Ice-Marginal and Periglacial Processes and Sediments (Geological Society Special Publication 354)

Understanding the sediments deposited by glaciers or other cold-climate processes assumes enhanced significance in the context of current global warming and the predicted melt and retreat of glaciers and ice sheets. This volume analyses glacial, proglacial and periglacial settings focusing, among others, on sedimentation at termini of tidewater glaciers, on hitherto not-well-understood high-mountain features, and on sediments such as slope and aeolian deposits whose clasts were sourced in glacial and periglacial regions, but have been transported and deposited by azonal processes. Difficulties are thus often encountered in inferring Pleistocene and pre-Pleistocene cold-climate conditions when the sedimentary record lacks many of the specific diagnostic indicators. The main objective of this volume is to establish the validity and limitations of the evidence that can be obtained from widely distributed clastic deposits, in order to achieve reliable palaeogeographic and palaeoclimatic reconstructions. At a more general level and on the much longer geological timescale, an understanding of ice-marginal and periglacial environments may better prepare us for the unavoidable reversal towards cooler and perhaps even glacial times in the future.

Ice-Marginal and Periglacial Processes and Sediments

The Geological Society of London Books Editorial Committee Chief Editor

Bob Pankhurst (UK) Society Books Editors

John Gregory (UK) Jim Griffiths (UK) John Howe (UK) Rick Law (USA) Phil Leat (UK) Nick Robins (UK) Randell Stephenson (UK) Society Books Advisors

Mike Brown (USA) Eric Buffetaut (France) Jonathan Craig (Italy) Reto Giere´ (Germany) Tom McCann (Germany) Doug Stead (Canada) Gonzalo Veiga (Argentina) Maarten de Wit (South Africa)

Geological Society books refereeing procedures The Society makes every effort to ensure that the scientific and production quality of its books matches that of its journals. Since 1997, all book proposals have been refereed by specialist reviewers as well as by the Society’s Books Editorial Committee. If the referees identify weaknesses in the proposal, these must be addressed before the proposal is accepted. Once the book is accepted, the Society Book Editors ensure that the volume editors follow strict guidelines on refereeing and quality control. We insist that individual papers can only be accepted after satisfactory review by two independent referees. The questions on the review forms are similar to those for Journal of the Geological Society. The referees’ forms and comments must be available to the Society’s Book Editors on request. Although many of the books result from meetings, the editors are expected to commission papers that were not presented at the meeting to ensure that the book provides a balanced coverage of the subject. Being accepted for presentation at the meeting does not guarantee inclusion in the book. More information about submitting a proposal and producing a book for the Society can be found on its web site: www.geolsoc.org.uk. It is recommended that reference to all or part of this book should be made in one of the following ways: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) 2011. Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354. Levy J. S., Head, J. W. & Marchant, D. R. 2011. Gullies, polygons and mantles in Martian permafrost environments: cold desert landforms and sedimentary processes during recent Martian geological history. In: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 167– 182.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 354

Ice-Marginal and Periglacial Processes and Sediments

EDITED BY

I. P. MARTINI University of Guelph, Canada

H. M. FRENCH University of Guelph, Canada

and A. PE´REZ ALBERTI Universidade de Santiago de Compostela, Spain

2011 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 10 000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc. org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ 44 (0)20 7434 9944; Fax þ 44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ 44 (0)1225 445046, Fax þ 44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2011. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of The Copyright Licensing Agency Ltd, Saffron House, 6 –10 Kirby Street, London EC1N 8TS, UK. Users registered with the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA: the item-fee code for this publication is 0305-8719/11/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-327-1 Distributors For details of international agents and distributors see: www.geolsoc.org.uk/agentsdistributors Typeset by Techset Composition Ltd, Salisbury, UK Printed by MPG Books Ltd, Cornwall

Contents Preface MARTINI, I. P., FRENCH, H. M. & ALBERTI, A. P. Ice-marginal and periglacial processes and sediments: an introduction

vii 1

´ . Fingerprints of Quaternary glaciations on Svalbard INGO´LFSSON, O

15

LØNNE, I. & NEMEC, W. Modes of sediment delivery to the grounding line of a fast-flowing tidewater glacier: implications for ice-margin conditions and glacier dynamics

33

LØNNE, I. & NEMEC, W. The kinematics of ancient tidewater ice margins: criteria for recognition from grounding-line moraines

57

LUKAS, S. & SASS, O. The formation of Alpine lateral moraines inferred from sedimentology and radar reflection patterns: a case study from Gornergletscher, Switzerland

77

PE´REZ ALBERTI, A., D´IAZ, M. V., MARTINI, I. P., PASCUCCI, V. & ANDREUCCI, S. Upper Pleistocene glacial valley-junction sediments at Pias, Trevinca Mountains, NW Spain

93

CARLING, P. A., KNAAPEN, M., BORODAVKO, P., HERGET, J., KOPTEV, I., HUGGENBERGER, P. & PARNACHEV, S. Palaeoshorelines of glacial Lake Kuray –Chuja, south-central Siberia: form, sediments and process

111

KELLER, M., HINDERER, M., AL-AJMI, H. & RAUSCH, R. Palaeozoic glacial depositional environments of SW Saudi Arabia: process and product

129

FRENCH, H. Frozen sediments and previously-frozen sediments

153

LEVY, J. S., HEAD, J. W. & MARCHANT, D. R. Gullies, polygons and mantles in Martian permafrost environments: cold desert landforms and sedimentary processes during recent Martian geological history

167

THORN, C. E., DARMODY, R. G. & DIXON, J. C. Rethinking weathering and pedogenesis in alpine periglacial regions: some Scandinavian evidence

183

GUGLIELMIN, M., FAVERO-LONGO, S. E., CANNONE, N., PIERVITTORI, R. & STRINI, A. Role of lichens in granite weathering in cold and arid environments of continental Antarctic

195

VANDENBERGHE, J. Periglacial sediments: do they exist?

205

VAN STEIJN, H. Stratified slope deposits: periglacial and other processes involved

213

OLIVA, M. & ORTIZ, A. G. Holocene slope dynamics in Sierra Nevada (south Spain). Sedimentological analysis of solifluction landforms and lake deposits

227

BROOKFIELD, M. E. Aeolian processes and features in cool climates

241

NEWELL, W. L. & DEJONG, B. D. Cold-climate slope deposits and landscape modifications of the Mid-Atlantic Coastal Plain, Eastern USA

259

Index

277

Ice-marginal and periglacial processes and sediments: an introduction I. PETER MARTINI1*, HUGH M. FRENCH2 & AUGUSTO PE´REZ ALBERTI3 1

School of Environmental Sciences, University of Guelph, Guelph, Ontario N1 G 2W, Canada

2

Departments of Geography and Earth Sciences, University of Ottawa, Ottawa, Ontario K1N 6N5, Canada 3

Departamento de Xeografı´a, Universidade de Santiago de Compostela, Santjago de Compostela, Spain *Corresponding author (e-mail: [email protected])

Abstract: The volume focuses on the analysis of glacial clastic sedimentary deposits, both ancient and recent. The papers range from reviews of glacial systems and cold-climate weathering products and processes to conceptual and field studies of specific ice-marginal and cold-climate sediments. Papers are included that deal with tidewater glaciers, mountain settings on Earth, permafrost areas on both Earth and Mars and detailed regional analyses of cold-climate sediments of Late Pleistocene and Holocene age. The identification of sedimentary facies allows an accurate reconstruction of many of the developmental processes that are involved in ice-marginal and periglacial environments. Lithostratigraphic characteristics of clastic deposits also constitute circumstantial evidence for the previous existence of ancient, and certainly pre-Quaternary, cold-climate systems. This is demonstrated by a study on putative Palaeozoic glacial deposits in Saudi Arabia.

This volume presents a number of papers that relate to both current and ancient ice-marginal and coldclimate environments. Studies of their sediments, weathering and transportation processes contribute to an understanding of the cryosphere. The cryosphere includes Earth’s surface areas that experience one or more of the following: snow cover, sea ice, glaciers, perennial and seasonal frost (Fig. 1). Here, we are concerned with the sediments and weathering processes that occur in the environments that are immediately adjacent to glaciers as well as the frost-dominated environments that characterize cold-climate settings in general. We include contributions that involve not only present-day cases but also those that occurred in the Pleistocene and, in minor measure, the more ancient geological past. In addition, and in anticipation of the future, we include a paper that summarizes recent progress in planetary (Martian) observations.

Glacial and periglacial environments Vast continental areas of Earth have been sculpted by glaciers and many regions are now covered by glaciogenic sediments. Remnants of

these Pleistocene-age ice sheets still exist today, the largest being in Antarctica and Greenland. These ice bodies and the many other smaller glaciers, together with their immediate pro-glacial or ice-marginal surroundings, constitute the glacial environments of today. Equally extensive, both now and in the past, are vast ice-free areas that have either experienced or currently experience cold-climate conditions. These may have lasted for thousands, and in some cases millions, of years. These areas constitute the so-called periglacial environments. Collectively, these two environments extend over approximately one-third of the Earth’s land surface; they undoubtedly occupied much more during the cold periods of the Pleistocene and even earlier during the cold events in ancient geological time. The extraordinarily high erosive and transportational power of glaciers has been well known for over 150 years. Prior to that, during the first half of the 1800s, the full potential of glaciers was not recognized although icebergs and the biblical flood were considered suitable agents for moving large erratic boulders and heterogeneous sediments over considerable distances. The early development of the glacial hypothesis encountered

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 1–13. DOI: 10.1144/SP354.1 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. (a, b) A schematic diagram that illustrates how geography and geomorphology interact with the related physical science disciplines and (c) the major constituents of the cryosphere. Studies of the sediments associated with either ice-marginal or periglacial environments lie within either the shaded or cross-hatched areas in (c) (from French 2007).

opposition but the ever-increasing evidence gradually converted the leading earth scientists of the time such as William Buckland and Charles Lyell (Chorley et al. 1964). A somewhat refined glacial hypothesis was developed by Louis Agassiz in 1840 but the first real scientific glacial study was published by Archibald Geikie in 1863 for Scotland (followed by several other publications that included the first edition of The Great Ice Age; Geikie 1874). By the turn of the century, the theory of Pleistocene ice ages was well established both in Europe and North America (Wright 1890; Geikie 1897; Daly 1934; North 1943). The periglacial concept is more recent in origin. The term was first used by a Polish geologist, Walery von Łozinski, in the context of the mechanical disintegration of sandstones in the Gorgany Range of the southern Carpathian Mountains (a region now part of central Romania). He described the angular rock-rubble surfaces that characterize the mountain summits as ‘periglacial facies’ formed by the previous action of intense frost (Łozinski 1909). Subsequently, the concept of a ‘periglacial zone’ was introduced (Łozinski 1912) to refer to the climatic and geomorphic conditions of areas peripheral to Pleistocene ice sheets and glaciers. Theoretically, this was a tundra zone that extended as far south as the treeline. In the mountains, it was a zone between the timberline and snowline.

Today, Łozinski’s definition is regarded as unnecessarily restricting; few, if any, modern analogues exist (French 2000). There are two main reasons. First, frost action phenomena are known to occur at great distances from both present-day and Pleistocene ice margins. In fact, frost-action phenomena can be completely unrelated to icemarginal conditions. Second, the term has been increasingly understood to refer to a complex of cold-dominated geomorphic processes. These include not only unique frost-action and permafrost-related processes but also a range of azonal processes, such as those associated with snow, running water and wind, which demand neither a peripheral ice-marginal location nor excessive cold. Instead, these processes assume distinctive or extreme characteristics under cold, non-glacial conditions. Studies of the ice-marginal and periglacial environments do not differ tactically from those of other Earth surface systems except for one important fact: one is dealing with environments in which an unusual mineral (ice, H2O) is very close to its melting point. It also experiences sublimation. As a result, the presence of snow and ice generates conditions and landscapes that are unusual and highly variable over both short and long time spans (night and day, seasonal and multi-annual, century, millennia). A number of texts cover the broad fields of ice, glaciology and glacial

INTRODUCTION

3

geomorphology (Souchez & Lorrain 1991; Paterson 1994; Benn & Evans 1998; Liestol 2000; Martini et al. 2001).

Ancient environments and geological contexts In the study of Earth systems, we are trained to learn from the present in order to interpret the past. However, we must be mindful of the very different settings that are involved and that some events are so rare they may not be observed directly during a lifetime and need to be inferred from the sediment/rock record they leave. Moreover, the geo(morpho)logic system is complex; a full understanding requires contributions from a myriad of sciences that have become increasingly complex in the last two to three decades. For example, the basic sciences such as physics, chemistry and biology must be applied to understand the main component of both the terrestrial glacial and periglacial systems and Martian geology, namely ice. The rheological behaviour of glaciers and the landscape, both erosive and depositional, that glaciers leave behind are also central concerns while an understanding of the freezing process, be it seasonal or perennial, is an essential but not defining aspect of periglacial geomorphology. There is also overlap with other subdisciplines; for example, in both glacial and periglacial environments, azonal processes such as running water, wind and gravity-induced displacements often assume critical importance. The same combination of processes must also be considered when inferring the nature of wind-related processes on the Martian surface. Since early times, Earth’s climate has experienced variations from cool long-lasting (‘Ice-house’) periods to warm (‘Greenhouse’) periods (Fig. 2). Humans have evolved and still live in the last Icehouse period, the Quaternary, a period punctuated by relatively short warmer stages when glaciers retreated (interglacial) and longer colder stages (glacial) during which glaciers advanced and snow and ice covered large expanses of the Earth’s surface. Within each glacial stage, smaller temperature variations determined colder periods when glaciers expanded and warmer periods when melting prevailed. Currently, Earth is in an interstadial stage and experiencing a global temperature increase.

Planetary environments The recent growth in the study of planetary environments presents special problems for the two disciplines of glacial and periglacial geomorphology. On Mars for example, temperatures fall to as low

Fig. 2. A graph showing estimated changes in global Earth temperature during geological time and alternating cold and warm periods (modified from Scotese 2008).

as 2250 K and the planet is correctly viewed as possessing not only a cryotic (temperature less than 0 8C) environment but also several Ice Ages (Head et al. 2003). It is highly probable that the Martian near-surface contains H2O in the form of buried icy bodies (Mellon & Jakowsky 1995) and there is morphological evidence that suggests the ephemeral occurrence of surface water in the geological past (Baker 2001). The weathering and

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landscape models associated with traditional (Earthbased) ice-marginal and periglacial processes and sediments are therefore uniquely challenged when totally cryotic environments are considered.

The glacier system By definition, glaciers form on land but may extend into large lakes and the ocean where they form ice shelves. They respond to accumulations of snow and ice in the upper part of their system by flowing under gravity across the surrounding land as a sort of gigantic debris flow. When armoured with rock and sediment acquired from surrounding exposed terrain or from the glacier substrate through various processes, they abrade and pluck sediment along the way and transport and release it elsewhere. The latter is achieved directly either by plastering on the substrate or in situ melt-out, or indirectly by providing water for gravitational mass movements such as debris flows or for overland, rill and channel fluid flows. Erosional features, from large-scale features such as glaciated valleys and tunnel valleys to smaller features such as striations on bedrock, may survive repeated glaciations. By contrast, the sedimentary records of older events may be partially or totally removed by younger glaciations. Different features form in different parts of the glacier and at the ice margin at different times. Glacial sediment sequences, often partially reworked and modified by proglacial processes, typically become visible upon retreat of a glacier. These sequences vary depending on the morphology and lithology of the substrate and the type of glacier that formed them: valley glaciers or large ice sheets. They may be either temperate or polar and either prevalently warm- or cold-based. The features of glaciers and glacial sediments have been well studied and do not need to be repeated. One exception is to mention the debate on whether features were formed by direct action of glacier ice or by other processes such as sediment gravity flows (mainly debris flows and turbidity currents) and canalized fluid flows. An example of this debate involves the origin of unsorted, usually polymictic, massive or poorly structured deposits. These are generally called ‘tills’ when released directly from the glacier ice or ‘diamicton’ when their origin is uncertain even if their material may have a glaciogenic source. Another example is the origin of drumlins. These may have various internal compositions ranging from massive diamicton to mostly stratified sand and gravel (Shaw & Kvill 1984; Menzies 1995, 1996). The uncertainty regarding depositional process becomes critical when the existence of pre-Quaternary glaciations and their

spatial extension needs to be established (Hambrey & Harland 1981; Deynoux 1985; Eyles 1993; Crowell 1999). It is obvious that no single feature representing a clearly defined process can determine a palaeoenviroment; rather, reliable interpretation must rely upon an assemblage of features, representing a reoccurrence of processes in repeating vertical and/or lateral successions and occurring in a wellestablished stratigraphic framework. To place the various contributions on icemarginal sediments and environments within an appropriate context, the following briefly summarizes several of the characteristic features associated with this environment. First, physical weathering by either armoured glacier ice or by meltwater flows under or in the proximal parts of a glacier leads to a progressive comminution of terrigenous material. Fracturing of particles under the weight of moving glacier ice generates characteristic microscopic and submicroscopic surface textures (Mahaney 1996; Whalley 1996). Pebbles and large clasts, transported at the base of the glacier and subject to vertical movement due to repeated pressure variations and phase change of the ice/water, are moulded into characteristic polished, striated, facetted iron-shaped (flatiron) forms (Fig. 3). The high occurrence of such clasts within a sedimentary deposit is a good indication of glacial origin. Furthermore, striations generated on bedrock surfaces may be a good indication of direct glacial activity and provide palaeoflow directions. Second, a variety of meso- to mega-scale erosional features are created by armoured ice or by subglacial meltwater flows. For example, swarms of partially to totally infilled large channels and tunnel valleys have been interpreted as indicators of ice-marginal proximity in many places in Europe

Fig. 3. Typical striated, facetted flatiron cobble transported at the base of a temperate Pleistocene glacier, S. Ontario, Canada (modified from Martini et al. 2001).

INTRODUCTION

5

Fig. 4. Map illustrating the distribution of buried uppermost Ordovician valleys around the Arabian Shield interpreted as tunnel valleys (modified from Aoudeh & Al-Hairi 1995; Le Heron et al. 2009).

(Piotrowski 1994; Jørgensen & Sandersen 2006), North America (Barnett et al. 1998; Russell et al. 2003; Hooke & Jennings 2006), South Africa (Visser 1988; Eyles & de Broekert 2001) North Africa (Ghienne & Deynoux 1998; Hirst et al. 2002) and the Middle East (Aoudeh & Al-Hajri 1995; Le Heron et al. 2009) (Fig. 4). Third, materials transported by a glacier usually retain the characteristics imparted by cold-climate weathering such as angular clasts and unsorted matrix, even when being moved in either supraglacial or englacial positions. Some of this material, frequently polymictic and with large erratic clasts, may be transported from distant and geologically different areas. Material transported at the base may be subject to polishing, rounding and sculpting but the sediment retains a generally poorly sorted, massive and compacted nature. Glacier movements during normal advances, surges or related repeated retreats and re-advances of the terminus can deform these deposits in a characteristic fashion. These deformations therefore provide useful information for interpreting Pleistocene and older putatively glacial deposits (Fig. 5) (Le Heron et al. 2005; Evans et al. 2006). Fourth, the presence of till or till-like deposits is one of the principal lines of circumstantial evidence for past glacier activity in Pleistocene and older successions (Crowell 1999). However, many processes contribute to the release, reworking and sedimentation of glaciogenic material at its terminus, particularly of temperate glaciers. These mainly include debris flows that generate diamicton (similar to tills in terrestrial settings), turbidity flows that move glaciogenic material into the deeper parts of lake and marine basins and water flows that generate a variety of fluvial sedimentary sequences generally of the braided river type in proglacial settings (Fig. 6). These processes may obscure and

sometime obliterate most of the direct evidence of glacial activity (Eyles 1987). However, some telltale features of glaciations may persist in sediments that allow a glaciogenic interpretation. When placed in the appropriate stratigraphic, palaeoclimatological and regional palaeoenvironmental contexts, they hint to past proglacial activity. Such types of evidence include the polymictic composition of clasts, the occurrence of erratics, the presence of deformations in sandy gravelly deposits (probably due to the melting of stranded or partially buried ice blocs; Price 1973; Fay 2002; Russell & Knudsen 2002) and lonestones that pierce or otherwise deform laminations in fine-grained marine and lacustrine deposits which can be interpreted as ice-rafted dropstones (Fig. 7).

The periglacial system Lozinski’s original concept of the periglacial zone was that of a northern mid-latitude mountain zone lying between timberline and snowline. The ‘zone’ reflected climatic zonation. When considered subsequently in the Pleistocene context, it was a proglacial zone peripheral to the mid-latitude ice sheets and glaciers. A complication is that so-called periglacial conditions often extend south of the latitudinal treeline and below the altitudinal timberline. This is because many areas of the northern boreal forest or taiga are underlain by relict permafrost and glaciers may extend below the timberline and into the forest zone in alpine areas. These various concepts are illustrated schematically in Figure 8. Today, the periglacial concept is slightly broader in definition and usually refers to a range of cold non-glacial processes (French 2007). Snow, ice and permafrost are central, but not defining, elements.

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Fig. 5. Idealized scheme of possible soft-sediment deformation generated by glaciers (modified from Le Heron et al. 2005).

It can be argued that typical periglacial regions do not exist and, if they do, lack well-defined boundaries. It is more realistic to envisage periglacial areas as being cold-climate ‘zones’ in which seasonal and perennial frost, snow and normal azonal processes are present to a greater or lesser degree. The reality is that most periglacial landscapes inherit the imprint, in varying degrees, of either glacial or non-glacial climatic conditions. The essence of both the current periglacial system and the proglacial or Lozinski’s Pleistocene ‘periglacial zone’ can best be illustrated with reference to an area of northwest Banks Island in the western Canadian Arctic. Located at latitude 748N (Fig. 9), not only is the area an obviously active periglacial environment but it also illustrates the nature of the ice-marginal proglacial environment (French 1972). Part of the region is shown in an aerial photograph (Fig. 10). During the Late Pleistocene, an ice lobe associated with the late Wisconsinan ice sheet impinged

on the north coast of Banks Island. A welldeveloped lateral moraine system was formed and, in the proglacial zone to the immediate south, a series of broad meandering ice-marginal channels were eroded. These are very clearly shown in Figure 10. Some channels appear to have been subsequently abandoned when they became plugged with material that either slumped or soliflucted off the moraine. The ice lobe also blocked northwards drainage and a number of ice-dammed lakes formed in the lower reaches of valleys. One such proglacial lake overflowed westwards, forming a striking spillway channel visible on the aerial photograph. A radiocarbon date of 8380 + 150 a BP provides a minimal age for the ice-dammed lake and hence a terminal date for when ice impinged upon the land in that area. In summary, northwest Banks Island was a classic Late Pleistocene–early Holocene proglacial environment. At the same time, NW Banks Island is today a classic active periglacial environment

INTRODUCTION

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Fig. 6. Schematic model of principal terrestrial and marine environments and sedimentary sequences formed during a single advance and retreat of a temperate glacier (from Eyles & Eyles 1992).

characterized by intense frost action and the presence of permafrost. It is the first of these two sorts of environments and its associated sediments that is the central focus of Part One of the volume. The

Fig. 7. Lonestone (dropstone) in laminated marine deposits, ‘Palaeozoic Itarare’ Formation, Brazil (modified from Martini et al. 2001).

second environment relates to Part Two of the volume. Prior to reading the various contributions in Part Two, it is instructive to describe the processes currently operating on northwest Banks Island. Temperatures rise to between þ5 and þ7 8C for approximately 3 months in the summer and fall to below 225 8C during the polar night; permafrost is estimated to be over 400 m thick. The landscape is being eroded by a combination of wind-induced and snow-related mass-wasting processes together with fluvial activity over frozen ground during the short summer months (French 1970, 1971). The surface is being dissected by west- and northwest-flowing streams in shallow valleys with dendritically arranged tributary valleys (Fig. 11). Preferential mass wasting (gelifluction) on northeast- and east-facing slopes reflects the dominant southwest winds in winter that deposit snow on lee (northeast-facing) slopes and keep exposed (southwest-facing) slopes and upland surfaces largely snow-free. This produces a striking

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Fig. 8. Schematic diagram illustrating the limits of the periglacial zone: (a) high latitudes and (b) alpine areas (from French 2007).

asymmetry of slopes and drainage patterns in which south- and southwest-facing slopes are steeper than north- and northeast-facing slopes. The surface of the plain is characterized by large thermalcontraction crack polygons. In summary, this Arctic island and other areas of the circumpolar region are classic examples of periglacial landscapes currently being fashioned by frost action and mass wasting (gelifluction) processes in conjunction with the operation of the azonal processes of wind, snow and fluvial activity. It is this sort of environment and the associated

processes and sediments that are central to Part Two of the volume.

Contributions in this volume Much is now known about glaciers and cold-climate environments. The transition from the ice-marginal setting to that of the cold but essentially non-glacial setting is however of particular interest during a period of active glacier retreat when the great variety of ice-marginal conditions can be observed.

INTRODUCTION

Fig. 9. Location map of northwest Banks Island in the western Canadian Arctic.

9

What we can learn from modern and Pleistocene settings can be used to interpret more ancient occurrences of cold-climate conditions. Whereas the occurrence of cold conditions and glaciers is readily recognized (at least for the Pleistocene), it is still difficult to establish with confidence the glacier margin; the occurrence and actual distribution of more ancient pre-Pleistocene glacial and periglacial systems is therefore especially ambiguous. Many of the contributions (a number derive from presentations made at a session devoted to glacial and periglacial deposits at the 27th Meeting of the International Association of Sedimentologists held in Alghero, Italy in September 2008; others are invited) published in this volume address some of these concerns. The volume is subdivided into two parts. The first part deals with ice-marginal environment and sediments. The first paper by Igo´lfsson presents a brief review of the glaciations of Svalbard. This provides a good, confined model for

Fig. 10. Aerial photograph of part of northwest Banks Island, providing a field example of the spatial overlap of a Late Pleistocene ice-marginal environment and a current periglacial environment (part of A17381-137, National Air Photo Library, Ottawa; produced under license from Her Majesty the Queen in Right of Canada, with permission from Natural Resources Canada).

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Fig. 11. An oblique aerial view of northwest Banks Island showing: the current fluvial dissection; the asymmetrical nature of the valleys; large-scale thermal-contraction crack polygons on the upland surface; and snow remaining on lee slopes and in valley bottoms and small gullies. The photograph was taken in early July.

both terrestrial and marine glacial processes, landforms and sediments. It is followed by two papers by Lønne & Nemec that examine the deposits of the end moraines associated with tidewater glaciers. These settings are impossible to deal with directly during their formation, but the products can be examined in emerged systems. The first paper examines the sedimentological/stratigraphic characteristics of one of these moraines and reconstructs the processes responsible for its formation. The second paper develops a sedimentological/ stratigraphic model that can be used to study and better understand the processes active at the termini of tidewater and other glaciers. The following three papers deal with glaciated mountain settings. The paper by Lucas & Sass analyses the development of high-mountain lateral moraines. It utilizes field observations and geophysical methods (ground-penetrating radar) to establish their evolution and develops a model of formation different from that of larger lateral moraines located further downvalley. The paper by Pe´rez Alberti et al. examines the Pleistocene deposits formed at the junction of two valley glaciers where one temporarily dams the valley of the other, faster-retreating glacier. This is a common occurrence in modern mountains where recurring breaks of the ice dam lead to local highly dissected sedimentary sequences. The paper presented in this volume examines sedimentologically the successions and, aided by OSL (optically stimulated luminescence) dates, establishes the relationships between the remnant parts of the dissected record

and reconstructs a glaciation model of the area during MIS 3–4 (marine isotope stages). The last mountain paper by Carling et al. examines a very large Upper Pleistocene lake dammed by a glacier in the Altai Mountains of Siberia. The lake developed in an unglaciated valley surrounded by a crown of glaciated mountains. The ice dams broke several times leading to megafloods. The study reported here examines the sequences of beaches and shoreline notches left by the lake along the flanks of the valley, and reconstructs and models the palaeohydrology and palaeowinds of the area. Part 1 ends with a paper by Keller et al. that shows how the previous existence of ancient pre-Pleistocene glaciers can be inferred with some confidence from detailed sedimentological/ stratigraphic analysis. The study deals with Palaeozoic petroleum-bearing horizons of SW Saudi Arabia. Although the normal diagnostic characteristics used to recognize the direct action of glaciers cannot be used with great confidence in this case, certain macrofeatures such as tunnel-valley patterns and their sedimentary fills constitute circumstantial evidence. The second part of the volume deals with periglacial settings, which fall into four groups. First, two papers discuss the typical permafrost-related features that develop in perennially frozen surficial materials: one in the relatively humid and warm humid Earth setting (French) and the other in the intense cold (cryotic) conditions that exist on Mars (Levy et al.). Although strong differences exist between the two systems because of these very

INTRODUCTION

different environments, there are sufficient similarities to warrant the application of remote sensing concepts and terrestrial knowledge (particularly in Antarctica) to interpretation of Martian surface and near-surface sediments. A second group examines weathering processes in cold-climate settings. Thorn et al. conclude that strong, active chemical weathering occurs in mountainous sub-arctic environments, further dispelling the traditional idea of the overwhelming efficacy of physical weathering in such settings. To place the physical v. chemical weathering discussion in an even better perspective, the paper by Guglielmin et al. is a case study of the role of biological weathering in the extreme cold-climate environments of Northern Victoria Land, Antarctica. They report on the role of lichens in hardening the exterior of the cupola of tafoni in Antarctica. The fact that these unusual weathering structures also develop in hot environments, and that saline conditions appear to be intimately involved, also highlights the lack of understanding of some aspects of the nature of cold-climate weathering. A third group focuses upon the stratified slope deposits that occur widely in Pleistocene periglacial environments and in today’s alpine environments of the middle and low latitudes. The initial paper by Vandenberghe discusses the apparently semantic but essentially fundamental problems associated with the recognition of so-called ‘periglacial sediments’. It is clear that there are cold-climate environments where particular types of weathering such as frost shattering are highly efficient and typical deposits, such as blockfields (the so-called ‘periglacial facies’ of Lozinski), are generated. However, in many cold-climate environments, the typical modes of transport and sedimentation do not produce clastic sediments that are fundamentally different from those produced in other climatic zones. It follows that the use of sedimentary facies alone are insufficient to infer ancient periglacial environments from ancient sediments. Instead, the existence of ancient periglacial (permafrost) environments must be inferred from the presence of post-depositional features such as frost-fissure casts and pseudomorphs. In the following paper, Van Steijn reviews stratified slope deposits. He concludes that although the component particles may have been generated in periglacial settings and may preserve their shape (for instance, forming breccias), other characteristics reflect azonal modes of transport such as rock falls, debris slides and mostly wet and dry debris flows and fluid flows. The third paper by Oliva & Go´mez Ortiz examines sediment movement on slopes in the current periglacial zone of Sierra Nevada (Spain). They ascribe the coarse-grained clastics alternating with organic-rich finer-grained

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sediments as well as the coeval alternation of coarser and finer grained sediments in an adjacent small mountain lake to variations in climatic (temperature and precipitation) conditions from the mid-Holocene onwards. The fourth and final group contains two papers dealing with cold-climate sediments on a broader scale. The first, by Brookfield, reviews aeolian deposits including the putative cold-climate deposits. Loess is included in the discussion. Again, except for certain particular features such as the freshness of the component particles indicating limited chemical weathering, the sedimentological character of wind-blown sediment does not specifically identify a cold-climate origin or specific periglacial processes. Indeed, large quantities of fines generated by glacial abrasion glacier are stored temporarily in proglacial/periglacial settings and are partly eroded and redistributed continent-wide by wind, both in periglacial and non-periglacial settings (Derbyshire & Owen 1996). The existence of periglacial settings is indicated primarily by post-depositional features that indicate frozen ground, such as the relatively well-known frost cracks and cryostructures but also by the less well-known secondary precipitates, neoformed clay minerals and fragipan layers, as described by French. The final paper by Newell et al. brings together many of the various concepts and preoccupations examined in the Special Publication and reports upon the nature and distribution of Late Pleistocene sediments on the Mid-Atlantic Coastal Plain of the eastern US.

References Agassiz, L. J. 1840. E´tudes sur le Glaciers. Jent & Gassmann, Neuchaˆtel, Switzerland. Aoudeh, S. M. & Al-Hajri, S. A. 1995. Regional distribution and chronostratigraphy of the Qusaiba Member of the Qalibah Formation in the Nafud Basin, northwestern Saudi Arabia. In: Al-Husseini, M. I. (ed.) Geo ‘94, The Middle East Petroleum Geosciences (1). Gulf PetroLink, Manama, Bahrain, 143– 154. Baker, V. 2001. Water and the Martian landscape. Nature, 412, 228–236. Barnett, P. J., Sharpe, D. R., Russell, H. A. J., Brennand, T. A., Gorrell, F., Kenny, G. & Pugin, A. 1998. On the origin of the Oak Ridges Moraine. Canadian Journal of Earth Sciences, 35, 1152– 1167. Benn, D. I. & Evans, D. J. A. 1998. Glaciers and Glaciation. Arnold, London. Chorley, R. J., Dunn, A. J. & Beckinsale, R. P. 1964. The History of the Study of Landforms or the Development of Geomorphology, Volume 1: Geomorphology before Davis. Methuen, London. Crowell, J. C. 1999. Pre-Mesozoic Ice Ages: Their Bearing on Understanding the Climate System. Geological Society of America Memoir 192. Daly, R. J. 1934. The Changing World of the Ice Age. Yale University Press, New Haven.

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Derbyshire, E. & Owen, L. A. 1996. Glacioaeolian processes, sediments and landforms. In: Menzies, J. (ed.) Post Glacial Environments: Sediment, Forms and Techniques. Butterworth– Heinemann Ltd., Oxford, 213– 227. Deynoux, M. 1985. Terrestrial or waterlain glacial diamictites? Three case studies from the Late Precambrian and Late Ordovician glacial drifts in West Africa. Palaeogeography, Palaeoclimatology, Palaeoecology, 51, 97– 141. Evans, D. J. A., Phillips, E. R., Hiemstra, J. F. & Auton, C. A. 2006. Subglacial till: formation, sedimentary characteristics and classification. EarthScience Reviews, 76, 115– 176. Eyles, C. 1987. Glacially influences submarine-channel sedimentation in the Yakataga Formation, Middleton Island, Alaska. Journal of Sedimentary Petrology, 47, 1004–1017. Eyles, N. 1993. Earth’s glacial record and its tectonic setting. Earth-Science Reviews, 35, 1 –248. Eyles, N. & Eyles, C. H. 1992. Glacial depositional systems. In: Walker, R. G. & James, N. P. (eds) Facies Models. Geological Associations of Canada, Memorial University, St. John’s, Newfoundland, Canada, 73– 100. Eyles, N. & De Broekert, P. 2001. Glacial tunnel valleys in the Eastern Goldfields of Western Australia cut below the Late Paleozoic Pilbara ice sheet. Palaeogeography, Palaeoclimatology Palaeoecology, 171, 29– 40. Fay, H. 2002. Formation of ice block obstacle marks during the November 1996 glacier-outburst flood (jo¨kulhlaup), Skeiðara´rsandur, southern Iceland. In: Martini, I. P., Baker, V. R. & Garzo´n, G. (eds) Flood and Megaflood Processes and Deposits: Recent and Ancient Examples. International Association of Sedimentologists, Wiley-Blackwell, Oxford, Special Publication, 32, 85–98. French, H. M. 1970. Soil temperatures in the active layer, Beaufort Plain. Arctic, 23, 229– 239. French, H. M. 1971. Slope asymmetry of the Beaufort Plain, northwest Banks Island, N.W.T., Canada. Canadian Journal of Earth Sciences, 8, 717–731. French, H. M. 1972. Proglacial drainage of northwest Banks Island, District of Franklin, NWT. The Musk-Ox, 10, 26–31. French, H. M. 2000. Does Lozinski’s periglacial realm exist today? A discussion relevant to modern usage of the term ‘periglacial’. Permafrost and Periglacial Processes, 11, 35–42. French, H. M. 2007. The Periglacial Environment, 3rd edn. Wiley, Chichester. Geikie, A. 1863. On the glacial drift of Scotland. Transactions of the Geological Society of Glasgow, 1, 1 –190. Geikie, J. 1874. The Great Ice Age, 1st edn. Ibister and Company, London, Daldy. Geikie, J. 1897. The Great Ice Age, 3rd edn. Edward Stanford, London. Ghienne, J. F. & Deynoux, M. 1998. Large-scale channel fill structures in late Ordovician glacial deposits in Mauritania, western Sahara. Sedimentary Geology, 119, 141– 159. Hambrey, M. J. & Harland, W. B. (eds) 1981. Earth’s Pre-Pleistocene Glacial Record. Cambridge University Press, Cambridge.

Head, J. W., Mustard, J. F., Kreslavsky, M. A., Milliken, R. E. & Marchant, D. R. 2003. Recent ice ages on Mars. Nature, 426, 797– 802. Hirst, J. P. P., Benbakir, A., Payne, D. F. & Westlake, I. R. 2002. Tunnel valleys and density flow processes in the upper Ordovician glacial succession, Illizi Basin, Algeria: influence on reservoir quality. Journal of Petroleum Geology, 25, 297–324. Hooke, Le B. R. & Jennings, C. E. 2006. On the formation of the tunnel valleys of the southern Laurentide Ice Sheet. Quaternary Science Reviews, 25, 1364– 1372. Jørgensen, F. & Sandersen, P. B. E. 2006. Buried and open tunnel valleys in Denmark-erosion beneath multiple ice sheets. Quaternary Science Reviews, 25, 1339– 1363. Le Heron, D. P., Sutcliffe, O. E., Whittington, R. J. & Craig, J. 2005. The origins of glacially related softsediment deformation structures in Upper Ordovician glaciogenic rocks: implication for ice sheet dynamics. Palaeogeography, Palaeoclimatology, Palaeoecology, 218, 75– 103. Le Heron, D. P., Craig, J. & Etienne, J. L. 2009. Ancient glaciations and hydrocarbon accumulations in North Africa and the Middle East. East-Science Reviews, 93, 47–76. Liestol, O. 2000. Glaciology, 2nd edn. Department of Physical Geography, University of Oslo, Unipub forlag, Olso. ¨ ber die mechanische Verwitterung Łozinski, W. 1909. U der Sandsteine im gema¨ssigten klima. Bulletin International de l’Academie des Sciences de Cracovie, Classe des Sciences Mathematiques et Naturelles, 1, 1– 25. (English translation: On the mechanical weathering of sandstones in temperate climates. In: Evans, D. J. A. (ed.) 1994. Cold Climate Landforms. Wiley, Chichester, 119– 134.) Łozinski, W. 1912. Die periglaziale fazies der mechanischen Verwitterung. Comptes Rendus, XI Congre`s Internationale Geologie, Stockholm 1910, 1-39-1053. Mahaney, W. C. 1996. Glacial crushing, weathering and diagenetic histories of quartz grains inferred from scanning electron microscopy. In: Menzies, J. (ed.) Moden Glacial Environments – Processes, Dynamics and Sediments. Butterworth-Heinemann, Oxford, 487–506. Martini, I. P., Brookfield, M. E. & Sadura, S. 2001. Principles of Glacial Geomorphology and Geology. Prentice Hall, Upper Saddle River, NJ. Mellon, M. T. & Jakovsky, B. M. 1995. The distribution and behaviour of Martian ground ice during past and present epochs. Journal of Geophysical Research, 100, E6, 11 781– 11 799. Menzies, J. 1995. Modern Glacial Environments – Processes, Dynamics and Sediments. ButterworthHeinemann, Oxford. Menzies, J. 1996. Past Glacial Environments – Sediments, Forms and Techniques. ButterworthHeinemann, Oxford. North, F. J. 1943. Centenary of the glacial theory. Proceedings of the Geologists’ Association, 54, 1 –28. Paterson, W. S. B. 1994. The Physics of Glaciers, 3rd edn. Pergamon, Oxford.

INTRODUCTION Piotrowski, J. A. 1994. Tunnel valley formation in northwest Germany – Geology, mechanisms of formation and subglacial bed conditions for the Bornho¨ved tunnel valley. Sedimentary Geology, 89, 107–141. Price, R. J. 1973. Glacial and Fluvioglacial Landforms. Hafner Publishing Company, New York. ´ . 2002. The effects of Russell, A. J. & Knudsen, O glacier outburst flood flow dynamics on ice-contact deposits: November 1996 jo¨kulhlaup, Skeiðara´rsandur, Iceland. In: Martini, I. P., Baker, V. R. & Garzo´n, G. (eds) Flood and Megaflood Processes and Deposits: Recent and Ancient Examples. International Association of Sedimentologists, Wiley-Blackwell, Oxford, Special Publication, 32, 67– 83. Russell, H. A. J., Arnott, R. W. C. & Sharpe, D. R. 2003. Evidence for rapid sedimentation in a tunnel channel, Oak Ridges Moraine, southern Ontario, Canada. Sedimentary Geology, 160, 33–55. Scotese, C. R. 2008. Paleomap project. http://www. scotese.com/climate.htm (modified by the Acer-acre

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org (Association for Canadian Educational Resources), Glacier periods and glaciers, http://www.acer-acre. org/ClimateChangeCD/sec4/421a.htm). Shaw, J. & Kvill, D. 1984. A glaciofluvial origin for drumlins of the Livingstone Lake area, Saskatchewan. Canadian Journal of Earth Sciences, 21, 1442– 1459. Souchez, R. A. & Lorrain, R. D. 1991. Ice Composition and Glacier Dynamics. Springer-Verlag, Berlin. Visser, J. N. J. 1988. A Permo-Carboniferous tunnel valley system east of Barkly West, northern Cape Province. South African Journal of Geology, 91, 350 –357. Whalley, W. B. 1996. Scanning electron microscopy. In: Menzies, J. (ed.) Past Glacial Environments – Sediments, Forms and Techniques. ButterworthHeinemann, Oxford, 357–375. Wright, G. F. 1890. The Ice Age in North America and its Bearings upon the Antiquity of Man. Appleton and Company, New York.

Fingerprints of Quaternary glaciations on Svalbard ´ . INGO ´ LFSSON O Faculty of Earth Sciences, University of Iceland, Sturlugata 7, Is-101 Reykjavı´k, Iceland and The University Centre in Svalbard (UNIS) (e-mail: [email protected]) Abstract: Marine and terrestrial archives can be used to reconstruct the development of glacially influenced depositional environments on Svalbard in time and space during the late Cenozoic. The marine archives document sedimentary environments, deposits and landforms associated with the Last Glacial Maximum (LGM) when Svalbard and the Barents Sea were covered by continental-scale marine-based ice sheet, the last deglaciation and the work of tidewater glaciers in interglacial setting as today. The terrestrial archives record large-scale Quaternary glacial sculpturing and repeated build-up and decay of the Svalbard– Barents Sea ice sheet. The fingerprinting of Quaternary glaciations on Svalbard reflects the transition from a full-glacial mode, with very extensive coverage by the Svalbard –Barents Sea ice sheet and subsequent deglaciation, to an interglacial mode with valley, cirque and tidewater glaciers as active agents of erosion and deposition. Conceptual models for Svalbard glacial environments are useful for understanding developments of glacial landforms and sediments in formerly glaciated areas. Svalbard glacial environments, past and present, may serve as analogues for interpreting geological records of marine-terminating and marine-based ice sheets in the past.

Svalbard is an archipelago in the Arctic Ocean that comprises all islands between 748N–818N and 108E–358E (Fig. 1). The principal islands are Spitsbergen, Nordaustlandet, Barentsøya, Edgeøya, Kong Karls Land, Prins Karls Forland and Bjørnøya (Bear Island). The total area of Svalbard is 62 160 km2. The West Spitsbergen Current, which is a branch of the North Atlantic Current, reaches the west coast of Svalbard, keeping water open most of the year. The present climate of Svalbard is Arctic, with mean annual air temperature of c. 26 8C at sea level and as low as 215 8C in the high mountains. Most of Svalbard is situated within the zone of continuous permafrost (Humlum et al. 2003). Precipitation at sea level is low, only c. 200 mm water equivalent (w.e.) in central Spitsbergen and c. 400 –600 mm w.e. along the western and eastern coasts of the island. The Svalbard landscape, in particularly the island of Spitsbergen, is generally mountainous with the highest elevation of c. 1700 m a.s.l. on north-eastern Spitsbergen. Large glacially eroded fjords are numerous, particularly at the northern and western coasts of Spitsbergen where the Wijdefjorden, Isfjorden and Van Mijenfjorden fjords have lengths of 108, 107 and 83 km, respectively. Some coastal areas are characterized by strandflat topography: low-lying bedrock plains often blanketed by raised beaches. About 60% of Svalbard is covered by glaciers (Hagen et al. 1993, 2003), with many outlet glaciers terminating in the sea. Svalbard ice caps and glaciers cover about 36 600 km2, with an estimated total volume of c. 7000 km3 (Hagen et al. 1993). Most of the ice volume is contained in the highland ice fields and ice caps on Spitsbergen and

Nordaustlandet, but large valley glaciers and cirque glaciers are frequent along both the west and east coasts of Spitsbergen. Small ice caps also exist on the eastern islands, Edgeøya and Barentsøya (Fig. 1). On Spitsbergen, glaciation is most extensive in areas near the eastern and western coasts, where many glaciers terminate in the sea. In contrast, glaciers in the central part of the island are smaller, mainly because of low precipitation (Humlum 2002). A significant number of glaciers in Svalbard are of the surging type. The surges are relatively short intervals (,1 to .10 a) of extraordinary fast flow which transfer mass rapidly down-glacier, punctuating much longer quiescent periods (,10 to .200 a) characterized by stagnation when ice builds up in an upper accumulation area forming a reservoir of mass for the next surge (Dowdeswell et al. 1991, 1999; Lønne 2004; Sund 2006). Lefauconnier & Hagen (1991) suggested that the majority of Svalbard glaciers surged. The mass balance of many glaciers in Svalbard is partly controlled by snowdrift during the winter (Humlum et al. 2005). The equilibrium-line altitude (ELA) rises on a transect from west to east across Spitsbergen (Fig. 1), reflecting the distribution of precipitation very well. On Prins Karls Forland and along the central west coast it lies at 300 m a.s.l., but reaches .700 m in the highlands of north-eastern Spitsbergen. There are two end-member modes of glacierization on Svalbard: a full-glacial mode, when Svalbard and the Barents Sea were covered by a large marine-based ice sheet, and an interglacial mode (like today) when the Svalbard glacial system is dominated by highland ice fields, ice caps and numerous valley and cirque glaciers. The

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 15– 31. DOI: 10.1144/SP354.2 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. The Svalbard archipelago with distribution pattern of the equilibrium-line altitude (ELA) given as 100 m contour intervals (modified from Hagen et al. 2003). The islands of Hopen (SE from the Svalbard archipelago) and Bjørnøya (midway between Norwegian mainland and Spitsbergen) are not on the map.

full-glacial mode leaves pronounced fingerprints on the continental shelf margins and slopes, and during deglaciation sediments and landforms are deposited on the continental shelf and in fjords around Svalbard. Most sedimentation occurs subglacially in fjords and on the shelf, and ice-marginally on the continental break and slope. There is prevailing erosion inside the present coast, but a strong signal of glacial isostasy in response to deglaciation where sets of raised beaches mark deglaciation and marine transgression. The interglacial mode is characterized by fjord and valley sedimentation below and in front of polythermal and surging glaciers. The interglacial mode of glacierization produces landform-sediment assemblages that can be related to the tidewater glacier landsystem (Ottesen & Dowdeswell 2006), the glaciated valley landsystem (Eyles 1983) and the surging glacier landsystem (Evans & Rea 1999). The glacial fingerprinting on Svalbard is primarily reflecting the transition from a full-glacial mode to an interglacial mode.

Full-glacial-mode sediments and landforms The timing of the onset of Cenozoic Northern Hemisphere high-latitude glaciations is not well known. Ice rafted debris (IRD) and foraminiferal data from Arctic basin deep-sea sediment cores suggests that episodical perennial sea ice might have occurred as early as the middle Eocene 47.5 million years ago (Ma) (Stickley et al. 2009). It is recognized that sea-ice cover existed in the central Arctic basin by the middle Miocene (Darby 2008; Krylov et al. 2008), but ice-sheet build-up over the Svalbard– Barents Sea region probably did not initiate until the Pliocene– Pleistocene, 3.6– 2.4 Ma (Knies et al. 2009). Sejrup et al. (2005) suggested that extensive shelf glaciations started around Svalbard at 1.6–1.3 Ma. The number of full-scale ice-sheet glaciations over Svalbard–Barents Sea is not known, but Solheim et al. (1996) suggest at least 16 major glacial expansion events occurred over the past 1 Ma. Laberg et al. (2010) reconstructed the

FINGERPRINTS OF GLACIATIONS ON SVALBARD

late Pliocene –Pleistocene history of the Barents Sea ice sheet, based on three-dimensional seismic data from the south-western Barents Sea continental margin. They inferred that a temperate Barents Sea ice sheet with channelized meltwater flow developed during the late Pliocene–Early Pleistocene. More polar ice conditions and a Barents Sea ice sheet that included large ice streams, with little or no channelized meltwater flow, occurred in the Middle and Late Pleistocene. There are both marine and terrestrial geological archives that highlight full-glacial-mode conditions and subsequent deglaciation.

Marine archives The dimensions and dynamics of the Last Glacial Maximum (LGM) Svalbard –Barents Sea ice sheet are reflected in the submarine sediments and landforms preserved on the seafloor of the deglaciated shelves and fjords (Ottesen et al. 2005). Marine archives that contain information on former ice-extent and ice dynamics include the following. Shelf bathymetry. Landforms include glacial troughs, submarine transverse ridges, mega-scale glacial lineations, elongated drumlins and rhombohedral ridge systems. These delineate the drainage of glaciers and show that the shelf areas have been shaped by erosion and deposition below and in front of moving outlet glaciers and ice streams. High-resolution seismic records. These show glacial unconformities and give information on thickness, extensions and architecture of sediments above basement rocks. These records signify the extent of glacial erosion and subsequent deposition on the shelf. Sediment cores. These include sedimentological and petrographic analyses for identifying tills and glaciomarine sediments. Sediment cores are used to verify seismic records. The tills are first-order evidence on former ice extent, and 14C dates from glaciomarine sediments provide constraining minimum dates for deglaciation of the shelf areas. The seafloor morphology of the Svalbard margin west and north of the archipelago is characterized by a series of deep fjord-trough systems separated from one another by intervening shallow banks. This is caused by the actions of ice sheets and ice streams during the Pleistocene, where the extent of the Svalbard –Barents ice sheet during peak glaciations was repeatedly limited by the shelf edge (Solheim et al. 1996; Vorren et al. 1998). Sejrup et al. (2005) concluded that the morphology strongly reflected that fast-moving ice streams had repeatedly entered the continental shelf areas, creating numerous glacial troughs/channels that

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are separated by shallow bank areas. Less dynamic ice probably existed on shallower banks (Landvik et al. 2005; Sejrup et al. 2005; Ottesen et al. 2007). Studies of large-scale margin morphology and seismic profiles have identified large submarine trough-mouth fans (TMF) at the mouths of several major cross-shelf troughs (Fig. 2) (Vorren et al. 1989; Sejrup et al. 2005). These are stacked units of glaciogenic debris flows interbedded with hemipelagic sediments displaying thickness maxima along the shelf edge, and reflect direct sediment delivery from an ice stream reaching the shelf edge (Vorren et al. 1989; Vorren & Laberg 1997). Andersen et al. (1996) defined five lithofacies groups from cores retrieved from the western Svalbard continental slope. Laminated-to-layered mud and turbidites reflect post-depositional reworking of the shelf banks, caused by eustatic sea-level fall during ice growth. Hemipelagic mud represents the background sediments and is evenly dispersed over the entire continental margin. Homogeneous and heterogeneous diamictons were deposited during glacial melt events (hemipelagic mud with ice-rafted debris) and during peak glaciation on the submarine fans (debris-flow deposits). Large-scale slope failures have affected the glaciogenic deposits along the western Barents Sea margin (Kuvaas & Kristoffersen 1996; Laberg & Vorren 1996). The largest TMFs occur in front of the Storfjorden and Bear Island trough mouths (Fig. 2), probably reflecting where the largest Svalbard–Barents Sea palaeo-ice streams entered the western shelf break (Faleide et al. 1996; Vorren & Laberg 1997; Andreassen et al. 2008). The oldest Storfjorden and Bear Island TMF sediments have been estimated to be c. 1.6 Ma (Forsberg et al. 1999; Butt et al. 2000). Whereas TMFs can be regarded as archives of numerous glaciations, most sediments and landforms on the shelf and in the fjords relate to the LGM and subsequent deglaciation. End-moraines have been identified at several locations on the shelf (Ottesen et al. 2005, 2007; Ottesen & Dowdeswell 2009), suggesting outlet glaciers and ice streams draining the Svalbard fjords and a shelf-edge glaciation along the major part of the margin during the LGM. Ottesen et al. (2005, 2007) and Ottesen & Dowdeswell (2009) recognized an assemblage of sediments and landforms that can be used to infer the flow and dynamics of the last ice sheet on Svalbard (Fig. 3). They distinguished between inter-ice-stream and ice-stream glacial landform assemblages, which reflect different glacial dynamics associated with ice streams in fjords and troughs and slower moving ice between the troughs and ice streams. They identified five subsets of landforms that make up the inter-icestream glacial landform assemblage, and labelled

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Fig. 2. Location of large submarine trough-mouth fans (TMF), reflecting where the largest Svalbard– Barents Sea palaeo-ice streams entered the western shelf break (modified from Vorren et al. 1989).

them 1 to 5 by their relative age of deposition (Fig. 3a). Landforms relating to ice advance to the shelf edge. These are glacial lineations orientated in the direction of ice flow across the shelf, and a well-defined linear belt of hummocky terrain inferred to represent the shelf-edge ice grounding zone (1 on Fig. 3a). The glacial lineations are sets of parallel subdued ridges that have amplitudes of less than 1 m and a wavelength of several hundred metres. The hummocky belt is a well-defined, continual and linear belt of irregular hummocky terrain about one kilometre in width, where hummocks and ridges have amplitudes of c. 5 m. The belt terminates abruptly at the shelf edge (Fig. 3a), and Ottesen & Dowdeswell (2009) suggest that this terrain represents the grounding zone of an ice margin. Landforms of ice retreat across the shelf during deglaciation. These are large and small transverse moraine ridges; small ridges are interpreted to be retreat moraines whereas the larger ridges probably mark stillstands during retreat of a grounded ice margin (2 on Fig. 3a). The lateral continuity of the ridges over a number of kilometres also implies systematic retreat along a wide ice front. Landforms of ice retreat from fjord mouths to fjord heads. These are arcuate moraines (suggesting possible glacial re-advance to fjord mouths and/or stillstands during deglaciation), crag-and-tail features and small transverse ridges (suggesting active ice

in fjords prior to deglaciation). The crag-and-tail landforms (3 on Fig. 3a) are streamlined features with an upstream core of bedrock and glacial sediments deposited in lee of the bedrock knob, produced at the bed of moving ice. The sediment-landform sets (4) and (5) (Fig. 3a) defined by Ottesen & Dowdeswell (2009) were produced during the Holocene and belong to interglacial-mode tidewater glacier sedimentlandform assemblages. These include basin fills within fjords (4), representing fine-grained sediment deposition linked to the discharge of turbid meltwater from tidewater glacier margins and submarine slides from steep fjord walls, demonstrating slope instability. Landforms of recent ice re-advance and retreat at fjord heads (5) include large terminal moraines within a few kilometres of present tidewater glacier margins, recording re-advance associated with the Little Ice Age and subsequent retreat marked by deposition of small, sometimes annual transverse ridges. The ice-stream glacial landform assemblage (Fig. 3b) of Ottesen & Dowdeswell (2009) recognizes sediment-landform subsets that characterize the action of active ice streams in cross-shelf troughs. Mega-scale glacial lineations and lateral icestream moraines. The mega-scale glacial lineations are streamlined linear and curvilinear submarine features elongated in the direction of the long axis of the depressions, observed in several major fjords and cross-shelf troughs on the Svalbard

FINGERPRINTS OF GLACIATIONS ON SVALBARD

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Fig. 3. Schematic models of submarine glacial landforms on Svalbard continental margins. (a) An inter-ice-stream glacial landform assemblage, located between fast-flowing ice streams. (b) An ice-stream glacial landform assemblage, where fast-flowing ice was fed from large interior drainage basins. The landforms are labelled by their relative age of deposition, where 1 denotes the oldest landform (modified from Ottesen & Dowdeswell 2009).

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margin (Ottesen et al. 2007; Ottesen & Dowdeswell 2009). They vary from hundreds of metres to more than 10 km in length and up to 15 m in height. The mega-scale lineations probably result from softsediment deformation at the base of fast-flowing ice streams (Dowdeswell et al. 2004). Lateral ice-stream moraines (1 on Fig. 3b) are individual linear ridges of tens of kilometres in length and up to c. 40 –60 m high that have been observed along some of the lateral margins of cross-shelf troughs in Svalbard. Ottesen et al. (2005, 2007) described linear ridges of tens of kilometres in length and up to 50 m in relative elevation running along the lateral margins of the Isfjorden and Kongsfjorden cross-shelf troughs as they approach the shelf break west of Svalbard. Sub-bottom profilers do not generally achieve acoustic penetration of these ridges, implying that they are made up of relatively coarse diamictic sediments. These extensive lateral ridges are interpreted to define the lateral margins of fast-flowing former ice streams (Ottesen et al. 2005, 2007). Grounding zone wedges and transverse ridges. Grounding zone wedges are large seafloor ridges orientated transverse to the direction of former ice flow and occur both at the shelf edge and in the troughs and fjords of Svalbard. The ridges are characteristically tens of metres high, up to several kilometres wide and tens of kilometres long. Acoustic stratigraphic records show that the ridges form sedimentary wedges lying above strong basal reflectors. Ottesen et al. (2007) concluded that although the sedimentary wedges sometimes only have relatively subtle vertical expression on the sea floor, they may contain a few cubic kilometres of sediments. Where these extensive ridges and underlying sedimentary wedges are found in the troughs and fjords of Svalbard (2 on Fig. 3b) they are interpreted as marking major stillstands of the ice margin during general deglaciation (Landvik et al. 2005; Ottesen et al. 2007), lasting for hundreds rather than tens

of years (Dowdeswell et al. 2008). The diamictic grounding-zone wedges were produced by continuing sediment delivery from the deforming beds of active ice during the stillstands (Dowdeswell et al. 2008). The transverse ridges that are observed on the continental shelf to the side of the troughs (Fig. 3b) have been interpreted to be recessional push moraines reflecting stillstands or wintersummer ice-front oscillations during deglaciation (Ottesen & Dowdeswell 2006). Individual ridges are up to 15 m high, are spaced a few hundred metres apart and usually occur in clusters rather than as isolated individual features. Dowdeswell et al. (2008) argued that the megascale glacial lineations were products of rapid ice retreat, whereas the grounding-zone wedges suggested episodic retreat. They interpreted suites of transverse ridges to be indicative of relatively slow retreat of grounded ice margins. Seismic record and sediment core data. These data (Fig. 4) concur with the bathymetric data on the glacial origin of landforms and sediments described above. Unconformities caused by glacial erosion provide strong reflectors (Solheim et al. 1996). When the marine sequence is penetrated by corers, stiff diamictons, interpreted to be subglacial tills deposited below grounded glaciers in the fjords and out on the shelf, are retrieved (Svendsen et al. 1992, 1996; Landvik et al. 2005). The diamictons are overlain everywhere by fine-grained marine or glaciomarine muds (Elverhøi et al. 1980, 1983; Sexton et al. 1992). Radiocarbon ages on subfossil shells from the muds give constraining ages for the muds as being of deglaciation ages and the diamictons having been deposited in connection with the LGM expansion of ice.

Terrestrial records While the marine archives contain evidence of repeated expansions of the Svalbard–Barents Sea ice sheet to the continental margin around Svalbard,

Fig. 4. Sketch of seismic section along Isfjorden, Svalbard. A moraine ridge at the shelf edge marks LGM extension of an ice stream in the Isfjorden trough, and stiff diamicton is interpreted to be till deposited by the last major glaciations (modified from Svendsen et al. 1996).

FINGERPRINTS OF GLACIATIONS ON SVALBARD

the terrestrial record of full-scale glaciations is more fragmentary because of the prevailing erosion at times of major ice-sheet expansion. Volume estimates of sediments offshore have been argued to indicate that 2–3 km of rock has been removed from central Spitsbergen since the Eocene (Eiken & Austegard 1987; Vorren et al. 1991). It has been suggested that at least half of this volume was removed during the Pleistocene glaciations (Svendsen et al. 1989; Dimakis et al. 1998; Elverhøi et al. 1998), and it has been assumed that the bedrock geomorphology of Svalbard is predominantly the result of Quaternary sculpturing (Hjelle 1993). The landscapes of Svalbard are characterized by extensive glacial carving of cols, valleys and fjords where the glaciers have enhanced pre-glacial fluvial and tectonic landscapes. Svendsen et al. (1989) pointed out that erosion of the major fjords below sea level requires large ice sheets with outlet glaciers at the pressure melting point at their base. They also concluded that the pronounced alpine landscape of Svalbard indicated that cirque and valley glaciers, rather than ice sheets, were mainly responsible for carving the valleys and other high-relief landforms and that glacial erosion by polythermal valley glaciers is the most important geomorphic process in the present climate. Evidence of more extensive ice cover than today during the LGM and previous glaciations is present on every ice-free lowland area around Svalbard outside the Neoglacial limits in the form of glacial drift, erratics and striations (Sollid & Sørbel 1988; Salvigsen et al. 1995). Directional evidence generally suggests ice flow offshore towards the shelf areas on western Svalbard (Kristiansen & Sollid 1987; Landvik et al. 1998). Evidence on ice thickness and ice movements during the LGM include ice-abraded ridge crests roche moutonne´es, striations, erratics and glacial drift on nunataks and coastal mountains. A number of studies have addressed the thickness of the Svalbard –Barents Sea ice sheet over Svalbard during the LGM. A long-standing debate exists on whether morphological data (such as the existence of pre-LGM sets of raised beaches and large rock glaciers) could be taken to suggest the existence of ice-free enclaves on the lowlands of western and northern Svalbard (Landvik et al. 1998, 2005; Andersson et al. 1999; Houmark-Nielsen & Funder 1999). There is a growing consensus that although some coastal mountains may have protruded as nunataks above the ice-sheet surface at LGM on the outer coast of northern and western Svalbard, there are very little data to support the existence of any lowland icefree enclaves (Landvik et al. 2003, 2005; Ottesen et al. 2007). Taken together, marine and terrestrial evidence suggest a LGM configuration of the

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Fig. 5. Reconstruction of the Svalbard– Barents Sea ice sheet and its fast-flowing ice streams (modified from Ottesen et al. 2005).

Svalbard–Barents Sea ice sheet that covered most of Svalbard and its shelf areas (Fig. 5). As there is overall erosion on land on Svalbard during repeated glaciations, the pre-late Quaternary (Saalian) glacial history of Svalbard lacks all details (Svendsen et al. 2004). There are a number of keylithostratigraphical sections that contain tills and marine sediments that have been dated or correlated to late Quaternary Svalbard–Barents Sea ice sheet oscillations (Mangerud et al. 1998) (Fig. 6): Kongsøya (Ingo´lfsson et al. 1995), Kapp Ekholm

Fig. 6. Location of key stratigraphic sites on Svalbard.

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(Mangerud & Svendsen 1992), Skilvika (Landvik et al. 1992), Linne´elva (Lønne & Mangerud 1991), Site 15 (Miller et al. 1989), Kongsfjordhallet (Houmark-Nielsen & Funder 1999) and Poolepynten (Andersson et al. 1999). Most stratigraphic key sites are on the west coast of Svalbard, but the recently described site from Murchisonfjorden, Nordaustlandet (Fig. 6) (Kaakinen et al. 2009) adds to our understanding of late Quaternary glacial events on Svalbard. One striking characteristic of the lithostratigraphical records from coastal Svalbard is that sections often reflect glaciation events in the form of repeated regressional sequences (Figs 7 & 8). Each cycle consists of a basal till (Fig. 8a) deposited during a regional glaciation large enough for isostatic depression to cause transgression and deposition of glaciomarine –marine sediments on top of till as the ice sheet retreats (Fig. 8b, c). Glacial unloading and isostatic rebound causes a coarsening-upwards sequence where sublittoral sediments and beach foresets reflect regression (Fig. 8d, e). This is particularly well expressed in the stratigraphic record from Kapp Ekholm (Fig. 7). Raised beaches around Svalbard can generally be regarded as isostatic fingerprinting of earlier expanded ice volumes compared to present. Postglacial raised beaches have been described from most ice-free coastal areas (Forman 1990; Landvik et al. 1998), and the elevation of the postglacial marine limit and history of relative sea-level changes are well known (Fig. 9) (Forman 1990; Forman et al. 2004). The isostatic fingerprinting (Fig. 10) reflects the heaviest glacial loading in the central Barents Sea and clearly expresses the differential ice load of the Svalbard–Barents Sea ice sheet at LGM.

Interglacial-mode sediments and landforms

Fig. 7. Composite stratigraphy of the Kapp Ekholm section. Each coarsening-upwards sequence reflects glaciation (till) and deglaciation (marine-to-littoral sediments) (modified from Mangerud & Svendsen 1992).

Svalbard did not completely deglaciate during the Holocene (Hald et al. 2004). Salvigsen et al. (1992) and Salvigsen (2002) documented warmer conditions in Svalbard during the early and mid Holocene compared to the present-day climate. Glacier volumes were probably considerably smaller than present (Svendsen & Mangerud 1997; Forwick & Vorren 2007) and some valley/cirque glaciers may have melted away completely. Because of the Neoglacial expansion of glaciers that started some time after mid-Holocene (Svendsen & Mangerud 1997) and culminated by the end of the Little Ice Age around 1890–1900 AD (Werner 1993; Mangerud & Landvik 2007), the timing, extent and volume of ice at the early Holocene glacial minima is not well known (Humlum et al. 2005). Interglacial-mode glacial landforms and sediments

FINGERPRINTS OF GLACIATIONS ON SVALBARD

23

Fig. 8. Examples of Svalbard glacial-deglacial sediments in coastal sections: (a) subglacial till, unit A, Kapp Ekholm (Figs 6 & 7) (pocket knife for scale); (b) dropstones in shallow-marine sediments (pocket knife for scale), site 15 (Fig. 6); (c) stratified shallow-marine sediments with subfossil kelp (35 cm scrape for scale), Poolepynten (Fig. 6); (d) a whale rib at the contact between sublittoral marine sediments and gravelly beach foresets (1 m stick for scale), site 15 (Fig. 6); (e) sublittoral marine sediments with in situ subfossil molluscs, Skilvika (Fig. 6). All photographs by ´ . Ingo´lfsson in 2008. O

on Svalbard primarily relate to the Neoglacial expansion of glaciers. Most glaciers in Svalbard are presently retreating from their 1890– 1900 AD

maxima, and many glaciers have retreated 1–2 km or more. It has been calculated that the net mass balance of Svalbard glaciers has been negative

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Fig. 9. Relative sea-level curves from Svalbard (modified from Forman et al. 2004).

Fig. 10. The pattern of postglacial raised beaches combined with well-dated relative sea-level curves fingerprints the isostatic depression caused by the Svalbard– Barents Sea ice sheet (modified from Bondevik 1996).

FINGERPRINTS OF GLACIATIONS ON SVALBARD

most years for the past .100 years, and that the glacial systems of Svalbard may have lost up to 30% of their volume since 1900 AD (Lefauconnier & Hagen 1990; Glasser & Hambrey 2003).

Tidewater glacier/fjord environments There are a number of conceptual models proposed for tidewater glaciers (Fig. 11a) (Elverhøi et al. 1980; Bennett et al. 1999), identifying and linking sedimentary processes, deposits and landforms. Plassen et al. (2004) proposed a model for sedimentation of Svalbard tidewater glaciers (Fig. 12a) based on high-resolution acoustic data and sediment cores and sedimentation patterns in four tidewater glacier-influenced inlets of Isfjorden, Svalbard. Their model shows glaciogenic deposition in proximal and distal basins. The proximal basins comprise

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morainal ridges and hummocky moraines, bounded by terminal moraines marking the maximum Neoglacial ice extent. The distal basins are characterized by debris lobes and draping stratified glaciomarine sediments beyond and, to some extent, beneath and above the lobes. Distal glaciomarine sediments comprise stratified clayey silt with ice-rafted debris content (Forwick & Vorren 2009). Ottesen & Dowdeswell (2006), Ottesen et al. (2008) and Kristensen et al. (2009) identified an assemblage of submarine landforms from the margins of several Svalbard glaciers that they linked to glacier surging into the fjord environments (Fig. 12b). The submarine landforms include: streamlined landforms found within the limits of known surges, interpreted as mega-scale glacial lineations formed subglacially beneath actively surging ice (1 on Fig. 12b); large transverse

Fig. 11. Svalbard glaciers: (a) Kongsvegen tidewater glacier, Kongsfjorden; (b) Comfortlessbreen glacier in surge; and ´ . Ingo´lfsson in 2008. (c) Pedersenbreen polythermal glacier, Kongsfjorden. All photographs by O

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Fig. 12. Svalbard tidewater glaciers: (a) a model for proglacial sedimentation by Svalbard polythermal tidewater glaciers (modified from Plassen et al. 2004); and (b) landform assemblage model for Svalbard surge-type tidewater glaciers (modified from Ottesen et al. 2008).

ridges, interpreted to be terminal moraines formed by thrusting at the maximum position of glacier surges (2a on Fig. 12b); sediment lobes at the distal margins of terminal moraines, interpreted as glaciogenic debris flows formed either by failure of the frontal slopes of thrust moraines or from deforming sediment extruded from beneath the glacier (2b on Fig. 12b); sinuous ridges, interpreted as eskers, formed after surge termination by the sedimentary infilling of subglacial conduits (4 on Fig. 12b); concordant ridges parallel to former ice margins, interpreted as minor push moraines probably formed annually during winter glacier re-advance (5 on Fig. 12b); and discordant ridges oblique to former ice margins and interpreted as crevasse-squeeze ridges, forming when soft subglacial sediments were injected into basal crevasses (3 on Fig. 12b). Ottesen et al. (2008) proposed that these submarine landforms were deposited in the following sequence based on cross-cutting relationships between them, linked to stages of the surge cycle (Fig. 12b): (1) mega-scale glacial lineations; (2a)

terminal moraines; (2b) lobe-shaped debris flows; (3) isolated areas of crevasse-fill ridges; (4) eskers and (5) annual retreat ridges.

Terrestrial polythermal and surging glaciers There are numerous studies of the depositional environments of Svalbard terrestrial polythermal and surging glaciers (Fig. 11b, c) which outline structural properties, landform-sediment associations and dead-ice disintegration (Boulton 1972; Bennett et al. 1996, 1999; Boulton et al. 1999; Hambrey et al. 1999; Lysa˚ & Lønne 2001; Sletten et al. 2001). Glasser & Hambrey (2003) gave an overview of sediments and landforms associated with glaciated valley landsystems on Svalbard (Fig. 13). Characteristics of this landsystem are rockfall debris supply, passive transport and reworking of a thick cover of supraglacial morainic till, combined with actively transported debris derived from the glacier bed. They identified moraine complexes produced by thrusting as the most common. The sedimentary composition of moraine

FINGERPRINTS OF GLACIATIONS ON SVALBARD

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Fig. 13. A landsystem model for terrestrial Svalbard polythermal glacier (modified from Glasser & Hambrey 2003).

complexes varies with source materials and ranges from reworked marine sediments to terrestrial diamictons and gravels. Original sedimentary structures or subfossil marine mollusks are commonly preserved as a slab of sediments which has been stacked by the glacier. The thrusted moraine complexes often show evidence of glaciotectonic deformations, including low-angle thrust faults and recumbent folds. Moraine complexes resulting from deformation of permafrost also occur on Svalbard. There, stresses beneath the advancing glaciers are transmitted to the proglacial sediments and can cause proglacial deformation of the permafrost layer. This may lead to folding, thrust-faulting and overriding of proglacial sediments. Glasser & Hambrey (2003) suggested that a typical receding Svalbard glacier has three zones within its forefield (Fig. 13) as follows. (1)

(2)

Outer moraine ridge. These are arcuate ridges rising steeply from the surrounding topography to heights of 15– 20 m. They are commonly ice-cored and may be either the result of englacial or proglacial thrusts or be a product of permafrost deformation. Some glaciers have large ice-cored lateral moraines. Moraine-mound complex (Fig. 13), often draped by supraglacial debris stripes. These are often present in the form of arcuate belts of aligned hummocks or mounds comprising

(3)

a wide variety of morphological types (often ice cored), linear ridges up to 100 m long or short-crested ridges of several metres and near conical mounds. Rectilinear slopes and stacking indicate that the moraine-mound complex is a result of thrusting in proglacial, ice-marginal and englacial position. Inner zone, between the moraine-mound complex and the contemporary glacier snout comprising various quantities of foliation-parallel ridges, supraglacial debris stripes, geometrical ridge networks, streamlined ridges/flutes and minor moraine mounds. Sediment facies are predominantly glacial diamicton, commonly being reworked by proglacial streams.

The most widespread deposit on the forefields of receding valley glaciers on Svalbard is diamicton (Glasser & Hambrey 2003) produced by basal lodgement processes or meltout. The diamictons are in turn reworked by fluvial processes and slumping where there is active down-wasting of dead ice (Schomacker & Kjær 2007). Christofferson et al. (2005) described landformsediment assemblages relating to surging Svalbard glaciers. They identified ice-flow parallel ridges (flutings), ice-flow oblique ridges (crevasse-fill features), meandering ridges (infill of basal meltwater), thrust-block moraines, hummocky terrain and drumlinoid hills. Kristensen et al. (2009) suggested

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that surging glacier ice-marginal landforms on land closely resemble the corresponding landforms on the seabed, including debris-flow mud aprons in front of surge moraines. They argued that both the submarine and the terrestrial mud apron were formed by a combination of ice push and slope failure.

particularly important for our understanding of the signatures of surging glaciers, where the recognition of palaeosurges within landform and sedimentary records is still somewhat capricious. Valuable and constructive suggestions from the journal reviewers are acknowledged. The paper was written during a sabbatical visit to Lund University, Sweden.

Conclusion Conceptual models have been developed that explain sediment-landform assemblages for Svalbard shelf-, ice-stream-, fjord-, surging- and terrestrialpolythermal glacial systems. Landsystem models are useful tools for the reconstruction of past environments and palaeoglacier dynamics from geomorphological, sedimentological and stratigraphical records (Evans 2003). Our understanding of the dynamics, processes and products of marinebased ice sheets is hampered by lack of data (Vaughan & Arthern 2007). The Svalbard models therefore have the potential to help clarify the genesis of glacial landforms and sediments in formerly glaciated areas and to help explain the geological record of ancient marine-terminating ice sheets such as the Upper Ordovician Saharan ice sheet (Le Heron & Craig 2008; Le Heron et al. 2010) or the Carboniferous –Permian Gondwana ice sheet (Visser 1989; Isbell et al. 2008). The stratigraphic record of Svalbard –Barents Sea glaciations, with recurring shallowing-upwards marine to littoral sequences separated by tills (Mangerud et al. 1998), could help in the recognition of transitions from full-glacial to interglacial situations recorded in ancient glaciogenic sequences. Epicontinental glaciogenic deposits are generally poorly preserved in the geological records (Eyles 1993) and, seen over an interglacial– glacial cycle, most interglacial deposits and landforms will be destroyed by an advancing/growing ice sheet as the glacial system shifts to full-glacial mode. It has been pointed out that because of the predominantly ice-cored nature of Neoglacial moraines on Svalbard and the very active dead-ice melting, together with the active reworking processes and cryoturbation, the preservation potential of terrestrial glacial landforms on Svalbard is probably poor (Evans 2009). The use of these moraines as modern analogues for ancient glaciated landscapes therefore may not be appropriate (Lukas 2005). However, geomorphological and sedimentological research on landforms and sediments resulting from the last deglaciation and Holocene oscillations of Svalbard glaciers can provide important analogues for palaeoglaciological reconstructions (Boulton 1972; Boulton et al. 1999). This is

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Lysa˚, A. & Lønne, I. 2001. Moraine development at a small high arctic valley glacier: Rieperbreen, Svalbard. Journal of Quaternary Science, 16, 519– 529. Mangerud, J. & Svendsen, J. I. 1992. The last interglacial glacial period on Spitsbergen, Svalbard. Quaternary Science Reviews, 11, 633–664. Mangerud, J. & Landvik, J. Y. 2007. Younger Dryas cirque glaciers in western Spitsbergen: smaller than during the Little Ice Age. Boreas, 36, 278–285 Mangerud, J., Dokken, T. M. et al. 1998. Fluctuations of the Svalbard –Barents Sea ice sheet the last 150 000 years. Quaternary Science Reviews, 17, 11– 42. Miller, G. H., Sejrup, H. P., Lehman, S. J. & Forman, S. L. 1989. Glacial history and marine environmental change during the last interglacial– glacial cycle, western Spitsbergen, Svalbard. Boreas, 18, 272–296. Ottesen, D. & Dowdeswell, J. A. 2006. Assemblages of submarine landforms produced by tidewater glaciers in Svalbard. Journal of Geophysical Research, 111, F01016. Ottesen, D. & Dowdeswell, J. A. 2009. An inter–icestream glaciated margin: submarine landforms and a geomorphic model based on marine-geophysical data from Svalbard. Geological Society of America Bulletin, 121, 1647–1665. Ottesen, D., Dowdeswell, J. A. & Rise, L. 2005. Submarine landforms and the reconstruction of fast-flowing ice streams within a large Quaternary ice sheet: the 2500-km-long Norwegian-Svalbard margin (578 –808N). Geological Society of America Bulletin, 117, 1033–1050. Ottesen, D., Dowdeswell, J. A., Landvik, J. & Mienert, J. 2007. Dynamics and retreat of the late Weichselian ice sheet on Svalbard inferred from high-resolution sea-floor morphology. Boreas, 36, 286–306. Ottesen, D., Dowdeswell, J. A. et al. 2008. Submarine landforms characteristic of glacier surges in two Spitsbergen fjords. Quaternary Science Reviews, 27, 1583– 1599. Plassen, L., Vorren, T. O. & Forwick, M. 2004. Integrated acoustic and coring investigation of glacigenic deposits in Spitsbergen fjords. Polar Research, 23, 89–110. Salvigsen, O. 2002. Radiocarbon-dated Mytilus edulis and Modiolus modiolus from northern Svalbard: climatic implications. Norwegian Journal of Geography, 56, 56–61. Salvigsen, O., Forman, S. & Miller, G. H. 1992. Thermophilous molluscs on Svalbard during the Holocene and their paleoclimatic implications. Polar Research, 11, 1 –10. Salvigsen, O., Adrielsson, L., Hjort, C., Kelley, M., Landvik, J. Y. & Ronnert, L. 1995. Dynamics of the last glaciation in eastern Svalbard as inferred from glacier-movement indicators. Polar Research, 14, 141– 152. Schomacker, A. & Kjær, K. H. 2007. Quantification of dead-ice melting in ice-cored moraines at the high arctic glacier Holmstro¨mbreen, Svalbard. Boreas, 37, 211–225. Sejrup, H. P., Hjelstuen, B. O. et al. 2005. Pleistocene glacial history of the NW European continental margin. Marine Petrology and Geology, 22, 1111–1129.

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Modes of sediment delivery to the grounding line of a fast-flowing tidewater glacier: implications for ice-margin conditions and glacier dynamics IDA LØNNE1* & W. NEMEC2 1

Villaveien 21, N-1440 Drøbak, Norway

2

Department of Earth Science, University of Bergen, N-5007 Bergen, Norway *Corresponding author (e-mail: [email protected])

Abstract: The study focuses on the terminal moraine of a fast-flowing, temperate tidewater glacier that protruded in Oslofjorden trough, southern Norway, during one of the re-advances of the receding Fennoscandian Ice Sheet in the Younger Dryas time. Allostratigraphic mapping is used to reconstruct the moraine’s morphodynamic development, showing how information on the dynamics of ancient glaciers can be derived from their grounding-line deposits. The Storsand moraine commenced its development in the latest phases of ice-margin advance and continued to grow during the stillstand phase, as long as the ice flux persisted. The thick moraine (.100 m) formed in a few decades, to be rapidly abandoned and later emerged by regional uplift. The study concludes that: (a) both meltwater and ice flow invariably supply sediment to the grounding line, and it is the varied preservation potential of ice-derived diamicton that results in misleading differences between moraines; (b) the glacier-front kinematics is asymmetrical with slow advances and rapid retreats; (c) no moraines can form during glacier retreat; (d) the front of an outlet glacier may stabilize while the adjacent ice margin is oscillating or virtually retreating; and (e) marine moraines are an important source of information about ancient ice margins and glacier dynamics.

Ice streams and outlet glaciers, particularly their marine segments, are crucial elements of an ice sheet as they discharge the main part of the bulk ice flux, affect ice-sheet stability and alter global sea level (Alley et al. 2004, 2005; Edwards 2007; Nick et al. 2009). The dynamics of marine outlet glaciers and the particular processes along their grounding line – the border of subglacial and proglacial realm – entail a number of contentious issues, referred to jointly as the ‘grounding-line problem’ (Benn et al. 2007). The controlling system of factors for the observed changes in glacier termini and their varied behaviour remains poorly understood, especially the coupling between internal and external forcing and the interaction between sedimentation, sea-level and glaciological factors (Bamber et al. 2007). Sedimentation is considered to play an important role (Alley et al. 2007) with the sediment delivery by meltwater outflow and by till release from the glacier sole. Sediment accumulation at the grounding line may effectively control the icefront thickness, advance and calving. Although the mechanism of ice-front movements remains to be debated (Alley 1991), subglacial availability of sediment is thought to be critical for the glacier to advance in deep water. The transfer of large sediment volumes to the grounding line is considered

to require a subglacial deforming bed (Boulton 1996a, b; Alley et al. 1997; Murray 1998; Evans et al. 2006). This mobile substrate is envisaged to be an irregular mosaic of deforming spots that are changing their spatial distribution. This results in a subglacial till layer although changes in local subglacial conditions may render the till discontinuous and inhomogeneous, varying considerably in its provenance, derivation mode, composition and thickness (van der Meer et al. 2003). Evans et al. (2006) recognized subglacial thrusting, pervasive shear and melt-out as processes of till formation but pointed out that most tills are heterogeneous, formed by a combination of these processes. Although all these are plausible physical notions, few of them can readily be verified on empirical basis. The number of available marine outlet glaciers is limited, the access to their fronts is restricted by calving and sea ice and it is practically impossible to study their grounding-line processes directly unless by robotic submarines (Powell et al. 1996). However, many conceptual notions can be evaluated on the basis of ancient marine moraines, which are important repositories of sedimentological information on the grounding-line processes and glacier-front behaviour (Lønne 1995, 1997a, b, 2001; Lønne & Syvitski 1997; Nemec et al. 1999; Lønne et al. 2001). The present study of the terminal

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 33– 56. DOI: 10.1144/SP354.3 0305-8719/11/$15.00 # The Geological Society of London 2011.

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moraine of a Younger Dryas outlet glacier in Oslofjorden, southern Norway, follows this line of research with a focus on the sediment delivery to the grounding line. The Oslofjord ice-drainage basin, more than 100 km long, played a significant role in the dynamics of the Fennoscandian Ice Sheet in the final phases of the last glaciation (Boulton et al. 2001b) when the present-day fjord was a shelf trough up to 400 m deep. This paper discusses the development of a submarine moraine at Storsand (Fig. 1) at the grounding line of a temperate, tidewater outlet glacier that advanced in the shelf trough during the Younger Dryas stadial. The fast-flowing glacier re-advanced from an uncertain landward position and moved southwards along the trough, climbing on to a bedrock sill at c. 200 m water depth, scouring sediment deposited during the previous ice-front recession. A prominent submarine moraine was formed, later emerged by regional isostatic uplift. The Storsand moraine, composed of sand and gravel, was subject to commercial mining from c. 1890. The paper is based on the pre-existing maps and photographs and on a detailed study of the open-pit mining sections over the last two decades. The outcrops are presently covered, but

the paper documents one of the most prominent and widest-excavated moraines in southern Scandinavia. The sedimentological study focuses on the moraine architecture and on the significance of subglacially derived sediment varieties and their diverse modes of deposition at the grounding line. Discussed is palaeoglaciological information derived from the moraine record of glacier advance, stillstand and rapid retreat, with broader regional implications for the dynamics of the Younger Dryas ice margin.

Regional geomorphic and palaeogeographical setting Oslofjorden (Fig. 1) is the southern part of an extensive topographic trough (Sigmond et al. 1984) incised in the Precambrian gneisses and Permian granites of southern Fennoscandia by fluvial erosion during the middle to late Miocene dramatic fall of the North Sea level (Galloway 2002; Huuse 2002; Japsen et al. 2002) and scoured further through the Quaternary glacial cycles. The fjord today has a maximum water depth of 85 m in its inner part surrounded by a bedrock landscape 200–300 m in altitude, and gradually deepens and

Fig. 1. The extent of Younger Dryas moraines in Scandinavia (after Andersen et al. 1995a) and the corresponding moraines in Oslofjord area (compiled from Andersen et al. 1995b; Bergstrøm 1999). Note the location of the study area with Storsand moraine.

SEDIMENT DELIVERY TO THE GROUNDING LINE

widens towards the Skagerak area of the North Sea (Fig. 1). The Younger Dryas stadial in Scandinavia was initiated by an estimated drop in mean surface temperatures of 7 8C with a fall of equilibrium-line altitude (ELA) to 600 –700 m, and was terminated by the temperature increase to its pre-stadial level (Boulton et al. 2001b). Modelling simulations of the Fennoscandian Ice Sheet suggest that the ice mass comprised fast-flowing zones separated by sectors with slower flow, all subject to asynchronous oscillations in time and space (Boulton et al. 2001b; Forsstro¨m et al. 2003). This scenario resembled the flow pattern of recent ice sheets, with ice streams 10– 100 km wide. The Oslofjord trough was occupied by fast-flowing ice, forming an ice stream under full glacial conditions and turning into an outlet glacier when the melting ice sheet thinned and high bedrock topography began to emerge. A large calving bay opened as the ice margin receded northwards, with the tidewater front shifting and leaving behind grounding-line moraines (see reviews by Andersen et al. 1995a, b; Bergstrøm 1999). The thick moraine at Storsand belongs to a distinct belt of ice-contact deposits (Fig. 1) with marine and terrestrial segments. To the west of Storsand, the belt is a series of modest ridges referred to as the Geitryggen –Ski moraine com˚ s –Ski plex (Bergstrøm 1999); the corresponding A moraine complex lying to the east (Andersen et al. 1995b) comprises swarms of small ridges intercalated with glaciomarine mud. The Storsand moraine (Fig. 2a) formed at the grounding line of a 4 km wide outlet glacier on a 100 m deep bedrock sill in Vestfjorden, the western arm of inner Oslofjorden, separated from the main fjord by a narrow bedrock ridge (island) 229 m in its present-day altitude. The local marine limit is at 197 m (Sørensen et al. 1990) and some of the adjacent shelf areas were already emerged and deglaciated, hosting deciduous forest vegetation. The fjord was deeper to the north and to the south of Storsand, reaching a depth of 300 m at an up-fjord distance of 4 km and a depth of 400 m at a down-fjord distance of 3 km. The studied outcrop sections extend 750 m parallel and 250 m transverse to the former glacier front (sections 1– 4 in Fig. 2b, c), with the moraine reaching an altitude of 112 m well below the marine limit.

Method Terminology and techniques The definition of ‘sedimentary facies’ and descriptive sedimentological terminology are after Harms et al. (1975, 1982) and Collinson & Thompson

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(1982). The terminology for sediment gravity processes on steep subaqueous slopes and criteria for their recognition are according to Nemec (1990), as used in our related previous papers (Lønne 1995, 1997a, b; Nemec et al. 1999; Lønne et al. 2001; Lønne & Nemec 2004). Following Blikra & Nemec (1998, 2000), the terms debrisflow and debrisfall are written as single words in semantic analogy to widely used terms such as mudflow, sandflow, grainflow and rockfall. The conventional method of lithostratigraphic logging has been used and the use of interpretive facies plot at the log margin, introduced by Nemec & Muszyn´ski (1982), is as in our earlier related papers (references above). The outcrop sections of the excavation pit have been studied in detail and mapped on an allostratigraphic basis with a focus on the bounding discontinuities, rather than facies homogeneity. Isotopic dates are given as non-calibrated C14 ages, for comparative consistency with the previous case studies of Norwegian moraines.

Allostratigraphic guide to moraine architecture As a synthesis of detailed case studies, Lønne (1995, 2001) proposed a generalized allostratigraphic model for the development of marine terminal moraines (Fig. 2d). Five allostratigraphic units are distinguished in a fully developed moraine: units A and B are deposits of ice-contact submarine fan formed during the glacier advance and subsequent stillstand; unit C comprises deltaic deposits formed after the fan had aggraded to the sea surface; and the moraine-draping units D and E are deposits formed after the moraine abandonment by glacier and during its subsequent emergence by regional isostatic uplift. Each unit is further divided into subunits corresponding to the main morphodynamic parts of the system: the ice-contact proximal part (1); the moraine crest zone (2); the ice-distal slope (3); and the slope toe zone (4) passing into basin plain (for example, see subunits A1 –A4 in Fig. 2d). The true ice-contact surface (TICS) corresponding to units A– C is distinguished from the apparent ice-contact surface (AICS) used in geomorphic mapping of moraines (Fig. 2d). The allostratigraphic method is based on architectural analysis and objectively recognizes morphodynamic changes in the system, without implying their causes. The allostratigraphic model has served as a basis for the mapping of outcrop sections and is given here as a guide to the following description of the Storsand moraine architecture and its genetic interpretation.

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Fig. 2. (a) The Storsand moraine and its topographic setting. (b) A plan-view sketch of the excavation pit in 1990 showing the location of outcrop sections 1– 4 and sedimentological logs 1 –3; the south-eastwards bedding dip refers to the grounding-line deposits and the north-westwards dip to postglacial littoral deposits. (c) Oblique aerial photograph of the moraine seen from the east (1989 image no. 292232 from Fjellanger-Widerøe A/S) with the location of outcrop sections and logs and the distribution of allostratigraphic units A–E (their colours correspond to the allostratigraphic model below); MGLP is the maximum grounding-line position. (d) Outcrop-based allostratigraphic model for a fully-developed marine moraine, built to the sea surface, with the five main phases of morphodynamic development recognized as units A– E (after Lønne 2001); for further explanation, see text.

SEDIMENT DELIVERY TO THE GROUNDING LINE

Description of the Storsand moraine Large-scale allostratigraphic architecture The main internal part of the Storsand moraine (Fig. 2c) consists of SE-dipping sand and gravel clinothems that form two large-scale foresets superimposed upon each other. They lack fluvial topset, occur far below the marine limit and are draped with fossiliferous marine facies; this indicates submarine deposition, apparently well below the fjord’s storm-wave base. The moraine’s lower foreset unit reaches an altitude of 95 m and is dominated by deposits of meltwater underflow which are increasingly gravelly upwards, accompanied by debrisflow beds and by diamicton clasts/blocks towards the top. The overlying foreset unit reaches an altitude of 112 m, consists of submarine debrisflow and meltwater outwash deposits, abounds in diamicton olistoliths (Fig. 3b) and shows glaciotectonic deformation (Fig. 3a). The two superimposed foreset units, attributed to the ice-front halting phase and subsequent stillstand (allostratigraphic units A and B in Fig. 2c, d), are truncated by an erosional surface dipping in the up-fjord direction and overlain by deposits with subglacial shear deformation (merged subunits A1/B1 in Fig. 3d). The moraine’s youngest deposits, with clinothems dipping in the up-fjord direction and coarsening upwards from mud to wave-worked sand and beach gravel (subunits D1 and E1 in Figs 2c, d & 3d, e), were formed during the moraine emergence. The basal mud (subunit D1) represents proglacial hemipelagic sedimentation after the moraine abandonment by the glacier. The moraine stoss-side topographic surface is the NW-sloping top of subunit E1 (AICS in Fig. 2d), whereas the TICS is the buried top of subunit A1/ B1 (Fig. 3d). The foreset subunit A3 is dominated by stratified, light-grey gravelly sand deposited from traction by sediment-laden meltwater underflow (sustained turbidity current sensu Kneller & Branney 1995) and intercalated with light-grey massive beds of mud-poor sandy gravel attributed to debrisflows (Fig. 3b). Clasts and small blocks of brownish-grey diamicton are common in the uppermost part of the foreset. The overlying foreset subunit B3 is richer in debrisflow facies and abounds in small to large olistoliths of cohesive, consolidated diamicton composed of mud-rich sand, brownish- or bluish-grey in colour, with an admixture of gravel (Fig. 3b). In the lower part of subunit B3, fragments of bluish diamicton – ranging in size from large pebbles to slabs several metres long – are commonly enveloped in brownish diamicton (Fig. 3a, b). Olistoliths of the bluish diamicton predominate in the upper part, however, occasionally bearing inclusions of

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brownish diamicton. Small fragments of either diamicton also occur scattered in the associated turbiditic and debrisflow facies (Fig. 3b). The two diamicton types show rigid-state interweaving, rather than mixing by mutual dissemination. Their boundaries are sharp shear surfaces (Fig. 3c), indicating that both diamictons were preconsolidated when intercalated by subglacial thrusting, discrete shear and fragmentation. The diamictons have a similar, well-mixed and nearly unimodal grain-size composition of their mud-rich sandy groundmass, with a comparable proportion of gravel and a clay content of 3 to 7 wt% (grain-size distributions in Fig. 4) although the brownish diamicton is slightly richer in sand. They differ markedly from the associated debrisflow deposits which have a bimodal and better sorted grain-size composition, with only 3.5 wt% of silt and little more than 0.5 wt% of clay (Fig. 4). The following description of outcrop sections gives special attention to the occurrence of the two diamicton varieties.

Outcrop section 1 The NE-trending outcrop section 1 (Fig. 2b) is c. 10 m high and 270 m long, curved towards the south at its SW end, and has been divided into five arbitrary segments (labelled a –e in Figs 2c & 4) for easier reference. The SW segment (section 1a, Fig. 4) is roughly transverse to the glacier front and shows the SE-dipping foreset deposits of allostratigraphic subunit B3 overlain by fan-top subunit B2; the underlying foreset deposits of subunit A3 are better exposed in sections 2 –4 and described later in the text. The other segments of section 1 are nearly parallel to the ice front and show mainly subunit B2 (sections 1b–d, Fig. 4). Subunit B3 consists of gravelly debrisflow beds, subordinate pebbly sand turbidites and sporadic debrisfall deposits (log 1, Fig. 4). The foreset succession abounds in diamicton olistoliths, mainly brownish-grey in colour, with occasional inclusions of bluish-grey diamicton (Fig. 5a). Large slabs of diamicton occur piled up on one another in the uppermost part of the foreset slope (Fig. 5a, b), whereas the lower slope deposits bear scattered diamicton olistoliths ranging in size from ,2 cm to lenticular blocks up to 10 m in length (Fig. 5a). The olistoliths are thought to have been dumped from the floating ice front and calving icebergs, as is common at the termini of tidewater glaciers (Fig. 5c, d; Dowdeswell et al. 1994). Some diamicton slabs or their packages intercalated with pebbly sand had apparently moved over short distances downslope as slide blocks, partly due to the fanhead collapses caused by glaciotectonic thrusting. The sudden emplacement of large diamicton blocks affected local sedimentation on the fan slope,

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Fig. 3. Sedimentary facies in the ice-proximal part of Storsand moraine. (a) Proglacially deformed deposits of subunit B2 (outcrop section 1d, Fig. 2b) overlying unconformably the foreset deposits of subunit B3 (transport direction to the right); note the large slabs of bluish-grey diamicton enveloped in brownish-grey diamicton (arrows).

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resulting in hydraulic jumps with cross-strata backsets (Fig. 3b) or entrapment of debrisfalls on the stoss side (Fig. 5b). Small intra-slope debrisflows were probably generated by gravitational winnowing (Postma 1984) due to intense filtrational flow of meltwater through the porous fan head; the larger debrisflows, commonly bearing diamicton clasts, were likely triggered by the aggrading fan-head collapses. Debrisfalls (Nemec 1990) were likely caused by excessive accumulation of coarse gravel as a lag at the fan apex, where single large clasts or their groups would lose stability and move rapidly downslope by rolling, bouncing and sliding according to their individual momentum (as can be seen, for example, on modern talus slopes; Fig. 6a). Debrisfall deposits, in the form of scattered or clustered outsized clasts (Fig. 6b) and upslope-fining gravel patches (Fig. 6c, d), are common in the steepening-upwards foreset of subunit B3. The foreset deposits of subunit B3 are truncated by the subhorizontal to gently inclined base of subunit B2, dominated by diamicton slabs with glaciotectonic thrusts and related synsedimentary deformation (Figs 2c & 4). This fan-top subunit B2 also shows several cross-cutting channels up to a few metres deep (Figs 3a & 7b), probably formed as extensions of subglacial meltwater conduits. The channels are filled with coarse-grained, amalgamated turbidites and debrisflow deposits, occasionally bearing diamicton blocks. Subunit B2 shows no evidence of subglacial shear, and its glaciotectonic thrusting and folding are attributed to the grounding-line bulldozing process. The thrusting of subglacial subunit A1/B1 (Figs 2d & 7c) indicates close proximity to the grounding line and shows an attempt by the glacier to advance over its evolving moraine.

Outcrop section 2 The outcrop section 2 (Fig. 2b, c) is 150 m long and roughly parallel to the direction of glacier advance. The outcrop in its ice-proximal part shows allostratigraphic subunit A3 overlain unconformably by the deformed and merged subunits A1/B1, covered with subunit E1 (Fig. 7a). The outcrop

39

central to the ice-distal part shows the foreset subunit A3 covered and overstepped by the foreset of subunit B3, with the fan-top subunit B2 represented by a channel-fill 5 m thick (Fig. 7a, b). They are truncated by a NW-inclined erosional surface onlapped by a package of well-consolidated, diamicton-rich deposits, up to 7 m thick (merged subunits A1/B1, Fig. 3d), with internal thrusts subparallel to the basal surface. This package comprises large imbricate slices and isolated blocks of both brownish- and bluish-grey diamicton, interspersed with lenses of water-washed sandy gravel. The bluish-grey diamicton predominates upwards. The package pinches out in both updip and downdip direction, and is covered by a downlapping foreset of NW-dipping wave-worked littoral deposits up to 7 m thick (subunit E1, Fig. 3d) which extend laterally to outcrop section 3 (Fig. 2d) and pinch out towards the SW. Subunit E1 (Figs 3d & 7b) consists of clinothems with tangential bases and shows an upwards coarsening from silty sand to parallel-stratified pebbly sand (shoreface facies) and well-rounded, sorted gravel (beach facies), with isolated massive beds of similar gravel (debrisflow facies). Clinothem geometry varies from sigmoidal to oblique, truncated at the top, and the foreset abounds in internal erosional discontinuities (reactivation surfaces, Fig. 7b). The sedimentary succession is interpreted to have formed by submarine fan build-up during the final phase of the glacier advance and its subsequent stillstand (foreset subunits A3 and B3), with the merged subunits A1/B1 emplaced under the sole of an overriding glacier. The heterogeneous subunit A1/B1, chaotically imbricated by thrusting, indicates a non-pervasive, discrete shear and thus a short distance of subglacial transport. The top of subunit A1/B1 is the true ice-contact surface (TICS-A/B in Fig. 7a, b). The weakly compacted channel-fill subunit B2 (Fig. 7b, c) was likely formed by meltwater underflow at the mouth of a subglacial conduit. The landward limit of the palaeochannel corresponds to the seaward pinch-out of subunit A1/B1, allowing the maximum advance position of the grounding line to be recognized (MGLP in Fig. 7b). The overlying subunit E1

Fig. 3. (Continued) (b) Transition from foreset subunit A3 to diamicton-rich foreset subunit B3 (outcrop section 4, Fig. 2b); transport direction to the right. (c) Close-up view of a sharp, sheared boundary between brownish-grey and bluish-grey diamicton. (d) The ice-proximal part in outcrop section 2 (Fig. 2b), showing foreset deposits of subunits A3 and B3 truncated and overlain by the subglacial deposits of amalgamated subunits A1/B1 with slices of brownish and bluish diamicton and evidence of subglacial thrusting (arrows); note the overlying littoral deposits of subunit E1 dipping in the up-fjord direction. (e) Portion of the strike-parallel outcrop section 3 (Fig. 2b), showing channel-fill deposits with diamicton blocks (subunit A1/B1) overlying the foreset deposits of subunit A3 (transport direction towards the viewer, obliquely to the right); the thin covering mud subunit D1 and littoral subunit E1 are dipping away from the viewer.

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Fig. 4. Sketches of outcrop sections 1a, 1b and 1d (Fig. 2c) showing diamicton-rich deposits of the stillstand-phase fan upper slope and top (subunits B3 and B2). Section 1c is poorly exposed, and section 1e is shown in Figure 7c. The lower-left diagram shows sediment grain-size distributions (discussed in the text). Log 1 (from section 1d) shows the component facies of subunits B3 and B2 and their genetic interpretation; HDTC and LDTC denote high- and low-density sustained turbidity currents; DMO denotes diamicton olistoliths.

represents sedimentation during the relative sealevel fall caused by regional glacio-isostatic uplift. The internal reactivation surfaces in subunit E1 are

attributed to erosive storm events, with the episodic debrisflows triggered by gravitational collapses of wave-undercut rising beaches. The topographic

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Fig. 5. (a) Sketch of the southern part of outcrop section 1 (Fig. 2b) showing transition of foreset subunits A3/B3 and the overlying, glacitectonically deformed subunit B2. (b) A pile of brownish-grey diamicton olistoliths up to 10 m long, some with inclusions of bluish-grey diamicton, in the uppermost fan-slope deposits of subunit B3. (c) An overturned modern iceberg in Kongsfjorden, Svalbard, with diamicton attached to its original sole. (d) Another overturned iceberg in the same fjord, showing linear scour striations on the original glacier sole from which basal till has already been dumped to the seafloor.

surface of subunit E1 is only an apparent ice-contact surface (AICS in Fig. 7a, b).

Outcrop section 3 The outcrop section 3 (Fig. 2b, c) is 270 m long and parallel to the depositional strike of the moraine,

traversing its thickest, axial part. The allostratigraphic units recognized and mapped here are subunits A3, A1/B1 and D1/E1 (Fig. 8). The lowest part of the outcrop shows multiple, gently mounded, coarsening-upwards packages of parallel-stratified gravelly sand attributed to meltwater underflow (sustained low-density turbidity

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Fig. 6. (a) Modern talus slope in the Storsand excavation pit, showing small and large patches of openwork debrisfall gravel; the largest one was incidentally arrested on the stoss side of a diamicton block. (b) Lenticular patches of debrisfall gravel embedded in deposits of meltwater underflow in the submarine-fan foreset subunit B3. (c) Deposits of closely successive debrisfalls embedded in a stratified pebbly sand of meltwater underflow, sandwiched between debrisflow deposits. (d) Debrisfall deposit, in the form of a thick patch of well-rounded, clast-supported gravel, arrested at the rear of debrisflow deposit and covered by another debrisflow deposit.

current), increasingly interspersed with debrisfall gravel and intercalated with gravelly debrisflow beds (lowest log 2, Fig. 8). These packages are interpreted to be offset-stacked depositional lobes formed at chute mouths at the toe of the submarine fan slope of an early-stage moraine. They are overlain by a broader lenticular package of debrisflow deposits interspersed with turbidites, up to 4 m thick (Fig. 8b), which is even richer in debrisfall gravel and has a broadly convex-upwards top (Fig. 6b). The SE-inclined deposits are thought to represent the broad, lobe-shaped depocentre of the prograding fan slope, probably related to the outlet of a major subglacial conduit. A succession of similar broadly lenticular packages, offset relative to one another, forms the higher part of subunit A3 (Fig. 8b, log 2) suggesting unstable, laterally shifting subglacial conduits. The overlying subunit A1/B1 (Fig. 8a), c. 6 m thick, consists of multiple channels; some are filled with outwash gravel and others with imbricate slices of brownish- and bluish-grey diamicton interlaced with sandy gravel (Figs 3e & 9). This subunit

is thickest in the central part of outcrop section and pinches out towards both NE and SW, but can be traced laterally to the base of subunit B3 in section 4. The deposits occur no more than 100 m landwards from the maximum grounding-line position (Fig. 2b), and the channels are likely grounding-line features overridden by the glacier. The top of subunit A1/B1 is therefore the true ice-contact surface (TICS-A/B in Fig. 8a). The top of subunit A1/B1 is draped with a mud layer up to 2 m thick (subunit D1, Fig. 9) that pinches out to both SW and NE (Fig. 3e), apparently due to erosion. It bears scattered gravel clasts, more common at the base, and shells of boreal to arctic molluscs (Hiatella arctica, Macoma calcarea, Mya truncata, Arctica islandica and Pecten sp.). The mud layer represents a period of suspension fallout on to the ice-contact surface directly after the moraine abandonment. The scattered debris and fauna content suggest that the glacier retreated rapidly by calving and the submarine moraine soon found itself in a rapidly shallowing environment. The overlying foreset of northwest-dipping,

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Fig. 7. (a) Outcrop section 2 (Fig. 2b) showing foreset subunits A3 and B3 (arrows indicate sediment transport direction) overlain unconformably by the merged subglacial subunits A1/B1 and covered by the NW-dipping littoral deposits of subunit E1. (b) The western part of outcrop section 2 showing the stratigraphic superposition of subunits A3, B3, B2, A1/B1 and E1 (AICS is the apparent ice-contact surface and TICS-A/B is the true ice-contact surface); note that the subglacially deformed subunit A1/B1 pinches out towards the moraine crest, where it overlies undeformed proglacial channel-fill deposits of subunit B2. (c) The northernmost part of outcrop section 1 (Fig. 2b), where the deposits of fan foreset and crest (subunits B3 and B2) are overlain by subglacially deformed deposits (merged subunits A1/B1) and covered by the oppositely dipping clinothems of littoral subunit E1.

coarsening-upwards clinothems of sandy shoreface to gravelly foreshore deposits (subunit E1) is attributed the forced regression caused by glacio-isostatic uplift.

Outcrop section 4 The SW-trending outcrop section 4 (Fig. 2b, c) is c. 120 m long, crossing the south-eastern part of

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Fig. 8. (a) Outcrop sections 3 and 4 (Fig. 2b), showing fan foreset subunits A3 and B3, erosional subunit A1/B1, thin muddy subunit D1 and littoral subunit E1 – 2. (b) Southern part of outcrop section 3, where the steeply inclined foreset of subunit A3 in strike section shows mounded bed packages that are offset relative to one another and markedly broader in the upper part; foreset dip is towards the viewer. Log 2 (from outcrop section 3) shows facies of subunit A3 and their genetic interpretation; HDTC and LDTC denote high- and low-density sustained turbidity currents.

the moraine and roughly parallel to the ice-flow direction. Allostratigraphic subunits A3, B223 and E122 (Fig. 2d) have been recognized and mapped in this section (Fig. 10).

The lowest exposed unit, up to 35 m thick, is a SW-dipping foreset of light-grey gravelly debrisflow deposits intercalated with turbidites and interspersed with debrisfall gravel and small diamicton

SEDIMENT DELIVERY TO THE GROUNDING LINE

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Fig. 9. Multiple channel-fill deposits of subunit A1/B1 in outcrop section 3 (Fig. 2b). Some channels were filled with a mixture of diamicton slabs intercalated with water-washed sediment and others with outwash gravel. The deposits were overridden by the glacier (TICS-A/B: ice-contact surface for units A and B), covered with a mollusc-bearing mud (subunit D1) and overlain by north-dipping littoral facies of subunit E1.

olistoliths (subunit A3 in log 3, Fig. 10). Brownishgrey diamicton predominates, but increasingly more common upwards are inclusions of bluish-grey diamicton; the latter predominates in the overlying foreset subunit B3, which is up to 25 m thick, pinching out updip towards the ice grounding line (Fig. 10). The diamicton blocks and slabs up to several metres in length, scattered in a succession of turbiditic underflow deposits, are thought to have been dumped from the sole of a floating ice terminus, often with some further sliding. The clinothem dip angle in unit B decreases significantly upwards (Fig. 10, top), which marks the transition from progradational subunit B3 to aggradational fan-crest subunit B2 (Fig. 2d) and indicates stabilization of the grounding-line position. No muddy unit D is recognizable in outcrop section 4, probably due to erosion during the relative fall of sea level. The thin wedge of northwestdipping littoral sandy clinothems downlapping subunit B3 at the ice-proximal end of section 4 (Fig. 10, top) can be traced laterally to subunit E1 in section 3 (Fig. 8a). This unconformable foreset extends a few tens of metres to the east, where it covers conformably the subhorizontal fan-top deposits of subunit B2 and is mapped as subunit E2 (Fig. 2d). Subunits E1 – 2 represent deposition during the moraine emergence, and the preservation of subunit B2 indicates negligible erosion in this area. However, the ice-distal side of the moraine lacks coeval littoral deposits (subunit E3, Fig. 2d) and instead shows an equivalent, smooth erosional surface indicating sediment sweep by marine currents. Similar lateral changes from pronounced deposition to sediment bypass or erosion during

emergence characterize other marine moraines (Lønne et al. 2001), reflecting local patterns of wave approach and water-mass circulation.

Morphodynamic evolution of the Storsand moraine In this section, we summarize briefly the physical scenario for the moraine origin. We discuss its morphodynamic development, with a special emphasis on the significance of the sedimentary characteristics of its various parts. The margin of a rapidly thinning Fennoscandian Ice Sheet had re-advanced during the Younger Dryas (Andersen et al. 1995b), forming a belt of terminal moraines in southern Norway (Figs 1 & 11) with marine and terrestrial segments. The outlet glacier in Vestfjorden, the western arm of ice-draining Oslofjorden bedrock trough (Fig. 11), advanced in a water depth of 300 m (profile I, Fig. 11); while considerable, this depth was much shallower than in the settings of analogous recent glaciers in Greenland and Antarctica where water depths commonly exceed 1000 m (Truffer & Echelmeyer 2003). The glacier came to a halt at the mouth of Vestfjorden by climbing on to a bedrock sill at 200 m water depth (profile I, Fig. 11), where the Storsand moraine was deposited as an ice-contact submarine fan during the latest phase of the glacier advance and its subsequent stillstand (Fig. 12). Most deposition occurred on the downfjord side of the bedrock sill, whereas the moraine’s upfjord side was subject to glaciotectonic deformation.

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Fig. 10. Sketch of outcrop section 4 (Fig. 2b) showing the transition from fan foreset subunit A3 to foreset subunit B3 and the fan-top subunit B2, with a relic of littoral subunit E2 at the top; sediment transport direction is to the right.

SEDIMENT DELIVERY TO THE GROUNDING LINE

About 90 –95% of the moraine volume consists of the submarine deposits of sediment-laden meltwater underflow (sustained turbidity current) and related debrisflows, accompanied by debrisfalls and by the dumping of diamicton olistoliths from a floating ice terminus and icebergs. The mud-rich diamicton with well-mixed and nearly unimodal grain-size composition is of intrabasinal subglacial origin, derived from the bottom of the inner, landward part of the trough, where proglacial mud had accumulated since the ice-margin retreat from the line of Ra moraines (Fig. 1). The observed stratigraphic change in diamicton type from brownishto bluish-grey is attributed to a gradual change in the provenance of diamicton and in the mechanics of its subglacial entrainment, as discussed later in the text (pp. 50– 51). The great thickness of Storsand moraine, 112 m above the present-day sea level and several tens of metres below it to the south, suggests deposition at the grounding line of a temperate, fast-flowing tidewater glacier. The notion of persistent ice flux is supported by the abundance of diamicton and the gradual change in its composition, implying progressive substrate scouring. Fast ice flow is generally attributed to soft substrate deformation (Paterson 1994; Bamber et al. 2007). The ice movement in a high-relief bedrock trough might not necessarily depend on this factor, however, since large parts of the glacier perimeter would effectively abut and shear against bedrock walls (Truffer & Echelmeyer 2003).

The record of ice-front latest advance As the outlet glacier at Storsand began to climb on to the bedrock sill into shallower water, the glacier terminus was thinning and its basal shear stress was reduced, allowing progressively thicker deposits to accumulate at the grounding line (Fig. 12). The top of the submarine-fan deposits of unit A is at a water palaeodepth of c. 100 m but lacks flat-lying beds and is overlain unconformably by subunit B2, which suggests that the fan-crest deposits (subunit A2) were eroded prior to the stillstand phase. Nevertheless, about 80 –90% of the moraine volume consists of deposits formed during glacier advance (unit A, Fig. 13). The exposed part of foreset subunit A3, showing middle- to upper-slope facies of the grounding-line

47

submarine fan (Fig. 13), consists of the deposits of high- to low-density turbidity currents, debrisflows and debrisfalls. The sediment-laden meltwater underflow issuing from subglacial conduits was eroding the fan crest and bypassing it through transient channels, accompanied by fan-head-derived debrisflows and debrisfalls. Scattered diamicton olistoliths indicate dumping of subglacial till from a floating ice terminus and its icebergs in the latest phase of glacier advance, with the brownish-grey diamicton increasingly interwoven with a different, bluish-grey diamicton variety.

The record of ice-front stillstand The upper fan-slope deposits in foreset unit B similarly are products of debrisflows, debrisfalls and sediment-laden meltwater underflow (sustained turbidity current). Tractional, stratified deposits of low-density turbidity current constitute c. 10–20% of the foreset thickness (log 3, Fig. 4). The volumetric proportion of diamicton olistoliths is from 30% in the fan axial part (Fig. 4) to c. 50% near its margin (Fig. 5a), and their abundance testifies to a persisting ice flux and intense calving during the stillstand phase. Short, unstable proglacial channels that routed underflow across the fan crest were filled with diamicton olistoliths and/or outwash sediment (Fig. 9) and partly overridden by the glacier. Series of debrisflow deposits in the fan foreset reflect relatively low meltwater discharges, with the outwash sediment accumulating in the fan-head zone and repeatedly collapsing. Debrisfall deposits are typically associated with turbidites and indicate high discharges, with only gravel lag accumulating at subglacial outlets and episodically falling down. The observed facies changes in foreset units A and B may thus reflect variation in meltwater discharge and local seasonal conditions. The increased proportion of sandy turbidites towards the top of unit B (Fig. 10) may be a record of increased axial meltwater runoff heralding the glacier retreat. The maximum grounding-line position (MGLP) for the outlet glacier was likely determined by the local change in the fjord topography, abruptly widening and deepening to the south, and by the volumetric rate of sediment supply to the grounding line. The MGLP in Storsand moraine (Fig. 13) is located between the southernmost extent of subglacial deposits (the pinch-out point of shear-merged subunits

Fig. 10. (Continued) Log 3 shows the facies succession of subunits A3 and B3, with HDTC and LDTC denoting high- and low-density sustained turbidity currents and DMO denoting diamicton olistoliths. The outcrop details show: (a) a diamicton olistolith and upslope-fining patches of debrisfall gravel in a succession of meltwater underflow turbidites; (b) slabs of bluish- and brownish-grey diamicton in the lower part of foreset subunit B3; (c) an olistolith of brownish-grey diamicton with inclusions of bluish-grey diamicton; and (d) a portion of foreset subunit A3 with debrisflow deposits and erosional relics of minor turbidites.

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Fig. 11. Map of the Younger Dryas terminal moraines in the Oslofjord region including Oslofjorden, Øyeren and Øgdern ice-drainage shelf troughs (compiled from Kjærnes 1986; Nordahl-Olsen 1990; Sørensen et al. 1990). ˚ s –Ski moraines (redrawn from Sørensen 1983; originally by The map detail (top) shows the swarms of small A Rosenfeld 1979). Profiles I– III (locations in the main map) show the moraines in present-day topography. The radiocarbon dates are compiled from various sources (Lønne et al. 2001).

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Fig. 12. Development of the Storsand moraine. (a) Longitudinal profile showing the outlet glacier advancing southwards onto the bedrock sill at Storsand and forming an ice-contact submarine fan (unit A, Fig. 2d); the present-day bedrock topography with sediment cover (vertical scale exaggerated). The subglacial till layer composed of brownish-grey diamicton was formed during the glacier advance, supplying olistoliths to the grounding line. (b) Close-up depiction of the glacier front during its stillstand with persisting ice flux, deposition of unit B and the brownish till increasingly replaced by a bluish-grey variety.

A1/B1 in section 2, Fig. 7) and the northernmost limit of proglacial deposits (such as the bulldozed subunit B2 in sections 1a –d, Fig. 4). On this basis, an approximate trace of the MGLP has been reconstructed for the better-exposed north-eastern half of the moraine, 750 m long (Fig. 2b). The lower

boundary of subglacial deposits suggests that the MGLP had a lateral relief of nearly 70 m. During the maximum advance, the height of the ice front from the grounding line to the contemporaneous sea surface would be c. 160 m in the central part of the fjord and only c. 90 m near its margin.

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Fig. 13. Allostratigraphic reconstruction of the Storsand moraine. The shaded part indicates exposure in the open pit. The upper surface is the present-day topography and the bedrock surface is approximate. Note the stratigraphic position of logs 1 –3.

The submarine fan crest from the stillstand phase is relatively well preserved (Fig. 10, top), little eroded during the moraine emergence. The horizontal distance between the fan crest and the reconstructed MGLP decreases from c. 100 m in outcrop section 4 to 0 in section 2 (Fig. 2b), which indicates a crest line roughly parallel to the grounding line. The evidence from outcrop section 1 (Fig. 4) suggests that small accretionary mouth-bars tended to form at the meltwater outlets during episodes of intensified calving, and were subsequently deformed and reworked during the ice-front re-advance. The grounding line might therefore be subject to minor oscillations. A number of factors may cause short-term fluctuations in ice flux and result in such oscillations (Alley et al. 2004). The maximum ice-front position (Fig. 12b) was likely negotiated by the ice flux and local calving rates. It was apparently reaching beyond the MGLP in the form of a floating terminus, with maximum excursions during periods of highest ice flux and sea ice cover. The floating terminus might have extended by little more than a few hundred metres, which would be sufficient to account for the delivery of diamicton olistoliths. If the ice-sole material is chiefly a deformation till with a slip plane at the ice/substrate interface (Evans et al. 2006), the loss of support from the moraine at the floating terminus would greatly enhance the release of the basal diamicton in this zone. The increase in the size and abundance of diamicton olistoliths at the transition of units A/B is

attributed to their early delivery mainly by the melt-out from icebergs and their later supply from a floating terminus (Fig. 12b). The terminus might episodically break into icebergs under the action of waves and tides, perhaps even seasonally, which would account for the comparable seaward distance of olistolith dispersal in units A and B. The concentration of olistoliths near the grounding line may suggest that icebergs tended to be entrapped near the front by either waves or sea ice, as the fjord freezes over even today. The calving of the tidewater ice front may involve several mechanisms (Benn et al. 2007), but separation by deep crevasses along the grounding line would probably be the main factor in the present case. An ice cliff standing 160 m from the grounding line to the sea surface might have produced icebergs up to 200–250 m in length, although they would be much smaller if the ice front was heavily crevassed.

The significance of diamicton olistoliths and subglacial deposits The subglacially deformed deposits (merged subunits A1/B1) on the moraine ice-proximal side are up to 7 m thick (Fig. 7), although their truncated top indicates some erosion during the moraine emergence. Their listric basal surface rises steeply towards the moraine crest (Figs 7 & 11), truncating the underlying foreset subunits A3 and B3 (Fig. 3d) and showing shear deformation. This suggests a

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uniform glacier advance and stable stillstand with a decreasing subglacial scouring of the moraine and no significant grounding-line oscillations. The subglacial deposits are highly heterogeneous and laterally varied, comprising imbricate thrust sheets of diamicton intercalated with waterwashed sand and gravel (Fig. 3d). The slivers of water-washed sediment were likely derived by subglacial thrusting from the deposits of abandoned subglacial canals and the moraine stoss side (Fig. 12b). The bulk style of deformation with the rigid-state interweaving of two diamicton varieties and slicing by thrusts renders the deposits similar to a glacitectonite (Boulton 1996a; Evans et al. 2006), whereas the texture of each diamicton variety indicates intense, ductile-state mixing by pervasive shear. Each diamicton is likely a product of combined substrate thrusting and bulk mixing by shear (Roberts & Hart 2005; Evans et al. 2006). The two different diamictons were delivered concurrently to the grounding line, yet they must have formed separately. The stratigraphic change in diamicton type suggests that the brownish-grey diamicton, bearing well-oxidized mud, was the first to arrive from the glacier upper reaches. A large part of it was likely piled at the foot of the bedrock sill at Storsand. As the ice flux persisted, the glacier scouring must have soon reached deeper, non-oxidized marine mud whereby the bluish-grey diamicton began to form and arrive at the sill foot, becoming interwoven with the brownish diamicton by thrusting and gradually replacing it (Fig. 12a). This interpretation is consistent with the notion that a high diamicton supply to the grounding line can persist only if sediment continues to be entrained by the deforming bed, which means substrate erosion and incorporation of newer sediment portions in the shearing layer (Cuffey & Alley 1996; Alley et al. 1997). The impressive growth of the submarine fan (Fig. 13) suggests a similarly continued growth of the deforming bed. The deforming bed involved material from the successive grounding-line sedimentary wedges cannibalized by the glacier (Fig. 12a), as well as large portions of pre-existing seafloor sediment sheared and entrained by glacial thrusting. The maximum thickness of a pure-shear deforming bed may be approximated by the size of its largest clasts (Dowdeswell et al. 1994; Glasser & Hambrey 2001), and is estimated here as little more than 20–30 cm on the account of scattered cobbles and boulders in the subglacial diamicton. However, the diamicton olistoliths released from the floating-ice sole are up to a few metres in thickness. The subglacial till produced by pure shear was considerably thickened by subglacial thrusting, particularly at the landward foot of the bedrock sill, whereby the diamicton dragged to the

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moraine crest and released there as olistoliths appears to be excessively thick.

The record of ice-front retreat and moraine emergence The mud layer separating the merged subglacial unit A1/B1 from the postglacial subunit E1 (subunit D1 in Figs 9 & 13) is the record of sedimentation that occurred directly after the moraine abandonment. Deposits formed during ice-front retreat often show an upwards fining from gravel-bearing sand to mud (Lønne et al. 2001), but the ice front in the present case must have retreated rapidly in the northwards-deepening fjord (profile I, Fig. 11). Intense calving occurred once the glacier lost its contact with the moraine, as is indicated by the scattered gravel clasts in subunit D1 (Fig. 9). The overlying, shallowing-upwards succession of littoral deposits (subunit E1, Fig. 9) is a forcedregression record of the moraine emergence by regional isostatic uplift. The littoral cover extends from the moraine crest to its ice-proximal side (subunits E2 and E1, Fig. 13), but is lacking on the icedistal side. The ice-distal slope of the moraine, exposed to an open sea, was apparently subject to marine erosion and recorded a non-depositional forced regression with an erosional surface equivalent to subunit E3 (Figs 2d & 13).

Implications for the dynamics of fast-flowing outlet glaciers There has been little research on the sedimentary record of glacier advance, as it was generally assumed to be negligible and a moraine was thought to form mainly during the glacier stillstand and retreat (Boulton et al. 2001b). The allostratigraphic study of the Storsand moraine points to the volumetric prominence of deposits formed during the glacier advance and to the palaeoglaciological significance of their facies architecture (Figs 12 & 13). The study supports the notion of asymmetrical ice-front kinematics with a relatively slow advance and rapid retreat (Meier & Post 1987; Trabant et al. 2003), and concurs with other case studies (Nemec et al. 1999; Lønne et al. 2001) which document rapid ice-front retreat in deepwater basins where calving rates can be catastrophically high. The sedimentary record of moraine growth includes the glacier advance surface whose character (single or multiple TICSs, Figs 2d & 13), geometry and related subglacial deposits bear important information on the glacier behaviour, as discussed in the previous section. In the Storsand moraine, the single listric surface (Figs 7 & 13) indicates a

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continuous glacier advance uninterrupted by significant oscillations; this implies a persistent ice flux, though not necessarily very high or constant. The ice flux was never sufficient for the glacier to advance across the submarine fan, but allowed development of a floating ice foot (perhaps seasonal) from which the diamicton blocks were released. The evidence of a quasi-steady ice flux is consistent with observations from some recent fastflowing tidewater glaciers, which appear to be little susceptible to summer melt and show only minor seasonal velocity fluctuations. This quasisteadiness of flow may seem unusual for a temperate glacier, but is attributed to the high rate of basal melting and well-maintained ice catchment (Truffer & Echelmeyer 2003; Alley et al. 2008). However, such conditions are by no means universal and the moraines of other temperate glaciers may show significant grounding-line oscillations (Lønne et al. 2001). Tidewater ice fronts in modern fjords with water depths comparable to that of the present case tend to be stable for no more than a few decades (Hagen et al. 1993; Weidick 1995). The radiocarbon dates of main ice-margin positions (Fig. 11, map) are therefore of little use when it comes to the glacier residence time at Storsand. However, the reconstructed trajectory of the grounding-line movement (Fig. 13) affords a dynamic resolution close to that of modern ice-front observations. The moraine is estimated to have formed in a few decades; its great thickness with the rapid burial of diamicton olistoliths therefore indicates a very high sedimentation rate (albeit difficult to quantify in reliable terms due to erosional discontinuities and a wide range of depositional processes). Little has thus far been known about the relative role of sediment delivery by meltwater outflow and the deforming-bed conveyor, and also about the depth of subglacial deformation and substrate erosion (Boulton et al. 2001a). The existing literature appears to be somewhat confusing, if not contradictory, in this respect. On the one hand, modelling studies indicate that both ice flow and deforming bed are required for a glacier to advance in deep water, with sediment accumulation at the grounding line (Alley 1991). On the other hand, glacial geologists in their field studies claim that the terminal marine moraines of temperate glaciers are formed during ice-front recession and chiefly by meltwater outwash (Hunter et al. 1996a, b; Seramur et al. 1997). The abundance of diamicton olistoliths in the Storsand moraine is strong evidence for the active role of a deforming bed during both the glacier advance and its stillstand. However, the majority of analogous moraines, even coeval with this one, contain little or no diamicton. May this fact mean

that outlet glaciers can advance and form moraines without an involvement of a deforming bed? It has been suggested by Powell & Alley (1997) that the difference between a diamicton-rich and an outwash-dominated marine moraine is due to the glacier thermal regime. However, the present and earlier case studies (Lønne 1993, 1995, 1997a; Nemec et al. 1999; Lønne et al. 2001) show that both such moraine varieties can be formed by temperate glaciers, even contemporaneously and at closely adjacent localities. Is there a clue to this apparent puzzle? The key factor is the varied preservation potential of glacier-supplied diamicton. The diamicton olistoliths derived from a deforming bed tend to be emplaced on the moraine crest and uppermost seaward slope, which have the lowest preservation potential. They tend to be eroded by the bulldozing glacier and by the powerful outflow of subglacial meltwater and are preserved only in extreme aggradational conditions (when the ice front is thin and the moraine is about to be abandoned), which means time of little or no diamicton delivery. A relatively thick ice front will cannibalize the grounding-line submarine fan, turning its diamicton-bearing subunits into the homogenized, sheared subunit A1 (Figs 2d & 12) where the diamicton olistoliths will effectively fade back into the subglacial till layer. Strong meltwater outflow may be destroying diamicton olistoliths, rather than burying them. The peaks of meltwater outflow and subglacial diamicton delivery are likely to alternate rather than coincide, and it is probably their relative magnitude and frequency as well as the ice-front thickness and flow rate that may render diamicton olistoliths preserved or virtually erased. The powerful meltwater underflow, capable of scouring fan-breaching channels, will inevitably erode and disseminate an obstacle pile of diamicton slabs if given sufficient time. The outlet glacier at Storsand had a modest ice-front thickness and a quasisteady but not too high ice flux, and generated mainly moderate meltwater discharges as indicated by the predominance of debrisflows (Figs 4, 8 & 10). The stronger underflows also tended to be heavily laden with sediment (sustained highdensity turbidity currents) and had limited entrainment capacity, whereas low-density underflows were much less frequent (10– 20% in terms of the fan thickness). These conditions rendered the local sedimentation rate and preservation potential of diamicton very high. Nevertheless, the isolated diamicton blocks and piles of diamicton slabs embedded in outwash deposits (Figs 4, 5 & 10) are probably little more than abundant relics since much diamicton was likely eroded. Depending on local conditions,

SEDIMENT DELIVERY TO THE GROUNDING LINE

diamicton relics may therefore be similarly abundant or be minor or virtually lacking in other, even adjacent, moraines. Poor preservation of diamicton is more common, resulting in an overrepresentation of outwash sediment in the moraine and leading to erroneous inferences. Accordingly, we contend that marine terminal moraines are formed during the glacier advance and evolve further during its stillstand, and that both meltwater outflow and the ice flow itself invariably supply sediment to the grounding-line depositional system. It is only the varied preservation potential of glacier-derived diamicton in the grounding-line systems that makes the misleading apparent difference between moraines.

The regional behaviour of Younger Dryas ice margin The moraine lines mapped in south-eastern Norway (Figs 1 & 11) indicate a highly dynamic ice-sheet margin, varying in spatial and temporal behaviour during its overall northwards recession in the Younger Dryas time (Sørensen et al. 1990; Andersen et al. 1995b; Bergstrøm 1999). Major ice re-advances occurred, and one of the most dramatic is afforded by the Ra moraine line truncating the Onsøy and Hvaler lines of older moraines to the south (Fig. 1). The thickness of moraines and the extent of glacier re-advances were controlled by topography with the thickest moraine segments corresponding to the deepest basins (shelf troughs or fjords) where the ice flux, accommodation space and sediment availability were highest. The moraines are considerably thinner in the coeval shallow-water segments, where they commonly form swarms of closely spaced monoepisodic moraines (inset map and profile II in Fig. 11), composed solely of ice-advance unit A (Fig. 2d). The study of the Storsand moraine adds to the regional observations on Younger Dryas ice-margin behaviour drawn from an analogous allostratigraphic study of the Mona moraine to the east (map and profile III in Fig. 11; Lønne et al. 2001). Moraines, as sedimentary ridges, invariably indicate an episode of ice-front advance even if brief and over a short distance; the traditional regional scenario of ‘a continuously retreating ice-margin during the Younger Dryas’ therefore needs to be revised. The allostratigraphic architecture of all studied moraines shows a core of ice-advance deposits (unit A, Fig. 2d) and the thicker the moraine, the more prominent the sedimentary record of its ice-advance increments. The Storsand and Mona moraines are 100 –150 m thick. The former shows an allostratigraphic succession of units A –B –D –E and the latter shows a succession

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of units A –B –C– D– E (profiles I and III in Fig. 11), possibly with one or more younger re-advance units A buried under glaciomarine mud on its iceproximal side. The Mona moraine formed on a bedrock sill at the mouth of a landwards-deepening shelf trough, where it reached the sea surface and evolved into an ice-contact delta with a short (160 m) fluvial distributary plain. A radiocarbon date of shell from the moraine subunit D1 (Lønne et al. 2001) indicates that the delta was abandoned 10.1 ka BP. However, the dates of two shells from the lowest ice-distal deposits are 300– 500 years older (profile III, Fig. 11) and suggest that a possible relic of an earlier moraine formed on the same sill, which would mean an erosional superposition of moraines. ˚ s–Ski area (profile II and map Moraines in the A detail, Fig. 11) are small ridges 1–6 m thick, formed on a shelf platform with water depth of 50 –100 m (Kjærnes 1986). Although formed during an overall ice-margin recession, the swarms of small moraines indicate numerous minor advances and retreats (possibly seasonal), with some of the re-advances removing previous ridges. Similar seasonal oscillations have been reported from modern tidewater glaciers, such as the Alaskan Hubbard Glacier oscillating by up to 100 m (Trabant et al. ˚ s–Ski moraines thus 2003). The development of A indicates a highly mobile broad front of a relatively thin ice mass with many pinpoints on the shallow shelf, responding rapidly to changes in ice flux, sea level and ablation rate. The moraines are covered with marine mud and littoral deposits formed during the ice-margin retreat and subsequent self emergence (units D– E in profile II, Fig. 11). Radiocarbon dates indicate that the ice margin had retreated from the Ra line to the Aker line of moraines (Fig. 1) in c. 1 ka, but the age resolution is far too low for a reliable correlation of ‘broken’ moraine segments and for an estimation of the temporal frequency of margin shifts and moraine accumulation time. The residence time of outlet glaciers forming the thickest moraines, such as the Storsand and Mona ridges, was probably of the order a few decades rather than a century. It is uncertain to what extent the mobile ice margin was overriding its earlier moraines. Further radiocarbon dating of the ridges and recognition of out-of-place dates on an allostratigraphic basis may help to assess the magnitude and frequency of such episodes. Future regional research should also verify the hypothesis that a major part of glacial deposits on the palaeoshelf may actually represent ice-margin advances, rather than retreat phases, perhaps with a considerable reworking of moraines by ice re-advances.

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Conclusions This study of the Storsand moraine used sedimentological facies analysis and Lønne’s (1995, 2001) method of allostratigraphic mapping to reconstruct in considerable detail the moraine morphodynamic evolution and to show how a wide range of important information on the dynamics of ancient tidewater glaciers can be deciphered from their grounding-line deposits. The Storsand moraine formed at the grounding line of an outlet glacier during one of the re-advances of the receding Fennoscandian Ice Sheet in Younger Dryas time. The outlet glacier was a temperate, fast-flowing ice protrusion in a shelf trough and came to a halt on local bedrock sill. The submarine moraine formed as an icecontact sedimentary wedge that escaped cannibalization and continued to grow during the grounding-line stillstand. Ice flux persisted, supplying abundant diamicton olistoliths of subglacial till to the grounding line. The ice front was relatively thin, ice flux not too high and meltwater discharge mainly moderate, which allowed large portions of diamicton to be preserved in the conditions of limited erosion and high sedimentation rate. The thick moraine is estimated to have formed in a few decades, before being rapidly abandoned and later emerged by regional uplift. The evidence from the Storsand moraine and related case studies of thick moraines suggests that both meltwater and ice flow invariably supply sediment to the grounding line, and it is the varied icemargin regime and diamicton preservation potential that result in the misleading apparent differences between moraines. The study confirms that the main part of a moraine is invariably formed during the glacier advance and not retreat; this supports the notion of asymmetrical ice-margin kinematics with relatively slow, sediment-piling advances and rapid, barely depositional retreats. This notion implies that a major part of deposits on a glaciated shelf may represent ice-front advances, rather than retreat phases. The study contributes to the regional evidence (Lønne et al. 2001) that the front of an outlet glacier may stand still, draining ice and forming a thick moraine, while the adjacent ice shelf may be oscillating or virtually retreating. This evidence indicates that regional palaeoclimatic inferences based on conventional time –distance diagrams of ice-margin shifts may be quite misleading. The study points to the importance of ancient marine moraines as a source of valuable information on ice-front behaviour and glacier dynamics, supplementing knowledge derived from recent glaciers and numerical modelling. The selection of modern analogues for ancient glaciers is highly problematic,

and the physical framework established for an ancient ice margin on a sedimentological basis may therefore be of crucial importance to an understanding of the dynamics of ancient ice sheet. We thank J. Winsemann and M. Ghinassi for their critical reviews of the manuscript, and I. P. Martini for his useful editorial suggestions.

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The kinematics of ancient tidewater ice margins: criteria for recognition from grounding-line moraines IDA LØNNE1* & W. NEMEC2 1

Villaveien 21, N-1440 Drøbak, Norway

2

Department of Earth Science, University of Bergen, N-5007 Bergen, Norway *Corresponding author (e-mail: [email protected])

Abstract: Modern studies indicate that the kinematic behaviour of tidewater glaciers is a crucial part of ice-sheet dynamics. A similar relationship may be expected for ancient ice sheets, but can the kinematics of ancient tidewater ice margins be recognized? The paper addresses this methodological issue by pointing to the high-resolution stratigraphic record of marine moraines, thus far little explored. On the basis of a series of case studies, a range of field criteria are proposed for the recognition of short-term grounding-line movement and possible oscillations in moraine outcrop sections. The method combines allostratigraphic mapping and architectural facies analysis of the moraine sedimentary units formed during glacier advance, subaqueous stillstand and eventual subaerial stillstand, with recognition of the successive ice-contact surfaces. The stacking architecture of the sedimentary units and their ice-contact bounding surfaces reveal the time–distance trajectory of the grounding-line positions. It is also possible to recognize changes in the mode and rate of subglacial sediment delivery, as well as fluctuations in the ice flux and meltwater discharge. This methodology invites detailed studies of marine moraines. Systematic case studies on a regional scale may provide new information on the behaviour of tidewater ice margins and lead to unprecedented insights into the dynamics of ancient ice sheets.

Modern glaciological research and realistic numerical modelling indicate that the behaviour of tidewater glaciers (Fig. 1a) does not only reflect the bulk ice-sheet dynamics, but may actually affect and control the latter (Alley et al. 2004, 2008; Nick et al. 2009). The grounding-line sedimentation and kinematics, particularly in fast-flowing ice sectors, may therefore be of crucial importance to an understanding of the present and future ice-sheet dynamics (Boulton et al. 2001; Forsstro¨m et al. 2003; Alley et al. 2007; Bamber et al. 2007). The same relationship can be expected for ancient ice sheets, to understand better their growth, oscillations, recession and demise. But can the existing models be readily used for this purpose? The glaciological observations and ice-sheet model calibration have necessarily been based on the two existing continental ice sheets in Greenland and Antarctica which occur at comparable latitudes; both have reacted spectacularly to interglacial conditions. However, the models have thus far failed to reproduce the observed ice-sheet responses to recent global warming (Alley et al. 2008) which indicates a need for a better understanding of the critical controlling factors, particularly the behaviour of glacier marine termini. The climatic conditions in Greenland and Antarctica have not been the same, and the response of their ice masses to the past and present climate changes has been

different (Rignot & Thomas 2002). The loss of ice mass in Greenland, which has doubled its rate in the last decade (Rignot & Kanagaratnam 2006), occurs mainly by surface melting and outlet calving. The ablation in Antarctica involves primarily calving and basal melting of ice shelves (Bamber et al. 2007). Models derived from one ice sheet may therefore not apply to the other, and hence also their usefulness as analogues for ancient ice sheets is uncertain. Furthermore, the present ice margins show great lateral variation in their behaviour (Alley et al. 2004; Joughin 2006; Holland et al. 2008; Van den Broecke et al. 2009), as can also be expected for their ancient counterparts (Lønne et al. 2001). An insight into the dynamics of an ancient ice sheet may therefore require an understanding of its margin variable and differential behaviour. However, can the kinematics of an ancient tidewater ice margin be recognized on a local basis and with a temporal resolution considerably higher than that of conventional time–distance plots? This important methodological issue, addressed originally by Lønne (2001), is discussed further in the present paper with an elaboration on the range of palaeoglaciological field criteria. The suggested methodology – advancing the concepts of allostratigraphic mapping, architectural elements and facies analysis in the study of marine moraines – derives

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 57– 75. DOI: 10.1144/SP354.4 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. (a) A modern stage for the processes discussed in this paper: the front of a tidewater glacier in a Spitsbergen fjord, Norwegian Arctic. (b) Processes accompanying the advance of a tidewater glacier, with the conveyor-belt sediment delivery by a subglacial deforming bed, sediment outwash deposition by meltwater flow and the formation and erosional recycling of a grounding-line wedge of deposits (submarine-fan unit A); the growth of the sedimentary wedge reduces calving and allows the glacier front to advance in deep water, until it grinds to a halt on a seafloor sill. (c) The deposition of a marine moraine (glacier front not shown; only its grounding-line positions (GLP) and the erosional and depositional effects of its advance and stillstand are indicated); the cannibalization of ice-contact fan declines as the glacier starts grinding to a halt (allostratigraphic unit A), the submarine fan grows during the stillstand (unit B) and evolves into a delta after reaching the sea surface (unit C); MGLP is the maximum (outermost) grounding-line position.

from glaciomarine deposits formed in high-relief settings during the last deglaciation, but may conceivably apply to well-exposed analogous deposits of any age.

The significance of marine moraines The term ‘moraine’ is used here in its general geomorphic meaning to denote a sedimentary ridge formed at the grounding line of a glacier. Moraines have traditionally been mapped as topographic ridges, with their ice-side slope considered to represent the ice-contact surface. On this basis and

with the aid of radiocarbon dates, the so-called time–distance plots of ice-margin shifts have been constructed (as exemplified by Walker 1995) to draw inferences about regional climatic fluctuations. An undeclared assumption in this approach is that the ice margin would advance and retreat consistently, though not necessarily uniformly. This assumption, however, is clearly at odds with the highly inconsistent pattern of documented recent and historical ice-margin oscillations reported from the coasts of Greenland, Svalbard and Alaska (Hagen et al. 1993; Williams & Ferrigno 1995; Weidick 1995). It has been shown that a marine moraine in its coeval lateral segments

KINEMATICS OF TIDEWATER ICE FRONTS

may not only have quite a different depositional history but may record some strikingly different styles of ice-margin behaviour, such as the margin advance in some segments occurring concurrently with its retreat in adjacent segments (Lønne et al. 2001). It has also been demonstrated that the moraine ice-side slope cannot really be regarded as the true ice-contact surface (Lønne 1995, 1997; Nemec et al. 1999; Lønne et al. 2001). In short, the geomorphic mapping and isotopic dating of moraines is one thing, but an understanding of the moraine’s development and its regional significance for ice-sheet dynamics is quite another. Recent sedimentological studies have shown that marine moraines afford unique yet little-explored archives of stratigraphic information on the shortterm behaviour of glacier front, unattainable from any other sources (Lønne 1993, 1997, 2005; Lønne & Syvitski 1997; Nemec et al. 1999; Lønne et al. 2001). Moraine thicknesses of up to a hundred metres or more indicate rapid sediment accumulation in front of temperate glaciers with high ice flux; the moraines therefore potentially bear a high-resolution proxy record of the glacier margin behaviour and grounding-line processes. This record can be deciphered by an allostratigraphic mapping of the moraine’s internal architecture and analysis of its sedimentary facies (Lønne 1995, 2001). The moraines in transverse cross-sections are simple or composite wedges of grounding-line deposits (Fig. 1b, c) formed during the glacier advance and associated minor oscillations. Our previous case studies of a range of marine moraines (see references above) have focused mainly on the foreset clinothems and deltaic topset deposits of an evolving ice-contact system. A more recent study of another moraine in south-eastern Norway (Lønne & Nemec 2011) has shown that the intricate architecture of the moraine ice-proximal and crest part may also provide valuable information on the grounding-line conditions and kinematics. In the present paper, we compile our re-assessed observations in order to offer a broader range of methodological guidelines for an analysis of moraine internal architecture and high-resolution dynamic stratigraphy. The essential part of analysis is an allostratigraphic mapping of the moraine internal architecture with the recognition of the true ice-contact surfaces and the successive grounding-line positions, their time –space trajectory and geometrical relationship to the foreset clinothems. This approach allows the successive sediment packages and their spatial organization to be recognized, and it also reveals important stratigraphic discontinuities recording subglacial sediment bypass and/or erosion.

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Allostratigraphic mapping of moraine architecture Norwegian moraines that formed during the last deglaciation, when the ice-sheet margin was very active and the relative sea level in innermost Oslofjorden basin reached 220 m higher than today, afford some uniquely instructive cases with great accumulated thicknesses and high potential stratigraphic resolution. On the basis of a series of detailed case studies combining sedimentological analysis and ground-penetrating radar (GPR) survey, Lønne (2001) proposed a generalized allostratigraphic model for marine moraines (Fig. 2). This model serves as a predictive tool for marine moraine internal stratigraphic mapping, as an interpretive guide to their development as icecontact sedimentary systems and as a basis for their genetic comparison and correlation. The outcrop-derived model offers a useful conceptual framework for the interpretation of processes and sediment dynamics along a grounding line and for the recognition, comparison and correlation of changes in moraine development. The model points out that a grounding-line moraine may evolve from a fully subaqueous system, with foreset deposits composed of outwash sediment and subglacial diamicton (a submarine ice-contact fan sensu Lønne 1995), into a fluvial-deltaic system with foreset and bottomset deposits composed of outwash sediment (an ice-contact to proglacial delta sensu Lønne 1995). Five allostratigraphic units (A– E) are distinguished in the model (Fig. 2) to recognize the succession of main morphodynamic changes in an evolving ice-contact system. Units A and B consist of the grounding-line deposits formed during the glacier advance and stillstand, respectively, whereas unit C is the deltaic part developed once the sedimentary system has aggraded to the sea surface. Unit D comprises deposits formed during the moraine abandonment and represents mainly hemipelagic sedimentation accompanying the glacier retreat. Unit E consists of littoral deposits formed during the moraine emergence due to regional glacio-isostatic uplift. Each unit is further divided laterally into four subunits, or front-parallel zones, corresponding to particular geomorphic parts of the system. For example, unit B represents a grounding-line stillstand submarine fan and its subunits B1 –B4 (Fig. 2b) correspond to the fan’s ice-proximal (B1) and top part (B2), ice-distal depositional slope (B3) and the slope base (B4) passing into basin plain. The five units A –E and their divisions are characterized by specific sedimentary facies/processes and bedding architecture (Lønne 1995, 1997, 2005; Nemec et al. 1999; Lønne et al. 2001), as summarized in Figure 3. The model

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Fig. 2. Allostratigraphic model for the development of marine moraines (modified from Lønne 2001). (a) Moraine formed by glacier advance (submarine-fan unit A) and short stillstand (submarine-fan unit B), draped with sediments deposited during the glacier-front retreat (unit D) and the subsequent emergence by regional glacio-isostatic uplift (unit E); a ‘monoepisodic’ moraine will lack unit B and signify a negligible stillstand episode. (b) A fully-developed ‘polyepisodic’ moraine with ice-advance submarine-fan unit A, stillstand submarine-fan unit B and stillstand deltaic unit C; the moraine-draping units D and E (not shown) will be as in the upper diagram. AICS and TICS are the apparent and true ice-contact surfaces; for further explanation, see legend and text. The sedimentary characteristics of the individual units are reviewed in Figure 3.

points out that the ultimate topographic surface of the moraine ice-proximal slope (Fig. 2a) is basically an apparent ice-contact surface (AICS), even if locally coinciding with the true ice-contact surface

(TICS), and that the TICS as such tends to shift with the moraine growth and becomes buried beneath the moraine-draping deposits (subunits D1 and E1).

KINEMATICS OF TIDEWATER ICE FRONTS

As discussed by Lønne (2001), a simple ‘monoepisodic’ moraine, composed of a solitary unit A draped with units D –E, is a product of brief icemargin advance and indicates a highly mobile margin without significant stillstand. Composite ‘polyepisodic’ moraines, comprising units A –B (Fig. 2a), A –B –C (Fig. 2b) or their more complex configurations covered with units D –E, indicate a short (Fig. 2a) or more prolonged (Fig. 2b) stillstand with a lasting ice flux. Polyepisodic moraines commonly bear the record of minor, short-term oscillations of the grounding line (Lønne et al. 2001), as discussed further in the present paper. The grounding line, defined as the boundary between subglacial and proglacial realm, may coincide with the ice-front position or be considerably offset from the latter in the case of a floating tidewater front (Fig. 1b) or an ice shelf. The grounding-line position (GLP) is determined allostratigraphically at the intersection of subglacial ice-contact surface with the apex of the underlying package of fan foreset clinothems (GLP-1 to GLP-6 in Fig. 1c). Special attention is given to the kinematic significance of the GLP space –time trajectory in moraine cross-section. The maximum grounding-line position generally coincides with the ultimate apex of the submarine fan (MGLP in Fig. 1c), which may be below the sea surface in the case of a short stillstand (Fig. 2a) or be at the sea level in the case of delta formation (Fig. 2b). However, the glacier front may oscillate and the MGLP may thus not necessarily be the last grounding-line position pre-dating the glacier retreat and moraine abandonment (Lønne et al. 2001).

Moraine as a sedimentary recorder of glacier dynamics A tidewater glacier (Fig. 1a) will advance as long as the calving process is outperformed by the ice flux. Calving is curtailed by sediment accumulation at the grounding line, and hence the availability of sediment is a critical factor (Alley et al. 1989). The grounding-line deposits tend to be eroded and recycled as the glacier moves forwards (Alley 1991). The general notion has therefore been that the glacier advance leaves little or no deposits and that the resulting sedimentary succession may look just like a reverse stratigraphic image of the succession left by glacier retreat (Powell & Alley 1997; Powell & Cooper 2002; Trabant et al. 2003; Corner 2006). The sedimentary record of glacier advance was then expected to be negligible and drew little attention. However, large volumes of sediment tend to be transferred to the grounding line by meltwater

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flow and in connection with subglacial deforming bed (Boulton 1986; Alley et al. 1997; Murray 1998). An inevitable consequence of the transfer and conveyor-belt recycling process is an increasing sediment accumulation at the grounding line which renders the ice front thinner, reduces erosion and makes the glacier climb over its own deposits. This process is greatly enhanced when the glacier in a shelf trough or a fjord encounters a seafloor topographic sill, such as a bedrock threshold. Allostratigraphic case studies have indeed demonstrated that a major part of marine moraines consists of sediment accumulated during the glacier advance (Lønne 1993, 1995, 1997, 2005; Lønne & Syvitski 1997; Nemec et al. 1999; Lønne et al. 2001; Lønne & Nemec 2011). Examples reviewed in Table 1 indicate that thick moraines consist chiefly of a solitary or multiple unit A, formed through the ice-margin advance (Fig. 2a). The stillstand succession of units B or B–C (Fig. 2b) is volumetrically less prominent and is virtually lacking in swarms of shelf-hosted smaller moraines (Lønne et al. 2001). The advance-episode deposits also bear an important record of icemargin kinematics. Allostratigraphic case studies (review in Table 1) indicate that thick marine moraines consist chiefly of a succession of solitary or multiple units A, formed through the ice-front advance (Fig. 2a). It was suggested by Evans et al. (2006) that a glacier with a stationary margin, as opposed to an oscillating margin, will quickly exhaust its supply of sediment to the grounding line. However, the evidence of significant sedimentation during stillstand episodes (Table 1) demonstrates that the grounding line may be stationary, while the ice flux and sediment delivery persist. The development of an icecontact delta and its considerable progradation (unit C in Fig. 2b) are clear proof of abundant sediment supply to a stationary ice margin (Nemec et al. 1999). One might argue that to remain stationary with persisting ice flux and sediment supply, the grounding line will require a delicate balance of several variable factors such as the ice-flux rate and the rate of subglacial sediment accretion, as well as the ice thickness and ablation rate. If so, the available evidence from Norwegian moraines (Table 1) indicates that such conditions were relatively common during the late Weichselian to early Holocene ice-sheet re-advances. Recent discussions of subglacial processes (Boulton et al. 2001; van der Meer et al. 2003; Evans et al. 2006) suggest that the glacier substrate is a mosaic of warm- and cold-base states, involving sliding, shear deformation and deeper sediment detachment, with the relative role of sediment deformation, erosion and accretion changing rapidly in

62 I. LØNNE & W. NEMEC Fig. 3. An illustrated review of the main sedimentary characteristics of moraine allostratigraphic units and subunits distinguished in the model (Fig. 2). Outcrop photographs are from Norwegian marine moraines formed during the last deglaciation.

KINEMATICS OF TIDEWATER ICE FRONTS

Fig. 3. Continued.

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Table 1. Comparison of some marine moraines studied on an allostratigraphic basis. Moraine classification into water-depth categories (discussed in the text) pertains to the basin-floor site bathymetry prior to the moraine emplacement. The categories of ice-contact zone (erosional, accretional, compound or aggradational) are as discussed in the text Basin water depth

Shallow water: 0– 175 m

Intermediate: 175 –350 m

Deep water: 350 to .1000 m

Moraine allostratigraphy Stacking succession of allostratigraphic units (and data basis)

Relics of older moraine

A– A– A–A –A –A (outcrops) A– B– C– C– D/E (outcrops) A– A– B– C– D–E (outcrops & GPR sections) C– A– D/E (outcrops) A– B– D– Ad– D–E Ad ¼ deformed unit A (outcrops) A– A– B– D/E (outcrops) A– A– A–A –B (seismic sections) A– A– B– A (seismic sections) A– A– A–B (seismic sections)

Moraine ice-proximal side

Moraine crest

Basin water depth behind the moraine ice-proximal side

Max. thickness of subglacially affected deposits and character of ice-contact zone

Dip of glacier advance surface

Evidence of ice-margin oscillations (O) or re-advance (R)

2

170 m

1 m depositional

Subhorizontal

þ

c. 175 m

Non-exposed

þ

c. 200 m

2

Ice-distal side

Water depth at seafloor sill

Moraine thickness

Upper fan-foreset deposits rich in diamicton or seismic signature suggestive of diamicton facies

–

c. 70 m

c. 70 m

In units A (30 –35 m)

Non-exposed

–

100 – 150? m

c. 175 m

Non-exposed

c. 5 m compound

10 –158

O

60 m

c. 60 m

Poorly exposed

Seafloor high

1 m depositional

Non-exposed

–

increasing to 18 m

5m

Non-exposed

2

c. 200 m

c. 5 m erosional

10 –158

R

290 m

c. 100 m

In unit A (30 –40 m) and unit Ad (10 m)

2

c. 300 m

10 –158

–

200 m

c. 110 m

2

DTA

c. 7 m aggradational DTA compound

DTA

O

c. 400 m

??

In unit A (c. 30 m) and unit B (25 m) Evidence of discordant lenticular bodies

2

DTA

DTA compound

DTA

O

c. 500 m

c. 70 m

–

þ

DTA

DTA compound

DTA

R

c. 600 m

c. 70 m

Evidence of discordant lenticular bodies

Abbreviations: GPR, ground-penetrating radar survey; DTA, difficult to estimate (seismic section).

Locality and reference

Skardmunken moraine, Norway (Lønne 1993; allostratigraphy added) Kregnes moraine, Norway (Nemec et al. 1999) Mona moraine, Norway (Lønne et al. 2001) Bolterdalen moraine, Svalbard (Lønne 2005; allostratigraphy added) Egge-Lynga˚s moraine, Norway (Lønne 1997) Storsand moraine, Norway (Lønne & Nemec 2011) Moraine 7, Lake Melville, Canada (Lønne & Syvitski 1997) Moraine 5, Lake Melville, Canada (Lønne & Syvitski 1997) Moraine 3/4, Lake Melville, Canada (Lønne & Syvitski 1997)

KINEMATICS OF TIDEWATER ICE FRONTS

both time and space. The coupling of substrate conditions with the glacier profile and grounding-line dynamics implies that the moraine record may reflect changes not only in the glacier front zone, but in the bulk state of the glacier and its subglacial conditions. For example, a seasonal increase in meltwater yield may affect sediment transfer and ice-flow rate which may further influence ice thickness, subglacial erosion and ice-margin behaviour. The sedimentary record of such dynamic pulses may be limited to the moraine’s ice-proximal part or may extend to its ice-distal side as changes in the architecture and facies composition of foreset clinothems (Fig. 2). In short, a whole range of valuable palaeoglaciological information can potentially be deciphered through a detailed analysis of suitable outcrop crosssections of marine moraines.

The moraine record of grounding-line behaviour Detailed allostratigraphic studies of marine moraines have indicated that various signatures of grounding-line behaviour are recognizable in three main parts of a moraine: the moraine-front foreset deposits, the subglacially affected ice-proximal part and the transitional zone of moraine crest (Table 1). The characteristics of these deposits are reviewed briefly in Figure 3 and are described in detail in our case studies (Lønne 1995, 1997, 2005; Nemec et al. 1999; Lønne et al. 2001; Lønne & Nemec 2011).

The record of moraine-front foreset deposits (subunits A3 – C3) The fan-slope foreset deposits emplaced during glacier advance (subunit A3, Fig. 2b) have a distinct architecture reflecting the upper, middle and lower slope segments, but may form a range of configurations depending on the grounding-line behaviour. A simple and well-preserved subunit A3 is characterized by steep to oversteepened upper-slope clinothems as a result of the submarine fan progradation coeval with the forward movement of the grounding line and subglacial deformation. The fan-head deposits thus tend to be truncated by the glacier (TICS-A in Fig. 2), although their youngest part may be preserved in the fan crest zone. The upper-slope part of the fan foreset is characterized by uneven and highly irregular bedding due to the alternation of depositional and erosional processes directly outside the grounding line, combined with glaciotectonic influence at the ice contact. Debrisflow mounds and diamicton deposits alternate with scours formed by hyperpycnal meltwater

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effluent. Ice-rafted debris (IRD) and shallow chutes with backset bedding formed by underflow current are common. In the downdip direction, these irregular beds interfinger with and pass into the more tabular beds of hyperpycnite and debrisflow deposits that predominate in the middle-slope part of the fan foreset; IRD and isolated chute-fill features add some irregularity to the even, parallel bedding. The lower-slope part is characterized by less inclined parallel bedding with backset-filled chutes, passing into the mounded fan-toe depositional lobes of sandy hyperpycnites at the transition to the subhorizontal bottom-set subunit A4 (Fig. 2b) composed of sandy turbidites interbedded with mud. There is a general downslope decrease in gravel/ sand ratio, although mounds and lenses of gravel derived from upper-slope collapses may occur even near the slope base. The configuration of the fan foreset clinothems reflects the sediment recycling process and the interplay between moraine aggradation and the related thinning/protrusion of the ice margin (Fig. 1b). As the ice thins and substrate erosion decreases, the moraine aggradation continues with the grounding line climbing onto the ice-proximal slope. The moraine architecture can then be used to assess the extent of subglacial erosion, as the changing depth of erosion is indicated by the degree of preservation of relic unit A beneath the glacier-advance surface (as shown schematically in Fig. 4). In addition, the stacking pattern of clinothem packages and the space –time trajectory of fan-toe depositional lobes (Lønne 2001) may serve to recognize changes in the aggradation and build-out rate of the submarine fan. The sedimentary succession recording glacier advance may form a simple monoepisodic moraine (Lønne 2001) with no major discontinuities within the fan foreset (subunit A3) and with welldefined ice-contact surface and fan-crest morphology. However, the foreset succession is commonly compound and shows a complex stratigraphic configuration with scoop-shaped internal reactivation surfaces, downlapping or onlapping bedding pattern and a back-stepping ice-contact surface (increments 1A– 3A in Fig. 2b). This evidence of slope failures and major fluctuations in sediment delivery signifies grounding-line oscillations. The fan foreset deposits formed during the stillstand episode (subunit B3, Fig. 2b) resemble those of subunit A3, but show a pronounced aggradation. As the grounding-line stabilizes, the fan crest grows higher and broader and the clinothems become increasing sigmoidal, flattening in the updip direction (Fig. 4), while the calving with IRD delivery gradually ceases. When the fan crest eventually reaches storm-wave base and becomes significantly

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Fig. 4. Key characteristics of an ice-contact zone passing from deep to shallow water, with the glacier front advancing and coming into a stillstand (ice not shown). GLP are the successive positions of grounding line and MGLP is its maximum (outermost) position. The allostratigraphic sedimentary units A– C are as shown in the model (Fig. 2), illustrated in Figure 3 and discussed in the text.

KINEMATICS OF TIDEWATER ICE FRONTS

influenced by marine background processes, its sediment storage capacity declines and the build-up of foreset subunit B3 recognizably shifts basinwards. The sedimentary facies of the upper fan slope and the relative thicknesses of foreset subunits A3 and B3 (Figs 2a & 4) may thus provide information on the changing mode and volume of sediment delivery to the grounding line (Lønne & Nemec 2011). Fan-head chutes/channels and the predominance of tractional deposits accompanied by cobbly to bouldery debrisfalls from gravel lag collapses indicate high meltwater discharges, whereas abundant debrisflow deposits indicate resedimentation of material accumulated at the fan crest during lower discharges. Diamicton is an important indicator of significant ice flux (Lønne & Nemec 2011). However, diamicton tend to be destroyed by meltwater jets; in addition, the diamictonbearing deposits of fan crest and adjoining upper slope (subunits A2/B2 and uppermost A3/B3 in Fig. 2b) are often eroded (Fig. 5c, d), which may give the false impression of little or no sediment delivery from the subglacial deforming bed. Special attention should then be given to the occurrence of diamicton relics scattered in the remaining part of foreset subunits A3/B3 (Nemec et al. 1999). Once the submarine fan has aggraded to the sea surface, an ice-contact delta develops with the initial mouth-bars evolving into a Gilbert-type foreset downlapping the pre-existing fan slope (subunit C3, Fig. 2b). The most distinct feature formed in this episode is the glaciofluvial topset (subunit C2). The contact of deltaic clinothems to the topset deposits may vary from sharply angular to tangential (Nemec et al. 1999; Lønne et al. 2001), recording the delta-front regime whether fluvial- or wave-dominated (Colella 1988).

The record of subglacially deformed deposits (subunits A1 – C1) The record of subglacial regime occurs on the ice side of MGLP, in the ice-contact zone of subglacial influence (subunits A1, B1, C1 and their subglacial bounding surfaces TICS-A to TICS-C, Fig. 2b). The ice-contact zone commonly involves multiple surfaces that may: form back-stepping en-echelon segments bounding successive clinothem packages; separate sediment wedges accreted at the icecontact surface; at least locally merge into a solitary surface; and show secondary glaciotectonic disturbances. The zone of subglacial influence (Table 1) bears some of the most important record of grounding-line behaviour. The subglacially affected deposits may form a simple homogeneous layer, as is often the case in

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monoepisodic moraines, or a thicker and more complex layer as is common in thick polyepisodic moraines. Although a basal till sheet is generally characteristic of temperate or polythermal ice masses, it may occasionally form at the base of cold ice (Evans et al. 2006) and thus is not environmentally diagnostic. As long as the ice flux persists, a till may form under a range of conditions such as during the grounding-line advance and retreat, at a low or high sediment-flux rate and onshore as well as in shallow or deep water. An analysis of the subglacially affected zone should focus on the character of its lower and upper boundaries, its thickness and possible lateral variation (Fig. 4). The lower boundary is commonly distinct, with a simple or complex truncation of the underlying foreset deposits. The character of the whole zone or its particular segments may be (Fig. 2b): (a) erosional, showing evidence of net subglacial erosion; (b) accretional, showing evidence of net sediment accumulation and formation of a simple till layer; (c) compound, with multiple ice-contact surfaces and a combination of accretion, erosion and thrusting; or (d) aggradational, with multiple ice-contact surfaces and a recognizable stacking of successive compound and/or accretional layers. Multiple ice-contact surfaces indicate grounding-line oscillations (Fig. 5). The ice margin may briefly retreat and deposit a new wedge of grounding-line sediment, which is then squeezed and pressed against the former ice-contact surface during the glacier re-advance. The result will be a plastered wedge-shaped aggradational package of deformed deposits with little or no corresponding growth of foreset clinothems (Fig. 5a), depending on the extent of the grounding-line oscillation. A thick advancing glacier is highly erosive, scouring the substrate and sweeping sediment from seafloor humps. Even if the sediment flux is high and the ice-contact fan reaches considerable thickness, the ultimate record of subglacial sedimentation may be little more than a till sheet. A thicker moraine wedged underneath the glacier front will render it thinner and reduce its erosional capacity whereby sediment may increasingly accumulate at the glacier sole, particularly in the upper part of the moraine ice-contact slope where the ice thickness approaches its contemporaneous minimum. Substrate scouring may continue in the deeper, upflow zone of the basin, whereas the thinning glacier margin in a feedback may allow rapid increase of the moraine thickness, volume and morphometric profile (Fig. 4). In thick moraines formed by a glacier advance ended in stillstand, unit B generally reflects the prevalent recycling of sediment from the ice-proximal to ice-distal side of the moraine (Lønne & Nemec 2011).

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Fig. 5. Architectural characteristics of ice-contact zone, with multiple ice-contact surfaces and changing of grounding-line positions. (a) Grounding-line oscillations on the stoss side of moraines formed in shallow-tointermediate water depths; the encircled numbers 1– 4 represent successive grounding-line positions. Note how the small submarine fan formed at the grounding-line position 3 (in diagram on the left) was squeezed and plastered onto the previous ice-contact surface (TICS 1A-2A) during the glacier re-advance (diagram to the right), forming a strongly deformed unit onlapping this surface, whereby the depositional record from position 3 is effectively destroyed. (b) Ground-penetrating radar section through a moraine that formed on a seafloor sill and evolved into an ice-contact delta (Lønne et al. 2001), showing a compound ice-contact zone close to sea level with a signature of grounding-line oscillations. (c, d) Interpreted seismic sections through moraines formed in a deep-water fjord (moraines 7 and 3/4 from Lønne & Syvitski 1997); moraine 7 (to the left) was deposited at a water depth of c. 450 m and moraine 3/4 (to the right) at a depth of c. 550 m; vertical scale exaggerated (cf. Fig. 7).

The subglacially emplaced sediment may derive from the cannibalization of the groundingline moraine and be mainly coarse-grained and well washed, or may abound in mud and diamicton derived by the glacier through deforming bed from distant areas. The composition of subglacial deposits is key to the sediment provenance and the upflow extent of subglacial conveyor-belt transport. If the glacier stillstand time is

considerable and ice flux declines, the iceproximal slope of the moraine will be increasingly eroded and its sediment transferred to the grounding line. If significant ice flux however persists, diamicton will continue to be delivered from the glacier sole to the moraine ice-contact slope and become partly entrapped there. This will be manifested by a predominantly diamictic and nonuniform composition of subglacially accreted

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deposits and by their significant thickness (Lønne & Nemec 2011).

The record of moraine crest deposits (subunits A2 – C2) The crest part of the moraine (subunits A2, B2 and C2, Fig. 2b) records the transition from subglacial to proglacial processes and sediment dynamics. The sediment storage potential here is at the lowest for the entire proglacial zone and decreases further as the moraine crest builds up, reaches the wave base and slowly aggrades to sea level. The crest is directly affected by the high-pressure meltwater jets, which sweep sediment and form crosscutting channels extending into slope chutes. The violent hydraulic regime combines with glaciotectonic deformation driven by glacier push causing proglacial folding, thrusting and vertical piling of sediment, most of which is then quickly removed by the prevalent erosion. Deposits preserved at the moraine crest as subunits A2 and B2 (Fig. 2b) thus provide a valuable record of the grounding-line conditions and increasing sediment-storage capacity (Lønne & Nemec 2011). Subunit C2 chiefly bears the record of proglacial fluvial drainage, but may include evidence of a transient activity of tidal currents within the narrow time interval between the moraine abandonment and its full emergence. There are cases where the uppermost part of the flowing ice had apparently decoupled itself from the main ice body and formed a protrusion overriding the moraine (Fig. 1b; Lønne 2001). If the moraine crest at this stage is at the sea surface, the ice protrusion will not obstruct the formation of delta but will necessarily affect the delta growth by scouring its fluvial distributary plain. If the moraine crest is still below sea level however, the ice protrusion may form a subhorizontal till layer at the flattruncated fan top (unit A, Fig. 3). This may then be mistaken for a deltaic topset in seismic or GPR sections and lead to an erroneous estimation of the local sea level. A till can readily be distinguished from stratified fluvial topset in an outcrop section, but an independent assessment of the local marine limit may be needed in a study based on remote sensing.

Bathymetric control on moraine allostratigraphy Case studies suggest that the allostratigraphic complexity of marine moraines increases with the contemporaneous water depth, as well as with the ice thickness and flux rate (Lønne 1993, 1995, 1997, 2005; Lønne & Syvitski 1997; Nemec et al. 1999;

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Lønne et al. 2001). The available data are too limited to draw any definitive conclusions, and the following suggestions are therefore little more than a preliminary attempt to assess the controlling role of basin bathymetry. The distinction of shallow (,175 m), intermediate (175–300 m) and deep water (300 to .1000 m) conditions is somewhat arbitrary, although based on case studies.

Moraine formation in shallow water The grounding-line depositional system in shallow water (,175 m) tends to be relatively simple, forming a small monoepisodic moraine (solitary unit A) or slightly more composite moraine (two or more units A), with or without minor stillstand unit B and seldom with an incipient deltaic unit C (Lønne 2001, fig. 2b –e). The advancing glacier in shallow water is relatively thin and its erosional capacity is low, whereby the fan foreset deposits tend to be well preserved. The glacier front is such settings is more mobile, seldom coming to a longer stillstand. Subglacial influence is usually restricted to the ice-proximal side of the moraine, but an ice protrusion may form and partly or fully override the moraine. The glacier-advance surface and subglacially affected deposits should therefore be carefully recognized and mapped in the moraine cross-section (Fig. 5a, b). In the very shallow water of littoral to sublittoral shelf (50 m), the glacier advances over a low-relief seafloor. The resulting moraine therefore tends to have a sheet-like geometry because the groundingline deposits are easily overridden and truncated by the glacier (Fig. 6a; examples in Landvik et al. 1992; Lønne 2005). If formed in somewhat deeper (.50 m) water, the moraine assumes a ridge-like topography and the glacier advance may culminate in a brief stillstand (Fig. 6b; Lønne 1993) as the expanding ice becomes increasingly pinned down to the seafloor topography. In general, the grounding-line advances in such settings leave a record of short-term oscillations with occasional minor stillstands (Fig. 5a). The moraine ice-contact surface may be undulating, with the convex-upwards parts representing sediment accretion after the grounding line reached its maximum position and underwent oscillations prior to retreat (Fig. 6c, d). An even and gently inclined ice-contact surface will indicate prevalent subglacial erosion and bulk sediment transfer to the moraine ice-distal side. Since the glacier in shallow water is relatively thin and has a limited erosional capacity, the preservation potential of older moraines overridden by the glacier in such settings may be relatively higher (Table 1). Special attention should be given to the

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Fig. 6. Schematic diagrams based on case studies showing diverse trajectories of grounding-line migration during the formation of marine moraine; only the ice-related units A, B and C are shown, whereas the post-retreat units D and E are omitted (Fig. 2). (a) Simple glacier advance in shallow water with low sediment supply, resulting in a flat-top moraine almost completely overridden by the glacier. (b) Glacier advance with two stillstand episodes and no aggradation to sea level, but with 2/3 of the broad and flat moraine crest overridden by the glacier. (c) Glacier advance

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recognition of possible moraine relics overlain by younger moraines (Fig. 1c).

Moraine formation in deep water Moraines formed in deep water (300 to .1000 m) are generally thick and polyepisodic, composed of multiple allostratigraphic units recording grounding-line advance and significant stillstand, but rarely showing the development of Gilbert-type delta. The ice-contact surface is concave upwards and commonly smooth (Fig. 6f, g), suggesting a simple continuous advance of the grounding line. Frequent minor oscillations can be expected, since rapid calving will occur in deep water as soon as the glacier loses its contact with the moraine (Benn et al. 2007). The ice in deep water is thick and its advance erosive, whereby scouring prevails on the icecontact surface and often removes the record of minor grounding-line oscillations. At the lower range of deep water (c. 300 –500 m), the ice will be thinner and less erosive hence allowing significant grounding-line oscillations to occur and be recorded even if no major ice-margin retreat by calving takes place (Fig. 5c, d).

Moraine formation at intermediate water depth Moraines developed at intermediate water depths may combine characteristics of the two previous categories, but are generally composite with multiple ice-advance units A and a compound or markedly aggradational ice-contact zone (Fig. 6e). The glacier front typically endures minor oscillations, with ice-mass recovery and reoccupation of the moraine ice-proximal slope. An onlapping geometry of till layers aggraded on the ice-proximal slope (Fig. 6e) indicates a glacier front that was retreating from the moraine crest, forming a grounding-line sedimentary wedge and then plastering it to the icecontact surface during re-advance (Fig. 5a).

Comparison of moraines The recognition of allostratigraphic units A, B and C (Fig. 2) and mapping of their stacking architecture are the primary basis for a comparative study of marine moraines. Other descriptive characteristics

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relevant to the comparison of moraines and their morphodynamic evolution are: (a) the number and sequencing of allostratigraphic units recognized (such as the succession A –A –A –A –B in Fig. 5c); (b) the moraine maximum thickness, as measured from the seafloor sill to moraine crest (c. 70 m in Fig. 5c); and (c) the vertical distance from the lowest point of a rising glacier-advance surface to the contemporaneous sea level – taken at the delta plain or local marine limit (c. 400 m in Fig. 5c) – as a proxy of the ice thickness. Table 1 summarizes the key characteristics of moraines thus far studied on an allostratigraphic basis. All these moraines contain single or multiple wedges of ice-advance deposits (unit A), although the allostratigraphic complexity seems to increase with the moraine thickness and hence broadly with the basin water depth and ice thickness. Wedges of stillstand deposits (unit B) occur both intercalated with units A and as a culmination of the latter. Some moraines lack unit B (Table 1), which means that they were abandoned by the glacier soon after its advance episode without a significant stillstand. The common occurrence of irregularly bedded, diamicton-rich facies in the head part of submarine fan foreset indicates sediment delivery from a subglacial deforming bed, and thus a high ice flux during the glacier advance. Particular attention should be paid to the moraine’s ice-contact deposits. For example, the GPR section in Figure 5b (110 m long) shows the iceproximal side of a moraine that evolved into ice-contact delta (Lønne et al. 2001). The upper boundary of the moraine foreset, at the top of the succession of subunits A3 –B3 – C3, is recognizably compound with subhorizontal or gently inclined lenticular bodies passing into packages of foreset clinoforms. Well recognizable is also the upper boundary of subglacially affected deposits, corresponding to the basal surface of subunits D1/E1 that has been correlated to outcrop section (Lønne et al. 2001, fig. 16) and postdates the ice detachment from the moraine. Notably, this part of the growing moraine was close to the sea level (Fig. 5b). The undulating boundary with truncated bedding reflectors and discordant lenses in the ice-contact zone indicate subglacial erosion and thrusting at the top of unit A, probably due to minor oscillations of the grounding line. For comparison, the interpreted seismic sections in Figure 5c, d are from two deep-water

Fig. 6. (Continued) with the grounding-line stillstand, minor retreat and re-advance and with the moraine aggrading close to sea level. (d) Glacier advance, brief stillstand and slow retreat with several grounding-line oscillations, with each new sediment wedge plastered to the previous ice-contact surface during the glacier re-advance. (e) Glacier advance with a high sediment supply and strong moraine aggradation, interrupted by a brief episode of retreat and re-advance, and culminating in a stillstand without aggradation to sea level. (f) Smooth erosional advance of a glacier grinding to a halt, with the moraine aggradation terminated by rapid ice-margin retreat. (g) Similar, but with a stronger aggradation and a stillstand episode (unit B).

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Fig. 7. Longitudinal cross-section through the Lake Melville fjord basin, showing its bottom topography and the interpreted seismostratigraphy of its deposits (modified from Lønne & Syvitski 1997). Seismostratigraphic units 2–13 were deposited during the fjord deglaciation. The fjord was deglaciated in c. 1000 years, and the ice-margin movement plot at the top is an interpretation of the main grounding-line shifts during this time.

moraines (Lønne & Syvitski 1997). They demonstrate multiple allostratigraphic units A and icecontact surfaces indicative of significant groundingline oscillations with a markedly steepening grounding-line trajectory and culmination in a stillstand episode (unit B), but far from reaching the sea surface. A systematic comparative study of ancient moraines will inevitably encounter the same difficulties that are met in all other geological field studies, namely the requirement of good outcrops and the limitations of exposure. Moraine outcrop sections generally afford the highest sedimentological resolution but limited lateral control as compared to GPR or shallow seismic sections, and the marine moraines exposed by regional glacio-isostatic uplift are most commonly those formed in shallow-to-intermediate water depths of the glaciated basin. Field studies therefore need to be carefully planned and conducted on the basis of best outcrop sections, while taking good account of the evidence from poorer outcrops and possibly remote sensing.

Interpretation of glaciomarine basin-fill successions In order to highlight the importance of glacieradvance episodes in the filling of glaciated marine

basins, the stratigraphy of Lake Melville fjord basin (Lønne & Syvitski 1997; Syvitski & Lee 1997) is reassessed here on the basis of the elaborated range of criteria. This Canadian fjord, 65 km long, was deglaciated in c. 1000 years and accumulated sediments with a total thickness of 400 m, comprising 13 distinguishable seismostratigraphic units (Fig. 7). Three prominent moraines have been recognized in units 3/4 (diffuse boundary), 5 and 7. Three episodes of glacier recession were recognized (as shown by the ice-margin movement plot in Fig. 7): a rapid retreat in the outer basin (unit 2), slow retreat in the middle basin (units 327) and full ablation with glacier demise in the inner basin (units 8213). The study by Lønne & Syvitski (1997) demonstrated the importance of allostratigraphic mapping by revealing the architecture and dynamic stratigraphy of ice-advance deposits. All three moraines were recognized to contain multiple units A (cf. Fig. 5c, d), and it was argued by Syvitski & Lee (1997) that unit 2 was likely overridden by the glacier. However, the glaciodynamic implications of these findings were not fully realized. It was concluded that the 65 km fjord was filled with sediment during the ice retreat, although what the data actually revealed was an ice recession interrupted by four major re-advances with a total distance of c. 62 km (35 km for unit 2, 15 km for moraine 3/4, 5 km for moraine 5 and another 5 km for

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moraine 7, as shown by the ice-margin movement plot in Fig. 7). The latter plot also points to a relatively slow build-up of the moraines compared to the much faster ice retreat, which is consistent with the asymmetrical pattern of ice-margin oscillations reported from contemporary tidewater glaciers (Meier & Post 1987; Trabant et al. 2003). The glacier in Lake Melville fjord, as a branch of the Laurentide Ice Sheet, was subject to dramatic thinning between 10 and 9 ka BP. It was, however, still sufficiently thick to erase the pre-existing and contemporaneous sedimentary topography when re-advancing to the outer basin with net erosion. In contrast, the subsequent re-advance to the middle basin was characterized by net deposition with a delicate balance of sediment recycling and deposition held at the grounding line, resulting in ice-advance sediment wedges intercalated with subordinate stillstand deposits. The inner basin primarily accumulated a blanket of proglacial deposits (sheet-like unit 8 in Fig. 7), indicating a thin and rapidly retreating glacier. The ice-advance deposits may therefore form a considerable part of deep fjord-fill sedimentary succession as seen in the middle part of the Lake Melville fjord basin; this requires an appropriate combination of ice flux, sediment supply and basinfloor topography. Otherwise, the moraine preservation potential in deep-water basin is low both during its deposition and afterwards since fjords and shelf troughs tend to be emptied by erosion during episodes of glacier advance.

Conclusions and implications Short-term oscillations of tidewater glaciers have been a key issue in the recent theoretical discussions and modelling of ice-sheet dynamics, with field observations from existing ice sheets. The evidence indicates that ice sheets differ in their behaviour, depending on the thermal regime, and that the kinematics of tidewater ice margins vary laterally and are both reflecting and affecting the ice-sheet dynamics. Much remains to be learned about these physical relationships, but it is clear that the icemargin behaviour bears a crucial signal and that an ice-sheet model will accordingly need to be tuned to give insights into ancient ice-sheet dynamics. For this purpose, the behaviour of former tidewater glaciers needs to be studied on a comparative local basis. Information on the local kinematic behaviour of an ancient tidewater ice margin can be derived from marine moraines, whose main part invariably consists of deposits formed during the glacier advance. The grounding-line moraine affords a highresolution allostratigraphic record of ice-margin

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dynamics, revealing changes occurring on a short timescale. In this paper we have suggested a range of field criteria for the recognition of ice-margin kinematics and short-term oscillations, based on a series of last-glacial moraine case studies. The case studies indicate that ancient outlet glaciers were far more kinematically active than had been earlier anticipated, and that the short-term oscillations and re-advances of glacier termini may therefore be of crucial importance to an understanding of ancient ice-sheet dynamics and for comparisons with modern ice margins. Allostratigraphic analysis of moraine facies architecture not only provides an insight into the icemargin dynamics, but also gives a common basis for comparison of marine moraines. Special focus should be on the ice-advance deposits (unit A); these were commonly unrecognized as such and attributed to the ice stillstand or initial retreat. They were virtually overlooked where buried under the stillstand deltaic deposits, as such moraines were simply considered to be glaciofluvial deltas. Particular attention should be given to: the recognition of true ice-contact surfaces, their geometry and spatial architecture; the textural composition (diamictic versus water-washed) of sediment bounded by these surfaces and preserved in moraine crest (subunit A2); and the geometrical pattern of submarine fan clinothems (subunit A3), their relationship to fan-crest deposits and the trajectory of fan-toe lobes (Fig. 6f, g). Stillstand sedimentation requires persistent ice flux and sufficient sediment supply, and such glacier-margin conditions are apparently short-lived, although the preservation potential of units B and C is expectedly somewhat higher. Where formed and preserved the stillstand deposits, particularly subunit B3 and early subunit C3, may provide important information on ice-flux fluctuations and possible changes in the mode and rate of sediment delivery to the grounding line. The methodological guidelines presented in this paper invite extensive detailed studies of marine moraines. Systematic case studies on a regional scale will provide new information on the behaviour of tidewater glacier termini and – in combination with appropriate modern ice-sheet models – may lead to unprecedented insights into the dynamics of ancient ice sheets. We thank E. Arnaud, I. P. Martini, J. Shaw and A. J. (Tom) van Loon for their critical reviews of the manuscript.

References Alley, R. B. 1991. Sedimentary processes may cause fluctuations of tidewater glaciers. Annals of Glaciology, 15, 119 –124.

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Alley, R. B., Blankenship, D. D., Rooney, S. T. & Bentley, C. R. 1989. Sedimentation beneath ice shelves – the view from Ice Stream B. Marine Geology, 85, 101– 120. Alley, R. B., Cuffey, K. M., Evenson, E. B., Strasser, J. C., Lawson, D. E. & Larson, G. T. 1997. How glaciers entrain and transport basal sediment: physical constraints. Quaternary Science Reviews, 16, 1070– 1038. Alley, R. B., Anandakrishnan, S., Dupont, T. K. & Parizek, B. R. 2004. Ice streams – fast, and faster? C. R. Physique, 5, 723–734. Alley, R. B., Anandakrishnan, S., Dupont, T. K., Parizek, B. R. & Pollard, D. 2007. Effect of sedimentation on ice-sheet grounding line stability. Science, 315, 1838– 1841. Alley, R. B., Fahnestock, M. & Joughin, I. 2008. Understanding glacier flow in changing times. Science, 322, 1061– 1062. Bamber, J. L., Alley, R. B. & Joughin, I. 2007. Rapid response of modern day ice sheets to external forcing. Earth and Planetary Science Letters, 257, 1–13. Benn, D. I., Warren, C. R. & Mottram, R. H. 2007. Calving processes and the dynamics of calving glaciers. Earth-Science Reviews, 82, 143– 179. Boulton, G. S. 1986. Push-moraines and glacier-contact fans in marine and terrestrial environments. Sedimentology, 33, 677 –698. Boulton, G. S., Dongelmans, P., Punkari, M. & Broadgate, M. 2001. Paleoglaciology of an ice sheet through a glacial cycle: the European ice sheet through the Weichselian. Quaternary Science Reviews, 20, 591– 625. Colella, A. 1988. Pliocene-Holocene fan deltas and braid deltas in the Crati Basin, southern Italy: a consequence of varying tectonic conditions. In: Nemec, W. & Steel, R. J. (eds) Fan Deltas: Sedimentology and Tectonic Settings. Blackie, London, 50– 74. Corner, G. D. 2006. A transgressive-regressive model of fjord-valley fill: stratigraphy, facies and depositional controls. In: Dalrymple, R. W., Leckie, D. A. & Tillman, R. W. (eds) Incised Valleys in Time and Space. SEPM, Tulsa, Oklahoma, Special Publications, 85, 161–178. Evans, D. J. A., Phillips, E. R., Hiemstra, J. F. & Auton, C. A. 2006. Subglacial till: formation, sedimentary characteristics and classification. EarthScience Reviews, 78, 115– 176. Forsstro¨m, P.-L., Sallasmaa, O., Greve, R. & Zwinger, T. 2003. Simulation of fast-flow features in the Fennoscandian ice sheet during the Last Glacial Maximum. Annals of Glaciolocy, 37, 383 –389. Hagen, J. O., Liestøl, O., Roland, E. & Jørgensen, T. 1993. Glacier atlas of Svalbard and Jan Mayen. Meddelelser, 129, 1 –141. Holland, D. M., Thomas, R. H., De Young, B., Ribergaard, M. H. & Lyberth, B. 2008. Acceleration of Jacobshavn Isbrae triggered by warm subsurface ocean waters. Nature Geoscience, 1, 659–664. Joughin, I. 2006. Greenland rumbles louder as glaciers accelerate. Science, 311, 1719–1720. Landvik, J. Y., Bolstad, M., Lycke, A. K., Mangerud, J. & Sejrup, H. P. 1992. Weichselian stratigraphy and paleoenvironments at Bellsund, Western Svalbard. Boreas, 21, 335–358.

Lønne, I. 1993. Physical signatures of ice advance in a Younger Dryas ice-contact delta, Troms, northern Norway: implications for glacier-terminus history. Boreas, 22, 59– 70. Lønne, I. 1995. Sedimentary facies and depositional architecture of ice-contact glaciomarine systems. In: Chough, S. K. & Orton, G. J. (eds) Fan Deltas: Depositional Styles and Controls. Sedimentary Geology, Special Issue, 98, 13–43. Lønne, I. 1997. Sedimentology and depositional history of an early Holocene ice-contact submarine fan: the Egge-Lynga˚s ‘end-moraine’, southern Norway. Norsk Geologisk Tidsskrift, 77, 1– 20. Lønne, I. 2001. Dynamics of marine glacier termini read from moraine architecture. Geology, 29, 199– 202. Lønne, I. 2005. Faint traces of high Arctic glaciations: an early Holocene ice-front fluctuation in Bolterdalen, Svalbard. Boreas, 34, 308– 323. Lønne, I. & Syvitski, J. P. 1997. Effects of the readvance of an ice margin on the seismic character of the underlying sediment. Marine Geology, 143, 81–102. Lønne, I. & Nemec, W. 2011. Modes of sediment delivery to the grounding line of a fast-flowing tidewater glacier: implications for ice-margin conditions and glacier dynamics. In: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 33 –56. Lønne, I., Nemec, W., Blikra, L. H. & Lauritsen, T. 2001. Sedimentary architecture and dynamic stratigraphy of a marine ice-contact system. Journal of Sedimentary Research, 71, 922 –943. Meier, M. F. & Post, A. 1987. Fast tidewater glaciers. Journal of Geophysical Research, 92, 9051– 9058. Murray, T. 1998. Assessing the paradigm shift: deformable glacier beds. Quaternary Science Reviews, 16, 995–1016. Nemec, W., Lønne, I. & Blikra, L. H. 1999. The Kregnes moraine in Gauldalen, west-central Norway: anatomy of a Younger Dryas proglacial delta in a palaeofjord basin. Boreas, 28, 454 –476. Nick, F. M., Vieli, A., Howat, I. M. & Joughin, I. 2009. Large-scale changes in Greenland outlet glacier dynamics triggered at the terminus. Nature Geoscience, 2, 110– 114. Powell, R. D. & Alley, R. B. 1997. Grounding-line systems: processes, glaciological inferences and the stratigraphic record. Geology and seismic stratigraphy of the Antarctic margin, Part 2. Antarctic Research Series, 71, 169– 187. Powell, R. D. & Cooper, J. M. 2002. A glacial sequence stratigraphic model for temperate, glaciated continen´ Cofaigh, C. tal shelves. In: Dowdeswll, J. A. & O (eds) Glacier-influenced Sedimentation on Highlatitude Continental Margins. Geological Society, London, Special Publications, 203, 215– 244. Rignot, E. & Thomas, R. H. 2002. Mass balance of polar ice sheets. Science, 297, 1502–1506. Rignot, E. & Kanagaratnam, P. 2006. Changes in the velocity structure of the Greenland Ice Sheet. Science, 311, 986– 990. Syvitski, J. P. & Lee, H. J. 1997. Postglacial sequence stratigraphy of Lake Melville, Labrador. Marine Geology, 143, 55–79.

KINEMATICS OF TIDEWATER ICE FRONTS Trabant, D. C., Krimmel, R. M., Echelmeyer, K. A., Zirnheld, S. L. & Elsberg, D. H. 2003. The slow advance of a calving glacier: Hubbard Glacier, Alaska, U.S.A. Annals of Glaciology, 36, 45– 50. Van den Broecke, M., Bamber, J. et al. 2009. Partitioning recent Greenland mass loss. Science, 326, 984–986. van der Meer, J. J. M., Menzies, J. & Rose, J. 2003. Subglacial till: the deforming glacier bed. Quaternary Science Reviews, 22, 1659–1685.

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Walker, M. J. C. 1995. Climatic changes in Europe during the last glacial/interglacial transition. Quaternary International, 28, 63– 76. Weidick, A. 1995. Greenland. In: Williams, R. S., Jr. & Ferrigno, J. G. (eds) Satellite Image Atlas of Glaciers of the World. U.S. Geological Survey, Reston, Virginia, Professional Papers, 1386. Williams, R. S., Jr. & Ferrigno, J. G. (eds) 1995. Satellite Image Atlas of Glaciers of the World. U.S. Geological Survey, Reston, Virginia, Professional Papers, 1386.

The formation of Alpine lateral moraines inferred from sedimentology and radar reflection patterns: a case study from Gornergletscher, Switzerland S. LUKAS1* & O. SASS2 1

School of Geography, Queen Mary University of London, Mile End Road, London, E1 4NS 2

Institute of Geography and Regional Science, University of Graz, Heinrichstraße 36, 8010 Graz, Austria *Corresponding author (e-mail: [email protected]) Abstract: Studies on the formation of alpine lateral moraines are rare; consequently, their internal structure and their modes of formation are relatively poorly understood. We present here sedimentological and ground-penetrating radar (GPR) data from a lateral moraine complex where an exposure allows radar facies to be compared to the field evidence. The moraine, which is slightly asymmetric with a slightly steeper distal slope, consists of alternating clast- and matrix-supported stratified diamicts and intercalated sorted sediment units which are all subparallel to the moraine surface. They are interpreted as subaerial debris flows and fluvial sediments, respectively. The GPR, using frequencies of 50, 100 and 200 MHz, allows tracing of very similar subparallel reflectors to c. 10 m into the subsurface. These are interpreted to represent the diamict units while the thickness of the sorted sediment units is below GPR resolution. Our data suggests that a two or three frequency approach supplemented by common-midpoint (CMP) measurements works well in glaciated high-mountain environments. Based on our sedimentological and GPR data we develop a conceptual framework for the formation of successive lateral moraines which involves ice-contact fan sedimentation, followed by collapse and incremental reworking of the proximal side during glacier retreat.

Studies combining both field sedimentological methods and ground-penetrating radar (GPR) to investigate Quaternary sedimentary successions have been a relatively recent addition to the pool of techniques Earth scientists can draw from. Such an approach has been popular in undulating lowlands covered by thick successions of Quaternary sediments to detect, for example, the fluvial architecture (Heinz & Aigner 2003; Bersezio et al. 2007) or deformation structures within glacial sediments (Busby & Merritt 1999; Jakobsen & Overgaard 2002; Bakker & van der Meer 2003; Helle 2004; Sadura et al. 2006; Burki et al. 2009). At the other end of the spectrum, GPR has been intensively used to understand the thickness and evolution of Alpine geomorphological features composed of coarse debris such as rock glaciers (Berthling et al. 2000; Degenhardt et al. 2003) or talus slopes (Leopold & Vo¨lkel 2003; Gerber et al. 2007; Sass 2007; Sass & Krautblatter 2007). GPR has not been extensively applied to glacial sediments in Alpine environments, however; consequently, the versatility of this method has never been explored systematically. We here present a combined approach that utilizes field sedimentology of Alpine lateral moraines and GPR measurements

at the same sites to fill those gaps. In particular, we aim to (a) test the suitability of GPR in finegrained and diamictic Alpine sediments and (b) establish the thickness and subsurface structure of sediments that make up lateral moraines commonly found in Alpine environments.

Study area Covering an area of 38.2 km2, Gornergletscher (78460 E, 458580 N) is the second-largest glacier in the Swiss Alps and is located in southern Switzerland east of the village of Zermatt (Fig. 1). This Alpine valley glacier, which currently extends from 4600 to 2200 m altitude, has been glaciologically investigated for over 30 years (Huss et al. 2007; Sugiyama et al. 2007). Gornergletscher reaches a maximum thickness of 450 m at the confluence with Grenzgletscher, and radio-echo soundings reveal that the glacier is largely temperate with only a few frozen spots (Huss 2005). The margins of the trunk glacier are largely debris-covered, which is partly the result of medial moraines being deflected by the larger and dominant trunk glacier (Fig. 1a). Geologically, the area surrounding Gornergletscher consists of tertiary metamorphic

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 77– 92. DOI: 10.1144/SP354.5 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. (a) Location of the study area in Switzerland and overall context of the plateau north of Gornergletscher; location of the study area is shown as a block dot in the upper left insert and black frame in (a). (b) Detailed geomorphological map of the study area showing the distribution and ages of lateral and annual moraines, location of GOR4 and GPR profiles on the moraines MOR 1 –3 and in ice marginal areas (IM1).

units of durable and resistant lithologies such as gneisses, garnet-rich mica-schist, quartzites and meta-granites (Bearth 1953; Ebert 2001). Aerial photographs that have been obtained by the Glaciology section of Eidgeno¨ssische Technische Hochschule (ETH), Zu¨rich at the beginning and end of each year’s ablation season since 1977 show that the northernmost subsidiary snout of Gornergletscher has formed annual moraines at the end of each winter season over this period. These features are typically 0.5–1.5 m high and up to 6 m

wide. They are constrained to a slightly undulating plateau just north of the trunk glacier (78470 3000 E, 458580 3000 N) where they appear to have been deposited on bedrock and Quaternary sediments, but the base of these forms and the total sediment thickness is unknown. Like the majority of Alpine glaciers (Maisch et al. 2000), Gornergletscher advanced in 1850, 1920 and 1980 to form prominent lateral moraines that effectively divide the plateau into three areas of annual moraines (Fig. 1b). The lateral moraines display distal slopes that are up to 6 m high

SEDIMENTOLOGY AND GPR OF LATERAL MORAINES

(AD 1850), but more commonly reach heights of up to 3 m at their distal side (AD 1920 and 1980); all of these display a slight asymmetry with steeper distal and slightly gentler proximal faces. The lateral moraines dated to AD 1920 and AD 1980 describe halfarcs that reach several tens to hundreds of metres in length, but can locally consist of a large number of ridge fragments that are 20 –50 m long and form part of longer chains (Fig. 2). These arcs, together with the more fragmented annual moraines on the adjacent plateau, trace the shape of the ice margin at the time of moraine formation.

Methods Geomorphology and sedimentology Geomorphological mapping was carried out at a scale of 1:10 000 enlarged from topographical base maps of 1:25 000 aided by panchromatic aerial photographs at a scale of c. 1:8000 and an extremely high ground resolution of c. 2 –5 m. The exposure through the lateral moraine reported here was created manually using trenching tools and spades. A measured drawing on square millimetre paper was made in the field, using marker cairns along the base of the section and a tape measure hung from the top of the exposure as a vertical scale. Prominent boulders and unit boundaries were drawn first, with increasing detail being added progressively. In order to increase planimetric accuracy, a photomosaic was taken and the field log later transferred. Individual sedimentary units were identified and distinguished on the basis of their visual physical properties including grain size range, sorting, compaction, sedimentary structures (depositional, erosional and deformation structures) and visual clast shape and roundness. The nature of contacts between individual units was

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also recorded. A slightly modified version of the lithofacies code introduced by Eyles et al. (1983) is employed for effective and rapid description in sedimentary logs (Evans & Benn 2004). The dip and strike of moraine surface slopes and selected units was measured using a Recta compassclinometer.

Ground-penetrating radar (GPR) The GPR method is based on an electromagnetic pulse emitted from a transmitter antenna, which is reflected at inhomogeneities and layer boundaries and received by a second receiver antenna after a measured two-way travel time. The reflectivity of the subsurface depends on contrasts in the dielectric permittivity which is primarily controlled by water and clay mineral content. Grain size or texture contrasts are usually related to different water contents and can thus be detected by GPR. The principles of the method are treated in detail by Bristow & Jol (2003), Moorman et al. (2003), Neal (2004) and Annan (2009). A RAMAC GPR device (Mala˚ Geosystems) was used for the measurements. The aim was both to detect the bedrock surface and to depict the internal structures and stratigraphy of the deposits that make up the moraines. For this purpose we chose frequencies of 50, 100 and 200 MHz. In general, lower frequency surveys result in a higher penetration depth and less clutter in the profiles. On the other hand, as penetration depth increases, the resolution of small objects and layers decreases. The vertical resolution of the antennas equals c. 0.25 times the wavelength which is 0.5 m for 50 MHz, 0.25 m for 100 MHz and 0.125 cm for 200 MHz at a propagation velocity of 0.1 m ns21. The subsurface propagation velocity of the radar waves was determined by five common-midpoint

Fig. 2. Photograph showing the 1980 lateral moraine with the trench of exposure GOR-4 in the foreground (arrow), annual moraines and Gornergletscher in the background (photograph taken on 14.06.2007). The 1980 lateral moraine distal slope (arrowed) is c. 3 m high.

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(CMP) measurements. In this array, the antennas are moved apart in equal increments. The ground wave velocity was determined at a first approximation from the distance/travel time diagram. The propagation velocities at greater depths were determined from velocity scans (semblance analysis) of the CMP data (Tillard & Dubois 1995) using the REFLEXW program. Jacob & Hermance (2004) report a typical accuracy of c. +0.01 m ns21. The REFLEXW program was used for data processing and interpretation. Only the smallest necessary number of filters was applied; typically, this involved static correction of the first onsets, bandpass frequency filtering and two-dimensional (2D) background removal. Due to the highly heterogeneous subsurface material, migration using a constant propagation velocity only worked for confined areas of the respective profiles; interfering artefacts were produced in other areas. To overcome this problem, migration was carried out at a range of velocities and it was carefully noted which reflections were caused by diffraction tails of point-shaped reflectors. These reflections were discarded for structural interpretation. The mean energy decay with depth was calculated by REFLEXW and an inverse, energy decay-dependent amplification was used for displaying the moraine profiles. Automatic gain control (AGC) was additionally used to verify the continuity of reflectors but is not shown in the profiles. Prominent reflectors were derived from the 50 MHz profiles and small-scale reflections (thin lines) from the 200 MHz profiles. Topography was obtained by measuring slope segment length and dip angle with a tape measure and compass clinometer during GPR measurements.

Study design We measured three profiles in 50, 100 and 200 MHz across the 1920 and 1980 lateral moraines (Fig. 1). Profile MOR1 crosses over a bedrock outcrop, a field of annual moraines, the 1980 and 1920 lateral moraines and finally up the steep slope of the 1850 lateral moraine. Profiles MOR2 and MOR3 only extended over the 1980 and 1920 ridges. The overall slope of these profiles is related to the position at the valley side. The 200 MHz profiles was chosen for depicting small-scale sedimentary and/ or deformation features in the sediments; 50 and 100 MHz profiles were measured to obtain supplementary information from the deeper subsurface. The step size was 0.2 m for the 100 and 200 MHz profiles and 0.5 m for the 50 MHz profiles. Additionally, five 200 MHz profiles (step size 0.2 m) were measured immediately at the ice margin (as of June 2007); each of these profiles started on the thin blanket of glaciogenic sediments some tens of metres from the glacier, ran perpendicular to the

glacier margin and extended some tens of metres onto the glacier surface. As the radargrams are widely similar, only one example is presented in the results section. Four of the five CMP profiles were located on profile line MOR1 (see Fig. 1) and the fifth was on the surface of Gornergletscher.

Results Lateral moraine sedimentology One large trench measuring 6.0 1.8 1.5 m (width, height, depth) was dug by hand perpendicular to the crestline of the 1980 moraine that runs SW –NE (2288 –488) in this location (Fig. 3). This moraine, like the other prominent lateral moraines in this catchment, is slightly asymmetrical with a steeper distal slope facing NW (left, c. 30 –358) and a slightly gentler proximal SE-facing slope (right, c. 25 –308). In some places, considerable microrelief causes a more complex moraine surface morphology and removes the overall asymmetry. Sediment characteristics allow two identical lithofacies associations to be distinguished in exposures on either side of the trench. The first (A in Fig. 3) consists of units of matrix-supported stratified diamictons (Dms) that reach thicknesses of up to 0.5 m; this facies is relatively easy to excavate, dips parallel to the surfaces on either side of the moraine crestline and dominates the exposure. Individual layers of this unit are often separated by laminated or massive fine-to-medium sand layers or lenses that commonly reach thicknesses of 0.01 –0.05 m. Few outsized granule to gravel-sized particles occur throughout these sorted sediment layers. Boulders reach a maximum a-axis length of 0.7 m and are found within the exposure and scattered on the moraine surface. Clusters of interlocking boulders can be found throughout the diamicton; usually these are subangular gneisses and mica-schists that are bullet-shaped and/or display faceting. The a-b-planes of platy boulders are orientated parallel to the moraine surface. A diamicton with similar characteristics (Dms), but an abundance of sorted openwork granule layers instead of finer-grained units (B in Fig. 3), overlies this unit dipping towards the SSE (right in Fig. 3). The second lithofacies association that occurred primarily in the upper half of the exposure consists of somewhat thicker and more continuous relatively well-sorted sediment units, including crude horizontally bedded and massive granule lenses (GRh and GRm, respectively) and laminated or massive fine-to-medium sand layers or lenses (Fl, Fm/Sm, Sm, Sl). They are commonly between 0.01 and 0.1 m thick (C in Fig. 3) but can occasionally reach thicknesses of 0.05– 0.1 m (D in Fig. 3). These layers dip parallel to the moraine surface

SEDIMENTOLOGY AND GPR OF LATERAL MORAINES

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Surface

2m

Clasts

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Sand or fine-grained sediment lenses/units

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/S

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Dm

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Fig. 3. Sedimentary log of section GOR4 through the crestline of the 1980 lateral moraine (Figs 1b & 2). A detailed explanation of the lithofacies codes and lettering is given in the text.

and thinner sorted sediment layers within the diamict units. One such layer bends round underneath the crestline.

Ground-penetrating radar CMPs and wave propagation velocity. The mean propagation speed in the 1920 and 1980 moraines was c. 0.08 m ns21 with the velocity being slightly higher on steep slopes (0.09 m ns21) due to drier conditions and slightly lower in visibly wetter places (0.06–0.07 m ns21). On bedrock, a value of 0.11 m ns21 was determined (Fig. 4). The CMP analyses show clearly that velocity changes with depth; in all moraine CMPs a distinct velocity increase to typical bedrock values is visible at a few metres’ depth (Fig. 4). The position of this velocity rise

coincides with prominent reflectors which are interpreted as bedrock surface in profile MOR1 (see below). At CMP4, a shallow zone of very low propagation velocity (,0.06 m ns21) is found directly above the supposed bedrock which is probably due to a heightened water content above the damming bedrock. The ice of the Gornergletscher stands out because of a much higher propagation velocity of 0.17 m ns21. As the glaciogenic sediments are the main target of the investigation, a value of 0.09 m ns21 was used for the runtime-depth conversion of all moraine profiles. Moraine profiles. In Figure 5, the 50 and 200 MHz radargrams of profile MOR1 are shown. The 50 MHz picture highlights the most prominent features while the 200 MHz picture is much more

Fig. 4. Profiles of vertical wave propagation velocity at CMP 1 –5 as derived from semblance analysis using REFLEXW.

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Fig. 5. Cross-section of the moraine complex derived from GPR measurements (profile line MOR1): (a) 50 MHz radargram; (b) exemplary section of the 200 MHz radargram (full profile not reproducible at the given size); and (c) combined interpretation. Thick lines: prominent reflections (50 MHz); thinner lines: 200 MHz reflections; grey dotted line: bedrock surface.

detailed. However, some structures which are very clear in 50 MHz are less obvious in 200 MHz. The intermediate 100 MHz profiles delivered no additional information and are thus not presented. Bedrock depth was interpreted mainly using the 50 MHz radargram in the combined interpretation sketch (Fig. 5c). The sediment thickness is mostly between 1 –5 m as indicated by the sharp reflections on the 50 MHz profile; data from CMPs 2 and 4 support the interpretation. In the centre of the section under the 1980 ridge no unequivocal bedrock reflection was found. CMP 3 indicates a sediment thickness of c. 3.5 m which is in agreement with the fragmentary reflections found in the vicinity. Along the profile, a bedrock outcrop occurs at the base (0–6 m distance) and a strong reflection marks its subsurface position at c. 10–30 m.

Between 30 and 60 m, a strong reflector extends diagonally downwards to c. 15 m depth. However, only the topmost part of this zone shows additional reflections which are probably associated with sedimentary deposits. We therefore attribute this reflector to an internal bedrock structure; the sediment thickness is probably rather low (c. 2 m, supported by CMP 2). Mostly fragmentary reflections were found in the shallow subsurface; only a few continuous reflectors could be traced. Particularly strong scattering was observed in the radargram area directly under the crestline of the 1980 moraine, which is probably due to topography effects in the vicinity of the steep ridge. Reflections associable to sedimentary structures are mainly found in the shallow depression near the north-western starting point of

SEDIMENTOLOGY AND GPR OF LATERAL MORAINES

the profile and underneath the 1980 ridge. At a profile distance between 5 and 20 m the reflections are fragmentary and surface-parallel, probably indicating glaciofluvial deposits. Under a small ridge at 22 m, small-scale reflectors are dipping to the right (SE), that is, away from the former ice margin. No clear reflections are visible underneath the annual moraines except for some subparallel to the surface. Under the 1980 ridge, information from the 50 and 200 MHz radargrams differs. In the surficial area of main interest, wavy, subhorizontal, interrupted reflections are found. Further upslope, the reflection patterns are unclear; very little information can be derived from the near-surface parts of the radargrams. At the locality MOR2 (Fig. 6), the position of the bedrock surface cannot be unequivocally traced; without the aid of CMP profiles there are at least

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two possible interpretations. The near-surface reflections are surface-parallel, and this holds particularly true for the ice-proximal SE slopes. On icedistal slopes, small-scale fragmentary reflection patterns are found (such as between 45 and 48 m distance) but the most prominent features dip away from the ice to the right (NW) side. At c. 36 m profile distance, the interface between the two moraine bodies dips diagonally down to the left. The migration process reveals that several hyperbolic reflectors can be traced back to point-shaped subsurface features (black dots in Fig. 6c). These are most likely boulders in the sediment layer, but can also be caused by protrusions of the bedrock surface or by intersecting joints within the bedrock. At MOR3 (Fig. 7), a subhorizontal reflector recognizable in both the 200 and 50 MHz radargrams is interpreted as the bedrock surface. Depth

Fig. 6. Cross-section of the moraine complex derived from GPR measurements (profile line MOR2): (a) 50 MHz radargram; (b) 200 MHz radargram; and (c) combined interpretation. Thick lines: prominent reflections (50 MHz); thinner lines: 200 MHz reflections; grey dotted line: bedrock surface; black dots: point-shaped reflectors.

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Fig. 7. Cross-section of the moraine complex derived from GPR measurements (profile line MOR3): (a) 50 MHz radargram; (b) 200 MHz radargram; and (c) combined interpretation. Thick lines: prominent reflections (50 MHz); thinner lines: 200 MHz reflections; grey dotted line: bedrock surface; black dots: point-shaped reflectors. Black arrow highlights the arched reflector discussed in the text.

to bedrock is 2–8 m in the left half of the profile. In the right half, the most probable position of the bedrock surface is supposedly at 5– 7 m as indicated in the figure. Both sides of the ridges show surfaceparallel reflections. The 50 MHz radargram displays a pronounced arched structure, onlapping the surface-parallel structure of the 1920 ridge. In the 200 MHz radargram, this structure cannot be recognized. Ice-marginal profile IM1. The profile lies in a wide depression filled with loose, moist glaciogenic

sediments. They consist of low ridges oriented parallel to the ice margin; from the geomorphology, these are interpreted as annual recessional moraines. The glacier base can be clearly recognized in the radargram (Fig. 8); the strong reflector causes an apparent second ‘multiple’ reflection at greater depth. There is a multitude of small-scale reflections in a c. 5 m thick layer in the foreland (0–30 m distance). These are fragmentary, mostly surface-parallel and probably correspond to the proglacial sediment cover. Many small-scale discordant structures can be observed and single reflectors

SEDIMENTOLOGY AND GPR OF LATERAL MORAINES

are noticeably wavy. The same type of reflections continues underneath the glacier snout. At the base of the mentioned reflections in the foreland, a rather weak reflector and the generally decreasing reflectivity might evidence the interface to bedrock. However, under the glacier the underlying bedrock cannot be depicted due to amplified data noise. Several pronounced reflections can be recognized within the ice body (dashed lines).

Interpretation Lateral moraine sedimentology The relatively low consolidation of the matrixsupported, stratified diamict units, their inclination subparallel to the ground surface and their gentle dip suggest that these units were deposited in a subaerial setting by cohesive debris flows to form a terrestrial ice-contact fan during the 1980 advance of Gornergletscher. Local clast concentrations within the diamicts might therefore represent either a concentration of heavier components at the base of flows during transport (basal traction gravels; cf. Lawson 1988; Benn 1992) or the sliding and interlocking of clasts downslope at the surface of thin debris flows (clast clusters; Lukas 2005; Benn & Lukas 2006). Indeed, the latter effect has commonly been observed as presenting obstacles that halted progress during excavation. The second lithofacies association, but also thinner sorted sediment units within the diamict, are interpreted as fluvial sediments deposited during thin sheetflows or in shallow rills, based on the presence of laminations

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and their laterally discontinuous arrangement (Lawson 1988). The alternating deposition of both lithofacies units can be explained by a combination of two factors: a stationary ice margin and differences in water content. The latter determines whether gravitational or fluvial processes dominate in fan formation at any given time (Krzyszkowski & Zielinski 2002). Thin layers and lenses of sorted sediments would most likely have resulted from overland-flow events during which sheetflows or water in shallow rills would have eroded, transported and deposited only finer fractions (Lawson 1982, 1988; Benn 1992; Krzyszkowski & Zielinski 2002; Lukas 2005; Reinardy & Lukas 2009). The presence of a second diamicton near the surface (B in Fig. 3) can be interpreted as evidence of localized reworking of the underlying debris flows either during rainfall events or after the glacier had retreated from this position and withdrawn its support from the proximal face. In general, the subparallel and gentle dip of units underneath the proximal side of the moraine indicates that retreat was gradual and that the proximal units were reworked successively by debris flows towards the glacier as the ice margin lowered. The proximal sides of this moraine thus represent an icecontact face where material is at the angle of repose. The steep dip of the distal (left) slope indicates that some post-depositional steepening of these units might have occurred as depositional slopes of such fans are usually somewhat less steep (Kjær et al. 2004; Lukas 2005). However, in the absence of thick fine-grained units, it is difficult to judge the degree of alteration by ice push. It is worth noting

Fig. 8. GPR profile Ice Margin IM1, 200 MHz. (a) Migrated radargram (2D-migration using a subsurface velocity of 0.16 m ns21 for glacier ice and 0.08 m ns21 for sediments and (b) interpretation. Black lines: GPR reflections; grey dotted line: supposed bedrock surface.

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that natural exposures occurring along the flanks of lateral moraines in the present study area, but also elsewhere in the Alps, display highly similar facies and sedimentary structures (Humlum 1978; Small 1983, 1987; Benn et al. 2003; Graf 2007). However, despite the similarity of the internal architecture of these lateral moraines and the surface-parallel dip of all internal units with published examples, the surface asymmetry appears to be reversed. In most published cases and those known to the authors, the steeper face usually represents the ice-proximal face while the ice-distal face is the gentler fan surface (Reinardy & Lukas 2009). However, it should also be noted that similarly steep distal slopes have been reported in Alpine settings (Humlum 1978; Ro¨thlisberger & Schneebeli 1979; Graf 2007; Curry et al. 2009). In those cases, however, the proximal slopes are considerably steeper and may not reflect a typical asymmetrical fan cross-section where the proximal slope is material at the angle of repose. Instead, it may represent glaciotectonized pre-existing sediment that has been overconsolidated during repeated glacier advances (Graf 2007; Lukas et al. unpublished data). In addition, the height of the three prominent moraines (AD 1850, 1920, 1980 advance) is as much as an order of magnitude less than for other glaciers in the Alps where moraines reach heights of up to 100 m. Therefore, the processes reconstructed from sedimentology and inferred from the surface morphology are apparently at odds with each other and warrant further examination. To place this in a wider context, four key potential controls will be discussed below. Firstly, the valley walls of the catchment under study are steep, with bedrock at or close to the surface across much of the catchment (Fig. 2). Secondly, the glacier that formed the moraines investigated here is a subsidiary tongue to the much larger Gornergletscher (Fig. 1) and has covered a much smaller area than the deeply excavated valleys occupied by much thicker Alpine glaciers elsewhere (Huss 2005; Farinotti et al. 2009). Thirdly, and possibly partly as a result of the above, this subsidiary tongue fluctuates much more dynamically than Gornergletscher itself and most other glaciers in the Swiss Alps. This can be inferred from the presence of annual moraines, which are very unusual landforms in a high-mountain context because of a more limited mass turnover (Benn et al. 2003) compared to temperate lowland settings where annual moraines form more commonly (Sharp 1984; Evans 2003; Bradwell 2004). Lastly, material supply to the snout (most notably the surface) of Gornergletscher appears to be much less compared to other glaciers in a similar lithological setting (Lukas et al. in press). This can be inferred from the patchy distribution of only a thin veneer of

sediment in the proglacial area and sparser emergence of debris-rich englacial septa which are much more frequent on many other Alpine glaciers (Graf 2007; Lukas et al. in press). We argue that limited material supply, combined with a shorter snout response time and the fact that deposition occurred against a steep bedrock slope, may explain the unusual shape and apparently reversed cross profile of lateral moraines at Gornergletscher. Limited material is not sufficient to form prominent moraines if the margin is not stationary for a considerable amount of time or is unable to push up large amounts of pre-existing material (Benn & Evans 1998; Benn et al. 2003). In addition, the steep bedrock slopes make sliding and continued reworking of material on the proximal slopes more likely after glacier retreat, resulting in a proximal slope that is slightly less steep than the norm. In addition, several Alpine glaciers continue to consolidate their proximal lateral moraine slopes by lateral drag, especially when such moraines hem in and channel glacier flow hence producing extremely steep proximal faces that can reach angles of 608 and more (Ro¨thlisberger & Schneebeli 1979; Graf 2007; Curry et al. 2009). Based on test pits in other parts of this and other lateral moraines in the catchment, we regard this one section as representative of the internal architecture. Furthermore, several lateral moraines in the Alps consist of similar or identical lithofacies associations (Lukas et al. unpublished data); we are therefore confident in the interpretation given above.

Ground-penetrating radar The main finding from profiles MOR1 –3 (Figs 5 –7) is that the sediment thickness ranges from 1–8 m. Depth to bedrock and sedimentary structures are partly unclear under the steep ridges. However, surface-parallel stratification is evident in all of the profiles. No clearly discernible reflectors are found underneath the annual moraines (which may be a problem of resolution) and there is no well-defined interface between diamict layers of different age. The relatively narrow range of sediment and bedrock velocities as derived from the CMPs means that there is not necessarily a distinct dielectrical contrast at the interface between two sedimentary units or at the sediment/bedrock interface, particularly when the sediment is dry. This problem was encountered in previous investigations on talus debris (Sass 2007) and explains the lack of a clearly defined bedrock reflection in some places (MOR2, IM1) while the reflection is more pronounced when local stagnant water is involved (MOR1 and surrounding of CMP4). The data noise in the vicinity of the steep ridges poses

SEDIMENTOLOGY AND GPR OF LATERAL MORAINES

additional problems for interpreting the bedrock surface from the radargrams. Regardless, the sediment thickness can be narrowed down at MOR1 to 4–5 m in the small depression, mostly less than 3 m under the annual moraines and under the slopes, probably less than 4 m under the 1980 ridge and 7 m or more under the uppermost part of the 1850 moraine slope. At MOR2 and MOR3 the results are less clear but they point to a similar range of thicknesses. The irregular reflection patterns of the shallow subsurface probably reflect the small-scale structures of the poorly consolidated glaciogenic sediments. There does not seem to be a well-defined interface between individual sedimentary facies making up moraines of different age. This observation may be due to very similar grain size distributions or to reworking of older deposits by subsequent glacier advances. Only at MOR3 an arched, onlapping structure gives evidence of a third, shallow ridge between the two main moraine ridges. The interface between the two moraine bodies dips diagonally down to the proximal side. This might indicate deposition by lateral-frontal stacking of debris flows against the older lateral moraine. During this stage, the swale between the outer, older moraine and that in formation could have been occupied by meltwater channels or small ponds (Benn & Owen 2002; Benn et al. 2003); subsequent incision could have led to the formation of the local depressions found in most radargrams. At all moraine profiles, both the proximal and the distal sides of the ridges show surfaceparallel stratification. This finding confirms the interpretations derived from the sediment exposure and extends them to a greater depth. At the ice margin site (IM1), the wavy, subhorizontal reflectors immediately under the ice are not typical of bedrock and clearly point to a sedimentary base of the glacier. Single reflectors within the forefield sediments are noticeably wavy which may indicate proglacial thrusting and folding or basal freeze-on of sediment slabs during seasonally driven glacier advances. The clearly recognizable reflections within the ice body are interpreted to be englacial shear planes or debris-rich layers.

Discussion and implications Suitability of GPR Despite a few methodological complications, which we highlight below, our data show that GPR is well suited to extend the information gained from shallow surface exposures into the deeper subsurface. However, even the vertical resolution of the 200 MHz antennas (roughly 10– 15 cm in the

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given environment) is not sufficient to detect thin beds and small-scale sedimentary structures. The radargrams therefore cannot resolve, or replace, the detail of field sedimentological investigations. Furthermore, the direct air and ground waves of the GPR blank out the uppermost 0.5– 0.8 m of subsurface information. A complete cross-validation between exposure and GPR data will therefore hardly ever be possible. The GPR however provided valuable information in establishing the overall internal structure and thickness of the glacial deposits, extending to the subsurface outcrop information, generally to the upper 5–10 m in these systems. This is invaluable in settings such as the present where natural exposures are rare and digging progress is slow due to the diamictic nature of the sediments. Applying two or three working frequencies proved to be rewarding in characterizing the main structures of interest. While 200 MHz profiles rendered it possible to depict internal sedimentary structures, the 50 MHz radargrams provided valuable additional information on the importance and continuity of reflections, making it easier to estimate the position of the bedrock surface. The vertical velocity profiles from the CMPs proved to be extremely valuable to underpin the interpretation of depth to bedrock. Irregular reflections are found in the radargrams of all frequencies, irrespective of the filtering strategies applied. Thus, the radargrams are thought to give a realistic picture of the general internal structures of the sediments of the moraines. However, it is evident that the radargrams reveal less clear information on the internal structure of the moraine complexes than earlier investigations (Bakker & van der Meer 2003; Helle 2004; Sadura et al. 2006). This result can be understood when considering the following points. Firstly, the studied moraines are one order of magnitude smaller than those previously investigated in the Netherlands, Norway and Canada; some of the internal structures that are seen in the exposures are therefore beyond the resolution of the 200 MHz antennas used. Secondly, the sediments are diamictic with frequent outsized clasts in a poorly-sorted matrix; it is therefore unlikely that an even higher antenna frequency would obtain better results. Thirdly, due to the production of rock flour by glacial abrasion, fine silt is abundant which leads to a strong attenuation of the GPR signal (Schwamborn et al. 2008). Burki et al. (2009) also encountered the problem of relatively weak and diffuse internal reflections and an unclear moraine base. Regardless of the actual thickness of the deposits, penetration depth is restricted to around 10 m. Despite these limitations, our approach worked reasonably well in boulder-rich diamictic sediments

S. LUKAS & O. SASS

Debris-flow dominated ice-contact fan

Glacier

dr oc k Be

(b) Initial phase of glacier retreat Collapse of proximal sediments: ice-contact face formation

Possible incorporation of dead ice

Be dr oc k

Glacier

(c) Complete deglaciation and moraine stabilisation

Be

dr

oc k

Ice-contact face formation complete

(d) Renewed advance and formation of inset lateral moraine Debris-flow dominated ice-contact fan

Onlap and possible meltwater channel formation

Glacier

oc k

The resolution of the GPR is not high enough to determine the presence of thin layers of englacial sediment. However, subglacial sediment thicknesses of up to 5 m can be carefully estimated for the subglacial realm, and this is visible on the radargrams (Fig. 8). Field observations higher upglacier and from aerial photographs show that a small number of englacial debris bands bring sediment to the glacier surface; this is a characteristic feature of many Alpine glaciers (Goodsell et al. 2005; Lukas et al. in press) and might be the most likely explanation of a steady supply of material that can be mobilized by water to form debris flows and thus the sediments observed in section. Medial moraines and marginal accumulations of supraglacial debris might form an additional source of material; however, the latter material would be expected to be dominantly angular, and this angularity has not been observed in exposures. It is worth noting that dead ice appears to be present in the immediate proglacial surroundings of the snout, and this is consistent with field observations and test diggings in annual moraines where sedimentary structures indicate collapse following dead-ice meltout (Lukas in press). However, no such evidence was found in the sediments exposed in the lateral moraine investigated here (Fig. 3) and geomorphological features associated with dead-ice meltout (Kjær & Kru¨ger 2001; Lukas et al. 2005) are absent. We present our conceptual model of lateral moraine formation based on our data below and in Figure 9. Both the proximal and distal slopes of the moraine show surface-parallel strata with a slight asymmetry reflected in a slightly steeper distal slope. This is probably the result of the stacking of supraglacially derived debris flows against a temporarily stationary ice margin (Fig. 9a). Although such debris flows have been reported to accumulate as a fan or ramp at a low depositional angle on distal slopes of moraines (Benn 1992; Lukas 2005; Reinardy & Lukas 2009), these depositional angles may be much steeper where cohesive debris flows are stacked up in high-mountain contexts,

Meltout of englacial sediment (originating from crevasse traces, shear planes)

dr

Glacial debris release and lateral moraine formation at Gornergletscher

(a) Initial phase of lateral moraine formation

Be

commonly encountered in glaciated (high) mountain environments. Together with the initial work of Lønne & Lauritsen (1996) in high Arctic lowland settings, this contradicts earlier suggestions by Beres & Haeni (1991) that boulder top reflectors might disguise sedimentary boundaries in diamictic sediments. Thus, our data encourage extending the use of GPR to diamictic sediment in such environments.

En gl a (s cia he l d ar eb pl ris an s es ep ?) ta

88

Overriding and possible glaciotectonisation

Fig. 9. Schematic diagram summarizing the key information on genetic processes of the lateral moraines at Gornergletscher as deduced from moraine sedimentology and GPR. All panels are vertically exaggerated and not to scale. For a full explanation see text.

frequently forming very large lateral moraines (Benn et al. 2003; Graf 2007; Curry et al. 2009). This steeper dip of units underlying the distal slope might also reflect post-depositional steepening by proglacial or ice-marginal push. However, the

SEDIMENTOLOGY AND GPR OF LATERAL MORAINES

inability of diamict units to accurately record deformation internally makes it difficult to analyse the degree and style of deformation at the macro-scale (Hart & Boulton 1991; Lukas 2005; Reinardy & Lukas 2009). Sadura et al. (2006) reported similar results: they found surface-parallel layered deposits, the proximal parts of which had been somewhat oversteepened by ice push. Larger-scale deformation structures such as thrusts or overturned folds, as observed elsewhere (Lønne & Lauritsen 1996; Bakker & van der Meer 2003), appear to be absent in the present study area. Although the evidence reported above can only be considered preliminary, studies conducted on large lateral moraines elsewhere in the Alps and other high-mountain environments show that large-scale brittle deformation is usually not widespread (Humlum 1978; Ro¨thlisberger & Schneebeli 1979; Small 1983; Graf 2007). As the glacier retreated and withdrew its support the proximal face would have collapsed, leading to the establishment of a new equilibrium by creating a steeper ice-contact face where material is at the angle of repose (Fig. 9b, c). At this stage, depending on the thickness of sediment cover and the speed with which the retreating glacier thins, dead ice may (theoretically) be incorporated into the base of a lateral moraine and meltout may trigger further readjustment of the proximal slope after glacier retreat (Lukas in press). As retreat is a gradual process, the proximal slope is often also stratified (surface-parallel stratification) but there is a relative paucity of fine-grained sediment lenses due to a lack of sheetflow processes (Reinardy & Lukas 2009; Fig. 3). Renewed oscillations of the glacier during subsequent readvances may add new material (Fig. 9d), and meltwater may collect in the depression between the lateral moraine upslope and that currently being formed. This may explain some of the onlapping younger units onto older ones which have been observed in some radargrams (Figs 7 & 9d). This also suggests that (re-)advances of the glacier onto older moraines did not significantly erode previously deposited sediment to a large extent, although some of this evidence may be covered by subsequent post-depositional slumping. In general, it can be said that the moraines form a clear morphostratigraphical succession with no evidence of cross-cutting or overriding/overtopping of previous glacier maxima that can be found elsewhere in the Alps (e.g. Ro¨thlisberger & Schneebeli 1979; Small 1983). The combination of sedimentological analyses and GPR has therefore enabled the modes of formation of a somewhat unusual lateral moraine sequence at Gornergletscher to be established. This adds to a growing body of literature seeking

89

to elucidate the formation and significance of moraines in an Alpine context.

Conclusions Four main conclusions can be made from this study. (1) Sedimentological analyses of an exposure through the lateral moraine of Gornergletscher, formed during the AD 1980 advance, have enabled the mode of moraine formation to be determined. The data show that the moraine is a result of the stacking of supraglacially-derived debris flows and associated waterlain deposits that form nearsymmetrical lateral moraines with slightly steeper proximal slopes. The cross-profile of the moraine investigated is somewhat atypical in a high-mountain context and probably a result of a combination of limited debris supply, extremely short snout response times and steep bounding bedrock walls. (2) Ground-penetrating radar (GPR) has been used along transects across bedrock outcrops, fields of small annual moraines and a succession of lateral moraines relating to the AD 1850, 1920 and 1980 advances. Radar patterns for the deeper subsurface provide a meaningful insight into the continuation of the overall geometries seen in the exposure and can be used to extend the observations to greater depths along the chosen profile lines. Common-midpoint velocity profiles are vital for a sound interpretation of depth to bedrock. The radar and sedimentological facies corroborate each other very well in the shallow subsurface where ground control exists. (3) GPR investigations using 50 and 200 MHz antennas proved to be a promising approach for assessing field sedimentology in boulderrich, diamictic sediments; we are not aware of this having been used in Alpine glacial environments previously. Reflectors reconstructed from the radargrams are interpreted to represent a realistic image of unit boundaries and larger scale dip and structural trends. Thus, our data encourage extending the use of GPR in such environments. However, even the vertical resolution of the 200 MHz antennas (roughly 10 –15 cm in the given environment) is not sufficient to detect thin beds and small-scale sedimentary structures; the radargrams therefore cannot resolve or replace the detail of field sedimentological investigations. (4) A genetic model based on sedimentological and GPR analyses strongly suggests that the

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AD 1980 lateral moraine of Gornergletscher formed through ice-marginal stacking of debris flows, followed by collapse of proximal units associated with glacier retreat. Subsequent advances of lesser extent appear to have formed new lateral moraines at lower altitudes by lateral accretion and without any visual signs of deformation of alreadyexisting moraines. This genetic model may have implications for the interpretation of lateral moraines in high-mountain environments elsewhere and encourages further research into these fascinating landforms. We would like to thank M. Funk, M. Huss, F. Walter and M. Werder (ETH Zu¨rich) for the much appreciated hospitality at the ETH’s summer camp among Gornergletscher’s moraines. M. Werder, F. Weber and N. Borsutzki are thanked for company in the field and help with enlarging the pit through the 1980 moraine. Part of the work was undertaken while SL was in receipt of an EU Marie Curie Intra-European Fellowship (023494-PACES) which is hereby gratefully acknowledged. Comments by A. Schomacker, B. Mahaney, I. P. Martini and two anonymous reviewers helped considerably to improve the message of this paper.

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Beres, M., Jr. & Haeni, F. P. 1991. Application of ground-penetrating-radar methods in hydrogeological studies. Ground Water, 29, 375– 386. Bersezio, R., Giudici, M. & Mele, M. 2007. Combining sedimentological and geophysical data for highresolution 3-D mapping of fluvial architecture elements in the Quaternary Po plain (Italy). Sedimentary Geology, 202, 230 –248. Berthling, I., Etzelmu¨ller, B. & Isaksen, K. 2000. Rock glaciers on Prins Karls Forland. II: GPR soundings and the development of internal structures. Permafrost and Periglacial Processes, 11, 357–369. Bradwell, T. 2004. Annual moraines and summer temperatures at Lambatungnajo¨kull, Iceland. Arctic, Antarctic and Alpine Research, 36, 502– 508. Bristow, C. S. & Jol, H. M. 2003. GPR in sediments: advice on data collection, basic processing and interpretation, a good practice guide. In: Bristow, C. S. & Jol, H. M. (eds) Ground Penetrating Radar in Sediments. Geological Society, London, Special Publications, 211, 9–27. Burki, V., Larsen, E., Fredin, O. & Margreth, A. 2009. The formation of sawtooth moraine ridges in Bødalen, western Norway. Geomorphology, 105, 182–192. Busby, J. P. & Merritt, J. W. 1999. Quaternary deformation mapping with ground penetrating radar. Journal of Applied Geophysics, 41, 75–91. Curry, A. M., Sands, T. B. & Porter, P. R. 2009. Geotechnical controls on a steep lateral moraine undergoing paraglacial slope adjustment. In: Knight, J. & Harrison, S. (eds) Periglacial and Paraglacial Processes and Environments. Geological Society, London, Special Publications, 320, 181– 197. Degenhardt, J. J., Jr, Giardino, J. R. & Junck, M. B. 2003. GPR survey of a lobate rock glacier in Yankee Boy Basin, Colorado, USA. In: Bristow, C. S. & Jol, H. M. (eds) Ground Penetrating Radar in Sediments. Geological Society, London, Special Publications, 211, 167–180. Ebert, A. 2001. Strukturgeologie und Petrographie im Gebiet von Zermatt an der Dent-Blanche-Decke, in der Combin-Zone, in der Zone von Zermatt-Saas Fee und in der Monte Rosa-Decke. Unpublished MSc thesis, Universita¨t Stuttgart, Germany. Evans, D. J. A 2003. Ice-marginal terrestrial landsystems: active temperate margins. In: Evans, D. J. A (ed.) Glacial Landsystems. Arnold, London, 12– 43. Evans, D. J. A. & Benn, D. I 2004. Facies description and the logging of sedimentary exposures. In: Evans, D. J. A & Benn, D. I (eds) A Practical Guide to the Study of Glacial Sediments. Arnold, London, 11– 51. Eyles, N., Eyles, C. H. & Miall, A. D. 1983. Lithofacies types and vertical profile models. An alternative approach to the description and environmental interpretation of glacial diamict and diamictite sequences. Sedimentology, 30, 393–410. Farinotti, D., Huss, M., Bauder, A., Funk, M. & Truffer, M. 2009. A method to estimate the ice volume and ice-thickness distribution of alpine glaciers. Journal of Glaciology, 55, 422– 430. Gerber, R., Salat, C., Junge, A. & Felix-Henningsen, P. 2007. GPR-based detection of Pleistocene

SEDIMENTOLOGY AND GPR OF LATERAL MORAINES periglacial slope deposits at a shallow-depth test site. Geoderma, 139, 346– 356. Goodsell, B., Hambrey, M. J. & Glasser, N. F. 2005. Debris transport in a temperate valley glacier: Haut Glacier d’Arolla, Valais, Switzerland. Journal of Glaciology, 51, 139–146. Graf, A. 2007. Genesis of alpine lateral moraines – an example from Findelengletscher near Zermatt (VS), Switzerland (in German). Unpublished MSc thesis, University of Bern. Hart, J. K. & Boulton, G. S. 1991. The interrelation of glaciotectonic and glaciodepositional processes within the glacial environment. Quaternary Science Reviews, 10, 335–350. Heinz, J. & Aigner, T. 2003. Three-dimensional GPR analysis of various Quaternary gravel-bed braided river deposits (southwestern Germany). In: Bristow, C. S. & Jol, H. M. (eds) Ground Penetrating Radar in Sediments. Geological Society, London, Special Publications, 211, 99– 110. Helle, S. K. 2004. Sequence stratigraphy in a marine moraine at the head of Hardangerfjorden, western Norway: evidence for a high-frequency relative sea-level cycle. Sedimentary Geology, 164, 251–281. Humlum, O. 1978. Genesis of layered lateral moraines: implications for palaeoclimatology and lichenometry. Geografisk Tidsskrift, 77, 65–72. Huss, M. 2005. Gornergletscher: Gletscherseeausbru¨che und Massenbilanzabscha¨tzungen. Unpublished MSc thesis, ETH Zu¨rich. Huss, M., Bauder, A., Werder, M., Funk, M. & Hock, R. 2007. Glacier-dammed lake outburst events of Gornersee, Switzerland. Journal of Glaciology, 53, 189–200. Jacob, R. W. & Hermance, J. F. 2004. Assessing the precision of GPR velocity and vertical two-way traveltime estimates. Journal of Environmental and Engineering Geophysics, 9, 143– 153. Jakobsen, P. R. & Overgaard, T. 2002. Georadar facies and glaciotectonic structures in ice-marginal deposits, northwest Zealand, Denmark. Quaternary Science Reviews, 21, 917–927. Kjær, K. H. & Kru¨ger, J. 2001. The final phase of dead-ice development: processes and sediment architecture, Ko¨tlujo¨kull, Iceland. Sedimentology, 48, 935–952. Kjær, K. H., Sultan, L., Kru¨ger, J. & Schomacker, A. 2004. Architecture and sedimentation of outwash fans in front of the My´rdalsjo¨kull ice cap, Iceland. Sedimentary Geology, 172, 139– 163. Krzyszkowski, D. & Zielinski, T. 2002. The Pleistocene end moraine fans: controls on their sedimentation and location. Sedimentary Geology, 149, 73–92. Lawson, D. E. 1982. Mobilization, movement and deposition of active subaerial sediment flows, Matanuska Glacier, Alaska. Journal of Geology, 90, 279–300. Lawson, D. E. 1988. Glacigenic resedimentation: Classification concepts and application to massmovement processes and deposits. In: Goldthwait, R. P. & Matsch, C. L. (eds) Genetic Classification of Glacigenic Deposits. Balkema, Rotterdam, 147–169.

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Leopold, M. & Vo¨lkel, J. 2003. GPR images of periglacial slope deposits beneath peat bogs in the Central European Highlands, Germany. In: Bristow, C. S. & Jol, H. M. (eds) Ground Penetrating Radar in Sediments. Geological Society, London, Special Publications, 211, 181 –190. Lønne, I. & Lauritsen, T. 1996. The architecture of a modern push-moraine at Svalbard as inferred from ground-penetrating radar measurements. Arctic and Alpine Research, 28, 488– 495. Lukas, S. 2005. A test of the englacial thrusting hypothesis of ‘hummocky’ moraine formation – case studies from the north-west Highlands, Scotland. Boreas, 34, 287– 307. Lukas, S. In press. Ice-cored moraines. In: Singh, V., Singh, P. & Haritashya, U. K. (eds) Encyclopedia of Snow, Ice and Glaciers. Springer, Heidelberg. Lukas, S., Nicholson, L. I., Ross, F. H. & Humlum, O. 2005. Formation, meltout processes and landscape alteration of high-arctic ice-cored moraines – examples from Nordenskio¨ld Land, central Spitsbergen. Polar Geography, 29, 157–187. Lukas, S., Coray, S., Graf, A. & Schlu¨chter, C. In press. The influence of clast lithology and fluvial reworking on the reliability of clast shape measurements in glacial environments – a case study from a temperate Alpine glacier. In: Bridgland, D. R. (ed.) Clast Lithological Analysis. Technical Guide. Quaternary Research Association, London. Maisch, M., Wipf, A., Denneler, B., Battaglia, J. & Benz, C. 2000. Die Gletscher der Schweizer Alpen: Gletscherhochstand 1850, aktuelle Vergletscherung, Gletscherschwund-Szenarien. Hochschulverlag ETH, Zu¨rich. Moorman, B., Robinson, S. & Burgess, M. 2003. Imaging periglacial conditions with groundpenetrating radar. Permafrost and Periglacial Processes, 14, 319–329. Neal, A. 2004. Ground-penetrating radar and its use in sedimentology: principles, problems and progress. Earth-Science Reviews, 66, 261–330. Reinardy, B. T. I. & Lukas, S. 2009. A comparison of the sedimentary signature of ice-contact sedimentation and deformation at macro- and micro-scale: a case study from NW Scotland. Sedimentary Geology, 221, 87–98. Ro¨thlisberger, F. & Schneebeli, W. 1979. Genesis of lateral moraine complexes, demonstrated by fossil soils and trunks: indicators of postglacial climatic fluctuations. In: Schlu¨chter, C. (ed.) Moraines and Varves. Balkema, Rotterdam, 387– 419. Sadura, S., Martini, I. P., Endres, A. L. & Wolf, K. 2006. Morphology and GPR stratigraphy of a frontal part of an end moraine of the Laurentide Ice Sheet: Paris Moraine near Guelph, ON, Canada. Geomorphology, 75, 212–225. Sass, O. 2007. Geophysical quantification of talus thickness and rockwall retreat in the eastern European Alps. Journal of Applied Geophysics, 62, 254– 269. Sass, O. & Krautblatter, M. 2007. Debris-flowdominated and rockfall-dominated scree slopes: genetic models derived from GPR measurements. Geomorphology, 86, 176– 192.

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Upper Pleistocene glacial valley-junction sediments at Pias, Trevinca Mountains, NW Spain ´ RCEL DI´AZ1, I. PETER MARTINI2*, AUGUSTO PE´REZ ALBERTI1, MARCOS VALCA 3 VINCENZO PASCUCCI & STEFANO ANDREUCCI3 1

Departamento de Xeografı´a. Universidade de Santiago de Compostela, Santjago de Compostela, Spain 2

School of Environmental Sciences, University of Guelph, Guelph, ON, Canada

3

Dipartimento Scienze Botaniche, Ecologiche e Geologiche, Universita` di Sassari, Sassari, Italia *Corresponding author (e-mail: [email protected]) Abstract: Sediments at Pias (Galicia) provide evidence of Upper Pleistocene glacial activities at a valley junction in the north-western mountains of Spain. The sedimentary sequence consists of lower, predominately fine-grained lacustrine deposits with few lonestones, overlain by poorly sorted, sandy gravels interstratified with massive diamicton deposited during mid-Weichselian times (MIS 3) (marine isotope stage 3). The lacustrine sediments were deposited in a glacial valley temporarily dammed by a confluent glacier. The presence of active ice is suggested by massive diamicton layers best interpreted as till, rafted sediments in lacustrine deposits and deformation structures indicative of loading and kettle formation. Frozen ground conditions are suggested by a few involution-like structures. The Pias area contains one of the few western-Spain sedimentary records of Upper Pleistocene glaciation at relatively low latitude (about 428N) and low altitude (less than 1000 m a.s.l.). A southern dip of the Polar front to 40–458N latitudes, as occurred during Last Glacial Maximum, could have cut moisture to the northern Fennoscandinavian ice sheets. At the same time, however, sufficient precipitations would have persisted in northwestern Spain to sustain extensive ice caps and their outlets to elevations as low as c. 900 m a.s.l.

In the last decade the knowledge of the Upper Pleistocene glaciations of southwestern Europe has greatly improved, particularly for Galicia (northwestern Spain) (Fig. 1; Pe´rez Alberti & Valca´rcel Dı´az 1998; Valca´rcel Dı´az 1998; Pe´rez Alberti et al. 1994, 2004; Cowton et al. 2009). Galicia (centred at about 428300 N latitude) has elevations that exceed 2100 m a.s.l. and was subjected to sufficiently high precipitation and cold temperature to allow extensive mountain glaciations during the Quaternary (Rodrı´guez Guitia´n & Valca´rcel Dı´az 1994; Pe´rez Alberti & Valca´rcel Dı´az 1998). It has been suggested that thick and extensive ice caps formed on abraded, flat surfaces above 1400 m a.s.l., and valley glaciers may have extended as thin ice tongues down to altitudes as low as 700 m a.s.l. (Pe´rez Alberti & Valca´rcel Dı´az 1998). Geomorphologic evidence of former glaciations is common in NW Spain, but exposures of glacial deposits are rare. One exception is near Pias, southeastern Galicia where sedimentary successions, preserved in a terrace at the bottom of a narrow valley, are exposed in sand and gravel pits. The Pias deposits were sedimented at the junction of two glacial valleys; they therefore constitute a rare ancient record of those ice dammed settings including temporary lakes, commonly observed in recent

mountain chains (Hambrey 1994). Because glacial dams tend to leak during melting periods or even collapse, sediments of the dammed valley experience several cycles of deposition and erosion. The end product is a highly dissected sequence. Furthermore, postglacial processes such as floods and slope-wash or sediment-gravity flows may further modify the sedimentary succession. In particular cases such as that of Pias, however, parts of the original sequences are preserved allowing reconstruction of the events that occurred during the Pleistocene. The objectives of this paper are therefore to discriminate between glacial and postglacial mountain-valley deposits and to reconstruct some of the processes, specific environments and events that occurred in the Pias area at the junction of two ancient valley glaciers.

Study design and methods The complex sedimentary succession at Pias was examined through a sedimentological facies analysis to establish the various settings and their occurrence in respect to the valley junction (Reading 1986; Walker & James 1992). Accordingly, processes for each layer were determined by its sedimentary facies, the environments of sedimentation

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 93– 110. DOI: 10.1144/SP354.6 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Location map of study site in Galicia, NW Spain (dotted line is the eastern boundary of Galicia).

at each locality were established by the facies associations (sequences) and (important for this study) the relationship of all these in relation to the valley junction was established by the facies associations distribution (facies architecture or litho-stratigraphic units) and reported in some details in this paper. The overall glacial activity of the area was determined by evident geomorphologic features, and the approximate time of formation of the preserved sediments at Pias was established by Optically Stimulated Luminescence (OSL) dating of sand layers. The glacial geomorphologic evidence was obtained from field surveys and analysis of topographic maps, air photographs and satellite imageries. The sedimentological work required measurements of numerous sections at different times as the excavations proceeded. Finally, valid OSL dates were obtained from three sand layers.

The samples for OSL dating were collected using opaque polyvinyl chloride (PVC) tubes (D ¼ 8 cm; L ¼ 40 cm) of freshly exposed sandy sediments. Analyses were conducted at the Nordic Laboratory for Luminescence Dating, Risø (Denmark) following the methodology described by Murray & Wintle (2000, 2003) and Bateman et al. (2007). For each sample, approximately 100 g of sediment was used to estimate the water content (Table 1) and to calculate the natural radioactivity (dose rate, Dr). Dose rate calculations are based on high-resolution gamma spectrometry following the methodology described by Murray et al. (1987), and the cosmic ray contribution to Dr follows the Prescott & Hutton (1994) algorithm. Measurements of present-day water content were between 1.5 and 6.2%, upon which a +5% error was added to incorporate seasonal and longer-term fluctuations in moisture that the samples may have

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Table 1. Summary of dosimetry, dose equivalent (De) measurements and luminescence quartz ages. The samples (all sand lens) are named according to exposures taken from locations in Figure 4 Sample (n)

Depth# (m)

Water8 (%)

PIAS 1 PIAS 4 PIAS 8

3 5 7

3.7 3.2 4.3

238 U (Bq kg21)

228 Th (Bq kg21)

40 K (Bq kg21)

226 Ra (Bq kg21)

Dr§ (Gy ka21)

De (Gy)

n

Age (ka)

21.5 + 1.9 19.3 + 0.7 511.2 + 8.1 20.3 + 0.7 2.83 + 0.03 75 + 4 23 27 + 2 20.9 + 3.2 18.1 + 0.5 503.7 + 7.1 19.7 + 0.4 2.40 + 0.03 80 + 5 26 33 + 3 76.6 + 8.6 93.5 + 1.4 738.5 + 14.2 78.4 + 1.1 4.61 + 0.03 144 + 10 15 31 + 3

# Corresponds to the depth of samples below the top of the outcrop. 8Corresponds to selected water content chosen for calculation (see text for details). § Represents the conversion factors from activity concentrations to dose rate (from Olley et al. 1996). Note: n is the number of individual aliquots (10 mm diameter stainless steel disc) contributing to De

experienced since burial. These values were selected as average lifetime water content since time of deposition.

weathered-out crystals in some clasts; otherwise, very little clay is present in the study area because it may have been transported to more distal intermontane basins (Vergnolle 1990; Pe´rez Alberti 1993a, b).

Geological and geomorphological settings The lithology and geological structure of the region affected the localization and erosional efficiency of the Upper Pleistocene glaciers and the type of sediments that were produced. Galicia is underlain by upper Precambrian to Palaeozoic metamorphic rocks and granites of the Hesperian Massif that was originally attached to North America (Fig. 2a; Matte 1968; Den Tex 1981). These rocks were strongly deformed during the Palaeozoic Hercynian Orogeny that imparted a characteristic arcuate, westward-vergent, structural pattern (virgation) to most of Galicia (Fig. 2a). Thrust faults and nappetype folds superimposed deep basinal deposits over shallow marine and continental deposits. The materials were variously metamorphosed to phyllite and gneiss (‘Ollo de Sapo’ orthogneiss) and intruded by acid magmas (Fig. 2b). During the Mesozoic, the Hesperian Massif detached from North America, drifted eastwards and was dissected by numerous normal and strike –slip faults. During the Tertiary, north-western Spain was involved in collision and subduction with the Euro-Asiatic plate in the Biscay Bay area (Fig. 1). Differential block uplifts occurred during the Alpine Orogeny (Eocene– Oligocene), partly reactivating older fractures. The main faults strike SW –NE and south–north and are intersected by secondary NW– SE-oriented faults. Uplift led to entrenchment of streams and the formation of deep, partly fault-controlled valleys that separate extensive highland plateaus. This dissected landscape was affected by intense weathering under warm, humid conditions during the Tertiary. Granite and gneiss produced considerable amounts of sand, boulders, cobbles and pebbles and phyllites, and schist produced mostly fine-grained material. Kaolinite is found replacing

The Pias landscape The Pias area is located on the southern flank of the Trevinca Mountains whose highest point crests at 2127 m a.s.l. (Fig. 1). The mountainous region exhibits a relatively flat summit and relatively deep valleys dissect its flanks, several containing welldefined cirques (Fig. 2c). Two of these valleys, the Bibei Valley (where the studied Pias deposits are located) and the adjacent Barxacoba valley, merge just downstream from the main study locality (Fig. 2c). Field observations indicate rounded inter-valley surfaces with small shallow lakes such as at Lagoa de Barxacoba (1400 m a.s.l.), polished bedrock, streamlined features and broad dry channels joining the Barxacoba Valley to the adjacent Bibei Valley (Pe´rez Alberti et al. 2002). In the uppermost parts of the western flank of the Bibei Valley a series of discontinuous, longitudinal ridges (lateral moraines) up to 6 m high extend over a distance of c. 10 km (Fig. 3a). They are composed of coarse-grained, poorly sorted gravel containing numerous, disseminated granitic and gneiss exotics (up to 5 m in diameter) (Fig. 3b). Valley bottoms generally lack exposures of sediments except in terraces at sheltered locations such as at Pias.

Sedimentary lithofacies of the Pias area Various types of diamicton, gravel, sand and sandsilt rhythmite are exposed near Pias in large active sand and gravel pits (Fig. 4). The different sedimentary facies and their interpretation are described here and summarized in Table 2.

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Fig. 2. Schematic geological map of (a) Galicia and (b) area of study near Pias (arrow and vertical-line pattern). (1, homogeneous Augen gneiss ‘Ollo de Sapo’; 2, Augen gneiss ‘Ollo de Sapo’; 3, homogeneous migmatites with quarzites and schists intercalated; 4, shales; 5, granite; 6, blastomylonite; 7, valley deposits (glacial, periglacial and postglacial); 8, major faults; 9, fluvial network, lakes and water reservoirs). (c) Topographic map of the Pias area with glacial geomorphic features and location of the sand and gravel pit of Pias (dashed pattern).

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Fig. 3. Photographs of series of (a) lateral moraines (white arrows) and (b) strewn exotics on the high slopes of the Bibei Valley, at Cepedelo above the Pias study site (see Fig. 2c for location).

Matrix-supported diamicton (Dmm) Matrix-supported diamicton (Dmm) occurs as two distinct subfacies: one exposed along the northern part of the pit exposure (Fig. 4: Loc. 1–5) and the other along the southern part of the pit closer to the valley junction (Fig. 4: Loc. 6, 7). The first subfacies (Dmm1) is mainly a massive, sandy-silty deposit with disseminated pebbles of various sizes, isolated large cobbles and a few boulders up to 2.5 m in diameter (Fig. 5a, b; Pe´rez Alberti & Covelo Aveleira 1996; Pe´rez Alberti et al. 2007). The coarse clasts (pebbles to boulders) are mainly composed of granite (58 –70% depending on locality), quartzite (10–30%) and gneiss (6–60%), mostly subrounded to rounded while some are bullet shaped and faceted. The coarse clasts show preferred a-orientation modes downvalley but irregular imbrication. The deposit is compact and rests with sharp, basal contact upon poorly sorted, massive sandy conglomerate (Gm, see below) and rarely over massive sand. The upper contact is erosive, with channel cuts lined with cobbles and coarse-grained pebbles (Fig. 4: Loc. 1, 4; Fig. 5a, b). This diamicton subfacies occurs primarily in the upvalley part of the exposure as two major, composite (amalgamated), continuous beds

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(each up to about 2 m thick) in the middle part of the succession and as lenses (Dmm) as thin as 30 cm in the upper part of the succession. All these layers are cut off downvalley by channelized sandy gravels (Fig. 4: Loc. 4, 5; Fig. 6). The second diamicton subfacies (Dmm2) has silty, sandy matrix with numerous disseminated pebbles, some cobbles and rare small boulders. Its main characteristic is the presence of angular to subrounded, clast-shaped moulds filled with black to dark grey, sand-size material, possibly a weathering residue of ultramaphic rocks or reworked sandy concretions (Fig. 4: Loc. 6, 7; Fig. 7a, b). Some of the solid pebbles are coated by black, iron– manganese oxides. Dmm2 is generally massive; however, locally, the disposition of the pebbles marks inclined pseudo-laminations. Dmm2 occurs primarily as a thin, continuous layer in the middle part of the succession in the downvalley part of the exposure. It rests with a sharp contact upon sandy, pebbly gravel, locally cross-bedded. This diamicton has an irregular, deeply eroded upper surface and is overlain by massive, to cross-bedded, sandy gravels (Fig. 7a, c). In the upper part of the succession, Dmm2 occurs in isolated lenses in sandy gravelly deposits. Interpretation. We interpret the main strata of Dmm1 as basal till, but the possibility that the isolated lenses of this facies in the upper part of the succession may be derived from slope remobilization of diamicton originally deposited higher on the valley flanks cannot be excluded. Dmm2 probably represents soliflucted, glacier-front deposits or, the upper smaller lenses, reworked (soliflucted) valley-flank diamict deposits. The different size and composition of the clasts of Dmm2 in respect to those of Dmm1 suggests a different origin, perhaps deriving from the Barxacoba Valley glacier.

Coarse-grained, massive sandy gravel (Gm) Coarse-grained, massive sandy gravel (Gm) is poorly sorted and characterized by a pebble to cobble framework with some small boulders and various amounts of sandy matrix (Figs 4b & 6). The clasts are of granite and schist. Locally, imbricated pebble-clusters indicate a palaeocurrent down the Bibei Valley. This facies occurs both in thick (several metres) massive amalgamated units and also in tabular, metre-thick beds, separated by thin gravelly sandy lenses. Interpretation. We interpret the lower (below the main Dmm1 layer) occurrences of this facies as proximal glaciofluvial deposits mainly because of their intimate relationship with the diamicton

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Fig. 4. Studied pit in the Bibei Valley. (a) Photographs showing the Pias sand and gravel pit and measured sections (1 to 8) shown in (b) and reported in this paper. The Barxacoba Valley is shown in the background; downvalley direction to the right. (b) Map and vertical successions of lithofacies at the Pias pit, in relation to confluent valley and inferred glaciers (arrows indicate possible ice movement; labels for lithofacies are explained in Table 2 and text).

(Fig. 4: Loc. 1). Such a setting does not normally allow good development of braided-stream structures and pattern. The upper occurrences of this facies in the successions may represent proglacial to postglacial deposits of possibly powerful, hyperconcentrated floods.

Trough cross-bedded gravel (Gt) Trough cross-bedded gravel (Gt) has a framework of pebbles to small boulders, generally with

sandy matrix, locally with well-developed openframework (openwork) pebble lenses. The clasts are mostly subrounded and there is some welldeveloped, upvalley imbrication. This facies is best developed as a crudely cross-bedded coarser subfacies in association with massive to poorly structured gravel (Gm) in a deep (up to 5 m) channel cutting through most other deposits (Fig. 4: Loc. 5; Fig. 6) and in large cut-and-fill structures in the upper part of the southern exposure (Fig. 4: Loc. 6, 7; Fig. 7c).

Table 2. Lithofacies characteristics Label Dmm

Particle size Matrix-supported diamicton

Clast-supported, massive, sandy gravel

Gt

Gs

Trough cross-bedded gravel Planar cross-bedded gravel Planar-bedded gravel

Sm

Massive sand

Sl

Laminated sand

St

Cross-bedded sand

Sr

Ripple cross-laminated sand Rhythmite

Gp

Fl

Dmm1: sandy-silty matrix, disseminated pebbles to boulders Dmm2: silty-sandy matrix, disseminated pebbles, cobbles and a few boulders. It contains few disseminated cobble-size black, rounded, granular clasts of uncertain origin Pebble-cobble framework, isolated boulders, various amounts of sandy matrix Pebbles to small boulders, sandy matrix, locally openwork Pebbles, openwork to sandy matrix

Poor

Pebbles, occasional cobble, sand matrix Coarse-grained sand, few isolated pebbles a. Medium- to coarse-grained

Moderate to poor

b. Fine-grained Medium- to coarse-grained sand, some pebbles Fine- to medium-grained sand, silt

b. Good Moderate good

Fine- to medium-grained sand, silt, few lonestones

Sedimentary structures Massive

Possible environment Basal till Primary till or, in some layers, soliflucted previously deposited till

Poor

Massive, locally some preferred imbrication

Ice-marginal fluvial setting

Moderate to poor

Trough cross-beds, cuts-and-fills Steeply inclined foresets Plane beds

Subaqueous dunes, cut-and-fill structures in braided streams Transverse bars; small-scale Gilbert-like delta Supercritical flow in fluvial setting Sediment-laden water flow

Moderate to good

Moderate to poor a. Moderate, few pebbles

Good Good

Massive to faintly laminated a. Laminations, few small-scale cuts-and-fills b. Laminations Trough cross-beds Cross lamination; drift cross-lamination Laminas to thin beds

Supercritical flow in (a) fluvial and (b) hyperphycal flow lacustrine settings

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Gm

Sorting

Subaqueous dunes Mostly lacustrine setting in the Pias case Glacio-lacustrine setting with few ice-rafted clasts

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Fig. 6. Succession in the eastern flank of the midvalley part of pit (Fig. 4: Loc. 4, 5) showing the diamicton-and-gravel dominated upvalley succession locally cut downvalley by thick very coarse, poorly sorted, sandy gravelly deposits.

Fig. 5. (a) Sedimentary succession in the upvalley part of pit (Fig. 4: Loc. 1) (composite figures with pictures taken at different times; the lower section having become covered) characterized by a lower laminated to ripple-marked, fine-grained sand and silt, overlain by interlayered sandy gravel, gravelly sand, sand and highly compacted diamicton layers (Dmm1). (b) Detailed view of the lower main diamicton.

Interpretation. This facies records fluvial flood deposits and locally, in the southern exposures, acquires well-structured braided-stream characteristics.

Planar cross-bedded, sandy, small-pebble gravel (Gp) Planar cross-bedded, sandy, small-pebble gravel (Gp) occurs in steeply inclined foresets, locally

Fig. 7. (a) A succession in the downvalley part of the pit showing interstratifications of diamicton (Dmm2), gravel, plane- and cross-bedded coarse-grained gravelly sand to sandy gravel and minor medium-grained sand. Deformations and downcutting of diamicton layers occur (Fig. 4: Loc. 7). (b) Detailed view of a concretion of unknown origin reworked within a diamicton layer (black spots in Fig. 7a). (c) Massive diamicton sharply overlain by trough cross-bedded sandy gravel and gravelly sand.

grading into both subhorizontal topsets and laminated sandy bottomsets (Fig. 4: Loc. 8; Fig. 8a, b). The bottomsets are composed of fine- to mediumgrained sand with lenses of small-pebble gravels. The foresets dip upvalley (northwards) and consist primarily of openwork gravel alternating with sandy gravel, with the coarser clasts concentrated toward the bottom. This gravel facies occurs in the southern exposures of the pit associated with sandysilty rhythmites (Fl) (Fig. 8c).

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Fig. 8. A successions in the downvalley part of pit (Fig. 4: Location 8) characterized by (a) fine-grained, laminated sand and silt locally interdigitating and overlain by (b) cross-bedded sandy gravel. (c) Cross-laminated sand (Sr) alternating with plane-laminated sand (Sl). (d) Sandy-silty rhythmites (Fl) with pebbly lonestone (dropstone) bending and piercing laminas.

Interpretation. This deposit is interpreted to represent a small Gilbert-like delta prograding up the Bibei Valley.

Planar-bedded sandy gravel (Gs) Planar-bedded sandy gravel (Gs) consists of pebbles with occasional small cobbles, moderately to poorly sorted, with a sandy matrix. It occurs in planar beds, generally lenticular. This facies is well developed in wide, shallow channels cut into the top matrixsupported diamicton (Dmm1), with boulder to coarse cobble concentrations at the base (Fig. 4: Loc. 1; Fig. 5a). It is occasionally present in a few other parts of the succession primarily along the northern and middle part of the pit. Interpretation. This facies records supercritical flow conditions in fluvial settings.

Massive sand (Sm) Massive sand (Sm) is coarse-grained and moderately to poorly sorted. The sand has quartzo-feldspatic composition. Isolated pebbles are present in some layers. The massive sand occurs in lenses up to 80 cm thick alternating with pebbly, sandy gravel and gravelly sand. It is found underlying Dmm1 diamicton (Fig. 6: Loc. 4, 5). Interpretation. This is a poorly exposed facies and may represent rapid deposition from sediment-laden (possibly hyperconcentrated) floods.

Laminated sand (Sl) Laminated sand (Sl) is medium- to coarse-grained and moderately sorted, locally with a few pebbles. In places, the laminations are recognizable

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because of a change in grain size, in others because of changes in both grain size and composition. In the latter case, in the south part of the exposure, light grey and brown laminas alternate irregularly with dark grey to black laminas (mainly Fe –Mn oxide precipitation). Small-scale cut-and-fill structures and wavy laminations are also present. These sandy layers are generally deformed either in open folds or locally in tight convolutions. This sand generally occurs above a matrix-supported diamicton layer (Fig. 4: Loc. 1, 7). Occasionally, a coarsergrained variety of the sand with poorly developed laminas occurs as erosional remnants in sandy gravel (Gm/Gt) layers. A fine-grained version of this laminated sand occurs in association with ripple cross-laminated sand in the lower part of the downvalley part of the outcrop (Sr, see below) (Fig. 8c). Interpretation. The coarse-grained version represents supercritical water-flow deposits in fluvial settings; the fine-grained version is interpreted to have formed from hyperpycnal flows in lacustrine settings.

Cross-bedded sand (St) Cross-bedded sand (St) is medium- to coarsegrained and moderately sorted with disseminated small pebbles in some foresets. It occurs primarily at the southern part of the pit (Fig. 4: Loc. 7).

and silt laminas (Fig. 8d). This facies is generally intensely deformed by folding and, locally, by microfaulting. It is well developed in the southern part of the pit. Interpretation. This facies represents lacustrine settings receiving turbid (hyperpycnal) flows. Isolated cobbles and pebbles and clusters of coarser-grained sand attest to active sediment ice-rafting.

Deformation features in the Pias area Deformations are present in the sediments previously presented. They affect single layers as well as successions of layers. The following are the principal types.

Folds Folds are best developed in the middle-southern part of the pit where lenses of silty sand in a predominantly sandy gravel succession are exposed. They consist of complex downvalley-oriented folding of the finer-gained layers mainly due to the sudden superimposition of sandy gravel (Fig. 9a).

Involution-type structures Few involution-like structures are present in a sandy gravelly succession along the uppermost part of the

Interpretation. This facies represents subaqueous, fluvial dunes or cuts-and-fills.

Ripple cross-laminated sand (Sr) Ripple cross-laminated sand (Sr) is fine- to mediumgrained, fairly well sorted, alternating with medium dark grey, fine-grained, silty sand in thinner laminas (Fig. 4: Loc. 1, 6, 8; Fig. 8c). Ripple-drift crosslaminations occur with various angles of climb, which alternate in a semi-regular rhythmic fashion with quasi-horizontal or slightly undulating laminas. It is exposed primarily in the southern part of the pit. Interpretation. This facies represents a lacustrine setting affected by sediment-loaded floods (Jopling & Walker 1968).

Sand-silt rhythmites (F1) Sand-silt rhythmites (Fl) consist of alternating laminated (Sl) locally cross-laminated (Sr) well-sorted sand and silt (Fl) (Fig. 4: Loc. 1, 6, 8; Fig. 8c). Some layers are heavily stained with Mn –Fe oxides. Isolated coarse pebbles (lonestones) occur, bending and in places piercing underlying sand

Fig. 9. Folds: (a) irregular folds associated with loading and push (downvalley direction to the right) and (b) involution-like structure and modern tree solution tongues (black inverted-triangular features) (midvalley part of pit).

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eastern exposure (Fig. 9b). Recent tree-root solution tongues are also present, but are not to be confused with the involutions. Interpretation. Because higher areas in the region still experience seasonal frost, these involution-like deformations may reflect previous conditions of either deep seasonal frost or even permafrost. However, the latter is unlikely given the location of the study area and the relatively low elevation of the sediments.

Faults Faults generally have a small scale. Normal faults are the most common type but thrusts are also present. They occur frequently in fine-grained sands and silts, the thrusts being nucleated by climbing ripples in the southern part of the pit.

Metre-sized funnel-shaped structure Metre-sized funnel-shaped structures occur in relatively well-sorted, plane- to cross-bedded sand and gravel succession and in single sandy gravel beds in the southern central part of the pit. These are of two types. The first type resembles ice-wedge pseudomorphs when seen on flat outcrop surfaces (Fig. 10a). However, upon excavation, the structures are shown to be funnel-shaped rather than planar, as would be expected for ice wedges. A possible explanation would still require frozen ground conditions in that the funnels may have been formed by water expulsion forced by cryostatic pressures as suggested by Gangloff (1974) for analogous features observed in Quebec, Canada. However, Dionne & Pe´rez Alberti (2000) have interpreted these sandy funnels as water-escape features forced by the sudden addition of coarse-grained material over watersaturated, fine- to medium-grained, sandy deposits. Similar interpretation of structures in glaciofluvial Upper Pleistocene deposits in Canada was given by Burbidge et al. (1988) and in Denmark by Jensen & Miller (1990). Given their position at the top of the sedimentary sequence of the terrace, a third interpretation may be associated with fluvial settings experiencing strong flood discharges generating steep slopes along stream banks. The funnel structures may therefore have formed by the infilling of sinkholes associated with water piping through the lower sandy layers. Such a piping can be commonly observed in Upper Pleistocene successions such as those exposed along the northshore bluffs of Lake Ontario, Canada (Martini & Brookfield 1995) and elsewhere (Kochanov 1999). The second type is developed entirely in sandy gravel layers bounded by undeformed massive

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diamicton (Dmm2) at the base and sandy gravel at the top (Fig. 10b). Its interpretation is more problematic because it cannot be analysed in three dimensions. In any case, it is likely to have been formed by filling of a kettle generated by melting stranded ice blocks in a glaciofluvial setting.

Large funnel-shaped feature A large funnel-shaped feature occurs in the southern part of the exposure. It involves layers of sandy gravel, poorly-sorted massive sandy gravel and diamicton (Fig. 10c). A likely interpretation is that this represents a kettle generated by buried dead ice. Similar features are well documented in flood deposits of sandar (outwash plains) of southern Iceland (Fay 2002; Russell & Knudsen 2002), and are typical of recent and ancient, pitted, glacial outwash everywhere (Price 1973). Another possible process may be collapse-filling of a pipinggenerated void by water escaping from under a glacier dam. Similar events, although not necessarily similar structures, have been frequently documented at the recent junctions of valley glaciers (Clement 1984; Liverman 1987).

Discussion Three major topics are discussed: first, the significance and distribution of the sedimentary features observed in relation to the valley junction; second, the possible glacial model responsible for the formation of these valley-junction glacial sedimentary successions; and finally, the regional climatic context of NW Spain where these mid-Weichselian glaciations occurred.

Lithostratigraphic units and interpretation Valley fill deposits are usually highly dissected and, except for a few key beds, the various fragments do not necessarily represent time-equivalent events. Abstracting from the bewildering juxtaposition of sediment types, the sedimentary facies associations at Pias occur preferentially as three major lithostratigraphic units (Fig. 6). (1)

The lower unit is dominated by fine-grained deposits (Sl, Sr, Fl) with minor coarsergrained sand and gravel (mainly St and Gp). It represents fluvial-lacustrine settings that occurred both in the northern and southern part of the pit, but most likely they did not form at the same time. The OSL dates (Table 1; Fig. 4) indicate that the lacustrine sequence of the northern part is older than that of the southern part (closer to the valley junction) because it has been overridden by a

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Fig. 10. Funnel-shaped structures. (a) Structure developed in relatively well-sorted, cross-bedded sand and sandy gravel. Note more recent normal fault (slump?) and organic-rich slope deposits capping the succession (middle part of pit). (b) Structure developed in poorly sorted, crudely cross-bedded sandy gravel lying over undeformed Dmm2 diamicton (downvalley part of pit). (c) Large-scale funnel shape structures involving several layers of gravelly sand, poorly sorted sandy gravel and diamicton (downvalley part of pit).

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(2)

(3)

basal till (Dmm1) older than 27–33 ka BP, whereas the southern sequence has been directly dated to about 31 ka BP. The middle unit is complex, and the portion exposed along the northern and middle parts of the pit differs from that exposed at the southern end. A characteristic of this unit in the northern and middle part of the outcrop is the presence of diamicton (Dmm1) both as a few thick, laterally continuous, lower layers at the outcrop scale and in thinner upper lenses, alternating with horizontally crudely bedded massive sandy gravels. This unit is considered to be of glacial origin, formed directly by or in proximity of a glacier. The presence of two relatively thick matrix-supported diamicton (Dmm1) layers suggests re-advance of the ice. In the southern part of the exposure, the diamicton (Dmm2) is encased in well structures of sandy, gravelly deposits (Gt, St, Sl and similar). Part of this latter diamicton may be of direct glacial origin, but it could also have formed through slope remobilization of previously deposited glacial and soil material. In either case the diamicton is associated with a fluvial-dominated environment. The upper unit is primarily characterized by waterlaid, generally coarse-grained deposits (Gs, Gt, Gm, Sl) and few lenses of diamicton (Dmm). It is capped by slope deposits which are partly organic-rich, soliflucted from the valley flanks. Along the middle portion of the pit very coarse massive (Gm) to crossbedded (Gt) sandy gravels cuts through the whole succession (Fig. 4: Loc 5). This upper unit can be interpreted as proglacial to postglacial fluvial settings at times experiencing strong floods (particularly in the middle part of the pit) and locally developing regular braided stream characteristics, particularly in the southern part of the exposure. Some of the topmost diamict lenses may record local slope soliflution deposits or blocks of previously deposited glacial material eroded by floods.

Numerous modern fluvial-lacustrine analogues for the Pias case can be observed today in mountain regions at the junctions of two valley glaciers. The processes and deposits of such settings have been extensively described in the literature (Theakstone 1976; Liverman 1987; Donnely & Harris 1989; Johnson & Kasper 1992; Hambrey 1994). The temporary proglacial lakes experience fast deposition of a variety of sediments, but are recurrently drained either through breaches of the icy dams or by water siphoning though tunnels under the dams.

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During each sedimentation-erosion cycle, only small remnants of the original sediments are preserved. The final preserved sequence is composed of juxtaposed deposits of different age. This is what the facies-association distributions of the lower and middle units of Pias are recording; the following features are notable. The southern lower lacustrine sequence near the valley junction includes a small Gilbert-like (Gp) delta that prograded up the Bibei Valley. The southern middle succession of the pit in the Bienei valley is composed of material different from the northern one which was most likely derived from the damming element of the valley, the Barchacoba Valley glacier. The large funnel-shaped structure of the southern section might further indicate melting of buried dead ice or even filling of sinkholes, possibly formed by water escape through breaches or tunnels within the damming ice.

Glacial model Another fundamental question is the timing of the events that moulded the morphology and the sediment deposition in the area. The available OSL dates indicate that major glacial events at Pias occurred during the middle Weichselian glaciation (MIS 3). Radiocarbon dates from the nearby Tera Valley system, located approximately 20 km NE of Pias, suggest however that later glaciations during the Last Glacial Maximum (LGM) may have extended to low altitudes in some areas of the Trevinca Mountains as indicated by the age (c. 14 – 15 ka 14C BP) of the earliest postglacial lacustrine sediments (not calibrated date, Cowton et al. 2009). The Upper Pleistocene glacial conditions in the north-western highlands of Galicia, including the Trevinca Mountains, may have started during marine isotope stages MIS 4– 3. Smaller glaciers persisted in the highest mountains during the cold, generally drier LGM (25–18 ka BP; Cowton et al. 2009) and this also conforms to the finding of Vieira et al. (2001) who dated glacial tardiglacial deposits (13.1 + 2.0 ka; thermoluminescence date) in the Serra de Estrela in Portugal at elevations of c. 1590 m a.s.l. This fits with the general findings that the ‘Last Ice Maximum Extent’ generally predates that of the ‘Last Glacial Maximum’ (Jime´nezSa´nchez & Farias 2002; Garcı´a-Ruiz et al. 2003; Hughes et al. 2006; Van Vliet-Lanoe¨ 2007). The events that occurred in the Pias area can be reconstructed as follows. (1)

During local maximum glaciations, valleys were filled with ice fed by an ice cap covering the abraded highlands of the Trevinca Mountains. The local Last Ice Maximum Extent stage occurred prior to 30 ka BP, mostly during MIS 4, 3.

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(2)

During deglaciation, glaciers became increasingly concentrated into the valleys and interplay occurred in the Pias area between the Bibei and Baraxoba glaciers (Fig. 11; Pe´rez Alberti et al. 2002). At times, the Baraxoba glacier dammed the Bibei glacier. This fostered development of temporary shallow glacial lakes in the lower part of the Bibei Valley. The deposits at Pias formed under and in front of the Bibei Valley glacier and in the ice-dammed lake(s), and have been partially reworked by glacial floods and postglacial events.

Regional climatic overview In the regional contest, Pias is one of the areas in Galicia (NW Spain) recording relatively low latitude (c. 42.088N) and altitude (c. 900 m a.s.l.) Upper Pleistocene glaciations in Europe. Now, as in the past, NW Spain has a special climatic condition driven by proximity to the Atlantic Ocean, the presence of dissected mountains inland and the relative proximity to the hot African continent. This is responsible for the present maritime climate characterized by high precipitation and low

thermal amplitude near the coast, and a modified continental-alpine climate characterized by slightly lower precipitations and higher thermal oscillations inland. Both climatic types occur in an area where cold northern air currents interact with hot, southern tropical air (Castillo Rodrı´guez et al. 1999). During the Late Pleistocene similar but colder conditions existed in NW Spain as the polar front and the winter sea-ice limit in the Atlantic (North Atlantic Polar Front) fluctuated and repeatedly moved southwards to Portugal, as happened during the LGM (20– 16 ka BP) (Ruddiman & McIntyre 1981). Furthermore, climatic variations may also have been associated with Heinrich events that marked the major stadial periods (Heinrich 1988; Bond et al. 1993; Andrews 1998; Rahmstorf 2002). Indeed, Salgueiro et al. (2010) were able to demonstrate the influence of Heinrich events on the Spanish region correlating information from higher latitudes with the relative abundance of benthic polar foraminifera [Neogloboquadrina pachyderma (sinistral)] in a core taken offshore north-western Spain (Fig. 12). Data from that site ‘. . . corroborated a strong impact of the glacial polar front movements and the southward penetration of subpolar to polarwater [water] masses . . .’ (Salgueiro et al. 2010, p. 693).

Fig. 11. Inferred deglaciation stages of the Pias area. (a) Active glaciers during formation of the lateral moraine preserved high on the valley flanks. (b) Possible damming of the Bibei Valley by the large, more active Barxacoba glacier and formation of temporary lakes. (c) Detailed view of Figure 11b with inferred generalized sedimentation settings.

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Fig. 12. Climatic proxy data from the North Atlantic area and core SU92-03 taken offshore NW Spain (modified from Salgueiro et al. 2010). (a) Correlation of Greenland ice cores d18O GISP (black: from Stuiver & Grootes 2000) and of d18O NGRIP in the older part (grey: modified from NGRIP-members 2004) (18O/16O ratio d18 expressed in parts per thousand (‰) relative to the standard mean oceanic water (V-SMOV): values of about and lower than c. 238‰ indicate glacial conditions). (b) Relative abundance of polar benthic foraminifera Neogloboquadrina pachyderma (sinistral). (c) Abundance of lithic fragments per unit dry dry-weight sample. (d) Summer surface water temperature (SSTsu) estimated with the analogue technique SIMMAX 28 (Pflaumann et al. 1996). H11: Heinrich events; MIS 3: marine isotope stage (14C dates are calibrated).

Whereas colder climatic conditions and southwards shifting of the North Atlantic Polar Front may have generated cold, dry conditions in northern Europe, sufficient moisture may have remained available to the north-western part of Spain throughout the Late Pleistocene for ice caps to form on highlands and for active outlet valley glaciers to extend to low altitudes.

Conclusions The following conclusions can be drawn from this study. (1)

Evidence of glaciation in the Pias area is provided by erosional (such as cirques, rock

(2)

(3)

(4)

basins and striated bedrock) and depositional landforms (such as moraines) high on the valleys flanks and by sedimentary deposits with diamicton (interpretable as tills) in valley bottoms. Proglacial settings are evidenced by glacial lake and valley-train deposits with some diamicton interlayers and by deformation structures possibly related to stranded ice blocks. Unsorted, fluvial deposits and relatively loose gravity-flow deposits at the top of the succession now covered by plant litter at Pias attest to postglacial reworking and deposition. The sediments at Pias offer a rare example of ancient interaction between two valley glaciers and the development of temporary

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(5)

lakes at their junctions. Lacustrine successions and adjacent glaciofluvial deposits are well preserved in a terrace. OSL dates indicate that the widest extent of glaciers (Last Ice Maximum Extent) most likely occurred during the middle Upper Pleistocene (MIS 4, 3) in the Pias area, prior to LGM (MIS2). This agrees with what has been reported from other parts of Galicia and Spain. Although there is no evidence for the presence of an active late glacier at Pias, LGM ice may have persisted in other parts of the mountainous area and (locally) may have eroded previous sedimentary records, as indicted in the adjacent glaciated Valley de Tera where the earliest postglacial lacustrine deposits date to 14–15 ka 14C BP.

We would like to thank R. Blanco Chao and B. Carballido for their help with the graphical design. J. C. Dionne, L. A. Owen and J. Rose read early versions of the manuscript and provided suggestions for improvement. We are grateful to H. M. French for editorial help. We are indebted to A. Murray and the NLL group for laboratory assistance. Financial support for this research was provided by the University of Santiago de Compostela (Spain) to APA & MVD, the Natural Science and Engineering Research Council of Canada (NSERC) to IPM, the Italian Ministry of University and Scientific and Technological Research (MURST) to VP and young researcher grant from Regione Autonoma Sardegna: PO Sardegna FSE 20072013 and L.R.7/2007 ‘Promozione della ricerca scientifica e dell’innovazione tecnologica in Sardegna’ (resp. Stefano Andreucci).

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Palaeoshorelines of glacial Lake Kuray –Chuja, south-central Siberia: form, sediments and process P. A. CARLING1*, M. KNAAPEN2, P. BORODAVKO3, J. HERGET4, I. KOPTEV5, P. HUGGENBERGER6 & S. PARNACHEV7 1

School of Geography, University of Southampton, Highfield, Southampton, SO17 1BJ, UK 2

HR Wallingford Ltd., Howbery Park, Wallingford, OX10 8BA, UK

3

Institute of Monitoring on Climatic and Ecological Systems, Siberian Branch of the Russian Academy of Sciences, Akademicheskii Prospekt 10/3, 634055 Tomsk, Russian Federation 4

Geographisches Institut, Universita¨t Bonn, D-53115 Bonn, Germany

5

Institute of Geology and Oil and Gas Industries, Tomsk Polytechnic University, 30 Lenin Avenue, 634050, Tomsk, Russian Federation 6

Geologisches Institut, Departement Geowissenschaften, Universita¨t Basel, Bernoullistrasse 32, 4056 Basel, Germany

7

Tomsk Petroleum Institute (TomskNIPIneft), 72 Mira Ave, 634027, Tomsk, Russian Federation *Corresponding author (e-mail: [email protected]) Abstract: Glacial lake Kuray–Chuja which occupied the Kuray and Chuja Basins during the Quaternary Period (Marine Isotope Stages 2 and 3) left distinct shoreline features around the basin margins. At the greatest extent the lake had a surface area of 2951 km2 with wind fetches up to 70 km. Wind waves constructed erosional, erosional-accumulative and accumulative strandlines, the latter including spits, tombolos, barrier beaches and offshore bars. Strandlines range in altitude between c. 1600 and 2100 m, the range in altitudes demonstrating lake level variations through time. The shoreline forms and stratigraphy are detailed. Particular attention is given to the presence of pocket beaches on rock-coasts, the alluvial strandlines of which have distinctive alongshore gradients. Although reasonably ascribed to regional differential tectonic uplift, the possibility that the local alongshore gradients are augmented as a process response to alongshore drift of sediment in a sediment-starved system of closed coastal cells is explored using a simple numerical model.

Ice-dammed lakes are associated with glacial environments and there is increasing evidence that the largest lakes may have influenced the climate of Quaternary glaciations (Obolkin & Potemkin 2006). The study of ice-marginal lakes can therefore provide information on quantities and timing of ice-melt events and catastrophic flood events, as well as regional climatic change information (Krinner et al. 2004). Moreover, a consideration of lake strandline-building processes can provide palaeowave and palaeowind information (Dupre´ 1984; Jewell 2007) at the local scale (Adams & Wesnousky 1998), adding insight into the palaeobeach process regimen. Deviation from the horizontality of strandlines has been used to measure tectonic and isostatic adjustments (Gilbert 1890; Crittenden 1963; Bills & May 1987; Adams 1997; Teller 2001). However, the natural variation in strandline altitudes due to strandline building

processes needs to be accounted for when considering tectonic deformation; this paper provides an important caveat in this respect. This study considers palaeolakes within the Altai-Gobi transition zone of central Asia (Grosswald & Rudoy 1996; Komatsu et al. 2001; Rudoy 2002; Gillespie et al. 2008; Komatsu et al. 2009). Often strandlines are preserved at multiple levels around these lake basins. Consequently, this paper provides an introduction to one of the largest palaeolakes within the Altai Mountains of southern Siberia and details the basic characteristics of the strandline assemblages. In particular, the paper considers whether the sedimentology, stratigraphy and shoreline morphology can provide information on the strandline building processes and notable palaeowave and palaeowind regimens and whether the alongshore drift of sediment can, in certain circumstances, be sufficient to induce a significant local

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 111–128. DOI: 10.1144/SP354.7 0305-8719/11/$15.00 # The Geological Society of London 2011.

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alongshore gradient in the shoreline height. Several factors control the development of lake strandlines including wind-wave regimen, lake level fluctuations and the local configuration of the lake basin shoreline geometry (Davies & Johnson 2006).

The study area The lake studied is the ice-dammed Kuray–Chuja palaeolake in the Kuray and Chuja Basins within the Altai Mountains (Fig. 1).

Geological, tectonic and geomorphologic context The lake basins have a complex geological and tectonic history (Bogachkin 1981; Ota et al. 2007). The key issue is the presence of two intermontane basins (the Chuja and Kuray basins) (Fig. 2), separated by the Chagan Uzun Interbasinal Uplift (CIU) (Rogozhin et al. 2007) but linked by a narrow valley. The surrounding mountain ranges, up to 4000 m in height, sustained icecaps and valley glaciers during the Quaternary up until today. During the Cenozoic, the area underwent considerable SSE–NNW compression associated with the Indo– Eurasian collision (Buslov et al. 1999; Babichev et al. 2009). The Kuray–Chuja basin developed initially as a middle Miocene –Lower Pleistocene graben, with vertical relative movement of the order of 2500 m primarily occurring during the Pliocene–Pleistocene transition. The area is dissected by numerous reverse and thrust faults and strike– slip faults with vertical components (Rogozhin et al. 2007). Further motion occurred again at the end of the Lower Pleistocene and during the Upper Pleistocene, and may have included

Fig. 1. Location map of the study area in southern Siberia (Russian Federation). The Kuray (K) and the Chuja (C) Basins are shown in the headwaters of the Chuja River.

isostatic movements due to deglaciation and lake drainage (Carling et al. 2002; Herget 2005). Cenozoic deposits infilling the Chuja and Kuray basins include Quaternary glacial moraine and glaciolacustrine deposits at the margins (Zol’nikov et al. 2010). Within the Chagun-Uzun River valley (Fig. 2b) at the western end of the Chuja Basin spectacular and extensive exposures of mid-Pleistocene lake sediment sequences can be found, which constitute parts of one of the most complete and important sections of the Cenozoic within central Asia (Zykin & Kazanskii 1995; Sheinkman 2002). Quaternary alluvial fans are also well developed, in places forming a piedmont, and the lower portions of these fan surfaces were sublacustrine at times of high lake stands. The flanks of the basins are delimited by highly fractured bedrock of a variety of lithologies. Ice-dammed lakes formed in both the Chuja and Katun tectonic depressions (Carling et al. 2002). Several authors have provided brief reports of the Kuray– Chuja Lake primarily because it is believed to have been a source of water for catastrophic megafloods (Baker et al. 1993; Carling et al. 2002; Herget 2005). At the maximum stand, glacial Lake Kuray– Chuja impounded at least 594 km3 of precipitation and glacial ice meltwater during the Late Glacial (40 –13 ka) and was 600– 700 m deep at the deepest points. The lake had a total surface area of 2951 km2 and maximum fetch-lengths of the order of 60–70 km in the Chuja Basin and 20– 25 km in the Kuray Basin when the lakes were ice-free (Borodavko 2003). This paper considers primarily the Upper Pleistocene conditions as are evident from multiple, wellpreserved strandlines that exist around the basin margins from between c. 1600 m and a maximum altitude of 2100 m (Fig. 3). On steep rocky slopes, the strandlines are often faint and are best seen under particular light conditions, seasonal vegetation development or light snow cover. The strandlines can be traced over many kilometres but stop near the town of Aktash (Fig. 2a) where glaciological studies indicate that three large glaciers converged and impounded the sole drainage outlet from the basin (Herget 2005; Carling et al. 2009). Strandlines are also absent from other locations within the basins, where it is known that Pleistocene glaciers extended into the lake waters (Herget 2005; Zol’nikov et al. 2010) including during the Late Glacial, but otherwise the basins were not glaciated at this latter time. At times of high water level there was one conjoined lake in the Kuray and Chuja basins. At very low stands during the Upper Pleistocene, however, it is evident that a separate low stand lake occurred in the Kuray Basin as faint accumulative strandlines in the southeast of the Kuray Basin curve around and cross the entrance

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Fig. 2. (a) Sketch map of the Kuray Basin showing primary locations and features and (b) sketch map of the Chuja Basin showing primary locations and features.

to the narrow connecting valley to the Chuja Basin. The basins now contain no substantive bodies of water. The timing of upper Quaternary regional lake highstands is uncertain (Komatsu et al. 2001) although three highstands have been inferred for Lake Kuray –Chuja at 40, 25 and 13 to 8.7 ka BP

(Borodavko 2003; Herget 2005). Strandlines at altitudes of 1860 and 1610 m on the northern flank of the Kuray Basin have been dated for this study using radiocarbon at 32 290 + 260 and 8700 a BP, respectively. Reuther et al. (2006), using cosmogenic dating, argue for a final draining at

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Fig. 3. Vertical air photograph of strandlines on the southern margin of the Kuray Basin.

15.8 + 1.8 ka BP of the Chuja –Kuray Lake. Rudoy (1988) reports the development of pingos on the basin floors from radiocarbon assay in the Kuray Basin from 10.8 + 0.080 ka BP and in the Chuja Basin from 3.8 ka + 0.105 ka BP, at which time the basins must have been dry.

Methods The locations and altitudes of significant shoreline features were recorded using a hand-held GPS. Plan-view locations are accurate to within 10 m, whereas altitudes are accurate to +5 m. The altitudes were calibrated using a barometric altimeter with reference to the altitude of spot heights on topographic maps. The fieldwork was supplemented by consideration of panchromatic 1:30 000 scale air photographs, the 1:200 000 map dated 1992, produced by the Federal Service of Geodesy and Cartography, Russia, the 1:1 000 000 pre-Quaternary geological map dated 1978, produced by the USSR Ministry of Geology, and 1 m resolution Ikonos satellite images. Notes were made of the stratigraphy and sedimentology of minor exposures and small pits were opened for the purposes of this study. Grain-size distributions were determined through sieve analysis. Topographical surveys (Geodolite) of selected strandlines were completed as were two groundpenetrating radar (GPR) profiles (Beres et al. 1995). Topographically corrected GPR profiles consisted of one transverse and one parallel traverse of a shoreline, close to an inspection pit. A pulseEkko 100 GPR system was employed with 100 MHz antennae and 2.0 m antennae separation, 0.25 m trace spacing, 0.8 ns sampling rate, 400 V transmitter and 128 vertical stacks. Profile lengths varied from 300 to 400 ns. The common midpoint method was used to estimate depths.

The shoreline model STWAVE (Smith et al. 1999; Smith 2007) was selected to explore strandline response to wind-wave climate. The model contains the main processes to simulate the gravel beach processes: depth-induced wave refraction and shoaling; current-induced refraction and shoaling; depth-induced and steepness-induced wave breaking; as well as diffraction. The model was driven using a Jonswap spectrum with the wave amplitude determined by the fetch-limited wave height predictor of Hasselman et al. (1973). As long as there is sufficient sediment available, the height of the beach is directly related to the wave energy of the highest waves reaching the beach (Cooper et al. 1996; Lorang 2002). STWAVE is therefore coupled with a model that links wave breaking energy with beach crest height for gravel beaches (Lorang 2002). The wave model calculates alongshore wave height variation, due to refraction along the strandline. These wave height variations are than translated into beach crest height variations. All calculations are based on steady-state uniform offshore wave boundary conditions and no tidal or winddriven flows. Wind-wave generation and wave dissipation due to bottom friction are neglected; which are fair assumptions on the small spatial scales (,1 km) investigated here. Further details of the model are given in the section ‘Pocket beach crestline model’ below.

General characteristics of the strandlines The strandlines were formed on bedrock and colluvial slopes, end moraine complexes, river terraces and piedmont fans around the basin margins. Many of the strandlines are formed against steep (reflective) rock slopes (Fig. 4a –d); each strandline tread therefore often consists of a combination of an erosional platform with an associated wedge of sediment, although locally the strandlines are purely erosional. However, in the east of the Chuja Basin especially, slopes are gentle and often alluvial. In these latter dissipative environments, tombolos (Fig. 4e) are found linking former islands as are barrier beaches closing former lagoons. At the lowest elevations in the southeast, series of indistinct offshore sand bars are evident. As the water level in the lake rose, or fell, strandlines were inundated or abandoned in subaerial positions. In the first case, the strandlines may have been reworked or preserved albeit with morphological modification by processes acting at or below the wave base. Similarly, abandoned strandlines may have been modified by subaerial processes but are largely well preserved. The excellent preservation must be related to the recent age of the strandlines and the coarseness of the beach gravels. In all cases a

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Fig. 4. (a) A series of strandlines can be seen to the left of the dark shadowed rock outcrop. Below the outcrop a series of accretionary capes are visible towards the foreground (three of which are arrowed). Alongshore drift direction left to right. (b) A series of stacked recurved spits (arrowed) are visible downdrift of the rocky headland centre left. (c) View downdip along tread of tilted shoreline (see also Fig. 6); note the beach gravel consists mainly of angular reworked colluvium. Steep risers are visible top left and bottom right. Tread is 30 m wide. (d) View from downdrift rock headland towards updrift rock headland of pocket embayment. Horizontal strandlines are visible at the back of the embayment top centre and left centre. These levels cut across a prominent tilted shoreline (centre). Black dot (centre) on the tread of the tilted strandline is a car for scale. Regional alongshore drift direction is out of the image across the embayment. (e) An offshore tombolo (arrowed) between two former islands. Field of view in all panels except (c) is 1 to 2 km.

period of time was necessary for stable new strandlines to develop at a given elevation, but the duration of stillstand of the lake level at any given strandline is unknown. The strandlines throughout the Kuray Basin and on the northern flanks the Chuja Basin are essentially horizontal. However, on the southern flanks of the Chuja Basin, sets of well-defined horizontal strandlines cuts across, but do not erase, a set of sedimentary bodies that traditionally have been interpreted as an earlier generation of strandlines

(Svitoch & Khorev 1975). These latter strandlines (see Pocket beaches below) are tilted with alongshore gradients up to 58 (Carling et al. 2002) rising to the southeast (Fig. 4d). The NW–SE extension of the basins (Figs 1 & 2) ensured that during periods of maximum lake level an uninterrupted fetch existed of c. 65 km in the case of the Chuja Basin and c. 30 km for the Kuray Basin. During the Quaternary, a cold anti-cyclone developed over central Asia (Yu & Wright 2001; Velichko 2005) such that the prevailing winds

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Fig. 5. Dropstone brought by iceberg in NE area of Chuja Basin, elevation 1890 m. Photograph by V.V. Sapozhnikov in 1911 near Korkorya (Fig. 2b) (Sapozhnikov was a notable pioneering Russian glaciologist in the region).

within the Altai were from the north and northwest (Velichko 1984; Back & Strecker 1998) and were funnelled to the southeast by the basin topography, making conditions favourable for the development of wave trains emanating from the north and NW.

Dropstones Iceberg-calving and ice-drifting left putative dropstones displaying glacial striae (up to c. 8 m blocks) in widely dispersed locations throughout the Kuray and Chuja basins (Fig. 5). The Kujektanar and Tydtugem glaciers (Fig. 2) are believed to have periodically blocked the valley connecting the Chuja and Kuray basins impeding drainage from the Chuja Basin. Evidence for glaciers within the connecting valley is moraines flanking the main valley at elevations of 18 –20 and 35–40 m. However, this evidence almost certainly post-dates the period of main lake development. As noted above, the basins were not glaciated in the Late Glacial. However, evidence for a dropstone origin include dropstones consisting of rock type found in the Kujektanar and Tydtugem valleys to the west as well as dropstones sitting on wellformed strandlines where the iceberg presumably grounded.

Classification of strandline types Using the approach of Leontiev (1963), the following basic types of strandline were mapped: erosional, erosional-accumulative and accumulative shores.

Erosional strandlines Erosional strandlines are defined by waves cutting knickpoints into the slope materials. As is detailed

below, individual strandlines in the Kuray –Chuja Basin can vary over short distances from erosional knickpoints cut in hard bedrock, fissile bedrock and compacted colluvium to accumulative beaches formed of accreted sand and gravel. The defined bedrock knickpoints indicate that although lake levels varied, stillstands were often long enough to erode bedrock. Laterally extensive erosional strandlines, with risers a couple of metres high, occur in competent bedrock only locally. Usually, they are found on steep friable outcrops that formed headlands in the former coastline. Elsewhere they are well developed (albeit with small riser heights) in end moraines, alluvial fans and debris cones. This type is well represented on the southeast and northeast margins of the Chuja Basin and on some western shores of the Kuray Basin. Erosional strandlines are noted particularly on the steeper slopes at the higher elevations associated with the higher lake stages including the maximum transgressions (1980– 2100 m). Here wave action attacked the bedrock directly and there are few topographic lows in which substantial accumulative beaches might form.

Erosional-accumulative strandlines These strandlines consist of variable quantities of sediments accumulated on erosional platforms and represent more than 50% of the palaeoshores. Often the sediment veneer is thin (,5 m) and formed either boulder– pebble, pebble–granule or sand beaches. The boulder– pebble beaches mostly occur at the highest elevations where the basin margins are the steepest. The treads are some 5 to 12 m in width with the offshore slopes of risers being up to 208. The steep reflective foreshore slopes ensure that any surf zone would be narrow or absent with the beach face being dominated by run-up. The large angular clasts are often very well packed and, as such, would have presented transport-resistant structures limiting sediment entrainment (Carter & Orford 1993). The pebble–granule beaches are more extensive and often occur downdrift of coarser beach units due to alongshore downstream fining. Sandy erosional-accumulative beaches are rare and mainly found in the east.

Accumulative strandlines Zones of prominent accumulation of alongshore drifted sediments include accretionary capes and spits, the latter occurring at changes in coastal alignment such as those which occur immediately downdrift of former headlands. Often a series of spits occur at a given location associated with different strandline elevations but with similar alongshore

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topography. These spits are easily seen on aerial photographs and satellite images. Spits at intermediate elevations can be very distinctive and are linear elongated forms (up to 250 m in length). The widths of spits are usually c. 15 –50 m and the heights are rarely more than 10 m. The shore-normal morphology is often asymmetrical, with the gentle slopes occurring to landward and very steep slopes occurring to lakeward. The surfaces of spits are usually flat and straight, but sometimes slightly wavy. Some spits are characterized by recurved heads and a few spits have bulbous heads representing coalescence of several indistinct minor spit heads. Generally, at low elevations where the slope towards the centre of the lake basin slope is slight, spits may still be present but are very indistinct with little surface expression. The linear extent of accumulative capes is significantly less than that of spits, and rarely extends more than 50 to 60 m from the headlands. The capes have more rounded platforms, but otherwise are structurally and texturally similar to the spits on account of similar genesis. In a few locations, tombolos are found which link former nearshore islands or coastal headlands. The fact that quite small and distinctive strandlines and delicate spits are often well preserved at differing elevations, but at the same general location, demonstrates that an active wave regimen was maintained for a considerable time from winds emanating from the same general quarter. It further hints that water levels were generally regressive with discrete periods of stabilization during the main periods of the formation of strandlines.

Pocket beaches In the south-eastern part of the Chuja Basin distinctive erosional-accumulative tilted strandlines are found within rock embayments along the basin margin (Figs 4 & 6). Whereas the individual horizontal strandlines reported above may be traced for many kilometres, the tilted strandlines are isolated accretionary units of limited lateral extent but of much greater bulk than the horizontal strandlines. Each beach is formed within a bay with terminations due to former rock headlands on either side. As such, they constitute ‘pocket beaches’ formed between headlands (Adams & Wesnousky 1998; Backstrom et al. 2008; Bowman et al. 2009). The tilted strandlines are often stacked one above the other and each may terminate in a spit (Fig. 4b) providing a suite similar in appearance to strandline accretion units described and illustrated by Gilbert (1890, plate XXV). Each Chuja Basin pocket beach is usually at a low altitude and is narrow updrift within the embayment. Close to the downdrift headland, however, it broadens and is

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Fig. 6. (a) View to south into embayment along the southern shoreline of the Chuja Basin south of the village of Tebelar below point 2075 (Fig. 2b). Note the prominent tilted accumulative strandline at the back of the embayment centre-right. This level can be traced into the left-hand embayment at a high elevation. (b) Close up of strandline in (a) showing horizontal strandlines cutting across the tilted strandline. Alongshore drift in both panels is right– left. Cattle provide scale.

at a higher altitude, exhibiting a bluff or slight recurved termination just downdrift of the headland (Fig. 4a, b). Immediately to the SE, in the lee of the headlands, strandlines are locally absent or are very poorly developed due to sediment starvation. The increase in beach volume downdrift is a characteristic of some modern pocket beaches (Carter & Orford 1993). In all cases, the spits lack compound or complex morphologies which is consistent with the dominance of one significant wind-wave direction (Krist & Schaetz 2001). Within embayed settings, a change in lake level not only implies a shift of the shoreline location but also results in a change in the effective embayment configuration in terms of the hydraulic conditions and beach sediment flux (Dashtgard et al. 2007). The latter should influence successive strandline alongshore gradients as lake levels varied. A later generation of faint, parallel and horizontal strandlines cuts across the pocket beach accumulations (Fig. 4d). The smaller features usually have risers of about 1 m and tread widths of a few metres, which are particularly evident on the treads of the pocket beaches although subaerial erosion has

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obscured them on the pocket beach risers. At the extreme, one of these pocket beaches rises 55 m in a distance of 700 m (0.078) but is otherwise undistorted; it would be perfectly horizontal if the regional trend was subtracted. It therefore appears that this generation of strandlines has been affected by regional uplift of tectonic blocks in the SE of the region (Zol’nikov et al. 2010). Other pocket beach alongshore gradients can be as small as 0.0052 but are usually in the range 0.24 to 0.056. However, the alongshore drift direction is also to the SE such that any alongshore systematic gradient changes due to sediment transport would compound or influence any tectonic interpretations obtained from the shoreline gradients. It is posited that the alongshore gradient of smaller pocket beaches in particular might reflect alongshore drift processes (Fredsoe & Deigaard 1995; Short 1999) as well as tectonic uplift. It is therefore hypothesized that alongshore drift entrains the sediment supplied within each embayment and deposits the material along the downdrift margins of each headland. As the supply of sediment within the embayments is exhausted, sediment is eroded from the updrift portions of the margin inducing the tread to narrow and fall in altitude. Downdrift, erosion is less and gives way to deposition near the headland apex; hence the broadening of the tread and the increase in altitude. The problem with this hypothesis is that there is little supporting analogous strandline development on modern coastlines, or other Pleistocene strandlines. Few, if any, prior studies have considered the alongshore gradients of pocket beaches. However, Zenkovich (1967) provides qualitative descriptions of depositional capes broadening towards headlands in the Black Sea with downdrift shore gradients rising by around 18. In coastal seas however, in addition to the alongshore drift, there are usually substantial additional transfers of sediment from offshore as well as zones of coastal erosion or river inputs to the coast to replenish sediment supply, countering the downdrift progressive exhaustion of beach sediments. The explanation for the local steep gradients might therefore lie in the fact that each embayment-headland along the margin of the Chuja Basin effectively formed a relatively small, closed, sediment transport cell. In this model transport is only induced by alongshore drift with very little additional supply of sediment from the steep offshore zone and a restricted supply from onshore sources, as each embayment generally has exceedingly small colluvial sediment source areas. Over time, sustained drift from NW to SE would induce the steep alongshore profiles by the sediment exhaustion model outlined above, with little sediment being transferred to the next cell along the coast.

Stratigraphy and sedimentology Kuray Basin strandline sediments consist of angular to subrounded fine to very coarse pebbles (mode 23 to 25 phi) occasionally with small cobble units or isolated cobble clasts present. Lenses of sand (2 to 3 phi) and coarse silt, a few centimetres thick, occur locally but only rarely form units of significant thicknesses (.100 mm). Chuja Basin strandline sediments can be coarser, including angular small boulders. The few shallow pits opened in the Chuja Basin have similar stratigraphy as described below for the Kuray Basin. Sections within erosional-accumulative shorelines in the SE of the Kuray Basin at elevations between 1600 and 1700 m are presented in Figure 7. Section 1 (Fig. 7a) is cut into moraine by waves emanating from the west and NW. The surface layer 1 consists of a well-sorted coarsegranule soil horizon. Layer 2 is water-washed medium sorted sandy-gravels, locally demonstrating graded bedding dipping offshore. Layer 3 is gravelly but rich in quartz-feldspar sands and silty sands which can be attributed to calm-weather deposition; these deposits are not usually preserved as thick units due to later erosion by storm waves. Well-sorted, coarse to fine sequences with relatively extensive bedding, exhibiting low dips, can be interpreted as individual shore-building swash/ back-swash dominated units termed ‘selection pavements’ by Bluck (1999). In Figure 7a only two major units of this kind can be identified: layers 8 to 7 fine upwards as do layers 6 and 5. Layer 4 contains abundant organic detritus in silty sand and might be interpreted as a jetsam depositional unit during raised summertime lake level regimen due to the annual melt cycle. Layers 3 and 2 are not typical of selection pavements; rather they contain lenses of fine gravel or of sand and outwards-dipping bedding with an abundance of poorly-rounded and angular gravel. These units might represent locally derived colluvium, little reworked by wave action. Although it is probable that water levels were sustained for some time at this elevation, the coarseness of the gravels and the limited accumulation ensures that little if any evidence of individual storm events or local fluctuations in water level are recorded. Section 2 is cut into moraine (Fig. 7b) and has better defined bedding than the strandlines described above (Fig. 7a). Several distinct selection pavements might be distinguished. Event 1: layers 10 and 9; event 2: layers 8, 7 and 6; event 3: layers 5 and 4; event 4: layer 3 and event 5: the inset layer 2. Layer 4 contains in situ rootlet traces and represents a temporary fall in water level during which colonization of the foreshore by herbs occurred. Section 3 (Fig. 7c) is similar to

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Fig. 7. Sketches of small sections cut in Kuray Basin strandlines on the south-eastern shore. Numbering refers to distinctive layers picked out primarily by differences in grain size (see text for detail).

Section 2 but is not as well developed. Three shorebuilding events are recorded: Event 1: layer 5; Event 2: layers 4 and 3; and Event 3: layer 2. Layer 2 consists of distinctly less-rounded gravel than the rest of the sequence and must reflect a short period

less-energetic depositional condition with short pebble transport distances. The cross-shore gravel-rich stratigraphy of an accumulative strandline at 1600 m altitude on the northern slopes of the Kuray Basin is presented in

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Fig. 8. Sketches of small sections cut in Kuray Basin strandlines on the northern shore (modified from Butvilovsky 1993). Distinctive layers are picked out primarily by differences in grain size and style of bedding (see text for detail).

Figure 8. Although the location in respect of the prevailing palaeowind differs, the setting is not otherwise dissimilar to that of the strandlines in Figure 7. The surface of the 30 m wide tread (Fig. 8) dips offshore. The upper edge of the tread is marked by an indistinct riser with a surface accumulation of larger well-rounded clasts, mixed with angular colluvial gravel particles. The tread at this point is incised into the sediments of the next strandline upslope. Distinct, steep, lake-facing unconformities reflect reactivation surfaces due to wave cutting, with lakeward accretion characterized by cross-stratified gravels lying above unconformities. Although offshore dipping strata represent wavebreak swash action cutting the outer margins, no distinct water level is evident from the strata packages and the preserved strata packages must represent only a small proportion of the layers that would have been deposited and eroded repeatedly during the growth of the beach height. The presence of the coarse-gravel/fine-gravel layering represents energetic wave conditions probably during storms, while sandy lenses represent calm weather transport with sorting and deposition of fine grains by small waves traversing stable gravel layers. Each of these described beach accumulations are no more than several metres thick; as such, they represent relatively stable lake levels, the beach thickness indicating the limited range of the water level fluctuations (Bluck 2010). As wave-cut ravinement surfaces are noted at some locations cut into the basement, it is clear that water levels were stable for considerable periods of time. The alongshore georadar profile (Fig. 9a) terminates downdrift about 8 m from a steep beach front tread which represent the termination of a bluff spit. The section therefore shows the internal structure of the strandline near the termination. A 4 m thickness of beach gravel overlays a bedrock abrasion surface.

In the section between 0 and 30 m within the updrift portion of the image, reflectors dip downdrift at angles of up to 138 as series of contiguous sedimentary packages each a few metres in length. Taken together, they amalgamate and represent a shore face prograding to the southeast (1288). The cause of the significant multiple at 30 m may represent the presence of a large boulder in the section, although there is no distinct radar parabola. In the downdrift portion, updrift dipping reflectors are evident at a distance of 32–36 m and again at 50 – 52 m, the latter location being towards the termination of the spit. Reflectors between 40 and 50 m are generally dipping downdrift. Both these basal sets of reflectors are overlain by a final indistinct fill sequence that varies between 1 m and a maximum thickness of 4 m at the 32 m distance marker. The basal sequences are interpreted to represent a propensity for the spit to prograde in the direction of the prevailing wave action but this is interrupted by occasional wave trains emanating from an opposing quarter which results in local beach scour, reworking and reflectors dipping against the regional trend. The cross-shore profile (Fig. 9b) extends across the full 54 m width of the tread giving a c. 4 m thickness of gravel above a near horizontal abrasion surface. Extended lakewards there are two superimposed units both with 5–88 dipping reflectors which can be traced as semicontinuous surfaces over distances of many metres, but locally reflectors are cross-cutting with lakeward dips of up to 188. However, dips are in the opposite direction at the 6 and 38 m distance marker with two trough-shaped reflectors which may reflect the development of a shoreface bar prograding landwards. Taken together, both images therefore reflect variable short-term shingle drift directions locally within a system that is otherwise prograding towards 1288 in the longer term.

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Fig. 9. Georadar images of gravel shoreline deposits on the southern margin of the Chuja Basin: (a) alongshore profile and (b) cross-shore profile (see text for details).

Wind-wave climate Typical wave heights (Fig. 10) and wavelengths for the range of palaeofetches commensurate with the known range in heights of strandlines for the Chuja and Kuray basins were determined by the formulae of Bogoslovskiy (1960) and of CERC (1984) for deep-water waves. The results using the two methods were comparable. Wave height is defined (Bogoslovskiy 1960): h ¼ 0:020W 5=4 D1=3

(1)

where h is significant wave height (m), W is the wind speed (m s21) and D is the length of wave fetch (kilometres). The maximum wave fetch in both the Chuja Basin (c. 65 km) and Kuray Basin (c. 25 km) is from the west to the east. There is no information on the temporal distribution of wind speeds during the Quaternary but it is evident from Figure 10 that substantial wave heights (he  5 m) with wavelengths of 100– 125 m and with periods (T ) of

between 8 and 9 s could develop for wind speeds of 30 m s21, which can develop after only 4 –5 h (CERC 1984). Sustained winds of 20 m s21 today are associated with katabatic winds and storms on the Iceland icecap as well as within European Alpine valleys (Obleitner 1994; Oerlemans et al. 1999), so the persistence of high-speed winds around the Altai Quaternary ice-mass is not unreasonable. Komar & Miller (1973, 1975) calculated the orbital threshold velocity needed to entrain sediment as coarse as 0.10 m beneath deep-water Airy wind waves. Consideration of the wave height and wave period noted above and application of the function of Komar (1976) demonstrates that waves in water depths of 30–50 m could readily entrain grain diameters of 0.01–0.10 m. The method of Komar does not consider coarser grain sizes such as boulders. Rather, Lorang (2000) developed a swash action theoretical entrainment function developed for steep (.78) boulder beaches that compared well with field data. Exploration of the wave climate parameter space for the applicable ranges of the coefficient values reported by Lorang

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Fig. 10. Variation in predicted significant wave height as wind speed and fetch is adjusted according to the altitude of the lake levels: (a) Chuja Basin and (b) Kuray Basin.

(2000) and Hartstein & Dickinson (2006) showed that, for steep beach slopes (7 –208), 5 m high waves could readily entrain 0.4 m blocks of quartzdensity rock. Coarser blocks would tend to become embedded. It is therefore evident that, given the long fetches, the wind-wave climate in the Chuja and Kuray basins was capable of eroding, transporting and depositing coarse sediment alongshore to contribute to the construction of strandline landforms.

Pocket beach crestline model In this section, the sediment transport of gravel on the pocket beaches is modelled to determine whether the modelling approach can reproduce the alongshore gradient as recorded for the tilted shorelines. The dynamics of gravel beaches differ significantly from sandy beaches and are poorly known, thus a simple modelling approach is advised to extract basic dynamic behaviour (Buscombe & Masselink 2006). On gravel beaches, the voids between the coarse gravel allow the water mass brought in by the wave to drain away through the beach. The waves move the gravel up the slope,

but there is little return flow to transport the gravel down again. The downslope motion is largely governed by gravity-driven rolling. As a result, gravel beaches are considerably steeper than sandy beaches as noted above. When there is sufficient sediment available, the height of the beach crest is directly related to the wave energy of the highest waves reaching the beach (Cooper et al. 1996; Lorang 2002). Lorang (2002) derived two models predicting the crest height of gravel beaches. In the first model, Lorang calculates the wave dynamics and the resulting sediment transport. The second model correlates the measured wave energy at breaking with observed crest heights. This latter model gave a better result than the more complex first model, given the few test data available. Lorang (2002) rightly questions the general validity of this model for coastal beaches subject to significant tidal amplitude. For lakeside beaches, however, this latter model does make sense from an energy perspective. Even although the beaches are steep, there is little reflection of the breaking waves. Breaking waves are very dissipative in general, but the porosity of gravel beaches further reduces wave reflection.

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The breaking waves will cause some onshore– offshore flow, but the amount of energy exchange involved is limited. For steep gravel beaches, the width of the surf zone is therefore relatively narrow and the roughness of the gravel bed limits the speed of the swash and back-wash, limiting sediment motion. In the case of pocket beaches, the geometry of the coastal embayment further reduces the potential of wave driven flows. Consequently, most energy dissipation occurs by the movement of sediment up the slope, turning potential and kinetic energy of the water waves into potential energy of the sediment. Hence the formula proposed by Lorang (2002) that balances wave energy with potential energy of the sediment is perfectly in order for steep pocket beaches consisting of gravel: mgHc ¼

rg 2 h 8 b

(2)

where Hc is the crest height of the gravel beach, hb the wave height at breaking, r the density of water and m the mass of the beach per unit area. The model was implemented with values of m appropriate for 5 –10 cm beach gravel ranging between 50 and 100 kg m22 (Day 1997). As the elevation of most Kuray–Chuja pocket beach shorelines implies a deep lake, all models were run with an offshore water depth of 30 m to accommodate deepwater waves. As equation 2 assumes shore-normal waves, a term is added to correct for the angle between the breaking waves and the local contour: Hc ¼

r 2 h cos(ac  ab ) 8m b

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typical beach shape with a beach length of 100 m (Fig. 11c, d) at a 458 angle. This wave would be achieved by a wind speed of 11 m s21 over a fetch of about 60 km. Such a wave condition gives a difference in the height of the beach crest of just over 6 m, giving a slope of about 0.06. A similar wave on a 50% longer beach of 150 m (Fig. 11e, f ) only marginally increases the crest level difference to 7 m, resulting in a beach slope of just over 0.04. When these waves approach a more symmetrical beach with typically steeper slopes at the side than in the middle, the crest differences reduce with the lowest point reached about a third of the way downdrift (not illustrated). With higher wave energy at the boundary, much higher beach crest variations can be achieved up to slopes more then 10 m for every 100 m. One of the observed beaches had a length of about 700 m on which a beach slope of almost 0.075 was observed. Such a steep gradient can still be achieved in the model, but only using more extreme conditions. The sediment mass is reduced to 70 kg m22, which is still a realistic value for uniformly sized gravel. With wind speeds up to a 16 ms21 gale force wind, wave heights of 6.2 m are generated. The angle between the wave direction and the beach orientation is increased to 608. The results (not illustrated) show a beach crest height difference over 50 m over a length of 700 m but, unlike the field example noted above which has a planar surface, the modelled beach has a parabolic surface elevation profile.

(3)

where ac is the normal of the beach contour and ab is the direction of the wave at breaking. As the beach crest height depends directly on the wave height at breaking, alongshore gradients in the wave height at breaking result in an alongshore gradient in the beach crest levels. The general topography of the embayment shown in Figure 6a was derived from the Advanced Spacebourne Thermal Emission and Reflection (ASTER) Global Digital Elevation Model (ASTGTM) with a spatial resolution of 15 m and simplified to provide the correct ranges of altitudes with symmetric or asymmetric beach plan profiles as in Figure 11a, b. To accommodate the natural range of pocket beach geometries, the modelled beach lengths were varied between 75 and 700 m with headlands extending from the bayhead by 50 –300 m. Within the model, waves of up to 6.2 m could be generated for the given fetches for wind speeds of 17 m s21, consistent with the calculations presented in ‘Wind-wave climate’. The first test reported here as an example used a 6.1 s period wave with 2.5 m height, approaching a

Discussion The Kuray –Chuja horizontal shorelines in any vertical succession are developed in pebble to coarser gravel and, as such, are not much subject to subaerial weathering and show little erasure through soil creep. Their consistent ‘crisp’ form, at any vertical succession, seems to indicate that they are all of a similar and recent age. If the shorelines dated from distinctively different epochs, one would expect differences in the degree of rounding of treads on finer sediment slopes and erasure on active soil creep slopes. In a similar vein, it would seem that slow rises in lake level would result in smoothing of steps through wave action. Although the staircase of strandlines may represent both periods of both rising and falling lake level, it therefore seems more likely that the suite of faint horizontal shorelines represents a steadily falling lake level wherein, if any rises occurred, they did not inundate higher level treads. Major rises could occur over long timespans by changes in the climatic water balance and through changes in the drainage properties of the ice dam. However, annual fluctuations probably also

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Fig. 11. Examples of generalized bathymetry of (a) a symmetrical embayment and (b) a skewed embayment. Wave approach from the north and east (right-hand side). (c) Effects of topography on wave refraction as waves cross the skewed topography and (d) the resultant beach crest variations along a beach of 100 m length. (e) Effects of topography on wave refraction as waves cross the skewed topography and (f ) the resultant beach crest variations along a beach of 150 m length. Beach crest height is relative to still water level (0 m) with 230 m offshore water depth.

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occurred as indexed by the thickness of the strandline sediment fill (Bluck 2010). As has been suggested for Lake Missoula (Alt 2001), the lake level probably rose rapidly during the summer snowmelt period preventing waves from carving distinctive shorelines. During the autumn, the lake level would stabilize as snowmelt stopped and storm waves could cut the strandlines. During the winter, if there was no continuous ice cover, there would have been a heavy pack ice that would dampen wave action and a well-developed shore-attached ice body, especially within shoreline embayments. Nevertheless, the orientation of the spits, downdrift fining and the stratigraphy shown in the georadar profiles show that alongshore drift was usually to the southeast in accordance with the prevailing winds; sometimes, however (GPR data), this direction was reversed when subdominant winds blew from opposing quarters. The presence of extensive horizontal strandlines around both the Kuray and Chuja basins imply that there was no large differential isostatic rebound (Teller 2001) across the basins during the upper Quaternary lifetime of the lakes. The tilted shorelines are largely due to tectonic motion and are believed to be Middle Pleistocene (Zol’nikov et al. 2010). However, it is evident from the modelling studies that the effect of alongshore drift of sediment on pocket beach crest height will also need to be considered in many cases where tectonic tilting is suspected. The presence of the large wedges of coarse sediment with offshore slopes of more than 208 and thickness of 20 m cannot be accounted for by rapid regression or transgression. Rather prolonged lake stillstands are indicated which allowed accumulation of these bodies; a conclusion reinforced by the presence of substantial ravinement surfaces cut into bedrock locally at the same elevations as the accumulated sediments. Stillstand allows for cutting of local bedrock to provide the coarsest beach material, whereas transgression would have left the shorelines below the wave base and blanketed by finer sediments transported from offshore (which is not the case). Rapid regression is also indicated; otherwise, the steep shore faces would have been degraded by wave cutting. The well-preserved shore faces are most evident where there are stacked spits which imply that the major beaches ‘young’ from high altitude to low altitude and represent a series of sudden drops in lake levels. It is intriguing to question whether the individual levels of (some) spits might correlate with individual catastrophic flood events as water levels fell rapidly when the ice-dam failed. Three flood events are recorded (Carling et al. 2002), the largest of which would require the lake to be at maximum elevation (Carling et al. 2010). However, the correlation of lake levels and

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megafloods must await further shoreline dating control. The shoreline modelling results show that it is feasible to achieve alongshore gradients in the beach crest height of gravel beaches, using the mechanism of sheltering and refraction/refraction of waves around headlands, using ranges of beach mass and wave height and angles that are realistic. The alongshore gradients in the breaking height of the waves has a quadratic impact on the crest height of the gravel beaches. In the model, the alongshore distribution of the sediment size has been assumed to be uniform. In reality the variation in wave energy along the beach and the obliqueness of the waves will cause a grading of sediment, resulting in spatial varying beach material sorting as observed in the field. However, insufficient information is available to determine the effect of such variations on the beach morphology. To achieve the extreme gradient of 7.5% for the 700 m long strandline, a significantly stronger wind is required than on the other beaches. This is seemingly in contradiction to neighbouring shorter beaches and can be explained by a more sheltered position of the smaller beaches by headlands or islands further upwind. This sheltering reduces the wave height under the extreme winds. The offshore wave height for the smaller beaches can be lower under the same wind climate as that which influenced the larger beach. The modelled alongshore gradient is distinctly sinusoidal, however, whereas the field example has a planar surface. In this latter respect it might be concluded that the alongshore regional slope of the majority of the larger strandlines is due to tectonic tilting as none exhibit sinusoidal profiles. However, the effect of alongshore systematic variation in beach elevation due to sediment transport processes should be accounted for in other examples of tilted shorelines worldwide.

Conclusions Glacial lake Kuray –Chuja occupied the Kuray and Chuja Basins during the Quaternary Period and left distinct shoreline features around the basin margins between 1600 and 2100 m. These features include erosional, erosional-accumulative and accumulative strandlines; the latter including spits, tombolos, barrier beaches and offshore bars. The consistent orientation of spits indicates a persistent palaeowind direction. The width of ravinement surfaces cut within colluvium or bedrock beneath strandlines demonstrates that lake levels, although variable, were often constant for sufficient periods of time to erode bedrock and to deposit up to 10 m thicknesses of beach deposits. Beach stratification and radar reflectors record small changes in lake

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levels which can be ascribed to seasonal effects and short-term beach accretion or erosion due to relatively short-term changes in wind direction. The ‘crisp’ nature of the strandlines suggests that lake levels were generally regressive during the formation of the preserved shoreline sequences. Some pocket beach strandlines are tilted, which can reasonably be ascribed to tectonics. However, the direction of alongshore drift is in the same direction as the presumed tectonic uplift. An alongshore drift model shows that the palaeowave regimen was sufficient to augment the alongshore gradient of the beaches in a sediment-starved system of closed coastal cells. Consequently, the drift process needs to be accounted for in other similar palaeolake systems where alongshore beach gradients may be initially related to tectonics or isostacy. Data were collected during expeditions between 1994 and 2006 which were sponsored by the German Research Foundation, the Royal Society of London and the Royal Geographic Society of London Gilchrist Expedition Award (1996). A. Kirkbride and R. Robbins are thanked for their assistance with surveying. The preliminary GPR results contributed to the undergraduate dissertation of R. Robbins. U. Moody is thanked for provision of two C14 strandline dates. Thoughtful reviews by P. Jewell and I. P. Martini helped the authors to focus the arguments on key issues.

References Adams, K. D. 1997. Late Quaternary pluvial history, isostatic rebound, and active faulting in the Lake Lhontan basin, Nevada and California. Unpublished PhD thesis, University of Nevada, Reno. Adams, K. D. & Wesnousky, S. G. 1998. Shoreline processes and the age of the Lake Lahontan highstand in the Jessup embayment, Nevada. Geological Society of America Bulletin, 110, 1318–1332. Alt, D. 2001. Missoula and its Humungous Flood. Montana Press Company, Missoula, MT. Babichev, A. V., Novikov, I. S., Polyanskii, O. P. & Korobeinikov, S. N. 2009. Modeling Cenozoic crustal deformation in Gorny Altai. Russian Geology and Geophysics, 50, 137–151. Back, S. & Strecker, M. R. 1998. Asymmetric late Pleistocene glaciations in the North basin of the Baikal Rift, Russia. Journal of the Geological Society of London, 155, 61–69. Backstrom, J. T., Jackson, D. W. T. & Cooper, J. A. G. 2008. Shoreface morphodynamics of a high-energy, steep and geologically constrained shoreline segment in Northern Ireland. Marine Geology, 257, 84– 106. Baker, V. R., Benito, G. & Rudoy, A. N. 1993. Palaeohydrology of late Pleistocene Superflooding, Altay Mountains, Siberia. Science, 259, 348 –350. Beres, M., Green, A., Huggenberger, P. & Horstmeyer, H. 1995. Mapping the architecture of glaciofluvial sediments with three-dimensional georadar. Geology, 23, 1087– 1090.

Bills, B. G. & May, G. M. 1987. Lake Bonneville: Constraints on lithospheric thickness and upper mantle viscosity from isostatic warping of Bonneville, Provo, and Gilbert stage strandlines. Journal of Geophysical Research, 92, 11 493– 11 508. Bluck, B. 1999. Clast assembling, bed-forms and structure in gravel beaches. Transactions of the Royal Society of Edinburgh: Earth Science, 89, 291–323. Bluck, B. 2010. Structure of gravel beaches and their relationship to tidal range. Sedimentology, doi: 10.1111/j.1365-3091.2010.01192.x. Bogachkin, B. M. 1981. Cenozoic History of Tectonic Evolution of the Gorny Altai, Nauk, Moscow (in Russian). Bogoslovskiy, B. B. 1960. Limnology. Moscow State University, Moscow (in Russian). Borodavko, P. S. 2003. Evolution of the Chuja– Kurai lake system during the Late Neopleistocene. Extended abstract of Candadian Science Dissertation (Geography), Tomsk State University, Tomsk. Bowman, D., Guille´n, J., Lo´pez, L. & Pellegrino, V. 2009. Planview geometry and morphological characteristics of pocket beaches on the Catalan coast (Spain). Geomorphology, 108, 191–199. Buscombe, D. & Masselink, G. 2006. Concepts in gravel beach dynamics. Earth-Science Reviews, 79, 33–52. Buslov, M. M., Zykin, V. S., Novikov, I. S. & Delvaux, D. 1999. The Cenozoic history of the Chuja Depression (Gonry Altai): Structures and geodynamics. Geologiya i Geofizika, 40-12, 1720–1736 (in Russian). Butvilovsky, V. V. 1993. Last Glaciation and Holocene Palaeogeography of the Altai: Cataclysmic Model. Tomsk State University, Tomsk (in Russian). Carling, P. A., Kirkbride, A. D., Parnachov, S., Borodavko, P. S. & Berger, G. W. 2002. Late Quaternary catastrophic flooding in the Altai Mountains of south-central Siberia: a synoptic overview and an introduction to flood deposits sedimentology. In: Martini, I. P., Garzon, G. & Baker, V. R. (eds) Flood and Megaflood Processes and Deposits: Recent and Ancient Examples. Special Publication of the International Association of Sedimentologists, 32, 17–35. Carling, P. A., Martini, P., Herget, J., Borodavko, P. & Parnachov, S. 2009. Megaflood sedimentary valley fill: Altai Mountains, Siberia. In: Burr, D. M., Carling, P. A. & Baker, V. R. (eds) Megaflooding on Earth and Mars, CUP, Cambridge, 243–264. Carling, P. A., Villanueva, I., Herget, J., Wright, N., Borodavko, P. & Morvan, H. 2010. Unsteady 1D and 2D hydraulic models with ice dam break for Quaternary megaflood, Altai Mountains, southern Siberia. Global and Planetary Change, 70, 24– 34. Carter, R. W. G. & Orford, J. D. 1993. The morphodynamics of coarse clastic beaches and barriers: a short- and long-term perspective. Journal of Coastal Research, Special Issue, 15, 158– 179. CERC (COASTAL ENGINEERING RESEARCH CENTER) 1984. Shore Protection Manual, 4th edn. 2 Volumes. Department of the Army, Waterways Experiment Station, Corps of Engineers, Coastal Engineering Research Center, Vicksburg, United States Government Printing Office.

PALAEOSHORELINES OF GLACIAL LAKE Cooper, N., Workman, M. & Coates, T. T. 1996. Field Monitoring of Shingle Beaches at Shoreham and Elmer, West Sussex. HR Wallingford Report, TR 8. Crittenden, M. D. 1963. New data on the isostatic deformation of Lake Bonneville. U.S. Geological Survey Professional Paper 454, E1–E31. Dashtgard, S. E., White, R. O., Butler, K. E. & Gingras, M. K. 2007. Effects of relative sea level change on the depositional character of an embayed beach, Bay of Fundy, Canada. Marine Geology, 239, 143–161. Davies, A. J. & Johnson, M. P. 2006. Coastline configuration disrupts the effects of large-scale climatic forcing, leading to divergent temporal trends in wave exposure. Estuarine, Coastal and Shelf Science, 69, 643–648. Day, R. W. 1997. Study of gravel deposit created by wave action. Environmental and Engineering Geoscience, 3, 443–448. Dupre´, W. R. 1984. Reconstruction of paleo-wave conditions during the Late Pleistocene from marine terrace deposits, Monterey Bay, California. Marine Geology, 60, 435– 454. Fredsoe, J. & Deigaard, R. 1995. Mechanics of Coastal Sediment Transport. Advanced Series on Ocean Engineering, 3rd edn. World Scientific Publishing, Singapore. Gilbert, G. K. 1890. Lake Bonneville. Monographs of the United States Geological Survey, US Government Print Office, Vol. 1, Washington. Gillespie, A. R., Burke, R. M., Komatsu, G. & Bayasgalan, A. 2008. Late Pleistocene glaciers in Darhad Basin, northern Mongolia. Quaternary Research, 69, 169–187. Grosswald, M. G. & Rudoy, A. N. 1996. Quaternary glacier-dammed lakes in the mountains of Siberia. Polar Geography, 20, 180– 198. Hartstein, N. D. & Dickinson, W. W. 2006. Wave energy and clast transport in eastern Tasman Bay, New Zealand. Earth Surface Processes and Landforms, 31, 703– 714. Hasselmann, K., Barnett, T. P. et al. 1973. Measurements of wind-wave growth and swell decay during the Joint North Sea Wave Project (JONSWAP). Deutsche Hydrographische Zeitschrift Supplement A 8. Herget, J. 2005. Reconstruction of Ice-dammed Lake Outburst Floods in the Altai Mountains, Siberia. Geological Society of America, Special Paper 386, Geological Society of America, Boulder. Jewell, P. W. 2007. Morphology and paleoclimatic significance of Pleistocene Lake Bonneville spits. Quaternary Research, 68, 421– 430. Komar, P. D. 1976. Beach Processes and Sedimentation. Prentice-Hall, Englewood Cliffs, NJ. Komar, P. D. & Miller, M. C. 1973. The threshold of sediment movement under oscillatory water waves. Journal of Sedimentary Research, 43, 1101–1110. Komar, P. D. & Miller, M. C. 1975. On the comparison between the threshold of sediment motion under waves and unidirectional currents with a discussion of the practical evaluation of the threshold: A reply. Journal of Sedimentary Petrology, 45, 362–367. Komatsu, G., Brantingham, P. J., Olsen, J. W. & Baker, V. R. 2001. Paleoshoreline geomorphology

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of Bo¨o¨n Tsagaan Nuur, Tsagaaan Nuur and Orog Nuur: the valley of Lakes, Mongolia. Geomorphology, 39, 83– 89. Komatsu, G., Arzhannikov, S. G., Gillespie, A. R., Burke, R. M., Miyamoto, H. & Baker, V. R. 2009. Quaternary paleolake formation and cataclysmic flooding along the upper Yenisei River. Geomorphology, 104, 143–164. Krinner, G., Mangerud, J., Jakobsson, M., Crucifix, M., Ritz, C. & Svendson, J. 2004. Enhanced ice sheet growth in Eurasia owing to adjacent ice-dammed lakes. Nature, 427, 429 –432. Krist, F. & Schaetz, R. J. 2001. Paleowind (11,000 BP) directions derived from lake spits in Northern Michigan. Geomorphology, 38, 1– 18. Leontiev, O. K. 1963. Kratky Kurs Morskoy Geologyi. Moscow University Press, Moscow. Lorang, M. S. 2000. Predicting threshold entrainment mass for a boulder beach. Journal of Coastal Research, 16, 432 –445. Lorang, M. S. 2002. Predicting the crest height of a gravel beach. Geomorphology, 48, 87–101. Obleitner, F. 1994. Climatological features of glacier ¨ tztal Alps, and valley winds at the Hintereisferner (O Austria). Theoretical and Applied Climatology, 49, 225– 239. Obolkin, V. A. & Potemkin, V. L. 2006. The impact of large lakes on climate in the past: a possible scenario for Lake Baikal. Hydrobiologia, 568, 249–252. Oerlemans, J., Bjo¨rnsson, H., Kuhn, M., Obleitner, F., Palsson, F., Smeets, C. J. P. P., Vugts, H. F. & De Wolde, J. 1999. Glacio-meteorological investigation on Vatnajo¨kull, Iceland, summer 1996: an overview. Boundary-layer Meteorology, 92, 3–21. Ota, T., Utsunomiya, A., Uchio, Y., Isozaki, Y., Buslov, M. M., Ishikawa, A., Maruyama, S., Kitajima, K., Kaneko, Y., Yamamoto, H. & Katayama, I. 2007. Geology of the Gorny Altai subduction-accretion complex, southern Siberia: Tectonic evolution of an Ediacaran-Cambrian intraoceanic arc-trench system. Journal of Asian Earth Sciences, doi: 10.1016/j.jseaes.2007.03.001. Reuther, A. U., Herget, J., Ivy-Ochs, S., Borodavko, P., Kubik, P. W. & Heine, K. 2006. Constraining the timing of the most recent cataclysmic flood events from ice-dammed lakes in the Russian Altai Mountains, Siberia, using cosmogenic in site 10Be. Geology, 34, 913–916. Rogozhin, E. A., Ovsyuchencko, A. N., Marakhanov, A. V. & Ushanova, E. A. 2007. Tectonic setting and geological manifestations of the 2003 Altai earthquake. Geotectonics, 41, 87–101. Rudoy, A. N. 1988. The regimen of intermontane ice-dammed lakes of South Siberia. Data Glacial Research, Academy of Sciences, USSR, 61, 36–42 (in Russian). Rudoy, A. N. 2002. Glacier-dammed lakes and geological work of glacial superfloods in the Late Pleistocene, Southern Siberia, Altai Mountains. Quaternary International, 87, 119 –140. Sheinkmann, V. C. 2002. The diagnostic age of glacial sediments of the Altai, tested using sections of the Dead Sea. In: Data of Glaciological Studies. Institute of Geography of the Russian Academy of Science,

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Palaeozoic glacial depositional environments of SW Saudi Arabia: process and product MARTIN KELLER1,2*, MATTHIAS HINDERER3, HUSSAIN AL-AJMI4 & RANDOLF RAUSCH2,3 1

Geozentrum Nordbayern, Abt. Krustendynamik, Schlossgarten 5, D-91054 Erlangen, Germany

2

Gesellschaft fu¨r Technische Zusammenarbiet-International Services (GTZ-IS), P.O. Box 2730, 11461 Riyadh, Saudi Arabia 3

Institut fu¨r Angewandte Geowissenschaften, TU Darmstadt, Schnittspahnstrasse 9, D-64287 Darmstadt, Germany

4

Ministry of Water and Electricity, Water Resources Development Department, Riyadh, Saudi Arabia *Corresponding author (e-mail: [email protected]) Abstract: Saudi Arabia is one of the few places in Gondwana where deposits of two Palaeozoic glaciations are well preserved. They were formed under similar conditions in a tectonically stable shelf setting. We studied the sedimentary facies and architectural geometries of Upper Ordovician and Permo-Carboniferous glaciogenic deposits in the Wajid Sandstone and developed a genetic model for both units. Common features and of both glacial episodes are (a) incision of glacial (tunnel?) valleys, (b) subsequent valley fill by proglacial deposits in front of an oscillating, polythermal ice shield; (c) intra-formational erosional events through repeated ice advance, (d) widespread and large-scale soft deformation due to glacial surge during deglaciation, and (e) marine transgression most probably due to eustatic sea-level rise following deglaciation. The general patterns and observations fit well with recently published observations on Upper Ordovician deposits in northern Africa, pointing to closely coupled glaciological processes of the North African– Arabian ice shield. Although the same general patterns apply for the Permo-Carboniferous glaciation, sedimentary styles and petrographical properties differ, probably due to a more heterogeneous ice-flow pattern controlled by a more pronounced topography after the Hercynian tectonic event.

During the Palaeozoic, global cooling led to several glacial episodes which are well constrained by glaciogenic deposits and glacial striations on the bedrock in northern and western Gondwana which at that time was in a southern polar position (Semtner & Klitzsch 1994). Recently, marine isotope records and studies of global eustacy gave additional insight into the timing, cyclicity and magnitude of the glacial periods during the Late Ordovician to early Silurian (Sutcliffe et al. 2000) and the Permo-Carboniferous, respectively. Although glaciogenic deposits are attractive targets for hydrocarbon reservoirs in North Africa and Arabia, detailed studies on the sedimentary facies and architecture of these rocks have been presented only recently by Le Heron et al. (2004, 2005, 2009, 2010) and Ghienne (2003) for northern Africa and by Gutierrez-Marco et al. (2010) for northern Spain. From these studies it becomes clear that the Ordovician glaciation was not a single event but that there were several cycles of glaciation and deglaciation, mainly during the Hirnantian. Evidence from carbonates of the northern hemisphere

indicates that the Ordovician glaciation might have started as early as the Katian and that it may have lasted into the Silurian (Saltzman & Young 2005; Buggisch et al. 2010; Keller & Lehnert 2010). Similarly, the Permo-Carboniferous glaciation is a succession of events with several glacial and interglacial stages. The exact spatial and temporal extent of each of these pulses is still a matter of debate. On the Arabian Platform, both the Late Ordovician and the Permo-Carboniferous glacial episodes are documented through glacial deposits. From seismic surveys (McGillivray & Al-Husseini 1992; Evans et al. 1997), valleys are known that cut down into Cambrian and earlier Ordovician strata up to 600 m. These successions provide the most important Palaeozoic productive oil reservoirs in central and northern Saudi Arabia. The genesis of these tunnel (?) valleys is still a matter of debate. Their systematic arrangement orthogonally to the former ice front in North Africa and Arabia, however, points to a subglacial or proglacial origin and a common origin and source area.

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 129–152. DOI: 10.1144/SP354.8 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Early summaries on glacial deposits in Arabia were given by McClure (1978, 1980), Al Laboun (1987) and McClure et al. (1988). Unequivocal evidence of Ordovician glacial sediments was first presented by Vaslet (1990) from the northern part of Saudi Arabia. Many questions remain with respect to the age and stratigraphic position of these mostly coarse-grained siliciclastic deposits, which cannot directly be dated through biostratigraphy because of the paucity of body or trace fossils. Recent palynological studies on the confining stratigraphic units, regional mapping and sequence stratigraphic correlations have greatly improved knowledge on the age of Palaeozoic glacial deposits on the Arabian Platform (Kruck & Thiele 1983; Kellogg et al. 1986; McClure et al. 1988; Vaslet et al. 1991; Stephenson & Filatoff 2000; Konert et al. 2001; Sharland et al. 2004; Stephenson 2004, 2008; Stephenson et al. 2007). This paper presents new observations on sedimentary facies and depositional architecture of outcrops in south-western Saudi Arabia, which belong to the siliciclastic succession of the Wajid Sandstone that spans almost the whole Palaeozoic from the Cambrian to the Middle Permian. Major unconformities within this succession are thought to be related to glacial episodes (Kellogg et al. 1986; Sharland et al. 2001, 2004) or to tectonic movements (Stump & Van der Eem 1996). From systematic field observations of facies associations, a genetic model for both glacial episodes is set up and compared with models recently proposed for other regions in northern Gondwana (Ghienne & Deynoux 1998; Ghienne 2003; Le Heron et al. 2004, 2005, 2006; Le Heron & Dowdeswell 2009).

Study area The study area is located in south-western Saudi Arabia (Fig. 1). The outcrops of the Wajid Sandstone cover an area of approximately 44 000 km2 from Sulayyil in the north to Najran in the south. To the east, the outcrops disappear beneath the sandy cover of the world’s largest sand desert, the Rub’ Al Khali. From there, equivalent strata are known from several wells (GTZ/Dornier Consulting 2009). Throughout the outcrop area, sediments of the Wajid Group are very unevenly distributed. Cambrian and Ordovician sediments are found in the north-western edge of the outcrop belt (Fig. 1), which elsewhere is dominated by the sediments of the Devonian –Mississippian Khusayyayn Formation. Presumed Permo-Carboniferous deposits have been mapped in the northern part of the area (Kellogg et al. 1986) as occurring in two broad channels (Figs 1 & 2). In addition, there is an isolated outcrop of Permo-Carboniferous strata in the SE (Fig. 1) at Bani Khatmah.

Stratigraphy of the Wajid Sandstone The Wajid Sandstone is a several hundred metres thick succession of sandstones and conglomerates that are sandwiched between the Precambrian basement of the Arabian Shield below and Permian sediments of the Khuff Formation above. The internal lithostratigraphical subdivision and the biostratigraphical assignment of the units within the Wajid Sandstone are still a matter of debate (Kellogg et al. 1986; Evans et al. 1991; Stump & Van der Eem 1995, 1996). We follow Kellogg et al. (1986) who subdivided the Wajid Group into formations: The Dibsiyah, Sanamah, Qalibah, Khusayyayn and Juwayl formations (Fig. 3). The Dibsiyah Formation is the oldest unit within the Wajid Group and of presumed Cambrian and Ordovician age (Fig. 3). This is assumed from a few Cruziana traces and from abundant Skolithos tubes (Tigillites), which are typical of Cambrian and Ordovician sedimentary environments. The Dibsiyah Formation mainly consists of sandstones and some conglomerates. Its lower part was deposited in marine tidal environments with adjacent braidplain deposits. The upper part consists of an alternation of tidal channel deposits and thick subtidal Skolithos-bearing sandstones and conglomeratic sandstones (Al Ajmi et al. 2008). The Sanamah Formation unconformably overlies the Dibsiyah Formation (Fig. 4a). The unconformity is represented by deep channels, several 100 m to several kilometres wide. Outside the channels, only a thin veneer of Sanamah sediments covers the adjacent highlands. The Sanamah Formation, part of the present study, consists of conglomerates, coarse-grained and medium-grained sandstones with characteristics attributable to a periglacial environment (Al Ajmi et al. 2008). In the outcrop area of the Wajid Group the Qalibah Formation is represented only by its lower member, the Qusaiba Member or Qusaiba Shale (Fig. 3; Evans et al. 1991; GTZ/Dornier Consulting 2009). In the outcrop area, there is a succession of fine-grained sandstones, siltstones and a few shales that are regarded as an equivalent of the subsurface Qusaiba Shale. The subsurface Qusaiba Shale has been correlated to similar deposits along the northern margin of Gondwana, where these sediments are attributed to the Silurian sea-level rise following the ultimate melting of the Hirnantian ice cap (Lu¨ning et al. 2000). No age data are available from these rocks in the outcrop area; their attribution to the Qusaiba Member and their interpretation as a post-glacial ablation unit is therefore solely based on lithology and the lithostratigraphical position. The Khusayyayn Formation is of assumed Devonian and probably Mississippian age (Fig. 3). It is a succession of medium- to coarse-grained, locally

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Fig. 1. Geological map of the Wajid Sandstone modified from Al-Husseini (2004), based on the lithostratigraphic subdivision of Kellogg et al. (1986). Black quadrangle is study area supplemented by some outcrops at Bani Khatmah c. 250 km SE from this area.

conglomeratic sandstones. The most conspicuous feature is large-scale cross-bedding interpreted as migrating subtidal sand dunes and tidally influenced (herringbone cross-stratification) horizons (Al Ajmi et al. 2008). The basal surface of the Khusayyayn

Formation is an unconformity that locally cuts down through the Qusaiba Shale into the Sanamah deposits. In contrast to Stump & Van der Eem (1995, 1996), the Khusayyayn Formation is interpreted as an entirely marine succession (Al Ajmi et al. 2008).

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Fig. 2. Enlargement of the geological map modified from Al-Husseini (2004) of the north-western study area with localities mentioned in the text (A, Jabal Ma’aleek; B, Jabal Atheer; C, Jabal Sanamah; D, Jibal Al Qahr; E, Jabal Overheat; F, Fard Al Ban; G, Bani Ruhayah). Note the NNW– SSE trending channel fills of the Juwayl Formation to the right (light green areas).

Yet another unconformity separates the Khusayyayn Formation from the overlying Juwayl Formation. Deposits of the Juwayl Formation are found in two elongate NW –SE trending channels (Kellogg et al. 1986). The depth of the channels is unknown; their width is of the order 20 km (Figs 1 & 2). In addition, Juwayl sediments are present in an isolated outcrop along the Tuwaiq Escarpment in the SE of the study area (Bani Khatmah, Fig. 1). In the following, only the Sanamah and Juwayl formations will be considered because they

contain glaciogenic to proglacial deposits of the Late Ordovician and Permo-Carboniferous glaciations, respectively.

The Sanamah Formation Distribution The outcrops of the Sanamah Formation are restricted to the western belt where it occurs in

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Age The age of the Sanamah Formation is not well constrained. No fossils have been reported from these strata. Within the underlying Dibsiyah Formation, large Cruziana sp. and abundant Skolithos pipe rock point to a Cambrian and Ordovician age; the glacial deposits of the Sanamah therefore might represent one episode of the Late Ordovician glacial cycle. Similarly, Evans et al. (1991) considered this interval to be the product of the Late Ordovician glaciation. This consideration was based on seismic lines which show striking similarities to seismic lines reflecting better-dated sequences from northern Saudi Arabia. In addition, the unconformity (sequence boundary) at the base of the Upper Ordovician deposits in the north could be traced into the Wajid subcrop area, thus adding in bracketing the age of the Sanamah deposits. In this paper, the presence of glacial features and their gross correlation to better-dated strata in the north is taken as evidence that the Sanamah Formation indeed represents one episode of the Late Ordovician glacial events (Fig. 3).

Thickness The channels at Jibal Al Qahr are hundreds of metres to probably 2 km wide (Fig. 2); a minimum depth of 80 m could be reconstructed. At Jabal Sanamah, erosion cut down into the lower part of the Dibsiyah Formation. As the upper Dibsiyah Formation in the vicinity is more than 100 m thick, the original depth of the channels must have been even greater. From the spatial relationships observed in the field, it is apparent that the channels locally might cut into the basement and that a minimum depth of the channels of more than 150 m has to be assumed.

The Juwayl Formation Fig. 3. Lithostratigraphic subdivision of the Wajid Group (‘Wajid Sandstone’) in SW Saudi Arabia (redrawn from Kellogg et al. 1986).

several isolated hills (Figs 1 & 2). The main outcrops are those of Jabal Sanamah, which is the type locality, and at Jibal Al Qahr. At Jibal Al Qahr (Fig. 2), the contact between the Dibsiyah Formation and the Sanamah Formation is exposed three-dimensionally in a system of canyons. Walking out the basal and lateral bounding surfaces, it becomes evident that the Sanamah Formation was deposited in large wide channels.

Distribution The Juwayl Formation is preserved in two distinct settings. In the north, two NW–SE trending largescale belts are present (Fig. 2). These belts have been mapped by Kellogg et al. (1986) who concluded that with a width of 20 km they represent the original depositional topography of broad channels. The channel structures are similar to those described from the Sanamah Formation and are cut into the underlying Khusayyayn Formation. In the SE, a well-bedded succession of periglacial sediments is present at Bani Khatmah, first described as glaciogenic deposits by McClure et al. (1988).

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Fig. 4. (a) Erosional unconformity of a palaeovalley incised into the shallow-marine Dibsiyah Formation and filled with glacial to proglacial deposits of the Sanamah Formation. (b) Sedimentary architecture of an Upper Ordovician palaeochannel fill showing low-angle prograding sediments of the basal conglomeratic unit S1. They are overlain by steeply downlapping clinoforms of the sandy unit S2. Clinoforms dip towards the south. (c) Low-angle cross-bedded lobes of conglomeratic unit S1 with erosional base (Sanamah Formation). (d) Low-angle cross-bedded medium- to coarse-grained sandstones with nests of pebbles in unit S2 (Sanamah Formation).

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Age The Juwayl Formation is of late Carboniferous and Permian age (Fig. 3). Its northern equivalent in Saudi Arabia, the Unayzah Formation, is better dated and the flora studied indicates a late Pennsylvanian and early Permian age (McClure 1980; El-Khayal & Wagner 1985; Al Laboun 1988). Due to the strongly erosional contact to the underlying strata and the differential onlap onto the postHercynian topography, the age assignments vary considerably between the different authors. Sediments of the Juwayl Formation have been briefly described by McClure (1980), Al Sharhan et al. (1993) and recently by Pollastro (2003), Melvin & Sprague (2006) and Melvin et al. (2010). Together with the interpretations from Oman (Osterloff et al. 2004), a large glacial lake across the southern Arabian Peninsula and adjacent north-eastern Africa has been reconstructed. To the north, this lake was bounded by a belt of aeolian sediments.

Thickness The thickest preserved succession of the channel fills is about 80 m thick. At Bani Ruhayah (Fig. 2), the upper part of the Juwayl Formation and the transition into the overlying Khuff Formation are exposed. They have a thickness there of 80 m. At Bani Khatmah (Fig. 1), about 50 m of the lake facies are exposed, which are erosionally overlain by the Lower Jurassic Tuwayq Mountain Formation. As the spatial relation between the different outcrops is not quite clear, and as the sediments were deposited across a tectonically induced rugged topography, thicknesses between several tens of metres and several hundreds of metres have to be assumed.

Lithofacies of the Sanamah Formation and the Juwayl Formation LF1: Conglomerate, clast supported This facies, only observed in the Sanamah Formation, contains abundant pebbles and cobbles; sorting is poor. The matrix is composed of medium- to coarse-grained sand. The majority of the clasts is of vein-quartz origin, many of the pebbles are rounded to well-rounded but of low sphericity. The conglomerates are mainly reddish coloured. The thickness of the individual beds varies considerably between 20 cm and 3 m.

LF2: Conglomerate, matrix supported The sediments of this facies have a sandy (and silty and clayey in the Juwayl Formation) matrix, in

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which pebbles and boulders are generally irregularly distributed; locally they are concentrated in nests and pavements (Figs 5c & 6b). Sorting is poor; the shape of the components is subangular to rounded. The colour of the matrix-supported conglomerates is grey to reddish. They form mostly structureless horizons and beds up to 6 m thick. In the Sanamah Formation, most of the clasts are of vein-quartz origin; some of them show striations and shatter marks. Locally, these strata contain slabs of the clast-supported conglomerates described above. In the Juwayl Formation, green to grey metagreywackes (Figs 5c & 7e), granite (Fig. 7d) and metamorphic rocks are among the pebbles and boulders. Clasts of the underlying sedimentary succession are present but relatively rare. Many of the clasts show glacial striations, the distinct shape of subglacially transported clasts, and shatter marks. Some of the horizons are structureless, whereas others are well bedded (Fig. 6a, b). According to their grain-size distribution, sedimentary features and clast properties, these sediments in the Juwayl Formation are, from a purely descriptive perspective, glacial diamictites.

LF3: Massive to cross-bedded coarse sandstone, conglomeratic These deposits are dominantly coarse-grained sandstones with a variable amount of pebbles. Clast shape is rounded to well rounded but of low sphericity. In the Sanamah Formation at Jabal Atheer, these reddish sandstones form individual graded layers 25 to 35 cm thick. They have a sheet-like to channelized geometry. Where beds are erosional, they incise up to 5 m into the underlying deposits. In the Juwayl Formation, these sandstones are white and show large tabular to 2D trough crossbedded structures. They are present in several metres thick, laterally continuous beds and show a persistent direction of foreset dipping to the South and SE.

LF4: Massive medium- to coarse-grained, sandstone, pebbly These rocks, restricted to the Sanamah Formation, form medium to thick horizons without apparent bedding. Only locally, a weak large-scale crossstratification can be observed. The pebbles are mainly randomly distributed throughout the sandy matrix but locally are concentrated in patches (Fig. 4d). In places the sandstones have an erosional base, in others they were deposited in channels some 10 m wide. Locally, slumping disrupted the beds and generated floating intraclasts. These rocks are

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Fig. 5. (a) Large-scale (see geologist to the right as scale), multistory ‘Sorbet’ facies at Fard al Ban (Juwayl Formation) with vertical channel wall (see arrow). (b) Thin-bedded, partly ripple cross-laminated sandstone facies of unit J3 of Juwayl Formation. (c) Thin-bedded, partly ripple cross-laminated sandstone facies of unit J3 of Juwayl Formation with strong, post depositional soft-sediment deformation. (d) Diamictite with dropstones and gravel pavements in unit J1 of Juwayl Formation. (e) Thin-bedded, partly ripple cross-laminated sandstone facies of unit J3 of Juwayl Formation with post-depositional local soft-sediment deformation.

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Fig. 6. (a) Diamictite with silty laminated beds and impacted by dropstones in unit J1 of Juwayl Formation. (b) Tabular, cross-bedded siltstone overlain by a gravel carpet. Weak erosional base points to ice scratching, probably by an oscillating ice margin. Unit J1 of Juwayl Formation. (c) Local thrust fold within diamictite most probably caused by scratching of an iceberg keel. Unit J1 of Juwayl Formation. (d) Local thrust fold within diamictite most probably caused by scratching of an iceberg keel. Unit J1 of Juwayl Formation.

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Fig. 7. (a) Downbending of clayey varve beds of unit J1 at the rim of a Permo-Carboniferous tunnel valley. (b) Silicified shales interpreted as varves (unit J1). (c) Slumped varve deposits with intercalated sandstone bed (unit J1). (d) Angular granitic block deposited as dropstone in varve deposits (unit J1). According to its geometry, this block is of supraglacial origin. (e) Striated greenish metagreywacke dropstone clast deposited in a silty matrix (diamictite, unit J1). According to its geometry, this rock is of subglacial origin.

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dominantly yellow to light brown and grey in colour.

LF5: Massive coarse-grained sandstone These sandstones are organized in irregular channel-like features with an erosional base and interfinger with LF6 and LF7 (Figs 5a & 8a –d). Locally, faint slab-like strongly concave internal cross-bedding is visible. Some of the erosive boundaries form vertical walls pointing to unusual environmental conditions. In addition, floating blocks with sizes up to several metres are present. These large slabs consist of the same material as the matrix they are associated with and of the medium- to coarse-grained pebbly sandstones described above. This lithofacies is only observed in the Juwayl Formation. The sediments are yellowish to brownish and occur in packages up to 30 m thick.

LF6: Cross-bedded medium- to coarse-grained sandstone, pebbly The sandstones are reddish to brownish, well sorted, with very little conglomeratic material or finergrained matrix. They show low-angle planar to trough cross-bedding and channelized geometries with erosional bases (Fig. 8a, c). Locally, climbing ripple cross-laminations have been observed. Packages of this facies reach thicknesses up to 50 m, but they are intimately associated and interfinger with the following LF7 facies (Fig. 5b). This lithofacies is observed in the Sanamah and Juwayl Formations.

LF7: Siltstones to fine-grained sandstones These sediments are thin-bedded; some of them show cross-bedding (Fig. 6a, b). They show strong and large-scale soft-sediment deformation, including truncated folds, dome-like and large-scale ball-and-pillow structures and contain outsized clasts (Fig. 6a, c, d). These structures are interpreted to result from slumping and loading. The siltstones and sandstones show intense red or locally yellow colours. This lithofacies is only observed in the Juwayl Formation.

LF8: Siltstones and shale The rocks of this lithofacies, only present in the Juwayl Formation, are variegated and occur as discrete horizons with thickness of the order centimetres to a decimetre thick. They also form distinct couplets of light and dark millimetrethick laminae (Fig. 7b). Stacks of these couplets may be several metres thick. In many places, these

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successions of stacked couplets have been silicified. Few outsized clasts up to boulder size are present (Fig. 7c, d). The siltstones and shales are widespread at the base of the channels and onlap their upper margins. Here, downbending of the strata can be observed: the deformed sediments of the Juwayl rest on undeformed horizontal sandstones of the underlying Khusayyayn Formation. In addition, soft-sediment deformation such as slumping occurs along the flanks of the channels (Fig. 7c). Undeformed fine-grained sediments drape the channels and their margins. There, the bedding is parallel to that of the underlying undeformed Khusayyayn Formation.

LF9: Fine- to coarse-grained sandstones, well-bedded These are present in thin- to medium-bedded units of white, red and grey colours. Slumping and dewatering structures are frequently observed in these sediments (Fig. 5d, e). Very rarely, bedding-plane parallel trace fossils have been found as well as current ripples. This lithofacies is observed in the Sanamah and Juwayl Formations.

LF10: Shale and ripple-bedded sandstone Shales are laminated to thin-bedded and multicoloured. They are intercalated with continuous thin beds of medium sandstones with wave ripples and mud drapes. Bedding-plane parallel feeder traces are common, and some vertical burrows have been observed. This lithofacies is only observed in the Juwayl Formation.

Facies associations, glaciogenic processes and depositional model Several of the described lithofacies are per se non-diagnostic of any particular environment: for the most part they record processes. However, when grouped into lithofacies associations, these associations characterize well-defined depositional environments. Three lithofacies associations (S1, S2, S3; Fig. 9a, b) have been established in the Sanamah Formation, and three in the Juwayl Formation (J1, J2, J3; Fig. 9c, d). Concomitantly, in the Sanamah Formation these facies associations also represent depositional sequences, separated by disconformities. By interpreting these facies associations, we develop a depositional and glaciological model for the Ordovician and PermoCarboniferous glacial cycles as preserved in southwest Saudi Arabia (Figs 10 & 11).

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Fig. 8. (a) Liquefaction of cross-bedded, channelized sandstones forms viscous grain flows resulting in lobe-like structures (‘Sorbet’ facies). Unit J2 of Juwayl Formation. (b) Large-scale multistory ‘Sorbet’ facies (arrows with ‘S’ show base of successive levels of ‘Sorbet’ facies at Fard al Ban). (c) Erosional base of ‘Sorbet’ facies. (d) Floating intraformational block of cross-bedded sandstones.

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Fig. 9. Schematic model of the sedimentary architecture of the fill of tunnel valleys and onlapping glacial to postglacial successions for the Late Ordovician (a, b) and the Permo-Carboniferous (c, d) glaciations in SW Saudi Arabia (Wajid Sandstone). S1–S3 and J1 –J3 correspond to the units described in the text. (a) and (c) are cross-sections through the valleys, (b) and (d) are sections parallel to the axis of the valleys (modified from Le Heron et al. 2004).

Massive to cross-bedded conglomerate facies association (Unit S1) This facies association consists of clastsupported conglomerates (LF1), matrix-supported

conglomerates (LF2) and coarse, conglomeratic sandstones (LF3; Fig. 4b) and make up the lower part of the valley fills (Fig. 9a). We interpret this as proximal successions of coarse outwash sediments near glacial meltwater outlets forming

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Fig. 10. Schematic depositional models for all six facies associations inspired by Le Heron et al. (2010). Sanamah Formation (Upper Ordovician). (a) Massive to cross-bedded conglomerate facies association (Unit S1). (b) Massive to cross-bedded sandstone facies association (Unit S2). (c) Siltstone to medium-grained sandstone facies association (Unit S3). Juwayl Formation (Permo-Carboniferous). (d) Diamictic siltstone and shale facies association (Unit J1). (e) Coarse-grained sandstone facies association (Unit J2). (f ) Well-bedded shale to sandstone facies association (Unit J3).

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Fig. 11. Glaciological model for the (a) Late Ordovician and (b) Permo-Carboniferous glaciation based on sedimentary facies and glaciogenic signatures in the outcrop area of the Wajid sandstone in SW Saudi Arabia. At the position of the studied outcrops, megascale coarsening and fining-upwards cycles and related facies association are indicated. Note that more glacial cycles may have been active, which are not preserved in the sedimentary record. See text for further explanation.

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temporal fan-like lobes, most probably in a sealocked, subaqueous environment and filling up drowned valleys during glacial recession (Fig. 10a). The lobes gently dip downvalley towards the SE (Fig. 9b). Geometries and thicknesses of these ice-proximal fans indicate ample accommodation space of several tens of metres; up to 80 m must be assumed. Strong turbulence in the flowing water together with discontinuous sediment supply is responsible for repeated erosional events and the frequent changes and transitions between water-lain sediments by sediment-laden, high-energy meltwaters and gravitationally induced mass transport reworking glacial tills. This is deduced from varying, mostly weak sorting and absent or weakly developed cross-stratification. No silt or clay is present in the sediments, pointing to a fractionation of the sediment load into underflow and overflow in the water column where fine particles become buoyant and are transported further downcurrent (Le Heron et al. 2004). Indications of direct glacier action by soft sediment deformation and pure glaciogenic deposits are lacking except for some small patches of till-like deposits at the bottom of the valleys. Non-cohesive intraclasts indicate that deposits were partly reworked and seasonally (?) frozen. We suggest that after subglacial erosion the valleys were subsequently filled up during glacial recession. We hypothesize that focused drainage in the valleys from a large continental ice sheet led to specifically high-current velocities and meltwater volumes, most probably preventing any deposition of pure till deposits. Episodic meltwater outbursts may also have played a role leading to some similarities to subaerial deposits of jo¨kulhlaups that have been described from modern Icelandic environments by Maizels (1991, 1993, 1997).

Massive to cross-bedded sandstone facies association (Unit S2) The facies association is composed of massive, pebbly, medium- to coarse-grained sandstones (LF4) and cross-bedded, pebbly, medium- to coarsegrained sandstones (LF6). This facies association is frequently organized in large clinoforms up to 50 m high and dipping at around 208 (Fig. 4b). They fill up the upper section of the valleys (Fig. 9a, b). The absence in many beds of internal bedding features points to depositions from sediment-laden flows and mass flows. We interpret this facies association as Gilbert-type deltas that prograded from a sandur plain into the water-filled valleys during glacial recession. In contrast to S1, which we interpret as subaqueous deposits directly deposited in front of meltwater outlets of the glacier, during deposition of S2 an extensive subaerial sandur plain established which

drained through the drowned valleys and rapidly filled them up (Fig. 10b). Similar architectures are known from Pleistocene tunnel valleys (Praeg 2003). The top of Unit S2 is erosional with striated and fluted surfaces, and is cut by wide channels up to 20 m deep filled with sediments of Unit S3. This erosional surface is interpreted as an indication of direct action of a re-advancing glacier over a frozen sandur or delta plain.

Siltstones to medium-grained sandstone facies association (Unit S3) This facies association constitutes massive to weakly cross-bedded sandstones, pebbly with intraclasts (LF4) and thin-bedded to cross-bedded siltand sandstones (LF7). LF4 is organized in channels with a width of about 10 m (LF4) which are stacked within LF7. Most of the sequence shows strong internal small-scale load casts as well as strong large-scale soft deformation including truncated folds, diapiric, dome-like and large-scale ball-andpillow structures. The general decrease of grain size in Unit S3, the preservation of silt, absence of mass-flow deposits despite some local slumps and fewer and lesspronounced internal erosional surfaces indicate less turbulent flow conditions and a decrease of transport energy when compared to Unit S1 and S2. Whereas S1 and S2 are restricted to the valleys, S3 shows widespread distribution (Fig. 9a). This succession is interpreted as a proglacial product of a subaerial, glaciofluvial plain, affected by one or repeated events of glacial shearing and/or loading. Following Le Heron et al. (2005), who found similar features in the Murzuq Basin in SW Libya, the soft-sediment deformations are interpreted as the result of deformation of unfrozen waterlogged sediments possibly by a relatively short-lived glacial re-advance (surge) during a fluctuating late stage of deglaciation (Fig. 10c). The top of Unit S3 is marked by a non-erosive, flat disconformity overlain by the Qusaibah Shale that records a major marine transgression following the Ordovician glaciation (Sharland et al. 2004). In our study area, the former subaerial sandur plain was flooded and fine-grained, marine clay and siltstones were deposited during landward stepping of the coastline.

Diamictic siltstone and shale facies association (Unit J1) This association shows matrix-supported conglomerates (LF2) alternating rapidly with laminated to thin-bedded siltstones and shales (LF8). These rhythmites are interpreted as deposits in an ice-distal lake setting (Fig. 10d). Strong silicification masks

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microstructures, hence we cannot decide whether they are subaqueous varve deposits or mud turbidites deposited from silty to very fine-grained milky meltwaters. At Bani Khatmah, rhythmites are interfingering with diamictites (Fig. 6a, b) pointing to an ice-proximal equivalent in a similar glaciolacustrine setting. Although a Permo-Carboniferous age is certain (McClure et al. 1988), no direct chronostratigraphic control exists and the facies association at Bani Khatma might represent another glacial stage and/or might be locked to another ice shield (see below). Nevertheless, both settings can be combined in terms of glaciolacustrine facies and processes. Outsized conglomeratic clasts and their arrangement as pebble carpets and nests found within the fine-grained lacustrine deposits at Bani Khatma can be best explained as ice-rafted debris (Fig. 6b, c). This interpretation is also supported by localized soft-sediment deformation structures, which are most probably formed by scratching of iceberg keels over soft sediment (Fig. 6c, d). Altogether a glaciolacustrine environment is assigned to this facies association similar to the interpretation of McClure et al. (1988). This interpretation is corroborated by observations by Pollastro (2003) and Osterloff et al. (2004). Apparently, a vast glacial lake covered much of south-eastern Saudi Arabia and adjacent Oman. As lake sediments are also found in Ethiopia (Bussert & Schrank 2007), this or a series of lakes probably extended across much vaster areas.

Coarse-grained sandstone facies association (Unit J2) This is composed of coarse-grained sandstone lithofacies (LF5 and LF6) in combination with finergrained sediments (LF7). The massive to weakly cross-bedded sandstones (LF5) show identical grain-size to the cross-bedded sandstone facies (LF6), similar intraclasts and both interfinger closely. It is most likely that facies LF5 was produced by localized, synsedimentary liquefaction of the surrounding cross-bedded sandstones and mainly mobilized by mass flows. Preservation of floating blocks of coarse sandstones and weak internal large-scale often steep trough-like crossbedding (Figs 5a & 8d) identify a particular type of resultant deposit named ‘Sorbet’ facies in this paper. The channelized cross-bedded sandstones (LF6) are interpreted as having been deposited in a large glaciofluvial plain (Fig. 10e). The meltwater streams were directed towards the pre-existing large-scale channels with their high accommodation space and rapid sediment accumulation. Towards the top, the finer-grained LF7 interfingers with LF6, possibly pointing to a more ice-distal glaciofluvial setting. The top of J2 shows widespread soft deformation similar to the top of the Sanamah Formation

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(S3), but less pronounced. Deformation is at smaller scale and does not reach as deep as during the Ordovician glacial cycle. Unit J2 is interpreted to represent a rapid progradation of glaciofluvial sediments into the (possibly) ice-dammed lake of J1. The sediments must have been exposed to a polythermal regime of repeated freezing and thawing causing extensive liquefaction and mobilization of the sands as mass flows. Although sand is non-cohesive, trough-like crossbedding and lobate structures as well as floating non-cohesive blocks point to high viscous, laminar flow. We hypothesize that these unusual features were caused by viscous flow of a water-ice-sand mixture. The high ice and water content as well as the necessary gradient for mass movement of the sands might be caused by locking to the vast glacial lake which, according to Pollastro (2003), was most probably triggered by short-term lake level fluctuations and/or glacial surge (Fig. 10e). We interpret the widespread soft deformation at the top as glacial shearing, most presumably during a glacial surge in the late stage of deglaciation. No tillite, however, is present.

Well-bedded shale to sandstone facies association (Unit J3) This facies association (LF3, LF10) represents a transgressive succession, most probably due to sealevel rise during deglaciation (Fig. 10f ). Above the basal, almost planar disconformity, the deposition of transgressive marine beds started under tidal control (LF10) and reduced sediment supply. Within tidal shallow-marine deposits of LF10, channels few metres deep and a few tens of metres wide are repeatedly observed towards the top of the succession (LF3); however, they become increasingly smaller and less frequent. Evidence for glacial processes is no longer visible in this succession, which terminates the Wajid Sandstone Group at the PreKhuff Unconformity (Sharland et al. 2001). Similar to the Qusaibah Shale above the Ordovician sediments, this can be related to a relatively fast eustatic, postglacial sea-level rise. Later on, in late Permian time, the first platform-wide marine transgression took place depositing the Khuff carbonates. This fundamental change of depositional style on the Arabian Platform was controlled by tectonic subsidence (Sharland et al. 2001).

Analysis and evaluation of selected structures as evidence of ancient glaciations Although typical till deposits are missing in southwest Saudi Arabia, much evidence of a glacial to

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proglacial setting of the deposits can be found. This includes specific lithofacies types described above, which can only be explained if sediments had been temporarily frozen (such as LF4, LF5). In addition, sand-gravel mixtures, massive to weakly cross-bedded without clay and silt, could be explained by unsteady, strongly sediment-laden currents since they typically occur by meltwater outbursts in proglacial settings (LF1, LF2, LF3). We describe some specific glaciogenic features in the Wajid Group of southwest Saudi Arabia in the following sections.

local obstruction associated with the development of adjacent soft-sediment deformation. All striations observed in the Sanamah and Juwayl formations were most likely formed on unconsolidated sediments. Soft-sediment striated surfaces have been observed in other African Palaeozoic units where they have been also interpreted as glacially induced (Deynoux & Ghienne 2004; Le Heron et al. 2005). Le Heron et al. (2005) interpret these features as formed by shearing along bedding planes of over-pressured sand packages, most probably by ice load.

Striations

Fluted surfaces

In both the Sanamah and Juwayl formations, striations have been observed on disconformable surfaces and bedding planes. In a single section (Jabal Ma’aleek, Sanamah Formation), five horizons have been observed in which the striae run parallel to those on the underlying and overlying horizons (Fig. 12e). In the vicinity of Jabal Atheer, such a horizon was also observed in the uppermost part of the non-glacial Dibsiyah Formation just beneath the unconformity with the Sanamah Formation. Other striations were observed near Bani Ruhayah in the upper part of the Juwayl Formation (Fig. 12c). Most of the striae are well preserved because they are diagenetically cemented by iron oxides. Striations in the Sanamah Formation show a very consistent direction between 1108 and 1308; in the Juwayl Formation, directions between 1308 and 1508 have been measured. In both formations, striations are more or less parallel to the axis of the channels in which the sediments were deposited. Locally in the Sanamah Fommation, two sets of highly divergent striations have been observed (Fig. 12a). Striations are characteristic of a glacial environment; some facts, however, indicate that these striations cannot have been created through direct glacier contact. In most known examples of true glacial striations, these striations are preserved in hard rocks at the base of the glacial succession. Magmatic and metamorphic rocks are suitable for the preservation of striae, but also quartz arenites and occasionally carbonate rocks. Secondly, in some horizons cross-cutting relationships have been observed between two generations of striations. This may be due to deviation of the ice movement by local obstructions, by changing movement direction of a single glaciers during its thinning and retreat or, if found extensively, formed by two glaciers advancing from different directions (Flint 1957; Benn & Evans 1998). There is no sufficient information on the distribution of divergent stations in the Sanamah Formation to confidently establish a specific cause except perhaps some

Within the Sanamah Formation there is an unconformity (S2/S3) along which well-developed iron crusts are preserved. This iron-cemented unconformity can be taken as a marker horizon and can be traced over large areas. On this surface, flute structures similar to those described by Le Heron et al. (2005) have locally been observed (Fig. 12b). Direction of the flutes coincides with those of the striations and show directions between 1108 and 1308.

Dropstones and boulder pavements Many sediments of the Juwayl Formation were deposited in a vast lake (Pollastro 2003). Within these fine-grained sediments, exotic clasts of pebble to boulder size are present. Magmatic rocks (granite), metamorphic rocks (especially metagreywackes) and sedimentary rocks of older Wajid Group units have been identified (McClure et al. 1988). Frequently, the clasts show faceting and striations, in particular the fine-grained pebbles of metagreywacke (Fig. 7e). They also deform the laminas of sediments they rest on (Fig. 6a). These clasts can be reasonably interpreted as ice-rafted dropstones. At Bani Khatmah, a succession of fine siltstones to sandstones containing dropstones are abruptly overlain by a clast-supported conglomerate whose clasts show striations (Fig. 6b). The accumulation of these boulders most likely represents remnants of a conglomeratic pavement associated with the oscillation of an ice front.

Local post-depositional sedimentary deformation At Bani Khatmah and at Jabal Overheat, the lacustrine sedimentary succession of the Juwayl Formation shows large-scale, post-depositional deformation structures consisting of folds and thrust faults (Fig. 6c, d). Some of the horizons are almost entirely deformed. These structures may be partly associated with pressure exerted by icebergs

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Fig. 12. (a) Iron-cemented multi-directional striations at the erosional unconformity between units 2 and 3 within the Sanamah Formation. (b) Fluted surface slightly above the erosional unconformity between Units S2 and S3 within the Sanamah Formation. (c) Unidirectional striations at the top of unit J3 of the Juwayl formation. (d) Flutes with oblique overprint of striations at the erosional unconformity between units S2 and S3 within the Sanamah Formation. (e) Iron-cemented unidirectional striations at the erosional unconformity between units S2 and S3 within the Sanamah Formation.

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drifting across stagnant water and occasionally touching down on the sediment surface.

Origin and possible significance of the large valleys Both the Sanamah Formation and the Juwayl Formation were partly deposited in broad valleys up to several kilometres wide and more than 100 m deep. In northern Saudi Arabia, similar Ordovician structures on seismic lines are up to 600 m deep (McGillivray & Al-Husseini 1992). Seismic mapping of these valleys in northern Saudi Arabia show a radial orientation and longitudinal shape with low sinuosity (Sharland et al. 2001). These features are therefore preferentially interpreted as tunnel valleys. However, other models also exist which have been reviewed by Sharland et al. (2001) for northern Saudi Arabia and Le Heron et al. (2004) for northern Africa. They consider a subglacial origin as well as pure fluvial incision and fluvial incision coupled with ice isostasy. A tectonic origin as proposed by Stump & Van der Eem (1995, 1996) is rather unlikely, given the tectonic quiescence at that time and the dimensions of uplift, erosion and subsidence required within a rather limited time interval. A radial pattern of valleys would also be unlikely. The lithofacies associations with vestiges of repeated outburst floods indicate that these valleys might have originated as tunnel valleys (Piotrowski 1994; Ghienne & Deynoux 1998; Le Heron et al. 2009). This interpretation implies subglacial erosion by meltwater during the ice maximum and shortly thereafter, when the hydraulic transmissivity of ice sheet stratum was insufficient to drain all basal meltwater and meltwater excess was evacuated via open channels at very high discharge rates and flow velocities (Piotrowski 1994). These sudden repeated outbursts caused efficient subglacial meltwater erosion close to the margin of the ice shield and may have prevented till deposition in the valleys as has been described from Pleistocene examples. The depth of our Palaoeozoic valleys of up to 100 m is surprisingly close to those reported from Pleistocene examples of northern Germany (Piotrowski 1994). One distinct property of subglacial tunnel valleys is the irregularity of their gradient. We could not decide whether this is true for our valleys, because valley floors were not exposed along longitudinal profiles and geophysical surveys have not yet been carried out in this area.

Large-scale soft-sediment deformation by glacier push Glaciogenic large-scale soft-sediment deformation structures are well known from Quaternary ice

sheets, but have rarely been recognized in Palaeozoic deposits. Le Heron et al. (2005) showed that structures such as large thrust-and-fold systems, giant ball-and-pillow structures and dome-like structures are widespread in front of the Upper Ordovician ice sheet in Libya and that they have important glaciological and stratigraphic significance. These structures were produced by glacier push and by load-induced diapirism. Similar features are present at the top of the Sanamah Formation and within the upper part of the Juwayl Formation, where these phenomena attain dimensions of up to 30 m. They are not restricted to the presumed tunnel valleys, and thus are valuable marker beds. The deformation is most probably caused by a short-term glacial surge during deglaciation and potentially of chronostratigraphic value (Le Heron et al. 2005).

Palaeozoic glaciological models for SW Saudi Arabia Based on observed facies associations, architecture and glaciogenic sedimentological features as described before, we present a conceptual model for the two Palaeozoic glacial episodes in SW Saudi Arabia (Fig. 11a, b). We propose threeand two-glacial advances for the Ordovician and Permo-Carboniferous, respectively, which include additional high-frequency fluctuations as deduced from the sedimentary deposits. The sedimentary record starts at or shortly after the ice maximum when subglacially eroded tunnel valleys began to be filled up. We suggest a polythermal regime for this time. In the Ordovician, a second advance led to erosion and created widespread striations in underlying beds (discontinuity S2/S3). In both the Ordovician and PermoCarboniferous cases, the upper part of the glaciogenic successions shows distinct horizons of widespread soft-sediment deformation followed by post-glacial flooding. The observed features of soft-sediment deformation could be explained by glacial shearing under mostly warm-based conditions. However, unequivocal till deposits are absent. We propose a short-term surge scenario for these late re-advances, probably initiated by the buoyancy of the ice sheet due to meltwater ponding in downwarped marginal areas and/or a post-glacial sea-level rise. Reduced ice viscosity due to temperature rise may also have played a role. In both cases, the glaciogenic succession is covered by pronounced marine transgressive sequences which correspond to primary maximum flooding surfaces in the sequence stratigraphic framework of the Arabian Plate (S10 and P20: Sharland et al. 2001).

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Significance of Palaeozoic glaciations of Gondwana The sedimentary record of the two glacial episodes in southern Saudi Arabia shows striking similarities and indicates rather similar glaciological conditions during the Late Ordovician and during the PermoCarboniferous. In addition, sedimentary facies and architecture of Upper Ordovician deposits of the Arabian Platform can be compared with those recently described from Mauritania (Ghienne & Deynoux 1998; Ghienne et al. 2003) and the Murzuq Basin in Libya (Le Heron et al. 2004). Le Heron et al. (2004, 2005) described many similar features and suggested a model of a polythermal regime at an oscillating ice front, followed by deglaciation and a late glacial surge. Finally, postglacial eustatic sea-level rise led to transgression and the deposition of marine beds above the glacial succession. The almost identical pattern of glaciological regimes over a distance of more than 1000 km along the northern margin of Gondwana supports the hypothesis of a large, coupled ice shield over at least North Africa and Arabia as has been suggested by Le Heron & Dowdeswell (2009), rather than an isolated ice sheet on the Arabian Shield. Some differences remain, however, concerning the number of ice advances and ice retreats. While Le Heron et al. (2004) found incision of the tunnel valleys into glaciogenic deposits of an earlier ice advance, and Ghienne et al. (2003) as well as Le Heron et al. (2006) suggested four glacial cycles from observations in Mauritania and Libya, Ordovician glaciogenic deposits in SW Saudi Arabia are restricted to the valley fill and thin onlapping on the valley margins. Together with a first glacier advance which eroded the subglacial valleys and deposited a first glaciolacustrine to deltaic sequence (S1), two further advances are indicated by a major intra-formational erosional unconformity (S2/S3) and widespread soft deformation at the top (S3) (Fig. 11b). Vaslet et al. (1991) concluded from observations on outcrops and seismic data from northern Saudi Arabia that two erosional episodes formed the tunnel valleys, which were subsequently filled by the Zarqa Formation and after the second erosional episode by the Sarah Formation. Both units together are the northern equivalent of the Sanamah Formation of the Wajid Group. Repeated ice advance-retreat episodes have also been reported from the subsurface of west-central Saudi Arabia (Melvin & Miller 2002). We suggest that these two erosional events in glaciogenic Upper Ordovician sequences of the northern and central Arabian platform can be correlated with the two unconformities observed in SW Saudi Arabia, and might be of high chronostratigraphic value. The second

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erosional event, however, is less pronounced in the SW and failed to cut down to the bottom of the valleys as has been reported from northern Saudi Arabia (pre-Sarah unconformity). The lack of pre-incision glaciogenic deposits on the Arabian Platform that are present in Libya might be explained by topographic effects such as a flatter morphology and the absence of prominent intracratonic basins on the Arabian Shield. This would have resulted in low accommodation and a lower potential for preservation of the sediments. Another explanation would be that tunnel valley formation was not exclusively restricted to one distinct glacial stage along the entire margin, but that it may have occurred repeatedly at an oscillating ice margin. It will probably never be possible to resolve these events biostratigraphically because of the absence of any fauna. Hence, a correlation to the North African events must necessarily remain speculative. As recent research (Saltzman & Young 2005; Buggisch et al. 2010) has shown that the Hirnantian glaciation was only the peak event of a succession of larger-scale glacial episodes since the Katian, a correlation entirely based on sedimentological criteria would be audacious. Hence it may be that not all of the events in Saudi Arabia correspond to the Hirnantian glaciation proper. The Permo-Carboniferous glacial episode is not well documented in northern Africa; it is, however, well constrained by biostratigraphical data from Oman (Pollastro 2003; Osterloff et al. 2004), Yemen (Kruck & Thiele 1983) and Ethiopia (Bussert & Schrank 2007). Here, glaciogenic deposits were deposited in glacial valleys cutting down to the Precambrian basement. For Ethiopia, Bussert & Schrank (2007) suggested that the sediments were deposited in a glaciolacustrine to glaciofluvial environment, similar to the model presented here. Similarily, Kruck & Thiele (1983) and Pollastro (2003) assume that large areas of the south-western Arabian Platform have been covered by a glacial lake during this time. We could confirm this hypothesis by describing typical glaciolacustrine facies associations from the Wajid group. Several indications exist that the PermoCarboniferous ice sheet was much more heterogeneous than that of the Late Ordovician. During the early Carboniferous, the Hercynian tectonic event led to a major unconformity on the Arabian Platform; up to several hundred metres of sediments were removed by erosion (Sharland et al. 2004). The presence of large angular basement blocks in the Juwayl deposits can only be explained by supraglacial transport and rock fall from a mountain range or nunataks. Such elevated areas seem to have existed not only around the Arabian Shield but also at the southern edge of the Arabian Platform in Oman,

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Yemen and Ethiopia (Kruck & Thiele 1983; Bussert & Schrank 2007). The glaciogenic deposits of Bani Khatma, 250 km south of the other outcrops (Fig. 1), show a spectrum of dropstone petrography not known from the presently exposed rocks of the Arabian Shield (McClure et al. 1988). The clasts were presumably shed from a southern source located at the opposite margin of the glacial lake. Kruck & Thiele (1983) found a SW-directed and Bussert & Schrank (2007) a north–south directed ice flow. Both are opposite to the model of a single huge Gondwana ice shield with a unidirectional north– NW directed ice flow. Therefore, Bussert & Schrank (2007) concluded that an independent ice cap must have existed in Eritrea and/or in southern or central Saudi Arabia, which was probably restricted to the area of the Hercynian uplift in this region. The incision of tunnel valleys, a polythermal setting in front of an oscillating ice sheet and extensive glacial surge during a late stage of glaciation reflects similar glaciological processes in the Carboniferous and in the Ordovician. This includes large ice sheets which discharge to sea- or lakelocked coastal areas. A remarkable difference is the quartzose petrography of the Ordovician conglomerates and accompanying sandstones compared to the Permo-Carboniferous deposits. Upper Ordovician glacial erosion mostly reworked mature shelf sediments poor in clay, which had been deposited across the Precambrian–Cambrian peneplain of the Arabian Shield (Avigad et al. 2005). No pebbles of basement rocks were found, which indicates that the Arabian Shield was not exposed or was poorly exposed during the Late Ordovician and that it did not act as a major sediment source. This drastically changed during the Hercynian tectonic event. In addition, less intense chemical weathering and enhanced production of clay minerals with the advent of land plants may be responsible for a quite different composition of the Carboniferous sediments (Keller & Lehnert 2010).

Conclusions Based on detailed analysis of sedimentary facies and architectural geometries of Upper Ordovician and Permo-Carboniferous glaciogenic deposits, a genetic model for both glacial episodes is proposed and compared with other locations in northern and north-eastern Gondwana. Saudi Arabia is one of the few places where both glaciations can be studied in well-preserved sedimentary succession, which formed under similar boundary conditions. Among these are (a) incision of subglacial tunnel valleys, (b) subsequent valley-fill by subaqueous to subaerial

proglacial deposits in front of an oscillating, polythermal ice shield, (c) widespread and large-scale soft sediment deformation due to glacial surge during deglaciation and (d) marine transgression as a consequence of eustatic sea-level rise following deglaciation. The general pattern fits well with observations recently published by Ghienne et al. (2003) and Le Heron et al. (2004, 2005, 2009) from northern Africa, pointing to closely coupled glaciological processes along the northern margin of the Gondwana ice shield during Upper Ordovician glaciation. Although the same general patterns apply for the Permo-Carboniferous glaciation, styles and petrographical properties differ most probably due to a more complex ice-flow pattern with a rougher topography after the Hercynian event and contrasting weathering. This paper benefited from thoughtful comments and suggestions by F. Ghienne, D. Le Heron and I. P. Martini.

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Frozen sediments and previously-frozen sediments H. FRENCH Departments of Geography and Earth Sciences, University of Ottawa, Canada (e-mail: [email protected]) Abstract: Frozen surficial sediments that are otherwise unconsolidated contain structures and characteristics that are different from those of the same sediments when in an unfrozen state. These differences are usually related to either the nature of the ice contained within the frozen sediment or to weathering processes and chemical precipitates that are associated with freezing and thawing. This paper summarizes (a) the manner in which ground freezes when a landscape experiences the onset of cold-climate conditions and (b) what happens when newly transported sediments freeze following deposition in that environment. In the absence of obvious morphological evidence, the recognition of previously-frozen sediments is problematic. Less well-understood evidence includes secondary precipitates, neoformed clay minerals, seasonal frost cracks and fragipans.

Unconsolidated surficial sediments can be frozen on diurnal, seasonal or perennial timescales. When perennial frozen, the sediments constitute permafrost, that is, earth material that remains at temperatures of below 0 8C for at least two years (Sumgin 1927; Muller 1943). The major difference between seasonally and perennially frozen sediments is that seasonally frozen sediments thaw during the summer whereas perennially frozen sediments experience progressive annual freezing; only the near-surface layer, the so-called active layer, thaws each year. A second difference is that sediments that only freeze seasonally do so from the surface downward while perennially frozen sediments experience twosided freezing, that is, both downwards from the surface and upwards from the underlying permafrost. The thawing of both seasonally and perennially frozen sediments occurs primarily from the surface downwards. Diurnal freezing, a common experience in the mid-latitudes, is not considered in this paper. The study of frozen sediments is part of cryostratigraphy, itself a branch of permafrost science or geocryology. Cryostratigraphy differs from traditional stratigraphy by explicitly recognizing the presence of ground ice within sediments (Katasonov 1962, 1975; Gravis 1974; Popov et al. 1985; Mel’nikov & Spesivtsev 2000; French & Shur 2010). This can be of varying amount, distribution and origin. The aims of cryostratigraphy are: first, to understand the genesis of frozen sediment and to infer the frozen history of earth material and second, to correlate sequences of ground ice in contemporary frozen sediments with horizons of former ground ice in sediments that are no longer frozen.

Epigenetic and syngenetic freezing If freezing occurs after sedimentation has occurred then freezing is termed epigenetic. The time lag between deposition and freezing may reach thousands and millions of years. By contrast, freezing is termed syngenetic when freezing occurs at the same time (approximately) as cold-climate sedimentation is occurring. It can be argued that syngenetically frozen sediments are cold-climate (periglacial) sediments sensu stricto because epigenetically frozen sediments may be non-periglacial in origin; and the latter would merely have acquired cold-climate characteristics following the onset of periglacial conditions. The concepts of epigenetic and syngenetic freezing are illustrated in Figure 1. Anti-syngenetic freezing should also be mentioned because it must occur wherever anti-syngenetic or burrowing wedges develop.

Ice within frozen sediments It is well known that pore and segregated ice form when ground freezes (Williams & Smith 1989). Pore ice forms in situ in pore spaces; segregated ice forms by the migration of unfrozen water towards the advancing freezing plane. Capillarity and cryosuction are the physical processes responsible for ice segregation. It is also well known that not all water contained within soil and rock freezes at 0 8C. Instead, moisture may exist in a super-cooled state. For example, it has been demonstrated that small amounts of unfrozen water can exist in fine-grained sediments at temperatures as low as 215 8C. Therefore, if the freezing plane

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 153–166. DOI: 10.1144/SP354.9 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Diagrams illustrating the concepts of (a) syngenetic and (b) epigenetic permafrost formation. The y-axis on both diagrams represents depth (m). Sediment accumulation in (a) includes peat accumulation. In (b) the active layer includes the transition zone (see Fig. 2a).

remains at approximately the same level for an extended period of time, lenses of segregated ice can form. Other types of ice may also exist within frozen sediment (Shumskii 1959; Mackay 1972; Popov 1973; Vtyurin 1975). The most common is vein ice. This is formed by the penetration of snow and/or surface water into thermal-contraction cracks. Vein ice ranges in size from single nearvertical veins to large v-shaped foliated ice bodies (ice wedges). Although such cracking is generally regarded as a diagnostic feature of perennially frozen sediments, similar cracking is also known to occur in seasonally frozen sediments. Other types of ice are less common but usually easier to interpret. Examples include ice bodies formed by the freezing of bulk water injected ahead of an advancing freezing plane, such as pingo ice or that trapped in underground erosional conduits such as thermokarst-cave (‘pool’) ice. Perhaps the most spectacular type of ground ice is the buried glacier ice that exists in parts of the Western Canadian Arctic and Western Siberia.

Frozen sediments and freezing history The nature and distribution of ice present within perennially frozen sediment can, in certain situations, allow one to infer either syngenetic or epigenetic freezing and, by extension, either the depositional environment or the freezing history of the sediment. There are several lines of evidence.

The interface between the active layer and permafrost The base of the active layer (the depth of annual thaw) can vary on annual, decadal and millennia timescales. An ice-rich zone characterized by icy lenticular layers marks the long-term position of the contact between the active layer, as defined above, and the upper part of the permafrost (Shur et al. 2005). This transition zone can be further subdivided into a short-term transient layer and a longer-term intermediate layer (Fig. 2). Identification of a well-developed ice-rich zone as being the active-layer–permafrost interface explicitly recognizes the presence of a stable ground surface and infers epigenetic freezing of sediments. Typically, the active layer is identified by field probing to determine the depth to frozen sediment at the end of the summer. The depth of the active layer is an important environmental indicator because, once identified, it can be interpreted in terms of the variables that influence ground thawing and freezing; these include air temperature, snow cover and vegetation and the presence or absence of organic and/or mineral soil. Moreover the Stefan equation (Johnston 1981), commonly used by cold regions geotechnical engineers, can be reformulated to use the thickness of the active layer to infer the mean annual air temperature (MAAT) and assist in palaeoenvironmental reconstruction. When permafrost degrades, the transient layer (the icy zone) at the permafrost table experiences thaw consolidation

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Fig. 2. The active-layer–permafrost interface. (a) The simple three-layer model (1, active layer; 2, transition zone; 3, permafrost). (b) A four-layer model shows two layers in the transition zone: 1, active layer (seasonal freezing and thawing); 2, transient layer due to variations over c. 30 years (the period defining the contemporary climate); 3, intermediate layer formed from part of the original active layer due to environmental changes, primarily ice-rich organic accumulation; 4, permafrost (freezing and thawing at century to millennial scales) (modified from French & Shur 2010).

and a compacted soil microstructure, or fragipan, can result (see p. 164). In environments of deep seasonal frost, no such zone of icy sediments develops at the base of the frozen layer because seasonal thaw, by definition, prevents the formation of permafrost. On the other hand, the zone of seasonal freezing may result in disaggregation (brecciation) of bedrock and friable soils, both caused by the formation of pore and segregated ice, and disruption of the near-surface caused by needle-ice formation. If low air temperatures prevail and there is deep frost penetration, or if the ground surface experiences rapid subzero temperature changes, seasonal frost cracks may form (see pp. 161– 164).

Cryostructures The amount and distribution of pore ice and lenses of segregated ice within frozen sediment can be described in terms of cryostructures. These are well known in the Russian and North American permafrost literature (Katasonov 1969; Zhestkova 1982; Murton & French 1994; Shur & Jorgenson 1998). They depend largely on the initial water content of the sediment and the extent of moisture migration during freezing. Some cryostructures, such as irregular and chaotic cryostructures, appear characteristic of epigenetic permafrost growth while others, such as lenticular and micro-lenticular cryostructures, appear

characteristic of syngenetic permafrost growth (Shur et al. 2004; Bray et al. 2006) (Fig. 3a–d). The best-known, but probably least common, cryostructure is that of an ice crust or rim around a rock clast; this is the crustal cryostructure that forms by localized ice segregation around frost-susceptible clasts just beneath the active-layer–permafrost interface. The orientation of lenticular cryostructures, generally termed ice lenses, typically reflects the orientation of the freezing front and/or the structural properties, such as bedding, of the sediment. Reticulate cryostructures are commonly interpreted to reflect the desiccation and shrinkage of finegrained sediments upon freezing. Glaciolacustrine sediments provide a good example of the variable cryostructures that form during syngenetic freezing and their subsequent thaw followed by epigenetic refreezing. For example, freezing occurs as sedimentation takes place; this gives rise to a layered cryogenic structure with continuous ice layers and an ice content that may be as high as 85%. When thaw consolidation occurs followed by refreezing, random or chaotic epigenetic cryostructures form and the sediment exhibits low measurements of soil water and low visible ice content.

Weathering profile differences Field and experimental studies indicate that one effect of freezing and thawing of unconsolidated

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surficial sediments is the preferential production of silty particles with grain sizes of between 0.05 and 0.01 mm in diameter. One hypothesis attributes this to the variable thickness of the layer of unfrozen water that is adsorbed on different primary minerals. Where the thickness of this layer is less than the dimensions of the various micro-fractures and defects that characterize the surface of the mineral particle, ice forms in micro-cracks and produces volume widening following repeated freeze –thaw. Russian studies indicate that, under these conditions, quartz particles exhibit a relatively high degree of instability relative to feldspar particles (Konishchev & Rogov 1993). A coefficient of cryogenic contrast (CCC), based on the relative proportions of quartz and feldspar grains in the size fractions of 0.05 –0.00 and 0.1–0.05 mm, has been developed that identifies sediments affected by cryogenic weathering. A value in excess of 1.0 is generally thought indicative of cryogenic weathering. For example, a detailed CCC study undertaken at Moscow State University of sediments of different ages in several of the major north-flowing drainage systems of central and eastern Siberia found that sediments of Quaternary age all indicated cryogenic weathering while earlier sediments, of Neogene and Palaeocene age, did not (Table 1). Other studies suggest physico-chemical changes occur under negative temperatures. For example, certain clay minerals experience aggregation and coagulation together with marked changes in Na and K ions after 100 freeze-thaw cycles (Datsko & Rogov 1988; Yershov 1998). There are also studies that indicate secondary precipitates of iron, calcite and sulphates with peculiar crystalline features may form during the freezing process (Vogt & Larque´ 2002). For example, on Seymour Island, Antarctica, the lower faces of cobbles and rock debris in the active layer are coated with acicular crystals, growing vertically downwards normal to the cooling surface (Vogt & Corte 1996). X-ray diffraction and SEM analyses show the coatings consist of needle-like calcite and finely-cleaved gypsum crystals. Although these precipitates could be ascribed theoretically to evaporation of water during the summer months, this is unlikely for two

Fig. 3. Cryostructures in poorly-lithified sediments. (a) A lenticular-layered cryostructure, with ice lenses 30– 80 cm long and 5 –10 cm thick, in glaciolacustrine silty clay, Mayo, Yukon Territory, Canada. The fallen spruce tree, bottom left, is about 1.5 m long.

Fig. 3. (Continued) (b) Reticulate ice veins formed in silty clay diamicton, Pleistocene Mackenzie Delta, NWT, Canada. Ice axe for scale. (c) Reticulate-chaotic cryostructure (location marker is 1.25  1.25 cm in size), CRREL permafrost tunnel, Alaska (photograph courtesy of Y. Shur). (d) Lenticular cryostructure (core diameter is 5 cm), CRREL permafrost tunnel, Alaska (photograph courtesy of Y. Shur).

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Table 1. CCC values* (.1.0 indicate cryogenic weathering) of sediments of different ages in central and eastern Siberia where Q1 –Q4, Quaternary sediments; N1– N3, Neogene sediments; P, Palaeogene sediments (Konishchev & Rogov 1993) Age

Yana/Indigirka

Kolyma

Chukotka

Average

Quaternary Q4 Q3 Q1–3

2.11/1.49/1.57 1.77/1.23/3.00 1.90/1.02/3.30

1.91 2.27/1.41 1.21

1.6 – 1.6

1.74 1.93 1.82

Pleistocene –Pliocene boundary N3–Q1 1.70/0.67

1.41

Neogene N2 N2–N3 N1

0.76/0.36 0.80 –

– – 0.78

– – 1.51

0.56 (0.80) 0.70

Palaeogene P

0.65

0.51

–

0.58

1.06

*CCC ¼ (Qi/Fi)/(Qii/Fii) where Qi, Fi and Qii, Fii are the quartz, feldspar content (%) in the 0.05 –0.01 mm and 0.01 –0.05 mm fractions, respectively.

main reasons. First, some precipitates occur below the layer of maximum frost penetration. Second, melting ice produces water deprived of solutes and so no precipitation takes place when it evaporates. Furthermore, evaporation of mixed solutions produces mixed assemblages of crystals whereas the aggregates in question are either pure calcite or gypsum. The more likely explanation is that the freezing front draws moisture upwards, segregating and concentrating solutes ahead of the freezing front in the unfrozen zone. This would be favoured in sediments situated in a saturated zone between an advancing cooling front and an underlying impermeable layer (possibly permafrost). Accordingly, the presence of these unusual precipitates is regarded by some as a clear indicator of either past or present permafrost (Vogt & Corte 1996; Vogt & Larque´ 2002).

Ice discontinuities If sediments are ice-rich, their thaw or partial thaw can sometimes be recognized by ice discontinuities; these are termed ‘thaw unconformities’. The activelayer –permafrost interface (see above) is one such thaw unconformity. In syngenetically frozen sediments, the successive positions of the active-layer – permafrost interface can sometimes be identified by the correspondence of the upper limits of frost cracks. For example, in the walls of the Cold Regions Research and Engineering (CRREL) permafrost tunnel, Alaska, successive horizons of small frost cracks are present in the Late Pleistocene age silty sediments. These sediments contain a preponderance of syngenetic cryostructures (Kanevskiy et al. 2008).

Detailed cryofacial studies indicate seven thin organic-rich horizons in the tunnel wall (Fig. 4). This exposure is at a depth of 12 –14 m below the ground surface. At depths of between 0.4 and 0.6 m below each horizon, distinct ice-rich layers and numerous thin cracks partially filled with ice (ice veins) extend downwards from the organic horizons and form polygons up to 0.5 m across in horizontal plane. The cryostratigraphy clearly illustrates the progressive accretion of sediment under permafrost conditions and the upwards migration of the active-layer –permafrost interface. Sometimes, a regional ice discontinuity may form during a period of climatic amelioration, only for sediments to refreeze during a subsequent climatic deterioration. In this case, the unconformity can be recognized because younger, typically epigenetic, cryostructures form above the unconformity. One such unconformity occurs widely today in the Western Canadian Arctic at a depth of c. 125– 150 cm (Burn et al. 1986; Burn 1997). This unconformity is thought to represent the maximum depth of the active layer that formed during a warm episode of the early Holocene. However, interpretation of such thaw unconformities must be undertaken with care. Although widespread truncation of icy bodies clearly indicates a regional thaw event, the truncation of more discrete ice bodies may merely indicate a local thaw-erosion event. For example, if thaw-erosion were preferentially concentrated along the axes of ice wedges, subsequent collapse and fill of the gully may result in the formation of thermokarst-cave ice or soil pseudomorphs (Murton & French 1993; Kanevskiy et al. 2008). The latter would contain epigenetic cryostructures.

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Fig. 4. Cryostratigraphic map of part of the wall of the CRREL permafrost tunnel, Alaska. It shows bodies of syngenetic permafrost separated by an erosional (thaw) boundary. Scale indicated is in metres (modified from Kanevskiy et al. 2008).

Recognition of mineralogical and weathering differences that occur above and below a thaw unconformity can also indicate prior freezing histories because the heavy mineral assemblages may indicate chemical weathering during thaw (Xing et al. 1980; Burn et al. 1986; Burn 1997). For example, two distinct heavy mineral assemblages are present within a perennially frozen diamicton in the Pleistocene Mackenzie Delta, Canada (Murton et al. 2005). The stratigraphically higher assemblage is within a palaeothaw layer. It is depleted in pyrite and siderite due to oxidation and solution and enriched in goethite, leucoxene, garnet and ilmenite relative to the lower assemblage at the base of the palaeothaw layer and in the underlying permafrost (Table 2). The preservation of pyrite and siderite in the basal part of the palaeothaw layer is attributed to the former occurrence of reducing conditions beneath a perched water table lying above the impermeable permafrost table at the time. There are several reasons that suggest mineralogical changes in the weathering profile may not occur or are not discernible. For example, the sediments may contain minerals that are relatively nonsusceptible to weathering and the thaw-related event

may have been relatively short lived. Moreover, if freezing were syngenetic and the ground surface was actively aggrading during the weathering period, profile differentiation may not occur.

Previously-frozen sediments Traditionally, the recognition of a previously frozen state in Pleistocene-age surficial sediments has been based upon permafrost-related structures such as ice-wedge pseudomorphs with secondary infill and surrounding collapse structures, sand veins and sand-wedge casts, and the remnants of aggradational permafrost landforms such as frost mounds (pingos and palsas). All are associated with perennially frozen ground and all are well documented in the recent Quaternary literature (Elias 2007, volume 3, 2138 –2256; French 2007). More ambiguous are the frost-disturbed structures (cryoturbations) that occur widely in the near-surface because, although clearly indicative of intense frost action, it is not clear whether they require seasonally or perennially freezing for their formation. Evidence that relates to the thaw of frozen sediments is also well documented in the same recent

Cryostratigraphy

Palaeothaw layer (melt-out till) Diamict (basal glacier ice; basal till)

Sediments

Height above thaw contact (cm)

Pyrite (Py)

Siderite (Sd)

Garnet (Gn)

Goethite (Go)

Ilmenite (Il)

Leucoxene (Le)

Sand and diamicton assemblage Layered assemblage Sediment-poor ice Sediment-rich ice

115– 125 60– 70 10– 30 30– 50 100– 120

0.7 0.3 47.0 38.0 40.0

0.7 0.3 6.3 5.7 8.3

8.7 8.6 2.7 10.0 5.0

22.6 26.0 33.0 30.7 28.4

9.6 7.0 0.7 2.3 1.7

3.3 3.0 0.7 2.3 2.3

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Table 2. Frequency percentages (based on 300 grain count) of selected heavy minerals from above and below a secondary thaw contact (palaeothaw unconformity; data from Murton et al. 2005)

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literature (see above). First, some of the most common takes the form of gravity-induced soil movement that develops sediment lobes and sheets. The latter are the result of either solifluction (movement of saturated material in seasonally frozen terrain) or gelifluction/plug-like flow (movement of thawed material above and across the permafrost table). Sometimes it is difficult to distinguish between these two modes of sediment movement when the ground is no longer frozen. Second, the large nondiastrophic deformations that occur at depth and take the form of load casts, pseudo-nodules, balland-pillow structures and diapirs are frequently interpreted in the context of degrading permafrost with water-saturated sediments occurring in the thawed sediments above a relatively impermeable permafrost table. The mechanisms that are usually suggested involve loading, buoyancy and water escape and fluidization. Unfortunately for the periglacial/permafrost interpretation, these mechanisms can operate not only in degrading ice-rich sediments but also in non-frozen settings. Another line of evidence concerns the veneer of slumped heterogeneous mass-flow deposits that mantle many low-angled slopes in mid-latitudes. Upon geotechnical analysis, some of these materials indicate a failure angle considerably lower than the equilibrium angle predicted for that slope. In many instances, the only explanation is to attribute failure to frozen-ground conditions in the past and, more specifically, to the thaw consolidation of icerich and fine-grained sediment bodies. This would have occurred widely in the mid-latitude lowlands that lay beyond the limits of the various Upper Pleistocene glaciations but which experienced frozen-ground conditions during the height of the glacial episodes. Essentially, the pore-water pressures associated with thaw and the progressive reduction in shear strength at the base of the active layer from gelifluction movement are invoked to cause failure on slopes at angles well below the normal threshold for that slope. Analogous processes in present permafrost environments result in

periglacial mudslides and active-layer-detachment failures. Again, these processes are well documented in the recent literature cited above. Unresolved problems associated with palaeocryostratigraphic reconstruction in the nowtemperate mid-latitudes centre upon (a) recognition of the palaeo-active-layer –permafrost interface, possibly indicated by the presence of an indurated layer such as a fragipan or a calcrete and (b) differentiation between previous conditions of permafrost and deep seasonal frost, possibly indicated by clays, secondary minerals and frost cracks. The evidence to date is summarized below.

Secondary precipitates, clay minerals and CCC values In the absence of the more traditional and obvious morphologic and stratigraphic evidence for previously-frozen ground in unconsolidated sediments, secondary precipitates, clay minerals and coefficient of cryogenic contrast (CCC) values may provide some of the best evidence. Outside Russia, there are few studies that utilize CCC methods. One is upon fine-grained sediments in the Pine Barrens of southern New Jersey (Table 3; Demitroff et al. 2007) where a general chronology of past permafrost events has been established (French et al. 2007). Three sediment samples taken from within a deformed sandwedge structure, optically-stimulated-luminescence (OSL)-dated to marine isotope stage (MIS)-3, indicates cryogenic weathering of the wedge infill while the enclosing sediments have not experienced similar conditions. Elsewhere in the Pine Barrens, the near-surface sediments in three shallow depressions (‘spungs’), interpreted as having formed by deflation during the late glacial, all showed relatively high CCC values. However, these data are limited in quantity and, until more systematic analyses have been undertaken, the validity of CCC as a diagnostic tool for the identification of previously-frozen sediments remains unproven.

Table 3. CCC values* (.1.0 indicate cryogenic weathering, calculated as for Table 1) of near-surface samples of surficial sandy sediments in the Pine Barrens, southern New Jersey, US (data from Demitroff et al. 2007 and age determinations from French et al. 2007) Age MIS-3 (Wisconsinan) OSL dated to c. 35 ka (UIC 1494) Late Pleistocene–Early Holocene (c. 12 ka)

*See Table 1 for definition of CCC.

Sediment

Location

CCC value

Gravelly sandy

‘Sediment-filled pot’

1.44

Silty sand

Deflation hollows (‘spungs’) (i) Lee Pond (ii) Debbie Pond

1.33 2.41

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In many parts of southern Europe and Eurasia, numerous studies have observed calcareous crusts in Quaternary sediments such as gre`zes lite´es, talus cones, loessic materials and alluvial terraces (Cailleux 1972; Vogt 1977). The crusts contain unusual calcite crystals and iron hydroxides; laboratory experiments suggest they are due to cryoosmosis. Calcretes with similar characteristics occur elsewhere in the mid-latitudes, as in Morocco and central Patagonia (Vogt 1989, 1992). For example, the calcretes that are widespread in southern Patagonia provide convincing evidence for past frozen ground because they infill and overlie a network of frost cracks, 1– 1.5 m deep and spaced at 1– 2 m apart (Vogt 1992; Vogt & Del Valle 1994). Although these cracks are referred to as ice-wedge pseudomorphs (Corte & Beltramone 1984), their relatively small depth and close spacing suggest they reflect seasonal rather than perennial frost (see below). As explained earlier, the precise mechanism concerning calcite crystallization under coldclimate conditions is still unclear. Equally problematic are the changes in grain size composition of clayey soils and associated physicochemical properties under freezing and thawing conditions (Yershov 1998). Several studies suggest that neoformed clays form under cryogenic conditions. For example, in the circum-Baikal region, typical kaolinite deposits of pre-Pleistocene age were probably formed under warm and humid conditions. However, in locally-derived Pleistocene sediments in the same area, the clay fraction

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contains a detrital mineral assemblage of not only chlorite, illite and kaolinite, as expected, but also smectite and sepiolite (Vogt & Larque´ 2002). The latter two minerals are not inherited from the prePleistocene rocks or soils and suggest that detrital clay minerals are destroyed under cryogenic conditions, allowing the neoformation of certain clays. If further studies confirm this conclusion, the presence of unusual clay minerals such as smectite and sepiolite may be a positive indicator of previouslyfrozen sediment.

Frost cracking: seasonal or perennial frost? Frost cracking is not restricted to permafrost environments because there are documented cases of active frost cracking occurring in areas of deep seasonal frost (Washburn et al. 1963; Friedman et al. 1971; Svensson 1988). In the Russian permafrost literature, these seasonal frost cracks are termed soil wedges or ‘ground veins’ (Danilova 1956; Katasonov 1973; Romanovskii 1973, 1978). They are considered part of a range of perennial frost-crack structures (Fig. 5). Unfortunately, it appears extremely difficult to differentiate between a seasonal and perennial frost crack, especially when observed in previously-frozen sediment. The essential difference is one of size; seasonal frost cracks are restricted to the seasonally-frozen layer. The problem is further compounded when we appreciate that the controls over frost cracking are complex. Field studies indicate that active frost

Fig. 5. Classification of frost fissure structures (modified from Romanovskii 1978).

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Fig. 6. Perennial frost cracks: ice-wedge pseudomorphs and sand-wedge casts. (a) Ice-wedge pseudomorph developed in Saalian-age fluvioglacial gravels, north of Ło´dz, central Poland. (b) A bundle of sand veins indicates the repeated thermal contraction cracking that allowed growth of a Late Pleistocene sand wedge by the filling of each crack by wind-transported sediment, Bełchatow opencast mine, central Poland. (c) Sand wedge cast, 2.7 m deep and 5– 20 cm wide, contains well-sorted wind-abraded fine sand OSL dated at .65 ka, Pine Barrens, southern New Jersey, eastern USA. (d) Deformed Saalian-age sand-wedge cast, Bełchatow opencast mine, central Poland.

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cracking occurs not only when air and ground surface temperatures are lower than c. 215 to 220 8C (French 2007) but also that the atmospheric cooling rate, the ground cooling rate at the surface prior to cracking and the ground thermal gradient (8C m21) may be important. Moreover, small-scale cracks, forming polygons 1–1.5 m in diameter, have been described from tundra regions and attributed to wind action and desiccation (Karavaeva 1969); linear cracks that have been observed in the Colorado Front Range and the uplands of southeastern Siberia have been attributed to differential frost heave and seismo-dislocation, respectively (Benedict 1970; Leshchikov 1999). It is clear, therefore, that the utility of frost cracks as a differentiating criterion between perennial and seasonal frost is highly problematic. In terms of field recognition, perennial frost cracks become progressively filled with ice, mineral soil or a combination of ice and wind-blown material. This results in the well-known primary wedges such as ice wedges and sand wedges. Typically, epigenetic wedges penetrate as much as 2.5– 3.0 m beneath the ground surface, are c. 0.5–1.5 m wide

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near the surface and, in plan, form large-scale polygonal or hexagonal patterns, sometimes as much as 10 –15 m in dimensions. Syngenetic features can be even larger, although many are also relatively small (French & Gozdzik 1988). The thawing of primary wedges results in deformation and collapse, the degree of modification being dependent upon the ice content present both within the wedge and the enclosing sediment (Fig. 6a–d). In the case of seasonal frost cracks, no such obvious thaw-related deformation structures are visible. Instead, the wedges are narrow; the primary fill, if present, is usually similar to the enclosing sediment and relatively little deformation occurs in the enclosing materials. The most obvious difference between seasonal and perennial cracks is that the depth of seasonal crack penetration is shallow, usually between 30 and 70 cm (Danilova 1956; Friedman et al. 1971; Mel’nikov & Spesivtsev 2000). A second characteristic is that seasonal cracks are relatively closely spaced, every 1–3 m. For example, several generations of intraformational wedges that are spaced 2–3 m apart have been described from southern Patagonia (Vogt & del Valle 1994).

Fig. 7. Seasonal frost cracks (soil wedges). (a) In central Iceland frost cracks are actively forming under conditions of deep seasonal frost. Numbers refer to the age of tephra layers associated with various Heckla (H) and Katla (K) eruptions. The tephra layers and associated ages demonstrate the syngenetic nature of the frost crack. (modified from Friedman et al. 1971.) (b) A relict seasonal frost crack, seen in oblique section and developed in a 30– 50 cm thick coversand unit, penetrates the upper part of an underlying Tertiary age sand and gravel unit in southern Maryland, eastern USA. An indurated layer (possibly either an upper-Tertiary age fragipan or the LGM palaeopermafrost table) exists at a depth of 1.0–1.5 m below the present ground surface.

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They are filled with calcareous debris derived from an overlying calcrete horizon (see above). These cracks were almost certainly syngenetic in origin because the authors admit that they constitute ‘. . . an intraformational phenomenon’ (Vogt & del Valle 1994, p. 66). Today, areas of active frost-crack formation in areas where permafrost is absent include central Iceland and northern Norway (Fig. 7a). In Pleistocene sediments, the recognition of seasonal frost cracks is one of the few criteria that allow the interpretation of deep seasonal frost conditions rather than permafrost. For example, based upon their occurrence in northern Maryland (Fig. 7b), a zone of Upper Pleistocene deep seasonal frost existed on the Mid-Atlantic Coastal Plain of the eastern US (French et al. 2007, 2009). This lay to the south of a zone of permafrost in southern New Jersey and northern Maryland and north of a zone in northern Virginia that never experienced deep seasonal frost.

Fragipans and the palaeopermafrost table The recognition of a near-surface fragipan with a platy microstructure is sometimes thought to reflect the palaeopermafrost table (Fitzpatrick 1956; van Vliet-Lanoe & Langhor 1981). This is because the icy nature of the transition zone (Fig. 2) suggests that an indurated layer would form at the palaeopermafrost table following thaw consolidation. Typically, ice segregation in clayey sediments causes consolidation and induration of individual platey fragments between the ice lenses. However, the presence of the transient and intermediate layers further complicates interpretation. Well-developed fragipans are widespread in temperate and tropical regions of the world. They are commonly interpreted as Upper Tertiary weathering phenomena (Schaetzl & Anderson 2005). It is highly probable that other processes, unrelated to cryogenic conditions, may also promote induration. Thus, a layer similar to a fragipan that occurs in the mid-latitudes today and which is clearly not being formed under today’s pedogenetic regime is highly problematic. For example, a widespread fragipan exists throughout the southeastern United States. This fragipan also extends northwards into the MidAtlantic Coastal Plain where seasonal frost cracks lie stratigraphically above the fragipan (Fig. 7b). This forces the unfortunately ambiguous conclusion that the fragipan could have either coincided with the palaeopermafrost table under earlier permafrost conditions, or acted like a permafrost table under later conditions of deep seasonal frost. We must conclude that both recognition of a fragipan in the mid-latitudes and its possible use as a palaeopermafrost indicator must be approached with caution.

Conclusions Although frozen sediments possess distinct characteristics that relate to either the ice contained within the sediments or to the freezing process which they experienced, recognition of a previouslyfrozen state of thawed sediment is more problematic. Where obvious morphological evidence is lacking, the presence of secondary precipitates, neoformed minerals, weathering and grain size differences, seasonal frost cracks and an indurated layer are possible, but currently ambiguous, criteria. Continued understanding of the nature and rate of thaw of previously-frozen ground has implications for the impact of climate change in high latitudes where perennially-frozen ground currently exists. Thanks are expressed to Y. Shur (UAF) for Figure 3c, d and M. Kanevskiy (UAF) for drafting Figure 4. Review comments by J. B. Murton and C. Thorn are appreciated.

References Benedict, J. B. 1970. Frost cracking in the Colorado Front Range. Geografiska Annaler, 52A, 87–93. Bray, M. T., French, H. M. & Shur, Y. 2006. Further cryogenic investigations in the CRREL permafrost tunnel, Fox, Alaska. Permafrost and Periglacial Processes, 17, 233 –243. Burn, C. R. 1997. Cryostratigraphy, paleogeography, and climate change during the early Holocene warm interval, western Arctic coast, Canada. Canadian Journal of Earth Sciences, 34, 912–925. Burn, C. R., Michel, F. A. & Smith, M. W. 1986. Stratigraphic, isotopic, and mineralogical evidence for an early Holocene thaw unconformity at Mayo, Yukon Territory. Canadian Journal of Earth Sciences, 23, 794–803. Cailleux, A. 1972. Fentes minces et de´poˆts chimiques secondaires en Iakoutie et au nord de Moscou. Biuletyn Peryglacjalny, 21, 21– 31. Corte, A. & Beltramone, C. 1984. Edad de la estructuras geocryoge´nicas de Puerto Madryn (Chubut). In: Corte, A. (ed.) Actas, Secundo Reunion del Grupo Periglacial Argentino, Universidad Nacional de San Juan, 2– 6 Abril, 1984, IANIGLA/CRICYT/ CONICET, Mendoza, 66–72. Danilova, N. S. 1956. Ground veins and their origins. Materialy k Osnovam Ucheniya o Merzlykh Zonakh Zemnoj Kory, 3, 109– 122 (in Russian). Datsko, P. S. & Rogov, V. V. 1988. Transformation of dispersed deposits under cycles of freezing-thawing. In: Microstructure of Frozen Ground, Moscow University Press, Moscow, 132–149 (in Russian). Demitroff, M., Rogov, V. V., French, H. M., Konischchev, V. N., Streletsky, D. A. & Lebederva-Verba, M. D. 2007. Possible evidence for episodes of Late Pleistocene cryogenic weathering, southern New Jersey, Eastern USA. In: Proceedings, International Conference on Cryogenic Resources of Polar Regions. Earth Cryosphere Institute, Russian

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Gullies, polygons and mantles in Martian permafrost environments: cold desert landforms and sedimentary processes during recent Martian geological history J. S. LEVY1*, J. W. HEAD2 & D. R. MARCHANT3 1

Department of Geology, Portland State University, 1721 SW Broadway, Cramer Hall Room 17, Portland, OR 97201, USA

2

Department of Geological Sciences, Brown University, 324 Brook Street, Providence, RI 02912, USA 3

Department of Earth Science, Boston University, 675 Commonwealth Ave, Boston, MA 02215, USA *Corresponding author (e-mail: [email protected]) Abstract: A range of cold desert landforms are found on the Martian surface that have been interpreted to indicate prevailing frozen and hyper-arid conditions for at least the past several million years. These cold desert conditions are punctuated by brief periods of localized surficial liquid water flow. Sediment transport pathways operate under these conditions of extreme cold and aridity and the processes involved generate permafrost landforms that are recognizable from spacecraft at local, regional and global scales. Thermal-contraction-crack polygons are associated with hemisphere-spanning mantle units that contain excess ice in the immediate subsurface. Sublimation is the dominant phase transition rather than melting under present Martian conditions. Evidence is presented for melting of near-surface snow, frost and/or ground ice in protected gully alcove microclimates during the most recent several million years.

Mars is a permafrost planet. The Martian surface supports a wide range of fluvial, volcanic and aeolian landforms analogous to features found on Earth (Chapman 2007; Carr & Head 2010). During most of the geological history of Mars (Laskar et al. 2002, 2004), the entire Martian surface and shallow subsurface have experienced mean annual temperatures well below 273 K (0 8C), commonly dipping below 220 K (Mellon & Jakosky 1993). Accordingly, the entire face of Mars meets the standard definition of a permafrost terrain (Gold & Lachenbruch 1973; Washburn 1973; French 2007). These permafrost conditions likely extend to a depth of several kilometres (Clifford & Parker 2001). Indeed, Mars may be considered a cryotic planet insofar as, at present, mean annual surface temperatures are below the melting temperature of several water-ice compounds and solutions (Yershov 1998). Is the permafrost terrain of Mars similar to that of Earth? Although permafrost conditions persist over c. 20% of the Earth’s land surface, much of Earth’s permafrost is found in the continental and maritime regions of the North American and Eurasian Arctic (French 2007). In these warmer climate zones, permafrost commonly experiences summertime

melting as the 0 8C (273 K) isotherm penetrates the frozen ground surface (Washburn 1973; Williams & Smith 1989; Yershov 1998; French 2007). This seasonally thawed portion of terrestrial permafrost is referred to as the ‘active layer’ and, when water-saturated (‘wet’), it is the horizon in which many of the classic permafrost landforms arise (Williams & Smith 1989; Vliet-Lanoe 1991; Yershov 1998). In contrast, Mars currently lacks a wet active layer, and has probably not experienced climate conditions permitting the widespread development of a wet active layer over at least the last 5– 10 Ma (Kreslavsky et al. 2008). Interestingly, though, many of the most dramatic Martian permafrost landforms (Fig. 1) including gullies, thermal-contraction-crack polygons and the latitude-dependent mantle (LDM), all formed more recently than c. 5 Ma (Mustard et al. 2001; Head et al. 2003; Milliken et al. 2003; Kuzmin et al. 2004; Riess et al. 2004; Levy et al. 2009a; Schon et al. 2009a) Accordingly, it is essential to consider Martian permafrost from a cold desert climate perspective in which wet active layers are rare or absent (Anderson et al. 1972; Gibson 1980; Marchant & Head 2007).

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 167–182. DOI: 10.1144/SP354.10 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Plot of permafrost landforms diagnostic of a range of morphogenetic climate regions on Earth and Mars (adapted from Baker 2001 and Marchant & Head 2007). Oval represents mean annual climate conditions typical of the Antarctic Dry Valleys. SUZ indicates the Antarctic Stable Upland Zone (Marchant & Head 2007). TD indicates Taylor Dome and LGM indicates the Last Glacial Maximum in interior Antarctica. Modern conditions at a range of latitudes on Mars and representative thermal-contraction-crack polygon populations typical of those latitudes are plotted, as are conditions modelled for ancient Mars at higher atmospheric pressures. For Martian polygons, field of view is c. 300 m in all cases (nomenclature from Levy et al. 2009d ). Flat-top small polygons are excerpted from PSP_001959_2485; peak-top polygons from HiRISE image PSP_001737_2250 and mixed-centre polygons from PSP_002175_2210. The field of view in the illustration of sublimation polygons in Beacon Valley, Antarctica is c. 200 m wide. Oblique aerial view of sand-wedge polygons in lower Beacon Valley, Antarctica, has a field of view c. 50 m wide. Composite-wedge polygons are illustrated in Wright Valley, Antarctica, cross-cut by a gully channel with a field of view c. 75 m wide. Aerial view of ice-wedge polygons in Taylor Valley, Antarctica has a field of view c. 75 m wide.

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On Earth, cold desert permafrost environments are more typical of the coldest Antarctic and Arctic environments than of warmer and more widely-studied (and inhabited) permafrost zones (Gibson 1980; Marchant & Head 2007; Levy et al. 2008). Extreme cold is a critical element for understanding terrestrial analogues for permafrost terrain on Mars; for example, the Phoenix lander (Smith et al. 2009) was sent to explore Martian permafrost near 688N latitude and reported peak summer air temperatures of only c. 245 K with atmospheric water vapour pressures of c. 1.8 Pa (Whiteway et al. 2009). These conditions are comparable to, but still colder and more arid than, the coldest and driest permafrost microclimate in the Antarctic Dry Valleys. There, mean annual temperatures are c. 251 K and mean annual water vapour pressure is c. 40– 50 Pa (Marchant & Head 2007). In extreme cold deserts, low-temperature, sublimation-driven processes dominate the geomorphologic record (Chinn 1981; Marchant et al. 2002; Marchant & Head 2007). This major difference, between wet and dry permafrost, guides much of the following discussion. Advanced studies of Martian permafrost incorporating future lander data may integrate other critical climate controls on permafrost morphology, such as annual positive degree-days and snow recurrence intervals (McKay 2008).

2008; Edwards et al. 2009). Particle sizes of Martian sediments range from boulders observable from orbit down to micron-scale dust particles observable with lander and rover microscopic imaging systems (Golombek & Rapp 1996; Pike et al. 2009). Boulders, cobbles, pebbles and finer sediments are common in Martian permafrost terrains, and the detailed analysis of sediment sorting or arranging is the subject of ongoing inquiry. Sorting by dry cryoturbation processes is suggested by Mellon et al. (2008), Heet et al. (2009) and Mellon et al. (2009b). Dry, non-churning permafrost processes are favoured by Levy et al. (2008, 2010a) (Fig. 2). This paper discusses three sedimentary landforms typical of Martian permafrost environments, gullies, polygons and mantling units. We explore how sediments are transported in these landforms and interpret mantling units as primarily resulting from atmospheric emplacement of ice and sediment, polygons as resulting primarily from sublimationdriven modification of mantling units and gullies as resulting from the top-down melting of nearsurface ice and entrainment of mantle-related sediments.

Sedimentary processes in Martian permafrost

What is particularly striking about the distribution of permafrost landforms on Mars is the fact that, despite the global occurrence of permafrost climate conditions (surface temperatures ,0 8C over inter-annual periods), ice-related landforms (gullies, thermal-contraction-crack polygons, etc.) have been shown to occur in latitude-dependent clusters (Kreslavsky & Head 2000; Mustard et al. 2001; Head et al. 2003; Milliken et al. 2003; Kostama

Observations of the Martian surface suggest that rocky, regolith-surfaced landscapes abound, making Mars an ideal laboratory for considering sedimentary processes in cold desert permafrost environments (Mutch et al. 1976, 1977; Golombek & Rapp 1996; Wyatt et al. 2004; Golombek et al.

Global-scale sedimentary processes: the Martian latitude-dependent mantle

Fig. 2. Sedimentary clasts on the Martian surface. (a) Dust, pebbles and cobbles at the Phoenix landing site. Dust patches commonly accumulate in polygon troughs, while polygon interiors are typically armoured by desert pavements of pebbles and cobbles (portion of Phoenix lander Surface Stereo Imager frame SS051EFF900731785_15C28R2M1-b). (b) A ‘boulder halo’ indicating the location of a buried impact crater on the Martian northern plains (Levy et al. 2008) (portion of PSP_001477_2470). (c) Boulder piles accumulated on polygonally patterned knolls near the Phoenix landing site (portion of PSP_001959_2485).

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et al. 2006; Soare et al. 2007; Levy et al. 2009a) (Fig. 3). For example, thermal-contraction-crack polygons (see next section) have been shown to form in geologically recent deposits that are craterdated to less than several Ma. These deposits drape

and smooth underlying terrain, and are present in stacked layers continuously from high latitudes equator-wards to c. 608. The deposits grow patchier and show signs of degradation from c. 608 to c. 308 (Kreslavsky & Head 1999, 2000, 2002; Mustard

Fig. 3. Key properties of the Martian latitude-dependent mantle (LDM) indicating that it is composed largely of massive, atmospherically-emplaced excess ice. (a) Medium-toned, polygonally patterned LDM material (right) is easily eroded by gully activity, and can be distinguished from darker-toned bedrock units (left) (portion of PSP_006794_1420). (b) LDM material fills craters, smoothing topographic variation (portion of PSP_006931_2530). (c) LDM material drapes underlying topography and landforms, and can accumulate to tens of metres thickness as indicated by shadow measurements. See detailed discussion of draping morphologies in Levy et al. (2009b) (portion of PSP_002175_2210). (d) ‘Windows’ can be eroded through LDM deposits, revealing pristine underlying landforms preserved beneath tens of metres of LDM deposits (portion of PSP_002175_2210). (e) Fresh impact crater, formed between 2004 and 2008 exposing and ejecting bright excess water ice at 468N latitude, beneath a lithic lag deposit (Byrne et al. 2009) (portion of PSP_010861_2265), with north to image top and illumination from the left. (f ) Re-imaging of the crater shown in part (e), 127 sols (Mars days) after the image in part (e) was collected. Note darkening of crater bottom and of ejecta. Sublimation rate modelling by Byrne et al. (2009) indicates a sediment/ice ratio of c. 1% sediment to c. 99% water ice.

COLD DESERT PERMAFROST LANDFORMS ON MARS

et al. 2001; Head et al. 2003; Milliken et al. 2003; Schon et al. 2009b) (Fig. 4). LDM deposits vanish equator-wards of c. 308 (Milliken et al. 2003).

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This unusual spatial distribution of ice-related features is generally consistent with the predicted stability depth for ice in the upper c. 1 m of the

Fig. 4. Schematic illustration of thermal contraction-crack polygon types observed on Earth and comparisons with Martian landforms. Block diagrams are adapted from Marchant & Head (2007) and show key morphological properties of ice-wedge, sand-wedge, composite-wedge and sublimation polygons. Levy et al. (2010a ) use a range of morphological characteristics to connect Martian thermal-contraction-crack polygons with genetic end-member types observed on Earth. Sublimation polygon variants are the dominant polygon type observed on Mars, grading into traditional sand-wedge polygons where excess ice is less abundant. Composite-wedge polygons may form in regions with occasional inputs of liquid water associated with gully activity. There is no definitive evidence of the presence of active ice-wedge polygons on the Martian surface. High-relief polygons are from a portion of PSP_001474_2520; flat-top small polygons are excerpted from PSP_001959_2485; irregular polygons are a portion of PSP_001959_2485; peak-top polygons are excerpted from PSP_01737_2250 and PSP_003217_1355; subdued polygons are a portion of PSP_003818_1360; and ‘gullygons’ are from PSP_002368_1275 and PSP_001846_2390.

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Martian surface based on neutron-, gamma-ray- and thermal emission-spectrometer results coupled with insolation-driven thermal modelling. For thermal modelling, see Jakosky & Carr (1985), Mellon & Jakosky (1995), Mellon et al. (2004, 2009a), Bandfield (2007) and Vincendon et al. (2010). For geophysical results, see Boynton et al. (2002), Feldman et al. (2002), Mitrofanov et al. (2002) and Kuzmin et al. (2004). The stability of subsurface ice alone cannot explain the presence of metres-thick, topographically high, crater-filling deposits that are surfaced by thermal-contraction-crack polygons and gullies. In light of global morphological observations, a number of authors have proposed the presence of a Martian latitude-dependent mantle (LDM), a metres-thick ice-and-dust layer which was deposited as atmospherically precipitated ice and lithic material during recent (c. 2–4 Ma) periods of high orbital obliquity (Mustard et al. 2001; Kreslavsky & Head 2002; Head et al. 2003; Laskar et al. 2004). The LDM model predicts the presence of massive, excess ice and dust beds that have undergone sublimation, allowing excess ice nearest the surface to be removed down to the depth of subsurface ice stability while simultaneously producing a thick, rocky, protective lag deposit at the surface (Head et al. 2003; Schorghofer & Aharonson 2005; Schorghofer 2007). The presence of nearly pure-water ice beneath a lithic lag deposit at Martian middle to high latitudes and the mapped range of the LDM was confirmed by (Byrne et al. 2009) through the detection of fresh impact craters that exposed a bright substrate spectroscopically identified as water ice (Fig. 3e). The bright, spectroscopicallydiagnostic ejecta and crater-bottom material faded to background brightness and spectroscopic parameters over a series of observations. Darkening time was shown to be consistent with an ice/rock mixing ratio of 99% ice to 1% sediment across the northern hemisphere study sites (Byrne et al. 2009). For comparison, terrestrial sublimation polygons in Beacon Valley, Antarctica, form in a buried glacier-ice substrate that is c. 97% water ice and c. 3% sand and rock (Marchant et al. 2002). Typically, interest in the LDM focuses on the presence of the massive, excess subsurface ice deposits – an interest based on the astrobiological importance of ice as a potential source of water and as a microbial habitat on Mars (Lederberg & Sagan 1962; Gilichinsky et al. 1992, 2007; Dickinson & Rosen 2003). Estimates for the ice content of the LDM suggest that it represents a reservoir of c. 3.9  105 km3 of ice (Head et al. 2003; Levy et al. 2010a). This is approximately one-tenth the volume of the current, residual polar caps (Smith et al. 1998; Zuber et al. 1998). Estimates of the

volume of ice in the LDM are most strongly affected by the spatial extent of the deposit which is well constrained by surveys of image data, and are secondarily affected by estimates of LDM thickness and by the mixing ratio of ice to dusty debris (Levy et al. 2010a). Turning these ice-reservoir calculations around, an estimate can be made of the volume of the lithic (primarily dust) component of the LDM. Using values reported by Head et al. (2003) and Levy et al. (2010a), if LDM deposits span c. 5  107 km2 of the Martian surface, are 10 m thick and have a ratio of ice to lithic fines of 4 (80% ice to 20% dust, assuming that some regions consist of pore-ice permafrost rather than the nearly pure ice observed by Byrne et al. 2009), then the LDM represents a global deposit of c. 1  105 km3 of dust. This is a global layer over half a metre thick and suggests that the LDM may represent a major sedimentary deposit on Mars. The thickness of this unit is particularly interesting given the extremely slow erosion rates observed at the Mars Pathfinder landing site: as little as 0.01 – 0.04  1029 m per year (Golombek 1999). In summary, the Martian latitude-dependent mantle (LDM) may represent a truly unique coldclimate, sedimentary landform in planetary permafrost science. This young, massive ice deposit is globally distributed at high latitudes and is the substrate in which a wide range of thermal-contractioncrack polygons form (Mustard et al. 2001; Head et al. 2003; Milliken et al. 2003; Mangold 2005; Levy et al. 2010a). The LDM is the permafrost layer underlying the eroded surface of Martian gullies (Levy et al. 2009b), and may be the source of gully sediments and even of some gully meltwater (Dickson & Head 2009; Levy et al. 2010b). The latitude-dependent mantle is the unifying substrate in which recent Martian permafrost landforms develop.

Regional-scale sedimentary processes: thermal-contraction-crack polygons Striking networks of tessellated, patterned ground are abundant at Martian middle and high latitudes (polewards of c. 308) (Mellon 1997; Malin & Edgett 2001; Seibert & Kargel 2001; Mangold 2005; Kostama et al. 2006; Mellon et al. 2008; Levy et al. 2009a) (Fig. 3). The relative importance, or even presence, of periglacial (freeze-thaw) sorting of sediments in Martian permafrost terrains is a subject of vigorous and ongoing debate. Some form of wet active-layer sorting is suggested by Balme et al. (2009), while a dry cryoturbation mechanism is preferred by Mellon et al. (2008, 2009b) and Heet et al. (2009). Dry, stable and minimally sorting processes are preferred by Levy et al.

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(2008, 2009c, 2010a). Here, we focus on small-scale (,c. 25 m diameter) thermal-contraction-crack polygons, a class of unsorted permafrost features diagnostic of ice-rich permafrost in terrestrial polar environments. By virtue of the processes involved in their formation, thermal-contraction-crack polygons represent a unique depositional environment for sediments in Martian permafrost terrains. Thermal-contraction-crack polygons form through climate- and substrate-dependent mechanisms. As a result, they can be used as markers of microclimate history and permafrost thermal conditions (Black 1976; Marchant & Denton 1996; Marchant & Head 2007; Levy et al. 2010a). Thermal-contraction-cracks form in ice-rich permafrost as it undergoes thermal contraction in response to cooling temperatures. When thermal tensile stresses at or near the ground surface exceed the tensile strength of the frozen ground fractures form orthogonal to the ground surface (the cooling plane) (Lachenbruch 1961, 1962; Mellon 1997; Plug & Werner 2001, 2002; Maloof et al. 2002). As fractures propagate parallel to the frozen ground surface they intersect to form the eponymous ‘thermalcontraction-crack polygons’ (Lachenbruch 1961, 1962; Plug & Werner 2001), forming closed polygonal shapes in map view. The size and shape of thermal-contraction-crack polygons is determined by complex interactions between ice content, cooling history and other mechanical properties of the soil, and is the subject of ongoing investigations (Lachenbruch 1961, 1962; Plug & Werner 2002; Mellon et al. 2008, 2009a). What makes thermalcontraction-crack polygons interesting as sedimentary features is the next step in polygon formation. Once fractures open in a frozen ground surface, infilling of fractures may occur as overlying material enters the fracture. Infilling processes are diagnostic of the climate conditions in which the fracture formed (Marchant & Head 2007). Repeated fracturing along the same plane of weakness in the frozen ground, coupled with repeated infilling, can lead to the formation of wedges of material underlying polygon troughs. Different permafrost climate conditions leave unique wedge structures in the stratigraphic record (Pewe 1963, 1974; Murton 1996; Murton & Bateman 2007). In warmer and wetter permafrost environments, in which a seasonally saturated active layer forms, meltwater can percolate through overlying peat, vegetation or regolith, filling thermal-contraction cracks with relatively pure liquid water that subsequently freezes, forming ice-wedge polygons (Leffingwell 1915; Lachenbruch 1962; Berg & Black 1966; Black 1982; Washburn 1973; Sletten et al. 2003; French 2007; Marchant & Head 2007). In cold and arid environments, in which either an active layer does not form or in which the active layer is water-free (‘dry’)

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(Bockheim et al. 2007), sand particles and other fines can winnow into open fractures from above, forming sand-wedge polygons (Pewe 1959; Berg & Black 1966; Murton et al. 2000; Sletten et al. 2003; Marchant & Head 2007; Murton & Bateman 2007). Some polygon-forming environments are too cold or too arid to regularly experience typical, widespread and saturated active-layer conditions, but do experience occasional, localized inputs of liquid water to the subsurface. This can occur, for example, due to ephemeral snowbank accumulation and melting. Alternating inputs of water and dry sediment to thermal-contraction cracks form composite-wedge polygons (Berg & Black 1966; Murton 1996; Ghysels & Heyse 2006). Finally, in select permafrost environments that are too cold to generate a seasonal active layer and that have abundant excess ice (ice exceeding available pore space) in the subsurface, sublimation polygons may form as ice sublimates preferentially along thermalcontraction cracks and is partially replaced by sieved fines winnowed from overlying tills (Marchant et al. 2002; Kowalewski et al. 2006; Levy et al. 2006; Kowalewski & Marchant 2007; Marchant & Head 2007). Sublimation polygons forming in buried or stranded glacier ice are most common in the Antarctic Dry Valleys (Marchant et al. 2002). The response of permafrost to wedge growth is diagnostic of polygon type (Marchant et al. 2007). Active ice-wedge and sand-wedge polygons commonly form broad, raised shoulders and low-lying centres as the ice-cemented soil adjacent to the wedges re-expands as the permafrost warms in summer. The increased subsurface volume (the ice- or sand-wedge) is accommodated by the wedge-adjacent permafrost deforming upwards towards the free surface at the ground– atmosphere interface (MacKay 2000). Ice-wedge polygons may become high-centred in response to thermokarst (melting) modification of the ice-wedge and rapid drainage of surrounding soils (MacKay 2000). Composite-wedge polygons may have slightly raised shoulders or may be flat-lying (Berg & Black 1966; Murton 1996; Ghysels & Heyse 2006). Active sublimation polygons are characteristically convex-up with high, domical centres that are underlain by relatively stable ice, surrounded by depressed troughs that lack raised shoulders (Marchant et al. 2002; Kowalewski et al. 2006; Kowalewski 2008). In the case of sublimation polygons, the addition of winnowed sediment to the subsurface is balanced (and often exceeded) by the preferential sublimation of buried ice, resulting in low-troughed and high-centred polygons (Marchant et al. 2002; Kowalewski et al. 2006). On Mars, thermal-contraction-crack polygons can be identified based on a range of morphological characteristics using both orbital and lander image

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data. Lander data is analysed by Heet et al. (2009), Mellon et al. (2009a), Levy et al. (2009a) and Smith et al. (2009). Orbital data analysis can be found in Mellon (1997), Malin & Edgett (2001), Seibert & Kargel (2001), Mangold (2005), Kostama et al. (2006), Levy et al. (2008, 2009a) and Mellon et al. (2008). Levy et al. (2009a) use multiple characteristics to identify thermalcontraction-crack polygons on Mars and to distinguish them from other polygonal landforms. These criteria include: (a) network morphology (indicating multiple episodes of fracturing), (b) polygon microtopography (showing raised rims or high, domical centres), (c) diameter (c. 25 m or smaller and comparable to terrestrial examples), (d) presence in latitude bands where active thermalcontraction cracking is modelled to presently occur (Mellon 1997), (e) presence on preferentially oriented slopes (which affects the depth and stability of ground ice), (f ) surface age (most polygon networks on Mars are very young, ,c. 2 Ma), (g) particle size and distribution (indicative of sublimation-driven rolling or slumping), (h) bedrock presence (permafrost-related polygons form in unconsolidated sedimentary units and not in bedrock), (i) associated landforms (suggesting permafrost processes, for example, scalloped terrain and mantling units) and ( j) albedo (polygons tend

to form in relatively low-albedo units). On the basis of multiple surveys of high-latitude datasets, thermal-contraction-crack polygons have been shown to be ubiquitous polewards of c. 508 latitude and to be very common polewards of c. 308 (Milliken et al. 2003; Mangold 2005; Levy et al. 2009a). What does the morphology of Martian thermalcontraction-crack polygons suggest about polygon type? Overwhelmingly, Martian polygons are highcentred with depressed boundary troughs and flat or domical interiors (Levy et al. 2009a). They (a) show evidence of possible orientation-dependent slope asymmetry indicating massive (structureless) excess ice (exceeding pore space) (Mangold 2005; Levy et al. 2009a), (b) are observed along with a range of landforms suggesting stable (unchurned) and sublimation-driven surface processes, and (c) do not commonly show morphological indications of melting typical of thermokarst-modified icewedge polygons (Levy et al. 2010a). When considered together, this evidence suggests the dominance of active sublimation polygons or sediment-starved sand-wedge polygons on Mars (Fig. 5) (Mangold 2005; Mellon et al. 2009a; Levy et al. 2010a). Polygons and fractures with raised shoulders are present in some locations on Mars (Lefort et al. 2009; Levy et al. 2009d ). However, the lack of definitive, accessory landforms suggesting

Fig. 5. Map showing the distribution of gullies, thermal contraction-crack polygons and gully-polygon systems across the Martian surface. Despite the global persistence of permafrost conditions on Mars, permafrost landforms such as gullies and polygons are latitudinally clustered. Both gullies and polygons are largely confined to the latitude-dependent mantle. Gullies are more common in the dissected mantle and polygons are common in both the dissected and the continuous portions of the LDM.

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active-layer conditions (Dundas & McEwen 2009; Lefort et al. 2009) and a climate history incompatible with recent active layer formation (Kreslavsky et al. 2008) suggest that these polygons are not ice-wedge polygons. Rather, evidence suggests that they are either sand-wedge polygons or, more likely, sublimation polygons. In the sublimation polygon case, they may have undergone topographic inversion as the once-stable ice in the high centre of the polygon collapsed due to ongoing sublimation (Levy et al. 2009d ). Liquid water volumes sufficient to produce ice-wedges seem not to be a major agent of geomorphologic work in recent thermal-contraction-crack polygon terrains. Where both gullies and polygons are present, the abundance of high-centred polygons suggests that, as modelled by Heldmann et al. (2005), liquid water involved in gully formation freezes and/or evaporates and does not initiate water-driven cryoturbation. This may produce composite-wedge polygons (Levy et al. 2009b). Rather than freezethaw phase transitions, sublimation appears to be the dominant player in determining the depth, stability and morphology of ice-related landforms on the Martian surface (Schorghofer 2007; Mellon et al. 2009a; Levy et al. 2010a; Sizemore et al. 2010).

Local-scale sedimentary processes: gullies Gullies are not only one of the most interesting features of the Martian surface, but also one of the most enigmatic (Fig. 6). Gullies are a class of young Martian landform that is typically composed of a recessed alcove, one or more sinuous channels and

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a fan or apron downslope of the channel mouth (Malin & Edgett 2000, 2001). Gullies are typically c. 1–2 km long from alcove apex to fan (Malin & Edgett 2000, 2001). Erosion, transport and deposition of particulate material are hallmarks of sedimentary processes, but how did the Martian gullies form given the extreme cold and aridity of the Martian surface? Several mechanisms have been proposed for the formation of gullies, ranging from water-free sediment flows (Treiman 2003; Shinbrot et al. 2004; Pelletier et al. 2008) to water-lubricated debris flows (Malin & Edgett 2000; Costard et al. 2002; Hartmann et al. 2003; Pelletier et al. 2008; Levy et al. 2010b), to water-rich sediment transport (fluvial or hyperconcentrated flow) and alluvial deposition (Heldmann & Mellon 2004; Heldmann et al. 2005; Dickson et al. 2007; Head et al. 2008; Dickson & Head 2009; Levy et al. 2009b). Given that mass movement of sediment appears to have occurred in gullied regions, or to be occurring (Malin et al. 2006), gullies represent an important sedimentary process in the Martian permafrost system. Since their discovery, evidence has continued to accumulate that implicates a water-related origin for Martian gullies. Malin & Edgett (2001) originally suggested that liquid water played a role in the formation of Martian gullies. Malin & Edgett (2001) argued for gully formation through a combination of overland flow, sapping and wet debris flow on the basis of channel morphology. They noted that gully channels are commonly sinuous, are branched or show anastomosing relationships, are commonly

Fig. 6. Relationships between gullies and thermal-contraction-crack polygons in Martian ‘gully-polygon systems.’ (a) Gully channels and alcoves incise polygonally patterned mantle material (portion of PSP_002054_1325). (b) Gully fan deposits overprinting thermal contraction-crack polygons. Note fine-scale polygonal patterning of the fan surface (portion of PSP_002368_1275). (c) Gully channels ‘annexing’ thermal contraction crack polygons (portion of PSP_001938_2265).

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flanked by levees and commonly display superelevated banking and incision. Likewise, Malin & Edgett (2000) report that gully fan morphologies are strikingly similar to alluvial fan morphologies on Earth, showing evidence of diverging lineations radiating from the channel mouth, lobate margins and distal thinning. More recent observations of gully morphology using image data of higher spatial resolution reveal additional features of some gullies that are more compatible with sediment transport by liquid water than by dry, granular flow. For example, Mangold (2010) used photogrammetric techniques to determine that asymmetries in levee height at bends in Martian gully channels are most consistent with a fluid viscosity typical of terrestrial water-lubricated debris flows, and would not be likely to emerge due to dry mass wasting. Likewise, high-resolution analyses of other gully deposits have revealed the presence of cutbanks, terraces, cut-off channels, incised fans and channel-fill deposits (Schon & Head 2009; Schon et al. 2009a) – all features largely consistent with alluvial-fan processes involving multiple episodes of fluvial activity (Fig. 6). If Martian gullies formed as a result of sediment transport by liquid water, what is the source of that water? Surface temperatures on Mars have been shown to only rise above the triple-point temperature of water (273 K) for a few tens of days of every Martian year and only at latitudes ,c. 308 (Haberle et al. 2001). In an apparently paradoxical relationship to this climate pattern, gullies are most abundant between c. 30 –558 latitude (Heldmann & Mellon 2004; Balme et al. 2006; Dickson et al. 2007; Dickson & Head 2009). Malin & Edgett (2000, 2001) hypothesized that liquid water involved in gully formation was stored in confined, subsurface, geothermally warmed aquifers that periodically ruptured the overlying permafrost, allowing water to flow over the surface. The flow of erupted, solute-free water entraining clastic materials was modelled by Heldmann et al. (2005). They were able to demonstrate that, while liquid water is not presently stable on the Martian surface, liquid water could flow a similar distance to the length of gully channels while undergoing evaporation and freezing. That is, liquid water is meta-stable on Mars (Hecht 2002). However, several observational idiosyncrasies of Martian gullies appear to be at odds with the confined aquifer model (for example, Fig. 6). These are: (a) gullies form at a range of elevations along Martian slopes and not always along an exposed bedrock (confining) layer; (b) gully channels and alcoves typically reach the apex of the slope on which they form, commonly meeting neighbouring gully channels and alcoves across a narrow topographic divide or at the apices of crater central

peaks; (c) gullies are absent from the lowest regions of the Martian surface, such as the Hellas Basin, where groundwater would be most likely to outcrop (Balme et al. 2006); (d) gullies are exclusively present polewards of 258 latitude and are most common between c. 30 –558 latitude (Dickson et al. 2007; Dickson & Head 2009); (e) gullies show a strong slope orientation preference, appearing on polewards-facing slopes at low latitudes (c. 25–408), equator-facing slopes at middle latitudes (c. 40– 558) and polewards-facing slopes at high latitudes (.558) (Christensen 2003; Dickson & Head 2009); and (f ) radar observations of gully sites originally inferred to have formed from groundwater release show no strong subsurface radar reflections indicative of the presence of liquid water reservoirs (Nunes et al. 2010). In light of these observations, a consensus is emerging among some Martian gully researchers that a top-down melting of near-surface ice and/or surface frost and snow may better account for the generation of gully meltwater than a groundwaterrelease mechanism (Costard et al. 2002; Christensen 2003; Dickson & Head 2009; Williams et al. 2009). This surface-ice melting process is consistent with recent modelling results showing that in some microclimates, during periods of high Martian orbital obliquity, water ice can accumulate by both atmospheric deposition (including frost emplacement and/or snowfall) and melt at the Martian surface. This emplacement and melting of surface ice occurs in protected gully alcoves, and can produce ephemerally present surface runoff sufficient to erode gullies at the precise latitudes, elevations, slopes and orientations at which they are observed (Costard et al. 2002; Hecht 2002; Williams et al. 2008, 2009). It follows that morphological observations in permafrost regions can be used to differentiate between the two primary gully formation models: top-down melting and confined aquifer. Martian gully deposits are commonly found on surfaces modified by thermal-contraction-crack polygons (see previous section and Fig. 7). Thermalcontraction-crack polygons form in soil surfaces that are not merely frozen (mean annual temperature ,273 K) but that are ice-rich. As a result, they are effectively impermeable on the timescales of water freezing/evaporation at the Martian surface (Heldmann et al. 2005; Levy et al. 2009b). Gully channels and alcoves commonly cross-cut thermalcontraction-crack polygons. In some locations gully fans overprint polygons and in other locations gully fans are cross-cut by thermal-contractioncrack polygons (Fig. 7). Levy et al. (2009b) interpret these stratigraphic relationships to indicate that widespread ice-cemented permafrost pre-dated the formation of the gullies, persisted through

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Fig. 7. Characteristic examples of Martian gully morphology indicative of multiple episodes of fluvial transport (modified from Schon & Head 2009). (a) Classic Martian gully on the interior slope of a crater. Gully elements – alcove, channel and fan – are illustrated. The gully extends all the way up to the crater rim. The sunlit side of the topographic divide is saturated with bright pixels in this contrast stretch (portion of PSP_001882_1410). (b) Sinuous and anastomosing gully channels (small arrows) and an eroded longitudinal bar (long arrow) downslope from a spur (modified from Schon & Head 2009, fig. 1). (c) Cut-banks, channel terraces and braided channels (portion of PSP_006593_1470). (d) A gully fan eroded by channels formed in subsequent flow events. A smaller channel has been abandoned and stranded at a higher topographic level from the large channel featured at image centre (portion of PSP_002292_1490).

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the period of gully formation and endures to the present. This suggests the continuous presence of metre-thick effectively impermeable material underlying gullies in these locales. Water pressures exceeding the c. 2 MPa required to fracture icerich permafrost (Mellon 1997) would result in catastrophic eruption of water sourced by a confined aquifer, likely producing dramatic scouring in gullies that is not observed. Accordingly, Levy et al. (2009b) favour a top-down melting mechanism. The erosion of ice-rich, thermal-contractioncracked permafrost during gully alcove formation suggests that some of the sediments involved in gully fan deposition are sourced in the underlying permafrost substrate (the LDM). This implies that ice-cemented, polygonally patterned permafrost may represent another critical element in the sedimentary system operating in the Martian cold desert.

Conclusions The above examples provide an introduction to the range of permafrost landforms currently being explored on the surface of Mars and an illustration of some of the key processes in the geological development of Martian cold-desert landforms. Martian permafrost terrains represent the extreme cold and dry end of the wet-to-dry permafrost landform spectrum on planetary surfaces. The development of thermal-contraction-crack polygons on Mars appears to be largely incumbent on the presence of LDM deposits that feature excess ice in the shallow Martian subsurface and on cold, dry conditions under which sublimation is the dominant phase transition, rather than melting. Melting of near-surface snow, frost and/or ground ice during the most recent several million years on Mars has been largely confined to protected microclimates in gully alcoves, from which flows of water-borne sediment have been transported into their present configuration, forming gullies. Connecting these exceptionally young permafrost deposits to the longer term rhythms of climate change on Mars remains a topic of great interest to the geomorphologists, climate modellers and astrobiologists who will help guide the next generation of exploration in Martian polar regions. Morphological analysis of fine-scale Martian features was made possible by the efforts of the HiRISE and Mars Reconnaissance Orbiter teams. Thanks to C. McKay and an anonymous reviewer for their comments. This work is partly supported by grant ANT-0851965 to JSL by the Antarctic Organisms and Ecosystems Program in the Antarctic Sciences Division of the National Science Foundation.

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Rethinking weathering and pedogenesis in alpine periglacial regions: some Scandinavian evidence C. E. THORN1*, R. G. DARMODY2 & J. C. DIXON3 1

Department of Geography, University of Illinois at Urbana-Champaign, 607 S. Mathews Ave., Urbana, IL 61801, USA 2

Department of Natural Resources and Environmental Sciences, University of Illinois at Urbana-Champaign, 1102 S. Goodwin, Urbana, IL 61801, USA 3

Department of Geosciences, University of Arkansas at Fayetteville, Fayetteville, AR 72701, USA *Corresponding author (e-mail: [email protected]) Abstract: Research in the Arctic and alpine regions of NW Sweden and south-central Norway reveals an active chemical-weathering regime: in Ka¨rkevagge, Swedish Lapland, pyrite weathers to produce an acidic environment with vigorous chemical weathering and active pedogenesis as indicated by locally well-developed Spodosols; in the Jotunheimen Mountains of Norway, soil horizons form slowly but cobbles develop weathering rinds within c. 250 years. Such results indicate that chemical weathering and pedogenesis may be quite advanced in alpine periglacial regions and is just one illustration that such regions may be profitably scrutinized using geomorphological and pedological conceptual frameworks other than the traditional periglacial framework.

When Łozin´ski (1909, 1912) introduced the term ‘periglacial’ he constrained it in a number of ways (Thorn 1992; French 2008). The most notable of these is: temporally (the Pleistocene); geographically (proximity to ice sheets); lithologically (coarse-grained sandstones); and by process (freeze –thaw weathering). These constraints largely evaporated as the term became broadly used to describe cold, but presently unglaciated, regions. However, the close association between the term periglacial and the notion of regional domination by freeze –thaw weathering lingered on throughout much of the heyday of periglacial geomorphology. Weathering is certainly one of the most fundamental geomorphological concepts and its role anywhere is central to a researcher’s view of how a particular portion of the landscape develops. In recent decades serious questions have been asked of the freeze– thaw weathering concept. For some the process itself remains in question (Hall 2006); for others it remains an established process (Matsuoka 2001; Murton et al. 2006). Whatever the researcher’s position it is likely that he/she no longer views periglacial weathering as entirely dominated by the freeze –thaw process, or at the very least recognizes that there is ample evidence to demonstrate that chemical and biological weathering are present in periglacial regimes. Among the first to suggest the significance of chemical processes in a periglacial

environment was Rapp (1960). More recently weathering studies in general (Viles 1995) and periglacial weathering studies specifically (Etienne 2002) have focused upon biological weathering, particularly the role played by lower life forms such as lichens. Today the term ‘periglacial’ embraces a much more diverse set of environments than when it was originally invoked; indeed French (2007, 2008) suggests demarcation of at least five categories of periglacial regime. We do not dispute that assertion; rather, we accept it as recognition of the disparity (probably disparities) that exist(s) among regions commonly lumped together as periglacial. Acceptance of French’s categorization means that it is unlikely that any field-based concepts are valid across the entire set of categories. For this reason, we emphasize that our remarks pertain only to ‘alpine periglacial’ Scandinavia. Although one of our sites, Ka¨rkevagge, is also technically an Arctic as well as alpine environment, it clearly does not meet French’s (2007) ‘High Arctic’ criteria. As a matter of brevity we refer to both sites as alpine. In more general terms, both these research areas should be considered as mild periglacial environments. The present paper is a summary of two multiyear studies of chemical weathering and pedogenesis undertaken in Scandinavia. In Ka¨rkevagge

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 183–193. DOI: 10.1144/SP354.11 0305-8719/11/$15.00 # The Geological Society of London 2011.

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(688240 N, 188200 E), Swedish Lapland, approximately a decade of research was devoted to attempting to understand the chemical weathering and, to a lesser extent, pedogenic processes that might underpin Rapp’s (1960) claim that the environment was dominated by chemical processes. At Storbreen, Norway (618350 N, 88210 E), the focus was primarily upon an assessment of the nature and rate of pedogenesis, although some rock weathering studies were also conducted. Here the work was built upon earlier research by Mellor (1987), Messer (1988) and others as well as upon numerous studies by John A. Matthews and co-workers (Matthews 1992). Instead of providing an exhaustive literature review (not possible because of limitations of space) we provide a summary of our own work, the individual original elements of which include such literature reviews. We also attempt to cite a series of entry points to the work of others which both supports and contradicts our own. Our overall objective is to demonstrate that in Scandinavian alpine periglacial environments the entire gamut of chemical weathering and pedogenic processes may be expected, sometimes with quite high rates. We have no information to offer on freeze –thaw weathering and while we readily accept that biological weathering may be important in our study areas, we have only limited data to offer on that matter. Given our stated objective we view chemical weathering and pedogenesis not from a ‘classical’ periglacial perspective (i.e. emphasizing domination by temperature and/or freeze-thaw oscillations specifically) but rather from a perspective more closely associated with temperate environments (i.e. tending to emphasize moisture supply). We suggest that this opens up the researcher’s view of things, permitting a much wider view of periglacial weathering itself as well as facilitating comparison with other climatic zones.

Weathering frameworks In his seminal study of weathering, Yatsu (1988, p. 1) asserts that ‘The clear understanding of a subject needs an exact definition of concepts . . .’ This is clearly the case, but here we use the widely accepted division of weathering into physical (mechanical) and chemical components. Note that while some researchers would add a biotic component, others would simply claim that both physical and chemical weathering may frequently be affected by biotic agents. Having established what is possible, it becomes essential to determine the rate(s) of any such process(es) and, eventually, relative rates among processes. Brunsden (1979, p. 75) notes ‘It is possible to construct world-wide distributions of weathering based only on the assumptions that the rate of

chemical weathering increases with temperature and water availability, and that the physical processes increase at lower temperatures’. It is the latter assumption that fundamentally underpins the claim for a distinctive periglacial weathering regime.

Mechanical (physical) weathering The theoretical underpinning of freeze –thaw weathering is surprisingly sparse. Work by Walder & Hallet (1985) and Hallet (2006) is probably the most satisfactory. However, empirical work both in the laboratory and field is much more plentiful. The primary shortcoming of empirical work is that it is almost impossible to produce an ‘experiment’ in which only the freeze –thaw process is active. Similarly, it is not possible to deal with large specimens in the laboratory or to mimic most natural moisture regimes realistically. Field studies in and around Storbreen led Berrisford (1991) to argue for accelerated mechanical fracturing beneath late-lying snow patches, while Matthews et al. (1986) and Shakesby & Matthews (1987) have argued for accelerated freeze –thaw weathering specifically around some lake margins in the Jotunheimen. Berrisford constrains his position on freeze –thaw weathering very carefully, while Matthews and co-workers appear to accept fragment angularity as a fundamental indicator of freeze– thaw weathering. The latter position seems difficult to defend because, in the absence of any one-to-one process-form link, the end product (angular fragments) cannot be taken as a definitive indicator of the formative process. Further afield, Matsuoka (2001) has also conducted field monitoring to investigate the products of freeze –thaw weathering. However, a criticism is of his measurements is that they are focused on the product (coarse debris) and not the process. Recent laboratory research by Murton et al. (2006) does appear to demonstrate rock shattering at the base of the active layer. In short, field research in alpine Scandinavia suggests that annual freezing coupled by excessive water availability may produce localized mechanical shattering. However, as fragment angularity may be the product of several factors (McGreevy & Whalley 1982) it alone cannot be invoked as a definitive indicator of freeze – thaw weathering.

Chemical weathering A number of comprehensive reviews (Loughnan 1969; Carroll 1970; Yatsu 1988) indicate that the underpinnings of chemical weathering are largely uncontroversial although it is clear that biological inputs may produce complex reactions. Chemical weathering processes such as oxidation, hydration

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and hydrolysis operate under the same rules in all environments; low temperatures however, such as those that characterize periglacial regions, may slow some processes down. Conversely, Williams (1949) has shown that carbon dioxide is more soluble in cold water and therefore there should be more carbonic acid available for weathering. Establishing an adequate framework for chemical weathering that embraces landscape variability is a challenge. Brunsden (1979) invokes Jenny’s (1941) well-known soil state factor equation. In broad terms, such a framework is applicable to both rock weathering and pedogenesis, although the organic influence is likely to be significantly more important in pedogenesis. In any environment, water is crucial for both chemical weathering and pedogenesis. In many circumstances, periglacial environments can be dry either because water is ice and therefore not chemically reactive or the landscape positions may be exposed to windy, desiccating conditions. This is certainly the case for the polar deserts of the world. The initial delivery of moisture to many periglacial regions is usually as snow. Typically, this is rapidly and sharply redistributed as seasonal snowcover meaning that many periglacial surfaces are commonly a sharply defined mosaic of exceedingly wet or dry fragments interspersed among a wide range of intermediate moisture conditions (Thorn 1978). The alpine regions of Scandinavia fall within this category. In addition, some periglacial soils are very coarse-textured and have low moisture-holding capacity, which makes them excessively drained and dry much of the time. For example, much of Iceland (a humid periglacial environment) may best be regarded as an edaphic desert. Finally, different species in the boreal forest regions of the sub-Arctic of the northern hemisphere influence snow distribution patterns at the micro-level and, ultimately, soil moisture temperature. There is considerable evidence from Scandinavia to support the presence of chemical weathering in its alpine periglacial regions. Two broadly convergent investigations, Andre´ (1995, 2002) in the Ka¨rkevagge region and Nicholson (2009) in the Storbreen region, have produced results in general agreement. Although there are some differences between the two studies, their consonance is much more obvious. Both Andre´ and Nicholson found post-glacial weathering rates to be low, both question the importance of freeze –thaw weathering and both emphasize the significance of biochemical weathering. The widespread occurrence of roches moutonne´es and similar forms in both regions provides an opportunity to integrate results from the two studies and make a series of general points. Roches moutonne´es appear to have experienced little post-glacial

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degradation of form in general, although some are fractured. As roches moutonne´es are generally convex, and consequently shed water, we would expect them to have experienced little weathering. What weathering has occurred appears to Andre´ and Nicholson to be associated with the widespread biochemical films found on them. Some of these features are heavily fractured although they retain their overall morphology, but the fractured forms are obviously within the same weathering environment as the unfractured forms. Consequently, it seems more appropriate to assign their fractured nature to inherent rock properties rather than to the weathering regime they are experiencing. It is undoubtedly true that once the fractures appear they may be exploited differentially, but Andre´ and Nicholson’s results are less uniform with respect to this specific matter. While Andre´ and Nicholson found generally slow rates of weathering over the c. 10 ka that have elapsed since deglaciation, McCarroll (1990) demonstrated that the chemical weathering of individual minerals can be quite rapid at Storbreen over a similar period. However, he also found the rates to differ significantly between unburied and buried contexts. The discrepancy led him to infer that the role of lichens might be significant in weathering processes and rates. Achim Beylich and co-workers (Beylich 2000; Beylich et al. 2004a, b, 2006) have pursued studies in eastern Iceland and at Latnjavagge which is in the immediate vicinity of our own work in Ka¨rkevagge. A primary focus of this work has been to examine the comparative contributions of various processes to the overall development of the drainage basins studies. A number of interesting findings have emerged; here we simply note some highlights. In the 10–12 ka since deglaciation, the periglacial regime has done very little to modify the larger landscape. At present, the fluvial and slope systems are largely decoupled while, within the slope system itself, water-driven processes rank first, followed by geochemical denudation. In Latnjavagge, Swedish Lapland, Beylich et al. found substantial variability of sediment and solute yields within the basin with such factors as late-lying snowcover, frozen subsurfaces and regolith depths having detectable impacts. Overall, chemical denudation slightly exceeded mechanical fluvial denudation. It is worth noting that our own work in Ka¨rkevagge was focused upon explaining Rapp’s (1960) claim for domination by chemical processes. This claim, highly unorthodox at the time, was underpinned by extensive work on mechanical processes but only skeletal chemical research. In contrast, the work of Beylich and co-workers is focused upon a ‘representative’ basin. In short, our results may well indicate what is possible chemically while Beylich et al.’s work

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may reflect what is regionally representative. Nevertheless, even Belyich et al.’s findings are a far cry from the traditional perspective emphasizing physical processes. Dixon & Thorn (2005) and Dixon et al. (2008) both represent attempts to view the implication of such findings in a broader perspective.

Biological weathering As noted earlier we are reluctant to make a tripartite, rather than bipartite, subdivision of weathering. However, a number of researchers have focused their weathering studies on biological inputs; consequently, we recognize this focus while again noting that weathering is fundamentally mechanical or chemical. Of course, such mechanisms commonly interact and are driven by biological inputs. For example, it is clear that lichens have both mechanical and chemical impacts on rock surfaces. Viles (1995) has championed the role of biological agents in weathering and at this time it is a widely accepted idea. It is of interest to note that much of her research has been conducted in deserts; clearly, many periglacial surfaces are also deserts albeit cold ones. Research from Storbreen (McCarroll & Viles 1995) attests to the rapidity of lichen-induced weathering, a perspective supported by work by Matthews & Owen (2008). Despite the claim by McCarroll & Viles that biological weathering is more significant than chemical weathering, comparative data is still lacking. Until it is available both processes are best viewed as important. This is likely broadly true in periglacial areas, as has been suggested by Etienne (2002). In its most generalized form the point is strongly supported by both Andre´ (2002) and Nicholson (2009) who invoke the widespread relative significance of biochemical processes vis-a`-vis freeze –thaw processes while noting the overall slowness of the periglacial weathering regimes they investigated.

remember that it is ground and not air climate that is really important in pedogenesis (Geiger 1950; Thorn et al. 2002; Schaetzl & Anderson 2005). Cryopedoturbation (mechanical disruption of soil horizons by freezing and thawing) is generally stronger in terms of frequency, duration and intensity in periglacial regions than in soils of temperate mid-latitudes, and clearly of no significance in soils of tropical climates. Depending on additional requirements, cryopedoturbation may also produce surficial forms such as patterned ground, frost polygons and solifluction lobes as well as subsurface profile features such as distorted horizons (Fig. 1), vesicular voids and frost-heaved coarse fragments. The other climatic feature unique to very cold soils is permafrost, which serves as an aquaclude inhibiting downwards flow of water, resulting in the hydromorphic soils commonly found in periglacial environments. Periglacial soils tend to be high in organic matter, particularly in wetter areas, because of both poor drainage and slowed decomposition due to cold temperatures. Wetness can also be enhanced by low rates of evapotranspiration. Conversely, if the soils are sandy or exposed to windy conditions, dryness can be a big factor. As a result, some periglacial soils exhibit similarities to desert soils including cryptogam crusts and vesicular structure.

Time Time as a factor in alpine periglacial pedogenesis is important in as much as many, but certainly not all, periglacial landscapes tend to be young. Where soils are young their youthfulness makes the nature and properties of the soil parent materials relatively more important than is the case with older soils. Weatherable minerals may be abundant in such periglacial soils. For example, the youthful moraines at Storbreen contain hornblende and biotite. However, there is evidence that mineral transformations are

Periglacial pedogenesis Jenny’s (1941) five factors of soil genesis, climate, time, organisms, relief and parent materials apply in periglacial regions as they do anywhere. Indeed, pedogenesis in periglacial environments is essentially the same as that found in less harsh environments. The same rules and processes apply but their rates, and consequently their importance, varies.

Climate Climate is the one factor that stands out in alpine periglacial regions because, by definition, the latter are cold regions. However, it is important to

Fig. 1. Aquic Cryorthent soil profile showing horizon distortions in a solifluction terrace, Ka¨rkevagge, Sweden.

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Fig. 2. Soil ‘poisoned’ by exudates from pyrite-rich boulder in Ka¨rkevagge, Sweden.

starting only 2–3 centuries since deglaciation (Darmody et al. 2005b). Another example is the bedrock exposed after deglaciation in Ka¨rkevagge, Sweden. There, pyrite (a mineral that is scarce in temperate climates) is responsible for acidifying soils due to its oxidation (Darmody et al. 2001). The resultant sulfuric-acid-contaminated soil is responsible for killing vegetation in areas where it is concentrated (Fig. 2). The nature or kind of clay minerals present has often been used as a relative age-dating metric in Quaternary research. Studies at both Ka¨rkevagge (Darmody et al. 2000b, 2004; Allen et al. 2001), and Storbreen (Darmody et al. 1987, 2005b; Darmody & Thorn 1997) illustrate that clay mineralogy is sensitive to both age and elevation while Egli et al. (2003a, b, 2007) and Mirabella et al. (2002) have demonstrated that parent material, aspect, climate, age and treeline all have distinguishable impacts on soil clay content in alpine periglacial areas. While working in the Okstindan Mountains of northern Norway, Ellis (1980) found chemical weathering rates to be low although he also found production of vermiculite and goethite in the same soils. Anderson et al. (2000) and Anderson (2007) have shown that age may be very important in the early stages of weathering of recently deglaciated terrain; they conclude that relatively high rates of carbonate dissolution and sulphide oxidation follow initial deglaciation and silica mineral weathering becomes dominant only after vegetation has become established.

Organisms Organisms as a factor in alpine periglacial pedogenesis generally have a lesser importance than in other environments, given the relatively short time they have to influence soil parent materials. However, cold-adapted plants, despite their slow growth, can still produce chelating agents that result in

Fig. 3. Typic Haplocryod soil profile from Ka¨rkevagge, Sweden (688250 14.230 N, 188170 50.820 E).

podzolization and prominent horizonation (Fig. 3). Pioneering plants, and those that comprise the cryptogram crusts which tolerate the extreme conditions that can exist in periglacial environments, are often nitrogen fixers. These early successional plants add N along with C to soils and produce fertile, dark-coloured A horizons, the first to be identified in the youngest of soils. In addition, the relative scarcity of burrowing animals, ants, worms, etc. limits faunal pedoturbation, and thus allows horizonation to continue where cryopedoturbation does not dominate.

Relief Relief in alpine periglacial environments is highly variable but, in general, the same weathering principles apply as in temperate soils. Erosion on steep slopes or slopes with solifluction activity can result in poorly expressed, buried or distorted soil horizonation (Fig. 1). Low-lying areas, as previously mentioned, typically have high organic matter content. Reducing conditions will predominate in these situations and organic soils will accumulate if the influx of mineral material is low enough.

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Parent materials Parent materials in general are a more important factor in pedogenesis in younger soils. The presence of easily weatherable minerals such as hornblende, biotite and pyrite make weathering rates of early stage pedogenesis more rapid. Soil texture of alpine periglacial soils is a highly important variable in as much as it controls soil moisture storage and fertility. In the alpine periglacial environments we have studied, the soils tend to be sandy with low moisture-holding capacity. This can impart moisture stress on plants unless the landscape position or the local ground climate keeps the soils moist.

Case studies: Ka¨rkevagge and Storbreen

Fig. 4. White precipitated Al compounds associated with natural pyrite oxidation in Ka¨rkevagge, Sweden.

Ka¨rkevagge Ka¨rkevagge (688240 N, 188200 E), a glacial trough in Swedish Lapland, was deglaciated c. 10 ka BP according to Andre´’s (1995) estimate of regional deglaciation or perhaps as much as 13 + 2.1 ka according to cosmogenic dating of some of the landscape elements within the valley (Darmody et al. 2005a). Valley walls expose horizontally bedded schists and interbedded marble units, some of which contain pyrite (Darmody et al. 2007). Pyrite oxidation is rapid by geological standards, providing a rapid burst of intense weathering in disturbed landscapes (such as surface mines) but disappearing and going largely un-noted in mature landscapes. The pyrite oxidation process in Ka¨rkevagge produces acidic waters of 4.5 pH, which vary around the valley depending upon bedrock type (Campbell et al. 2001). This acidity produces a series of clear indicators of chemical weathering in the valley: (a) individual boulders that are ‘rotten’ and kill/ poison the surrounding vegetation (Fig. 2) (Rapp 1960; Darmody et al. 2001); (b) a series of whitecoated stream beds along the valley, the coatings comprised of predominantly amorphous basaluminite and related minerals (Fig. 4); and (c) secondary sulphate minerals such as jarosite and gypsum (Fig. 5) (Darmody et al. 2007). This assemblage attests both to intense chemical weathering in Ka¨rkevagge and the significance of ground microenvironmental variability. From 1994 to 2004 a ‘potential’ weathering study was run across a series of microenvironments within Ka¨rkevagge and on the abutting ridge crest to the east. The study contained two elements, namely: (a) machine-polished disks of dolomite, granite and limestone buried at depths up to 60 cm (Thorn et al. 2006a); (b) mesh bags of crushed granite and limestone pinned to the surfaces immediately above the buried disks (Dixon et al. 2001; Thorn et al. 2006b). Salient results include the fact that the weathering

order among the disk types was limestone . dolomite . granite with granite disks exhibiting weathering rates too slow to be statistically distinguishable among sites. However, the two other rock types exhibited statistical distinguishable variability among pH and drainage classes, but not by soil horizon or burial depth. Much of this kind of variability was caught by a vegetation classification wherein heath categories (xeric environments) weathered slowly relatively to willow-covered surfaces (acidic and wet). Comparison of the surficial bags revealed them to exhibit spatial patterns of weathering generally correlative with the buried disks, although deterioration of the mesh bags limited the statistical comparison. An important aspect of the study was the determination that mass loss due to chemical weathering differed by as much as an order of magnitude in percentage terms over a five-year period. Furthermore, the second five-year period yielded significantly higher rates of loss than the first five-year period. In the absence of

Fig. 5. Sulphate salts associated with pyrite oxidation in Ka¨rkevagge, Sweden.

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long-term ground moisture and temperature records it can only be noted that, in terms of regional air climate, the second study period was both warmer and wetter than the first. A final element of the Ka¨rkevagge chemical weathering dataset is a series of water quality measurements (Campbell et al. 2001). However, it should be noted that water quality is always a secondary measure of chemical weathering and one that, by its homogenization across a catchment or portions of a catchment, does not capture the details seen in direct rock studies. The limited distribution of the schist units in which the pyrite is disseminated means that only a portion of the valley’s waters are dominated by sulphate while a combination of atmospherically derived CO2 and that derived from the weathering of local carbonate units dominates some portions of the valley. In broad terms there is a trade-off between solute intensity and discharge within the valley. The ultimate story to emanate from valley waters is that the minor mineral constituent pyrite and water availability, that is, seasonal snowcover, both have the capacity to influence denudation patterns in such an Arctic-alpine periglacial environment.

Storbreen The glacier foreland of Storbreen, Norway (618350 N, 88210 E), provides insight into the much earlier stages of post-glacial weathering, around 250 years, of resistant rocks dominated by pyroxene gneisses (Thorn et al. 2007).

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Soil development is slow, but detectable, in this environment. In particular, the complex embryonic stages of the chemical weathering of surficial and near-surficial cobbles on dated moraines (Matthews 1974) is readily detectable under a scanning electron microscope. Using the percentage porosity of thin slices of plagioclase grains at and near cobble surfaces as a surrogate measure of weathering (Thorn et al. 2007), the early stages of chemical weathering were determined from three microenvironments and several ages. The microenvironments were lichenfree surficial cobbles on moraine surfaces compared to lichen-covered cobbles in the same position and, thirdly, shallowly buried (c. 40 cm) cobbles from the same moraine locations. The age factor was created by examining multiple cobbles from the valleyside slope beyond the glacier foreland, a surface that is thought to be c. 10 ka old, and those from the 1750, 1810, 1870 and 1928 moraines. Three important trends emerge: (a) porosity (chemical weathering), admittedly at a microscopic scale, is readily apparent on glacial deposits within 200 a even though the environment is alpine periglacial and the rock type quite resistant; (b) buried cobbles are always significantly more weathered than surficial cobbles of both categories; and (c) it takes c. 190 a before there is a detectable increase in the weathering rate of lichen-covered cobbles over lichen-free cobbles (Fig. 6). Of course, given the slowness of lichen establishment and the lifespan of lichens, it is not possible to assume that the lichen-covered versus lichen-free distinction has been stable throughout the identified time periods.

Fig. 6. Porosity of cobbles recovered from Storbreen glacial foreland, Jotunheimen, Norway, from beneath the soil and on the surface, with and without lichen cover versus the age of the moraine. Moraine ages were determined by Matthews (1974) using lichenometry. Reproduced from Thorn et al. (2007) with the permission of Wiley-Blackwell, Oxford, England.

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Discussion In addition to Jenny’s (1941) soil model, soils may be examined in terms of a process model that invokes additions, translocations, transformations and losses (Simonson 1959). In terms of additions, and because many periglacial environments are young, the soil is initially simply rock debris. Subsequently, surfaces are quickly colonized initially by simple non-vascular plants making up a cryptogram crust followed, if conditions are favourable, by vascular plants. The latter change the parent material by increasing weathering due to the organic acids they produce and by adding organic matter. At Storbreen, the accumulation of organic matter in soils over a c. 250 a time span was measured; there was an exponential increase in both the total soil organic matter content and rate of soil organic matter content accumulation (Darmody et al. 2005b). This is because the pioneering plants improve the soil fertility, and subsequently allow better plant productivity. N content, a measure of soil fertility, is initially very low on deglaciated terrain, but pioneering plants are typically N fixers and they increase the N content of soils. However, the rate of N increase decreases exponentially as later non-N fixing successional plants take advantage of the increased soil fertility (Darmody et al. 2005b). In Ka¨rkevagge, soil materials that are essentially the same age may have very different horizonation because of the additions of organic matter and organic acids brought on by the plants growing locally. Microclimatic variability largely controls plant growth in alpine periglacial environments. For example, shelter from winds or a reliable moisture supply can allow plants to grow vigorously relative to exposed or dry sites. In extreme instances, similar-aged materials on comparable slope angles may exhibit strong or no horizonation dependent upon such variables. Variation in additions of plant-derived organic compounds or in translocation of associated chelated materials can produce prominent horizonation and well-developed soil profiles (Fig. 3). Harsh climatic conditions at this high latitude, as would also be the case at high elevations, restrict the kind of vegetation that can grow on any given spot and thus the soils found there. In reference to specific vegetation assemblages found in Ka¨rkevagge, podzolization in more sheltered locations is produced by the chelating action of plant exudates from white birch (Betula pubescens). By contrast, on exposed locations, soils have very poor horizonation and only a cryptogram crust with perhaps a few Least Willows (Salix herbacea) and Dwarf Birch (Betula nana). This vegetative cover is insufficient to produce the necessary chelating agents to develop

strongly expressed soil horizonation in the time since the surface stabilized. The previously mentioned oxidation of pyrite in Ka¨rkevagge as an important driver of both weathering and pedogenesis is one example of transformation. The aggressive sulphuric acid accelerates weathering and acidifies the soil. Likewise, the production of interstratified secondary minerals from primary mafic minerals in Storbreen is another example of a transformation (Darmody et al. 2005b). As weathering and pedogenesis proceeds, materials are lost to ground water. In Ka¨rkevagge, water chemistry is quite different depending on which part of the valley it originates due to the irregular distribution of pyrite (Darmody et al. 2000a; Campbell et al. 2001, 2002). The presence of permafrost is an important attribute in distinguishing soils in periglacial regions for the purpose of classification (Soil Survey Staff 1999). Soils that have permafrost within 1 m of the soil surface, or permafrost within 2 m and gelic materials within 100 cm, are considered within a unique soil order: the Gelisols. Gelic materials may be either mineral or organic but must show evidence of cryoturbation. This includes ice segregation and/or thermal-contraction cracking. Field evidence for cryoturbation includes irregular and broken horizons, involutions, organic matter accumulation on top of and within the permafrost, oriented coarse fragments and silt-enriched layers. Ice segregation is manifested by ice crystals, ice lenses, vein ice and ice wedges. Soil structure may be granular, platy or vesicular; subsurface horizons can be blocky, prismatic or structureless (massive) (Bockheim et al. 1997). Despite the cold climate, many different soil orders can be found in alpine periglacial regions. For example, in the limited area of Ka¨rkevagge, Mollisols, Gelisols, Inceptisols, Histosols, Entisols and Spodosols have all been identified (Darmody et al. 2000b). Conceptually, soil orders found in periglacial regions could also include Andisols (and maybe Vertisols under special conditions). This leaves out only Oxisols, Ultisols and Alfisols from those thought theoretically possible to occur in periglacial regions. Weathering and pedogenesis in alpine periglacial environments are closely related. Cold temperatures combined with a youthful landscape and materials would seem to be the dominant themes, but water dominates the salient processes. Cold temperatures cause water to freeze promoting cryopedoturbation and potentially frost shattering. However, frost shattering may in many cases be a product of self-destructive mineralogy as oxidation and hydrolysis exploit weatherable mineralogy (Fig. 2). It is also probably a secondary mechanism in many instances, occurring where ice crystals and/ or veins are able to grow in (and therefore exploit)

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fractures produced by innumerable other processes. Moreover, young materials are generally more weatherable than older materials so, despite the low temperatures, chemical weathering rates can be high in cold climates (Anderson et al. 1997). Undoubtedly, cold temperatures can slow many chemical processes but this does not change their fundamental nature, only their rate. Furthermore, in an environment dominated by coarse material (as opposed to fines), simple chemical principles other than temperature will also slow chemical weathering (i.e. surface area limitations and rapid drainage). In the realm of pedogenesis, cold temperature limits plant growth. This lessens its impact on pedogenesis because of the lower production rates of N and C fixation and lower rates of organic acid and chelating agent production. All these foster weathering and horizonation and eventually lead to a recognizable soil. Conversely, the widespread occurrence of cold temperature, poor drainage and soil saturation limit organic matter decomposition; alpine periglacial soils can therefore be high in organic matter. In Scandinavia at least, larger soil-dwelling animals are not as prevalent as in temperate soils, so faunal pedoturbation is less common in periglacial regions. In summary, the following fundamental messages are fairly simple as regards weathering. (1)

(2)

(3)

(4)

There is no field evidence to establish that freeze –thaw weathering dominates alpine periglacial environments. An equally plausible explanation would be to claim that the embryonic breakdown of bedrock and coarse debris in such areas is dominated by rock characteristics. It is undoubtedly true that freeze–thaw weathering occurs, but the extent to which it is limited by initial rock properties, is a secondary exploiter of fractures produced by other mechanisms or its relative importance visa`-vis chemical weathering have not been established. Alpine periglacial surfaces are commonly ‘young’ in the fundamental sense, recently deglaciated being the most common example. However, if both the rates of mechanical and chemical weathering are slow, they may also be young on that account. Alpine periglacial surfaces are enormously fragmented because their effective delivery of moisture is disproportionately skewed by seasonal snowpack accumulation and its subsequent melt. This situation may be further exacerbated by the high degree of variability in the coarseness of periglacial surfaces which influences their drainage to a very high degree.

(5)

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Moisture variability is likely to be at least as important in controlling alpine periglacial weathering as freezing and thawing.

All of these statements are equally applicable to pedogenesis in alpine periglacial environments. However, pedogenesis is rendered more complicated by the variability in botanical inputs.

Conclusions It is instructive to view alpine periglacial weathering and pedogenesis from the perspective of their commonalities with other climatic regimes, rather than to view them exclusively through the lens of their temperature regimes. Weathering and pedogenesis vary little in kind compared to their variation in intensity at the Earth’s surface. The intensity is undoubtedly controlled by temperature, water supply and botanical inputs at a first approximation. While a focus upon the temperature controls may vindicate treating cold regimes separately, a focus upon water supply (and perhaps biological inputs) supports seeing them as part of a continuum. Single cause– effect relationships in geomorphology, in this case temperature and weathering, have generally proven unsatisfactory. Alpine periglacial environments are not necessarily dominated by mechanical or freeze –thaw weathering as is often assumed; in fact, this fundamental process is so poorly understood that it should be invoked with more care than is usually the case. Many alpine periglacial environments are young and have various parent materials which influence the nature of weathering. The primary influence of an alpine periglacial climatic regime is on the rates of chemical weathering and pedogenesis, rather than on their fundamental nature. In short, to assign all alpine periglacial environments to dominance by traditional cold region mechanical or freeze –thaw weathering is an oversimplification. Certainly, the data available today argue for a much more nuanced approach to alpine periglacial weathering than is generally invoked.

References Allen, C. E., Darmody, R. G., Thorn, C. E., Dixon, J. C. & Schlyter, P. 2001. Clay mineralogy, chemical weathering, and landscape evolution in Arctic-Alpine Sweden. Geoderma, 99, 277–294. Anderson, S. P. 2007. Biogeochemistry of glacial landscape systems. Annual Review of Earth and Planetary Sciences, 35, 375–399. Anderson, S. P., Drever, J. I. & Humphrey, N. F. 1997. Chemical weathering in glacial environments. Geology, 25, 399–402. Anderson, S. P., Drever, J. I., Frost, C. D. & Holden, P. 2000. Chemical weathering in the foreland of a

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retreating glacier. Geochimica et Cosmochimica Acta, 64, 1173– 1189. Andre´, M. F. 1995. Postglacial microweathering of granite roches moutonne´es in northern Scandinavia (Riksgra¨nsen area, 688N). In: Slaymaker, O. (ed.) Steepland Geomorphology. John Wiley & Sons, Chichester, 103–128. Andre´, M. F. 2002. Rates of postglacial rock weathering on glacially scoured outcrops (Abisko-Rikgra˚nsen area, 688N). Geografiska Annaler, 64A, 139 –150. Berrisford, M. S. 1991. Evidence for enhanced mechanical weathering associated with seasonally late-lying and perennial snow patches, Jotunheimen, Norway. Permafrost and Periglacial Processes, 2, 331– 340. Beylich, A. A. 2000. Geomorphology, sediment budget, and relief development in Austdalur, east Iceland. Arctic, Antarctic and Alpine Research, 32, 466– 477. Beylich, A. A., Kolstrup, E., Thyrsted, T., Linde, N., Pedersen, L. B. & Dynesius, L. 2004a. Chemical denudation in arctic-alpine Latnjavagge (Swedish Lapland) in relation to regolith as assessed by radio magnetotelluric-geophysical profiles. Geomorphology, 57, 303–319. Beylich, A. A., Kolstrup, E., Thyrsted, T. & Gintz, D. 2004b. Water chemistry and its diversity in relation to local factors in the Latnjavagge drainage basin, arctic-oceanic Swedish Lapland. Geomorphology, 58, 125– 143. Beylich, A. A., Sandberg, O., Molau, U. & Wache, S. 2006. Intensity and spatio-temporal variability of fluvial sediment transfers in an Arctic-oceanic periglacial environment in northernmost Swedish Lapland (Latnjavagge catchment). Geomorphology, 80, 114– 130. Bockheim, J. G., Tarnocai, C., Kimble, J. M. & Smith, C. A. S. 1997. The concept of gelic materials in the new Gelisol order for permafrost-affected soils. Soil Science, 162, 927–939. Brunsden, D. 1979. Weathering. In: Embleton, C. & Thornes, J. (eds) Process in Geomorphology. John Wiley & Sons, New York, 73–129. Campbell, S. W., Dixon, J. C., Darmody, R. G. & Thorn, C. E. 2001. Spatial variation of early season surface water chemistry in Ka¨rkevagge, Swedish Lapland. Geografiska Annaler, 83A, 169– 178. Campbell, S. W., Dixon, J. C, Thorn, C. E. & Darmody, R. G. 2002. Chemical denudation rates in Ka¨rkevagge, Swedish Lapland. Geografiska Annaler, 84A, 179– 185. Carroll, D. 1970. Rock Weathering. Plenum, New York. Darmody, R. G. & Thorn, C. E. 1997. Elevation, age, soil development, and chemical weathering at Storbreen, Jotunheimen, Norway. Geografiska Annaler, 79A, 215– 222. Darmody, R. G., Thorn, C. E. & Rissing, J. M. 1987. Chemical weathering of fine debris from a series of Holocene moraines: Storbreen, Jotunheimen, southern Norway. Geografiska Annaler, 69A, 405–413. Darmody, R. G., Thorn, C. E., Harder, R. L., Dixon, J. C. & Schlyter, P. 2000a. Weathering implications of water chemistry in an arctic-alpine environment, northern Sweden. Geomorphology, 34, 89–100. Darmody, R. G., Thorn, C. E., Dixon, J. C. & Schlyter, P. 2000b. Soils and landscapes of Ka¨rkevagge,

Swedish Lapland. Soil Science Society of America Journal, 64, 1455– 1466. Darmody, R. G., Allen, C. E., Thorn, C. E. & Dixon, J. C. 2001. The poisonous rocks of Ka¨rkevagge. Geomorphology, 41, 53–62. Darmody, R. G., Thorn, C. E. & Allen, C. E. 2005a. Chemical weathering and boulder mantles, Ka¨rkevagge, Swedish Lapland. Geomorphology, 67, 15–170. Darmody, R. G., Allen, C. E. & Thorn, C. E. 2005b. Soil topo-chronosequences at Storbreen, Jotunheimen, Norway. Soil Science Society America Journal, 69, 1275– 1287. Darmody, R. G., Thorn, C. E. & Dixon, J. 2007. Pyrite enhanced chemical weathering in Ka¨rkevagge, Swedish Lapland. Geological Society of America Bulletin, 119, 1477–1485. Darmody, R. G., Thorn, C. E., Schlyter, P. & Dixon, J. C. 2004. Relationship of vegetation distribution to soil properties in Ka¨rkevagge, Swedish Lapland. Arctic, Antarctic, and Alpine Research, 36, 21–32. Dixon, J. C. & Thorn, C. E. 2005. Chemical weathering and landscape development in mid-latitude alpine environments. Geomorphology, 67, 127–145. Dixon, J. C., Thorn, C. E., Darmody, R. G. & Schlyter, P. 2001. Weathering rates of fine pebbles at the soil surface in Ka¨rkevagge, Swedish Lapland. Catena, 45, 273–286. Dixon, J. C., Thorn, C. E. & Darmody, R. G. 2008. Spatial scale and chemical weathering in Ka¨rkevagge: influences on landscape evolution. Zeitschrift fu¨r Geomorphologie, 52, 27– 49. Egli, M., Mirabella, A. & Fitze, P. 2003a. Formation rates of smectites derived from two Holocene chronosequences in the Swiss Alps. Geoderma, 117, 81– 98. Egli, M., Mirabella, A., Sartori, G. & Fitze, P. 2003b. Weathering rates as a function of climate: results from a climosequence of the Val Genova (Trentino, Italian Alps). Geoderma, 111, 99– 121. Egli, M., Mirabella, A., Sartori, G., Giaccai, D., Zanelli, R. & Plo¨tze, M. 2007. Effect of slope aspect on transformation of clay minerals in Alpine soils. Clay Minerals, 42, 373– 398. Ellis, S. 1980. An investigation of weathering in some arctic-alpine soils on the northeast flank of Oksskolten, north Norway. Journal of Soil Science, 31, 371–385. Etienne, S. 2002. The role of biological weathering in periglacial areas: a study of weathering rinds in south Iceland. Geomorphology, 47, 75–86. French, H. M. 2007. The Periglacial Environment. 3rd edn. John Wiley & Sons, Chichester, UK. French, H. M. 2008. Periglacial processes and forms. In: Burt, T. P., Chorley, R. J., Brunsden, D., Cox, N. J. & Goudie, A. S. (eds) The History of the Study of Landforms or the Development of Geomorphology. Volume 4: Quaternary and Recent Processes and Forms (1890– 1965) and the Mid-century Revolutions. The Geological Society, London, 621– 676. Geiger, R. 1950. The Climate Near the Ground. Harvard University Press, Cambridge. Hall, K. 2006. Perceptions of rock weathering in cold regions: a discussion on space and time attributes of scale. Ge´omorphologie, Relief, Processus, Environment, 3, 187 –196.

WEATHERING IN PERIGLACIAL SCANDINAVIA Hallet, B. 2006. Why do freezing rocks break? Science, 314, 1092–1093. Jenny, H. 1941. Factors of Soil Formation. McGraw-Hill, New York. Loughnan, F. C. 1969. Chemical Weathering of the Silicate Minerals. Elsevier, New York. ¨ ber die mechanische Verwitterung Łozin´ski, W. 1909. U der Sandsteine im gema¨ssigten Klima. Bulletin International de l’Acade´mie des Sciences de Cracovie class des Sciences Mathe´matique et Naturelles, 1, 1–25. Łozin´ski, W. 1912. Die periglaziale fazies der mechanischen Verwitterung. In: Evans, D. J. A. (ed.) Cold Climate Landforms. 11th Internat. Geol. Congress, Stockholm 1910, Comptes Rendus, 1039–53. Also translated by T. Mrozek (1994). Wiley, Chichester, UK, 119 –134. McCarroll, D. 1990. Differential weathering of feldspar and pyroxene in an arctic-alpine environment. Earth Surface Processes and Landforms, 15, 641–651. McCarrol, D. & Viles, H. A. 1995. Rock weathering by the lichen Lecidea auriculata in an arctic-alpine environment. Earth Surface Processes and Landforms, 20, 199–206. McGreevy, J. P. & Whalley, W. B. 1982. The geomorphic significance of rock temperature variations in cold climate environments: a discussion. Arctic and Alpine Research, 14, 157 –162. Matsuoka, N. 2001. Microge´livation v. macroge´livation: towards bridging the gap between laboratory and field frost weathering. Permafrost and Periglacial Processes, 12, 299– 313. Matthews, J. A. 1974. Families of lichenometric dating curves from the Storbreen gletschervorfeld, Jotunheimen, Norway. Norsk Geografisk Tidsskrift, 28, 215–235. Matthews, J. A. 1992. The Ecology of Recently Deglaciated Terrain. Cambridge University Press, Cambridge. Matthews, J. A. & Owen, G. 2008. Endolithic lichens, rapid biological weathering and Schmidt hammer R-values on recently exposed rock surfaces: Storbreen glacier foreland, Jotunheimen, Norway. Geografiska Annaler, 90A, 287–297. Matthews, J. A., Dawson, A. G. & Shakesby, R. A. 1986. Lake shoreline development, frost weathering and rock platform erosion in an alpine periglacial environment. Boreas, 15, 33– 50. Mellor, A. 1987. A pedogenic investigation of some soil chronosequences on Neoglacial moraine ridges, southern Norway: examination of soil chemical data using principal components analysis. Catena, 14, 369–381. Messer, A. C. 1988. Regional variations in rates of pedogenesis and the influence of climatic factors on moraine chronosequences, southern Norway. Arctic and Alpine Research, 20, 31– 39. Mirabella, A., Egli, M., Carnicelli, S. & Sartori, G. 2002. Influence of parent material on clay minerals formation in Podzols of Trentino, Italy. Clay Minerals, 37, 699–707.

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Role of lichens in granite weathering in cold and arid environments of continental Antarctica MAURO GUGLIELMIN1*, SERGIO E. FAVERO-LONGO2, NICOLETTA CANNONE3,4, ROSANNA PIERVITTORI2 & ANDREA STRINI5 1

University of Insubria, Department of Structural and Functional Biology, Via Dunant 3, Varese 21100, Italy

2

3

University of Torino, Department of Plant Biology and Centre of Excellence for Plant and Microbial Biosensing, Viale Mattioli 25, Torino 10125, Italy

University of Ferrara, Department of Biology and Evolution, Corso Ercole I d’Este 32, Ferrara 44100, Italy

4

Present address: Department of Chemical and Environmental Sciences, Insubria University, Via Lucini, 3, 22100, Como, Italy 5

PNRA, c/o University of Insubria, Department of Structural and Functional Biology, Via Dunant 3, Varese 21100, Italy *Corresponding author (e-mail: [email protected]) Abstract: The mechanical and chemical effects of lichens on the outer and inner surfaces of tafoni features were investigated through a multidisciplinary approach at two locations (Oasi 748420 S, 1648070 E, 40–250 m a.s.l.; Mount Keinath, 748320 S; 1638580 E; 850 m a.s.l.) close to the Italian Antarctic station (Mario Zucchelli). Outer tafoni roof surfaces show low values of effective porosity coupled with pervasive hyphal penetration and an extensive reddish-brown weathering rind. Inner tafoni surfaces show higher values of effective porosity, which correspond with an almost absent weathering rind and low hyphal penetration. Our observations indicate that: (a) iron oxyhydroxides, particularly concentrated where hyphal patches and bundles contact biotite, consist of hematite; (b) the microcosms of lichen hyphae and their precipitates fill voids to form case hardening on outer surfaces; and (c) on inner surfaces biological action is less active, most likely because of more intense thermal stress and salt action.

In Antarctica, as in other cold environments, mechanical weathering processes have traditionally been considered to be dominant (Campbell & Claridge 1987; Matsuoka 1995). However, some studies have recognized the importance of either chemical or biological processes (Friedman & Weed 1987; Hall & Otte 1990; Etienne 2002; Hall et al. 2002; Guglielmin et al. 2005). Among rockdwelling organisms, lichens have been identified as efficient deterioration agents. They are associated with both biomechanical and biochemical processes (Chen et al. 2000; St Clair & Seaward 2004; Gadd 2007). Biomechanical deterioration produced by lichens includes rock disaggregation upon hyphal penetration through cleavage planes and fractures; following the expansion and contraction of the thallus associated with its hydration state and with freezing and thawing; and the incorporation of mineral fragments into the thallus during lichen growth (Chen et al. 2000). Biochemical effects of lichens depend on respiratory CO2 and on the

excretion of oxalic acid and of other metabolites characterized by acidic and chelating functions, which are able to induce mineral solubilization and mineral neoformation (Chen et al. 2000; Gadd 2007). Lichens may exert another effect on rocks known as ‘rock bioprotection’ (Carter & Viles 2005; Favero-Longo et al. 2009). This is much less well known. The equilibrium between biodeterioration and bioprotection mainly depends on the weathering ability of the species and on the physicochemical features of the lithotypes (Adamo & Violante 2000; Chen et al. 2000; Gazzano et al. 2009a). The occurrence of biological crusts or films has been observed in cryotic environments such as Antarctica (Viles 1995; Andre´ & Hall 2005; Guglielmin et al. 2005), but its associated biogeomorphological action (sensu Naylor et al. 2002) has been neglected. In continental Antarctica where extreme aridity and cold limit soil formation and debris supply, physical and chemical weathering of granite by lichens, mainly endoliths, has been described

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 195–203. DOI: 10.1144/SP354.12 0305-8719/11/$15.00 # The Geological Society of London 2011.

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(Wierzchos et al. 2003; De los Rı´os et al. 2005). However, the lichen effects on cavernous weathering (tafoni), which is one source of sediment supply and one of the faster morphogenetic processes active (French & Guglielmin 1999; Guglielmin et al. 2005) has not been considered to date. This paper assesses the role of lichens on granite weathering and the development of tafoni using both geomorphological and biological techniques.

Study area The study area is located in Northern Victoria Land around the Italian Mario Zucchelli Antarctic Station (MZS; 748410 S, 164870 E) (Fig. 1) where large ice-free areas occur. Most areas are underlain by Ordovician and Precambrian granites (Granite Harbour Formation: Carmignani et al. 1988). The climate is cold and arid. Mean annual air temperature is around 214 8C in the coastal area, ranging between –38 and 8 8C (Guglielmin et al. 2009), while it is colder inland (Mean Annual Air Temperature (MAAT) is 222.6 8C at Mount Keinath) (Strini et al. 2008). Precipitation is less than 200 mm per year (Caprioli et al. 1997). Permafrost is continuous and the active layer in unconsolidated sediments ranges between 20 and 30 cm (Guglielmin 2006). Vegetation is composed exclusively of cryptogams and is dominated by lichens and mosses (Kappen 1985; Schwarz et al. 1992; Castello & Nimis 1995; Seppelt et al. 1995, 1996; Seppelt & Green 1998; Lewis Smith 1999; Cannone 2004; Cannone & Seppelt 2008). Lichen epilithic communities are dominated by Lecanora physciella (Darb.)

Hertel., Buellia frigida Darb., Lecidea cancriformis C.W. Dodge & G.E. Baker. Pleopsidium chlorophanum (Wahlenb.) A. Massal. and Umbilicaria decussata (Vill.) Zahlbr. are also reported. Species zonation depends on the morphology of the weathering landforms (Guglielmin et al. 2005). Observations were made at two sites: Mount Keinath and Oasi. Mount Keinath (748320 S, 1638580 E, c. 850 m a.s.l.) is a granitic nunatak protruding through the Antarctic ice cover that is characterized by welldeveloped tafoni (singular: tafone) and other weathering features such as grooves and pits (French & Guglielmin 2002a, b). The rock surface is coated by a reddish weathering rind (which can exceed 2 cm in thickness) and sometimes by a red rock varnish. The rock is a medium-fine grained monzogranite, composed of zoned plagioclase, quartz, microcline, biotite (apatite, titanite, zircon, magnetite as accessories) with maximum crystal diameters of 2 cm (Lanza & Tonarini 1998) and rare joints almost parallel to the slope (Guglielmin et al. 2005). Vegetation is exclusively epilithic and composed of lichens, dominated by Lecanora physciella in a strongly anitrophytic environment. The epilithic communities include other crustose lichens (such as Lecidea cancriformis, Pleopsidium chlorophanum), but only one species of foliose lichen (Umbilicaria decussata). Oasi (748420 S, 1648070 E, 40–250 m a.s.l.) is located in the immediate vicinity of the MZS and within a few hundred metres of the sea. The site is a gentle slope rounded and scoured by Cenozoic glaciations (Orombelli et al. 1990). The glacial sediments are referred to as the ‘Terra Nova Drift’. The surface shows pale yellow or orange discontinuous weathering rinds (Guglielmin et al. 2005). Several tafoni are developed both on granite bedrock faces and on erratic boulders. The bedrock is coarsegrained granite (with K-feldspar crystals exceeding 10 cm in length) with a large number of randomlyoriented joints. The vegetation is sparse and mainly consists of epilithic lichens (such as Buellia frigida, Lecidea cancriformis, Acarospora gwynnii C. W. Dodge & E. D. Rudolph, Umbilicaria decussata, Pseudephebe minuscula (Nyl. ex Arnold) Brodo & D. Hawksw., Xanthoria elegans (Link) Th. Fr.) with a relatively higher abundance of Buellia frigida, indicating a more nitrophytic environment (Guglielmin et al. 2005).

Field and laboratory procedures Field investigations

Fig. 1. Localization of the study area.

Field investigations at each site included a survey of the characteristics of the granitic outcrops (joint width, frequency and orientation and crystal size)

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and a visual description of the weathering rind and any other type of weathering phenomena that was present (Guglielmin et al. 2005). The spatial distribution of standing or running water, snow and congelation ice within and outwith all tafoni was also recorded. Samples of rock and sediment associated with tafoni features were collected for thin section, scanning electron microscopy (SEM) and microchemical analyses. A total of 35 rock samples were collected for subsequent analyses in Italy (Milan, Varese and Turin labs). A vegetation survey focused on describing the colonization patterns for each weathering feature. In particular, the floristic composition and spatial distribution of lichen species and their relation to the occurrence of weathering products (rock varnish, rind, grus) and water were examined. The nomenclature of lichen species used follows Øvstedal & Smith (2001) and Castello & Nimis (2000).

These latter were also analysed with microRaman spectroscopy (LabRam HR micro-Raman – Jobin Yvon – equipped with Olympus BX41 microscope) to assess their mineralogical identity. The analyses were performed in situ on the same thin sections, and the acquired spectra were compared with those contained in the Raman Spectra Database of Minerals and Inorganics (http://riodb.ibase.aist. go.jp/rasmin/). Mineral fragments scraped at the inner and outer surfaces of Mount Keinath tafone and grus samples (w , 125 mm fraction) collected on the floors of tafoni in both sites were powdered and analysed using X-ray powder diffractometry (XRPD) technique with copper anticathode X-ray tube (Siemens D5000 with u –2u geometry). The spectra obtained were identified by comparison with those contained in the Joint Committee of Powder Diffraction Standard (JCPDS) archives.

Analysis of weathering rinds and varnish

Surface disaggregation

In order to observe weathering changes between the surface and the rock interior, thin sections were cut perpendicular to the topographic surface to depths of 2 cm. Thin sections were photographed at different magnifications under an optical microscope by reflected and transmitted light microscopy. Any extension of weathering rind of a selected Mount Keinath tafone was quantified by image analysis of polished cross-sections (sensu FaveroLongo et al. 2005), which were also used to characterize hyphal penetration. These sections were observed under reflected light using a stereomicroscope. High-definition images were acquired using a scanner and processed through colour-based pixel classification by WinCAM Pro 2007 (Regent’s Instrument). This software package was trained to identify the colours of reddish-yellow weathering rind, thus allowing the quantification of rind extension. Results were statistically analysed by Analysis of Variance (ANOVA) with post hoc Tukey test. SYSTAT release (version 5.2; SYSTAT, Evanston, IL, USA) was used to perform the test ( p , 0.05).

A pull-up test using adhesive tape (3M 396, 3Mw, Italy) was applied to evaluate the disaggregation of the Mount Keinath tafone surfaces according to a protocol currently used on manmade substrates (Favero-Longo et al. 2009). Tape was applied to surfaces of average spatial extent 50 cm2. The snatched material underwent gravimetric analysis and results were statistically analysed by ANOVA with post hoc Tukey test.

Analysis of mineral weathering and neoformation Scanning electron microscopy and energy dispersive spectroscopy (SEM-EDS) elemental analyses were performed on carbon-coated thin sections with a Stereoscan S360 Cambridge Electron Microscope equipped with an Energy 200 Oxford Instruments EDS apparatus. Analyses were performed on iron-rich silicates (i.e. biotite) contacted by lichen hyphae within the rock and on mineral deposits filling intergranular and intragranular discontinuities.

Porosity Total porosity in different positions of the Mount Keinath tafone roof was analysed by image analysis of SEM back-scattered electrons (BSE) images (Campbell 1999; Thorn et al. 2007). The previously described thin cross-sections were carbon coated and observed using a Stereoscan S360 Cambridge Electron Microscope. Pores, including cracks, intragranular and intergranular voids, appear black in the BSE-images while minerals show different tonalities of grey depending on chemical composition. The extension of porosity and minerals (black and grey areas respectively) was quantified through greyscale-based pixel classification using WinCAM software. Results were statistically analysed by ANOVA with post hoc Tukey test. Effective porosity was quantified through an immersion gravimetric analysis of cross-sections. The weight increase was evaluated after immersion within de-ionized water.

Analysis of the hyphal penetration component The polished cross-sections of Mount Keinath tafone (the same used to measure the extension of

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the weathering rind) were stained using the Periodic Acid Schiff’s (PAS) method (Whitlach & Johnson 1974) in order to highlight the hyphal penetration component (HPC, sensu Favero-Longo et al. 2005). The spread of hyphae penetrating within the substrate (%) was quantified according to Gazzano et al. (2009b): 400 dpi images of stained sections were acquired and analysed through colour-based pixel classification using WinCAM. Results were tested for significance by means of ANOVA with post hoc Tukey test.

Results General At both sites (Mount Keinath and Oasi) tafoni hollows do not exhibit any statistically significant preferred orientation. In general, tafoni development post-dates other weathering features such as pits or grooves, usually destroying or truncating them (French & Guglielmin 2002a, b). Tafoni floors are covered by angular rock fragments (‘grus’), sometimes lying on snow. A slightly reddish weathering rind (7.5YR 6/8) is sometimes present on the inner tafoni walls. Epilithic lichens generally colonize the external roof of tafoni, with similar floristic composition to that of the surrounding rocks. Epilithic vegetation is absent on all internal tafoni surfaces. Instead, the latter are colonized by endolithic algae and hyphae. Scales and flaking generally break away between the hyphal layer and the underlying algal layer (Fig. 2). SEM samples of grus collected within tafoni did not show etching or pitting and grain surfaces and borders appeared fresh. X-ray diffractometer analysis of grus samples did not reveal any

Fig. 2. Breakdown of a rock sliver along the algal layer from the outer surface of the tafone roof at Oasi. It is possible to observe the whitish hyphal network and the underlying green algal layer.

evidence of mineral formation produced by subaerial weathering, such as clay minerals or salts, with the exception of some gypsum in a few samples of the ,125 mm fraction (Guglielmin et al. 2005).

The case study of Mount Keinath The outer surface of tafoni at this location are distinctive because they are coated by a yellow-red weathering rind and, locally, by a reddish rock varnish. Epilithic colonization is absent on the inner surface of the roof while just a few thalli of Lecidea cancriformis, up to 2 cm wide, colonize the outer roof surface. These grow along the microfissures between crystals and display white hyphal patches where crystals break off from the surface (Fig. 3a). Sparse epilithic colonies of microcolonial fungi (MCF), up to 250 mm in diameter, also occur on the outer tafone surface. A continuous occurrence of hyphae characterizes the 1.2 cm (+0.2) deep superficial layer of the outer roof (Fig. 3b). Patches of mycelium c. 150 mm thick occur between 0.5 and 1.5 mm in depth. They usually include scattered and/or clustered green algae. Similar patches also occur down to 2.5 mm in depth (rarely including algae) and to 3.5 mm (never displaying algal occurrence). Only rare, scattered hyphae occur within the superficial layer of the inner roof surface. The HPC is more pervasive on the outer roof (mean 15% of the PAS-stained polished crosssections) than in the inner roof (0.6%). The weathering rind is also more pervasive on the outer side (27%) than on the inner (3.2%) of the tafone roof. Figure 4 shows an example of image analysis of PAS-stained polished cross-sections through colour-based pixel classification to evaluate the extension of Weathering Rind (WR) (yellow) and HPC (violet) on the outer roof. Light microscopy observations indicate that the weathering rind goes down to 22.5 + 12.4 mm on the tafone outer roof surface. Rock varnish is localized only where the weathering rind reaches the deepest extension from the surface (31.0 + 10.8 mm). Light microscopy also indicates that reddish (oxidation) colours are mainly spread around the iron-rich minerals as biotite and magnetite, especially where they are in contact with hyphal patches and the penetrating hyphal network (Fig. 5a). However, EDS analysis of biotite in contact with the hyphae shows a slight depletion of K in proximity of the hyphae (from 10.3 to 8.1%; p , 0.003), while Fe shows a (relative) increase (from 21.4 to 25.1%; p , 0.043). Far from any iron-rich minerals, reddish colours are localized along intergranular (and sometimes intragranular) discontinuities where an iron-rich

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Fig. 3. Lichen colonization of the outer surface of the tafone roof. (a) Thalli of Lecidea cancriformis (detail in inset) growing along microfissures between crystals and showing white hyphal patches where crystals break off. Scale bars: 1 cm (2 mm in inset). (b) Hyphal penetration component, displaying hyphal patches which include algae in the superficial layer of rock and a hyphal network below, down to c. 1.2 mm. Scale bar: 200 mm.

Fig. 4. Image analysis of PAS-stained polished cross-sections through colour-based pixel classification (a) microphotograph aquired in stereomicroscopy; (b) recognition of WR, yellow, and HPC, violet, by WinCAM software (scale bar: 1 cm).

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neoformed minerals on the inner and outer Mount Keinath tafone roof. Micro-Raman spectroscopy (Fig. 5c), when applied to areas where an iron-rich precipitate is abundant, suggests hematite (Fe2O3) as the main mineral constituent. The inner surface of the tafone does not display any weathering rind, and low values of reddishyellow colours were detected by image analysis. No neoformation minerals were detected by XRPD and micro-Raman spectroscopy. Pull-up tests highlight a 20-fold higher disaggregation attitude for the inner surface of the tafone roof with respect to the outer (average fragment removal of 7.3 and 0.4 mg cm22, respectively). The inner surface of the tafone also shows a higher total porosity than at the outer (9.2 versus 7.0%). Moreover, all the pores were saturated on the outer roof (effective porosity 1.6%) while disaggregation prevented the possibility of measuring effective porosity on the inner side.

Discussion The hardening of the outer surface of the tafoni roof and a softening of the inner part have been observed in a number of different climatic and geomorphological conditions (Conca & Rossmann 1985; Mellor et al. 1997; Campbell 1999; Strini et al. 2008). Lichen colonization influences tafoni development and grus production in two ways.

Hardening of outer surface

Fig. 5. Weathering rind at the external surface of the tafone roof. (a) Reddish weathering rind localized where the endolithic lichen Lecidea cancriformis contacts biotite. Scale bar: 0.5 mm. (b) Reddish iron oxi-hydroxides filling porosity (intragranular void) within plagioclase. Scale bar: 12 mm. (c) Micro-Raman spectrum of hematite detected in correspondence of the weathering rind (x, Raman shift; y, intensity; AU, arbitrary units).

precipitate coats the mineral boundaries, often completely filling the void (Fig. 5b). The X-ray diffraction patterns did not highlight any occurrence of

Lichen colonization contributes to the hardening of the external surface of the roof of tafoni features. A correlation between the occurrence of cryptoendolithic communities and case-hardening was also suggested for sandstones of southern Jordan (Viles & Goudie 2004; Smith 2009). In Northern Victoria land, the very low values of our pull-up tests (0.4 mg cm22) and effective porosity measurements (1.6%) upon the outer tafoni surface indicate surface hardening despite the relatively high total porosity (7.0%). This hardening appears related to the denser and more pervasive HPC and weathering rind composed of iron oxides, mainly hematite (Fe2O3). These results suggest that hyphal structures increase by growing within cracks and along mineral boundaries and intragranular planes of weakness (as foliation), occupy rock voids (thus increasing the total porosity by expanding them) and, at the same time, decrease effective porosity. The biochemical action of Lecidea cancriformis in producing iron-rich precipitate as hematite (Fe2O3) is important because it fills the voids in a similar fashion to that observed by Wierzchos

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et al. (2003) for the same lichen, but within sandstone in Antarctica. Iron oxide is also known to act as cement within sandstone in a Mediterranean environment (McBride & Picard 2004). Analyses by De los Rı´os et al. (2005) on the weathering of biotite in granites of the Granite Harbour Formation by Lecidea cancriformis appears to be associated with a relative increase in iron. A relative increase of iron within biotite may be derived by the reprecipitation of the iron oxide deposit on the same surface of the weathered silicate. This is because of a templating action of adsorbed organic molecules (Beech et al. 2005) (as previously reported for other iron-containing bioweathered silicates; Turci et al. 2007) while K does not re-precipitate. Biotite in granite contacted by hyphae of Rhizocarpon geographicum display a decrease in both the K and Fe contents and leave a silica-rich relic which, when mobilized, fills fractures and pores beneath the interface (Lee & Parsons 1999). Likewise, Omelon et al. (2007) describe similar ion mobilization by cryptoendolithic communities both for Arctic and Antarctic cold environments. In summary, the filling of discontinuities with hyphae and their organic secretes-driven mineral precipitates (hematite in particular) are most likely to cause the case-hardening of the outer roof surface of the examined tafone.

Disaggregation of inner surface Our data show that the inner surface of the tafone roof has the highest value of total and effective porosity. The pull-up test indicates a high degree of disaggregation, despite a very poor occurrence of lichens at Mount Keinath (but not at Oasi where cryptoendolithic hyphae appear common). The absence of epilithic colonization on the inner surfaces suggests that abiological processes are more intense, as indicated by flaking and granular disaggregation. According to Strini et al (2008), short-time temperature fluctuations are larger on the inner surface than outside and therefore thermal stress (stress induced in crystals by contraction/ dilation cycles due to temperature fluctuations) is more intense than on the outer surface. Moreover, salt action (especially mechanical) is more pervasive on the inner surface because the lower wind speeds within the tafoni cavities induce a more abundant snow accumulation (Strini et al. 2008).

Conclusions Analyses carried out on tafoni developed in granites at two sites (Oasi and Mount Keinath) stress the role of lichens in the weathering process. A multidisciplinary approach indicates that weathering

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processes differ between the outer tafoni roof surface and the inner surfaces. The outer tafoni roof surfaces show lower values of effective porosity coupled with a more pervasive hyphal penetration and a more extensive and deeper reddish-brown weathering rind. The inner surfaces show higher values of effective porosity corresponding to an almost absent weathering rind and a very low hyphal penetration. Our results indicate that (a) iron oxy-hydroxides, particularly concentrated where hyphal patches and bundles contact biotite, consist of hematite and (b) the microcosms of lichen hyphae and their precipitates fill voids thereby forming a case hardening on the outer tafoni surface. On the inner surface, it seems that biological action is less active, most likely due to the more intense thermal stress and salt action. Research was funded by the Programma Nazionale di Ricerche in Antartide (PNRA) within the framework of the project entitled ‘Permafrost and Climate Change in Antarctica’. Special thanks to PNRA for logistical field support and to two anonymous referees for their revision of the manuscript.

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BIOWEATHERING IN ANTARCTICA Øvstedal, D. O. & Lewis Smith, R. I. 2001. Lichens of Antarctica and South Georgia. Cambridge University Press, Cambridge. St Clair, L. L. & Seaward, M. R. D. 2004. Biodeterioration of Stone Surfaces. Kluwer Academic Publishers, Dordrecht. Schwarz, A. M. J., Green, T. G. A. & Seppelt, R. D. 1992. Terrestrial vegetation at Canada Glacier, Southern Victoria Land, Antarctica. Polar Biology, 12, 397–404. Seppelt, R. D. & Green, T. G. A. 1998. A bryophyte flora for Southern Victoria Land, Antarctica. New Zealand Journal of Botany, 36, 617–635. Seppelt, R. D., Green, T. G. A. & Schroeter, B. 1995. Lichens and mosses from the Kar Plateau, Southern Victoria Land, Antarctica. New Zealand Journal of Botany, 33, 203–220. Seppelt, R. D., Green, T. G. A. & Schroeter, B. 1996. Additions and corrections to the lichen flora of the Kar Plateau, Southern Victoria Land, Antarctica. New Zealand Journal of Botany, 34, 329–331. Smith, B. J. 2009. Weathering processes and forms. In: Parsons, A. J. & Abrahams, A. D. (eds) Geomorphology of Desert Environments, 2nd edn. Springer Science and Business Media B.V., Berlin, 69– 100. Strini, A., Guglielmin, M. & Hall, K. 2008. Tafoni development in a cryotic environment: an example

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Periglacial sediments: do they exist? JEF VANDENBERGHE Institute of Earth Sciences, VU University, De Boelelaan 1085, 1081 HV Amsterdam, The Netherlands (e-mail: [email protected]) Abstract: ‘Periglacial deposits’ have often been reported in the Quaternary geological literature. Sometimes their identity has been confused with the periglacial processes, landforms and environments with which these deposits are linked. However, the question arises: what are the diagnostic characteristics of such ‘periglacial’ deposits? A number of examples illustrate that ‘periglacial deposition’ follows the general geomorphological and sedimentological laws, although it may be influenced by specific conditions of vegetation and sediment availability. These may be completely or partly dependent on periglacial climatic conditions. Some marginal conditions are specific to cold environments, for instance the scarcity of vegetation that has a clear effect on transport and deposition of aeolian sediments. In addition, the intensity of some processes is governed by climate-derived conditions, for instance in the case of fluvial and slope processes. In this sense, the answer to the question in the title is negative. ‘Periglacial deposits’ do not correspond to any particular sedimentary process. This conclusion is, however, only valid when post-depositional weathering and deformations at micro- and macro-scale are excluded.

The term ‘periglacial’ has often been used to characterize landscapes, land-forming processes and vegetation that form under ‘cold’ conditions. Likewise, one may find in the literature the term ‘periglacial sediments’. The central question in this paper is whether there is a single azonal sedimentology (the study of sediments) or a specific zonal sedimentology; for instance, ‘periglacial sedimentology’ was put forward by Clark (1988) for ‘periglacial hydrology’. It is striking that the term ‘periglacial sediments’ does not occur in sedimentological literature but emerges from time to time in descriptions on periglacial environmental processes (French 2001). Before analysing this question, we have to start at least from generally acceptable definitions. An acceptable definition of ‘sediment’ might be, without going into details and limiting ourselves to terrigenous sediments, ‘a body of solid material, mostly removed over some distance and subsequently deposited or re-deposited at a new place’. Like each material, a sediment may be characterized by its overall physical and chemical properties (such as cohesion, compaction) and organization (sedimentary structures) and by the properties of its constituents (composition and texture of particles). In contrast to ‘sediment’, the definition of the term ‘periglacial’ is disputed (French 2007) and sometimes its use even questioned (Linton 1969). We have to define what typifies the periglacial (or cold) nature of such sediments and ask the question why they should be different from ‘tropical (or warm) sediments’, a term which is in

fact never used. Tricart & Cailleux (1967) and Derruau (1972) emphasized the processes of freezing and thawing. In order to avoid some vagaries in terms such as ‘freezing’, ‘thawing’ and their ‘oscillations’, French (2007) proposes the general definition ‘cold, non-glacial’ when referring to periglacial processes and landforms. When referring to periglacial climate, he specifies the upper mean annual temperature limit at þ3 8C. In the present discussion on periglacial sediments a simple definition is used that accounts for ‘distinctive freezing and/or thawing influences’, without any regional significance or absolute temperature values. This definition is process oriented, as are the definitions by Tricart & Cailleux (1967) and Derruau (1972). Obviously, such processes occur under ‘cold, nonglacial’ conditions. Adopting a certain temperature limit is difficult to handle when talking about sediments and sedimentary processes, however. Periglacial sediments should be those that possess distinct characteristics related to either frost (either seasonal or perennial) or its subsequent thaw. It is the primary objective of this paper to question whether (terrigenous) periglacial sediments have an identity or not on the basis of their physico-chemical properties and structures. The answer to this question does not necessarily need to be the same as for periglacial geomorphology and weathering processes and environments. First, we must investigate whether those processes, forms and environments show a proper periglacial identity. Second, this analysis must be followed by the central question regarding ‘periglacial sediments’.

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 205–212. DOI: 10.1144/SP354.13 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Periglacial conditions and processes Periglacial climate and environment Climate conditions (mainly specific temperatures) control frost and thaw. By the accepted definition of ‘periglacial’, they therefore also have a definite significance in the concept of ‘periglacial’. It means that a ‘periglacial climate’ exists, that is, a climate that produces conditions that induce or favour the development of features that are dominated by frost and/or thaw. French (2007) divides the periglacial climate domain into one where frost action dominates with mean annual air temperature less than 22 8C, and one where frost action is present (but not dominating) with mean annual air temperature between þ3 and 22 8C. Apart from the absolute temperature, the rate of temperature change and annual temperature ranges also characterize the periglacial climate. However, the essence of the periglacial climate should not be limited to heat conditions since a whole range of other factors (such as humidity and wind activity) have to be considered when distinguishing different kinds of periglacial climate. Furthermore, periglacial climate conditions determine the complete environment (such as vegetation development, soil formation, hydrology and runoff). We can therefore distinguish ‘periglacial environments’ and ‘periglacial climates’ and typify them. In this sense it is understandable that several books have appeared under the title ‘periglacial environments’ (French 2007).

Periglacial landforms Frost and thaw play a major role in the nature and intensity of several geomorphological processes. Frost is directly involved in specific landforms that otherwise would never form because ice formation is essential to their development (such as pingos and palsas). Thaw degradation is responsible for specific forms such as hollows that develop after pingos or palsas degrade. Other geomorphological processes are only favoured by frost or thaw conditions. This means that such forms and processes may occur outside the periglacial realm, but are intensified by frost or thaw. Many slope processes such as block creep, mud flows or slumping are typical for periglacial environments but also occur in other environments. An illustrative example is provided by gelifluction and solifluction. Both process names have suffered from many definitions (Embleton & King 1975) and are used here to signify the slow, plastic flow of sediments excluding frost creep. Gelifluction received a proper name as a special kind of solifluction that involves the enhancement of the process by

the thaw of a frozen soil. Similarly, detachment slides should be considered as a special kind of epidermic debris sliding, where an active layer is making use of the sliding plane at the top of an ice body or frozen subsoil (Lewkowicz 1992). Those processes (gelifluction, detachment slides) may be called periglacial as they could not take place without the presence of frozen and thawed ground. Their related forms, however, cannot be distinguished from similar forms that developed without any frost or thaw. Solifluction, debris sliding and mud flows are general azonal processes. The periglacial singularity of fluvial hydrology and processes has been extensively discussed before (Clark 1988; Vandenberghe & Woo 2002). The claim that braided rivers are typical of cold environments has been opposed by many examples. Fluvial processes are operating at all latitudes but with different intensities, and cannot be distinguished from those that develop in other environments. In special cases, however, frozen rivers may provoke flow outbursts due to icing, a river flooding process that can only occur in periglacial conditions. Cryopediments should be regarded as typical periglacial forms, but the allusion to the semi-arid pediments already shows that the morphological appearance is not unique. What is unique is the intensity of the geomorphological processes that are responsible for their development and the frozen state of the substratum that enhances this intensity (Vandenberghe & Czudek 2008). It may be concluded that periglacial geomorphology has its specific identity that, according to French & Thorn (2006), should focus on the effects of ground ice and freezing processes. Once again, several books with this title have been published (Dixon & Abrahams 1992). Care should be taken not to overestimate its significance, however. We must take into account the fact that landforms often attributed to periglacial environments have equivalents in other environments, although they are often less pronounced.

Periglacial weathering and soil formation Temperature is a main factor in the process of weathering and organic matter production and decomposition, but is it more than just a determinant of the intensity of those processes? Frost shattering seems to occur only in periglacial conditions, although the exact processes are not always well known. Traditionally, frost– thaw fragmentation as a result of ice segregation and volume increase at the water-to-ice transformation phase has been considered as a typical process of ‘periglacial weathering’ (Yershov 1990; Murton et al. 2006). Its relative importance as compared to chemical weathering in

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cold environments has been questioned by Hall et al. (2002). The latter authors emphasize the role of moisture rather than temperature and conclude that ‘there is no zonality to cold environment weathering’. Consolidation of sediment as a result of alternating frost and thaw may be considered by definition as a process of ‘periglacial pedogenesis’. Such a consolidated soil horizon is called a ‘fragipan’ (Van Vliet-Lanoe¨ 1985). However, similar consolidation processes in soil horizons have also been reported from non-periglacial environments and thus may be due to processes other than cryogenic. Prismatic micro-jointing of otherwise unconsolidated sediments is also a phenomenon that is attributed to frost –thaw alternations by Mol et al. (1993), while reticulate structures are due to cryodesiccation after migration of unfrozen water at subzero temperatures (Yershov 1990). Tundra gleysols are frequently reported from periglacial environments. They require reduced soil conditions while the production of dissolved elements is hindered by low temperatures. Periglacial marshes therefore provide suitable conditions as they are waterlogged due to the presence of a (temporarily) frozen soil layer and do not show substantial migration of soil constituents due to limited chemical weathering. There is therefore some weathering due to the effect of frost and thaw, but again there is a warning against interpreting all weathering that takes place in cold regions as periglacial.

Periglacial sediments? If we consider that (with certain restrictions) there are periglacial landforms and land-forming processes, periglacial weathering and periglacial environments, it is tempting to consequently derive the existence of equivalent periglacial deposits. To answer the question whether sediments may also have their own periglacial identity, we must discuss some groups of sediments that were traditionally thought to be formed by genetically related, periglacial processes and, therefore, often called ‘periglacial sediments’.

Aeolian deposits Coversands. In their most typical form, coversands are described as mainly evenly laminated, rather well-sorted, slightly silty sands with a mean grain size of c. 150 mm (Ruegg 1983; Schwan 1986). They represent a characteristic element of the Pleistocene depositional system in the European lowlands, and occur as extensive aeolian sand sheets in periglacial environments. Deposition is locally

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interrupted by deflation and wind scouring. Occasionally, features indicative of a periglacial environment are found as, for instance, subvertical joints and seasonal frost cracks. Partial liquefaction and local water-escape structures may point to perched water tables that may have been caused by locally and seasonally frozen ground. The typical horizontal lamination is formed by the alternation of coarser and finer grained beds. While coarse-grained particles may result from saltation and creeping over short distance over a dry surface, the fine-grained sands and silts are probably transported over long distances in near-surface and high suspension clouds, respectively (Vandenberghe 1991). According to Schwan (1988) and Kasse (2002) this main coversand facies involves stepwise tractional transport over long distances under conditions of rare topographic barriers, sparse vegetation cover and a high ratio between wind energy and sand availability. One of the only coversand deposits reported from present-day periglacial environments (south Alaska) is from Lea (1990) who argues that coarse- to fine-grained couplets should represent different wind strengths. In the perspective of this paper, it is striking that the sedimentary conditions proposed by the above-mentioned authors are purely on the basis of geomorphological, sedimentological and wind characteristics. Scarcity of vegetation, strong winds and high sand availability may of course be induced by cold and dry conditions, but do not involve frost – thaw action. It is also illustrative that in today’s temperate regions, even during the Late Glacial and Holocene when the landscape was generally covered by vegetation, coversands were blown up locally, especially from vegetation-free areas such as elongated braided-river floodplains. Sand dunes. Sand dunes extend over much more restricted areas than coversands, that is, in places where vegetation is locally interrupted by bare surfaces (inland dunes and river dunes). They possess slope facies characterized by foreset cross-bedding and often occur in parabolic shape. From a sedimentological and geomorphological point of view, they cannot be distinguished from present-day coastal dunes (Fig. 1). High sand availability and the absence of vegetation in the source area are required, while the sand may be captured by the presence of vegetation or topographic barriers. A periglacial environment is certainly not necessary. Niveo-aeolian sands. Niveo-aeolian sands were reported from the middle of last century onwards, especially from ancient periglacial landscapes (Edelman 1947; Tricart & Cailleux 1967). Originally, the sediments that are indicated by this term are the finely laminated coversand facies that

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Asia, Smalley & Krinsley 1978; Ruth et al. 2003; Nugteren & Vandenberghe 2004; Vandenberghe et al. 2006). It has also been shown that loess deposition was ongoing in interglacial times in monsoonal climates (Liu 1985). The removal of loess by wind requires at least the presence of loosely packed silt-sized particles on a barren surface. This means, more specifically, absence of vegetation at a periodically dry surface and abundance of weathered (unconsolidated) material. This weathered material may have been assembled primarily by rivers in large alluvial fans and extensive lacustrine deposits before experiencing deflation (Lehmkuhl & Haselein 2000). The environmental conditions described are therefore not necessarily periglacial. However, availability of suitable sediment and appropriate surface conditions are essential.

Slope deposits

Fig. 1. (a) Drift sands blown up some 1000 years ago after general deforestation (top layer 1.5 m thick), overlying mid-Holocene podzol. They have exactly the same lithological characteristics as (b) the brown-coloured Younger Dryas dune sands (2 m thick) that formed in a periglacial environment occurring between the mid-Holocene podzol (top of the section) and an Allero¨d organic floodloam (middle of the section). Location: the Maas valley, south-eastern Netherlands (Bohncke et al. 1993).

we described above. The research of Schwan, Lea and others demonstrates that the role of snow in the formation of those deposits is non-existent. On the other hand, Koster & Dijkmans (1988) describe sand deposition on top of snow, probably transported by wind. After subsiding due to snow melting, these sands lost their original aeolian structures and become homogenized. For this origin of sediment, the term ‘niveo-aeolian’ is indeed applicable. The role of thaw is clear in defining the ultimate sediment properties (absence of structures and loose compaction), but this was only after deposition by wind. Loess. In the older literature, source regions from where loess has been blown up were assumed to have been located in the barren proglacial zones surrounding ice sheets (Smalley 1966; Mu¨cher 1986; Schwan 1988). However, it has been shown that loess may also find its origin in several other environments such as large floodplains (Ruhe 1975) and mid-latitude deserts (such as in Central

Geomorphological processes such as gelifluction and detachment sliding are discussed above. The associated sediments may be discussed in the same way for their assumed periglacial origin. The basic operating processes of gelifluction are completely identical to those of solifluction, although with different intensity due to different marginal conditions. Usually, soliflucted deposits cannot be distinguished from geliflucted deposits. Likewise, deposits resulting from detachment slides cannot be distinguished from the much more general epidermic debris sliding products. Theoretically, slope deposits that move by frost creep (as a result of frost heave) could be considered periglacial sediments, on the condition that they show properties that are typically the result of that periglacial process. The latter condition is, however, rarely the case. Similarly, particular frost-coated clast flows (He´tu et al. 1995) can only develop under repeated freezing, but are in fact only an expression of ubiquitous, cohesionless and matrixfree debris flows that are favoured under conditions of low temperature, absence of vegetation and abundance of fine debris. Frost-coated clast flow deposits cannot, in general, be distinguished from other dry grain flow deposits. Stratified slope deposits (gre`zes lite´es) were also traditionally interpreted as periglacial deposits, especially as the role of snow has originally been emphasized (Guillien 1951). They result from a large number of azonal slope processes (for instance grain and debris flow, slush flow), mostly involving surface runoff or even only gravitational grain flow without water to transport (van Steijn et al. 1995, 2002). The presence of snow is not essential as rain water or simply gravitational instability will provoke deposits with the same properties. Snow

PERIGLACIAL SEDIMENTS: DO THEY EXIST?

or frost may again initiate the transport process, but essentially that is not different from other factors such as the absence of vegetation and the availability of unconsolidated grains of appropriate grain size.

Fluvial deposits Fluvial processes, landforms and sediments, and especially their relation with climate, have been discussed intensively (Vandenberghe 2003). Coarsegrained deposits within terrace series have often been attributed to periglacial conditions. However, since it has been shown that there are no specific river types that may be connected with periglacial conditions, it is also obvious that there are no specific periglacial fluvial deposits (Vandenberghe & Woo 2002). Icing deposits are flood deposits that are characterized by their heterogeneous particle-size distribution and overall catastrophic appearance and, as such, have no equivalents in other than periglacial environments. It is a rare example of a truly periglacial deposit.

Thermokarst deposits Thermokarst depressions are typical periglacial landforms. Primordially, they are caused by the melting of ground ice leaving a depression filled with water. The process is undoubtedly periglacial in nature. Sedimentation in the depression, however, has no relation to the origin of the depression. Often the lake is filled after the cold period in which the ground ice was formed and during the warm period that was the cause of the ice melting. Examples are the pingo remnants that formed at the end of the last glacial and filled during the Late Glacial and Holocene. Such sediments have no single periglacial characteristic and are not different from sediments filling other kinds of hollows.

Cryogenic in situ weathered deposits In the sense that a sediment layer may be characterized by certain physical-chemical properties due to specific periglacial weathering processes, it may be called a periglacial sediment. As stated above, it is often difficult to define the precise weathering mechanism even if the weathering process is zonal or azonal. Frost-shattered rock fragments forming a sediment layer, for instance a block field, may be an example. Thus, the so-called ‘periglacial facies’ of Lozinski (French 2000) might be considered a periglacial sediment. Such sediment layers have undergone a complete transformation from solid rock to unconsolidated rock fragments and, as such, have attained a completely new

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identity. In that sense, rock fragmentation is not post-depositional but syn-depositional. Postdepositional weathering or soil formation that is not essentially changing the sediment characteristics should not be considered as periglacial, as discussed below. The transition between ‘complete transformation’ and ‘essentially unchanged sediment characteristics’ is however gradual, for instance in the case of brecciated chalk deposits (Murton et al. 2006). In addition, it is difficult to separate frost-induced from non-frost-induced rock fragmentation and, even more, rock layers that were weathered by frost during and after deposition by means of their sediment properties.

Discussion The global occurrence It appears that deposits that have frequently been considered as ‘periglacial’ show only the effects of depositional processes that may occur globally. River, wind and slope processes are indeed ubiquitous. Nevertheless, the fact that certain deposits occur preferentially in cold environments cannot be neglected. Aeolian deposits, as coversands, may be an example. Apparently, the formation of such deposits is favoured by periglacial conditions. For wind action, barren surfaces are required in combination with a certain amount of unconsolidated sediment with a suitable grain size. This is in addition to certain atmospheric and topographical conditions. Of course, vegetation cover may be poorer in periglacial areas (the so-called polar deserts) than in tropical regions. However, that only means that certain (aeolian) processes are favoured in periglacial environments, but not caused by periglacial (frost –thaw) processes. Similar arguments hold for slope and river deposits. For instance, gelifluction and frost creep are only special cases of solifluction and soil or rock creep that are physically well-defined, widely occurring processes. The fact that those general processes are more frequent in the presence of a slip face (on top of an ice body) or meltwater does not change the essence of the process. More important for the present paper is that even the properties of the transported deposits are not different: for instance a deposit due to gelifluction or frost creep cannot be distinguished from one due to solifluction or general creep, respectively. It may be concluded that the described sediments are not unambiguously typically periglacial in nature.

Post-depositional transformation On some occasions specific sediments seem to show characteristics that point to periglacial conditions.

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Those sediments have a definite post-sedimentary periglacial imprint in common. Some of them are weathered by frost –thaw processes or are affected by soil formation under cold conditions, while others are compacted due to thaw consolidation or homogenized after deposition due to snow melt (niveo-aeolian deposits). Deformed sediments such as cryoturbated layers (not discussed above) can be added to this type of change. It is a crucial point of discussion whether sediments that obtained their periglacial imprint after deposition should be called ‘periglacial sediments’. As an extreme example, it would feel very strange to call exhumed interglacial coastal deposits that were cryoturbated or penetrated by thermal-contraction wedges during a later cold period a ‘periglacial sediment’ (Fig. 2a, b). Sediments that change their properties as a result of ice growth or decay may be treated similarly. It cannot be denied that, for instance, sediment consolidation and permeability may change temporarily or definitively by ice growth and melting. The presence of ice lenses, and certainly their thaw, is however post-depositional. It is tempting to call a geliflucted deposit full of ice a periglacial deposit. But to describe, for instance, a former Tertiary beach deposit that is now full of ice as a periglacial deposit is as questionable as to designating such a deposit affected by a network of ice-wedge casts as a periglacial deposit. The frozen state of a sediment certainly affects the sediment properties,

but this is no different from the varying degrees of water saturation of a sediment layer between completely dry and over-saturated. Post-depositional periglacial processes should therefore not be included in the definition of ‘periglacial sediments’. This is in contrast to the list made by French (2001) that includes all kinds of post-depositional periglacial processes as being responsible for periglacial sediments.

Conclusions It is argued that periglacial conditions, landforms and processes do exist, even when using different definitions of ‘periglacial’. The existence of ‘periglacial sediments’ is a more subtle question, however. Many geomorphological processes and forms may be active in cold climates, while not being restricted to these environments. Their degree of activity (expressed for instance in terms of magnitude-frequency combinations) may show a periglacial imprint (van Steijn et al. 2002) and is sufficient for French (2001) to designate the resulting deposits as ‘periglacial sediments’. It is obvious that the answer to the question raised in the title of this paper is dependent on the definition of ‘periglacial’. If ‘periglacial deposits’ are defined as and restricted to ‘sediments that possess distinct characteristics related to either frost or its subsequent thaw’ (as argued in the Introduction), it may in fact be concluded that ‘periglacial sediments’

Fig. 2. (a) Pliocene sands (Scheemda Formation) deposited in warm-temperate conditions are penetrated by Pleistocene ice-wedge cast (northern Netherlands). (b) Estuarine clays of Early Pleistocene age are heavily cryoturbated during the last glacial maximum (Meerle, northern Belgium) (Vandenberghe 2006). Are they periglacial deposits?

PERIGLACIAL SEDIMENTS: DO THEY EXIST?

do not exist with the exception of a few specific cases (e.g. icing deposits and layers composed of frost-fragmented rock debris). It also means that the described sediments possess globally valid properties that are not diagnostic for periglacial environments. This statement assumes that, in the definition of ‘periglacial sediment’, post-depositional processes such as micro- and macro-deformations (for instance cryoturbations, sediment consolidation) and in situ weathering or diagenesis are excluded. In contrast, incorporating all ‘cold, non-glacial processes’ in the production of ‘periglacial sediments’ would also mean including post-depositional processes and hence would broaden the concept of ‘periglacial sediments’. Hugh French is warmly thanked for fruitful discussions.

References Bohncke, S., Vandenberghe, J. & Huijzer, A. S. 1993. Periglacial palaeoenvironments during the Weichselian Late Glacial in the Maas valley, the Netherlands. Geologie en Mijnbouw, 72, 193–210. Clark, M. J. 1988. Periglacial hydrology. In: Clark, M. J. (ed.) Advances in Periglacial Geomorphology. Wiley and Sons, Chichester, 415– 462. Derruau, M. 1972. Pre´cis de Ge´omorphologie, 5th edn. Masson et Cie, Paris. Dixon, M. J. C. & Abrahams, A. D. 1992. Periglacial Geomorphology. Wiley & Sons, Chichester. Edelman, C. 1947. Les limons et les sables de couverture des Pays-Bas-La ge´ologie des terrains re´cents. Session Extraordinaire Socie´te´ Belge de Ge´ologie, Bruxelles, 1946, 303–310. Embleton, C. & King, C. A. M. 1975. Periglacial Geomorphology. Arnold, London. French, H. M. 2000. Does Lozinski’s periglacial realm exist today? A discussion relevant to modern usage of the term ‘periglacial’. Permafrost and Periglacial Processes, 11, 35–42. French, H. M. 2001. Periglacial sediments. In: Matthews, J. A. (ed.) The Encyclopaedic Dictionary of Environmental Change. Arnold, London, 474– 475. French, H. M. 2007. The Periglacial Environment, 3rd edn. Wiley, Chichester. French, H. M. & Thorn, C. E. 2006. The changing nature of periglacial geomorphology. Ge´omorphologie: Relief, Processus, Environnement, 3, 165– 174. Guillien, Y. 1951. Les gre`zes lite´es de Charente. Revue ge´ographique des Pyrene´es et du Sud-ouest, 22, 153–162. Hall, K., Thorn, C. E., Matsuoka, N. & Prick, A. 2002. Weathering in cold regions: some thoughts and perspectives. Progress in Physical Geography, 26, 577–603. He´tu, B., Van Steijn, H. & Bertran, P. 1995. Le roˆle des coule´es de pierres se`ches dans la gene`se d’un certain type d’e´boulis stratifie´s. Permafrost and Periglacial Processes, 6, 173– 194. Kasse, C. 2002. Sandy aeolian deposits and enviroments and their relation to climate during the last glacial

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maximum and Lateglacial in northwest and central Europe. Progress in Physiscal Geography, 26, 507– 532. Koster, E. & Dijkmans, J. W. A. 1988. Niveo-aeolian deposits and denivation forms, with special reference to the Great Kobuk Sand dunes, Northwestern Alaska. Earth Surface Processes and Landforms, 13, 153– 170. Lea, P. 1990. Pleistocene periglacial eolian deposits in southwestern Alaska: sedimentary facies and depositional processes. Journal Sedimentary Petrology, 60, 582– 591. Lehmkuhl, F. & Haselein, F. 2000. Quaternary paleoenvironmental change on the Tibetan Plateau and adjacent areas (Western China and Western Mongolia). Quaternary International, 65/6, 121– 145. Lewkowicz, A. G. 1992. Factors influencing the distribution and initiation of active-layer detachment slides on Ellesmere Island, Arctic Canada. In: Dixon, J. C. & Abrahams, A. D. (eds) Periglacial Geomorphology. Wiley & Sons, Chichester, 223– 250. Linton, D. L. 1969. The abandonment of the term ‘periglacial’. Palaeoecology of Africa & of the Surrounding Islands & Antarctica, 5, 65–70. Liu, T. 1985. Loess and the environment. China Ocean Press, Beijing. Mol, J., Vandenberghe, J., Kasse, C. & Stel, H. 1993. Periglacial microjointing and faulting in Weichselian fluvio-aeolian deposits. Journal of Quaternary Science, 8, 15–30. Mu¨cher, H. J. 1986. Aspects of loess and loess-derived slope deposits: an experimental and micromorphological approach. PhD thesis, University of Amsterdam. Murton, J., Peterson, R. & Ozouf, J.-C. 2006. Bedrock fracture by ice segregation in cold regions. Science, 314, 1127– 1129. Nugteren, G. & Vandenberghe, J. 2004. Spatial climatic variability on the central Loess Plateau (China) as recorded by grain size for the last 250 kyr. Global and Planetary Change, 41, 185– 206. Ruegg, G. 1983. Periglacial eolian evenly laminated sandy deposits in the Late Pleistocene of NW Europe, a facies unrecorded in modern sedimentological handbooks. In: Brookfield, M. & Ahlbrandt, T. (eds) Eolian Sediments and Processes. Elsevier, Amsterdam, 455–482. Ruhe, R. V. 1975. Geomorphology. Houghton Mifflin Co., Boston. Ruth, U., Wagenbach, D., Steffensen, J. & Bigler, M. 2003. Continuous record of microparticle concentration and size distribution in the central Greenland NGRIP ice core during the last glacial period. Journal of Geophysical Research, 108, 4098, doi: 10.1029/2002JD002376. Schwan, J. 1986. The origin of horizontal alternating bedding in Weichselian aeolian sands in north-western Europe. Sedimentary Geology, 49, 73–108. Schwan, J. 1988. The structure and genesis of Weichselian to Early Holocene aeolian sand sheets in western Europe. Sedimentary Geology, 55, 197– 232. Smalley, I. J. 1966. The properties of glacial loess and the formation of loess deposits. Journal of Sedimentary Petrology, 35, 669–676.

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Smalley, I. J. & Krinsley, D. H. 1978. Loess deposits associated with deserts. Catena, 5, 53–66. Tricart, J. & Cailleux, A. 1967. Le modele´ des re´gions pe´riglaciaires. SEDES, Paris. Van Steijn, H., Bertran, P. & Francou, B. 1995. Models for the genetic and environmental interpretation of stratified slope deposits: review. Permafrost and Periglacial Processes, 6, 125–146. Van Steijn, H., Boelhouwers, J., Harris, S. & He´tu, B. 2002. Recent research on the nature origin and climatic relations of blocky and stratified slope deposits. Progress in Physical Geography, 26, 551– 575. Van Vliet-Lanoe¨, B. 1985. Frost effects in soils. In: Boardman, J. (ed.) Soils and Quaternary Landscape Evolution. Wiley, Chichester, 117– 158. Vandenberghe, J. 1991. Changing conditions of aeolian sand deposition during the last deglaciation period. Zeitschrift fu¨r Geomorphologie, 90 (Suppl.), 193–207.

Vandenberghe, J. 2003. Climate forcing of fluvial system development: an evolution of ideas. Quaternary Science Reviews, 22, 2053–2060. Vandenberghe, J. 2006. Cryoturbation structures. In: Elias, S. A. (ed.) Encyclopedia of Quaternary Science. Elsevier, Amsterdam, 2147– 2153. Vandenberghe, J. & Woo, M. K. 2002. Modern and ancient periglacial river types. Progress in Physical Geography, 26, 479–506. Vandenberghe, J. & Czudek, T. 2008. Pleistocene cryopediments on variable terrain. Permafrost and Periglacial Processes, 19, 71–83. Vandenberghe, J., Renssen, H. et al. 2006. Penetration of Atlantic westerly winds into Central and East Asia. Quaternary Science Reviews, 25, 2380–2389. Yershov, E. D. 1990. General Geocryology (English translation, Williams, P. J. 1998). Cambridge University Press, Cambridge.

Stratified slope deposits: periglacial and other processes involved H. VAN STEIJN Department of Physical Geography, Faculty of Geosciences, Utrecht University, Utrecht, The Netherlands (e-mail: [email protected]) Abstract: Stratified slope deposits occur on or at the base of slopes within a wide range of hilly or mountainous environments. Gelifraction, a lack of vegetation, and snowmelt-derived slopewash are thought to be important factors in their development. The relatively fine-grained gre`zes lite´es found at many sites within the French Charente region are the prototype of this kind of deposit. The original, strictly limited and genetically intended definition has been enlarged by mainly non-French authors causing considerable confusion. Other assemblages of cold-climate slope deposits are more heterogeneous, sometimes less clearly stratified, and show much larger variation in grain sizes and sedimentary structures than gre`zes lite´es. The most important of these are stone-banked solifluction (sheets or lobes), rock fall, (dry) grain flow, frost-coated clast flow, debris flow, snow avalanching and, to a lesser degree, rain-generated overland flow and (niveo-) aeolian activity. Many of these processes are azonal and occur under a wide range of environmental conditions. A periglacial context may nevertheless influence process activity in these cases, for instance with regard to their magnitude and/ or frequency.

Stratified slope deposits are heterogeneous in all senses – parent rock, granulometric properties, degree of stratification, stratification types, processes involved, depositional environments and built-in climate signal. They are found in highmountain environments but also in gently sloping hilly landscapes, in both cases generally as aprons on the lower parts of the slopes. Since Guillien’s first publication (Guillien 1951) about the fine-grained gre`zes lite´es found at many sites within the French Charente region, a periglacial connotation has been associated with this type of deposit. Earlier, Cailleux (1948) introduced the term e´boulis ordonne´s to describe similar deposits showing a regular alternation of coarse-grained, generally openwork units and fine-grained units consisting of a sandy to silty matrix containing small gravels. It was Tricart (1967) who implicitly stressed the periglacial significance of stratified slope deposits. The term ‘periglacial’ covers a wide range of conditions for ground frost (seasonal or perennial) and snow, together with several azonal processes (French 2007). All may be active within a given region, although with potentially different intensities and timing. A related term is ‘paraglacial’, a term initially used by Ryder (1971) to describe accelerated geomorphological process activity in response to retreating glacier ice. Church & Ryder (1972) subsequently refined the concept to deal with processes of a non-glacial nature directly conditioned by the former presence of ice. Such conditions exist when glacial ice disappears from a region, described as a ‘paraglacial period’

(Ballantyne 2002). The influence of paraglacial conditions on bedrock shattering is important.

The historic context Guillien (1951) proposed the first genetic interpretation: highly active gelifraction, lack of vegetation and snowmelt-induced slopewash were thought to be the most important factors in the development of gre`zes lite´es (Figs 1 & 2). Gre`zes lite´es are a finetextured material consisting of small, angular limestone clasts (maximum size 25 mm with a mode of c. 2.5 mm) and a fines fraction (,0.5 mm) that amounts to 10–20% by weight (Ozouf et al. 1995a). In the Charente region, the deposits have formed over carbonate bedrock (Bertran et al. 1992). They are several metres thick, formed by a regular alternation of matrix-rich layers with matrix-supported clasts (lits gras) and openwork layers (lits maigres). The thickness of individual units ranges between 2 and 25 cm, but are mostly between 10 and 15 cm. Layers dip at angles of between 128 and 168. The upper boundaries of the matrix-rich layers are flat and sharp, whereas their lower boundaries are less sharp and often festooned. Inverse grading is often found within these layers. Normal grading may occur within the openwork coarser grained upper layer, but both normal and inverse grading is found in places. The original, granulometrically limited and genetically intended definition by Guillien (1951) has been subsequently enlarged, causing considerable confusion. The e´boulis ordonne´s concept

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 213–226. DOI: 10.1144/SP354.14 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Techniques and terminology

Fig. 1. General view of gre`zes lite´es deposits, gravel pit near Verteuil, Charente, France. Visible scale length about 2 m.

(Cailleux 1948) for instance was found inapplicable in the case of the Charente gre`zes lite´es, essentially because of differences in grain size and slope steepness (Ozouf et al. 1995a). Dewolf (1988) summarized ideas existing about the origin of stratified slope deposits at that time. More recent research has shown that other processes, not strictly periglacial in nature, have been involved in the origin of many relict stratified slope deposits (Van Steijn et al. 1995; Nemec & Kazanci 1999). As stated by French & Thorn (2006), the ‘periglacial’ concept has become broader and includes both permafrost (the ‘cryological’ side of the periglacial domain) and seasonal frost. It is the non-permafrost context that dominates the development of stratified slope deposits. As a consequence, the role of azonal processes has to be evaluated. Azonal processes include overland flow, landslides or debris flows. The interplay of (periglacial) zonal and azonal processes is probably the most difficult question with regard to the interpretation of stratified slope deposits.

Fig. 2. Gre`zes lite´es, gravel pit near Verteuil. Succession of lits gras and lits maigres. Longitudinal section. Length of compass 10 cm.

From the sedimentological perspective, exposures both at the surface and in-section should be recorded and discussed. Studies such as those by Blikra & Nemec (1998) and Nemec & Kazanci (1999) provide examples of detailed sedimentological analysis that combine graphic logs, block diagrams and photographs. Additional data may be obtained from (macro) fabric analysis and micromorphological observation (Van Vliet-Lanoe¨ 1987; Bertran et al. 1997; Harris 1998; Bertran & Texier 1999). Recently, geophysical methods have been used to recognize stratification within slope deposits, although at a relatively low resolution (Vo¨lkel et al. 2001; Vo¨lkel & Mahr 2001; Leopold & Vo¨lkel 2003; Leopold et al. 2008). Radiocarbon dating is often applied to older deposits (Blikra & Nemec 1998; Nemec & Kazanci 1999; Garcı´a-Ruiz et al. 2001; Eaton et al. 2003) whereas dendrochronology and lichenometry may be applied to relatively young material (McCarroll 1993; McCarroll et al. 1995; Jomelli 1997; Lafortune et al. 1997; He´tu & Gray 2000). Palaeobotanical data are used to verify general climatic trends obtained from stratified slope deposits (Blikra & Selvik 1998; Garcı´a-Ruiz et al. 2001). With regard to terminology, the terms ‘colluvium’ or ‘stratified colluvium’ are sometimes used for the slope deposits discussed in this article. These are synonyms for ‘(stratified) slope deposits’ and ‘(stratified) scree-slope deposits’. ‘Colluvial cones’ are equivalent to ‘scree slopes’ and ‘talus’ while ‘colluvial aprons’ are landforms resulting from coalescent colluvial cones. Blikra & Nemec (1998) summarize the different terms in use in the sedimentological and geomorphological literature.

Debris production and stratified slope deposits Recent reviews of relict slope deposits have concentrated on diagnostic sedimentary characteristics. These allow recognition of the processes involved in sediment transport and deposition rather than the climatic information that these sediments contain (Van Steijn et al. 1995, 2002; Bertran et al. 1995a; Bertran 2005). Blikra & Nemec (1998) and Blikra & Selvik (1998) focus on snow avalanches and debris flows to interpret relict deposits of Holocene or older ages in Norway. Nemec & Kazanci (1999) discuss several types of grain-flow and overland-flow processes in Anatolia, Turkey. These studies emphasize the palaeoclimatic significance of stratified slope deposits. Gelifraction is thought to be an important process for the production of material found in

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stratified slope deposits (Cailleux 1948; Guillien 1951; Tricart 1967; Jomelli et al. 2003, 2004). Gelifraction, or frost shattering, is a complicated process dominated by ice formation and bedrock properties. The role of ice segregation in bedrock weathering is discussed by Hallet et al. (1991) while Matsuoka (1991) and Matsuoka et al. (1997) discuss the integral effect of a number of factors. Hall (1999) and Hall et al. (2002) treat the role of thermal stress in rock fatigue. Water from rain or snowmelt also contributes to the detachment of particles that are already loosened by other processes (Luckman 1988; He´tu & Vandelac 1989; Pe´rez 1998). This is reflected, for example, by increased rockfall activity during rainy periods (He´tu 1990, 1991; Pe´rez 1998). A similar role in clast detachment is played by strong winds, as reported by He´tu (1990, 1991). Bedrock properties are of paramount importance to the nature of stratified slope deposits. Lithological influences such as bedding or grain size are most often reported (Guillien 1951; Tricart 1967; Karte 1983; He´tu & Vandelac 1989; Gardner et al. 1991). Weakening by tectonic effects or paraglacial stress release is frequently mentioned. He´tu & Vandelac (1989) stress the role of outcropping thin shale beds as well as a dense fracture pattern. A paraglacial aspect is stress release within bedrock due to debuttressing of glacially steepened rock walls (Ballantyne 2002); in formerly glaciated regions this may be one factor in frost shattering during the paraglacial period (Anderson & Harrison 2007). All these factors contribute to the production of angular clasts in a wide range of grain sizes. There are other ambiguities. For example, Bertran et al. (1995b) observed a decrease in clast size in a downslope direction in Charente gre`zes lite´es. They suggest this decrease is caused by ongoing fragmentation of the debris due to lengthening of the transit surface. Second, in the case of the rhythmically stratified slope deposits described by Dylik (1960, 1969), the gravels, sands and silts are derived from older glaciofluvial deposits. Likewise, the material described by Francou (1988, 1990) and Bertran et al. (1995a) originated from moraines and talus. In these cases there is no direct relationship between the production of sediment and the development of stratified slope deposits. Third, for the stratified colluvium studied by Nemec & Kazanci (1999) in Anatolia, slope-process activity during the latest Pliocene and Early Pleistocene period was characterized by a warm-humid climate with pronounced phases of dryer conditions. This means that, at least during that early period of local colluvium development, gelifraction would have been absent or weak. Finally, many accumulations of stratified slope deposits show different facies (He´tu 1991; Bertran

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et al. 1995b; Blikra & Selvik 1998; Nemec & Kazanci 1999; Eaton et al. 2003). This suggests different environments connected to the depositional history of the deposit in question (Marion et al. 1995; Blikra & Selvik 1998; He´tu & Gray 2000).

Sediment types observed on modern scree slopes Based on their occurrence in the field and on climate and climate-dependent sedimentary processes, three groups of stratified deposits may be distinguished: (a) modern cold-climate (‘periglacial’) and other deposits; (b) relict periglacial deposits; and (c) deposits from azonal processes. The following illustrates the high diversity of processes that contribute to talus accumulations and stratification patterns and the wide range of environments within which stratified slope deposits may develop. Several processes are active on a scale that is not ‘suitable’ to deliver sediments to taluses or comparable accumulations. This is especially true for most mass movements such as slides, slumps or earthflows. Deposits of rock avalanches are observed at the internal base of several talus cones, upon which subsequent processes finally build slope deposits showing stratification. For example, Nemec & Kazanci (1999) mention chaotic accumulations of very heterogeneous boulders and gravels, sometimes with boulders of several metres length, present at the base of colluvial aprons.

Frost creep and solifluction: stone-banked sheets and lobes Francou (1988, 1990) studied active stratified slope deposits at altitudes of c. 5000 m in the Central Andes where the mean annual air temperature (MAAT) is near 0 8C. The climate of these regions is characterized by a large number of shallow freeze –thaw cycles and an often-occurring but ephemeral snow cover. Melting snow favours needle-ice development but also causes overland flow. Permafrost is absent (Francou 1989). Francou (1988) used the term nappe-coule´e (‘stonebanked solifluction sheet or lobe’, Benedict 1976; Francou 1990) to describe these deposits (Fig. 3) that generally extend some tens of metres in both transverse and longitudinal directions. Their thickness is variable; individual sheets consist of a matrix-rich layer c. 10 cm thick which covers an openwork unit of a variable thickness from 5 to 50 cm. Matrix-rich layers show a sharp upper boundary and a clear to diffuse lower boundary while openwork layers show grading which is usually normal but in some cases inverse. Francou

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Fig. 3. Stone-banked sheets, La Cumbre, 5000 m altitude, Bolivian Andes (from Bertran & Coutard 2004).

(1988, 1990) explained the stratification by the advance of the matrix-rich unit which progressively overruns the frontal stone bank. Meanwhile, the stone bank is being reconstituted by frost sorting of the coarse clasts and eluviation of fines. Movement is caused mainly by frost creep during the dry season and gelifluction in the wet season. In this way the frontal part is overrun, during which its original inverse grading may become reversed. Slope gradients vary between 48 and 358 (Francou 1988; Bertran et al. 1995a). The sheets develop from talus accumulations and morainic material, initially derived from highly gelifraction-sensitive rock types (Francou 1989). Figure 4 schematically

shows the internal organization of stone-banked solifluction sheets, together with the comparable internal structure of alpine stone-banked lobes (see below). At higher latitudes, for instance in the European Alps, stratified slope deposits showing a comparable alternation of matrix-rich and openwork units take the form of stone-banked lobes (Bertran et al. 1993, 1995a; Todisco et al. 2000). These lobes are much smaller than the sheets discussed above, probably due to the lower frequency of freeze –thaw cycles and longer lasting snow cover. Bertran et al. (1995a) note that lobes are imbricated, having a length of less than 15 m and widths of between 0.5 and 3 m; cross-sections reveal lenticular stratification.

Rockfall The contribution of rockfall to stratification in slope deposits is difficult to identify. Nevertheless, rockfall is often mentioned as one of the processes participating in modern sediment transfer on scree slopes (Luckman 1988; Pe´rez 1998). For example, He´tu & Vandelac (1989) found that large seasonal differences in rockfall activity in Gaspe´sie, Canada were caused not only by the different roles of snow, rain and wind, but also by clast sizes and forms. Comparable observations are reported by Pe´rez (1998) who describes how rockfall intensity on Californian talus slopes is influenced by moisture

Fig. 4. Schematic sections: (a) Alpine stone-banked lobe and (b) Andean stone-banked sheets. Matrix-rich layers in white. Coarse, openwork layers often show grading in two directions caused by their composite origin. Matrix-rich layers are lenticular in frontal sections. In the case of the Andean stone-banked sheets, matrix-rich layers are continuous and their lower boundary displays folds in longitudinal sections and a wavy pattern in frontal sections (from Van Steijn et al. 1995).

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supply and daily temperature fluctuations on slopes with northern exposures. Dorren (2003) reviews models of rockfall mechanics and their consequences for runout distances and sorting effects while small-scale laboratory experiments by De Blasio & Saeter (2009) show single-particle pathways, the development of normal grading, a general fining in the upslope direction and the subsequent triggering of (dry) grain flows (see below).

Grain flow (cohesionless debris flow) Under subaerial conditions, grain flows are generally characterized as dry when the interstitial fluid is mostly air. Modified grain flows occur where a fine-grained matrix and/or interstitial water participate in the transmission of the particles. Cohesionless and cohesive flows are two forms of sediment gravity flows (Lowe 1976) that participate in the development of stratified colluvium. Common types of cohesionless flows are (dry) grain flows and frost-coated clast flows (He´tu et al. 1994). Cohesionless grains stay dispersed against gravity by particle collisions. In cohesive flows, the strength of the fine matrix helps to keep the coarser particles separated.

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Dry grain flows are typical of the upper part of a talus where slope angles are steepest (generally exceeding 358) and clast accumulations are unstable. Short-lived flows covering limited distances are characteristic. They are generally triggered by the impact of individual rock falls on previously accumulated clasts at the base of a rock wall. The elongated flows are generally only a few metres wide and they stop moving en masse as soon as the slope angle falls below 34–358 because of interparticle friction. During transport, the heterogeneous material is reorganized as described by He´tu et al. (1995). Kinematic sieving (Allen 1972; Carniel & Scheidegger 1974) and dilatancy due to particle collisions are thought responsible for the changing internal structure of a dry grain flow. Together with the vertical velocity gradient present within the moving flow, these processes cause the development of a fine-grained slip plane at the upslope end of a dry grain flow (He´tu et al. 1995). Figure 5 schematically shows this development. Frost-coated clast flows are modified dry grain flows. They were first mentioned by He´tu & Vandelac (1989) and described by He´tu et al. (1994). They owe their special character to ice coatings around

Fig. 5. Schematic longitudinal section showing internal organization of a dry-grain-flow deposit. Different parts and materials and some examples of macrofabrics (circle diagrams) are shown. GF1– 3 represent individual grain flows. (a– c) Schematic vertical sections at indicated parts of the deposit (arrows). Numbers in (b) represent textures with 1, matrix-rich bed with finer clasts; 2, coarser, openwork material; 3, tail of fine gravel deposited at the end of the event on top of the already immobilized coarser material. Upper fabric diagram, lobe from frontal zone of an artificial flow at Belesten (French Pyrenees); lower fabric diagram, from slip plane in artificial flow in quarry near Remollon, French Alps (modified from Van Steijn et al. 1995).

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the clasts, which reduce the friction between clasts and between clasts and the base over which they are moving. In this way, the material remains mobile on relatively gentle slopes. For example, He´tu et al. (1994) report that, on taluses in Gaspe´sie, motion only stopped when the slope angle had diminished to 24 –268. The climate of the region is characterized by long and cold winters and short warm summers. Snow amounts to 38% of the yearly precipitation with mean January temperature of 210 8C, and there are about 41 days per year with at least one freeze –thaw cycle. Mean annual air temperature (MAAT) is 3 8C. According to He´tu et al. (1994), meteorological conditions that favour the occurrence of frost-coated clast flows are cold winters during which short periods of thaw are accompanied by rainfall amounts that exceed 10–20 mm/day. Furthermore, heavy winds blow the unvegetated parts of scree slopes free of snow. Clasts may develop fragile and partly interconnected ice coatings under such conditions. Relatively small disturbances, most often rockfall impacts, may trigger the onset of a frost-coated clast flow. Morphologically, frost-coated clast flow deposits are characterized by shallow but elongated channels bordered by levees comparable to debrisflow levees. Generally, the flows end with a welldeveloped terminal lobe. Examples of morphological elements and corresponding sedimentological characteristics are given in Figures 6 and 7. All three morphological elements show characteristic clast fabrics (He´tu et al. 1994; Bertran et al. 1997). He´tu & Gray (2000) conclude that, on many of the Gaspe´sian scree slopes, frost-coated clast flows are an important sediment-transfer process.

Fig. 6. Frost-coated clast flow deposits: channel flanked by levees seen in upslope direction. Coarse clasts dominate the levees, while mean gravel size on the channel bottom is smaller. Slope angle is about 278. Scale bar is in decimetres.

Debris flows Debris flows are a common phenomenon of many cliff-talus slope systems. These rapidly flowing masses of poorly sorted granular solids, water and air leave elongated, relatively narrow tracks along the talus down which they are running. Debris-flow rheology may be described as either visco-plastic (Johnson & Rodine 1984) or by a model combining dispersive pressure caused by particle collisions and turbulence (Takahashi 1978, 1980). The former model is suitable for cohesive debris flows (Lowe 1976) moving high sediment concentrations with high clay content. Because of the plastic component inherent in this model, the central part of such a debris flow shows uniform velocity and is visible as a ‘rigid plug’. Blikra & Nemec (1998) mention a few examples of what they call ‘high-viscosity debris flows’ where the geometry of the related deposits is described as broad lobes (Blikra & Nemec 1998). A further characteristic of this type of debris flow is the presence of large ‘floating’ cobbles or boulders on the lobes or lateral levees (Johnson & Rodine 1984; Blikra & Nemec 1998). The Takahashi model (Takahashi 1978, 1980) seems suitable to describe debris-flow activity on cliff-talus systems. In this model, larger clasts tend to move towards zones of low shear at the outer and front parts of the flow. Debris-flow tracks show a limited number of morphological units. The upper part of the track on the talus is generally incised. The channel bottom may contain sediments from recent events, mainly along the channel sides but sometimes in the form of lobes on a part of the channel floor. Further downslope, lateral levees develop that finally join to form a terminal lobe. These morphological units together with the

Fig. 7. Terminal lobe of frost-coated clast flow in a longitudinal view. Openwork clasts showing strong imbrication (clasts dipping in upslope direction) due to ‘braking’ effect when the flow stopped. Slope angle at lobe base about 268. Upper scale in decimetres (from Van Steijn et al. 1995).

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Fig. 8. Variations of sediment properties along the track of debris-flow deposits. Example from Vallon Laugier, French Alps (Bertran & Texier 1994). (a) Proximal sections show well-developed openwork lenses deposited as outer parts of lateral levees or channel-lag deposits. (b) Matrix-rich deposits represent the central part of a flow. (c) Sections in distal parts mostly show lenticular, diamictic units, while openwork layers are rare. Alternation of frontal lobes with palaeosols indicate the infrequent nature of debris-flow activity in this example (modified from Van Steijn et al. 1995).

often-shifting position of the debris-flow tracks on a given talus cause an intricate lenticular assembly of sediments (Van Steijn et al. 1988; Bertran & Texier 1994; Blikra & Nemec 1998; Nemec & Kazanci 1999). Figure 8 schematically shows an arrangement of debris-flow components. Matrix-rich and openwork units, both showing different types of grading and clast fabrics, occur close to each other within debris-flow deposits. Inner levee sides tend to show matrix-containing gravely material while their outer parts are openwork accumulations of coarser clasts, as observed by Nieuwenhuijzen & Van Steijn (1990). Channel-bottom deposits are diamictic (when directly deposited during an event) or openwork (when left as lag deposit or caused by subsequent eluviation). Fabric analysis may lend support to interpretation of the deposits (Bertran et al. 1997).

Snow avalanches Snow avalanches (or snowflows) have been analysed by Blikra & Nemec (1998) and Blikra & Selvik (1998). They combine observations of modern snow avalanches with that of relict deposits visible in gravel pits. The rheological behaviour of snow is complicated by the degree of compaction, or its

thermodynamic instability. Mechanical properties also vary. Snow is regarded by Blikra & Nemec (1998) as a plastic material, the shear strength of which ranges from mainly frictional to mainly cohesive. For example, dry, powdery or granular snow behaves mainly as a plastic material (at least at temperatures lower than 25 8C) while cohesive forces are more important in damp or wet snow. As a consequence, different types of snow avalanches and related deposits exist (Blikra & Nemec 1998). Dense snow avalanches usually start as translational slides, moving a broad slab downslope, and often disintegrate into gravel-sized clods or snow balls if the snow is cohesive. These avalanches move like cohesive debris flows, revealing high shear strength and thus large ‘rigid plugs’. This is reflected in the way they transport debris, which is comparable to that of debris flows. Avalanches of powder snow move mainly in a turbulent manner. They develop high velocities; frontal velocities of 12– 60 m s21 are reported by Blikra & Nemec (1998). Snow-avalanche activity depends on climate and on mountain morphology. Precipitation characteristics, temperature and wind are important factors, as are altitudinal range and slope aspect. The latter plays an important role in relation to prevailing winds that determine preferred triggering sites for

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avalanches (Jomelli 1997; Blikra & Nemec 1998). Several debris sources may contribute to the sediment content of snow avalanches. Rockfall from adjacent cliffs, but also material eroded from the slope or chute along which an avalanche moves, are important contributors as is erosion of the apex of the talus on which the avalanche is running out. Sometimes, aeolian transport supplies material. All sources may contribute in highly variable amounts. The physical properties of avalanche snow and its highly variable debris load cause highly variable deposits (Blikra & Nemec 1998). Fresh deposits are elongate covers of pre-existing topography, often only one clast thick. Bed geometry is highly uneven and discontinuous; often loosely packed coarse clasts left by a snow avalanche are surrounded by finer material, usually due to post-avalanche infilling. A further characteristic is the presence of large clasts when compared to bed thickness (Blikra & Nemec 1998; Blikra & Selvik 1998). Jomelli (1997, 1999a, 1999b), Jomelli & Bertran (2001) and Bertran & Jomelli (2004) analysed comparable snow-avalanche accumulations in the French Alps. The effect of snowmelt on the fabric properties of dirty snow-avalanche deposits was that, after snowmelt, clast orientation was still random and slope-parallel reorientation of clasts was not observed.

Overland fluid flow Overland fluid flow or surface runoff occurs when water supply exceeds infiltration potential. Slopewash is the downslope transport of loose material, detached by rainsplash or by running water (Morgan 1995). As mentioned by French (2007), water supply by snowmelt is more important than rainfall in non-permafrost periglacial regions. Under permafrost conditions, overland flow may occur due to the presence of water from melting of the active layer and snow. Outside periglacial environments, rainfall is the most important source of water to generate overland flow. Overland flow takes the form of sheetflow or rillflow. Harris & Prick (2000) studied the role of summer rain in the development of modern stratified slope deposits near Kluane Lake, Yukon Territory, Canada. The deposits show an alternation of openwork, angular fine gravel layers and matrixsupported layers in which small, platy clasts are present within a sandy to silty matrix. The alternating units resemble gre`zes lite´es to some extent. The local climate is described as seasonally cold and dry. Frost action produces debris in sand to gravel size up to about 5 cm. This material moves downslope as dry grain flows. The coarsest material is transported to the bottom of the slope, while the finer sandy material is left behind. During summer rains, the

fine material on the upper slope may be saturated and develops into thin debris flows which cover the openwork clasts. The matrix-supported debrisflow deposit then dries and hardens, thereby stabilizing the openwork layer. The steep slope angles (35–458) and the sedimentary properties of the different layers provide diagnostic properties which preclude their interpretation as gre`zes lite´es. Traces of overland flow on a steep, active scree slope in the Pre´alpes, southern France, were observed by Van Steijn & He´tu (1997) on a partly vegetated scree slope developed below a cliff formed in massive limestone beds and marly limestones alternating with marls. On the talus surface, sloping at about 388, dry grain flows showed characteristic properties: coarse-grained openwork clasts in their terminal lobes, decreasing grain size in upslope direction as well as inverse grading and a fine-grained, silty to sandy slip plane at the most proximal part. On the slip planes, small earth pillars (up to 4 cm high at the downslope side and capped with a small clast) were separated by shallow rills. Soil aggregates were spread over the relatively impermeable slip plane and a sharply defined miniature debris flow (indicated as ‘aggregate flow’ by Van Steijn & He´tu (1997) because of the presence of soil aggregates) covered previous dry grain flow deposits. It was concluded that overland flow due to a high-intensity rainstorm triggered this small debris flow. Van Steijn & He´tu (1997) concluded further that the initial slopewash developed into a flow with higher density that was able to detach and transport coarser clasts. The flow developed into a miniature debris flow moving more or less according to the model of Takahashi (1978, 1980). The present-day climate of the region is Mediterranean, modified by the Alpine relief. Summers are rather dry but high-intensity rainstorms occurring in summer and autumn, when peak intensities of 100–130 mm h21 for 2–5 min are observed (Descroix 1989; Blijenberg 1998), are also characteristic. The association of dry grain flow deposits and those of miniature debris flows was frequently found in two sections of a gravel pit lower on this scree slope.

Aeolian and niveo-aeolian transport and deposition The contribution of wind action to the development of stratified slope deposits has not yet been studied in detail (Van Steijn et al. 2002), although He´tu (1990, 1991, 1995) examined ‘active’ stratified slope deposits in Gaspe´sie, Canada. In the Gaspe´sie, wind may transport sediment in two directions along cliff-talus systems – either downslope or upslope – depending on the wind incidence direction in relation to cliff morphology

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(He´tu 1991). The quantities of material displaced depend on lithology and wind conditions. Cliffs composed of weak rocks are most sensitive, especially when they disintegrate into granular particles or fine platy clasts with a large surface compared to their weight. He´tu (1991) describes in detail the transport performance of a strong blizzard that occurred on 5 and 6 January 1988. For almost 13 hours, wind velocity remained over 72 km h21 with gusts of nearly 100 km h21. Gravel chips (median diameter 15 mm) were transported mainly by saltation over the snow cover on the talus. Chips up to 30 mm wide and weighing nearly 5 g were seen to be moving. After their detachment by the wind from the cliffs above the talus, many of the chips were blown tens of metres before reaching the talus snow cover. He´tu (1991, 1995) also describes the contribution of niveo-aeolian sedimentation to a modern scree slope. The combination of moderate wind (,50 km h21), powder snow and abundant fine clasts (mainly sand, median values between 0.3 and 0.7 mm) causes cross-bedded niveo-aeolian sedimentation up to 1.5 m thick at protected sites at the concave base of the talus cones and nearby forest fronts. After snowmelt, the sand content is concentrated in one single bed without internal structure (He´tu 1991). During strong blizzards such as that of 5 and 6 January 1988, however, fine platy gravels are also mobilized; for example, He´tu & Vandelac (1989) report an average cover of c. 2 –5 cm thickness on the scree-slope base as a whole, and up to 20 cm in the depression between two adjacent taluses. The final result of both types of transportation and after snowmelt is a stratification that resembles gre`zes lite´es to some extent (He´tu & Gray 2000).

Relict stratified slope deposits Relict periglacial stratified slope deposits are reported from many mountainous areas in different regions (Bertran et al. 1995a). They include areas in France (Bertran 2005), Mediterranean Europe (Garcı´a-Ruiz et al. 2001; Texier & Meireles 2003), North America (Eaton et al. 2003; Smoot 2004) and Antarctica (Mori et al. 2007).

Gre`zes lite´es The relict deposits known as gre`zes lite´es in western France are presently understood as having a periglacial origin, notwithstanding the initial interpretation given by Guillien (1951) as permafrost-related. The more recent interpretation of these deposits, as proposed by Francou (1988, 1989), has been accepted by French researchers (Bertran et al. 1994, 1995a;

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Ozouf et al. 1995b). As mentioned by Ozouf et al. (1995b), some exposures in the Charente region reveal older gre`zes lite´es below those that were dated as Weichselian in age. A palaeosol, provisionally dated as Eemian in age, separates the two levels. It follows that the lower gre`zes might be of Saalian age. Typically, gre`zes lite´es occur on slopes facing NE to SE, suggesting a snow cover associated with prevailing westerly winds. Bertran et al. (1995a) argue that the Andean stone-banked solifluction sheet model provides the most satisfying genetic interpretation. This model involves the absence of permafrost and a high number of freeze –thaw cycles. Bertran et al. (1995a) think such conditions existed in southwest France at the time the gre`zes lite´es were formed; attenuated, oceanic conditions resulted in gre`zes lite´es deposition while more severe conditions (possibly permafrost) resulted in massive gre`zes. The presence of gre`zes lite´es in Charente and neighbouring regions in France depends also on carbonate lithology and gentle relief: other kinds of stratified slope deposits are found in nearby areas with other types of bedrock and steeper slopes. Bertran et al. (1995b) mention deposits mainly from debris flows, but also from dry grain flows and rock fall. This means that gre`zes lite´es (sensu stricto) are a relatively rare deposit. It also leads to the conclusion that a periglacial environment is essential for their development. More precisely, the climate would have been characterized by a high frequency of freeze –thaw cycles, a thin snow cover was often renewed and there was an absence of permafrost.

Deposits similar to gre`zes lite´es Relict stratified slope deposits that possess similar properties to those of gre`zes lite´es appear to be infrequent. However, five recent examples can be cited. (1)

(2)

In eastern France, deposits are reported from Champagne (Laurain et al. 1995) that consist of coarse sand and fine gravel (up to 1– 2 cm) within a silty matrix derived from local chalk bedrock. Openwork and matrixsupported layers alternate. The sediments are caused in part by aeolian activity, slopewash and solifluction. The age of the deposits is not given. Harmand et al. (1995) describe deposits resembling gre`zes lite´es from the Meuse valley about 20 km south of Verdun. Again, grain sizes are mainly in the range of fine gravels, 2–12.5 mm, whereas the amount of fines (,0.05 mm) is always less than 13%. The deposits are interpreted as having been caused by slopewash, active during cold

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(4)

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periods of the Pleistocene. Three formations are recognized: the lowermost probably dates from the Saalian and the top level is Weichselian in age. Pappalardo & Spagnolo (1999) describe a ‘peculiar stratified slope deposit’ in the southern Maritime Alps, Italy, heavily eroded, but interpreted along the Andean model of Francou (1988) and Bertran et al. (1995a). A regular alternation of matrix-rich layers (matrix about 50%) and clast-supported layers is mentioned, but the grain sizes are considerably larger than in gre`zes lite´es deposits. Moreover, erosional channels are mentioned, found at the boundaries between the layers of the deposit. The deposits are not dated. Gardner et al. (1991) describe deposits in the Valley and Ridge Province of Central Pennsylvania, USA. They consist of well-stratified beds of platy, subangular shale chips ranging in size from very fine to very coarse gravel. The beds show matrix-supported and clastsupported facies. The deposits are dated as late Wisconsinan (late Weichselian). Frost shattering apparently caused the clast properties while snowmelt-related slopewash deposited thin layers of openwork, matrix-free gravel layers. Laterally extensive beds of matrix-supported material are interpreted as having been deposited by thin, low-viscosity debris flow. In conclusion, only some of the processes involved are periglacial sensu stricto. Eaton et al. (2003) describe stratified slope deposits resembling gre`zes lite´es in the Blue Ridge Mountains, Virginia, USA. The deposits are between 27 400 and 15 800 a BP in age (Last Glacial Maximum). Although the descriptions given by the authors do not allow us to verify the slopewash interpretation, their age does not support an interpretation of gre`zes lite´es sensu stricto.

sedimentary processes involved are, at least in part, azonal. For example, Garcı´a-Ruiz et al. (2001) concluded that the stratified scree they observed in the central Spanish Pyrenees was formed not only by gelifluction but also rockfall and debris flows. Likewise, Pleistocene or Holocene snow-avalanche deposits (Jomelli 1997; Blikra & Nemec 1998; Blikra & Selvik 1998) clearly require periglacial conditions. Both Blikra & Nemec (1998) and Blikra & Selvik (1998) discuss the palaeoclimatic signal of snow-avalanche activity in high-relief mountains. For the Norwegian mountains it was concluded that the impact of regional climatic change ‘overpowered’ the effect of local slope conditions. They suggest that higher snow-avalanche activity reflects an increase in snowstorm frequency, whereas the debris-flow record reflects milder climatic conditions and an increase in high-intensity rainfall. An emerging problem is whether debris-flow activity depends only on a cold climate or is fully azonal. Theoretically, debris flows occur everywhere provided certain conditions are met (only one of which is climate). The latter need not be periglacial. It appears virtually impossible to obtain a climate signal from debris-flow deposits.

Deposits from azonal processes An extensive literature describes azonal processes involved in the development of talus. Three examples illustrate the role of these processes. (1)

Other periglacial stratified slope deposits In their investigation of Anatolian Quaternary colluvium, Nemec & Kazanci (1999) inferred different climatic depositional environments that deposited periglacial sediments consisting of (very) coarse and commonly pebbly sand. This formed up to 90% of the succession of cold-climate sediments. A periglacial sheetflow interpretation was based on the absence of rill erosion or gullying. This is an argument for prolonged and modest runoff fed by meltwater from late-lying snowpacks. While some relict stratified slope deposits were certainly formed in a periglacial environment, the

(2)

Quaternary colluvial deposits in the Taurus Mountains, Southwest Turkey, are almost exclusively interpreted as caused by azonal slope processes (Nemec & Kazanci 1999). The oldest bedded, red-brown colluvium, dated about 1.4 Ma, consists of alternating deposits from rockfall, debris flows, other grain flows and some sheet wash. These deposits developed under a warm-humid to warm-dry climate. In overlying grey-coloured accumulations, deposition from concentrated runoff appeared to be more important. At a stratigraphically still higher level, snowavalanche, slushflow, and sheetflow deposits occur, alternating with rockfall and debrisflow deposits. The uppermost, youngest depositional complex again consists of grainflows, rock fall and debris-flow deposits. The conclusion is that the sedimentological evidence of snow-avalanche and slushflow deposits allows recognition of a short periglacial interval between much longer periods of Mediterranean or even warmer climatic conditions. A more restricted study was carried out in the French Pre´alpes by Van Steijn & He´tu (1997).

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(3)

An exposure showed a repetition of three different elements: (i) heterogeneous material, containing abundant angular gravels within a silty matrix, interpreted as debris flows; (ii) openwork gravel layers, showing coarseningupwards grading but sometimes coarsening downwards, interpreted as dry grain flow deposits; and (iii) complex units, 6–10 cm thick, consisting of a basal gravel layer, a middle layer characterized by silty to sandy aggregates and a top layer of very fine gravel. These complex units have the characteristics of miniature debris-flow deposits. The conclusion was that all processes inferred from the section were azonal and that coldclimate conditions are not needed to explain the succession of processes. Pe´rez (1998) reports on morphology and sedimentology of talus aprons formed following volcanic debris avalanches in Lassen Volcanic National Park, California, USA. Processes inferred are both periglacial and azonal: weathering of the (structurally weakened) bedrock is mainly by frost shattering, whereas transport processes are dominated by rock fall. Snow avalanching also occurs, as well as minor debris-flow activity.

In other cases, recent research caused a revised interpretation of stratified slope deposits that were previously seen as periglacial. An example is provided by Texier & Meireles (2003) who discussed relict mountain slope deposits of northern Portugal. Three main processes are reported as responsible for the emplacement of regional relict slope deposits: runoff, debris flows and dry grain flows. The conclusion from the cases described above is that many complexes of relict stratified slope deposits are known in which non-periglacial processes prevailed.

Conclusions Relatively few stratified slope deposits appear to have a fully periglacial origin. A general consensus has emerged about the periglacial origin of the Pleistocene gre`zes lite´es following acceptance of the stone-banked solifluction sheet model. The applicability of this model to the gre`zes lite´es in France leaves some points unresolved. For instance, the depth of freezing associated with the Andean stone-banked solifluction model appears to be always less than 15 cm whereas the depth of freezing associated with the Charente gre`zes lite´es probably reached at least 40– 50 cm. Furthermore, the duration of snow cover appears to have been much longer in the case of the Charente than in that of the Andean

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study sites. However, absence of permafrost is a common element for both regions. The investigation of relict snow-avalanche deposits establishes the periglacial background of these sediments. Other deposits are known from present-day cool temperate climates where frost or snow plays an important role. Frost-coated clast flows and niveo-aeolian deposits belong to this group. Azonal processes such as rock falls, debris flows or (dry) grain flows may also be active under such conditions, and this may complicate the climatic interpretation of relict deposits. There are several kinds of stratified slope deposits for which frost is no prerequisite. For example, most debris flows and other grain flows, aeolian transport and runoff due to rainfall belong to this category. Although these processes may be characterized as azonal, periglacial conditions nevertheless influence their intensity. Questions that require future research include the following. (1) What is the influence of frost, rain and wind upon clast production in different lithological and structural situations? (2) What are the cold-climate (periglacial) factors involved in the magnitude–frequency relationships of azonal processes that contribute to the development of stratified slope deposits? (3) What effect has Holocene climate change had upon the development of talus slopes in different regions? (4) What is the influence of paraglacial conditions on the initial phases of talus development? (5) Can one apply knowledge about periglacial stratified slope deposits to planetary process studies? Photographs shown in Figures 1– 3 were made available by Pascal Bertran.

References Allen, J. R. L. 1972. Intensity of deposition from avalanches and the loose packing of avalanche deposits. Sedimentology, 18, 105–111. Anderson, E. & Harrison, S. 2007. Late Quaternary paraglacial sedimentation in the Macgillycuddy’s Reeks, southwest Ireland. Irish Geography, 39, 69–77. Ballantyne, C. K. 2002. Paraglacial geomorphology. Quaternary Science Reviews, 21, 1935– 2017. Benedict, J. B. 1976. Frost creep and gelifluction features: a review. Quaternary Research, 6, 55– 76. Bertran, P. 2005. Stratigraphie du site des Peyrugues (Lot), une coupe de re´fe´rence pour le dernier Ple´niglaciaire en Aquitaine. Quaternaire, 16, 25–44. Bertran, P. & Texier, J. P. 1994. Structures se´dimentaires dans un coˆne de flot de de´bris (Vars, Alpes

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Holocene slope dynamics in Sierra Nevada (south Spain). Sedimentological analysis of solifluction landforms and lake deposits ´ MEZ ORTIZ2,3 M. OLIVA1,3* & A. GO 1

Institute of Geography and Territorial Planning, University of Lisbon, Alameda da Universidade, 1600-214 Lisbon, Portugal 2

Department for Physical and Regional Geography, University of Barcelona, Montalegre 6-8, 08001 Barcelona, Spain

3

Landscape Research and Mediterranean Mountain Palaeoenvironments Research Group, University of Barcelona *Corresponding author (e-mail: [email protected]) Abstract: The activity of geomorphic processes in the current periglacial belt of Sierra Nevada during the Holocene has been strongly controlled by climate variability. We examined two sedimentary records showing several cycles with varying intensity of slope processes in the massif. Chronostratigraphic studies of solifluction landforms have revealed seven cycles of enhanced slope activity in Sierra Nevada during the last 7 ka. Lake sediments show eight periods with increased slope activity over the last 6 ka. Depending on temperature and moisture regime, erosion or conditions of slope stability prevailed. Cold and wet periods triggered slope processes and favoured deposition of coarse-grained sediments into the lakes, whereas warm phases were prone to soil development and fine-grained sedimentation in the lakes. The lake sediments also show evidence for an increasing aridity trend since the Holocene Optimum Warm Period in Sierra Nevada, which induced a shift of solifluction processes to higher elevations accompanying late-lying snow patches.

Periglacial research has contributed to a better understanding of Holocene climate variability with the study of sedimentary deposits both in polar regions and in high mountain environments. Mountain areas have been defined as suitable environments to study climate variability due to the strong gradients of climate parameters that increase the effectiveness of geomorphic dynamics (Wanner et al. 2000). Our research is focused on the periglacial belt of Sierra Nevada (Spain), where we have reconstructed the mid –late Holocene activity of geomorphic processes from two natural archives: solifluction landforms and mountain lakes. A complete overview of present-day solifluction processes in different areas from polar regions and alpine environments was carried out by Matsuoka (2001). Several papers have focused on the development of solifluction landforms, trying to uncover why solifluction is active in some periods and inactive in others (Benedict 1966; Gamper 1983; Veit 1988; Elliot & Worsley 1999; Matthews et al. 2005; Kinnard & Lewkowicz 2006). However, the complexity of solifluction mechanisms and the wide range of environments where solifluction landforms are present make it difficult to define the key climate factors controlling their movement.

These include seasonality, intensity and type of precipitation, snow cover, water availability, seasonal frost and permafrost. In former papers we have already analysed the morphometry and distribution of more than 200 solifluction landforms existing in Sierra Nevada (Oliva et al. 2008, 2010a) and studied their present and past dynamics (Oliva et al. 2008, 2009). Solifluction processes in this massif are conditioned by a complex interaction between environmental factors (slope, vegetation cover, texture) and other parameters (ground thermal regime, length and thickness of the snow cover, water supply). Feedback mechanisms among these variables make it difficult to understand the key factors involved in present and past solifluction dynamics. Monitoring of lobes in different sites indicates the most favourable environmental conditions for solifluction displacements, which are highly dependent on the soil water content derived from snow melting. Lobes situated near water channels and those influenced by water supply from late-lying snow patches show higher displacements of c. 0.5 cm a21 (Oliva et al. 2009). In mountain areas lake sediments have been used to infer the activity of cold geomorphological processes during the Holocene such as glacial erosion

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 227–239. DOI: 10.1144/SP354.15 0305-8719/11/$15.00 # The Geological Society of London 2011.

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and deposition processes, landslides and debris flows (Schmidt et al. 2002; Irmler et al. 2006; Daigle & Kaufman 2009). Zielinski (1989) and Paasche et al. (2007) reconstructed periods of rock glacier activity in the Rocky Mountains and Scandinavia, respectively, by studying the properties of the sediments of the lakes located downvalley. The structure of this paper consists of three sections: solifluction records, lake sediments and the coupling of the geomorphic activity inferred from both records. The objectives of our research are: to determine the Holocene chronology for slope processes in Sierra Nevada deduced from solifluction landforms; to infer active slope phases from lake sediment records; to examine if terrestrial and aquatic records report a synchronous timing for geomorphic activity; and to assess how Holocene climate variability has conditioned slope processes in the area.

Study area Sierra Nevada is the range with the highest peaks in the Iberian Peninsula (Mulhace´n, 3478 m a.s.l.; Veleta Mt, 3398 m). Our study area is located above 2500 m in the westernmost part of Sierra Nevada, where the landscape is strongly shaped by the Pleistocene glaciations and post-Last Glacial Maximum (LGM) periglacial processes. During the Little Ice Age (LIA), Sierra Nevada hosted the southernmost glaciers in Europe in the Veleta cirque. They finally melted during the second half of the 20th century. Nowadays, island permafrost is present in the highest northern cirques above

3000 m with active rock glacier dynamics closely related to remnants of dead ice (Go´mez Ortiz et al. 1999). Seasonal frost penetration in the ground is highly variable and strongly controlled by snow thickness (Oliva 2009). Lithology is composed basically by schists, with a homogeneous composition in the highest catchments. At 2500 m the mean annual temperature is 4.4 8C and precipitation totals 700 mm, concentrated from October to April. Sierra Nevada lies in a border position between different influences (Fig. 1): geographical (Europe/Africa); maritime (Atlantic Ocean/Mediterranean Sea); and climatic (mid-latitude westerlies/subtropical high pressure belt). Vegetation cover is very scarce above the timberline (c. 3% over 2200–2300 m), mostly restricted to the valley floors and the shorelines of the lakes (c. 2900– 3000 m). Vegetation growth in the present-day periglacial environment is strongly conditioned by the lack of water availability during the summer. Our research area comprises the San Juan and Rio Seco valleys, where hundreds of solifluction landforms and several mountain lakes are located (Figs 1 & 2a). San Juan is an open U-shaped glacial valley at the northern flank of Sierra Nevada with solifluction phenomena concentrated at three elevation bands: 2470–2550 m (San Juan valley, sector C; SJC), 2790–2840 m (SJB) and 2840–2910 m (SJA). Rio Seco is a large cirque at the southern side of Sierra Nevada with solifluction lobes distributed in two sectors: 2930–2950 m (Rio Seco cirque, sector B; RSB) and 2990–3010 m (RSA). Sierra Nevada hosts the highest lakes in Western Europe, most of them distributed over 3000 m. In

Fig. 1. Location of Sierra Nevada in the southern part of the Iberian Peninsula. (a) Satellite image showing the study areas in the massif and (b) geographical context of Sierra Nevada range.

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Fig. 2. Two different landforms studied in this research: (a) solifluction lobes (6 m long and 5 m wide) and (b) Rio Seco mountain lake (110 m long and 45 m wide).

this paper we will mainly focus on the Rio Seco lake (Fig. 2b), located upvalley from the solifluction landforms developed in RSA and RSB.

Materials and methods Thirty solifluction landforms from the San Juan valley and the Rio Seco cirque have been analysed and sampled for chronostratigraphic purposes. Seven cores were collected from three lakes: Aguas Verdes, Rio Seco and Rio Seco lagoon. A total of 21 samples from solifluction landforms and lake sediments were processed for Accelerator Mass Spectrometer (AMS)-dating in order to identify the chronological framework of the environmental evolution of the area. Datings were carried out at the Angstrom Laboratory of Uppsala (Sweden). Radiocarbon ages have been converted to calendar ages with the CALIB 5.0.2 calibration program (Reimer et al. 2004).

Solifluction landforms Sampled units were selected according to the internal sedimentary structure and to encompass the different typologies of solifluction landforms described in these two valleys (Oliva et al. 2008), although most of the landforms analysed in this paper are solifluction lobes. Pits were excavated until a depth of c. 0.7 –1 m where large boulders appeared. A detailed study of the stratigraphy of each landform was performed in the field according to the FAO (2006) procedures. We collected c. 0.5 kg from each lithostratigraphic unit for laboratory analyses. Texture measurements and organic matter content determination were performed on dry, sieved samples (,2 mm). Grain size was quantified with a Micromeritics SediGraph-5100

system for clay (,2 mm) and silt (2 –63 mm), whereas sands (.63 mm) and gravels (.2 mm) were determined by wet sieving. Organic carbon (OC) was measured with an Elemental Analyser vario Macro. Because of the large proportion of modern roots in most of the samples and to avoid possible modern contamination by recent plant fragments, we did not date bulk sediments. Pollen grains from organic layers of solifluction lobes were concentrated for AMS-dating. We present 15 datings that provide a tentative chronology for slope processes and soil formation periods in the current periglacial belt of Sierra Nevada during the last millennia, although more datings should provide a better accuracy. We have assumed constant sedimentation rates in these erosion-dominated environments in order to provide a continuous chronology. Therefore, dating results must be considered as approximate estimations. Matthews et al. (1993) stated that an accuracy of better than +200 14C years for solifluction lobes is unlikely and that errors of +500 14C years are usual, with errors up to +1000 14C years possible.

Lake sediments Thanks to the shallowness of the lakes (c. 3 m), we collected cores from the bottom of the lakes by inserting 5 m long PVC tubes (40 mm inner diameter) from an inflatable boat. Cores were stored and preserved at 5 8C until they were subsampled at intervals of 1 cm. Organic carbon (OC) and C/N ratios were determined with an Elemental Analyser on dry, sieved samples (,2 mm). For grain size measurements, samples were pre-treated with H2O2 to remove the organic matter and quantified individually with a Malvern laser grain size analyser.

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Present solifluction landforms and former activity The fact that present-day monitoring of solifluction processes has revealed a very weak activity pattern indicates that the existing solifluction landforms may have developed during other more favourable climate conditions.

Examples of sections of solifluction landforms Sedimentological profiles of solifluction landforms show an alternation of soil horizons (E1 –E8) and mineral-rich units (S1 –S7), suggesting important environmental changes in the headwaters of the highest valleys of Sierra Nevada during the Holocene (Fig. 3). Figure 3 shows two examples of the stratigraphy of solifluction lobes that can be taken as representative profiles of the landforms produced during the mid– late Holocene (Oliva 2009). Lobe SJB.3 is composed of an abundant weathered gravel layer overlain by organic-rich beds alternating with thin lenses of coarse-grained horizons (Fig. 3a). The basal layer is a mineral-rich unit (S6) with a high percentage of gravels (.40%) and a sandy matrix (50–60%). A more organic layer, characterized by an increase in OC content (5%) and a finer matrix (E6), lies over the basal layer. The overlaying units (E5 –E3) are organic-rich fine-grained sediments (OC .20%) that alternate with thin lenses with less organic matter content (OC c. 10%) and coarsegrained slope deposits (S5 –S3). In the upper part of this section there are two layers (S2 and S1) with more gravel content (20–30%) and a significant decrease in the OC values (6 –7%). Between these periglacial deposits, a light increase of the OC content is recorded (13%) with a clear diminution of gravel deposition and a silty matrix (E2). The topsoil contains high percentages of organic matter (OC .20%) and a silty texture (E1). We performed three datings in lobe RSB.1 that indicate that this large solifluction landform developed during the last millennium (Fig. 3b). The basal organic layer dated 1120 + 120 14C a BP (Esteban 1995) has high OC values (20%) with a silty matrix (E3). Soil development was interrupted by a period of enhanced slope instability characterized by an increase in gravels and sands with a clear diminution of the OC content (S3). Subsequently, another phase of soil formation took place with very high OC values (.30%) and an increase of the fine fraction (E2). This level has been dated 619 + 64 14C a BP. Immediately after, a phase with gravels deposition (15%) and a decrease of the organic matter content is recorded (S2). The present topsoil is constituted by high OC percentages (20%) and a silty matrix (E1).

Interpretation of the sections What we interpret as solifluction deposits are characterized by low OC values, with a sandy matrix and high gravel content. Solifluction deposits are generally poorly sorted, with a thickness varying from 10 to 40 cm and a clear downslope clast orientation parallel to the slope. These coarse stony deposits are composed by fine angular rock fragments that result from the mechanical weathering of the schist bedrock. The proportion of gravels in the mineral units ranges from 10 to 50% depending on the proximity of the lobes with respect to the talus-covered slopes. In some lobes, such as in SJB.3, we have also detected several depositional events of terrigenous clastic sediments interstratified with more organicrich layers within the lobes (S5, S4 and S3). These events are characterized by significant increases in sand content and a low proportion of gravels and could be related to sediment gravity flows, such as grain flows or snowflows (Blikra & Nemec 1998; Bertran & Jomelli 2000). By contrast, soil horizons are characterized by significant proportions of organic carbon (OC up to 25– 30%), a silty matrix and low gravel content. The basal layer of lobe SJB.3 has many coarse pebbles and fine cobbles with the long axes imbricated upslope and a significant preferred orientation parallel to the slope direction that gives evidence that the downslope movement was due to solifluction (Nelson 1985; Bertran et al. 1997). The date near the base of the pit suggests that the formation of the SJB.3 lobe started before 3309 + 92 14C a BP, when a thin regosol developed. After a short period of slope activity a thick histosol (33 cm) formed. The uppermost level of this organic unit (13 cm depth) may have developed during the Mediaeval Warm Period (MWP) since it was dated 1020 + 126 14C a BP. Two sandy mass-flow deposits interfinger with this organic-rich layer, although we are not able to establish accurately the chronology of these events due to the lack of more datings in this profile. The uppermost dating suggests that the last two periods with coarser material deposition by solifluction processes took place during the last millennium, probably during the LIA, which could correspond to the glacier advances reconstructed by Schulte (2002) from 210 Pb dating of glaciolacustrine sediments in the nearby Veleta cirque. Although lobe RSB.1 shows a higher dynamism of geomorphic processes with respect to SJB.3 due to more water availability in this area, a synchronous timing is observed for mineral unit formation and soil development phases in both lobes over the last millennium. During the MWP, warm-climate conditions were favourable for histosol formation in both landforms.

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Fig. 3. Surface view, pit face and stratigraphic profile of two solifluction lobes from (a) San Juan valley and (b) Rio Seco cirque.

Slope deposits indicate that the deterioration of these warm conditions triggered two active periglacial phases that may correspond to the LIA cold pulses. In between these phases, a period of relative

soil stability occurred with regosol formation in SJB.3 and histosol development in RSB.1. In this lobe, the higher water supply favoured the formation of an organic-richer layer. Finally, the rising

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temperature trend since the mid-19th century has promoted the development of histosols in both solifluction lobes and has spread the grass cover over the valley floors.

Chronological framework of slope processes in Sierra Nevada inferred from solifluction lobes Chronostratigraphic profiles of other solifluction lobes also reveal seven different phases of

geomorphic activity in the highest western cirques of Sierra Nevada (Oliva 2009). Sedimentological records from solifluction landforms cover the last 7 ka BP in San Juan valley and the last 2 ka BP in Rio Seco. Figure 4 resumes all the datings performed on samples from solifluction lobes and synthesizes the chronologies for active slope phases and soil development periods in both valleys. In the valley floor of the lowest area in San Juan (SJC), some solifluction lobes record geomorphic activity during the early– mid Holocene, although most of the landforms of the highest part in this

Fig. 4. Chronology of slope processes and soil development during the Holocene in the present-day periglacial belt in Sierra Nevada (modified from Oliva 2009).

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valley record active slope processes during the mid –late Holocene. Nevertheless, chronology for slope dynamics in San Juan is only tentative from 7 to 4 ka BP, approximate between 4 to 2 ka BP and reliable for the last 2 ka BP (Oliva 2009). The aggradation of the lobe SJC.54 started during a solifluction phase previous to the buried soil (E8) dated 7098 + 60 a BP (Fig. 4). Two AMSdatings in lobe SJC.39 give evidence of different solifluction phases during the mid –late Holocene. At c. 2500 m in San Juan, according to lobes SJC.54 and SJC.39, soil formation (E7) was dominant from 7 to 5 ka BP (including the Holocene Optimum Warm Period, HWP) but the degradation of these warm conditions from 5 to 4 ka BP promoted slope instability (S7). Between 4 and 3.6 ka BP, soil development prevailed in San Juan (E6) prior to a new period of active slope processes from 3.6 to 3.4 ka BP (S6), which is reflected in a thick gravel deposit in unit SJB.3 (Fig. 3). In this lobe a regosol developed between 3.4 and 3.2 ka BP (E5), but immediately after another geomorphic phase took place with sands deposition (S5). The last three cycles of slope instability and soil development have a similar timing in both San Juan and Rio Seco valleys. During the Roman Warm Period (RWP) and the MWP, several lobes show evidence of soil formation (E5 to E3) with thick organic layers in both valleys and peat growth in areas with larger water supply. These were only interrupted by brief periods of dominance of depositional processes (S4 and S3). The LIA determined two phases of active solifluction in Sierra Nevada that occurred between the 7th and 8th centuries (S2) and between the 17th and the 19th centuries (S1), with regosol development during the 14th and 16th centuries (E2). Finally, the post-LIA warming trend initiated since the last pulses of the LIA has extended vegetation cover and induced soil formation in the highest parts of the massif (E1).

Geomorphic activity inferred from lake sediments: results and interpretation High mountain lakes in Sierra Nevada have not been glaciated since the LGM in the southern slope and probably since the Younger Dryas cold period in the northern cirques (Oliva 2009). Three lakes were sampled in the southern slope of the massif: Aguas Verdes, Rio Seco and Rio Seco lagoon. In this paper we discuss results from Rio Seco lake, which is a representative example of lake sediments in Sierra Nevada (Oliva et al. 2010b). Rio Seco lake is one of the highest lakes ever cored in southern Europe (3023 m a.s.l.). It is a small glacial lake with a mean depth of 2.5 m and

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a catchment of 12.2 ha. Lakes in Sierra Nevada are very oligotrophic showing very low biological production rates (Pulido-Villena et al. 2006). We extracted two cores from this lake and three datings were performed on the longest one (69 cm). It is a fairly homogeneous mud core with many decomposed roots and plant fragments (Sphagnum) and little changes in colour properties. However, sedimentological studies of Rio Seco lake sediments reveal several cycles of varying texture and organic matter contents (Fig. 5). Instability periods are characterized by a higher proportion of coarse sediments (sands) and low OC and C/N ratios, whereas more stable periods are represented by increases in the organic matter content and a lower percentage of sands. The fact that OC and C/N ratios define parallel curves indicates that the organic matter present in the sediments of these oligotrophic lakes has a terrestrial origin (Meyers & Teranes 2001). According to the chronology established from solifluction landforms (Fig. 4), soil formation was dominant during the HWP in the summits of Sierra Nevada (Fig. 4), so peat probably developed around the lakes during this period. Since the midHolocene, wave action may have eroded the margins of the lake and has deposited vegetal remains of this peat layer in the bottom of the lake. Precisely, we have dated fragments of Sphagnum. These redeposition processes may explain the inversion of our chronology and the reduced temporal range of the three datings (5.3–4.8 ka). Furthermore, we inferred a continuous sedimentation rate taking into account the youngest dating (4887+87 a BP), which is the dating that indicates the minimum age for peat formation. Thus, environmental evolution inferred from Rio Seco lake sediments provides a relative chronology for landscape changes but not a precise timing for the onset/offset of slope processes in Sierra Nevada. Taking into account the dating existing at 55 cm depth (4887 + 87 14C a BP), the vegetation cover in the lake catchment must have been significantly higher during the HWP as suggested by the high OC and C/N values. However, phases with enhanced erosion and deposition processes were also recorded before 4.2 ka BP (F8 –F6) with diminutions of the organic proportion in the sediments and more sand deposition. Since this moment, a clear decrease of the OC and the C/N ratio indicates that environmental conditions changed to be more difficult for vegetation growth and grass cover in the catchments gradually tended to be sparser. In accordance with other regional proxies that report more arid conditions over southern Iberian Peninsula and northern Africa since the HWP (Demenocal et al. 2000; Gasse 2000; Pantaleo´n-Cano et al. 2003), this trend started around 4.2 ka BP in Sierra

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Fig. 5. Lithostratigraphical properties of Rio Seco lake sediments.

Nevada. Despite the fact that water availability has progressively reduced during the last millennia in this semiarid range some periods have still been propitious for geomorphic activity, as reflected by increases in sand deposition and diminutions of the organic matter content (F5 –F1).

Discussion Solifluction landforms and lacustrine sedimentary records show evidence of significant environmental changes during the Holocene related to climate variability. Depending on temperature and moisture changes, geomorphic processes were more or less intense. We discuss separately the results from

solifluction landforms and mountain lakes and couple their geomorphological response to climate variability.

Holocene slope processes and palaeoenvironmental implications deduced from solifluction landforms Not only is slope instability affected by changing temperatures and precipitations, but it also depends on the seasonality of these changes. According to our findings, late spring is the most favourable period of the year for solifluction in Sierra Nevada today (Oliva et al. 2009). Water availability is the crucial factor controlling

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solifluction and other depositional slope processes in Sierra Nevada. Colder and/or wetter phases imply a thicker snow mantle in winter and postpone the melting period. The persistence of a widespread snow cover in late spring and early summer in the summits of the massif prolongs the seasonal frozen layer by several weeks. The abundance of late-lying snow patches is decisive in lengthening the runoff water period and in maintaining solifluction processes (Morin & Payette 1988; Matthews et al. 2005). This delay would also shorten the vegetation growing season, reduce grass cover at this altitude and enhance slope processes during late spring and early summer (Oliva et al. 2009). Although presently solifluction landforms are mostly covered by vegetation, which is also a physical impediment for erosion and mass wasting, during more active geomorphic periods they must have had a sparser vegetation cover on their surface that was more favourable for erosion processes. After cold periods prone to slope processes, grass cover became denser and mitigated water erosion and the formation of the mineral horizons as temperatures raised. Vegetation cover of these landforms in Sierra Nevada is also affected by other feedback mechanisms: fewer late-lying snow patches; thinner and shorter ground frost; lower water supply during the melting season; and lower albedo effect for snow-free surfaces that contribute to slow down geomorphic processes in the massif. The oldest deposits analysed are situated in San Juan at 2500 m (SJC) and correspond to the early –mid Holocene. At altitudes ranging from 2600 to 2900 m (SJB and SJA) several geomorphic phases took place during cold and wet periods within the mid –late Holocene. San Juan and Rio Seco valleys show a similar chronology for slope instability and soil formation throughout the late Holocene. Active slope processes in Sierra Nevada with enhanced erosion were recorded during the Neoglacial period (S6: 3.6– 3.4 ka BP) and the LIA (S2 and S1: 850 –700 and 400 –150 a BP). These phases have a similar timing as other analogous studies carried out in the Alps (Steinmann 1978; Gamper 1983; Veit 1988) which reported active solifluction during colder periods and soil formation prevailing during warm phases within the last millennia (Oliva et al. 2009). Other periods with unstable slopes in Sierra Nevada related to other sedimentary processes (grain flows or snowflows) took place between 5–4 (S7), 3–2.8 (S5), 2.5–2.3 (S4) and 1.8 –1.6 ka BP (S3). By contrast, edaphic processes evolve preferably during warm periods. In areas with high water availability soils are well-developed, while in areas where water supply is scarce soils are weakly developed. Dry periods mostly favour regosol formation whereas wet phases promote well-structured

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histosols. Regosols only developed during 3.4– 3.2 ka BP (E6) and 700–400 a BP (E2). By contrast, several periods during the Holocene were suitable for histosol formation: 7–5 (E8), 4–3.6 (E7), 2.8– 2.5 (E5), 2.3–1.8 ka BP (E4), 1600–850 (E2) and 150 a BP onwards (E1). Concerning the LIA (the last phase of active solifluction in Sierra Nevada from which more information is available), the existence of thick solifluction deposits corresponding to this period at altitudes below 2500 m (when solifluction at present-day is very weak at 2900–3000 m) suggests that colder and wetter conditions prevailed during that period (Oliva 2009). In fact, documentary sources from the 17th to 19th centuries confirmed the existence of small glaciers in the northern cirques and more snow patches remaining during the summer season (Go´mez Ortiz et al. 2009). Several researchers have reconstructed the temperature increase in the Iberian mountains since the coldest pulses of the LIA in a range between 0.8 to 1.7 8C (Martı´nez-Cortizas et al. 1999; Lo´pez Moreno 2000; Gonza´lez Trueba et al. 2008). Considering a mean temperature gradient of 0.65 8C/100 m, temperature by itself does not explain the altitudinal difference of c. 500 m between geomorphological processes dominating during the LIA and today. Precipitation must therefore have played a determining role in activating solifluction at altitudes of 2500 m during phase S1 (Oliva 2008). The historical reconstruction of precipitation in Andalusia since 1500 by Rodrigo et al. (1999) indicates that the wettest and coldest phase of the LIA corresponds to the period between 1590 and 1650, possibly related to more recurring and persistent negative phases of the NAO-WeMO (North Atlantic Oscillation-Western Mediterranean Oscillation) teleconnection patterns (Oliva et al. 2006) which may induce propitious climate conditions for solifluction activity in Sierra Nevada.

Slope processes and palaeoclimate conditions inferred from lake sediments The tentative chronology derived from lake sediments in Sierra Nevada suggests eight active slope phases during the last 6 ka BP with an approximate timing of: 6.2?– 6 (F8), 5.8– 5.6 (F7), 5.3–4.6 (F6), 3.7 –3.1 (F5), 2.5–2.2 (F4), 1.8 –1.6 (F3), 1.2– 0.9 ka BP (F2) and 650–200 (F1) a BP (Oliva et al. 2010b). Geochemical and texture properties of the sediments show an alternation between periods with more stable slopes and a dense vegetation cover developing around the lakes and others with enhanced geomorphic activity and less vegetation cover in the catchments.

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Fig. 6. Correlation between active slope phases derived from solifluction landforms and lake sediments.

Moist climate conditions both in winter and spring combined with lower temperatures in late spring and early summer must have been especially effective in triggering slope processes in Sierra Nevada. A thicker snow mantle implied that further snow patches remained during the melting season, which determine a longer and larger water supply that transported and deposited more material into the lakes. By contrast, periods with warm and dry conditions were prone to slope stability in the present-day periglacial environment, with less material deposited into the lakes and a light increase of the vegetation cover in the catchments. High temperatures (especially in summer) combined with wetter conditions would reinforce and expand a denser grass recover and soil formation surrounding the lakes, which would also have increased their bioproductivity. On the other hand, the clear decrease in OC and C/N in the Rio Seco lake sediments since the midHolocene reflects a diminution of the aquatic and terrestrial biological production in the catchment. The aridity trend initiated after the relatively moist HWP in the southern Iberian Peninsula (Carrio´n et al. 2007) hinders vegetation growth in the headwaters of the highest western cirques. The lower water runoff during the snow-free season tends to decrease sediment transfer and slope activity in southern valleys of Sierra Nevada and only permits a very scarce vegetation cover.

Coupling solifluction records and lake sediments data The chronology of slope instability inferred from solifluction landforms provides an estimated timing

between 7 and 4 ka BP, covers roughly the last 4 ka and has a fairly good resolution for the last 2 ka (Oliva 2009). On the other hand, lake sediments provide a relative chronology for active slope phases in the summits of the massif for the last c. 6 ka (Oliva et al. 2010b). The overlapping of both sedimentary archives allows us to deduce if solifluction landforms and lake sediments report geomorphic activity during the same periods. Figure 6 highlights a clear synchrony between both records: periods with enhanced slope activity recorded in solifluction lobes correlate with phases of higher coarse-grained input into the lakes. Colder and/or wetter conditions must have triggered intense slope instability in Sierra Nevada. Higher precipitations in winter/early spring combined with lower temperatures in late spring/ summer postpone the melting season, determine more late-lying snow patches that provide a larger runoff in late spring and early summer and shorten the vegetation growing time, reducing grass cover and enhancing potential mass movements. By contrast, warm periods in general favour soil formation and vegetation cover (more or less dense according to moisture availability) and impede slope processes and sediment migration downvalley.

Spatial distribution of solifluction landforms during the Holocene Nowadays, climate conditions are not suitable to produce active solifluction processes in Sierra Nevada. According to the oldest buried soil layers dated in each solifluction landform the oldest solifluction records are distributed in the lowest areas of San Juan and correspond to the early –mid

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Fig. 7. (a) Satellite image of the study areas in San Juan and datings of the oldest soil layers developed over solifluction deposits. (b) The distribution of these landforms reproduces the migration of solifluction upvalley in response to more arid conditions.

Holocene, whereas in those landforms situated at higher elevations the oldest deposits correspond to the mid –late Holocene. The same distribution pattern of solifluction landforms in relation to altitude may reflect the aridity trend inferred from lake sediments dominating in Sierra Nevada since the mid-Holocene (Oliva et al. 2010b): solifluction remained active progressively at increasing elevations where there were more late-lying snow patches (Fig. 7). In these areas the melting season was longer and water availability was higher, which is crucial to maintain solifluction processes in the semiarid environment of Sierra Nevada (Oliva et al. 2008). The inactivity of solifluction processes in the lowest areas has preserved the oldest slope deposits while those located at higher altitudes could have been eroded. However, the shift of solifluction to higher elevations is interrupted during the LIA and solifluction reactivated at elevations below 2500 m.

Conclusions We studied two different types of sedimentary records of the periglacial belt in Sierra Nevada that have revealed important landscape changes during the last millennia. Solifluction landforms and lake sediments show evidence of phases with varying intensity of geomorphic processes, with more or less slope activity according to moisture and temperature conditions. The observed geomorphological pattern in relation to climate variability suggests that colder and/or wetter periods are favourable for slope processes, whereas warm phases induce soil formation and promote the extension of grass cover in the valley floors above 2500 m in Sierra Nevada. Analogous studies performed in the periglacial belt of the Swiss and Austrian Alps report a similar pattern, with more stable slopes during warm phases and enhanced periglacial activity during cold periods.

The seasonality of the climate variability is certainly critical in understanding the geomorphological processes taking place in the summits of Sierra Nevada. Periglacial slope processes (e.g. solifluction) are more intense during periods with increased precipitation in winter and early spring that imply more snow and a longer water runoff in late spring and early summer. Otherwise, edaphic processes are more effective when significant precipitations (either in winter or summer) are combined with warm temperatures in late spring and summer. According to the lake records, geomorphic processes had to adapt to more arid conditions over the southern Iberian Peninsula from the midHolocene onwards (Carrio´n et al. 2007), probably conditioned by the reinforcement of the northern fringe of the high pressure subtropical belt over the western Mediterranean basin and northern Africa (Oliva 2009). As water availability was lower in the massif, solifluction became progressively inactive at lower altitudes (c. 2500 m) and remained only active at higher elevations (2600 to 2900 m), accompanying late-lying snow patches. This research was funded by the research project of the Spanish Government ‘Permafrost and fossil ice degradation and climate change in Sierra Nevada’ (018-2007) and by a postdoctoral grant of the Fundac¸a˜o para a Cieˆncia ea Tecnologia of Portugal. We would like to thank our colleagues J. Nofre (New University of Lisbon, Portugal) and L. Frago (University of Barcelona, Spain) for their help during the field work campaign. Laboratory facilities were kindly provided by H. Veit (University of Bern, Switzerland) and F. Burjachs (University Rovira i Virgili, Tarragona, Spain).

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HOLOCENE SLOPE ACTIVITY IN SIERRA NEVADA SPAIN Distribucio´n y morfometrı´a de los lo´bulos de solifluxio´n en los altos valles nevadenses. Nimbus, 23– 24, 133–148. Oliva, M., Go´mez Ortiz, A. & Schulte, L. 2010b. Tendencia a la aridez en Sierra Nevada desde el Holoceno Medio inferida a partir de sedimentos lacustres. Boletı´n de la Asociacio´n de Geo´grafos Espan˜oles, 52, 27– 42. Paasche, Ø., Dahl, S. O., Løvlie, R., Bakke, J. N. & Nesje, A. 2007. Rockglacier activity during the Last Glacial-Interglacial transition and Holocene spring snowmelting. Quaternary Science Reviews, 26, 793–807. Pantaleo´n-Cano, J., Yll, E. I., Pe´rez-Obiol, R. & Roure, J. N. 2003. Palynological evidence for vegetational history in semi-arid areas of the Western Mediterranean (Almerı´a, Spain). The Holocene, 13, 109–119. Pulido-Villena, E., Reche, I. & Morales-Baquero, R. 2006. Significance of atmospheric inputs of calcium over the southwestern Mediterranean region: high mountain lakes as tools for detection. Global Biogeochemical Cycles, 20, GB2012, doi: 10.1029/ 2005GB002662. Reimer, R. W., Remmele, S. et al. 2004. IntCal04 Atmospheric radiocarbon age calibration, 26-0 ka BP. Radiocarbon, 46, 1029–1058. Rodrigo, F. S., Esteban-Parra, M. J., Pozo-Va´zquez, D. & Castro-Dı´az, Y. 1999. A 500-year precipitation

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record in southern Spain. International Journal of Climatology, 19, 1233–1253. Schmidt, R., Karin, A. K., Thompson, R. & Kamenik, C. 2002. A multi proxy core study of the last 7000 years of climate and alpine land-use impacts on an Austrian mountain lake (Unterer Landschitzsee, Niedere Tauern). Palaeogeography, Palaeoclimatology, Palaeoecology, 187, 101– 120. Schulte, L. 2002. Climatic and human influence on river systems and glacier fluctuations in southeast Spain. Quaternary International, 93–94, 85– 100. Steinmann, S. 1978. Postglaziale Reliefgeschichte und gegenwa¨rtige Vegetationsdifferenzierung in der alpinen Stufe der Su¨dtiroler Dolomiten (Puez-und Sellagruppe). Landschaftsgenese und Landschaftso¨kologie, 2, 1–93. Veit, H. 1988. Fluviale und solifluidale morphodynamik des Spa¨t- und Postglazials in einem zentralalpinen Flusseinzugsgebiet (su¨dliche Hohe Tauern, Osttirol). Bayreuther Geowissenschaftliche Arbeiten, 13, 1– 167. Wanner, H., Holzhauser, H., Pfister, C. & Zumbu¨hl, H. 2000. Interannual to century scale climate variability in the European Alps. Erdkunde, 54, 62–69. Zielinski, G. A. 1989. Lacustrine sediment evidence opposing Holocene rock glacier activity in the Tempel Lake Valley, Wind River Range, Wyoming, USA. Arctic, Antarctic and Alpine Research, 21, 22–33.

Aeolian processes and features in cool climates MICHAEL E. BROOKFIELD Institute of Earth Sciences, Academia Sinica, P.O. Box 1 –55 Nankang, Taipei 1159, Taiwan (e-mail: [email protected]) Abstract: This review concentrates on the characteristics of wind processes and products in cold lowland deserts. Its main aims are: to show how difficult it can be to separate wind-generated from water-generated processes and features without detailed study of lamination types; to suggest cold aeolian interpretations for some ancient pre-Quaternary rocks; and to encourage students to revisit many ancient deposits with such interpretations in mind. Although upper Quaternary and Recent cold-climate aeolian deposits have been studied for many years there are very few older examples described. Martian features are included since Mars is now one of the most exciting areas of study for geologists.

Winds affect landscapes in all climates, but the fundamental requirements for significant effects are winds strong enough to erode and transport material and the lack of any protective covering to the surface. On Earth, normal and cyclonic winds are usually sufficient, at least annually, to erode and transport sediment anywhere that there is inadequate protective vegetation (Yaping Shao 2000). Semiarid to extremely arid (hyperarid) conditions affect about one-third of the present Earth’s surface and are planet-wide on other bodies in the solar system (Carr 2006). Extremely arid conditions (with no rainfall for periods of one year or longer) cover only about 4% of the Earth (Goudie 2002), but could have extended over much greater areas within supercontinents in the past. On Earth, there are now only three areas that regionally have inadequate vegetation because of dryness: the deserts around the subtropics 308N and 308S of the equator, where air in the equatorial Hadley circulation begins to descend becoming denser and drier; the temperate deserts of Central Asia; and the polar deserts where there is rarely any liquid water (Goudie et al. 2000; Seppala 2003) (Fig. 1), though locally such areas also occur around alpine glaciers. In pre-Silurian times, the absence of structured plant communities means that wind action would be important everywhere and soil (and all other) processes very different (Dott 2003; Went 2005; Corenblit & Steiger 2009). On other planets and moons of the solar system with atmospheres, wind action is much more extensive in the absence of vegetation (Greeley & Iversen 1985). On Mars for example, winds are normally too gentle (given the thin atmosphere) to move even dust although when strong winds do occur they can affect the entire planet (Fenton 2003) (Fig. 2). Since aeolian processes are most important in deserts, we need to state what we mean by a

desert. A simple definition is that deserts are dry areas with extreme temperatures and sparse vegetation that receive less than 200 mm precipitation (rain or snow) annually, although this masks great variations in mean, annual, weekly and daily temperatures. Warm deserts are those where annual surface temperatures are above 0 8C (such as the Sahara). Cool deserts are where annual surface temperatures are below 0 8C (such as Antarctica) and temperate deserts are those with annual changes from cool to warm (such as the central Asian deserts). On these definitions, Mars and Titan are cool (really very cold) deserts and Venus is a warm (really very hot) desert. These climatic definitions create some problems in distinguishing warm, temperate, and cool deserts in the stratigraphic record, as we will see. There is fundamentally no difference between the aeolian process and features in warm and cool deserts, apart from temperature which affects the density and viscosity of air (with a larger effect than that previously recognized, McKenna-Neuman 2004a) and the rate of chemical breakdown of minerals (Honda & Shimizu 1998). This review concentrates on the characteristics of wind processes and products in cool lowland deserts and evaluates criteria for distinguishing aeolian features in cool deserts from those in temperate and warm deserts (though the latter depends on mostly non-aeolian features). The main aims are to show how hard it can be to separate windgenerated from water-generated processes and features without detailed study of lamination types; to suggest cool aeolian interpretations for some ancient pre-Quaternary rocks; and to encourage students to revisit many ancient deposits with such interpretations in mind. Although late Quaternary and Recent cool-climate aeolian deposits have been studied for many years (Koster 1988; Koster &

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 241–258. DOI: 10.1144/SP354.16 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Arid zones on Earth (modified from Goudie 2002).

Dijkmans 1988; French 2007); there are very few older examples described. Martian features are included since Mars is now one of the most exciting areas of study for modern geologists.

Wind production and patterns Winds in cold climates (apart from those due to altitude) are generated by high pressure over ice

Fig. 2. (a) Mars north polar ergs (courtesy of NASA); (b) detail of north polar erg, Mars (courtesy of NASA/JPL/ Arizona State University).

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caps, which produce anticyclonic circulations and katabatic winds and by cyclones. In katabatic winds, the lower part of the atmosphere cools against the ice cap and the cold air then flows gravitationally downhill and outwards below the warmer air above (Heinemann 1999). These strong cold winds often warm as they flow over the surrounding landscape or ocean, giving rise to unstable convective conditions that increase surface sand transportation (Frank & Kocurek 2006), and develop large longitudinal vortexes (visualized as cloud streets) (Scorer 1978). Such longitudinal vortexes can explain the alignment of many linear dunes with predominantly unidirectional winds. The observed sideways movement of longitudinal vortexes, often brought up as an objection to the hypothesis (Lancaster 1995), can be easily solved by considering the ‘washboard effect’ whereby the vortexes bounce, controlled by kinematic waves at the boundary between the atmospheric boundary layer (ABL) and the troposphere (Taberlet et al. 2007). So far, however, there has not been much interest (except from those studying extraterrestrial examples) in the influence of the varying structures of the ABL under different conditions on sediment transport and deposition.

Particles Production, composition, grain size Sedimentary particles in cold climate can be of solid water (ice, snow), rock or minerals. Snow and ice particles are produced by precipitation and recrystallization. These range from small broken dust-sized needle-shaped crystals to coarse rounded recrystallized firn. Rock and mineral particles are produced by weathering and erosion of pre-existing rocks and sediments, and may inherit the character of the

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supplying material (e.g. platy feldspars from granites and rounded quartz from fluvial sandstones). The composition reflects available source materials, climate and transportation distance. Many sand dunes on Mars are dark because they consist of coarse basalt grains (about 0.5 mm diameter) rather than quartz sand (Carr 2006). Since chemical breakdown is slower under cold conditions, the initial sediment may be immature in both texture and composition (Smith 2009); aeolian sands derived from glacial outwash are more immature and have less rounded grains than warmer climate aeolian sands, where sand is often transported much longer distances with continual reworking in large deserts (Fig. 3). Silt and clay may consist of a greater mixture of less weathered grain compositions than in warmer climates. Large amounts of fine sediment may be removed by the wind from glacial outwash and former proglacial lakes, such as mega-Lake Agassiz in Canada, and transported long distances (Fig. 4a). Although the relative importance of glacial grinding and chemical action in the formation of fine sediment (loess) is still debated (Smith 2009), unstable minerals in these loess soils are now breaking down under warmer and more humid conditions to supply abundant inorganic elements needed for plant growth (Fig. 4b). Quaternary loess deposits form some of the most fertile soils in the northern hemisphere in the great belt from the American midwest though central Europe to the Ukraine and northern China (Pe´csi 1991). After production, the particles are then modified by wind abrasion and transport.

Transportation and deposition Wind transports sand as a creeping carpet moved forward by grain impacts and by saltation, whereas

Fig. 3. Fine- to medium-grained aeolian dune sands. (a) Well-rounded, mature quartz sands of warm deserts. (b) Subrounded, submature sands with ilmenite (dark), felspar (cloudy) and quartz (translucent) grains of temperate to cool deserts.

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Fig. 4. Loess. (a) Dust storm in Siberia and (b) loess section near Lanzhou, China. Redder areas are soils (interglacials/ interstadials). Inset shows thin section of soils – note angular quartz (white), rock fragments (grey) and iron oxides (black) in clayey matrix.

dust (silt and clay) is normally carried in suspension (Fig. 5a) (Nickling & McKenna-Neuman 2009). Gravel is rarely moved except under exceptional conditions and is often sandblasted to form ventifacts, though gravel megaripples occur in Antarctica and South America (Milana 2009). Sand-sized grains impact at smaller sizes in air compared to water (due to the lack of viscosity-cushioning effects). Snow and ice crystals are broken and shattered: rock and quartz grains are broken and rounded faster (and at smaller sizes) and unstable minerals with cleavages are shattered into dust particles that can then be rapidly decomposed chemically. Cold-climate sands are usually more compositionally immature than warm-climate sands (Fig. 3). Although rippled interdune sands may consist of coarse sand, most interdune sands are bimodal in the sand fraction and have higher silt and clay contents when compared with adjacent dune samples; this is either due to infiltration or protection from erosion by the larger grains (Folk 1968; Warren 1971). On Earth, sand transport rates may be as much as 70% greater in cold deserts than in warm deserts for equivalent wind speeds (McKennaNeuman 2004b). Experiments in Antarctica show that threshold wind velocities for sand movement are lower (5.5 m s21) in cold deserts than in warm deserts (6– 7.5 m s21), which may explain why the Antarctic dry valleys have high sand flux rates (Lancaster 2002). Furthermore, ice particles reach a hardness approaching that of feldspar (7) at subzero temperatures and can erode softer rock significantly (Koch & Wegener 1930). On Mars, the wind speed needed to move quartz sand in saltation is less than on Earth because of its

lesser gravity. This and higher wind speeds causes higher rates of transport on Mars than on Earth (Fig. 5b; Carr 2006). Both sand and dust may erode underlying materials. In the following two sections, modern cool-climate aeolian erosional and depositional features are described and possible ancient examples are suggested. Allowance has to be made for the very different densities of snow, ice, sand and dust when comparing such features. On Earth, direct evidence of cool conditions such as snow and ice deposits may not survive for geologically long periods. On Mars and elsewhere, however, such conditions (and features) may persist for aeons. Since cool-climate features often require identification of ice-related accompanying structures, it may be confusing where younger processes superimpose features on much older coolclimate materials: for example, modern ventifacts in a temperate climate are now developing on pebbles embedded in eroding Permo-Carboniferous tillites (Wilson & Edwards 2004).

Erosional features Erosional features are the result of abrasion by sand- and dust-sized particles (snow, ice rock or mineral) in the wind (Allen 1965; Matsuoka et al. 1998). Although we focus on smaller features here, wind erosion is possibly forming angular unconformities in the Antarctic ice sheet as it deforms over subglacial mountains (Welch & Jacobei 2005). † Fluted surfaces occur on surfaces hard enough to be polished by sandblasting. They consist of

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Fig. 5. Sediment transport. (a) Modes of sand transport: creep, saltation and suspension – there is no sharp boundary between these (courtesy of Ann Vermillion); and (b) fields of transport modes for quartz sand on Earth and Mars: frictional shear velocity and mean wind speed plotted against grain diameter (courtesy of Joseph Veverka, Cornell University/NASA).

scalloped interlocking surfaces with patterns indicating the average direction of sand-moving winds. Fluting is associated with zones of sand transport. It is also found to be associated with zastrugi (see below) and on some bedrock surfaces underlying ancient aeolian sandstones, and is common on Mars (Carr 2006). † Pans are shallow depressions with smooth rounded outlines and normally develop in finer grained, commonly lake, sediments: water collects in them during warmer periods and they may develop evaporitic crusts as well as crescentic dunes on their downwind margins. Evaporites

can develop under cool conditions, as in the Antarctic Dry Valleys, so this is not a diagnostic feature of warm deserts (Doran et al. 2004). † Regs (or serirs) are stone lags that form by removal of sand and dust and can form intricately interlocking stone pavements that armour the surface and protect underlying sediment from wind erosion. Many periglacial regs develop on outwash fans and have ventifacts associated with them (Fig. 6). † Ventifacts are pebbles facetted by sand-laden winds and may show two or more facets (Fig. 6b; Knight 2008). Like fluted surfaces and

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Fig. 6. Ventifacts (a) on outwash plain, Iceland and (b) basalt ventifacts on (a) (courtesy of Till Nierman).

yardangs (see below), they tend to preserve one predominant wind direction (presumably the strongest). Where well preserved they can be used to determine ancient wind directions, for example on proglacial Quaternary outwash plains (Mackay & Burn 2005; Gillies et al. 2009). Pre-Quaternary ventifacts are recorded from many ancient warm deserts, usually reworked into water-laid deposits (Brookfield 2000). They should also be common in ancient cool desert deposits. † Yardangs are wind-abraded ridges formed of cohesive material such as silts or rock. The classic form is a streamlined teardrop shape (like drumlins) with a steep upwind slope, a gentler downwind slope, and smooth sides running up to a sharp ridge. They vary in size from one to thousands of metres in length and vary greatly in morphology. Yardangs, especially in softer material, can show the average wind direction during their formation as well as the local airflow around them. Their orientation relative to sand streaks (and their orientation on hard and soft materials) may differ, indicating changing wind patterns with time (Fig. 7a). PreQuaternary yardangs occur at a (warm desert) Permo-Triassic unconformity in Utah, USA (Tewes & Loope 1992). † Zastrugi are fundamentally yardangs in snow and ice consisting of wind-parallel grooves or furrows (Fig. 7b). Since snow and ice particles may cohere, overhanging cornices may develop. † Ridges and swales are large features marking zones of aeolian sand transport. They occur between the large sand seas on Mars and in warm deserts on Earth (Mainguet 1978). The rapid alternation of climatic conditions and

abundant fluvial transport presumably minimizes their development in Quaternary periglacial deserts on Earth.

Depositional features Modern depositional features consist of all grain sizes from gravel to clay. Gravel deposits are primarily lags (see above) although very strong winds can also pile small pebbles into transverse ridges (or megaripples), as in the Antarctic dry valleys. Sandsized deposits accumulate behind obstacles and in topographic lows to form accumulations of various thicknesses and extent. Silt and clay deposits accumulate in topographic lows and where wind velocities decrease; they form extensive deposits marginal to Quaternary glaciers. Because similar bedforms can form in both air and water (Nin˜o et al. 2002), the first thing to do is identify aeolian layering; this is easiest for sands.

Aeolian sand laminae Aeolian sand laminae differ from subaqueous laminae formed in similar ways because of the different characteristics of the fluid, and are of four main types (Hunter 1977). Planebed lamination is produced by wind velocities too high for ripple formation and is analogous to upper-regime planebed in aqueous deposits. It occurs only in coarse sands and granules and can be used to recognize sand sheet deposits. Climbing ripple lamination closely resembles aqueous varieties, but the laminae coarsen upwards and the ripple foresets are difficult to recognize because of the low relief of the moving ripples. Hunter (1977) distinguished two main types: rippleform strata occur when the ripple foresets can be

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Fig. 7. (a) Yardangs on Mars. The two different orientations (arrows) in this multiple-layered, eroded terrain suggest changing wind patterns over time (courtesy of Steven Williams, NASA). (b) Zastrugi in Antarctica (courtesy of Kay Welch).

identified, whereas translatent strata occur when only the bounding surfaces between migrating ripples are visible. Sets in both types are inversely graded and relatively closely packed (average porosity 39%). Rippleform strata indicate very high rates of sedimentation relative to rates of migration, producing steep angles of climb. Grainfall lamination is produced by deposition from suspension, typically in the lee of obstacles such as dunes. Grain segregation is relatively poor and laminae often difficult to see. Packing (average porosity 40%) is intermediate between those closely packed traction laminae of planebed and climbing ripples and the loosely packed sandflow strata described below. Distinctive features of grainfall lamination are gradual thinning, or tapering downwind (down a slipface and across an interdune area) and extreme variability of thickness, from less than 1 mm (wind gusts of a few seconds) to 10 cm or more (sustained gusts). Sandflow lamination (avalanche cross-bedding) is caused by slumping and consequent grainflow down slopes of sand supplied by creep, saltation and grainfall to the brink of slipfaces. Sandflow cross-strata are loosely packed (average porosity 45%), interfinger with grainfall laminae near their base and form lenses parallel to the slipfaces. Examples of all these types are in Hunter (1977) and Brookfield & Silvestre (2010). Identification of one or more of these lamination types is usually sufficient to distinguish wind from water as the

agent of deposition – and should be the first step in any study. Three other useful features are: adhesion ripple pseudolamination (with the laminae dipping upwind) and crinkly parallel lamination (both form by sand sticking to damp surfaces; Kocurek & Fielder 1982); and interbedded ice and snow layers which are diagnostic of at least occasional freezing conditions and melt to form structureless sand layers (Bourke et al. 2009). Many of these can be identified in the predominantly aeolian sand sheets deposited between 14 and 12.4 ka in NW Europe south of the Weichselian ice sheets (van Huissteden et al. 2001) (Fig. 8). The term ‘niveo-aeolian’, derived from such deposits, refers to mixed deposits of wind-driven snow, sand and any other particles (Cailleux 1972).

Aeolian bedforms Aeolian bedforms occur at three scales: ripples, dunes and draa (Wilson 1972). There are interesting and ongoing studies on scaling laws for the varying sizes of bedforms. For example, by using the drag length rs/(rf d) where rs is the grain density, rf is the fluid density and d is the grain diameter, Claudin & Andreotti (2008) show that ripples and dunes from water to various atmospheres plotted linearly on a graph of increasing drag length against increasing ripple wavelength (Fig. 9). Experimentation is difficult above ripple

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Fig. 8. Annotated lacquer peel of sedimentary structures in dominantly aeolian Weichselian coversands in The Netherlands (modified from van Huissteden et al. 2001).

sizes, and larger bedforms show such enormous lag effects that we can be sure that they are not adjusted to the current wind regime (Wilson 1972; Allen 1974). Another problem is that wind-tunnel experiments do not resemble natural flows, since the former are basically pipe-flow while natural flows have a free surface (the top of atmospheric boundary layers). It may sometimes be difficult to determine whether some bedforms are depositional or erosional or a combination of both processes. For example, even large longitudinal dunes in warm deserts often rest on lag deposits and may be very ephemeral features at times (Mainguet & Chemin 1983). In Antarctica, the large megadunes with wavelength exceeding 5 km seem to result from a combination of erosion and deposition (Fig. 10). Wind ripples. Wind ripples are transverse to flow and have wavelength/ripple height ratios much greater than in water (aeolian c. 15; water c. 9).

This can help distinguish them on bedding planes, and they may show climbing ripple lamination in section (Sharp 1963). Measurements from Mars suggest a wavelength/height ratio of c. 6.7, twice as steep as terrestrial wind ripples; the reason for this is not known. They are identifiable on exposed surfaces and as translatent strata in sections (see above). Dunes. Dunes may be small or large, simple or complex and linear (transverse or longitudinal) or three-dimensional (star dunes). The most important difference between ripples and dunes is that dunes scale with the flow that forms them; in other words, deep flows make large dunes (provided, of course, that there is enough time, space and sand for the dunes to develop fully). The flow depth for aeolian dunes is the height of the atmospheric boundary layer, commonly between 1 and 2 km on Earth and around 5 km on Mars (Scorer 1978;

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Fig. 9. Wavelength plotted against drag velocity for water ripples on Earth and dunes for various planetary bodies (modified from Claudin & Andreotti 2008).

Carr 2006). Dunes may be classified into barchan, transverse, longitudinal, parabolic and star varieties, with barchanoid for irregular transverse examples (Fig. 11). Transverse and longitudinal are commonly grouped together as linear with variation from transverse to the wind through oblique to longitudinal; parabolic seem to require vegetation pinning the arms to form (Lancaster 1995). Coolclimate dunes on Earth tend to be relatively simple, possibly because of the relatively short geological time in which they are active, and barchan and parabolic dunes seem to be the most common (Koster 1988). Dunes in Victoria Valley, Antarctica migrated at a significantly lower rate than warm desert dunes because migration is retarded by entrained ice and a reversing wind regime (Bourke et al. 2009). The most diagnostic feature of coolclimate dunes are interbedded snow and ice layers, which form structureless sand layers after melting. Without these as well as climate information, the dune shown in Figure 12a could be in a warm desert after a cold night, in a temperate desert in winter or in a cool desert. It is actually from the permanent interbedded sand, ice and snow layers (Fig. 12b) in Antarctica.

Although different types of aeolian laminae occur on different parts of dunes and can be used to recognize their different components, there is a problem in the erosion of all but the basal parts in ancient bedforms. In some cases, however, the basic distribution of laminae types in a complex three-dimensional dune can be recognized even after deflation. For example, Hunter (1977) noted the slipfaces (sandflow cross-strata), the saddles (climbing translatent strata) and the passage between the slipfaces and saddles (grainfall lamination) on a planed-off temperate coastal dune. Draa. Draa (singular and plural) are large sand bedforms between 20 and 450 m high, characterized by the superimposition of smaller dunes on them. As for dunes they scale with depth of flow and the thickness of the ABL, although the dune–draa relationship is also dependent on the structure of the ABL. Both transverse and longitudinal draa occur with superimposed dunes in modern sand seas, but these draa are in disequilibrium with existing wind regimes since they take hundreds to thousands of years to readjust to changing conditions (Wilson 1972).

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Fig. 10. Megadunes in Antarctica (courtesy NASA): (a) general view showing locations; (b) detail of area shown on (a); (c) oblique air photo of megadunes; and (d) detail of zastrugi.

Internal structures Identifying the larger primary internal structures of modern dunes and draa is mostly limited to groundpenetrating radar (GPR) studies (Bristow et al. 1996). Radar has good penetration in dry sand, although it also picks up hydrological boundaries (Van Overmeeren 1998). Together with thermoluminescence dating, GPR can identify different stages in bedform growth and destruction. We cannot be sure, however, that the deeper preservable structures relate to the modern wind regime, so such studies need to be supplemented with computer modelling of the structures produced by the diversity of migrating bedforms (Rubin 1987; Schwammie & Herrmann 2004).

Large-scale cross-bedding, consisting mostly of sandflow and grainfall strata, is often considered characteristic of large aeolian bedforms. The size of internal cross-bedding, however, is determined by the size of the bedforms that form it and also by the rate of climb of the bedforms. Both aeolian and subaqueous bedforms form similar structures in shelf seas and deep rivers (Rubin & Hunter 1982). Only the most rapidly climbing aeolian bedforms preserve more than a small part of the lower lee-slopes; bedform climbing needs to be considered before any interpretation can be made of ancient aeolian sandstones (Kocurek 1988). In topographically complicated areas, dune migration need not be consistent as shown by cross-bedding. In the cold-climate dome-shaped and transverse

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Fig. 11. Dune types (courtesy Ahmed Hemden).

dunes in Antarctica, three dune fields within the same valley experience different wind regimes due to topographic steering of the winds and drag against the valley walls (Bristow et al. 2010a, b). The internal structures of these dunes resemble those of warm deserts, although the common occurrence of (presumably cryoturbation) sand wedges extending upwards from their bases may be a distinguishing character (Fig. 13) Bounding surfaces separate bundles of strata in all bedforms. Draa can show two main orders of bounding surface, with downwind dipping surfaces caused by dune migration cut by more horizontal first-order surfaces due to migration of the draa (Brookfield 1977; Kocurek 1988). Reversing dunes also show such second-order bounding surfaces (Fig. 14).

Interdunes and interdraa Interdunes and interdraa are an integral part of aeolian bedform systems. In deserts with a limited sand supply, interdunes consist of lag deposits, coarse sand sheets, small isolated dunes and temporary lakes. Because of the way in which the sand is transported through the systems, longitudinal bedforms tend to have coarse lag and coarse sand sheets and dunes in the interdune areas. Transverse bedforms tend to have sabkhas and fine sand dunes between them. The size of the interdune (and interdraa) areas is also dependent on sand supply and on the stage of development of the sand sea in which they occur (Brookfield & Silvestre 2010). Sand sheets consist of flat areas of plane-bedded coarse sand usually overlying finer material and are

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Fig. 12. (a) Barchanoid dune in Dry Valleys, Antarctica (see also Fig. 14); (b) planebed interlayered ice and sand with snow above (a, b courtesy of Kathy Welch); (c) interbedded grainfall sand and ice (coastal dune northern Ontario, Canada).

Fig. 13. Bedding and bounding surfaces in Antarctic McMurdo Valley dunes (modified from Bristow et al. 2010a). (a–d) Radar cross-section at noted locations on view of dome-shaped dune showing net unidirectional migration with south-easterly downwards-inclined (second order) bounding surfaces resting on subhorizontal unconformity (first-order surface) and sand wedges extending upwards to various heights within dune.

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Fig. 14. (a) View to south of transverse dune in McMurdo Valley with (b) interpreted radar cross-section (modified from Bristow et al. 2010b).

basically lag deposits of coarse sand in the zone of transport and bypassing. Although sand sheets tend to be small, some are very extensive and may have subdued superimposed bedforms. For example, the late Weichselian (14–14 ka) coversands of Western Europe consist of a vast plain of sand between the glacial front and the more southern belt of loess (Kasse 1997). The only bedforms present are low sand ridges and small dunes of coarse sand (zibar) identifiable in sections by plane bedding and low-angle lamination in coarse sands (Fig. 8). On Mars, sand sheets cover enormous areas (Carr 2006). Silt deposits (loess) cover vast areas south of the Quaternary ice sheets, generally south of the coversands. Successive palaeosols indicate interglacials and interstadials, with the silt being deposited primarily during the glacials by katabatic winds although transport and reworking is still taking place (Roe 2009). Layered ice-rich silt deposits also form widespread plains in the Arctic and are easily eroded by water (Fig. 15). Since the unstable minerals decompose over time – significant chemical weathering is taking place in subarctic loess (Muhs et al. 2008) – the silt may be transformed from a porous framework to a muddy silt and eventually a greywacke. Without associated periglacial features such as ice wedges

and diagnostic fossils, it is probably impossible to identify cool-climate loess from loess redeposited in arctic, temperate and other lakes, though some examples have been claimed (Soreghan et al. 2008) and warm-climate loessite has been recorded (Kessler et al. 2001).

Fig. 15. Ice-rich layered silt deposits with ice wedges, Bolshoi Lyahovsky island (748N, 1428E; courtesy of M. Grigoriev).

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Post-depositional features Post-depositional structures may give important information on environments if carefully evaluated. Deformed aeolian cross-bedding is often found where there is seasonal melting of ice in coldclimate dunes (Koster 1988). It also occurs in temperate climate dunes, for example in the Sonoran Desert of Mexico (Fig. 16) and in the tropical Jurassic Navajo Sandstone of the western USA (Doe & Dott 1980), and is thus not diagnostic of cool-climate dunes. Sinkholes form where snow and ice melt to form large amounts of water which then drain downwards through porous aeolian sands to form depressions and ‘chimneys’ of disrupted sand. Such sinkholes are common on periglacial aeolian accumulations (Cailleux 1972; Koster & Dijkmans 1988). Since winter snow layers melt during the summer in temperate (and subartic) climates, deformed layers, layers of structureless sand and sinkholes should be common in ancient cold-climate aeolian sands. Burrows and traces of various organisms, including animal burrows and tracks and rootlets, may occur at some horizons where they indicate less arid conditions. Hoof-prints of large ice-age herbivores occur in late glacial aeolian sand sheets in Alaska (Lea 1996). Soils and palaeosols form between depositional episodes and often under different climatic conditions; for example, the palaeosols interbedded with loess often mark warmer conditions (Fig. 4b) even if only interstadial conditions (Haesaerts et al. 2005), so they cannot be used without further evaluation as indicators of climate during aeolian deposition. Furthermore, soils have changed with the evolution of land biotas through time, and only post-Miocene palaeosols (with the evolution of grasses) can be compared directly with recent soils. The interpretation of palaeosols

Fig. 16. Large slumps on star dune, El Gran Desierto, Sonora, Mexico.

requires not only knowledge of modern soils and the processes that form them, but also their changing biological make-up and processes that have evolved through time (Retallack 2001).

Distinguishing cold from warm aeolian deposits In order to distinguish between cold and warm Aeolian deposits, we require evidence of temperature that is difficult to obtain. On Earth, direct aeolian evidence of cold conditions is rarely preserved, is limited to snow and ice layers within sediments and may reflect only seasonal or diurnal freezing. The best evidence comes from non-aeolian features in interbedded deposits; for example, pseudomorphs after ikaiite (known as glendonites) in shales (Selleck et al. 2007) or ice-wedge sand casts infiltrated by aeolian sand although even these are disputable (Murton 2003).

Examples of ancient cool aeolian deposits Quaternary cool-climate aeolian deposits can often be identified by their distribution around the margins of former ice sheets and their preserved geomorphology. Pre-Quaternary cool-climate aeolian features need to be identified from associated coolclimate glacial, fluvial and lacustrine features; even this can be a problem, however. Older glacial periods probably had the same geologically rapid alternation of glacial and interglacial phases and aeolian deposits interbedded with glacial and glacial stream deposits, which may have accumulated during relatively warm (temperate) phases. There are few carefully described and identified cool-climate aeolian deposits: the most convincing are those associated with the Permo-Carboniferous ice age. In southern Australia, the upper Carboniferous Merrimelia Formation of the Cooper Basin consists of glacial tillites, outwash, glacial lakes deposits and aeolian sands (Williams & Wild 1984). The aeolian sands form a thick (up to 232 m thick) succession of interbedded medium- to coarse-grained sandstones dominated by wind-ripple and sandflow lamination and are interbedded with outwash fan deposits (Williams et al. 1987). They resemble the late Quaternary coversand deposits of northern Eurasia, formed on outwash plains and dominated by small dunes and sand sheets. Although the Cooper basin was at .508S latitude in the earliest Permian and there are tundra soils developed in south Australia just before (Retallack 1999), no structures such as sand wedges or structureless layers that can be directly related to snow and ice were described from the Merrimelia Formation (Williams et al. 1987).

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In Saudi Arabia, the lower Permian Unayzah B Formation consists of glacial sediments including deformed push moraines, subaqueous outwash fans and laminites (varved shales) overlain by fine-grained red-beds with intercalated fluvial and aeolian sandstones (Al-Laboun 1987), the latter identified by the lamination types. Again, there are no specific diagnostic cold-climate features and the aeolian sandstones could be part of a warmclimate post-glacial erg, especially as it passes up into thick tropical carbonates. Some of the Ordovician sandstones associated with the Saharan glaciation could be aeolian but these have not been investigated in enough detail, although this is changing with modern oil exploration (Le Heron et al. 2010). It is, however, hard enough to interpret the so-called tillites as glacial deposits ( personal observations in Mauritania to southern Algeria). A number of Proterozoic aeolian sandstones have been described (Stewart 2005) but these do not have any cold-climate characteristics and their occurrence during the wild climatic swings of the ‘Snowball Earth’ may make it difficult to identify cold-climate examples (Hoffman et al.

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1998). Fossil biotas would be of no help in the pre-Silurian. Curiously, the best established ancient coolclimate aeolian deposits are on Mars, and they are certainly the oldest being probably at least 3.5 billion years old (Fig. 17).

Conclusions Aeolian sandstones can now be easily distinguished from fluvial sandstones by the detailed structure of their layering. Recognizing cool-climate aeolian sandstones requires identification of ice- and snowrelated structures such as ice- (now sand-) wedges, structureless sand layers (formed by melting of snow/sand mixtures) and interbedding with other cool-climate sedimentary facies such as tillites, outwash and glacial lake deposits (although the latter can also be difficult to identify in ancient deposits). The few described pre-Quaternary coolclimate aeolian sandstones are not entirely convincing given that cool-climate structures can form in winters in temperate climates, that rapid climatic

Fig. 17. Strata and interpreted section at Meridianum Planum, Mars (modified from Grotzinger et al. 2005).

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changes occur during glacial/interglacial periods in temperate latitudes and there are no truly polar ancient examples. I appreciate the financial and intellectual support of the Institute of Earth Sciences, Academia Sinica, Taipei, Tawian (via its director, B.-M. Jahn) over the last two years and the comments of H. French regarding an earlier draft of the manuscript.

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Went, D. J. 2005. Pre-vegetational alluvial fan facies and processes: an example from the Cambro-Ordovician Rozel Conglomerate Formation, Jersey, Channel Islands. Sedimentology, 52, 693–713. Williams, B. J. P. & Wild, E. K. 1984. The Tirrawarra Sandstone and Merrimelia Formation of the southern Cooper basin, South Australia – the sedimentation and evolution of a glaciofluvial system. Journal of the Australian Petroleum Exploration Association, 24, 377– 392. Williams, B. P. J., Wild, E. K. & Suttill, R. J. 1987. Late Palaeozoic cold-climate aeolianites, southern Cooper Basin, South Australia. Geological Society Special Publication, 35, 233–249. Wilson, I. G. 1972. Aeolian bedforms – their development and origin. Sedimentology, 19, 173–210. Wilson, P. & Edwards, E. J. 2004. Further examples of ventifacts and unusual patterned ground from the Falkland Islands, South Atlantic. Geografiska Annaler, 86A, 107– 115. Yaping, S. 2000. Physics and Modelling of Wind Erosion. Academic Press, Boston, USA.

Cold-climate slope deposits and landscape modifications of the Mid-Atlantic Coastal Plain, Eastern USA W. L. NEWELL* & B. D. DEJONG U.S. Geological Survey, Mail Stop 926A, 12201 Sunrise Valley Drive, Reston VA 20192, USA *Corresponding author (e-mail: [email protected]) Abstract: The effects of Pleistocene cold-climate geomorphology are distributed across the weathered and eroded Mid-Atlantic Coastal Plain uplands from the Wisconsinan terminal moraine south to Tidewater Virginia. Cold-climate deposits and landscape modifications are superimposed on antecedent landscapes of old, weathered Neogene upland gravels and Pleistocene marine terraces that had been built during warm periods and sea-level highstands. In New Jersey, sequences of surficial deposits define a long history of repeating climate change events. To the south across the Delmarva Peninsula and southern Maryland, most antecedent topography has been obscured by Late Pleistocene surficial deposits. These are spatially variable and are collectively described as a cold-climate alloformation. The cold-climate alloformation includes timetransgressive details of climate deterioration from at least marine isotope stage (MIS) 4 through the end of MIS 2. Some deposits and landforms within the alloformation may be as young as the Younger Dryas. Southwards along the trend of the Potomac River, these deposits and their climatic affinities become diffused. In Virginia, a continuum of erosion and surficial deposits appears to be the product of ‘normal’ temperate, climate-forced processes. The cold-climate alloformation and more temperate deposits in Virginia are being partly covered by Holocene alluvium and bay mud.

During the last decade, evidence for past permafrost (marine isotope stage or MIS 4) and deep seasonal frost (MIS 2) on the Mid-Atlantic Coastal Plain south of the limit of the Laurentide Ice Sheet (LIS) has been documented (French et al. 2003, 2007, 2009) (Fig. 1). On the Mid-Atlantic Coastal Plain, outcrops are sparse and excavations are commonly ephemeral. Site-specific studies of rare exposures (Lemke & Nelson 2004; Smoot et al. 2009; Losco et al. 2010) continue to yield information about past permafrost. Spatial variations of frozenground-related phenomena across 38 of latitude can be compared and interpreted from a detailed surficial geological map coverage from New Jersey to Virginia. Other maps, including recent LIDAR (Light Detection and Ranging) images, ranging in scale from 1:24 000 to 1:250 000, facilitate spatial morphostratigraphic analysis from southern New Jersey, the central Delmarva Peninsula, southern Maryland and Tidewater Virginia. Recent mapping of surficial deposits on the Atlantic Coastal Plain loosely followed a paradigm of integrating alluvium, colluvium, aeolian, marine, littoral and deltaic deposits of variable thickness into morphologic sequences that link materials, depositional environments and landforms. The morphologic sequence, or morphosequence concept (Koteff 1974), was originally developed for mapping glacial –fluvial deposits underlying distinctive

assemblages of landforms. Stratigraphic relationships are interpreted from cross-cutting or overlapping morphosequences. The morphosequence concept can be applied to mapping of other geomorphic systems, including periglacial and estuarine environments. As an extension of the morphosequence concept, a continuous cover of complex surficial deposits can be defined as an alloformation of contemporaneous facies bounded by common unconformities. The morphosequence strategy is particularly useful when working with surficial deposits that have been mixed from common sources. Initially, this paper describes the distribution of sediments and source materials for the surficial deposits of the Mid-Atlantic Coastal Plain. A regional overview is then presented of the surficial deposits and geomorphology south of the Potomac River estuary where a distinctive cold-climate signature is generally absent. This is followed by a description of the region most proximal to the Wisconsinan (LIS) terminal moraine. New Jersey presents a landscape apparently sculpted by numerous cycles of cold-climate erosion and warmclimate weathering as interpreted from nested morphosequences. In comparison, Delmarva and southern Maryland are characterized by more subtle landscapes mantled with slope and windblown sediments that generally obscure complex terrain and

From: Martini, I. P., French, H. M. & Pe´rez Alberti, A. (eds) Ice-Marginal and Periglacial Processes and Sediments. Geological Society, London, Special Publications, 354, 259–276. DOI: 10.1144/SP354.17 0305-8719/11/$15.00 # The Geological Society of London 2011.

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Fig. 1. Index map of surficial deposits on the Mid-Atlantic Coastal Plain with location of detailed sites and figures discussed for the region. Dot pattern, upland on oldest Miocene fluvial-marine deltaic deposits (area of ‘Pine Barrens’); gravel pattern, deeply weathered, Miocene –Pliocene upland gravel; grey, Pleistocene marine highstand coastal and estuarine terrace deposits. Arrows indicate palaeotransport directions. Blue Ridge Mountains and Palaeozoic sources for Virginia upland gravels are west of the map area. Major rivers include C, Choptank; D, Delaware; H, Hudson; P, Potomac; Pc, Pocomoke; Pt, Patuxent; R, Rappahannock; S, Susquehanna. Site-specific areas of following figures include 2, Stratford Cliffs; 3, Colonial Beach South 1:24 000 scale quadrangle; 4, generic badlands diagram for New Jersey Coastal Plain uplands; 5, Haines Corners; 6, Snow Hill 1:24 000 quadrangle; 7, Blackwater National Wildlife Refuge (Golden Hill and Taylors Island 1:24 000 scale quadrangles); 8, Ewing Pit; 9, Dependable Materials Pit; 10, nivation hollows on Delmarva uplands; 11, Upper Patuxent River Valley exposures and profile.

deposits of earlier erosional events. The Pleistocene history of the Chesapeake Bay, which has eroded and filled during numerous transgressive– regressive events, has also been obscured by an ubiquitous blanket of Late Pleistocene deposits and subsequent Holocene alluvium and bay mud (Cronin 2000; Halka 2004; Newell et al. 2004).

Regional geological setting The distribution of surficial deposits across the MidAtlantic Coastal Plain is best understood in the context of the antecedent geological framework including provenance of sediments, basin analysis of sediment deposition and the subsequent history

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of weathering, erosion and development of the Plio-Pleistocene landscape. The Mid-Atlantic Coastal Plain region of the United States between latitudes 378N and 418N includes the onshore margins of a region of prolonged subsidence known as the Baltimore Canyon Trough (BCT) (Poag 1985; Poag & Sevon 1989). The BCT is bounded by the western edge of the coastal plain in Virginia and North Carolina, and to the north by a platform underlying the Long Island moraines. A roughly north–south hinge line west of the modern Chesapeake Bay separates thin coastal plain sequences on the west from thickening, more continuous coastal plain deposits to the east. The sediments that have accumulated in the BCT range from Lower Cretaceous fluvial to marine packages, through Palaeogene and Neogene marine sequences. The last depositional basin, filled with Miocene and Pliocene marine sediments (of the Chesapeake Group), is known as the Salisbury Embayment (centred under the Delmarva Peninsula). Collectively, formations of the Chesapeake Group represent the accumulated sediments stripped from several kilometres of the evolving topography of the Piedmont, Blue Ridge and Appalachian mountains since early –mid-Miocene (Poag & Sevon 1989; Pazzaglia & Gardner 2000). The inner edge of the Virginia coastal plain is covered by Miocene –Pliocene littoral deposits at altitudes of 60–100 m. North of the Potomac River estuary to New Jersey, the inner edge includes irregular upland eroded remnants of Cretaceous and Palaeogene deposits. Across the middle of New Jersey, Cretaceous and Palaeogene marine sediments are in covered, linear, NE trending contact with Piedmont metamorphic rocks and early Mesozoic basin sedimentary rocks. This juxtaposition suggests that the nearshore and littoral equivalents must have occurred farther inland and at higher elevations but have been subsequently eroded. The sediments at the northern edge of the coastal plain have also been partly eroded and covered by Pleistocene glacial deposits at the terminus of Laurentide glacial ice. Figure 1 shows the distribution of three distinctive lithologies of Cenozoic surficial deposits that are extensive across the Mid-Atlantic Coastal Plain region. These include the mid-Miocene delta deposits, Miocene –Pliocene upland gravel deposits and subsequent lower elevation marine and estuarine terrace deposits from multiple Pleistocene transgressive–regressive sea-level oscillations. The deposition and subsequent weathering, erosion and re-deposition created antecedent landscapes that were further altered by cold-climate processes in a zone that extended more than 200 km south of the LIS terminal moraine. The distribution of continental sediments, known informally as the ‘upland gravels’, defines the upper

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Pliocene and Pleistocene events that filled the BCT from New Jersey to Virginia. Initially, McGee (1888) included all upland gravels in his Columbia Gravel (later termed a Group by Jordan 1964) and considered them analogous and synchronous with the Lafayette and Citronelle Gravels of the midcontinent and Gulf Coast. However, the Columbia has been superseded by definition of discrete upper-delta-plain deposits that are in a facies relationship with Miocene and Pliocene marine formations of the Chesapeake Group. These continental deposits on the margins of the Salisbury embayment are evidence for episodes of uplift in the source areas. Across New Jersey and Delmarva, Miocene and Pliocene fluvial gravel deposits (the Bridgeton and Pensauken Formations) were sourced by a proto-Hudson –Delaware River of long duration and enormous discharge (Owens & Minard 1979; Stanford 2010). The sedimentary petrology of other similar fluvial gravels (Brandywine Formation) indicated palaeo-Potomac River origins across southern Maryland (Schlee 1957). Southwards in Virginia, each Piedmont-Blue Ridge draining river has also left its deltaic footprint across the inner coastal plain margin. Petrology of upland gravel indicates Palaeozoic sources had extended west of the Blue Ridge Mountains since early Miocene (Newell & Rader 1982). The end result of this geological history is that a common sequence of coarse-grained, permeable deposits (‘upland gravels’) overlies fine-grained tightly packed marine shelf deposits (Fig. 2). The deltaic lobes of fluvial gravel have been harder to erode than the finer grained marine shelf sediments. Erosion of this common sequence has produced complex topography with gullies, erosional scarps, colluvium, smaller-scale alluvial fans and floodplains that, in turn, are armoured with reworked gravel. The major rivers have cut deep valleys between the footprints of their original deltas. As the modern fluvial system became established during the PlioPleistocene, each valley became an archive of cutand-fill events associated with the changing sea level. Sequences of deposits associated with each event define transgressive and regressive phases of sea level. These can be fully described and timed by parsing the stratigraphy of terrace deposits into systems tracts (Miall 1997). Identification of systems tracts permits interpretation of dominant events and rates that can be correlated with details of climate change as interpreted from deep-sea cores and other climate proxies. Early recognition of marine-cut terraces and the role of sea-level and climate change were complicated by a lack of accurate topographic data. In addition, the underlying estuarine and marine deposits were not fully understood. Cooke (1930)

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Fig. 2. Stratford Cliffs, Virginia; a 50 m exposure showing upland gravel deposits over marine Miocene and Pliocene shelf deposits. Recent alluvial fan is being deposited below the spring.

fitted a series of marine terraces on the Coastal Plain landscape to represent the interstadials spaced between each great glacial epoch. This strategy has been amplified and adjusted in more recent decades (beginning with Oakes & Coch 1963). However, increasingly precise topographic data, new agedating techniques and the paradigm of parsing systems tracts from sediments in cores and terrace exposures combine to caution that: (a) not all sea-level rise/fall sequences are represented everywhere and (b) a unique event, if datable and mappable, may not be found everywhere at the same elevation. In fact, some valleys with tidal sediments may be too young to include terraces of great age. In others, younger events may have totally eroded older deposits and replaced them with new, re-worked sediment.

Weathering Prolonged chemical weathering has altered the mineralogy and texture of the upland gravels on the coastal plain across New Jersey, Delmarva, southern Maryland and Virginia. Much like saprolite on crystalline rocks of the adjacent Piedmont, all or most of the labile constituents have been converted to a residuum rich in clay minerals (kaolin, illitesmectite, montmorillonite, bauxite), oxides of iron, titanium and aluminium and enriched volumes of resistant minerals including magnetite, rutile and zircon (Owens et al. 1983). Quartz grains are deeply etched and feldspars may be non-existent to depths of several metres. Original sedimentary structures have been obliterated and overprinted with textures and structures resulting from soil-forming processes. Most of these original deeply weathered deposits are Miocene or Pliocene. The duration of surface exposure approaches the equilibrium rate for the dissolution of quartz, suggesting that chemical

weathering alone can lower an ancient landscape. Actual preservation of original Miocene depositional surfaces is unlikely assuming no additions of sediment to the original depositional surface. As equilibrium thicknesses for regolith on such deposits are approached, material may also be eroded from the surface at rates equal to the rate of regolith formation at depth (Pavich & Obermeier 1985; Pavich et al. 1989). The end result is upland gravel that is different from that originally deposited; it is better able to retain soil moisture and trap illuviating materials in a downwards-thickening soil B horizon. Subsequent surficial deposits, as well as marine and estuarine terrace deposits, have recycled these weathered products from the original regolith.

Geomorphic comparisons of unfrozen and previously frozen coastal plain landscapes Virginia A portion of the Virginia coastal plain between the Potomac River and the Rappahannock River, originally mapped for documenting neotectonic effects (Newell 1985), was selected for the study of Holocene sediment production, transport and storage in the Chesapeake region. The geologic map of the Colonial Beach South Quadrangle (1:24 000 scale; Newell et al. 2005) includes all the physiographic settings and geologic framework typical of most Chesapeake coastal plain landscapes. Figure 3 shows these relationships on a perspective diagram derived from the geologic map. This area in Tidewater Virginia is useful as a control when examining the cold-climate evidence from the coastal plain that is south of the widely recognized LIS terminal moraine shown on Figure 1. The geomorphology and surficial geology of the map area adjacent to the Potomac includes

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Fig. 3. South view, oblique perspective from the Colonial Beach South 1:24 000 scale quadrangle. Local relief ranges from 0 –60 m. ‘R’ is eroded, deeply weathered regolith; untagged grey areas are eroded Miocene– Pliocene marine substrate; ‘Cr’ is colluvial– alluvial deposits eroded from ‘R’; ‘C’ is younger colluvium from mid-Pleistocene to Recent; ‘CC’ is 25 m estuarine terrace deposits (Charles City Formation, MIS 11); S is 12 m estuarine terrace deposits (Shirley Formation, MIS 7?); ‘T’ is 3– 4 m estuarine terrace deposits (Tabb Formation-Kent Island Formation, MIS 5e); and ‘A’ is modern alluvium including Potomac River beach and tidal wetlands.

deeply-weathered upland surfaces at 60 m elevation underlain by upland gravels, intermediate slopes covered with surficial deposits, steep headwater ravines that expose Miocene-age marine-shelf deposits and lowland sequences of estuarine terraces. The oldest terrace is 25 m in altitude with a thick fill of rising-stage systems tract deposits that are deeply weathered. This terrace is thought to be about 0.5 million years old, possibly MIS 11. Additional terraces descend in elevation to about 3–4 m; the lowest (youngest) is attributed to MIS 5e. Slope deposits and fan deposits are graded across the inner edges of each terrace which were originally at the base of wave-cut cliffs (Fig. 2). The common pattern for landscape development in these settings begins with springs that discharge from the interface of upland gravels and underlying marine substrate. These springs mark the heads of gully and ravine erosion (Fig. 2). Sediments are transported down first-order valley floodplains by alternating reaches of anastomosing channels and meandering trunk channels. Changes in gradient due to either sea-level change or climate-driven (meteoric) discharge have regulated the volumes and rates of sediment transport. The floodplains terminate in small deltas at tidewater. Slope deposits on the sides of ravines are commonly graded to floodplains with local undercuts and slope failures. On upland slopes, the locally continuous colluvial cover may be very old and graded to the

oldest terrace surface. Soil descriptions from the full array of topographic settings document increasing complexity of soil profiles as a function of increasing altitude and age (Markewich et al. 1987). For example, soils on low, Upper Pleistocene terraces of the Rappahannock River valley (location R, Fig. 1) are thin and organic-rich, whereas soils on the higher terraces and uplands are deeply oxidized and leached; surface erosion of the top of the regolith is common. As shown in Figure 3, many upland and intermediate slopes are covered with thin slope deposits (unit ‘Cr’) eroded from the proto-regolith of the highest areas. In comparison with descriptions of the more northern regions that follow, it must be noted that indicators of past permafrost are singularly lacking and the deeply ravined landscape and flat terraces are nearly devoid of dune fields and windblown sand sheets. All the present geomorphic processes and deposits are a continuum with those active during cold-climate stages. Further south, massive dune fields are common downwind of all major Coastal Plain river valleys from southern Virginia to Florida.

New Jersey When mapping the surficial geology in New Jersey (Newell et al. 2000), a strong correlation between age/altitude and complexity/depth of soil profile

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formation had initially been expected as seen in Virginia. Exposure by exposure across southern New Jersey, it gradually became apparent that the materials immediately below the surface to depths of 1 –1.5 m were commonly diamictic, even at very low-angled slopes and on nearly flat plains. Large clasts were matrix-supported. There were admixtures of windblown deposits, and the surficial deposits were commonly homogenized. Locally, where the physically rearranged surficial deposit was disconformable on a weathered gravel bed, the long axis of large pebbles had been rotated into near-vertical orientations from low imbrication angles of original fluvial deposition. Beneath the homogenized layer, deep regolith commonly appeared to have been stripped of near-surface horizons. Locally, the truncated regolith had been deformed beneath the homogenized layer. Normal sequences of leached, superficial horizons were missing from the tops of ancient weathering sequences (as seen south of the Potomac River estuary). In New Jersey, county soil map units suggest that modern soil-forming processes are superimposed on the reworked surficial sediments. Previously, geologists in New Jersey had noted the ubiquitous patterns of relict dune fields, closed depressions and deformed sediments. These had been interpreted in the context of a tundra-like landscape formerly underlain by permafrost (Wolfe 1953, 1956, 1977). Unfortunately, this concept did not gain much traction and was rejected in some cases (Black 1983). Today, the criteria for accepting past permafrost and periglacial processes south of the LIS terminal moraine have been satisfied (French et al. 2003, 2005, 2007, 2009). Other patterns of major broad valleys with severely under-fit streams also emerged. Some broad valleys were braid plains; others were blanketed by windblown sand or by Holocene peat and swamp deposits. Large low-gradient alluvial fans were mapped that appeared to have issued from the mouths of tributaries onto the floodplains of broader trunk streams. The modern floodplains are inset and narrow with meander bends scaled to discharges that are orders of magnitude less than discharges indicated by the broad valleys with braid plains. Modern sediment loads reported from Pine Barrens rivers are small (less than 10 tons per square mile). Today’s rivers that drain the Pine Barrens terrain (Fig. 1) of the New Jersey Coastal Plain are driven by shallow groundwater base flow. It is unlikely that these modern low-powered rivers have eroded kilometre-wide valleys inset with extensive deposits. The geomorphology suggests that the palaeohydrology of these river valleys was different; the substrate must have been less permeable when frozen (Newell & Wyckoff 1992).

Most of the upland interior of the New Jersey Coastal Plain can be interpreted as a palaeolandscape, created by past geomorphic processes. Evidence for cold-climate processes that shaped the palaeolandscape includes frost wedges, involutions and other deformations, patterned ground, extensive slope deposits on low gradients, windblown sand sheets and relict dune fields (Newell et al. 2000; French et al. 2003, 2005, 2007, 2009). The landscape tells the story of a formerly active system of seasonal discharge and mass movement on low slopes overlain by windblown coversand. The geomorphic system produced extensive deposits now dormant on the landscape. The evidence for widespread permafrost explains the scale and distribution of these deposits and the palaeohydrology. However, timing of permafrost aggradation and degradation is only a part of the story because effective erosion and transport peak during brief seasons of active layer thaw and during climatic warming and degradation of the permafrost (French et al. 2009). The Holocene warming climate fluvial system has only partially altered the palaeolandscape. Superficial cold-climate processes have given way to groundwater infiltration and minimal surface erosion. Although there is an extensive blanket of coldclimate slope, windblown and fluvial deposits across the surface of the New Jersey Coastal Plain landscape, not all aspects of the erosional topography are the result of stripping and deposition from Pleistocene climate oscillations (French et al. 2007). During surficial geologic mapping (Newell et al. 2000), local excavations revealed that intermediate upland gravel deposits, eroded from original Miocene fluvial, deltaic and marine sediments, had been re-deposited as fans, pediments or floodplain deposits. As headward erosion of valleys proceeded, new deposits were created from downwasting and erosion of old surficial deposits as well as the original Miocene substrate. Erosional remnants of old gravel deposits underlie isolated knobs and benches that are staged within the evolving topography. The resulting topographic features and their underlying deposits can be interpreted as systems of morphosequences (Fig. 4). During the mapping, dating technologies were insufficient to determine the actual ages of deposits; one upland watershed might include three sequences, another two or five. The surficial deposits in each morphosequence were therefore classified only by spatial relationships between landforms and deposits within each watershed. Within the morphosequence model for the New Jersey coastal plain, all ages are described in the relative context of cross-cutting relationships. The model for the evolution of this erosional landscape can be interpreted as a dominantly warm-climate

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Fig. 4. Block diagram of upland geomorphology of the New Jersey Coastal Plain. Map units show the domain of complex slope deposits eroded from upland gravels and transported down slope to broad valleys. Geomorphic system has been active during climate degradation leading to cold-climate periods of active layer processes and eventual thaw of past permafrost; warm pluvial periods may have been less effective (Newell et al. 2000). Tch is marine Miocene (Cohansey Formation); light grey is marine Miocene fluvial-deltaic deposits; R is deeply oxidized and leached regolith on eroded upland; 2, 3 and 4 are old eroded remnants of colluvial and alluvial surficial deposits (‘Qca’); and 1 is floodplains with braided and meandering streams from the LGM to the Younger Dryas.

fluvial process model or a dominantly cold-climate mass-wasting model. Either way, the New Jersey coastal plain morphosequences reflect erosion and aggradation that may be interpreted as broadly synchronous cyclic artefacts, or they may be interpreted as random in time and space. Sufficient data have yet to be found to argue either interpretation to the exclusion of the other. A continuum of fluvial processes (only occasionally overprinted by thin deposits of cold-climate landscapes) might favour the random temporal/spatial interpretation. In contrast, a climate-change-driven model would favour a series of broadly synchronous events. Neither interpretation is probably sufficient to exclusively interpret the remnant surficial deposits of great age and preservation back into the Pliocene. Younger, more extensive morphosequences may be suggestive of contemporaneous cyclic origins. In coastal watersheds for example, some of the younger morphosequences are lower in elevation being either at grade with the tops of marine terraces or at grade with valleys that are cut through marine terraces. In southern New Jersey, a complete climate cycle can therefore be represented by marine terrace deposits emplaced during warm-climate stages and by upland erosion and slope deposits emplaced during cold-climate low sea-level stages. Intervening periods of falling sea level, occurring during climate deterioration, may have been dominated by increased fluvial erosion stimulated by steeper local gradients. Corresponding periods of warming climate would have been characterized by sediment storage as systems of slopes and streams lowered

their gradients in adjustment to rising sea level and changing runoff. In fact, this is the pattern in Virginia. A return to warm climate and lower gradients would correspond to periods of relative stability and weathering. Pleistocene cold climates have been characterized for decades as periods of vigorous sediment transport and landscape modification (Washburn 1980; French 2007). Periglacial slope deposits in Pennsylvania (Braun 1989; Ridge et al. 1992) indicate widespread cold-climate slope processes during the Late Pleistocene, synchronous with New Jersey coastal plain surficial deposits. Opinions vary concerning the presence of vegetation (Russell & Stanford 2000) and its effectiveness for inhibiting sediment movement in periglacial New Jersey. Empirical evidence from New Jersey outcrops and from numerous high-latitude analogues indicates that sparse patches of boreal forest and tundra did not inhibit mass-wasting and aeolian transport on the past permafrost terrain of New Jersey (French et al. 2009). The spatial complexity and deep weathering of the oldest erosional remnants suggest that the New Jersey coastal plain uplands may include parts of morphosequences that are older than the Pleistocene. Some high-level remnants may be Pliocene in age and correlative with the transport and deposition of upper Miocene and Pliocene fluvial gravel into the Delmarva area (Pensauken and Beaverdam formations) (arrows on Fig. 1). All the New Jersey morphosequences are covered by a variable mixture of windblown and diamictic sediments

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deposited during the cold climates of the Late Pleistocene. Details of such a sequence have been described by Newell (2005) and French et al. (2007) from a 1986 excavation at Haines Corners (Fig. 5). This sequence includes locally derived slope deposits interbedded with windblown sand sheets and dunes overridden by thick debris-flow gravels from the top of the hill. A small sand-filled wedge with sand-filled veins extends downwards and penetrates most of the sequences. Seasonally frozen ground might have been sufficient for its formation. Although such exposures are rare, the coldclimate slope deposits at Haines Corners may be typical of the complexity within the cover of thin, diamictic and subtle bedded deposits that mantle the low-gradient surfaces of the New Jersey coastal plain. Over 100 years ago, during early stages of US Geological Survey geologic mapping in the Philadelphia–Trenton corridor (Bascom et al. 1909a, b), Salisbury & Knapp (1917) found it necessary to create a map unit on the New Jersey Coastal Plain that was thin, poorly sorted yet locally bedded and generally continuous. It obscured most of the Cretaceous and Tertiary outcrops and covered extensive low-gradient areas. They described it as they saw it and named it the ‘unclassified deposits’. Over 50 years later, Minard & Rhodehammel (1969) and Owens & Minard

(1975) attributed these materials to cold-climate fluvial processes. More recent mapping of the New Jersey coastal plain surficial geology (Newell et al. 2000) found these ‘unclassified deposits’ of Salisbury & Knapp (1917) to be accurate representations of the Quaternary mixed colluvium and alluvium map units (‘Qca’ as indicated in Fig. 4 caption) that define morphosequences of surficial deposits.

The interface between unfrozen and previously frozen coastal plain landscapes Coversands of Delmarva The Delmarva Peninsula extends south and west of New Jersey between the Delaware Bay and the Chesapeake Bay (Fig. 1). The highest elevations of the Delmarva uplands are barely 25 m above sea level and are probably enhanced by Wisconsinan sand dunes. The upland gravels of Delmarva include the Pensauken Formation, the Beaverdam Formation and, locally in Delaware, the Columbia Group (Jordan 1964). These upper delta plain deposits, largely deposited by the proto-Hudson/ Delaware system, are younger than the Bridgeton Formation in New Jersey and are considered late Miocene and probably Pliocene in age. Curiously, estuarine terrace deposits occur at the same altitude

Fig. 5. (a, b) Photos from 1986 foundation excavation at Haines Corner, New Jersey show assemblage of diamictic deposits interbedded with windblown deposits and deformed by slope and cold-climate processes. Uphill source to right of photos (from Newell 2005).

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in the Potomac estuary valley (Fig. 3) where they have been mapped as the Charles City Formation (Mixon et al. 1989; Newell et al. 2005), a unit considered to be mid-Pleistocene (possibly MIS 11). This stage is thought to be the warmest, highest sealevel stand in the mid-Pleistocene. The accordant elevations of these deposits of different ages indicate that the Delmarva area records cumulative subsidence while the Tidewater Virginia area records cumulative uplift toward the Fall Line (Newell 1985). Rasmussen & Slaughter (1955) originally noted ancient dune fields and elliptical, shallow, swampy flats rimmed by low sandy ramparts. These were visible on aerial photographs but indistinct on the coarse (1:24 000 scale) topographic maps of the time. They were thought to be similar to Carolina Bays (Thom 1970) and termed Maryland Bays. An early interpretation attributed the ‘bays’ to the melt-out of grounded icebergs on the old marine terraces. However, terrace altitudes represent marine highstands during warm-climate stages. ‘Bays’ on the older higher terraces would have been degraded and younger ‘bays’ would be more distinct. Recent imagery shows that they are all part of the same allcovering blanket of surficial sediment. Subsequent mapping in the late 1960s and 1970s noted the distribution of the ‘bays’ and the extensive cover of windblown surficial deposits. The thickest dune field accumulations were termed the Parsonsburg Formation (Owens & Denny 1978; Denny & Owens 1979; Denny et al. 1979), although it was also noted that uncommon exposures included stream-sorted alluvial deposits and poorly sorted colluvium. Peat beds and wood preserved within the map unit provided C14 dates (13 –30 ka) well within the Wisconsinan and pollen analyses indicated spruce forest (Sirkin et al. 1970; Denny et al. 1979). As for southern New Jersey, the upland core of Delmarva is covered with a ubiquitous blanket of thin surficial deposits that are not generally being eroded or redeposited today. New LiDAR base coverage with a Digital Elevation Model (DEM) of 20 –25 cm contour intervals now permits further re-interpretation of the geomorphology of Delmarva. For example, the area around Snow Hill, MD is shown in Figure 6 (Newell & Clark 2008). Elliptical to round to irregular, rimmed, flat-floored basins are shown on old estuarine terrace surfaces. The pattern is one of spatially dense forms with new basin rims overprinted upon older ones. The Parsonsburg dune fields are revealed as stacks of imbricated, prograding sand sheets and dunes with a NW to SE direction of transport. In the NW quarter of the image, SE-advancing sand sheets have overrun the closed basins. A preliminary interpretation is that the basins were shallow ephemeral ponds perched on silt and clay estuarine terrace

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deposits. Tundra ponds and thaw lakes that characterize regions such as the North Slope of Alaska and the lowlands of the Canadian Arctic are suggestive modern analogues. Freezing and thawing of the basins on Delmarva was likely, similar to the spungs and interdune basins of New Jersey (French & Demitroff 2001). However, no evidence of patterned ground or deformation of the basin bottom sediments has been observed to indicate scattered permafrost at this latitude (near the southern limit). The dune fields migrated during and after the formation of these ephemeral basins. All of these details are contained within the stratigraphic domain of the Parsonsburg Formation. A second illustration of surficial deposits and geomorphology from LIDAR coverage is interpreted from the area of the Blackwater National Wildlife Refuge (NWR) in Maryland (Fig. 7). The immediate issue at the Blackwater area is subsidence, coastal flooding and the loss of fresh to brackish wetlands habitat for waterfowl. As shown on Figure 7 from work in progress, dune fields with strong linear trends have been truncated by rimmed flat-bottomed closed basins. Smaller, lesswell-defined parabolic dunes are overprinted across the basin bottoms and a large-scale fluvial channel meanders across a low-gradient surface. All these features occur on a terrace generally less than 1 m above sea level. The northern portion of the DEM is a more-featureless upland surface generally less than 4 m above sea level. The wetlands area trending west to east across the centre of the map is underlain by an early buried channel of the Chesapeake. Coring and augering from the northern uplands southwards across the buried valley and onto the lower terrace surface, now barely above tide level, indicates a continuous cover deposit of 1 –5 m of mixed sand and gravel. Interpreting the new geomorphic details coupled with subsurface stratigraphy (from cores), age dates and pollen spectra permits re-definition of the Parsonsburg as a classic alloformation. Optically stimulated luminescence (OSL) ages from work in progress on recent cores at Blackwater NWR have yielded an age of 43 + 3.11 ka well within MIS 3 for the beginning of the alloformation. Blanketing all, the Parsonsburg is a continuum of cold-climate landforms constrained in space and time by a basal unconformity across valley bottoms, terrace and upland deposits alike. It extends to, and includes, synchronous, glacial-outwash gravels of the Susquehanna River below the base of the modern Chesapeake deep channel. In a similar manner, the upper bounding surface is being overlain by Holocene transgressive deposits. Modern floodplain deposits, fan delta deposits, tidal marsh and littoral and thick bay mud are accumulating over the Parsonsburg alloformation.

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Fig. 6. LIDAR-based shaded relief image of part of the Snow Hill 1:24 000 scale quadrangle on the Pocomoke River valley of Maryland. Note arrays of elliptical-to-round rimmed, flat-floored basins and overprint of SE-advancing dune field (map from Plate II of Newell & Clark 2008).

The Parsonsburg alloformation north of the Choptank River Ephemeral upland excavations (Fig. 8) indicate that early slope failures from upland gravel sources have produced massive, deformed allochthonous blocks similar to active layer detachment failures as described by Lewkowicz & Harris (2005). A date of 25.3 ka from peat below an allochthonous block is congruent with the range of dates for the Parsonsburg reported by Denny et al. (1979). These slope failures are covered by thick accumulations of later slope deposits produced by sheet flows, debris flows and windblown coversand. In most cases, the overprint of these later cover sequences has been so complete that no surface expression of earlier mass wasting of slopes remains; everywhere

simple catenary surfaces descend from broad hill crests to valley bottoms with small, entrenched ephemeral stream channels. Nearby upland excavations at about 20 m elevation provide a shallow but broad exposure that shows the typical source and facies relationships of the cover sequence (Fig. 9). The excavation barely penetrates the marine sediments of the Chesapeake Group (Calvert Formation). However, overlying gravel of the Pensauken Formation is thin, deeply eroded and weathered. It occurs as an erosional remnant that has been the source of both sorted and diamictic sediments downslope to the north. Windblown deposits occur above and below the gravel and organic-rich sediments are preserved in a former blowout depression near the top of the pre-excavation surface. This sequence is similar

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Fig. 7. LIDAR-based geomorphology of the Blackwater NWR area (includes Golden Hill and Taylors Island 1:24 000 scale quadrangles). ‘D’, cross-cutting relationships of dune fields; ‘B’, flat-floored ephemeral ponds; ‘P’, point bars of a meandering river channel provide a sequence stratigraphy of depositional and erosional events. ‘H’ is Holocene Peat and alluvium; ‘S’ is sand plain; and ‘PU’ is Parsonsburg Formation undifferentiated. (Map courtesy of DeJong).

to, but not as complex as, the New Jersey exposure illustrated in Figure 5. Macro-topography suggestive of niveo-aeolian features is also common on the uplands. An example is shown in Figure 10. These features can now be defined on a detailed DEM and can be distinguished from the more general fabric of subtle dunes and interdune basins. On Figure 10, irregular hollows are distributed around the upland perimeters. Commonly, circular, elliptical or irregular flat-floored hollows are inset below hill crests. They range in size from 30 –300 m in diameter. Confining ramparts are generally less than 1.5 – 2.0 m high with less than 2% gradients. Drainages from these features mark the point of headward erosion for small, entrenched, ephemeral channels

and the flat bottoms are underlain by silty, massive- to thin-bedded fine sand and organic material. Collectively, the morphology, sediments and topographic setting suggest that these hollows were modified by repeating seasonal or perennial accumulations and melt-out of snow that included dustings of interbedded windblown sediment. The spatial extent of these landforms and sediments suggests that a belt of heavy perennial snow cover existed late in the assemblage of the Parsonsburg alloformation.

Southern Maryland The upland counties of Maryland west of the Chesapeake Bay and south of Washington, DC are known

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Fig. 8. (a, b) Parsonsburg Formation at the Ewing Pit includes a slump block detached from Pensauken gravel and covered by subsequent sand, gravel and windblown deposits. Peat bed under slump block carbon dated at c. 25.3 ka includes spruce, pine and birch pollen indicating boreal forest. Photo (8b) is an orthogonal view of slump block showing arched bedding and weathering horizons cut by subsequent sand-filled cracks (at hands of investigator). Windblown cap has been excavated.

geographically as Southern Maryland, a region of former intense erosion from tobacco farming. Early studies by Hack (1955) and Nikiforoff (1955) in southern Maryland indicated previous periglacial conditions and French et al. (2009) suggest the possibility that extensive fragipan soil horizons were

also of periglacial origin. In addition to Hack’s report on Southern Maryland, particularly the Brandywine area, Glaser (1971, 1995) noted that it was difficult to resolve anticipated terrace sequences with exposures of estuarine sediments because of subsequent erosion and burial by slope and wind

Fig. 9. Cross-section of surficial deposits at the Dependable Materials pit on Delmarva uplands, north of the Choptank River. Surficial stratigraphy similar to the Haines Corner exposure (Fig. 5a, b). This excavation shows Calvert Formation, ‘Tc’; Pensauken Formation, ‘Tps’; weathered slope gravel, ‘g’, eroded from the Pensauken Formation; interbedded windblown coversand, ‘cs’; and a small remnant of a surface depression with humic sand, ‘hs’.

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Fig. 10. Flat-floored hollows possibly filled with niveo-aeolian sediment (grey) embedded in upland margins at the heads of first-order drainages on Delmarva. Arrows indicate direction of discharge from each hollow to ephemeral streams.

deposits. Thick colluvium and windblown deposits have concealed terraces and deposits in many tributary valleys. Old reconnaissance notes and photographs provide a valuable framework for understanding the internal stratigraphy of the Parsonsburg alloformation above tidewater in the Patuxent River Valley. A composite view of valley-side slope deposits is illustrated schematically in Figure 11, revealing the dilemma of locally buried Pleistocene terraces. The profile of slope and terrace deposits from the Patuxent Valley (Fig. 11) was compiled (by W. Newell) during the summer of 1997 and the exposure is now completely filled with construction and demolition waste. As exposed in 1997, the south wall of the section presented as much as 15 m of windblown sand and slope deposits punctuated by a sequence of humic layers. The surficial materials

were discontinuous and rested unconformably on old, weathered estuarine terrace deposits. In turn, the estuarine terrace deposits rested unconformably on underlying Eocene marine beds. The adjacent topography showed little or no expression of the buried estuarine terrace. The altitude and weathering of the terrace deposits suggest that it may be a remnant of one of the more dominant high sea-level stands; possibly MIS 11 or 7 as interpreted on the Virginia side of the Potomac estuary (Fig. 3). In Figure 11, the details of the toe of the slope are from a former nearby exposure. Directly west of both exposure locations, the Patuxent River floodplain is above tide and fluvial. The gravel pits visited during 1997 in the low terrace deposits exposed a complex fluvial sequence topped by a 2 m thick peat accumulation. Superposed above the peat is an alternating sequence of decimetre-thick sand beds that include planar and discontinuous

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Fig. 11. Cold-climate deposits on Patuxent Valley upland slopes. Generic deposits shown are identified as: ‘PS’, humate horizons; ‘E’, aeolian sand; ‘D’, diamict gravel; ‘NEA’, niveo-aeolian alluvium; ‘P’, peat and peaty sand (back swamp or abandoned channel); ‘A2’, low discharge alluvial braid plain sequence; ‘A1’, high discharge alluvial braid plain sequence; ‘TES’, tidal estuary sequence; ‘N’, Eocene Nanjemoy Formation. Patuxent River is at grade above tide and locally forms a braid plain on the Holocene alluvial fill.

cross-beds; these may replicate patterns of niveoaeolian discharge (Vandenberghe 2001, 2008; French 2007). A humic horizon, a diamictic gravel deposit and a thin cover of windblown deposits cap the terrace sequence. Typically, the complex slope and windblown sequence on upper Patuxent valley flanks in Southern Maryland attenuates downslope. It is truncated or buried by younger slope deposits that prograde over the lowest terrace sequence. The entire exposed sequence in the lower terrace deposits can be interpreted as subaerial, fluvial deposits capped by slope and windblown sediments. Part or all of this sequence may have cold-climate affinities that correlate with MIS 3. Local remnants of the same low terraces replicate the downvalley gradient of the modern Patuxent floodplain.

Local geomorphic response to glacial forebulge dynamics During the later cold stage of each Pleistocene climatic cycle, sea level was much lower than present. Thus, during falling-stage systems, low-stand systems and the early part of rising-stage systems, stream gradients were steeper, promoting periods of rapid headward erosion as channels adjusted to

new equilibrium profiles. The efficiency of these adjustments can be evaluated by examining the advance of young drainage systems eroding into terraces that had been deposited as bay bottom muds and littoral sands during a preceding sea-level highstand (Colman 1983). In Virginia, the distribution of Plio-Pleistocene fluvial, estuarine and marine terrace deposits and of Miocene marine littoral and fluvial deposits (Mixon et al. 1989, 2000) indicates a systematic, cumulative increase in altitude with increasing age through a range of 0–100 m. Newell (1985) interpreted these data to indicate that the elevation of more recent highstand systems tracts are eustatic artefacts. Highstand indicators of increasing age exceed the range of eustatic effects and their systematically increasing elevations are evidence of cumulative tectonic uplift. Thus, increasing age and cumulative uplift provide continuing potential energy for erosion and sediment transport into tributaries of the Chesapeake Bay. This scenario is not replicated everywhere around the Bay. As noted, thick weathered estuarine terrace deposits along the Potomac estuary in Virginia occur at 25 m and have been interpreted to be c. 450–500 ka (Mixon et al. 1989; Newell et al. 2005). The northern uplands of Delmarva occur at about the same elevation but are underlain by

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upper Miocene– Pliocene sediments that are fluvial weathered gravels. Similar deposits in Virginia occur at 60 –80 m elevation. Clearly, Delmarva has been subsiding. Recent evidence and modelling of inundation at the Blackwater NWR on Delmarva (Kearney et al. 2002; Larsen et al. 2004) used an average rate of 3 mm a21 which is within a 1– 4 mm a21 range for the Chesapeake Bay recently documented by Barbosa & Silva (2009). Low terraces in the 3–5 m elevation range around the mid-Bay area have been generally interpreted as artefacts of the last major sea-level highstand, MIS 5e. During original mapping of the Kent Island estuarine terraces on the eastern shore counties of the mid-Chesapeake Bay region, Owens & Denny (1979) collected numerous wood and peat samples that gave ages clustered in the range of 28– 30 ka. The dates were widely discounted because they were not synchronous with sea-level highstands as then described; the flora and fauna indicted a cool climate, however. Finkelstein & Kearney (1988) produced similar dates from Atlantic-facing terraces of Virginia. These data also correlate with MIS 3, a highstand commonly considered to have occurred below modern sea level (Wehmiller et al. 2004). The emerging uncertainties of terrace deposit ages and elevations that are incongruent with global sea-level records require a re-evaluation of the possible effects of a dilating and collapsing glacial forebulge (Potter & Lambeck 2003; Wehmiller et al. 2004; Scott et al. 2010). If the mid-Chesapeake area was uplifted during the LGM and is now collapsing, such phenomena as seemingly anomalous terrace ages and elevations followed by locally anomalous rates of sea-level rise may be explained (Scott et al. 2010). The net effect is that cold-climate processes may have been enhanced by higher relief of currently unknown dimensions. Even the domain of anomalies that could indicate a forebulge is not certain. However, the concept of a rising forebulge may explain locally accelerated erosion and the unique distribution of complex sequences of surficial deposits. Subsequently, local apparent sea-level rise rates may be suggestive of forebulge collapse for these same unique areas.

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inactive landforms include dune fields and rimmed flat-floored basins. Common fluvial parameters of large-scale terrace remnants are beyond the competence of the magnitude and frequency of modern geomorphic processes. Sediments within the alloformation are both near- and far-sourced. Interbedded deposits commonly include sand and gravel reworked and deposited from upslope sources mixed with fine-grained sand and local loess deposits blown back uphill from valley bottom deposits. Unweathered sediments derived from terraces and floodplains of outwash-bearing rivers, such as the Delaware and Susquehanna, are locally mixed with deeply weathered, polycyclic mature sediments from nearby Coastal Plain uplands. Homogenized surficial horizons may all appear massive and indistinguishable under different morphosequences. Coring and excavations at numerous localities indicate that the Parsonsburg alloformation is continuous beneath subaerial fill (Holocene peat and muck) and under Holocene bay mud under the Chesapeake Bay (Cronin 2000). Sub-bottom profiling indicates a possibly similar history in Delaware Bay (Knebel & Cerce 1988). OSL dates from work in progress indicate the beginning of the alloformation was within MIS 3, which correlates well with evidence of Markewich et al. (2009) that high volumes of aeolian sand were transported in the Eastern US during MIS 3 and continuing into MIS 2 through the Younger Dryas early in the Holocene. As described in southern Maryland, the recognition of periglacial features and the mappability of this evidence within a regional alloformation require a re-examination of local Pleistocene terrace stratigraphy. LIDAR images indicate that slopes between marine terraces generally do not reveal uniform strand lines of falling sea level with littoral sediments. Rather, they are covered by complex patterns of slope and dune deposits that suggest morphologies of sheetwash, slump, debris flow and windblown coversands. The assemblages of these materials are random from place to place on any slope. Similarly, accordant terrace top elevations cannot be assumed from place to place because original horizontal surfaces may have been eroded, deformed or altered by inset younger deposits and/or overprinted by coversands.

Discussion The nearly continuous sandy veneer of surficial deposits (coversand) that occur north of Tidewater Virginia, across Southern Maryland, Delmarva and New Jersey includes all manner of cold-climate morphosequences that are mappable as a classic alloformation. The morphosequences record erosion, transportation and deposition of surficial sediments on low-angle to nearly flat slopes. Presently

Conclusions A sequence of Plio-Pleistocene and Late Pleistocene events can be interpreted from the surficial geology of the Mid-Atlantic Coastal Plain. In Virginia, a continuum of slope and floodplain deposits through warm- and cold-climate cycles appear to have been unmodified by distinctive periglacial features.

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To the north, early upland deposits in central and southern New Jersey indicate a multistoried array of erosional and depositional events. The weathering and recycling suggests that the earliest effects of this degradation and storage of sediment extend back into the Plio-Pleistocene. These erosional and depositional events are thought to be most effective during climate degradation leading to cold-climate, low-stand systems tracts. Subsequent extensive surficial deposits covered and modified the landscape during thaw degradation of past permafrost. In Maryland, most of the Parsonsburg alloformation is an artefact of climate degrading to the LGM (MIS 3– 2) followed by morphosequences developed in part on local permafrost and deep seasonal frost. The fragipans of southern Maryland would have been enhanced and deformed at this time. Morphosequences at the end of the LGM represent thawing and continuation of valley erosion and deposition of windblown sediments, even on presently submerged slopes before sea-level rise. Aeolian processes and ephemeral basin freezing may have been active during the Younger Dryas followed by warming and local apparent sea-level rise over the last 7500þ years. The niveo-aeolian deposits and landforms that cap the Delmarva uplands could have been active during the end of the Wisconsinan as well as the Younger Dryas. Across the mid-Chesapeake area, a glacial forebulge of presently indeterminate dimensions may have changed local gradients and rates of erosion during isostatic response to Pleistocene glacial oscillations. Subsequent Holocene collapse may explain the local apparent acceleration of sea-level rise rates. We thank H. French and M. Demitroff for stimulating and far-ranging discussions in the field. We thank B. Landacre, J. McGeehin and T. Rittenour for laboratory support in producing pollen analyses, radiocarbon ages and OSL ages, respectively. The manuscript has been critically reviewed by S. Stanford of the New Jersey Geological Survey and M. Demitroff at the Department of Geography, University of Delaware. We thank them for their perspectives and suggestions which have helped us to focus this overview.

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the Delaware Valley, New Jersey and Pennsylvania. Geomorphology, 4, 319–345. Russell, E. W. B. & Stanford, S. D. 2000. Late-Glacial environmental changes south of the Wisconsinan terminal moraine in the Eastern United States. Quaternary Research, 53, 105– 113. Salisbury, R. D. & Knapp, G. N. 1917. The Quaternary formations of southern New Jersey. New Jersey Geological Survey Final Report, 8. Schlee, J. 1957. Upland gravels of Southern Maryland. Geological Society of America, Bulletin, 68, 1371– 1410. Scott, T. W., Swift, D. J. P., Whittecar, G. R. & Brook, G. A. 2010. Glacioisostatic influences on Virginia’s Late Pleistocene coastal plain deposits. Geomorphology, 116, 175 –188. Sirkin, L. A., Owens, J. P., Minard, J. P. & Rubin, M. 1970. Palynology of some Upper Quaternary peat samples from the New Jersey Coastal Plain. U.S. Geological Survey Professional Paper 700-D, 77–87. Smoot, J. P., Newell, W. L. & DeJong, B. D. 2009. Investigation into the origin and character of surficial sedimentary deposits at the Midshore Regional Solid Waste Facility, near Easton, Maryland. U.S. Geological Survey, Open File Report 2009–1002. Stanford, S. D. 2010. Onshore record of Hudson River drainage to the continental shelf from late Miocene through the late Wisconsinan deglaciation, USA: synthesis and revision. Boreas, 39, 1– 17. Thom, B. G. 1970. Carolina Bays in Horry and Marion Counties, South Carolina. Geological Society of America Bulletin, 81, 783– 814. Vandenberghe, J. 2001. A typology of Pleistocene coldbased rivers. Quaternary International, 79, 111– 121. Vandenberghe, J. 2008. The fluvial cycle at cold-warmcold transitions in lowland regions: a refinement of theory. Geomorphology, 98, 275– 284. Washburn, A. L. 1980. Geocryology, a Survey of Periglacial Processes and Environments. John Wiley & Sons, New York. Wehmiller, J. F., Simmons, K. R. et al. 2004. Uraniumseries coral ages from the US Atlantic coastal plain-the ‘80 ka problem’ revisited. Quaternary International, 120, 3 –14. Wolfe, P. E. 1953. Periglacial frost–thaw basins in New Jersey. Journal of Geology, 61, 133–141. Wolfe, P. E. 1956. Pleistocene periglacial frost– thaw phenomena on the New Jersey Coastal Plain. Transactions, New York Academy of Sciences, 18, 507–515. Wolfe, P. 1977. The Geology and Landscapes of New Jersey. Crane, Russak and Company, Inc., New York.

Index Page numbers in italic refer to Figures. Page numbers in bold refer to Tables. acidification 187, 188, 190 aeolian processes and products 207–208, 209, 220 –221 cool lowland deserts 241–256 ancient deposits 254– 255 depositional features 246–254 bedforms 247 –249 internal structures 250– 251 interdunes and interdraa 251, 253 post-depositional features 254 sand lamination 246–247 erosional features 244–246 evidence of temperature 254 particles 243–244 Agassiz, Louis, glacial hypothesis 2 Aker moraine line 53 allostratigraphy, marine moraines 59– 63, 70 and bathymetry 66, 68, 69, 71 comparison 64, 71– 72 Storsand moraine 35–45, 50, 64 alongshore drift, Kuray– Chuja palaeolake 116– 118 Alps, Swiss see Gornergletscher Altai Mountains, palaeolakes 111 –126 andisols 190 Antarctica lichens, role in granite weathering 195– 201 loss of ice mass 57 permafrost landforms 168 Arctic, Cenozoic glaciations 16–17 ˚ s-Ski moraine complex 34, 35, 48, 53 A avalanches, snow 219–220 Baltimore Canyon Trough 261 Bani Khatmah dropstones 146, 150 periglacial sediment 133, 135, 145 deformation 146 Banks Island, periglacial system 6 –8, 9, 10 barchans 249, 251, 252 Barents Sea ice sheet see Svalbard-Barents Sea ice sheet barrier beaches, Kuray–Chuja palaeolake 114 Barxacoba Valley 95, 96, 98, 106 basin-fill, glaciomarine 72– 73 bathymetry control on moraine allostratigraphy 69, 71 intermediate and deep water 66, 68, 70, 71 shallow water 66, 68, 69, 70 shelf, Svalbard 17 Bear Island trough-mouth fan 17, 18 Beaverdam Formation 266 bedforms, aeolian 247 –249, 250 –253 bedrock properties, and slope deposits 215 Bełchatow opencast mine, frost cracking 162 Bibei Valley 95, 96, 97, 101, 106 bioprotection, lichens 195 Blackwater National Wildlife Refuge, dune fields 267, 269 Blue Ridge Mountains 261 Bolterdalen moraine, allostratigraphy 64

boulder halo, Mars 169 boulder pavements, Juwayl Formation 146, 149, 150 bounding surfaces, aeolian bedforms 251, 252 burrows and traces 254 Calvert Formation 268 calving Storsand moraine 47, 49, 50, 51 tidewater glaciers 63, 71 Cenozoic, glaciation, Arctic 16– 17 Chagan Uzun Interbasinal Uplift 112 channels ice-marginal 4, 5 Banks Island 6, 10 Storsand moraine 39, 42, 45 Sanamah Formation 133, 134 see also tunnel valleys Charles City Formation 267 Chesapeake Bay 261 geomorphology, glacial forebulge 272–273, 274 Chesapeake Group 261 Chuja Basin, strandlines 112 –126, 113 clast flows, frost-coated 217– 218 climate New Jersey coastal plain 265 and pedogenesis 186 periglacial 205, 206 climbing ripple lamination 246 clinothems 61, 62, 65 Storsand moraine 37, 39, 43, 45 coefficient of cryogenic contrast 156, 157, 160 colluvium see slope deposits Columbia Group 266 Comfortlessbreen surging glacier 25 composite-wedge polygons, Mars 174–175, 177 conglomerate Sanamah and Juwayl formations 135 facies association 141, 142, 144 coversands 207, 209 Delmarva Peninsula 266– 267, 273 Netherlands 248 crag and tail features, Svalbard 18, 19 cross-bedding, aeolian 250–251 deformed 254 CRREL permafrost tunnel, Alaska 157, 158 cryodesiccation 207 cryopediments 206 cryopedoturbation 186, 190 cryosphere 1, 2 cryostratigraphy 153 cryostructures 155, 156 cryotic environment, Mars 3 –4 cryoturbation 158, 210 cryptogams 186, 187, 190 Antarctica 196–201 debrisfall 67 Storsand moraine 37, 39, 40, 42, 44, 47

278 debrisflow 67, 208, 218– 219, 220, 222 cohesionless see grain flow Gornergletscher 87, 88 Storsand moraine 37, 39, 40, 42, 44, 47 Svalbard 26, 28 Takahashi model 218–219, 220 deformation glacial 5, 6, 27 Pias area 102–103, 104 post-depositional, Wajid Group 148 soft sediment 20, 47 glacier push, Wajid Group 148 subglacial deforming-bed 33, 58 grounding-line behaviour records 67–69 Storsand 51, 52 deglaciation, Svalbard 18, 19, 20, 22 Delmarva Peninsula, coversands 266– 267, 273 delta, Gilbert-type 35, 59, 60, 62, 67, 71 Pias area 101, 105 deserts, cool lowland aeolian processes and products 241–256 particles 243– 244 detachment slides 206 diamictite Juwayl Formation 136, 137 facies association 142, 144–145 diamicton 5, 60, 61, 67 matrix-supported, Pias 97, 99, 100, 105 origin 4 Storsand moraine 37–40, 45, 47, 49, 50– 53 Svalbard 17, 20, 27 Dibsiyah Formation 130, 131, 133, 134 dilatancy 217 distortion, horizon 186, 187 draa 249, 251 dropstones 5, 7 Juwayl Formation 136, 137, 138, 146, 150 Kuray and Chuja basins 116 Svalbard 23 drumlins, origin 4 Dryas, Younger, Scandinavia 34–35, 45, 48, 53 dunes aeolian 207, 208, 243, 248– 249, 250 –253 Delmarva Peninsula 267 Mars 243, 248, 249, 255 types 248–249, 251 dust, Mars 171–172 Earth arid zones 242 climatic variation 3 e´boulis ordonne´s 213–214 Egge-Lynga˚s moraine, allostratigraphy 64 entisols 190 erosion subglacial 65, 67 wind 244 –246 erratics, Svalbard 21 eskers, Svalbard 26 fan, submarine 35, 58, 59, 60, 61 crest 60, 61, 65, 67 Storsand moraine 45, 46, 50 grounding-line behaviour records 65, 67

INDEX Storsand moraine 37, 39, 42, 45–47, 50, 51 Fard al Ban, Juwayl Formation, ‘Sorbet’ facies 136, 140 faults, Pias area 103 Fennoscandian Ice Sheet 34, 35, 45 fjords Spitzbergen 15 Svalbard 17, 18, 19, 20, 25 flatiron forms 4 fluid flow, overland see overland flow fluting glacial 6 Sanamah Formation 146, 147 Svalbard 27 wind erosion 244–245 fluvial deposits 209 fluvial processes 206 folds, Pias area 102 forebulge, glacial, Chesapeake Bay 272–273, 274 foresets fan 61, 62, 65, 67 Storsand moraine 36–39, 42– 47 fragipan 155, 164, 207 Southern Maryland 270, 274 freeze–thaw weathering 156, 183–184, 186, 206– 207 Mars 172 stratified slope deposits 215– 216 freezing epigenetic 153, 154, 154, 155 seasonal 153 syngenetic 153, 154, 155, 157 see also sediment, frozen frost cracking 155, 161–164 frost creep 208, 209 stratified slope deposits 215– 216 frost shattering 190, 206, 209 see also gelifraction funnel-shaped structures, Pias area 103, 104 Galicia geology and geomorphology 95, 96 Upper Pleistocene climatic overview 106–107 glacial model 105– 106 valley-junction sediments 93–108, 94 Gaspe´sie, stratified slope deposits 216, 218, 220–221 Geikie, Archibald, glacial studies 2 Geitryggen–Ski moraine complex 34, 35 gelifluction 160, 206, 208, 209 Banks Island 7, 8 gelifraction 214 –215 see also frost shattering gelisols 190 geomorphology, periglacial 206 glaciers early hypotheses 1 –2 outlet 33 Oslofjorden 35–54 polythermal, Svalbard 25, 26–27 sediments 4– 5 surging, Svalbard 15, 25, 25, 27– 28 systems 4– 5, 6, 7 tidewater grounding line sedimentation

INDEX and ancient kinematics 57–73, 58 Storsand moraine 33–54 ice-flux 51–52 Spitzbergen 58 Svalbard 25– 26 glendonites 254 gleysols 207 Gondwana, Palaeozoic glaciation 28, 129, 130, 149 –150 Gornergletscher 77, 78 lateral moraines 77– 90 formation 88–89 geomorphology and sedimentology 79, 80–81, 85– 86 GPR 79–80, 81– 84, 86–88 ice-marginal profile 84– 85 grainfall lamination 247, 250, 252 grainflow 208, 217– 218, 220 granite, lichens and weathering, Antarctica 196 –201 gravel massive sandy, Pias 97–98, 99, 100, 105 planar cross-bedded, Pias 99, 100–101, 103 planar-bedded, Pias 99, 101, 105 pocket beach crestline model, Kuray-Chuja palaeolake 122–123 trough cross-bedded, Pias 98, 99, 100, 105 gravels, upland Mid-Atlantic Coastal Plain 261, 262 Delmarva Peninsula 266 weathering 262 greenhouse periods 3 Greenland, loss of ice mass 57 gre`zes lite´es 208, 213– 214, 220, 221–222, 223 ground veins 161 grounding line ‘problem’ 33 sediment accumulation 33, 58, 63 Storsand moraine 39, 47, 49– 50 grounding line behaviour moraine record 65–69, 66, 70 intermediate and deep water 66, 68, 70, 71 shallow water 68, 69, 70 grounding line position 63, 66 maximum 58, 60, 63 Storsand moraine 47, 49– 50 grounding-zone wedges, Svalbard 18, 19, 20 gullies, Mars 169, 171, 173, 175 –178

pore 153, 155 within frozen sediments 153–154 ice crust 155 ice discontinuities 157– 158 ice flux Storsand moraine 47, 51– 52 tidewater glaciers 63, 67 ice lenses 155, 156, 190, 210 ice segregation 153– 154, 155, 190, 206 ice streams 33 Oslofjord trough 35 Svalbard, landforms 17, 18, 19, 20 ice veins 154, 156, 157, 190 ice-contact zone intermediate and deep water 66, 68, 70, 71 morphodynamic evolution 59–63, 66, 67, 68 Storsand moraine 35– 39, 41, 43, 50, 51 shallow water 66, 68, 69, 70 ice-front advance 58, 61 Storsand moraine 47 ice-front position 63 ice-front retreat 59, 60, 62 ice-front stillstand 58, 59, 60–62, 65 Storsand moraine 47, 49– 50 ice-house periods 3 ice-margin ancient, kinematics 57– 73 Gornergletscher, GPR 84– 85 grounding zone, Svalbard 18, 19, 20 retreat, Svalbard 18, 19 ice-rafted debris 65 ice-wedge casts 210 ice-wedge polygons 161, 162, 163, 190 Mars 172, 174– 175, 177 icebergs, overturned 41 see also calving Iceland, frost cracking 163, 164 icing deposits 206, 209 impact crater, Mars 169, 171 inceptisols 190 interdraa 251, 253 interdunes 251, 253 involution structures, Pias area 102 –103 Isfjorden lateral moraine ridges 20 seismic section 20 tidewater glacier 25

Haines Corner, cold climate slope deposits 266 Hirnantian glaciation 149 histosol 190, 230, 231, 232 Holocene slope dynamics, Sierra Nevada, Spain 227–237 Svalbard 22 Holocene Optimum Warm Period 233 hummocks, Svalbard 18, 19, 26, 27 hyperpycnite 65

Jibal Al Qahr, channels 133 Juwayl Formation 130, 131, 132, 133, 135 age 135 boulder pavements 146, 149 deformation, soft-sediment 148 dropstones 146, 149, 150 folds and faults 146 lithofacies 135–139 facies associations 139– 145 striations 146, 147 thickness 135 tunnel valleys 148

ice buried glacier 154 Mars 171– 172 needle 155 pingo 154 pool 154

Kapp Ekholm lithostratigraphy 21, 22 sediments 23

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Ka¨rkevagge 183 alpine pedogenesis 187, 190 chemical weathering 185, 188 –189, 190 horizon distortion 186 Khusayyayn Formation 130– 131, 133 Kongsfjorden lateral moraine ridges 20 overturned icebergs 41 Kongsøya, lithostratigraphy 21 Kongsvegen tidewater glacier 25 Kregnes moraine, allostratigraphy 64 Kuray Basin, strandlines 112– 126, 113 Kuray– Chuja palaeolake strandlines 111– 126 characteristics 114– 116 classification 116– 118 drop stones 116 pocket beach crestline model 122–123 stratigraphy and sedimentology 118– 120 wind-wave climate 115–116, 120, 121– 122 lag deposits 245, 246, 251 Lake Agassiz, aeolian processes 243 lake deposits, Holocene, Sierra Nevada, Spain 227– 237 Lake Melville fjord basin 72– 73 glacier 73 Moraine 3/4 64, 68, 72 Moraine 5 64, 72 Moraine 7 64, 68, 73 lakes glacial see Kuray– Chuja palaeolake; palaeolakes proglacial, Banks Island 6 lamination aeolian sand 246– 247, 248 climbing ripple 246–247, 248 crinkly parallel 247, 248 grainfall 247, 249, 250, 252 planebed 246, 252 sandflow 247 landforms, periglacial 206 Svalbard inter-ice stream 17–18, 19 marine archives 17–20 terrestrial records 20– 22 Lanzhou, China, loess 244 Lassen Volcanic National Park, stratified slope deposits 223 Last Glacial Maximum Galicia 105 Svalbard 17, 19, 21, 22 latitude-dependent mantle, Mars 169, 170, 171– 172, 178 Latnjavagge, chemical weathering 185 Laurentide Ice Sheet 73, 259, 261 Lecidea cancriformis 196, 198, 199, 200– 201 lichens Antarctica, role in granite weathering 195– 201 and weathering 186, 189 lineations, mega-scale, Svalbard 18, 19, 20, 25, 26 Linne´elva, lithostratigraphy 21, 22 lithostratigraphy, Svalbard 21–22 lits gras 213, 214 lits maigres 213, 214 Little Ice Age

Sierra Nevada, Spain 228, 231, 233, 235 Svalbard 18, 19, 22 lobes solifluction, Sierra Nevada, Spain 229, 230–233, 234–235, 236–237 stone-banked 215– 216 loess 208, 243, 244, 253 lonestones see dropstones Łozinski, Walery von, periglacial concept 2, 5 Mackenzie Delta, heavy mineral assemblages 158 mapping, allostratigraphic 59– 63, 64 Storsand moraine 35–45, 64 marine terraces, Mid-Atlantic Coastal Plain 261–262, 265 Mars aeolian deposits 255 sand 244, 246, 255 dunes 243, 248, 249, 255 wind ripples 248 yardangs 247 dust 171– 172 environment 3 –4 ice volume 171–172 north polar ergs 242 permafrost 167–178 landforms 168, 169–178 gullies 169, 171, 173, 175–178 impact craters 169, 170, 171 latitude-dependent mantle 169–172, 178 polygons 169–172, 174–175, 177 sedimentary processes 169 local-scale, gullies 173, 175 –178 regional-scale, polygons 172, 174– 175, 177 Maryland Bays 267 Mediaeval Warm Period, Sierra Nevada, Spain 230, 233 meltwater outflow 33, 58, 63 Storsand moraine 37, 52– 53 meltwater underflow 61 Storsand moraine 37, 39, 41, 47, 52 Merrimelia Formation, aeolian deposits 254 micro-jointing, prismatic 207 Mid-Atlantic Coastal Plain cold-climate slope deposits 259– 274 landforms New Jersey 263– 266 Virginia 262– 263 regional geological setting 260–262 weathering 262 minerals clay and CCC values 160–161 Mid-Atlantic coastal plain 262 periglacial pedogenesis 187 heavy, cryostratigraphy 158, 159 mollisols 190 Mona moraine 48, 53 allostratigraphy 64 moraines aggradation 65 arcuate, Svalbard 18, 19, 27 crest 60 sediment records 69 emergence 59, 62

INDEX grounding line, Storsand 34–54 lateral Bibei Valley 95, 96, 97 Gornergletscher 77–90, 79 formation 88– 89 GPR 79–80, 81– 85, 86–88 sedimentology 79, 80– 81, 85–86 Svalbard 18, 19, 20 marine 58–59 allostratigraphy 59–63, 64, 71– 72 bathymetric control 66, 68, 69, 71 comparison 64, 71– 72 Storsand moraine 35–45, 50, 64 as record of glacier dynamics 63– 69 grounding line behaviour 65– 69, 70 monoepisodic 63, 65, 69 polyepisodic 60, 63, 67, 71 retreat, Svalbard 18, 19, 20, 26 Svalbard 18, 19, 20, 26–28 terminal, Svalbard 18, 26 transverse, Svalbard 18, 19, 20 morphosequences 259 New Jersey coastal plain 264–266, 273 Mount Keinath, tafone, lichen colonization 196, 198 –201 mud apron, Svalbard 28 mud drapes 62 Storsand moraine 37, 42, 47, 51 Murchisonfjorden, lithostratigraphy 21, 22 nappe coule´e 215 needle-ice 155 New Jersey coastal plain climate 265 landscapes 263– 266 morphosequences 264–266, 273, 274 nitrogen fixation 187, 190 niveo-aeolian sands 207–208, 221 Delmarva Peninsula 269, 271, 274 nunataks, Svalbard 21 Oasi, tafone, lichen colonization 196 Øgdern ice-drainage shelf trough 48 olistoliths, diamicton, Storsand moraine 37, 38, 40, 45, 47, 50–52 Ollo de Sapo orthogneiss 95, 96 Ordovician, Late, glacial sediments Saudi Arabia 129 –130, 149, 150 glacial model 143, 148 Sanamah Formation 132–133 organisms burrows and traces 254 role in pedogenesis 187, 190, 191 Oslofjord trough 34– 35, 45 Oslofjorden, tidewater glacier, Storsand moraine 34– 54, 48 overland flow 220 Øyeren ice-drainage shelf trough 48 palaeoenvironments 3, 4 palaeolakes Saudi Arabia 135 strandlines, Altai Mountains 111– 126 palaeopermafrost table, fragipans 164

palaeosols 254 palaeothaw 158, 159 Palaeozoic glacial deposits, Saudi Arabia 129– 150 glaciation, Gondwana 149– 150 pans 245 paraglacial, terminology 213 Parsonsburg alloformation dune fields 267 north of Choptank River 268– 269, 270, 273, 274 particles, sedimentary cold climate 243– 244 transport and deposition 243– 244 Patuxent River Valley, slope deposits 271– 272 Pedersenbreen polythermal glacier 25 pedogenesis, periglacial 207 alpine climate 186 organisms 187 parent materials 188 and permafrost 190 relief 187 Scandinavia 186– 188 time 186– 187 Sierra Nevada, Spain 230– 233 Pensauken Formation 266, 268 periglacial, terminology 2, 5, 183, 205, 213 periglacial environments 1, 2, 206 periglacial processes 205, 206 post-depositional 209– 210 periglacial sediments 205– 211 periglacial systems 5 –8 periglacial zone 2, 5 –6 permafrost 5, 8, 153 active layer 153, 154–155, 157, 167 Banks Island 7 CRREL permafrost tunnel, Alaska 157, 158 deformation, Svalbard 27 interface with active layer 154–155, 157 Mars 167 –178 landforms 168 and pedogenesis 190 see also palaeopermafrost; sediment, frozen Permo-Carboniferous, glaciation Saudi Arabia 129, 149, 150 glacial model 143, 148 Juwayl Formation 135 Phoenix landing site, sediments 169 Pias, Galicia landscape 95 Upper Pleistocene valley-junction sediments 93–108 climatic overview 106– 107 deformation features 102 –103, 104 glacial model 105–106 lithofacies 95, 97–102, 103, 105 Pine Barrens, NJ CCC values 160 geomorphology 264 sand wedges 162 pingos 154, 209 Kuray and Chuja basins 114 planebed lamination 246, 252 plants, and pedogenesis 187, 190

281

282

INDEX

Pleistocene glaciation Banks Island 6–8, 9, 10 Svalbard 17 ice ages 1–2 slope deposits, Mid-Atlantic Coastal Plain 259– 274 Upper, valley-junction sediments, Pias 93– 108 pocket beaches 117–118 crestline model 122–123 podzolization 187, 190 polygons thermal-contraction-crack Banks Island 8, 10 Earth 168, 177 Mars 168, 169, 170–172, 174–175, 177 Poolepynten lithostratigraphy 21, 22 sediments 23 Potomac River 261, 262 geomorphology 262 –263 Pre´alpes, stratified slope deposits 222–223 proglacial zone 5, 6 pseudolamination, adhesion ripple 247, 248 Pyrenees, stratified slope deposits 222 pyrite, oxidation, and pedogenesis 187, 188, 190 Qalibah Formation 130, 131 Qusaiba Shale 130, 133 Ra moraine line 53 radar, ground-penetrating 71, 77 Gornergletscher lateral moraines 79–80, 81– 85, 86–88 Kuray–Chuja palaeolake 114, 120, 121 sand dunes 250 raised beaches, Svalbard 22, 24 Rappahannock River, terraces 263 REFLEXW program 80 regolith, New Jersey coastal plain 264 regosol, Sierra Nevada, Spain 230, 231, 233, 235 regs 245 relief, and pedogenesis 187 rhythmite sand-silt, Pias 99, 101, 102, 103 siltstone-shale, Juwayl Formation 142, 144–145 ridges aeolian 246 crevasse-fill 26, 27 meandering, Svalbard 27 moraine linear, Svalbard 19, 20 retreat, Svalbard 26 transverse, Svalbard 18, 19, 20, 26 Rio Seco lake, sediments 229, 233– 234, 235– 236, 237 Rio Seco Valley, solifluction landforms and lake sediments 228–237 ripples, aeolian sediment 247 –248, 249 roches moutonne´es Scandinavia 185 Svalbard 21 rockfall, stratified slope deposits 216 –217 Roman Warm Period 233 runoff, surface see overland flow

Salisbury Embayment 261 saltation, sand 207, 243, 245 San Juan Valley, solifluction landforms and lake sediments 228– 237 Sanamah Formation 130, 131, 132–133, 134 age 133 deformation, soft-sediment 148 flute structures 146, 147 lithofacies 135, 139 facies associations 139– 144 striations 146, 147 thickness 133 tunnel valleys 133, 134, 148 sand aeolian 243– 244 laminae 246– 247, 249 transportation 244, 245 cross-bedded, Pias 99, 100, 102, 103, 105 laminated, Pias 99, 100, 101 –102, 103, 105 massive, Pias 99, 101 niveo-aeolian 207–208, 221 Delmarva Peninsula 269, 271, 274 ripple cross-laminated, Pias 99, 100, 102, 103 sand dunes see dunes sand sheets 251, 253 sand veins 162 sand-wedge polygons 161, 162, 163, 174–175, 177 sandflow lamination 247, 250 sandstone fine and coarse grained Sanamah and Juwayl formations 139 facies association 142, 145 massive to cross-bedded Sanamah and Juwayl formations 135, 136, 139 facies association 142, 144 –145 ripple-bedded, Juwayl Formation 136, 139 Sarah Formation 149 Saudi Arabia, Palaeozoic aeolian deposits 255 glacial depositional environments 129– 150 evidence of glaciation 145 –148 boulder pavements 146, 149, 150 dropstones 146, 150 fluted surfaces 146 folds and faults 146, 148 soft sediment deformation 148 striations 146 tunnel valleys 5, 148, 149, 150 glacial models 148 stratigraphy, Wajid Group 130– 132, 133 Dibsiyah Formation 130, 131, 133, 134 Juwayl Formation 130, 131, 132, 133, 135 Khusayyayn Formation 130–131 Qalibah Formation 130, 131 Qusaiba Shale 130, 133 Sanamah Formation 130–133, 134 Scandinavia pedogenesis 186– 188 weathering 183–186 scree, modern, sediment types 215– 222 sea-level Chesapeake Bay area 272– 273, 274 influence of glaciers 33 Oslofjorden 59

INDEX Storsand moraine 34, 45, 47 Svalbard 22, 24 sediments 4 –5 deformation 5, 6 fluvial 5 frozen 153–158 active layer 153, 154 freezing history active layer/permafrost interface 154– 155, 157 cryostructures 155, 156 ice discontinuities 157–158 weathering profile differences 155– 157 ice 153– 154 Mars 167– 178 see also permafrost; sediment, previously frozen full-glacial-mode, Svalbard 16– 22 interglacial-mode, Svalbard 16, 22–28 Mars, permafrost landforms 169– 178 periglacial 205– 211 previously frozen 158– 164 clay minerals and CCC values 160–161 subglacial 33, 67– 69 Storsand moraine 39, 50– 51 transport 4, 5 valley-junction, Upper Pleistocene 93–108 sediment cores, Svalbard 17 sedimentation grounding line 57, 58 Storsand moraine 33–54 sedimentology Gornergletscher lateral moraines 79, 80– 81, 85–86, 88– 89 Kuray–Chuja palaeolake 118– 120 seismic records, Svalbard 17, 20 shale Juwayl Formation 139 facies associations 142, 144– 145 Siberia, palaeolakes 111– 126 Sierra Nevada, Spain 228–229 Holocene slope dynamics 227–237 lake sediments 229, 233– 234, 235– 236 solifluction landforms 229, 230– 233, 234– 235, 236 –237 sieving, kinematic 217 sills, bedrock Mona Moraine 53 Storsand Moraine 34, 45, 47, 49 silt, aeolian see loess siltstone, Juwayl Formation 137, 139 facies association 142, 144– 145 sinkholes 254 Site 15, Svalbard lithostratigraphy 21, 22 sediments 23 Skardmunken moraine, allostratigraphy 64 Skilvika lithostratigraphy 21, 22 sediments 23 slope deposits 208–209 Mid-Atlantic Coastal Plain 259–274 stratified 213– 223 aeolian and niveo-aeolian deposits 220–221 bedrock properties 215 debrisflow 218–219

grainflow 217– 218 overland flow 220 relict 221–223 azonal 222– 223 gre`zes lite´es 208, 213–214, 220–223 rockfall 216–217 role of water 215 snow avalanches 219– 220 stone-banked sheets and lobes 215– 216 slopewash 220, 221 snow avalanches 219–220 Snow Hill, MD 267, 268 snowmelt 220 soil contamination 187 horizonation 187, 190 post-depositional 254 see also pedogenesis soil wedges 161, 163 solifluction 160, 186, 187, 206, 208, 209 landforms, Sierra Nevada, Spain 227– 237 stratified slope deposits 215– 216, 221 ‘Sorbet’ facies, Juwayl Formation 136, 140, 145 Southern Maryland, slope deposits 269– 272, 273 spits, accumulative strandlines 115, 116–117 Spitzbergen, glaciers 15 spodosols 190 star dunes 248, 251, 254 stone-banks 215– 216, 221 Storbreen alpine pedogenesis 186–187, 189, 190 biological weathering 186 chemical weathering 185, 189 physical weathering 184 Storfjorden trough-mouth fan 17, 18 Storsand moraine, sedimentation 34– 54, 36 allostratigraphy 35– 45, 50, 64 deforming-bed 51, 52 ice flux 51–52 ice-front latest advance 47 ice-front retreat 51 ice-front stillstand 47, 49– 50 maximum grounding line position 47, 49– 50 maximum ice-front position 49, 50 morphodynamic evolution 45, 47– 51, 49 outcrop section 1 36, 37, 38, 40, 50 outcrop section 2 36, 38, 39–41, 43, 50 outcrop section 3 36, 38, 41–43, 44 outcrop section 4 36, 38, 43–45, 46, 50 subglacial deposits and diamicton 50– 51 strandlines, Kuray–Chuja palaeolake 111– 126 characteristics 114–116 classification 116–118 accumulative 116–117 erosional 116 erosional-accumulative 116 pocket beaches 117–118 crestline model 122–123 stratigraphy and sedimentology 118– 120 wind-wave climate 121 –122 striations 4 iceberg keel, Juwayl Formation 137 Svalbard 21 Wajid Group 146, 147

283

284 STWAVE shoreline model 114 sublimation polygons, Mars 171, 174 –175, 177 Svalbard lithostratigraphy 21– 22 Quaternary glaciations 15– 28, 16 full-glacial-mode sediments and landforms 16–22 submarine 17–20 terrestrial 20– 22 interglacial-mode sediments and landforms 16, 22–28 submarine 25–26 terrestrial 26– 28 surging glaciers 15more Svalbard –Barents Sea ice sheet 16–17, 20–21, 22, 24 swales 246 systems tracts, Mid-Atlantic Coastal Plain 261– 262, 272 tafoni, Antarctica, and lichen 198 –201, 196 Takahashi model 218–219, 220 talus 215–223 Storsand moraine 42 Taurus Mountains, stratified slope deposits 222 thaw 206 evidence for 158, 160 thaw unconformities 157–158, 159 thaw-erosion 157, 158 thermokarst deposits 209 pool ice 154 till 5 origin of 4 subglacial 33, 49, 51, 67 Svalbard 21, 22, 23 time, and pedogenesis 186– 187 tombolos, Kuray– Chuja palaeolake 114, 115 transgression/regression Mid-Atlantic Coastal Plain 261– 262 Svalbard 22 tree-solution tongues 102, 103 Trevinca Mountains see Pias, Galicia trough-mouth fans, Svalbard 17, 18 tundra, gleysols 207 tunnel valleys 4, 5 Saudi Arabia 129, 133, 134, 138, 141, 148– 150 turbidites 61, 65 Storsand moraine 41–42, 44, 45, 47 Tuwayq Mountain Formation 135

INDEX Unayzah B Formation, aeolian deposits 255 uplift glacio-isostatic 22, 40, 51, 59, 60 Virginia 272 varves, Juwayl Formation 138 vegetation, Holocene, Sierra Nevada, Spain 233–234, 235, 236 ventifacts 244, 245 –246 Vestfjorden, Storsand moraine 35, 45 Virginia coastal plain landscapes 262–263, 273 systems tracts 272 Wajid Group 130–150, 131, 133, 143 water, role in slope deposits 215, 220 weathering biochemical, Scandinavia 185 biological 183 Scandinavia 186 chemical 156–157, 158, 183, 191 Mid-Atlantic Coastal Plain 262 Scandinavia 183–186, 187 Ka¨rkevagge 185, 188–189, 190 cryogenic 155– 157, 160 in situ weathered 209 freeze– thaw 156, 183 –184, 186 Mars 172 periglacial 206– 207 physical 4 Scandinavia 184 post-depositional 209– 210 tafoni 195– 201 wedges grounding-zone 59 Svalbard 18, 19, 20 Mars 168, 172, 174 wind erosion features 244 –246 production and patterns 242–243 see also aeolian processes and products wind ripples 248 wind-wave climate Kuray–Chuja palaeolake 115– 116, 120, 121–122 pocket beach crestline model 122– 123 yardangs 246, 247 Zarqa Formation 149 zastrugi 245, 246, 247, 250 zibar 253