latitude snow line records. The calculations are based on a Russian 1:200,000 map and a 1:100,000 map of the Chinese Geological Survey.

4.2.3 Results

Results are given in Tables 4.1 and 4.2 and Fig. 4.4. No neon analysis was carried out for the Naisa samples. In accordance with their relative position in the field, the samples from the Puluo 1 and Naisa a and b moraines yield similar exposure ages between 14.1 and 24.8 kyr. In contrast, the Puluo2 samples are significantly younger (10.0 to 10.7kyr). The neon ages generally agree with the beryllium data, which argues against significant periods of pre-exposure. However, the uncertainties of the neon data are very large. Therefore, the conclusions given in this paper are based on the radionuclide data alone.

101 4 Tibet

Table 4.1: Be exposure ages for the Puluo and Naisa moraines.

Sample elevation thickness 10Be minimum 10Be 10Be ages 10Be ages

(ma s 1 ) correction (106 at g-1) ages (kyr) e=l mm/kyr e=5 mm/kyr (%)

Puluo 1 Ny9 4301 2 1 009 ± 0 059 154 ± 0 9 156 ± 09 16 7±1 0 NylO 4301 3 1 615 ± 0 129 24 9 ± 2 0 25 6± 2 1 28 7±2 4 Nyll 4344 2 n a

Puluo 2 Nyl2a 4261 3 0 681 ± 0 041 10 9 ± 0 7 110 ± 07 11 6±0 7 Nyl2b 4261 4 0 655 ± 0 030 10 5 ±0 5 10 6 ±0 5 11 1 ± 0 5 Nyl3a 4273 7 0 667 ± 0 025 10 9 ± 0 4 110 ± 04 11 5 ±0 4 Nyl3b 4273 7 0 703 ± 0 059 11 5 ± 1 0 11 6 ± 10 122 ± 1 1 Nyl4 4265 4 0 624 ± 0 046 10 0 ± 0 7 10 1 ± 0 8 10 5 ± 0 8

mean 10 6 ± 0 4

Naisa a Nai4 4280 5 4 519 ± 0 048 71 0 ± 4 6 76 7 ± 5 0 125 5 ± 8 1

Nai5 4280 2 1 505 ± 0 066 22 9 ± 1 0 23 4± 1 0 26 0 ± 1 1

Naisa b Nail 4289 4 0 938 ± 0 057 14 6 ± 0 9 14 7 ± 0 9 156 ± 1 0

Nai2 4308 5 n a

Nai3 4343 6 0 871 ± 0 048 14 2 ± 0 8 14 4 ± 0 8 153 ± 0 9

Given uncertainties are 1er including analytical and blank errors Nyll did not contain quartz and was therefore not analysed Bold numbers are the base for further discussion Typical blanks yielded ratios 10Be/9Be = 1-4 x 10~14, sample sizes were 20-40 g pure quartz and 0 3-0 5mg of 9Be-carner were added (Spex-standard 1 mg/ml Be) e is erosion rate Samples labelled "a" and "b" represent duplicate measurements from the same rock Ages are calcu¬ lated using the sealevel/high-latitude 10Be production rate of 5 latg-1 yr-1 given in Stone (2000) as well as the scaling procedure given in that manuscript We used the following parameters for thickness correction and erosion rate incorporation attenuation length L=150 g/cm2, t1/2(10Be)=l 5Ma

102 4.2 Tibet: The MIS 2 advance

30

- Puluo 1

MIS 3

C ) A/a/sa a

O

20 - Nai5

MIS 2 Naisa b ^ $ 0 $ Ny9* Ny1' Piilml Nai1 Nai3 Younger Dryas $ 10 - OB $

- Ny12 Ny13 Ny1 4 MIS1

Figure 4.4: Beryllium exposure ages of the Puluo and Naisa erratics. Boxes represent

different moraine ridges. The grey shaded band indicates the Younger Dryas period. With respect to the position of the ice, Puluo 1 is more

external than Puluo 2, and Naisa a more external than Naisa b.

103 4 Tibet

Table 4.2: 21Ne exposure ages for the Puluo moraines.

Sample T (°C), 20Ne 21Ne/20Ne 22Ne/20Ne 21Neea;c 2-'-Ne age time (mm) (109 at g-1) 10~3 (106 at g-1) (kyr)

Puluo 1 Ny9 400, 45 3 44 ± 0 10 3 986 ± 0 341 0 1023 ± 0 0024 5 047 ± 3 026 19 6 ± 11 7 Ny9 600, 45 3 56 ± 0 20 3 385 ± 0 173 0 1036± 0 0033 Ny9 1750, 15 5 87 ± 0 22 3 428 ± 0 189 0 0998 ± 0 0023 Nyll 400, 45 2 85 ± 0 06 5 879 ± 0 222 0 1036 ± 0 0027 27 740 ± 2 290 105 2 ± 8 7 Nyll 600, 45 3 68± 0 062 8 228 ± 0 288 0 1067 ± 0 0021 Nyll 1750, 15 5 26 ± 0 03 7 241 ± 0 169 0 1059 ± 0 0014

Puluo 2 Nyl2 400, 45 4 88 ± 0 28 3 340 ± 0 123 0 1035 ± 0 0033 3 888 ± 3 624 15 5 ± 14 5 Nyl2 600, 45 4 48 ± 0 14 3 412 ± 0 261 0 1021 ± 0 0044 Nyl2 1750, 15 10 85 ± 0 13 3 623 ± 0 269 0 0995 ± 0 0018 Nyl3 400, 45 13 36 ± 0 12 3 138 ± 0 257 0 1021 ± 0 0014 8 524 ± 9 493 318 ± 35 4 Nyl3 600, 45 34 45 ± 0 50 3 137 ± 0 114 0 1015 ± 0 0016 Nyl3 1750, 15 10 56 ± 0 12 3 832 ± 0 154 0 1029 ± 0 0019 Nyl4 400, 45 7 08 ± 0 24 3 154 ± 0 216 0 1014 ± 0 0021 5 147 ± 6 592 20 9 ± 26 8 Nyl4 600, 45 12 97 ± 0 44 3 249 ± 0 190 0 1021 ± 0 0014 Nyl4 1750, 15 7 47 ± 0 29 3 049 ± 0 202 0 1019 ± 0 0027

Uncertainties are 2

In a monsoonal environment, where freeze-thaw cycles dominate the weathering of the relatively weak lithology of the samples (biotite-rich gneiss), erosion rates may be substantial, even on the relatively short timescales discussed here. To give an impression for the sensitivity of exposure ages to erosion, we present the ages in Table 4.1 not only as minimum ages, assuming zero erosion, but also incorporating steady state erosion rates of 1 and 5 mm kyr-1, respectively. In the companion paper, we calculate maximum erosion rates as low as 1.5 to 3 mm kyr-1 from boulders of the same lithology. We therefore assume that an erosion rate of lmmkyr"1 is realistic, even if it is on the low side of the range of typical erosion rate for bedrock in alpine environments of 0 to 20 mm kyr-1 presented by Small et al. (1997).

Ages of glaciations

The exposure ages derived from the three samples from the Puluo 2 moraine agree excellently (Tab. 4.1). The mean value of the three minimum ages is 10.3±0.3kyr. This agreement basically excludes the possibility of pre-exposure, since identical pre¬ exposure periods for three boulders are highly unlikely. However, there are several processes potentially making the exposure age younger than the true deposition age, such as exhumation, snow and vegetation cover. Since we are not correcting for these processes, we consider the oldest age within a cluster of exposure ages the most reliable,

104 4.2 Tibet: The MIS 2 advance

and therefore assign the exposure age of Nyl3 of 11.5 ±1.0 kyr as the most likely age for the Puluo 2 event. Conservatively, the Puluo 2 moraine was deposited between 10.2 kyr (the low eror limit of the mean of the minimum ages) and 13.3 kyr (the upper error limit of the 10Be age of Nyl3 including an erosion rate of 5 mm/kyr).

This means that the Puluo 2 glaciation occurred prior to the Holocene. It is slightly younger than, but within uncertainties consistent with, the timing of a glacial event from Eastern monsoonal Tibet reported in an earlier study (Tschudi et ah, 2003), relating Late Glacial advances in Tibet to the termination of the Younger Dryas cold reversal some 11.5 kyr ago (Alley, 2000). The slightly differing exposure ages can be an effect of the different lithologies, yielding distinct erosion rates in the Nyalam biotite-gneiss compared to the quartz-rich granite in Kanding county (Tschudi et ah, 2003).

The two samples from Puluo 1 moraines analyzed so far for 10Be yield minimum ages of 15.3 ± 0.9 and 24.8 ± 2.0, i. e. consistent with the timing of MIS 2. Taking into account our best guess value of an erosion rate of 1 mm kyr-1 (Small et ah, 1997) yields ages of 15.6 ±0.9 kyr and 25.6 ±2.1 kyr, respectively. These two ages most likely represent two sub-advances during MIS 2 which did not result in distinct ridges, most likely due to the narrow glacier bed in Puluo Valley. This assumption is supported by consistent ages from Naisa Valley, where the glacier tongue ended in open terrain (see below).

The two boulders Nail and Nai2 from the innermost and therefore youngest Naisa Valley moraine display internally consistent minimum ages of 14.1 ± 0.8 and 14.4 ± 0.8 kyr

Those numbers are within uncertainties consistent with the age of Ny9 of 15.3 ± 0.9 kyr BP. Nai5, from the outermost right lateral moraine yields a minimum exposure age of 22.8 ± 1.0 kyr BP, indistinguishable from the minimum age of NylO. Adopting a steady state erosion rate of lmmkyr"1 increases the ages only slightly. Nai4, with a minimum exposure age of 71.0 ±4.6 kyr BP is an outlier, most likely due to exposure prior to deposition on the moraine.

Corresponding snow line depressions

We calculated snow line depressions corresponding to the glacial advances Puluo 2 and Naisa. Since the terminal moraines of the more extensive Puluo 1 stage is not preserved, the results represent a conservative minimum.

Our calculations yield snow line depressions 100 ± 50 m for the Puluo 2 moraine, 230 ± 60 m for the Puluo 1 event, and 650 ± 100 m for the Naisa glaciation, respectively. Applying an adiabatic lapse rate of some 0.7°C/100m, which seems appropriate under the humid conditions of Southern Tibet, the recorded snow line depressions correspond to temperature changes of 0.4 to 1.0 °C (Puluo 2), 1.2 to 2 °C (Puluo 1) and 3.9 to 5.3 °C (Naisa), respectively.

105 4 Tibet

4.2.4 Discussion

This study reports cosmogenic dating of glacial moraines in Nyalam County indicating that the timing of the climate changes during the last déglaciation in the monsoonal

" Himalayas is indistinguishable from the classic" northern hemisphere déglaciation pat¬ tern, the latter is characterized by a dominant LGM and a Late Glacial advance. We also present first estimates for snow line depressions for the Late Glacial advance in Tibet. These numbers imply that the amplitudes of the MIS 2 and the Late Glacial advances were about a half to two third of those in northern mid-latitudes, respectively (e.g. Swiss Alps (Maisch, 1987), Norway (Dahl and Nesje, 1992), Southern Alps of New Zealand (Porter, 1975), Chilean Lake District (Denton, 1999)). On the other hand, these studies report a ratio of full glacial snow line depression to Late Glacial snow line depression of approximately 4 to 6. This is consistent with our reconstructions within the given uncertainties. The temperature reduction inferred here for the MIS 2 advance (Naisa glaciation) nearly matches the cooling in tropical areas during the LGM of approximately 5°C (Beyerle et ah, 2003; Stute et ah, 1995). A striking correlation has been recorded between atmospheric cooling as recorded in the Greenland ice cores on the one side and decrease in monsoon intensity for both the East Asian monsoon (Wang et ah, 2001) and the Indian monsoon (Yuan et ah, 2004) on the other. These correlations were identified by comparing high-resolution oxygen isotope records in absolutely and precisely dated speleothems from China to Greenland ice core oxygen isotope data. This reduction of moisture supply during cold periods is the most probable reason for the limited amplitude of Tibetan glaciations. Two other reasons are the following:

(i) Compared to other glaciated areas, the Himalayas have very large distance to the next major moisture source, in the case of Nyalam about 2000 km to the Arabian Sea and 1000 km to the Bay of Bengal (compared e.g. to the European Alps: 0-200 km to the Mediterranean Sea; Rocky Mountains: < 300 km to the Pacific; New Zealand: < 100 km to the Southern Ocean; South American Andes: < 300 km to the Pacific);

(ii) Various effects of high altitudes diminish snow accumulation significantly: the de¬ creasing moisture content of the air, strong winds that remove snow from the

accumulation areas, and sublimation rates which increase drastically with wind- speed. All these effects become stronger with increasing elevation.

In the companion paper, evidence is presented against the case of a local LGM in Tibet during MIS 3 that preceded the last major glaciation of the northern hemisphere. Instead, rapid moraine degradation as a consequence of increased precipitation dur¬ ing the phase of strengthened monsoon in MIS 3 is suggested. Together, these papers propose that the timing of Late Pleistocene glaciations in monsoonal Tibet is within un¬ certainties indistinguishable from the "classic" northern mid-latitude LGM/Late Glacial

106 4.2 Tibet: The MIS 2 advance

pattern. The glaciations most likely occurred during periods of cool climate, accompa¬ nied by moderate to low monsoon intensity. This explains the smaller amplitude of late Pleistocene glaciations in monsoonal Tibet due to reduced moisture supply to the glacier accumulation areas. A contradicting interpretation of the Tibetan glacial record is supported by various studies (e.g. Finkel et al. (2003); Owen et al. (2002b,a); Phillips et al. (2000)). They advocate the idea that the monsoon is the dominant driver of glaciations in Tibet and Late Pleistocene glaciations are driven by the same forces than elsewhere in the northern hemisphere. The two studies presented here contradict that hypothesis. As suggested by various recent studies (Alley, 2003; Oerlemans, 1994, 2001), we propose that the dominant pa¬ rameter influencing glacial mass-balance on time scales of 1000 yrs and more, even in monsoonal areas, is the temperature during the ablation season, the influence of which by far exceeds the impact of changes in precipitation patterns on longer time scales.

4.2.5 Conclusions

This study reports absolute age determinations of glacial events which are within over¬ all uncertainty indistinguishable from the MIS 2 and Late Glacial event reported from northern mid-latitudes. The timing of the late Pleistocene glacial events is also consis¬ tent with the timing of cooler and less humid conditions as reported from speleothem records in both, East Asian and Indian monsoon areas. The amplitude of both events is reduced compared to northern mid-latitudes. This is demonstrated by estimations of snow line depressions for three different glacial events. The most likely reason for this limited glacier extent is reduced moisture supply compared to other glaciated areas, which indicates that the monsoon is not efficient enough in transporting moisture to the high-altitude accumulation areas of Tibet's glaciers to sustain significant glaciations. Together with a companion paper that explains the absence of glaciations in Tibet dur¬ ing MIS 3, this article presents independent arguments contradicting the hypothesis that Tibet's glaciations are driven by increased monsoon intensity.

107 4 Tibet

4.3 Late Pleistocene glaciations in the monsoonal Himalayas Part 2: Glaciation or denudation event in MIS 3?

Peter Oberholzer"'*, Joerg M. Schaefer6, Christian Schliichterc, Zhizhong Zhaod, Susan Ivy-Ochs6, Peter W. Kubik6'-^, Heinrich Baur", Rainer Wieler"

a : Isotope Geology, ETH Zürich, Sonneggstrasse 5, 8092 Zürich, Switzerland b: Lamont-Doherty Earth Observatory, Columbia University, Palisades NY 10964, USA c: Institute of Geology, University of Bern, Switzerland d: Chinese Academy of Geological Sciences, Institute of Geomechanics, Beijing, PR China

e : Particle Physics, ETH Zürich, 8093 Zürich, Switzerland f : Paul Scherrer Institute, Vilhngen, Switzerland *: corresponding author: [email protected]

Abstract Glacial erratics from three morainic ridges near Nyalam, Tibetan Himalaya, yield

a cluster of low minimum exposure ages (48 to 65 kyr) and a group of ages distributed

between 100 and 300 kyr. The data scatter among different blocks is considerable, but

the agreement of 10Be and 21Ne ages on single erratics is excellent. By means of shielded

control samples it is shown that this data scatter is not caused by pre-exposure. Instead,

we suggest significant variations in landscape denudation rates and resulting variations

in rates of boulder exhumation, triggered by changes in monsoon intensity to explain

the data. Thus, the concentration of ages between 48 and 65 kyr may reflect accelerated

landscape evolution rather than a glacial advance. The deposition of the moraines took

place before 174 ± 16kaBP by a glaciation that was of intermediate extent, as indicated by

a snowline lowering of 500 ± 100 m. These data indicate that the exposure ages from other

regions in the Himalaya that have led to the postulation of a "local Last Glacial Maximum

(LGM)" in Tibet during MIS 3 reflect a denudation event rather than a glaciation. The

absence of a glaciation during the phase of intensified Indian monsoon in MIS 3 suggests

that the monsoon is not a major force in triggering glacial advances in Tibet. This finding

is supported by the MIS 2 age of moraines in the same glacial system (Schaefer et al., this issue).

4.3.1 Introduction

The Tibetan Plateau is a unique geological and topographical feature. Its glaciers and rivers feed the hydrological system of one of the most densely populated areas on earth.

! To be submitted to Quaternary Science Reviews, coordinated with section 4 2

108 4.3 Tibet: The MIS 3 denudation event

Understanding Tibet's climate is thus of outstanding importance to develop future sce¬ narios for Tibet's hydrography, particularly in view of the growing anthropogenic impact on our climate system. It is unclear today how Tibet is climatically coupled to the rest of the planet. Did prominent climate changes such as the Last Glacial Maximum affect Tibet in a similar way as they did mid-latitude, low-altitude sites on the Northern Hemispheres? How significant is the monsoon system as a driver of glaciations in Tibet? A very effective strategy to improve our understanding of climate change is the recon¬ struction of past climate. We present here reconstructions of glacial advances, distinct witnesses of climate changes in the past. We focus on an area that is dominated by the Indian monsoon due to an orographic low in the Himalaya mountain chain. Our goal is to evaluate the relative importance of the monsoon system on one hand and the westerlies on the other for glaciations in Tibet. This will contribute to resolving the recent controversy whether or not the patterns of late Pleistocene glaciations in Tibet are similar to those of other areas in the Northern Hemisphere (e. g. Finkel et al. (2003)). This work represents an extension of earlier studies on the eastern margin of the plateau, influenced by the East Asian monsoon system (Schäfer et ah, 2002; Tschudi et ah, 2003). We present our results in two coordinated manuscripts (Schaefer et ah, this issue, and this work). Due to its large area and its high elevation, Tibet potentially is an active player in global climate change and may have been a major factor in initiating Quaternary glacia¬ tions in the Northern Hemisphere (Raymo and Ruddiman, 1992). It was shown that a persistently snow or ice covered plateau with a high albedo weakens the Indian summer monsoon circulation (Vernekar et ah, 1995; Anderson, 1993). The glacial conditions on the Tibetan Plateau therefore influence the climate on an supraregional scale. The patterns of glaciation in the Himalayas and on the Plateau emerging so far are complex and marked by strong local differences, leading to strongly diverging interpre¬ tations of evidence regarding the Last Glacial Maximum (LGM). A number of articles reported a "local LGM" in early marine isotope stage (MIS) 3 or late MIS 4, 40 to 60 kyr ago, out of phase with the global LGM in MIS 2 (Finkel et ah, 2003; Richards et ah, 2000; Benn and Owen, 1998; Owen et ah, 2003a). Early MIS 3 was a period of intensified Indian summer monsoon (Prell and Kutzbach, 1987), but an interglacial on the Northern Hemisphere. This raises the question whether a strong monsoon alone can cause ice advance in the Himalaya and on the plateau, while the rest of the Northern hemisphere undergoes interglacial conditions. On the other hand, the North Atlantic climate events of the Younger Dryas (Tschudi et ah, 2003) and the LGM (Schäfer et ah, 2002; Benn and Owen, 1998; Owen et ah, 2000) have been recorded in Tibet, although with a smaller amplitude than in other parts of the Northern Hemisphere. In the Tanggula Shan in the arid centre of the plateau, the most extensive ice advance occurred in early MIS 6, or probably MIS 7 (Schäfer et ah, 2002), but no advance took place in MIS 3. Thus, the relative importance of north Atlantic and regional climate mechanisms in triggering ice ages in Tibet is disputed. This underlines that precise absolute dating of

109 4 Tibet

climatic events is indispensable.

Surface exposure dating (SED) is an appropriate tool in cold environments, where organic matter for radiocarbon dating is often sparse. The exposure age of a rock surface reflects the moment when the surface became exposed to cosmic radiation. In the ideal case - and this is an assumption that most SED studies make implicitly - this moment coincides with the development of the landform to be dated, e. g. with the deposition of a moraine. However, a number of processes can lead to an apparent exposure age that are younger than its real one. For a moraine boulder, these are: uncovering of the sampled block some time after its deposition; overturning after deposition; temporary shielding by snow, sediment etc.; and erosion of the rock surface. The only process that results in too old apparent ages is exposure of the surface to cosmic radiation at a different location prior to its final deposition.

In addition to the common control strategies such as sampling of several boulders from the same moraine and analysis of more than one cosmogenic nuclides with different half-lives to identify long-term pre-exposure, we collected a sample from the bottom side of nearly every boulder, largely shielded from cosmic radiation, as a control for non-cosmogenic background and pre-exposure.

4.3.2 Sampling

Glacial geological setting

The southern side of the massif around Mt. Xixabangma (8012 m a. s .1., Fig. 4.2) is char¬ acterized by high mountains and deeply incised valleys. Zheng (1988) lists five major glaciations in this region since the Pliocene (from youngest to oldest): Little Ice Age, Neoglacial, Puluo (or Qomolangma II), Fuqu (or Qomolangma I), and Nyanyaxiongla (or Nyalam). Glacial deposits are excellently conserved in the Fuqu Valley upriver from the village of Nyalam (3750 ma. s .1.). Harper and Humphrey (2003) questioned a glacial origin of the Nyanyaxiongla moraines at the type site 45 km north on Yaruxiongla Pass. Therefore, this glaciation is referred to as Nyalam in this paper.

Geomorphology

The Nyalam moraines south of Nyalam consist of a complex array of overlapping walls. They are lateral moraines from the right side of Fuqu glacier that terminates today 15 km further up the valley at 4450 m a. s.l.. The moraine ridges rise 300 metres above the present-day valley bottom. Crests are rounded and slopes are gentle. This indicates significant weathering of the landscape (Fig. 4.5). In contrast, the moraines from the younger Puluo stadium in a tributary valley in the south (Fig. 4.2; Schaefer et ah, this issue), have a clearly fresher appearance.

110 4.3 Tibet: The MIS 3 denudation event

Figure 4.5: The erratic Ny6, the largest block sampled.

Ill 4 Tibet

Sample description

Very large boulders are abundant on all three walls, well-embedded in the soil, while blocks smaller than one metre are almost exclusively found as fragments of larger ones. The sizes of the sampled erratics range from 1.5 to 4.5m height, and up to 7m length (Fig. 4.6). They are all of gneiss lithology which appears to be less resistant to erosion than, e.g., middle to fine-grained granite. Some of the large blocks have broken apart and considerable parts have fallen off. Many show weathering pits up to 30 cm diameter and 20 cm depth on the top surface. Samples for exposure dating were taken from the top of the most stable looking part of a boulder, between weathering pits if present. To rule out overturning of the erratics to the extent possible, blocks on the very crests were sampled that were deeply embedded in the soil.

4.3.3 Analytical

Moraine ages were determined by measuring cosmogenic nuclides 10Be and 21Ne in eleven moraine boulders. The principles of surface exposure dating are described, e. g., in Gosse and Phillips (2001) and Niedermann (2002). Quartz separates prepared according to Kohl and Nishiizumi (1992) and Ochs and Ivy-Ochs (1997) were analysed for 10Be and 21Ne. The geochemistry for 10Be was done at the SED laboratory at Lamont-Doherty Earth Observatory, and the measurements were carried out on the PSI/ETH tandem AMS facility at ETH Zürich (Synal et ah, 1997). Noble gas concentrations were measured with a 90° sector field static noble gas mass spectrometer at the noble gas laboratory at ETH Zürich (Beyerle et ah, 2000). This spectrometer features a modified Baur/Signer ion source equipped with a compressor that enhances the sensitivity for helium and neon by two orders of magnitude (Baur, 1999). Ages were calculated using the following sea level, high latitude production rates: ö.latg-Vr"1 10Be (Stone, 2000) and 20atg"1yr~1 21Ne (Niedermann, 2000). For the scaling of these values to the sampling location, we applied the procedure as described by Stone (2000). The use of other scaling formalisms would change the exposure ages by up to 10%, but this would alter none of the conclusions drawn in this study.

Evaluation of potential periods of pre-exposure

Measuring two nuclides with different half-lives is a common way to identify pre-exposure and non-cosmogenic contribution in a rock surface (for short referred to here as "non- in-situ"). An alternative approach is to collect a control sample that in its present position is partially shielded from cosmic irradiation, e. g. from underneath a boulder, along with a fully exposed sample that is taken for dating (see also Lifton et al. (2001) for completely shielded control samples). If the nuclide contribution from in-situ irradiation in the bottom sample can be calculated, the difference to its measured concentration

112 4.3 Tibet: The MIS 3 denudation event

Figure 4.6: The block Nyl6; the bottom sample was taken at the left hand side of the block, behind the stone wall.

113 4 Tibet

basically allows to quantify pre-exposure of the boulder and possibly to recognise other inherited contributions such as nucleogenic 21Ne (see paragraph 4.3.5). This was done for six of the eleven boulders dated in this study. These bottom samples were analysed for cosmogenic neon, but it was not possible to collect shielded samples large enough for beryllium measurement.

4.3.4 Results

Isotope data from exposed samples

Isotope data, calculated exposure ages and erosion rates are given in figures 4.7 and 4.8, tables 4.3 and 4.4, and on-line table 4.5. All ages are strict minimum ages not corrected for erosion.

In figure 4.8, the beryllium and neon ages for all Nyalam samples and their uncertainty are plotted. Four main observations are made in this figure:

1. The agreement of beryllium and neon ages is excellent. In all samples but one for which both nuclides were analysed, the ages agree well within the given uncertain¬ ties (Fig. 4.8). This is a strong indicator for high reliability of the measurements

and inconsistent with substantial non-in-situ neon contributions.

2. Minimum exposure ages from ten boulders on the three Nyalam moraine ridges range from 48 to 281 kyr (Fig. 4.8).

3. A cluster of ages exists between 48 and 65 kyr (late MIS 4 and early MIS 3).

4. The rest of the boulder ages are distributed between 100 and 300 kyr.

The ages around early MIS 3 are in agreement with datings of glaciations in other regions of the Himalaya (Finkel et ah, 2003; Richards et ah, 2000; Benn and Owen, 1998; Owen et ah, 2003a). However, also higher ages are found on the same moraines here, some even at lower elevations. This is difficult to explain with straightforward deposition of the blocks by successive glacial advances. An alternative explanation will be presented in the discussion section.

Isotope data from partly shielded control samples

The shielded samples show a neon excess over air ranging from 10 to 89% of that detected in the corresponding surface sample (Tab. 4.6). The agreement with the calculated neutron flux ratios (see discussion section) is very good for two samples, Ny8 and Nyl6, and moderate for the others.

114 4.3 Tibet: The MIS 3 denudation event

Table 4.3: Neon isotope data for all samples from the Nyalam moraines. Uncer¬ tainties are 2

errors. Errors due to uncertainties of calibration gas amounts are not included but expected to be less than 3 %.

Sample T (°C), 20Ne 21Ne/20Ne 22Ne/20Ne time (mm) (109 at g-1) (10-3)

Nyl 400, 45 3 005 ± 0 12315 5 270 ± 0 121 0 1027± 0011 600, 30 2 688 ± 0 13564 4 310 ± 0 131 0 1028 ± 0 0012 1750, 15 3 683 ± 0 17349 3 244 ± 0 064 0 1019 ± 0 0008 Ny2 600, 45 1 856 ± 0 11768 9 372 ± 0 315 0 1084±0 0024 800, 20 1 315 ± 0 05745 3 185± 0 201 0 1033± 0 0021 1750, 15 1 194 ± 0 04982 3 622± 0 222 0 1010± 0 0029 Ny3 400, 45 2 26 ± 0 0968 10 716 ± 0 391 0 1096 ± 0 0021 600, 30 2 349 ± 0 11647 6 122 ± 0 494 0 1025 ± 0 0010 1750, 15 2 411 ± 0 10533 3 092 ± 0 109 0 0996 ± 0 0013 Ny4 400, 45 1 911 ± 0 10633 8 090 ± 0 176 0 1096 ± 0 0010 600, 30 2 145 ± 0 11635 4 289 ± 0 097 0 1011 ± 0 0027 800, 30 2 318 ± 0 08554 3 187 ± 0 118 0 0983 ± 0 0016 1750, 15 3 84 ± 0 1332 3 019 ± 0 096 0 1018 ± 0 0009 Ny5 400, 45 2 334 ± 0 11702 14 474 ± 0 532 0 1119 ± 0 0034 600, 30 1 282 ± 0 08946 7 069 ± 0 418 0 1045 ± 0 0027 1750, 15 182 ± 0 1186 3 407 ± 0 272 0 1016 ± 0 0041 Ny6 400, 45 2 777 ± 0 27031 17 006 ± 0 918 0 1177 ± 0 0035 600, 30 3 071 ± 0 33113 6 525 ± 0 374 0 1064 ± 0 0037 800, 30 3 478 ± 0 26134 3 475 ± 0 395 0 1005 ± 0 0040 1750, 15 2 817 ± 0 32651 3 605 ± 0 280 0 1027 ± 0 0036 Ny7 600,45 14 456 ± 0 54968 4 141 ± 0 099 0 1033 ± 0 0010 800,20 11 014 ± 0 46542 3 360 ± 0 078 0 1019 ± 0 0014 1750,15 4 168 ± 0 16304 3 673 ± 0 153 0 1040 ± 0 0022 Ny8 400, 45 1 557 ± 0 05771 19 900 ± 0 176 0 1210 ± 0 0016 600, 30 1 5 ± 0 117 6 157 ± 0 168 0 1028 ± 0 0015 800, 30 1 808 ± 0 10024 2 976 ± 0 101 0 1016 ± 0 0017 Nyl5 400, 45 2 471 ± 0 11613 11040 ± 0 322 0 1205 ± 0 0159 600, 30 2 434 ± 0 10102 4 554 ± 0 289 0 1042 ± 0 0010 1750, 15 2 728 ± 0 11284 3 903 ± 0 072 0 1020 ± 0 0007 Nyl6 400, 45 2 029 ± 0 12887 15 736 ± 0 956 0 1151 ± 0 0028 600, 30 2 465 ± 0 20595 5 579 ± 0 240 0 1053 ± 0 0034 800, 30 15 302 ± 0 65306 3 064 ± 0 086 0 1027 ± 0 0014 1750, 15 3 84 ± 0 2082 9 105 ± 0 270 0 1000 ± 0 0029 Nyl7 600,45 27 027 ± 1 08181 4 325 ± 0 081 0 1029 ± 0 0007 800,30 2 788 ± 0 10964 3 264 ± 0 188 0 1021 ± 0 0020 1750,15 1 424 ± 0 06172 2 925 ± 0 309 0 1020 ± 0 0028 4 Tibet

0.140

D 4Q0°Q45' A 600-O3Q' 0.130 O 800*030'

y=(1.120±0.021)x+0.098 -

J i I i L 0.0040 0.0080 0.0120 0.0160 0.0200

0.112

0.108 -

0.096 -

0.0040 0.0060 0.0080

21Ne/2°Ne

Figure 4.7: Neon three-isotope plot of all Nyalam top samples. Most of the data points from 400 and 600 °C steps plot on the atmospheric / cosmogenic mixing line. This makes non-cosmogenic contribution unlikely and allows

the calculation of exposure ages. The lower panel shows the data points

within the rectangle in the upper panel.

116 4.3 Tibet: The MIS 3 denudation event

Table 4.4: Beryllium and neon minimum exposure ages.

21 Sample elevation 21NeCTC 10Be Ne age 10Be age ages corrected for erosion1 21 (ma s 1 ) (107 at g-1) (106 at g-1) (kyr) (kyr) Ne (kyr) 10Be (kyr)

Nyl 3902 1 06 ± 0 12 2 77 ± 0 11 51 1 ± 5 4 52 4 ± 2 0 53 6 ± 5 4 56 6 ± 2 1 Ny2 3902 1 19 ±0 10 2 53 ± 0 13 57 5±5 0 48 0 ± 2 4 60 7±5 0 518 ± 24 Ny3 3902 2 50 ± 0 24 5 43 ± 0 27 121 9 ± 118 104 0 ± 5 2 137 6 ± 11 9 1193 ± 53 Ny4 3840 1 27 ± 0 12 3 33 ± 0 17 62 5 ± 5 7 64 5 ± 3 3 66 3 ± 5 7 70 9 ± 3 3 Ny5 3885 3 21 ± 0 25 7 97 ± 2 79 156 6 ± 11 9 152 2 ± 5 3 184 0 ± 12 1 186 5±5 4 Ny6 3941 5 00 ± 0 73 15 12 ± 0 56 236 6 ± 34 8 280 8 ± 10 5 308 2 ± 35 9 450 0 ± 10 6 Ny7 3976 1 71±0 19 2 82 ± 0 20 812±8 9 52 5 ± 3 8 87 8±9 1 57 2 ± 3 9 Ny8 3905 3 12 ± 0 14 7 78 ± 0 34 150 4 ± 6 7 147 1 ± 6 3 175 3 ± 6 7 180 4 ± 6 4 Nyl5 3890 2 39 ± 0 19 5 60 ± 0 24 1160 ± 9 3 106 8 ± 4 5 130 1 ± 9 4 1236 ± 46 Nyl6 3893 3 24 ± 0 38 9 69 ± 0 67 155 6 ± 18 3 182 5 ± 12 6 182 5 ± 18 6 237 9 ± 12 9

Nyl7 3891 3 69±0 34 - 172 2±15 7 - 206 0±6 1 -

Be ages calculated using 5 latg-1yr-1, Ne ages calculated using 20atg-1yr-1 for sealevel and high latitude Scaling based on Lai (1991), but including m-capture fraction of f(u)=0 03 (Stone, 2000), 10Be uncertainties represent 1er analytical errors, 21Ne age uncertainties represent the propagated 2a error 21 of the Ne excess above atmospheric neon compostion 1 erosion rate of 1 mm kyr-1, see text

Calculations of ELA depressions

The depression of the equilibrium line altitude (ELA) of Fuqu Glacier during the deposi¬ tion of the moraines was calculated to obtain an independent measure for the amplitude of the glaciation. Calculations are based on Chinese 1:100,000 and Russian 1:200,000 maps, using the THAR (Toe-to-Headwall Altitude Ratio) method with a THAR of 0.4 (Meierding, 1982). This method is the most robust one if no detailed map base is avail¬ able. The present-day ELA is 5400 m a. s .1.During the Nyalam stage, it was lowered by no more than 500±100 m, i.e. the equilibrium line was at 4800 to 5000 m a. s .1.This is a smaller depression than reported from other mountain areas (e.g., the Alps in the LGM: 1300 m, Hantke (1978)) or from Pakistan (1200 m in the Hindukush, Owen et al. (2002c)). A similar depression was determined in Nepal (Duncan et ah, 1998). The ELA depression of the Puluo stage of about 200 m (see companion paper) is even smaller.

4.3.5 Discussion

Control on pre-exposure through shielded samples

The distinct grouping of four erratic ages between 48 and 65kaBP (Fig. 4.8) could be interpreted to represent the time of moraine deposition, particularly as in other regions of the Himalaya moraines with similar ages were found (Owen et ah, 2002b; Finkel et ah, 2003; Owen et ah, 2003b). The ages of all older boulders in the Nyalam moraines would then have to be explained by pre-exposure. In the following we will show that at least two of the oldest ages are not biased by non-in-situ neon contributions. This implies

117 4 Tibet

that the old boulders yield valid minimum ages for the moraines, which therefore must be considerably older than MIS 3. To find out if the cosmogenic excess found in the shielded samples originates either from in-situ exposure on the Nyalam moraines or from pre-exposure or nucleogenic pro¬ duction, we first need to estimate the expected amount of in-situ produced cosmogenic neon in the bottom sample relative to the surface sample. If the expected concentration agrees within errors with the measured concentration, substantial non-in-situ produced neon can be excluded. We used a simplified boulder geometry to calculate the secondary cosmic ray neutron flux J at the position of the shielded sample relative to the top. The mean shielding depth d due to overlying rock was estimated for 5° sectors above the horizontal plane (around one horizontal axis only, i. e. assuming infinite extension of the block along this axis). For each sector, the neutron flux J was calculated as

J(d, 9) = Jo • e~Ad svnm 9 (4.1)

d is depth from rock surface, 9 is the angle measured from the horizon, J0 is the flux in vertical direction, and A is the attenuation length (150 g cm-2). The exponent m has been taken as 2.3 (Nishiizumi et ah, 1989). The fluxes from all sectors were then summed up and normalised to the flux at an unshielded position. Finally, the total flux at the shielded location was corrected for particles backscattered from the surrounding ground. These corrections are less than 5 %. They were deduced from figure 1 in Masarik and Wieler (2003), assuming that the decrease of production rate with the column length is caused by the decreasing contribution of backscattered particles. Because the determination of shielding depths is rather uncertain, the overall error of the flux calculations is estimated to be on the order of 30 %, except For Nyl5 where the

error was assumed to be 50 % . In table 4.6, the ratio of the calculated fluxes of the shielded and the exposed sample (Jbottom/Jtop) is compared to the ratio of the measured concentrations (Cbottam/Ctap). For samples Ny 8 and Ny 16, the calculated flux ratio is identical to the measured concen¬ tration ratio within the given uncertainties (Tab. 4.6). The neon excess in the shielded samples of Ny8 and Nyl6 is hence cosmogenic and built up in situ, and substantial non- in-situ production can be ruled out for these two samples. Therefore, the neon ages of Ny8 and Nyl6 are true minima. Thus, these two samples alone prove that the Nyalam moraines were deposited clearly before MIS 3, and that the wide age range observed is not the expression of pre-exposure of numerous blocks. The bottom samples of the boulders Ny4 and Nyl5 contain more excess 21Ne than would be expected, although for Nyl5, expected flux ratios and measured nuclide concen¬ trations marginally overlap (within uncertainties, Tab. 4.6). This suggests the possible presence of cosmogenic neon from pre-exposure or nucleogenic neon (corresponding to 14±9kyr for Ny4 and 50±51 kyr for Nyl5). Note that in this case the in-situ expo¬ sure ages of the two samples would become 48 ± 16 and 66 ± 60 kyr, resp., which would not alter any conclusion drawn in this study. On the other hand, pre-exposure of, or

118 4.3 Tibet: The MIS 3 denudation event

external internal 300 - 0 Be-IOsges I • N&21ag£B MIS 8

-_ 250 - t »

MIS 7

[- 200 J

g MIS 6 150 - _ è * *{*

_ T t ,* 100 - $ MIS 5 Ny8 Nyl5 Ny16 Ny17 * CD| MIS 4 ^ 50 - — ®i©* ffi " Ny4 Ny5 Ny6 Ny7 MIS 3 Ny1 Ny2 Ny3

- MIS 2

"

Figure 4.8: Neon and beryllium exposure ages of the Nyalam erratics. Boxes repre¬ sent different walls. Ages do not show a trend from the inner (right box) to the outer (left box) walls. A cluster of ages around 50 kyr is clearly visible, while the rest of the ages are distributed uniformly between 107 and 281 kyr.

119 4 Tibet

Table 4.5: Neon isotope data for the shielded samples from Puluo and Nyalam

moraines.

22 Sample Temperature (°C), 20Ne 21Ne/20Ne Ne/20 Ne 21Necos time (mm) (109 atg-1) (106 atg-1)

Ny3 400, 45 7 21 ± 0 34 3 27 ± 0 24 0 1010 ± 0 0013 2 21 ± 2 10 600, 30 4 65 ± 0 27 3 38 ± 0 19 0 1012 ± 0 0031 1 97 ± 1 41 1750, 15 27 59 ± 1 08 3 04 ± 0 08 0 1016 ± 0 0011 2 16 ± 3 08

Ny4 400, 45 3 07 ± 0 14 4 00 ± 0 07 0 1014 ± 0 0012 3 20 ± 0 45 600, 30 2 69 ± 0 14 3 94 ± 0 05 0 1016 ± 0 0010 2 65 ± 0 48 800, 30 2 31 ± 0 11 3 83 ± 0 07 0 1020 ± 0 0011 2 02 ± 0 34 1750, 15 3 98 ± 0 16 3 37 ± 0 06 0 1031 ± 0 0008 1 64 ± 0 41

Ny6 600,45 28 ± 0 13 5 26 ± 0 15 0 1040 ± 0 0025 6 44 ± 0 65 800,20 1 8 ± 0 08 3 02 ± 0 25 0 0984 ± 0 0020 0 11 ± 0 55 1750,15 2 92 ± 0 16 3 09 ± 0 18 0 1003 ± 0 0014 0 39 ± 0 77

Ny8 400, 45 2 32 ± 0 24 5 08 ± 0 09 0 1045 ± 0 0028 4 91 ± 1 42 600, 30 2 84 ± 0 19 3 66 ± 0 31 0 1032 ± 0 0039 1 99 ± 1 28 800, 30 2 75 ± 0 40 3 19 ± 0 21 0 1019 ± 0 0041 0 65 ± 2 15 1750, 15 4 55 ± 0 25 3 03 ± 0 25 0 1018 ± 0 0025 0 33 ± 1 53

Nyl5 400, 45 6 79 ± 0 28 4 67 ± 0 20 0 1033 ± 0 0019 11 61 ± 1 64 600, 30 4 28 ± 0 22 4 30 ± 0 23 0 1021 ± 0 0023 5 73 ± 1 34 1750, 15 5 45 ± 0 22 3 65 ± 0 24 0 1009 ± 0 0029 3 79 ± 1 52

Nyl6 400, 45 4 96 ± 0 22 4 30 ± 0 22 0 1004 ± 0 0016 6 63 ± 1 34 600, 30 55 93 ± 2 15 2 90 ± 0 05 0 0972 ± 0 0009 -3 37 ± 4 44 800, 30 33 13 ± 1 15 2 97 ± 0 06 0 0990 ± 0 0010 0 52 ± 2 57 1750, 15 23 88 ± 0 97 2 99 ± 0 06 0 0982 ± 0 0014 0 83 ± 2 33

Table 4.6: Comparison of calculated neutron fluxes J and measured nuclide concen¬ trations C in shielded and exposed samples.

( Jbottom, \ (cbttom\ Sample 21Necos \ JtoP ) \ ^top J (106 at g-1) (calc)1 (meas )2

Ny3 4 17±3 51 0 47±0 14 0 17±0 18 Ny4 5 85±0 93 0 24±0 07 0 46±0 09 Ny6 6 44±0 65 0 32±0 11 0 13±0 02 Ny8 6 90±2 70 0 33±0 10 0 22±0 09 Nyl5 17 35±2 98 0 47±0 23 0 89±0 20 Nyl6 3 26±5 78 0 13±0 04 0 10±0 18

1 calculated with the model described in the text

2 for the bottom samples, the data of the same T steps as for age calculation were used, not regarding their position in the three-isotope plot

120 4.3 Tibet: The MIS 3 denudation event

nucleogenic neon in these two samples is not indicated by the excellent agreement of their 10Be and 21Ne ages. Therefore, the measured ages of Ny4 and Nyl5 are regarded as minimum exposure ages and considered for discussion. For Ny3 and Ny6, finally, the calculated flux ratio is higher than the measured one. This is most likely due to changing of the shielding situation of the bottom sample after deposition of the block. Thus, these ages are strict minima. The data presented here are the first systematic test whether nucleogenic neon and/or neon from pre-exposure of a block may be evaluated by means of partly shielded control samples. The Nyalam moraines offered well-suited erratics for such a methodological study. The test is regarded as successful because it was possible to exclude non - in - situ components unequivocally for two of the five blocks, which indicates that the exposure ages are true minimum ages. Note, however, that the cases where the agreement between modelled flux ratios and measured nuclide concentrations is not very good are probably to be explained by uncertainties in the model (in particular the rock geometry assumed) rather than by non-in-situ neon contribution because in all cases, 10Be and 21Ne ages of the unshielded sample agree well with each other.

Denudation model

As shown above, the wide age scatter and the cluster of ages around early MIS 3 are difficult to explain by pre-exposure. We therefore propose as an alternative explana¬ tion that most of these ages do not reflect the deposition of the respective blocks and moraines. Instead, we suggest a model of moraine degradation and denudation. Precip¬ itation, as the main agent of moraine degradation (Allen, 1997), can erode the surface of a moraine by tens of meters within as little as 100 kyr (Hallet and Putkonen, 1994). During this process, loose material is removed from the moraine, boulders successively emerge from the till and may become exposed to cosmic radiation and weathering well after deposition. This process is illustrated in figure 4.9. Accordingly, the exposure ages of boulders found on the surface of a moraine range from zero up to the moraine age.

We conclude that the age pattern observed on the Nyalam moraines may reflect such continuous moraine degradation, as described by Hallet and Putkonen (1994). The rate at which boulders become unburied increases at times of intensified precip¬ itation. Vice versa, several boulders unearthed within a short time interval indicate a period of intensified moraine degradation. We therefore interpret the observed age clus¬ ter between 48 and 65 ka BP to represent a phase of accelerated landscape denudation, whereas at times before and after, boulders became exhumed at lower rates. According to proxy data from the Arabian Sea, the Indian monsoon was particularly strong during MIS 3 (Prell and Kutzbach, 1987). This is in good accordance with our explanation for the age cluster.

Similar clusters of surface exposure ages in MIS 3 were described in other regions of Tibet. They were interpreted to reflect glaciations, and it was postulated that the LGM in Tibet had occurred as a local event, earlier than in the the rest of the Northern

121 4 Tibet

m . : :

2 0 B # b

Q^ °û 10 . c

• • 0 0 • 0 • 0

time of MIS 4 MIS 3 present deposition

" time

Figure 4.9: Sketch of surface evolution on a degrading moraine. Boulders succes¬ sively become unearthed as the removal of till by surface runoff pro¬

gresses. At the time some blocks reach the moraine surface, others have been exposed to weathering and cosmic irradiation for long periods. The

ages of the blocks found at the moraines surface therefore range from

zero to the age of moraine deposition.

122 4.3 Tibet: The MIS 3 denudation event

hemisphere (Finkel et ah, 2003; Richards et ah, 2000; Benn and Owen, 1998; Owen et ah, 2003a). In view of our data, however, it seems possible that these earlier data reflect the same denudation event proposed here.

Deposition of the Nyalam moraines

From the arguments listed above it becomes evident that the deposition of the Nyalam moraines took place most probably before 174±16kaBP. This is the error-weighted mean of the 21Ne and 10Be ages of Nyl6, the oldest sample for which a shielded control samples allows to rule out significant pre-exposure. This number is underpinned by three more samples whose ages are identical within errors (Ny5, Ny8 and Nyl7) and the fact that only one sample in this data set is older. The true moraine formation age may, however, be considerably older than 174 kyr, on the one hand because this minimum age neglects any erosion of the boulder (see next paragraph), but also because it is far from certain that the oldest boulders in this data set have been exposed at the moraine surface since its deposition. In fact, the only boulder in the data set that is older than Nyl6, Ny6, points towards deposition earlier than 174 kyr ago. Its shielded sample does not indicate pre-exposure, but temporary shielding, making the 21Ne age of 237 ± 35 ka an absolute minimum age. Temporary shielding, on the other hand, is not in agreement with the 10Be age that is higher than the 21Ne age, although the uncertainties of these numbers show a slight overlap. For this reason, we do not use Ny6 for the discussion of the moraine ages, but to stress the minimum character of the Nyalam moraine age of 174 ka.

Sensitivity of exposure ages to erosion rate

Exposure ages discussed so far neglect the effect of erosion. In the following, the sensi¬ tivity of exposure ages to erosion rates will be illustrated.

The Nyalam data fall into the area of the 21Ne/10Be vs. 10Be plot ("erosion island plot") (Graf et ah, 1991) where all erosion rate curves converge and thus the plot is not sensitive to changes in erosion rate. However, a maximum erosion rate can be calculated from the concentration of a cosmogenic nuclide assuming infinite exposure.

Maximum erosion rates for the Nyalam samples range from 2 to 11 mm kyr-h This is consistent with erosion rates of 2 to 19 mm kyr-1 determined for granitic surfaces in alpine environments (Small et ah, 1997). For illustrative purposes, let us conservatively assume a moderate erosion rate of 1mmkyr-1. The age of Nyl6 then increases by 30 to 50 kyr. Note that applying a higher erosion rate leads to higher age increases. This again illustrates that the moraine age of 174 ± 16 kyr given above is a very conservative minimum value.

123 4 Tibet

Triggering mechanisms of Tibetan glaciations

We have hypothesised that during early MIS 3, when the Indian summer monsoon was considerably stronger than today, no major glaciation has occurred in the monsoonal Nyalam region. This would imply that an intensified summer monsoon alone does not cause significant ice advances in the Himalaya. The limited influence of the monsoon is underpinned by the low ELA depression (and hence the moderate extent) of even the largest ice advance in this monsoon-dominated area. This may have two main reasons. First, the monsoon is a seasonal circulation. It delivers moisture to Tibet in summer time, when most precipitation falls as rain even at high elevations, and thus snow accumulation is limited. In winter time, during the accumulation season of other mountain ranges, the dry winter monsoon winds prevail in Tibet, and accumulation is limited again. Second, the distance to the Arabian Sea, where the Indian Ocean summer monsoon winds originate, is very large, about 2000 km, and the air masses are relatively dry when they reach the Himalaya. Other mountain ranges such as the Alps, the Rocky

Mountains or the Andes are located much closer to their moisture source.

If the monsoon plays a minor role in driving glaciations in the southern Himalaya, it is expected that its influence is even smaller in the arid centre of the Tibetan plateau. There, moraines have been dated to more than 170kyr (Schäfer et ah, 2002) which is in agreement with the moraine age of around 174 kyr presented here. This points toward the North Atlantic climate as a major agent governing Tibetan climate. Strong support for this hypothesis comes from a second set of moraines near Nyalam, as discussed in a companion paper (Schaefer et ah, this issue).

4.3.6 Conclusions

Cosmogenic nuclides 10Be and 21Ne show that Nyalam stage moraines in monsoonal Nyalam County, Tibetan Himalaya, have been deposited at least 174 ± 16 ka BP. Glaciers did not reach a similarly large extension after. The cosmogenic neon content of partially shielded control samples from the dated boulders was successfully used to prove that pre-exposure of at least two surfaces is absent.

The scatter of ages between 100 kyr and 300 kyr is interpreted to be the result of suc¬ cessive exhumation of blocks from the moraines, and not of multiple glacial advances. A "denudation event" caused by an intensified Indian summer monsoon in MIS 3 acceler¬ ated exhumation and led to an accumulation of low boulder ages between 48 and 65 kyr. The last glacial maximum in Tibet thus probably has not occurred as a "local LGM" in MIS 3 that predated the ice sheet expansion in the Northern Hemisphere. Instead, a limited glacial advance during the "classic" LGM in MIS 2 is indicated in Nyalam (Schaefer et ah, this issue). The absence of a glacial advance during the MIS 3 phase of intensified Indian summer monsoon suggests that the monsoon alone is not a driving force of glaciations in the Tibetan Himalaya, and consequently on the Tibetan Plateau. The Tibetan glaciers thus

124 4.3 Tibet: The MIS 3 denudation event

appear to respond rather to changes in temperature than in precipitation. The relatively moderate amplitude of the glaciation can be explained with the seasonal character of the monsoon that brings moisture in summer, and with the large distance to the moisture source of the Indian Ocean. It is not clear, however, what caused the observed ice advance prior to 174 ky BP. To better assess the influence of North Atlantic climate, the age of this glaciation must be better confined, and its spatial distribution further investigated. To better understand the role of the monsoon, its variations before 150kaBP ought to be known.

Acknowledgements

This work was supported by the Swiss National Science Foundation and the Climate Center of Lamont-Doherty Earth Observatory. JS thanks the Gary Comer Science and Education Foundation for support.

125 4 Tibet

126 5 Exposure dating test studies from Europe

127 5 Europe

5.1 An attempt to date the Saale glaciation in Lower Saxony

5.1.1 Introduction

One of the largest ice sheets of the world, and the largest on the European continent, during the ice ages is the Scandinavian Ice Sheet. At its maximum extent, it reached as far south as 50 °N. Similar to the Alps, this maximum was not reached again at the Last Glacial Maximum (LGM), at least not on the European mainland (Smed and Ehlers, 2002). The extent of Scandinavian ice in Central Europe is marked by large erratic blocks, accumulations of sand, deposited at the former margin of the ice sheet, and the typical and widespread Old Moraine Topography (Smed and Ehlers, 2002), characterised by moraines that display a very smooth topography due to the long-term action of solifluction processes.

While the extent of LGM and - to a lesser degree - Pre-LGM glaciers into the foreland of the Alps and of the Scandinavian ice sheet into Northern Germany is fairly well known and mapped (e.g.Andersen (1981)), its timing is still contentious. It must be assumed that an ice mass so much bigger than that covering the Alps responses differently to climatic fluctuations. There might have been a time lag between the advance of the Alpine glaciers and that of the Scandinavian ice sheets due to the inertia of the vast ice mass of the latter, but it is not clear how large it was. Since former ice margin positions on the continents directly reflect global climate oscillations, they are the key to deciphering the time structure of past climate changes. This is particularly true for the Scandinavian ice, as it was located directly in downwind direction of the North Atlantic and therefore is an excellent climate archive for changes in this important area.

The chronology of the older glaciations is less constrained than that of the LGM. The fundamental task is to differentiate in time the LGM from the older cold climate events, i.e. to define the last interglacial. A very suitable area to study the southern margin of the Scandinavian ice sheet is Northern Germany where moraines and erratics of both LGM and Pre-LGM age are mapped (e.g. Eissmann (1997), and references therein).

Three glacials are known for Northern Germany, these are, from oldest to youngest, the Elster (480 to 430ka BP), Saale (240 to 180ka BP) and Weichsel (120 to 10ka BP) glaciations (Eissmann, 1997). Whereas in the Eastern part of Central Europe, the Elster glacial was the most extensive, Saale ice reached further south in the West (Eissmann, 1997). For this reason, Northern Germany was selected as the test area for a pilot study to numerically date the Pre-LGM ice advance of the Scandinavian Ice Sheet. Due to the low production rates and presumably strong erosion at this latitude and altitude, cosmogenic excesses are often difficult to detect.

128 5.1 Lower Saxony

5.1.2 Sampling

Erratics are not hard to find in Northern Germany. Generations of Geologists have searched them, catalogued them, and traced back their journey from Scandinavia using their lithology. Many of these erratics are tourist attractions today and under protection. We had the opportunity to sample two of them for exposure dating:

- Giebichenstein in Nienburg, the largest erratic of Lower Saxony (Fig. 5.1). It is 7.5 x 4.5 m wide and 2.8 m high (Meyer, 1999).

- Erratic from Ventschau, 4x4m wide and about 3 m high.

Both are of granitic lithology. As they are well-embedded in the soil and located on ground frequented by the public for leisure, it was not possible to take shielded samples. The analysis of the Ventschau erratic has not been carried out yet. Results are only available for the Giebichenstein.

5.1.3 Analytical

The sample Gie from Giebichenstein was analysed for 21Ne and 10Be. Quartz was sepa¬ rated with physical methods as described in section 2.4 for noble gas analysis (separate Giel). For the radionuclides, extraction of beryllium followed the procedure of Kohl and Nishiizumi (1992), modified after Ivy-Ochs (1996), and was done by Susan Ivy-Ochs at the Institute of Particle Physics at ETH (separate Gie2). The latter separate was also used for a second neon analysis. All measurements were carried out as described in chapter 2.

5.1.4 Results and Discussion

Results are given in tables 5.1, 5.2 and Fig. 5.2. Cosmogenic neon is expected to be released from Quartz below 600 °C (Niedermann et ah, 1993). According to Fig. 5.2, neither of the 400°C steps from Giel and Gie2 is cosmogenic when an atmospheric composition of trapped neon is assumed. Instead, the 1750 °C steps are cosmogenic, with a considerable neon excess over air of 43 % from Giel. The additional steps at 800 and 1000 °C from Gie2 have no significant neon excess.

It seems therefore advisable not to use the 400 °C step for age calculation. The 1750 °C step, however, is dubious as well, because at this temperature, trapped 21Ne is expected to leave the mineral, and a trapped neon component of atmospheric composition indeed is probable. A conservative interpretation would be to use the minimum exposure age from the 600 °C step, which is 237 ± 33 kyr for Giel and 391 ± 51 kyr for Gie2. The 10Be age is 103 ±4 kyr, which does not support the neon data and gives no indication which temperature steps represent the in-situ cosmogenic fraction in the neon data. A large difference between neon and beryllium ages of an erratic can have the following reasons (compare section 2.2.1):

129 5 Europe

Figure 5.1: The Giebichenstein near Nienburg .

130 5.1 Lower Saxony

Table 5.1: Neon isotope data for Giebichenstein. Giel and Gie2 are from the same sample, but quartz for Giel was separated using physical meth¬ ods, whereas Gie2 was separated chemically for Be analysis (see text). Uncertainties are 2

ination errors. Errors due to uncertainties of calibration gas amounts are not included but expected to be less than 3%.

Sample (°C, min) 20Ne 21Ne/20Ne 22Ne/20Ne 21Ne (109 at g"1) (10"3 ) (106 at g"1)

Giel (400,45') 2.554 ± 0.042 6.650 ± 0.379 0.1021 ± 0.0020 16.985 ± 1.008 Giel (600,45') 2.992 ± 0.029 4.644 ± 0.200 0.1033 ± 0.0026 13.894 ± 0.613 Giel (1750,15') 1.972 ± 0.028 5.204 ± 0.387 0.1024 ± 0.0028 10.262 ± 0.777

Gie2 (400,45') 1.319 ± 0.038 8.802 ± 0.266 0.1030 ± 0.0027 11.608 ± 0.485 Gie2 (600,45') 1.472 ± 0.072 8.617 ± 0.395 0.1056 ± 0.0024 12.688 ± 0.850 Gie2 (800,30') 1.106 ± 0.028 3.198 ± 0.244 0.0997 ± 0.0036 3.535 ± 0.284 Gie2 (1000,30') 0.689 ± 0.248 3.506 ± 0.148 0.1020 ± 0.0042 2.416 ± 0.876 Gie2 (1750,15') 0.751 ± 0.262 4.738 ± 0.166 0.1034 ± 0.0030 3.559 ± 1.246

Table 5.2: Beryllium and neon data for Giebichenstein. Neea;c is excess above at¬ mospheric composition.

Sample 10Be 21Neea;c 21Neea;c 21Ne Minimum Minimum (°C, min) (105 at g"1) (106 at g"1) (%) released 10Be age 21Ne age (%) (kyr) (kyr)

Giel (400,45') 9.43 ± 1.15 56 50 443 ± 54 Giel (600,45') 5.04 ± 0.71 36 27 237 ± 33 Giel (1750,15') 4.43 ± 0.86 43 23 208 ± 41

Gie2 (400,45') 7.71 ± 0.62 66 43 362 ± 29 Gie2 (600,45') 8.33 ± 1.08 66 46 391 ± 51 Gie2 (800,30') 0.26 ± 0.37 8 1 12 ± 15 Gie2 (1000,30') 0.38 ± 1.61 16 2 18 ± 76 Gie2 (1750,15') 1.34 ± 2.02 38 8 63 ± 95

Gie2 59.21±2.01 — — 103±4

Be-ages calculated using 5.1 atg yr, Ne-ages calculated using 20atg yr for sealevel and high latitude. Scaling based on Lai (1991), but including muon-capture fraction of f(/x)=0.03 (Stone, 2000);

10Be-uncertainties represent 1er analytical errors; 21Ne-age uncertainties represent the propagated 2a- error of the 21Ne-excess above atmospheric Ne-istopic compostion.

131 5 Europe

0.108 1 1 1

- Giebichenstein 0.106 ^~-^-

F 0.104 600 °C

750 °C | 0.102 si : *^ AIR 0.100

y=1.120x+0.098 .

0.098 i I 0.0040 0.0060 0.0080 21Ne/20Ne

Figure 5.2: Neon three isotope plot for Giebichenstein with the cosmogenic- atmospheric mixing line. Open squares: Giel; filled squares: Gie2. 5.1 Lower Saxony

- pre-exposure of the surface

- post-depositional overturning of the block

- trapped neon

- erosion

- temporary shielding

As the Scandinavian Ice Sheet was wet-based at its margins, i.e., where the entrain- ment of the European erratics took place, basal erosion would most probably have wiped out any cosmogenic nuclides from pre-exposure before the glacial transport. Post- depositional overturning can be ruled out by the large size of the boulder. As erosion affects both isotopes, it would not cause such a large difference between neon and beryl¬ lium ages. If the block was buried for a substantial amount of time, a considerable portion of 10Be had decayed, indeed leading to a lower 10Be than 21Ne age. However, the required period of burial would be more than 1.5 Myr, which is far more than the age of the oldest glaciation known in Northern Germany. The most probable explana¬ tion for the differences between beryllium and neon ages is therefore a non-atmospheric composition of trapped neon. The only way to test this hypothesis is on-line crushing and analysing the released gas (see chapter 2 and section 5.2 below). From the data presented here, a minimum age for the deposition of Giebichenstein of 103 ±4 kyr can be inferred, together with the indication that the real age can be considerably older, as indicated by the neon data.

5.1.5 Outlook

The neon data from the pilot study of a sample from Giebichenstein are not fully under¬ stood. By crushing in-vacuo, the composition of the trapped noble gas fraction can be determined and used for a reliable age calculation. Hopefully, this will ultimately reveal whether the high neon age is the result of trapped neon or of substantial temporary shielding. Analysis of the second test sample from Ventschau will, together with an improved dating of Giebichenstein, yield valuable information on the glacial history of lower Sax¬ ony.

133 5 Europe

5.2 The landslide of Kofels: An attempt to calibrate the 21 Ne production rate

Not only the scaling procedure of nuclide production rates to the geomagnetic latitude and elevation of the sampling site is a much debated issue. The absolute production rate, usually expressed for sea-level and high latitude, is under discussion, too. Basically, three different approaches to determine production rates exist: Physical modelling, exposing artificial targets, or measuring samples from independently dated surfaces. The first two approaches are beyond the scope of this thesis, but appropriate samples were available for the third one.

If the exposure age T of a surface can be determined independently and the concentra¬ tion N is measured, the equation T=N/P (eq. 2.3) for stable nuclides can be resolved for the production rate P (the same is true for radionuclides, with a different mathematical formalism, see chapter 2). Appropriate calibration sites are difficult to find, however. The lithology of the surface has to be resistant to erosion and suitable for the measure¬ ment of as many cosmogenically produced nuclides as possible. The exposure history must be doubtlessly simple. And, of course, the age of the surface must be reliably and precisely known from an independent dating method. Nishiizumi et al. (1989) have calculated production rates for 10Be and 26A1 in glacially polished bedrock from the Sierra Nevada. Niedermann et al. (1994) have determined the 21Ne production rate from aliquots of the same samples. As the age of the respective glaciation was re-calculated from Ilka to 13ka, the production rates had to be re¬ evaluated as well (Niedermann, 2000). Kubik et al. (1998) has calculated production rates for 10Be and 26A1 from the large, radiocarbon-dated landslide in Kofels (Austria), and recently revised his results (Kubik and Ivy-Ochs, tted). Many studies have been dedicated to the 3He production rate in olivine, as fresh basalt lava flows make good calibration sites. Only two papers deal with a production rate calibration for pyroxene, however (Cerling and Craig, 1994b; Cerling, 1990).

In order to obtain a 21Ne production rate calibration for mid-latitudes and interme¬ diate elevations (1680 m a. s.l.), aliquots of the above mentioned Kofels samples were analysed for cosmogenic neon. It was expected that, because of the geographic location of Kofels and the relatively low age of the landslide (9800 cal. 14C a determined in buried wood, Ivy-Ochs et al. (1998)), the cosmogenic excess over atmosphere would be small and difficult to detect.

A first attempt was made with a stepwise heating procedure. The sample KÖ5 (Kubik et ah, 1998) was crushed off-line to 100 //m and loaded to the extraction line. It was then heated to 400 °C, 600 °C, 800 °C and 1750 °C. The gas amounts released are clearly higher than the machine blank, but none of the temperature steps shows a significant excess over atmospheric composition (Tab. 5.3). Only the 21Ne/20Ne and 22Ne/20Ne ratios of the 1750 °C step are off the atmospheric composition in the neon three isotope plot, but the data plot below the atmospheric-cosmogenic mixing line, indicating a weak

134 5.2 The landslide of Kofels

Table 5.3: Crusher data

*>We 21Ne 21Ne 21Ne/20Ne 22Ne/20Ne exc e (109 at g"1) (lu"3) (106 at g"1) (%)

Stepwise heating KÖ5 (400,45) 18.85 ± 0.11 2.910 ± 0.045 0.1009 ± 0.0005 -0.91 ± 1.21 T7 KÖ5 (600,30) 42.09 ± 0.15 3.009 ± 0.029 0.1010 ± 0.0004 2.11 ± 1.75 1.7 KÖ5 (800,30) 16.64 ± 0.05 2.913 ± 0.055 0.1021 ± 0.0006 -0.76 ± 1.08 -1.6 KÖ5 (1750,15) 4.88 ± 0.06 3.131 ± 0.099 0.1004 ± 0.0006 0.84 ± 0.69 5.5 On-line crushing Heat blank (580,45) 235.75 ± 3.62 2.990 ± 0.143 0.1008 ± 0.0017 7.38 ± 46.03 1.0 Crusher blank (20,15) 59.22 ± 1.66 2.825 ± 0.474 0.1055 ± 0.0050 -7.94 ± 33.37 -4.7 Crush 1 (20,10) 22.41 ± 0.16 2.953 ± 0.098 0.1017 ± 0.0012 -0.12 ± 2.74 -0.2 Crush 2 (20,10) 7.93 ± 0.10 2.984 ± 0.071 0.1005 ± 0.0016 0.20 ± 0.94 0.9 Crush 3 (20,10) 5.97 ± 0.08 2.903 ± 0.136 0.1028 ± 0.0020 -0.34 ± 1.09 -1.9 Heat (580,45) 43.64 ± 0.17 2.998 ± 0.056 0.1016 ± 0.0004 1.71 ± 2.98 1.3

nucleogenic contribution. The goal of a second experiment therefore was to lower the amount of atmospheric contamination by reducing the surface area of the sample exposed to the atmosphere before loading it to the high vacuum. For this aim, an aliquot of the same sample KÖ5 was ground to 315//m instead of 100 /im, and the crushing to release nucleogenic neon was done in the UHV extraction line (see paragraph 2.6.2). This procedure was repeated two more times. The 21Ne/20Ne and 22Ne/20Ne ratios of the gas released by crushing were atmospheric in all three crush steps. In order to release cosmogenic neon, the sampling chamber was then heated to 580 °C for 45 minutes by a hot air gun. Again, the results show no cosmogenic imprint: the isotopic ratios are atmospheric within 2

135 5 Europe

D stepwise heating KÖ5 crusher heat step 0.104 online crushing

0

0.102

0

0.100

0.098 _L _L _L 0.0028 0.0030 0.0032 21Ne/20Ne

Figure 5.3: Neon three isotope plot for both experiments with sample KÖ5.

136 6 Conclusions and Outlook

137 6 Conclusions

6.1 Conclusions

Antarctica In four studies from the Ross Sea section of Antarctica, evidence from glacially polished bedrock, glacial drift sheets and periglacial bedrock was dis¬ cussed. This evidence comes from two main regions: The Dry Valleys and Northern Victoria Land. The studies from Northern Victoria Land represent an important step outside the much-discussed Dry Valleys. The goal was to extend the knowl¬ edge gained from the Dry Valleys and to test if it is valid in larger areas of the continent.

- In the Dry Valleys, data was produced that confirm and extend findings of previous studies. 14 samples related to buried ice relics in Beacon Valley were analysed for 3He and 21Ne (section 3.2). Their ages confirm that the ice is of Pliocene age and that it has continuously thinned since at least 2.6 Ma. 3He and 21Ne ages from eight boulders in Vernier Valley testify for a simple deposition history of three parallel lateral moraines besides Ferrar Glacier (section 3.3). The ice level change was only 125m within the last 3.4Ma, the thinning continuous, and the related ice surface change of Taylor dome was

even as small as 50 m.

- The first exposure ages in the region of Terra Nova Bay were determined in Deep Freeze Range near the Ross Sea coast (section 3.4). Erratics and glacially polished bedrock samples were analysed for 3He, 10Be and 21Ne.

This mountain range is ice-free since at least 5 Ma BP. Ice surface elevation in the Pleistocene was around 780 m higher than at present. A set of 14 bedrock samples from nunataks at the inner side of the Transantarc¬ tic mountains show that the ice-level variations there have been much smaller

than near the coast, less than 250m in the last 3.6Myr (section 3.5).

It is concluded that the main findings from the Dry Valleys - long-term climate stability and very restricted ice-volume changes of the East Antarctic Ice Sheet

since at least the lower Pliocene - are indeed valid for the entire Ross Sea section of the Transantarctic Mountains, and most probably for the whole continent.

Tibet The main goal from a study of valley moraines in the monsoonal Tibetan Hi¬ malaya was to clarify the relative importance of the Indian Summer Monsoon and the North Atlantic climate in triggering Tibetan glaciations. Samples were collected on five moraine ridges at two locations near Nyalam south of Mount Shishabangma.

- Three large moraine ridges yielded 10Be and 21Ne ages from ten erratics, ranging from 48 to 280 kyr (section 4.3). They were most likely deposited before 186 ka BP, and moraine degradation continuously unearthed erratics from the till. In early MIS 3, the intensified monsoon enhanced denudation,

138 6.2 Outlook

and an uncommonly high number of boulders was washed out within twenty

thousand years.

- A second set of moraines was dated to be 15 to 24 and around 11 kyr old, con¬ sistent with the North Atlantic climate signals of the Last Glacial Maximum and the Younger Dryas (section 4.2).

Together, these two studies indicate a minor role of the monsoon in causing glacier advances in Tibet. Instead, they suggest synchroneity to the glaciations in the rest of the Northern Hemisphere.

Surface exposure dating Surface exposure dating has repeatedly proven to be an ex¬ cellently suited and versatile means of dating glacial landforms at high latitude or high elevations.

- The best use can be made of the method by measuring more than one nuclide,

particularly a noble gas and a radionuclide. This has been demonstrated with the determination of erosion rates for granite bedrock in Antarctica using the erosion island plot (section 3.4); with the evaluation of pre-exposure along with the dating of erratic under difficult conditions in Tibet (section 4.3).

- With a large sample set of dolerite rocks it was shown that pyroxenes are not necessarily retentive for cosmogenic helium (section 3.5). One parameter that determines the retentivity of helium and the degassing pattern of neon is the Mg content of the mineral. Whereas cosmogenic and non-cosmogenic neon Mg-rich pyroxenes can be efficiently separated by heating the sample to 1500

°C, this is not possible with low-Mg pyroxenes. It is therefore advisable to analyse samples of one type of pyroxene and determine which one it is prior to the mass spectrometric analysis.

- Dating with cosmogenic noble gases has a lower temporal limit. Even at 4000 ma. s .1. where production rates are high, surfaces exposed for less than 50kyr

are difficult to date with sufficient accuracy if the conditions are not optimal (section 4.3). Atmospheric neon contribution is a serious problem for samples of comparably low exposure ages. On the other hand, there does not seem to

be an upper end to the time range of applications, except the availability of geological samples.

- In an environment of moderate to high erosion and weathering rates or high

uplift, a series of exposure ages from the same landform have to be regarded in context (section 4.3). The interpretation of single samples can be misleading.

6.2 Outlook

In this thesis, paleochmatic interpretations based on evidence from different localities have been made. Some of these localities have not been investigated before with absolute

139 6 Conclusions

dating methods, and some interpretations revise present views. For both cases, of course, is it is desirable to obtain a confirmation of the conclusions drawn here that is supported by a broader data base. Therefore, more studies with similar goals as those presented here are necessary. In particular, this is true for the concept of a denudation event in MIS 3 in Tibet, as well as for the identification of MIS 2 cold phases in that area. The analysis of samples from the vicinity of Nyalam that presumably have that age is in progress. It will also be worth to re-examine data from different parts of the Himalayas and the Tibetan Plateau under the light of a possible MIS 3 denudation phase. This hopefully will make the overall picture of local glacial chronologies in Tibet more homogeneous and improve our knowledge about the relative influence of westerly climate signals and the monsoon on the climate system in southeast Asia. In Northern Victoria Land, Antarctica, widespread glacial drifts await their dating.

The correlation of different occurrences of these drift sheet among each other and to those in the Dry Valleys will supply important insights into the glacial chronology of the Pleistocene, a period that in other parts of the Earth was characterised by strong climate variations. First samples from these drifts are collected, some are analysed, but more have to be done for a concise interpretation.

As the acceptance for the method of surface exposure dating has grown within the Earth scientists community, the applications have become more and more advanced.

This has led to an increased demand for accuracy of age determinations. The crucial parameter in this context is the production rate of cosmogenic nuclides at the Earth's surface and its spatial variation. Shortcomings in the theoretical and empirical foun¬ dation for calculating accurate production rates affect accuracy and applicability of the method. A substantial improvement of this situation is expected from the Europe and US-based CRONUS (Cosmic-Ray produced NUclide Systematics) project that attacks at the same time the geological calibration of production rates, modelling efforts and physical experiments. For cosmogenic noble gases in particular, it should be interest¬ ing to follow the establishment of additional cosmogenic nuclides for exposure dating. Furthermore, the techniques of separating cosmogenic from non-cosmogenic noble gas fractions ought to be refined. This will provide access to the dating of samples that have been exposed for shorter durations or at low production rates.

140 Dank

Zum Schluss möchte ich allen Leuten von Herzen danken, die mir in den letzten vier Jahren geholfen haben. Das waren ziemlich viele. Einige haben das absichtlich gemacht, andere wohl eher unbewusst. Ein paar Namen möchte ich erwähnen:

Zuerst Ramer Wieler. Er schafft es immer, seine Doktoranden so zu kritisieren, dass sie sich ernst genommen fühlen und die Gewissheit bekommen, dass ihre Arbeit etwas taugt. Man hat dann das Gefühl, dass man nicht seinem Professor gegenüber sitzt, sondern seinem Mit-Autor, und das ist sehr motivierend. Ich danke ihm auch für seine verständnisvolle und entgegenkommende Art als Vorgesetzter. Ein Makel an Rainer ist aber leider, dass er die ZSC Hamsters moralisch und durch seine gelegentliche Anwesen¬ heit im Hallenstadion unterstützt. Christian Schlüchter möchte ich für viele Einsichten über Quartärgeologie und Feldar¬ beit sowie für eine grossartige Reise nach Tibet danken. Es ist nicht immer einfach, einen Termin bei ihm zu bekommen, aber wenn man das geschafft hat, ist er voll und ganz bei der Sache. Ausserdem beherrscht er den passenden Wechsel zwischen Spannung und Entspannung bestens, wovon man sich auf jeder Exkursion überzeugen kann. Jörg Schäfer hat wohl alle Aspekte meiner Arbeit am meisten geprägt, weil er mich in die ganze Methodik eingeführt hat, von der Probennahme bis zur Präsentation an einer Konferenz, und mir unterwegs unzählige Male gedanklich auf die Sprünge geholfen hat. Seine Tipps zum Networking im Labor und an Tagungen, aber auch die Stunden in der Beiz, waren sehr wertvoll für mich.

Da Carlo Baroni ho imparato molto sul lavoro nel campo, la geologia glaciale e la cultura italiana. II suo modo di avvicinarsi all'Antartide con curiositä e con rispetto e di vederla corne un luogo molto speciale e non solo come oggetto scientifico mi ha impressionato profondamente. Grazie mille e "In bocca al lupo"!

Ich bedanke mich herzlich bei Philip Allen, der bereitwillig das Koreferat übernommen hat. Er hat gegen Ende der Dissertation neue Sichtweisen auf meine Ergebnisse einge¬ bracht. Das hat besonders das Tibet-Kapitel beeinflusst. Hein Baur hat eine scheinbar endlose Geduld mit den Benutzern seiner Spektrometer. Er hat es verstanden, meinen anfänglichen Schiss vor Tom Dooley in gesunden Respekt umzuwandeln, indem er mir in seiner ruhigen und humorvollen Art alles erklärte, auch mehrmals (bloss wusste er immer genau, was er mir schon erklärt hatte). Dafür danke ich ihm herzlich. Im Rückblick kann ich dank Heiri sagen: Massenspektrometrie ist gar nicht so schwer. Das Potentialfeld in der Ionenquelle ist nämlich eine Weide, die Ionen sind Kühe, die muss man nur noch aus dem Stall rausscheuchen, und wenn sie draussen sind, rennen sie immer bergab. Brut Meyer wurde einmal sehr treffend als die

Geheimwaffe unseres Institutes bezeichnet. Sie weiss einfach über alle administrativen

Belange Bescheid und organisiert neben ihrer ganzen Arbeit auch noch gesellschaftliche

Anlässe am Institut. Vielen Dank! Die Zweiradfreaks Urs Menet und Andreas Süsh

141 waren immer zur Stelle, wenn es um kreative Lösungen für mechanische Probleme oder extravagante Sonderwünsche an der Extraktionslinie ging. Bruno Rutsche und Do- nat Niederer halfen mir im Kampf gegen Bill Gates und Kabelsalat, sowie mit träfen Sprüchen in der Kaffeepause. RoKi sei gedankt für viel Wissenswertes über Helium, und seinen generösen Umgang mit Maschinenzeit.

Von Florian Kober konnte ich nicht nur dank seinen Abonnements des deutschen "Na¬ tional Geographie" und der NZZ viel lernen. Obwohl er zeitlich ein halbes Jahr hinter mir lag - mit der Diss und mit der Familienplanung - war er mir gedanklich nicht selten einen Schritt voraus. Dass es heikel ist, im NO-Gebäude Scooter - Rennen zu fahren, ahnte ich allerdings schon vor ihm. Ingo Leya danke ich für manche (eigentlich viel zuviele) spannende Stunde an der Eichgas-Abfüll-Anlage. Zusammen haben wir zahlreiche Varianten des Scheiterns ausgelotet. Ausserdem ist Ingo ein geduldiger Nach¬ hilfe-Lehrer in Physik und ein unterhaltsamer Mensagänger. Ich bin froh, dass er mir das IGMR erst nach mir verlassen hat! Obwohl ich das viel zu selten genutzt habe, war Susan Ivy- Ochs immer da für Fragen aller Art, und dafür bin ich ihr sehr dankbar. Su¬ san hat auch viele meiner Beryllium-Proben aufbereitet und alle zur Messung gebracht, obwohl sie genug zu tun hatte ohne mich. Schliesslich hat sie auch noch bewiesen, dass Wissenschaft und Familie zusammen gehen können. Mit Matthias Brennwald konnte ich jederzeit über das Doktorandendasein schimpfen, wobei es weniger ums Klagen (wir wussten ja um die Annehmlichkeiten unserer Situation) als um dessen formale Varia¬ tionen ging . Adi Gull, Martin Burkhalter und Pascal Hägeh danke ich fürs Baden,

Skifahren, Lachen und Fernsehen ganz herzlich. Urs Joss danke ich sehr für seine kom¬ petente und unendlich geduldige Unterstützung mit DTßX, und der Leser sollte das auch tun. Dank ihm ist dies hier ein lesbarer Text und nicht eine Aneinanderreihung von Ze¬ ichen. Veronika Heber, Philipp Heck, Ansgar Grimberg danke ich für ein angenehmes Klima im Büro, viele Gefallen und kreative Hilfe mit meinen zahlreichen Computer- und administrativen Problemen. Nadia Vogel hat meine Deutschkenntnisse beträchtlich erweitert, obwohl ich glaubte, Deutsch sei meine Muttersprache. Ausserdem fällt einem das Arbeiten immer leichter, wenn man eine so fröhliche Person im Rücken hat. Meinem Nachdoktoranden Stefan Strasky danke ich für seine tatkräftige Unterstützung bei vielen Abschlussarbeiten und wünsche ihm viel Spass mit seinem Doktorat.

Bei meiner ganzen Familie möchte ich mich bedanken für die ständige Unterstützung während meinen endlosen Schul- und Studienjahren; aus diesem Büchlein erfahrt ihr nun, dass ich mich nicht mit Katzenstreu beschäftigt habe.

Die ersten sind hier wieder mal die letzten: Nicole und unsere Tochter Elm haben mich immer wieder auf die richtige Distanz zu meiner Arbeit gebracht, indem sie mich gezwun¬ gen haben, irgendwann mit Arbeiten aufzuhören, auch wenn ich noch nicht fertig war (im einen oder anderen Sinn des Wortes). Nicole hat mich immer wieder neu motiviert und in der Schlussphase meine familiären Pflichten grösstenteils übernommen. Ohne sie wäre diese Arbeit nicht möglich gewesen.

142 Curriculum Vitae

Angaben zur Person

Name Peter Oberholzer

Geburtsdatum 14. August 1969

Geburtsort Uster, Schweiz

Bildungsgang

2000--2004 Doktorat am Institut für Isotopengeologie und Mineralische Rohstoffe, ETH Zürich 1993--1998 Fachstudium am Departement Erdwissenschaften, ETH Zürich 1985--1990 Bündner Lehrerseminar, Chur 1982--1985 Sekundärschule, Chur 1980--1982 Primarschule, Chur 1976--1980 Primarschule, Pfäffikon ZH

Berufliche Tätigkeit

1990-1993 Primarlehrer, Sils i. D. 6 Conclusions

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154 A Appendix

155 A Appendix

A.l Additional samples

A.1.1 Helium data

Table A.l: Helium isotope data of samples not listed in the main part of this the¬

sis. Errors are 2 a and include uncertainties in mass discrimination and

sensitivity; uncertainties due to calibration gas concentrations are not included but expected to be smaller than 3 %.

Sample (°C, mm) 4He 3 He 3 He/4 He R/RA (1012 at g"1) (106 at g-1) (io-6

NZ19 (600,30') 0 836 ± 0 027 0 32 ± 0 03 0 38 ± 0 03 0 28 ± 0 02 MEX1 Px (850,30) 0 000 ± 0 034 2 12 ± 0 13 SULA438 (600,45') 7 898 ± 0 003 114 66 ± 2 33 14 52 ± 0 85 10 49 ± 0 61 SULA438(1500,15') 2 512 ± 0 003 0 20 ± 0 10 0 08 ± 0 04 0 06 ± 0 03 SULA 245 (600,45') 22 923 ± 0 005 11 32 ± 0 83 0 49 ± 0 05 0 36 ± 0 03 971210 01 (600,45') 4 166 ± 0 013 1 13 ± 0 08 0 27 ± 0 02 0 20 ± 0 01 971222 02 (600,45') 9 408 ± 0 034 1 56 ± 0 09 0 17 ± 0 01 0 12 ± 0 01 971210 02 (600,45') 4 297 ± 0 013 2 72 ± 0 12 0 63 ± 0 03 0 46 ± 0 02 971210 10 (600,45') 10 059 ± 0 037 2 41 ± 0 09 0 24 ± 0 01 0 17 ± 0 01

A.1.2 Neon data

Table A. 2: Neon isotope data of samples not listed in the

main part of this thesis. Errors are 2 a and in¬

clude uncertainties in mass discrimination and

sensitivity; uncertainties of calibration gas con¬ centrations are not included, but expected to be smaller than 3 %. (TE: total extraction; olc: online crushing; MW: microwave)

Sample (°C, mm) 20Ne 21Ne/20Ne 22Ne/20Ne 21Ne time (mm) (109 at g-1) (io-3) (IO6 at g-1)

Luxl TE (1750,15) 3508 69 ± 176 16 2 996 ± 0 174 0 1017 ± 0 0014 10512 74 ± 806 59 BV9714 (600,45) 6 99 ± 0 36 10 854 ± 0 640 0 1110 ± 0 0021 75 82 ± 5 90 BV9714 (1750,15) 6 05 ± 0 33 3 184 ± 0 444 0 1025 ± 0 0045 19 26 ±2 89 AL9713-A (600,45) 4 21 ± 0 22 6 079 ± 0 879 0 1061 ± 0 0036 25 57 ± 3 93 AL9713-A (1750,15) 2 48 ± 0 14 4 150 ± 0 868 0 1036 ± 0 0082 10 29 ± 2 23 AL97-11A (600,45) 6 48 ± 0 33 15 707 ± 0 793 0 1145 ± 0 0026 101 77 ± 7 34 TAN 4 (600,45) 1105 ±0 57 6 375 ± 0 719 0 1061 ± 0 0031 70 47 ± 8 74 TAN 4 (1750oC, 15') 5 67 ± 0 31 4 204 ± 0 855 0 1025 ± 0 0070 23 85 ± 5 02 TAN 7 (600,45) 73 45 ± 3 68 3 380 ± 0 110 0 0980 ± 0 0007 248 30 ± 14 85 TAN 7 (1750,15) 3 85 ± 0 20 2 944 ± 1 105 0 0979 ± 0 0083 11 33 ± 4 29 NZ 17 (1750,15) 4 22 ± 0 03 4 936 ± 0 329 0 1065 ± 0 0021 20 81 ± 1 40 NZ 17 (600,45) 13 23 ±0 07 3 293 ± 0 139 0 1039 ± 0 0014 43 55 ± 1 86

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Table A. 3: Weights of all samples listed in this thesis as

loaded to the noble gas mass spectrometer.

Sample Run No Origin Weight (mg)

AL97-11A Allan Hills Antarctica 118 4 jsl8 , AL9713-A Allan Hills Antarctica 238 3 jsl8 , 971210 01 3 Anderson Antarctica 30 0 Ridge , 971210 02 3 Anderson Antarctica 30 1 Ridge , SULA 438 2 Arena Saddle Antarctica 202 0 , ALX-98-11-1A 16 Beacon Antarctica 233 0 Valley , ALX-98-11-2B 16 Beacon Antarctica 144 0 Valley , ALX-98-2A base 16 Beacon Antarctica 215 0 Valley , ALX-98-2A 16 Beacon Antarctica 123 0 top Valley , ALX-98-2B Beacon Antarctica 1910 jsl6 Valley , ALX98-2C Beacon Antarctica 211 9 jsl8 Valley , ALX-98-5A base 18 Beacon Antarctica 249 6 Valley , ALX-98-5A 16 Beacon Antarctica 140 0 top Valley , ALX98-5B Beacon Antarctica 160 1 jsl8 Valley , BV9714 Beacon Antarctica 208 2 jsl8 Valley , DLE-98-14S Beacon Antarctica 193 0 jsl6 Valley , DLG-98-14X Beacon Antarctica 195 0 jsl6 Valley , DLX98-06-Y Beacon Antarctica 67 5 jsl8 Valley , DLX-98-H Beacon Antarctica 136 0 jsl6 Valley , 971210 11 Black Antarctica 42 0 jsl8 Ridge , 971210 12 Black Antarctica 160 3 jsl8 Ridge , 971210 14 Black Antarctica 170 9 jsl8 Ridge , 971210 zehn 3 Black Antarctica 34 3 Ridge , CH2 15 Chisholm Hills Antarctica 31 9 , CH3 15 Chisholm Hills Antarctica 31 3 , CH4 12 Chisholm Hills Antarctica 31 7 , CH5 12 Chisholm Hills Antarctica 24 6 , 1000202 05 ss02 Ford Peak Antarctica 214 , 1000202 06 ss02 Ford Peak Antarctica 25 0 , 1000202 04 Ol ss02 Ford Peak Antarctica 43 6 , 1000202 04 Px ss02 Ford Peak Antarctica 42 0 , For2 12 Ford Peak Antarctica 27 9 , For4 12 Ford Peak Antarctica 24 1 , Bowl Px 14 Mt Bowen Antarctica 28 2 , BowlR 15 Mt Bowen Antarctica 19 2 , Bow2 8 Mt Bowen Antarctica 23 5 , Bow3 8 Mt Bowen Antarctica 24 6 , Bow3-2 15 Mt Bowen Antarctica 33 3 , Bow4 12 Mt Bowen Antarctica 22 7 , Bow6 12 Mt Bowen Antarctica 26 4 , 971229 04 5 Mt Keinath Antarctica 65 4 , 971229 06 1 Mt Keinath Antarctica 83 0 , 971229 06 Mt Keinath Antarctica 1210 jsl9 , 971229 08 Mt Keinath Antarctica 219 6 jsl8 , 971229 12 1 Mt Keinath Antarctica 718 , Mt Keinath Antarctica 61 9 Kl/2 jsl8 , Mt Keinath Antarctica 113 7 Kl/3 jsl8 , DXP-99-37 4 Pivot Peak Antarctica 66 4 , DXP99-38 Pivot Peak Antarctica 249 0 jsl8 , DXP99-39 Pivot Peak Antarctica 260 1 jsl8 , DXP-99-39Z Pivot Peak Antarctica 149 0 jsl9 , DXP-99-40 5 Pivot Peak Antarctica 65 0 ,

continued on next page A.2 Sample weights

Tab. A.3 (cont.): Weights for all samples.

Sample Run No Origin Weight (mg)

DXP-99-41 Pivot Peak Antarctica 163 0 jsl9 , DXP-99-42 1 Pivot Peak Antarctica 205 0 , DXP-99-46 5 Pivot Peak Antarctica 70 6 , DXP-99-48 1 Pivot Peak Antarctica 194 0 , RHPl 8 Ricker Hills Antarctica 23 8 , RHP2 8 Ricker Hills Antarctica 18 2 , RHP9 8 Ricker Hills Antarctica 42 9 , 971222 02 3 Terra Nova Antarctica 39 2 Bay , SULA 245 2 Antarctica 126 0 Wright Valley , Koe5 13 Koefels Austria 14 2 , TE Luxor 115 0 Luxl, top*H jsl8 , Egypt MexlPx 1 Mexiko 52 3 Citlaltepetl , 1 Mexiko 24 2 MexlQz Citlaltepetl , NZ17 Lake Pukaki New Zealand 120 0 jsl9 , NZ19 1 Lake Pukaki New Zealand 139 0 , del 14 Giebichenstein Niedersachsen 32 0 , Gie2 ss02 Giebichenstein Niedersachsen 54 1 , Cosl 8 LDEO Standard 15 1 , SA1 13 Saastal Switzerland 11 3 , Lit2olc 5 Tibet 0 6 Litang , 5 Tibet 56 7 Lit4<100 Litang , 5 Tibet 51 7 Lit4<315 Litang , 13 Tibet 22 2 Nyl Nyalam , NylO ss04 Nyalam Tibet 32 5 Nyll 14 Nyalam Tibet 31 0 Nyl2 11 Nyalam Tibet 13 0 Nyl3 11 Nyalam Tibet 13 6 Nyl4 11 Nyalam Tibet 18 9 Nyl5 13 Nyalam Tibet 25 5 Nyl5 bottom 14 Nyalam Tibet 14 9 Nyl6 10 Nyalam Tibet 32 8 Nyl6 bottom 1 10 Nyalam Tibet 21 2 Nyl6 bottom 2 10 Nyalam Tibet 32 6 Nyl7 ss02 Nyalam Tibet 35 2 Ny2 ss04 Nyalam Tibet 51 3 Ny3 bottom 14 Nyalam Tibet 16 0 Ny3 13 Nyalam Tibet 24 9 Ny4 9 Nyalam Tibet 27 8 Ny4 bottom 9 Nyalam Tibet 31 9 Ny5 14 Nyalam Tibet 29 4 Ny6 ik07 Nyalam Tibet 24 1 Ny6 bottom ss04 Nyalam Tibet 34 2 Ny7 ss04 Nyalam Tibet 43 4 Ny8 9 Nyalam Tibet 28 1 Ny8 bottom 10 Nyalam Tibet 21 1 Ny9 11 Nyalam Tibet 18 9 Ny9 bottom 13 Nyalam Tibet 15 8

Tan5b-MW 5 t Tibet 58 6 Tangguk , A Appendix

A.3 Sample codes of Mt. Keinath and Black Ridge samples

Table A.4: Original labels of the samples from Mount Keinath and Black Ridge and the sample codes used in the text and tables of chapter 3.4.

Original label Short code

971229.02 Kl Kl/2 K2 Kl/3 K3 971229.04 K4 971229.08 K5 971229.06 K6 971210.12 Bl 971210.14 B2 971210.11 B3

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