landscapes of transition: landform assemblages and transformations in cold regions

Landscapes of Transition

The GeoJournal Library

Volume 68

Managing Editor: Max Barlow, Concordia University, Montreal, Canada

Founding Series Editor: Wolf Tietze, Helmstedt, Germany

Editorial Board: Paul Claval, France R.G. Crane, U.S.A. Yehuda Gradus, Israel Risto Laulajainen, Sweden Gerd LOttig, Germany Walther Manshard, Germany Osamu Nishikawa, Japan Peter Tyson, South Africa Herman van der Wusten, The Netherlands

Landscapes of Transition Landform Assemblages and Transformations in Cold Regions

edited by

KENNETH HEWITT

MARY-LOUISE BYRNE

MICHAEL ENGLISH

and

GORDON YOUNG

Wilfrid Laurier University, Waterloo, Ontario, Canada

SPRINGER-SCIENCE+BUSINESS MEDIA, BV.

A C.I.P. Catalogue record for this book is available from the Library of Congress

ISBN 978-90-481-6037-2 ISBN 978-94-017-2037-3 (eBook) DOI 10.1007/978-94-017-2037-3

Cover iIIustration: Collapsed blocks on Nechelic Channel

Printed on acid-free paper

AII Rights Reserved © 2002 Springer Science+Business Media Dordrecht Originally published by Kluwer Academic Publishers in 2002 No part of this work may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, microfilming, recording or otherwise, without written permission from the Publisher, with the exception of any material supplied specifically for the purpose of being entered and executed on a computer system, for exclusive use by the purchaser of the work.

CONTENTS

Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. vii

Acknowledgments .................................................... ix

Introduction, Kenneth Hewitt. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. 1

PART ONE

GLACIAL AND IDGH MOUNTAIN ENVIRONMENTS. . . . . . . . . . . . . . . . . .. 9

CHAPTER ONE Development of Landform and Sediment Assemblages at Maritime High-arctic Glaciers, Michael 1. Hambrey and Neil F. Glasser ........ 11

CHAPTER TWO Proglacial and Paraglacial Fluvial and Lacustrine Environments in Transition, Peter G. Johnson ................................... 43

CHAPTER THREE Postglacial Landform and Sediment Associations in a Landslidefragmented River System: the TransHimalayan Indus Streams, Central Asia, Kenneth Hewitt . ................................. 63

CHAPTER FOUR Fluvial Sediment Transfer in Cold Regions, Michael Church . . . . . . . .. 93

PART TWO

COLD LOWLAND AND COASTAL ENVIRONMENTS. . . . . . . . . . . . . . . .. 119

CHAPTER FIVE Where on Earth is Permafrost? Boundaries and Transitions, Michael W. Smith and Dan W. Riseborough ................................. 121

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CHAPTER SIX Typical Aspects of Cold Regions Shorelines, Mary-Louise Byrne and Jean-Claude Dionne ........................................ 141

CHAPTER SEVEN Landform Development in an Ar ~tic Delta: The Roles of Snow, Ice and Permafrost, H. Jesse Walker ... .............................. 159

CHAPTER EIGHT The Search for an Arctic Coasta Karren Model in Norway and Spitzbergen, Joyce Lundberg and Stein-Erik Lauritzen . . . . . . . . . . . . .. 185

CHAPTER NINE

Index

Sedimentary Characteristics, Biological Zonation and Physical Processes of the Tidal Flats of Iqaluit, Nunavut, Janis E. Dale, Shannon Leech, S. Brian McCann and Glenda Samuelson . . . . . . . . . . .. 205

235

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PREFACE

This volume had its origins in an international symposium organised by the Cold Regions Research Centre, and held at Wilfrid Laurier University in November, 1999. The chapters are modified from a selection of the papers at the meeting, and reflect reviews and revisions in light of discussions then.

The original idea for the meeting was to address certain questions that the organisers were encountering in their own work, and that we felt had received limited attention in the recent literature. The two broad issues we wanted to address were: the complex associations of actual landforms and processes in cold regions, and how the almost universal legacies of past, different cold environments of the late Quaternary affect these landscapes in the present.

The former involves the problem of identifying landform and sediment complexes, and the interrelations of relevant processes. We sought to identify this in terms oflandform and sediment assemblages appropriate to regional and field-oriented concerns.

The second main concern involves the ways in which present day processes and landform development reflect patterns of adjustment away from past conditions and towards later and contemporary conditions. However, we were only indirectly considering reconstructions of former conditions and the landforms they had produced, or chronologies of changing environments from then to now. These are obviously important but we chose to emphasize the adjustments within and among present-day landforms and processes. We adopted the term 'transitions' to convey how so many landscapes are at certain stages of change that are not adjusted either to past, intervening, or present conditions. Even more important, from a landscape perspective, is the extent to which they are constrained by specific geomorphic response characteristics. We must consider that any given landscape is at a unique stage in distinctive temporal and spatial processes and an incomplete reorganisation of energy and sediment fluxes.

In recent years, the focus has been on processes peculiar to cold regions and related sub-specialties, especially glacial, periglacial, nival, or biogenic processes, iceinfested waters, or attendant microclimatic, hydrological, cryogenic or sedimentation processes. These concerns have improved our understanding enormously. The authors have worked mainly within such specialisations. Nevertheless, there is a certain sense of diminishing returns for landscape investigations, and various new concerns suggest the need to return to more holistic or eclectic, comparative and regional frameworks. These would address the actual complexity of given cold landscapes rather than seeking to classify and separate them in terms of unique processes. The growing demands to understand the impact of climate change, for example, require an ability to disentangle its consequences from the many processes of change in cold environments, including transitions as defmed above. These, too, mainly involve the interrelations of a range of processes and associations of landforms.

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Initially, of course, progress is as likely to come from redirecting work on specific problems in existing work on, say, glacial or coastal processes. And we had to identify researchers already pursuing relevant themes or promising developments. In some cases, such as the paraglacial with respect to the transitions theme, or sediment assemblages, existing work is directly concerned with the themes of the volume. In several chapters significant departures arose from research findings that highlighted the limitations of existing notions.

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ACKNOWLEDGMENTS

The symposium from which this book evolved was sponsored by the Cold Regions Research Centre and a Conference grant from The Office of Research, Wilfrid Laurier University. The editors of this book are indebted to Jo-Anne Horton of the Geography and Environmental Studies Department at Wilfrid Laurier University for her significant role in preparing the text, to Pam Schaus for her work on many figures, Julie Pocock for her help in preparing the index, and to John Barlow for his role in assisting in the organization of the symposium. We thank Drs. I. Brookes, J. Gardner, B. Luckman, and H. Saunderson for commentary on the papers and themes of the symposium.

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INTRODUCTION: LANDSCAPE ASSEMBLAGES AND TRANSITIONS IN COLD REGIONS

Kenneth Hewitt Cold Regions Research Centre and Department o/Geography and Environmental Studies Wilfrid Laurier University 75 University Avenue West Waterloo, Ontario Canada N2L 3C5

Abstract

This essay introduces the varieties and distinctive characteristics of cold regions. A landscape approach is adopted, emphasising associations ofland forms and sediments and the importance of spatial and temporal transitions in earth surface processes. The broadest classes of cold region are the "zonal" or higher latitude examples, and the "azonal" ones based on elevation, continentality and air mass regimes. Vast and singular subregions are also characterised by cold tolerant ecosystems such as heathlands and bogs, boreal or montane forests. These playa major role in landform development and appearance in their respective zones.

Work relating to earth surface processes tends to separate glacierised (i.e., icecovered) and non-glacierised areas, in the latter, periglacial and nival regimes, and cold coastal conditions. Cold regions geomorphology, in recent decades, has focused on processes peculiar to low temperatures and freeze-thaw. Much of our understanding has developed out of specialised investigations and experimental work on distinctive processes of these regimes, and the search for more abstract, general models of them. But a regional and comparative perspective must also recognise the large role played by weathering, fluvial, lacustrine, marine or aeolian processes shared with other regions if modified in cold contexts. A landscape and comparative interest directs attention to the variety of landform associations and related sediment assemblages, and how they represent the complexities of on-going and historical development.

Most cold regions landscapes record past and on-going environmental change. They contain more or less extensive legacies of clima-hydrological, ecological and geotectonic changes in the Quaternary. They involve glaciation and deglaciation, isostatic crustal adjustments, fluctuations in sea levels and in the extent of marine and freshwater ice, changes in snow covers, seasonal ground ice, and the patterns and thickness of permafrost. These legacies, and how they are removed or transformed, are integral to the

K. Hewitt et al. (eds.), Landscapes a/Transition, 1-8. © 2002 Kluwer Academic Publishers.

2 Kenneth Hewitt

interpretation of existing cold region landscapes. What we observe are geomorphic ''palimsests'' in which relict, overlapping and replacement forms are interwoven.

The notion of "transition" addresses the sense in which landscape development is not merely chronological and linear, or simply a "lagged" response to climatic and tectonic changes. There are diachronous episodes of (incomplete) readjustment to the cessation of past conditions, and towards later conditions, of which those at present are only one set. There are distinctive spatial and temporal patterns of adjustment, including "self-adjustment" specific to the earth surface processes at work. The paraglacial is a classic example. It is suggested that such temporal and spatial responses in earth surface processes apply much more generally as part of landscape transformation in the Quaternary.

Introduction

"Cold regions" are those areas of the Earth's surface where sub-zero temperatures, snow and ice, freezing and thawing, and freeze-tolerant organisms, are perennially or seasonally present. The emphasis here is on landscapes that are formed or constrained by cold conditions. In particular, we look at landform morphology, earth surface processes and sediments reflecting the presence of frost, ice and melt water, and the influence upon landscape of cold-adapted plants and animals. The main concerns of the volume are landform and sediment assemblages in cold regions, and how they depend upon transitions over time and space in the processes generating them. First, however, a brief overview .of cold regions will situate these concerns within the larger context.

Relative abundance of water is a decisive fact of planet Earth's physical environment. However, at least as important for geomorphic and life processes, are temperature and pressure conditions close to, and continually fluctuating across, phase change parameters for water. (On other planets where water exists, but locked up in a liquefied atmosphere, perennially frozen oceans, or permafrost at depth, life forms are not known or clearly present. The "land forms" depend largely or wholly on the impacts and accumulation of debris from space, sometimes on internal tectonic forces.)

Solar radiation and Earth's gaseous atmosphere promote continuous cycles of evaporation, condensation and precipitation of liquid moisture. These dominate the presence and influence of moisture at the Earth's surface. They are second only to solar radiation in controlling thermal environments.

However, extensive land and ocean surfaces are also frozen, or have temperatures that regularly cross the freeze-thaw threshold. Here, the temperature relations of moisture availability, its phase and behavior, introduce distinctive physical, chemical and biotic processes. Along with a mobile crust, such cold-related conditions make for diverse and, often, exceptionally dynamic climageomorphic environments. Moreover, cold environments have expanded and contracted enormously over the past 30,000 years or so. This has left the imprint of cold conditions on large areas no longer cold, and brought more or less drastic changes in those that are.

Introduction 3

The Scope and Variety of Cold Regions

Four primary geographical regions with cold conditions can be identified. They relate to latitude, altitude, continentality, and air mass regimes. High latitude, or polar, sub-polar and cool temperate regions, are sometimes referred to as "zonal ", a direct consequence of (low) solar altitude and cold winters. The other three influences are "azonal ". They overlap with and can intensify or moderate the latitudinal effect. In particular, however, they bring sub-zero temperatures or frozen precipitation to areas closer to the Equator, otherwise identified with mild, temperate, and even hot, tropical conditions. They involve, respectively, intensification of perennial or seasonal cold and snowfall by elevation or the orographic effect; radiative cooling ofland surfaces in winter, unmoderated by inflows of milder air; or air masses that bring heavy winter snow and cool, cloudy summers.

Broad, secondary divisions arise within cold regions due to the intensity and duration of cold conditions and moisture regimes, and among cold high plateaux or mountain lands as a function of latitude, air mass regime and ruggedness of terrain. Equally important regional landscapes are associated with major vegetation types such as peatlands, boreal or montane forest types, or cold mires (Specht 1979; Gore 1983). These characterise and influence the development and morphology of vast areas. Plants and animals affect the forms and variety of cold regions through their influence on patterns and processes of hydrology and sedimentation, through surface and soil processes, and organic build-ups (see Walker, this publication). We should also mention the constructional land forms and related regulation of hydrology by vegetation; the importance of vegetation in stabilising mountain slopes and unconsolidated sediments; and of fire ecology in certain cold forest, heath and peat environments. Organisms in cold lakes, rivers, deltas, or coastal waters, exercise a huge influence on patterns and rates of erosion and sedimentation (see Dale et aI, this publication).

High latitude cold oceans are seasonally or perennially ice-covered. Freshwater lakes and rivers in cold regions are seasonally frozen. The floating ice promotes and constrains geomorphic and ecological processes along coasts and river valleys. This applies when the ice is in place, during its break-up and melting. The configuration of the continents means that perhaps one third of the world's marine coastlines are affected by floating ice (see Byrne and Dionne, this publication). In the vast subarctic regions of Eurasia and North America are river systems where seasonal freezing, break-up, vast ice jams and inundations, exert a huge influence on land forms, sediment delivery and ecology.

Among the many local and regional subdivisions, zones of transition or ecotones, have received much attention. They often reflect sharp transitions in key climageomorphic processes and habitats, and some are visually as well as functionally well-defined. Examples include snow lines and timber lines. In the mid latitude and tropical high mountain ecotones between zones of increasingly intensive cold conditions - but moderated by influence of moisture deficit or excess - are stacked, one above the other, in altitudinally organised climageomorphic zones (Kowalkowski and Starkel 1984; Hewitt 1993; Sarmiento 2000). Cold coasts involve series of ecotones surrounding and beneath the water. The latter reflect water depth, exposure to waves and currents, modified by a

4 Kenneth Hewitt

variety of attendant processes according to the presence, duration and behavior of floating or fast ice.

At present, about half the world's land surfaces have seasonal snowfall as a significant climatic, hydrological and physiological factor. About one quarter have frozen ground too extensive for weak: or short warm seasons to fully melt it, although the snow cover may melt. Here frost, ground ice or permafrost are decisive for geomorphic processes and identified as periglacial regimes. Snowy conditions without significant ground freezing extend nival effects into milder areas. Conversely, where snow persists and builds up from year to year, glaciers cover the land. The Antarctic ice sheet dominates total world glacier ice, and covers about 12 million km2 (Williams and Ferrigno 1998). The present glacier cover of the Northern Hemisphere is about. 2.3 million km2, of which the Greenland Ice Sheet comprises almost 75% (Field 1975, 3).

Glacial and nival melt waters, wind and wave action, are also profoundly important, although their influence is not confmed to cold regions. Mountain glaciers, especially those outside polar and subpolar regions, comprise a small fraction oftoday's ice. However, the mid- and low latitude glaciers, and the movement of glacial meltwaters and sediments from them to surrounding foothills and lowlands, have a significance out of all proportion to their share of the global ice cover. That applies to both landscape development and human affairs. Indeed, the water, sediment and dust carried into surrounding ''warm'' regions and the world ocean are major, indirect cold region influences on landscapes and sedimentation. This includes the loess that covers vast areas and has come mainly from past or present cold regions. Some of the largest submarine cones of sediment have been built of materials from the larger, present-day or formerly glacierised mountain ranges.

We will concentrate on conditions in, and as they affect, present day cold regions. But, it is not implied that exchanges between them and other environments are less important.

Regional Landscapes and Landform Associations

Most of the chapters which follow adopt a view, as it were, from the landscape, and in terms of landscape-forming processes. The influence of broader atmospheric and geological changes are considered as they inform that perspective. It is accepted that the various styles of cold environment and forms of ice, freeze-thaw and melt waters, are associated with distinctive climageomorphic regimes. Along with ice-infested waters and cold-adapted ecosystems, they help to create distinctive landscapes.

In recent decades the emphasis in cold regions geomorphology has been on conditions special to them, and process-defmed subregions. Geomorphologists differentiate mainly between glacial, periglacial including permafrost, nival and cold desert regimes, and concentrate their efforts on their distinctive processes and forms. A great deal of work has examined singular features of periglacial, glacial and nival conditions - palsas or pingoes, drumlins and eskers, cirques or avalanche boulder tongues - and how they record specific processes (Clark 1988; Williams and Smith 1989; Hambrey 1984; Fitzgerald and Rosen 1987; French 19 ). This has certainly served to

Introduction 5

inform and even to revolutionise our understanding of the activity behind observed landforms.

However, the goal of describing, interpreting and comparing regional landscapes requires a more holistic approach. In geographical thought, the regional approach is classically distinguished from a "systematic" one. The difficulty is to incorporate understandings obtained by each approach into the other, given that we reject the view that one of them is superior, or can be reduced, to the other. Since the systematic approach has dominated geomorphology over recent decades, some of the salient problems at present are those of revising the regional approach with the aid of new insights developed in process geomorphology, sedimentology and cognate earth science fields.

Regional or "chorological" concerns involve the distribution, the interactions and mutual constraints among, the features and varieties of processes in any given landscape. This is usually combined with a comparative assessment of the differing and common features in different landscapes, or their changing balance as we move across the earth's surface. Such an approach is developed in the overview of cold coastal landforms by Byrne and Dionne, (this publication). In general, the chapters below not only emphasise land surface features in themselves, and their diversity in given settings, but address mainly regional, field oriented questions.

The idea of landform associations seeks to address the mutual fertilisation of regional description and systematic investigation. The idea of "landform assemblages" refers, more specifically, to the varieties of processes and their interactions that compose any actual landscape or distinctive class oflandscapes. Since a large fraction oflandforms in cold regions, as elsewhere, consist of or are cut in superflcal sediments, the understanding of landform assemblages almost invariably involves some study of sediments. And, in keeping with the assemblage notion, the varying composition and facies organisation or "architecture" of sediments comprise important diagnostic features.

Of course, such concerns have not been absent hitherto. The "landsYstems" approach has sought to address very similar issues (Eyles 1983). Work on sediment assemblages and sedimentary basin studies have developed techniques and concepts appropriate to the recognition of landform-sediment complexes (Miall 1990 ). These approaches explore ways to deal with heterogeneous processes. They examine the complexities of temporally and spatially varying developments, and the kinds of technical and intellectual questions arising out of, but inadequately dealt with by, process-specific work. They offer important tools for, and steps towards, site and region specific geomorphologies.

Several of the chapters below directly address such questions, They are developed in relation to the glacial context by Hambrey and Glasser, and by Johnston (chapters 2 and 3 ), in the high mountain context by Hewitt (chapter 4), for an Arctic delta by Walker (chapter 8) and in a cold coastal setting by Dale et al (chapter 9).

6 Kenneth Hewitt

Landscapes infof Transition

Equally, a regional and field based approach is obliged to recognise how nearly all actual landscapes bear the imprint of past conditions. Some of these conditions may no longer apply at all, others may have undergone great changes in intensity or scope. Thus, cold region conditions have expanded and contracted geographically over large areas. Almost all existing cold regions landscapes involve legacies of past and different "cold" conditions. Their landscapes record how climageomorphic controls have changed in pattern and strength over time.

The Russian geographer Peter Kropotkin seems to have been the first to recognise, in the 1870s in northern Eurasia, landscapes that reflect not merely an "Ice Age" and the imprint of glaciers, but a succeeding "Lake Age". It is associated with vast areas of lake sediments recording the melt water ponded in glacially excavated basins, or by drainage interruptions and realignments, by glacial deposits or remnant ice masses. In the Eurasian continental interior Kropotkin also recognised how his Lake Age was followed by climatic dessication and a trend towards more arid processes.

Similar developments were later recognised in North America, as well as the widespread replacement of ice sheets by permafrost and periglacial conditions, and modifications due to isostatic rebound, sea level rise, and colonisation by plant and animal communities (ed. Fulton 1989).

It is generally accepted that temporal transitions in cold regions have two important geographical aspects. On the one hand, there has been change in environments through the expansion and contraction of climatic, tectonic, biotic and anthropogenic conditions. There have been large changes in the extent of glaciers, permafrost, periglacial activity, snowfall, dryness and colonisation by cold-tolerant vegetation. On the other band, although they may have disappeared long ago, the work of past cold conditions is, to a greater or lesser extent, imprinted on the ground today. The legacies range from those of huge features like the Fennoscandian Ice Sheet or Glacial Lake Agassiz, to purely local and scattered ones, such as cirques, relict patterned ground, small patches of dunes or loess layers.

These phases in the past history of cold regions can be treated in terms of at least three analytical strategies. They may be used to esablish geochronologies, to reconstruct conditions in past times, or to help interpret the development of the present landscape. Work on Quaternary reconstruction has tended to combine the search for morphostratigraphic sequences and datable materials for historical chronologies, with reconstruction of conditions at particular time horizons within those chronologies. There has been a steady stream, sometimes revolutionary surges, of work showing that areas formerly thoughtnotto have undergone, say, glaciation or periglacial modification did so, and vice versa. This depends upon knowledge of, and skill in identifYing, the relevant processes, and adequate diagnostic sedimentary or morphological features.

Such questions are addressed in relation to permafrost by Smith and Riseborough (chapter 5). They show how a changed understanding of the regional distribution and varieties of permafrost now, especially the constraints governing it, is a necessary basis for reconstructing past changes in that distribution, or predicting its future

Introduction 7

response to climate change. The chapters by Hambrey and Glasser, Hewitt, and Lundberg and Lauritzen show how the recognition and reconstruction of past conditions through appropriate models and field investigations, involve on-going and not readily resolved problems of cold regions geomorphology.

However, the interpretation of regional landscapes introduces a third and, in some ways, more crucial and difficult aspect. Here it is identified with transitions or phenomena of transition. We know changes in geomorphic conditions have happened, and are going on at present. Exactly how they change, as opposed to sequences of landscapes attributed to particular regimes, is much less certain. Two concerns are at issue. First, there are the landfoms and sediments that reflect different past cold conditions in present-day cold regions. They would include such obvious cases as glacial molded rock outcrops, kettles or moraine systems in areas far from present-day glaciers, raised or drowned beaches and remnants of large-scale patterned ground in areas where periglacial processes are absent or weak. But these are not merely relict forms. They operate as constraints upon present-day developments. They are often associated with seemingly anomalous or disturbed relations of present-day processes to the landforms and landscapes in which they occur. Examples include the chaotic pattern of streams in formerly ice-scoured, rocky lowlands, ''misfit'' streams; shorelines where wave action is weak but having spectacular cliffs that record an emerged coast or glacial oversteepening.

Secondly, however, these are parts oflandscapes, hydrology, vegetation covers and pattems of sedimentation that are in incomplete transition from past conditions. The notion of temporal "transition" is employed here to address the ways in which environmental changes are uniquely expressed through adjustments in landforms and earth surface processes. The processes are driven by contemporary heat and moisture conditions, available relief, lithologies. But they are not only or directly responding to them. Conversely, they may have received impetus or initial conditions from past events and changes in external forcing conditions of climate or tectonics, but they evolve according to a "logic" of their own. Geomorphic transitions have, or tend to generate, their own temporal and spatial shape. They involve mechanisms orpattems of adjustment .distinct from, and not readily obvious in, the mere chronology of, say, climate change or tectonics.

Few aspects of landscape respond instantaneously to climatic and geological conditions. This is quite apart from the fact that the latter involve more or less drawn out and irregular episodes of change, rather than sudden, region-wide or unidirectional shifts. Most landforms, themselves, record the resistence of solids to mechanical or chemical change, not merely erosional forces. They will remain as they are until subjected to stresses sufficient to deform or remove them. When such thresholds are surpassed, the processes may have to operate for millenia and more, to fully remove or transform features generated under past conditions.

Meanwhile, we must challenge the implication that, when geomorphic processes are modifying past landscapes, the adjustments merely carry them directly from one "equilibrium" to another, perhaps complicated only by the rate at which this occurs. Indeed, there is little possibility of such a straight-forward impact-response relation to external forcing conditions. Rather, there are combinations of intervening constraints, self-

8 Kenneth Hewitt

adjusting mechanisms or "epicycles" peculiar to the earth surface processes affected. This gives them specific temporal shapes and landscape imprints of their own.

Paraglacial sedimentation is a "classic" transitional regime in cold regions. Church (this publication) gives an extended analysis and discussion of its spatial and temporal form. He shows how it involves the redistribution of former glacial sediments and generates an evolving pattern of related landforms. Isostatic rebound in heavily glacierised areas can also be described in such terms. Hewitt (this publication) suggests that catastrophic slope failure in formerly glaciated high mountains initiates distinctive transitional disturbances of fluvial development, as well as slope adjustment. Byrne (this publication) indicates that distinctive developments of this kind occur in aeolian dune systems of cold regions. It seems that similar arguments would be relevant to the development of emergent and submergent coastlines, lakes and other freshwater wetlands, constructional landforms of vulcanism, or active tectonic features in cold regions, since these phenomena of landscape change articulate with, and are recorded in the landform and sediment assemblages. Thus the two themes explored in this volume are seen as part of a single problem framework.

References

Clark, M.J. (1987) Geocryological Inputs to the Alpine Sediment System in Glacio-fluvial Sediment Transfer, in A.M. gumell and M.J. Clard, eds., Glacio-flavial Sediment Systems: an Alpin Approach, Wiley, New York, 33-58

Eyles, N. (1983) Glacial geology: and landsystems approach, in Eyles,N. (ed.) Glacial geology: And Introduction/or Engineers and Earth Scientists, Pergamon, Oxford.

Field W.O. (Ed.) (I 975) Mountain Glaciers o/the Northern Hemisphere 2 vols + Atlas, Cold Regions reserach and Engineering Laboratory, Corps of Engineers, U.S. Army, Hanover, New Hampshire.

Fitzgerald, D.M. and Rosen, P.S. (Eds.) 1987 Glaciated Coasts Academic Press, New York. French, H. (1987) Periglacial Geomorphology in North America, Progress in Physical Geography, 11/4,533-

547. Fulton, R.J. (ed.) (1989) Quaternary Geology o/Canada and Greenland Geology of Canada, no.l. Geological

Survey of Canada, Ministry of Supply and Services Ottawa. Gore, A.J.P. (ed.) (1 983) Mires: Swamps, Bog, Fen, and Moor: General Studies Ecosystems of the World 4A,

Elsevier, Amsterdam. Hambrey, M.J. (1994) Glacial Environments, University of British Columbia Press, Vancouver, B.C.

Miall, A.D. 1984 Principles 0/ Sedimentary Basin Analysis Springer-Verlag, New York. Hewitt, K. (1993) The Altitudinal Organization of Karakoram Geomorphic Processes and Depositional

Environments, in J.F. Shrader, Jr. ed., Himalaya to the Sea: Geology. Geomorphology and the Quaternary, Routledge, New york, 159-183.

Kowalkoski, A. and Starkel, L. (1984) Altitudinal Belts of Geomorphic Processes in the Southern Khangai Mts. (Mongolia), Studie Geomorphologia Carpatha-Balcanica (Krakow) XVll, 95-116.

Mial1, A.D. (1990) Principles o/Sedimentary Basin Analysis, 2ad edition, Springer-Verlag, New York. Sarmiento, F.O (2000) Breaking mountain paradigms: ecological effects on human impacts in Man-aged

Tropandean landscapes, Ambio 'l9n, 423-431. Specht, R.L. (ed.) (1979) Heathlantls and Related Shrublands: Descriptive Studies Ecosystems of the World

9A, Elsevier, Amsterdam. Williams. P.J. and Smith, M.W. (1989) The Frozen Earth: Fundamentals 0/ Geocryology Cambridge

University Press, New York .. Williams, Jr., R.S. and Ferrigno, I.G. (1998) Areal Extent of Present-day Glaciers of the World Satellite

Imoge Atlas o/Glaciers o/the World United States Geological Survey, Professional paper 1386-A.

PART I

Glacial and High Mountain Environments

DEVELOPMENT OF LANDFORM AND SEDIMENT ASSEMBLAGES AT MARITIME IDGH-ARCTIC GLACIERS

Michael 1. Hambrey and Neil F. Glasser Centre/or Glaciology Institute o/Geography and Earth Sciences University 0/ Wales Aberystwyth, Ceredigion SY233DB United Kingdom

Abstract

Detailed studies of sedimentary processes and landform development at modem glaciers are an essential pre-requisite for the interpretation of Quaternary glacial sediments and landforms. Recent work on polythermal glaciers in Svalbard has provided new insight concerning the processes responsible for glacial sediment/landform assemblages. Although the landforms associated with Svalbard glaciers are not in themselves unique, the particular assemblages and proportions of sedimentary facies differ markedly from those in temperate and cold glacier systems. The main conclusion is that deformation within glacier ice, as debris is entrained and subsequently transported, is the primary control on the nature of landform/sediment assemblages in the proglacial areas of Svalbard valley glaciers. The most important landform-creating modes of debris entrainment are:

(1) Incorporation of angular rockfall material within the stratified sequence of snow, fIm and superimposed ice, followed by folding with flow-parallel axes; the resulting medial moraines are preserved in the proglacial area as linear debris trains;

(2) Entrainment of debris at the bed to form a several metre-thick basal ice layer, which is released as a sheet of basal till;

(3) Incorporation of basal debris within longitudinal foliation, resulting in landforms referred to as foliation-parallel ridges;

(4) Thrusting, whereby basal and subglacial sediments are uplifted towards the glacier surface, and ultimately released as individual mounds within a large end-moraine complex (often referred to as 'hummocky moraine');

(5) Subglacial upright folding with transverse axes and faulting also producing large end-moraine complexes;

(6) Reworking of thrust- or fold-derived glaciofluvial material to produce longitudinal debris ridges in the ice, although their translation into landforms is poor.

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K. Hewitt et al. (eds.), Landscapes of Transition, 11-42. © 2002 Kluwer Academic Publishers.

12 Michael J. Hambrey and Neil F. Glasser

The principal sedimentary facies associated with these polythermal glaciers is diamicton of basal glacial origin, followed by sandy gravel of glaciofluvial origin. These facies, reworked by thrusting in glacier ice, dominate the end-moraine complexes.

These observations and inferences have been applied to areas such as Britain, where glaciers no longer exist, in the interpretation of Pleistocene landforms and sediments.

1. Introduction

The development oflandforms associated with late Quaternary glaciers and ice sheets has often been based on temperate alpine or Icelandic glaciers. However, this has sometimes resulted in an erroneous perception of the processes responsible for these landforms. Only in the last decade has the Arctic received the attention that it warrants, and a range of new models has now emerged for the development ofice-marginal constructional landforms that have considerable bearing on how landforms developed in areas no longer occupied by ice. In these investigations, it has been recognised that the only way to understand landform genesis is to establish the glaciological basis for debris-incorporation and transport. Many landforms, then, are controlled by the way debris is associated with ice deformational structures within the glacier.

This contribution examines the various processes of debris incorporation, transfer and deposition in High-Arctic glaciers, focusing especially on the archipelago of Svalbard (including Spitsbergen) at the NW edge of the European continental shelf (Figure I). In this climatically sensitive polar-maritime environment, the glaciers are believed to be mainly of the poly thermal type, that is they consist of warm ice in their deeper, upperreaches, and cold ice (frozen to their beds) in their marginal and snout areas. The polythermal glacier thermal regime gives rise to a distinctive suite of landforms which, if identifiable in formerly glaciated areas, can help constrain palaeoclimatic reconstructions in those areas.

2. Climatic and Glaciological Regime in Svalbard

The Svalbard archipelago (77° N to 80° N) lies at the northern extent of the mild Norwegian Current, a branch of the Gulf Stream, and enjoys a relatively mild climate for its northern latitude. On the western coast, the average annual temperature is -6°C. The average temperature on the west coast in the warmest month (July) is SOC whilst in the coldest month (January) it is -ISOC. Although there are contrasts between the maritime west coast and the more continental interior, precipitation in Svalbard is generally low. Typical values at sea level are 400 to 600 rom annually, falling to half these values inland. Precipitation in the more mountainous regions is increased by orographic effects, but even on the glaciers snowfall of more than 2-4 m is rare (Hagen et al. 1993). Ice-free land areas are underlain by permafrost to depths of between 100 and 400m.

Landfonn and Sediment Assemblages

Kronobroen/Kongsvegen

2 Auslre Or0ggerbreen, Veslre Lovenbreen, Mldl,e Loviinbreen, Auslre Lov6nbreen

3 Uv6rsbroon. Comforllossbrocn

4 Holmstrombreen

5 Sofstrombreen

6 Bal<enlntreen

7 Flnsterwalderbreen . Hessbr .... n

Figure I. Location map of glaciers in Svalbard referred to in the text

13

The archipelago is currently 60% glacierised. The largest volumes of ice are accounted for by the highland icefields and ice caps of eastern Spitsbergen, Edge0}'a, Barents0}'a and Nordaustlandet (Figure I). These ice masses cover the low mountainous areas, and their outlets are divided into individual glaciers by mountain ridges and nunataks. Many of these reach the sea to fonn large calving glaciers. Smaller cirque glaciers are also common, particularly on the more alpine terrain in western Svalbard.


Many of the larger glaciers in Svalbard are "subpolar" or polythermal, with extensive areas of temperate ice beneath their accumulation areas, but with their termini frozen to the bed (Hagen and Saetrang 1991; Hagen et al., 1991; 0degArd et al. 1992; Bj6msson et al. 1996). Most of the smaller cirque glaciers are probably cold based throughout. The hydrology of polythermal glaciers is poorly understood relative to alpine glaciers (Bamber 1989; Hagen and Saetrang 1991; Hagen et al. 1991; Vatne et al. 1996), although recent advances have been made in this field (Hodgkins 1996).

The debris structure of polythermal glaciers is also relatively little known in comparison with that of temperate glaciers (Weertman 1961; Swinzow 1962; Boulton 1970a; 1978; Hooke 1973; Clapperton 1975; Hambrey and MOller 1978). It is only in recent years that a clearer picture of the debris structure within polythermal glaciers has emerged. This work has confirmed the importance of thrusting in elevating basal debris within polythermal glaciers (Bennett et al. 1996a,b; Hambrey and Huddart 1995; Hambrey et al. 1996; Murray et al. 1997). It has also highlighted the significance offolding of debris-rich stratification in organising both supraglacial, basal and glaciofluvial debris (Hambrey and Dowdeswell 1997; Glasser et al. 1998a, 1998b; Hambrey et al. 1998).

An estimated 35% of the glaciers on Svalbard are surge-type (Hamilton and DowdeswellI996). These glaciers are therefore prone to dramatic increases in velocity and rapid frontal advances, followed by periods of quiescence during which velocities are generally low. Surge-type glaciers in Svalbard typically have relatively long quiescent phases between surge events (Dowdeswell et al. 1991). Surges have now been documented at numerous Svalbard glaciers including Usherbreen (Hagen 1987, 1988), Bakaninbreen (Murray et al. 1997), Kongsvegen (Melvold and Hagen 1998; Bennett et al. 1999), Seftstrmnbreen (Boulton et al. 1996), Holmstrmnbreen (Boulton et al. 1999), BrAsvellbreen (Solheim and PfJrman 1985) and Fridtjovbreen (Glasser et al. 1998c).

3. Structural Glaciological Controls on Debris Transport

Debris-entrainment in High-Arctic glaciers takes place by a variety of mechanisms which can be linked to the evolution of ice structures (Hambrey et al. 1999). The most significant structures in this respect include stratification, the deformed basal ice zone, foliation and thrusting, commonly associated with folding. Several modes of debris-entrainment are outlined below.

3.1. INCORPORATION OF ROCKFALL MATERIAL

Primary stratification is inherited from the build-up of snow and superimposed ice in the accumulation area, and is sometimes supplemented by rockfall material from the glacier headwalls (Hambrey et al. 1999). This material is typically angular, and is incorporated englacially as a result of burial and folding of the stratified sequence. Converging flow, where multiple accumulation basins supply a narrow tongue, promotes folding. Component flow units may be reduced in width by more than 50%. Typically, fold styles range from open "similar" to less common chevron and isoclinal types, each commonly associated with an axial-plane foliation. Fold axes typically are parallel to flow and plunge gently up-

Landform and Sediment Assemblages 15

glacier. As debris-laden folded stratification intersects the glacier surface in the tongue, "medial moraines" emerge from a single point-source or multiple point-sources, producing increasingly broad spreads of debris down-glacier (Figure 2a). As in temperate glaciers (e.g. Meier and Post 1969, Lawson 1996), this relatively simple structure is complicated by surges, giving rise to "looped" moraine structures in which the debris layers originate from the stratification (Hambrey and Dowdeswell 1997).

3.2. ENTRAINMENT OF DEBRIS AT THE BED

Maritime High-Arctic glaciers carry a large basal debris load, commonly reaching a thickness of several metres with debris concentrations in some zones approaching 100%. Although few studies have been made of the basal debris layer in such glaciers, by analogy with other glaciers, the principal processes of entrainment are regelation, flow through the vein system, bulk freezing-on, folding and shearing (Knight 1997). Layers of ice facies are commonly subjected to repeated shearing and isoclinal folding. The typical coarse clear ice crystal character of basal ice, combined with varying proportions of debris allow it to be distinguished from coarse bubbly ice derived from snow, when glaciotectonically transferred to higher level positions within the glacier.

3.3. INCORPORATION OF DEBRIS WITHIN LONGITUDINAL FOLIATION

Longitudinal foliation is a structure contmon to all glaciers, being the product of tight folding, or simple or pure shear. In glaciers dominated by converging flow, longitudinal foliation pervades the width of a glacier (e.g. White Glacier on Axel Heiberg Island, Canada: Hambrey and MUller 1978). Svalbard glaciers commonly have pervasive nearvertical foliation throughout their widths.

Debris is associated with foliation in two ways: (i) Moderately well-sorted angular debris of supraglacial origin which has been

folded so tightly that the original stratification has been transposed into foliation. Alternatively, the debris is disposed around the hinge of a more open fold, which is intersected by an axial place foliation (Hambrey et aZ. 1999).

(ii) Debris showing basal characteristics (i.e. a wide variety of clast shapes, clast surface features such as striations and facets on appropriate fine-grained lithologies, and poorly sorted texture), disseminated through, or layered within, coarse-clear ice. This debris is isoclinally folded or sheared within foliation, especially in ice-marginal areas and occasionally within the lower reaches of medial moraines. The folding mechanism which allows basal debris to reach the surface is obscure, but it has been suggested that in the thicker parts of the glacier a deformable bed of debris is folded within the overlying ice in the zone of converging flow (Hambrey et aZ. 1999). Again the folding is in axial planar relationship with the foliation.


2c

Landfonn and Sediment Assemblages 17

Figure 2. Debris-bearingice structures in Svalbard glaciers. (a) Emergence offolded debris-layer of supra glacial origin and its development into a medial moraine near the snout of Austre Lovenbreen. Note also the debristrains extending from the moraines across the proglacial area. (b) Basally-derived debris associated with longitudinal foliation near the snout of Vestre Lovenbreen (view downglacier). Note extension of ridge onto water-covered proglacial area. (c) Basal debris entrained by thrusting and recumbent folding, observed in a 2-3 m high longitudinal cross-section at Hessbreen. (d) Basally derived debris emerging at the intersection of a thrust with the surface of Finsterwalderbreen. ( e) Longitudinal sediment dyke on Marthabreen.


3.4. DEBRIS-ENTRAINMENT BY THRUSTING

Thrusting is the most controversial mechanism of debris-entrainment, following early suggestions that this was a valid mechanism for entrainment in the margin of the West Greenland ice sheet (Swinzow 1962), and the expression of an alternative view that incorporation of debris was a passive process as ice flowlines turned upwards towards the ice-frontal margin (Weertman 1961; Hooke 1973). The weight of evidence from Svalbard glaciers suggests that thrusting is indeed a valid process, having now been documented in numerous examples (e.g. Glasser et al. 1998; Hambrey et al. 1996,1999; Hambrey and Dowdeswell1997; Murray et al. 1997).

Thrusts observed in White Glacier (Canadian Arctic ) are clearly linked to previous structural inhomogeneities, notably transverse crevasse traces (Hambrey and Miiller 1978), but those in Svalbard glaciers tend to be new structures, first appearing within several hundred metres of the snout. Thrusts may be recognised on the basis of several criteria (Hambrey et al. 1999), but in the context of the contemporary glaciological regime of strongly receding ice fronts, probably few are active today (except in surging glaciers). Thrusting is associated with incorporation of whatever material lies on the bed or within the basal zone, be it basal till, glaciofluvial or even glaciomarine sediment. Sediment thicknesses associated with thrusting may reach several metres, and original sedimentary structures can be preserved. At the opposite end of the spectrum there may simply be thin layers of debris-rich ice of basal origin or films offme debris along a well-defmed plane.

In terms of geometry, thrusts show an asymptotic relationship with the bed and emerge at the surface typically at angles ranging from 15 to 70° (Figure 2c). Most layers with prominent debris-ridges dip at-30°. Thrusts may be laterally continuous for several hundred metres, although significant quantities of debris usually only occur on a fraction of this length (Figure 2d). Intersecting thrusts tend to promote accumulation of prominent debris heaps on the glacier surface (Bennett et al. 1999).

Debris-bearing thrusts are particularly well-developed in polythermal glaciers, and it is likely that the thrusting process is facilitated by the transition from an upper sliding bed to one that is frozen, a condition described from Trapridge Glacier (Yukon) by Clarke and Blake (1991) and inferred for polythermal glaciers in Svalbard (Hambrey et al. 1999). At such locations, water could facilitate the upward-displacement process of basal debris and subglacial debris. Thrusting is not confined to polythermal glaciers, but large-scale debrisincorporation in temperate glaciers generally requires particular topographical conditions, such as ice-flow acting against a reverse slope.

3.5. INCORPORATION AND MODIFICATION OF GLACIOFLUVIAL SEDIMENT IN ENGLACIAL CHANNELS

Flow-parallel ice structures can also incorporate large quantities of glaciofluvial sediment, helping to redistribute sediment within the glacier. On Marthabreen, for example, two types of longitudinal debris-rich structures have been described on the glacier surface: longitudinal sediment dykes and longitudinal ridge accumulations (Glasser et al. 1999). Longitudinal sediment dykes are ridges of sand and gravel, commonly 1-6 m long and 0.5 to 1.0 m high. They vary in width from 0.05 to 0.15 m and are always sub-parallel to the


foliation. Their sediment fill consists of fine sand and granule gravel, with a sedimentary structure that varies laterally along the dyke. In some locations it fonns a massive and structure less unit, while in other locations it fonns numerous graded units of fme granule gravel to fme sand with isolated small pebble clasts. These graded units dip downglacier along the central axis of the dykes and in some locations are also intensely folded in the fonn of tight isoclinal folds. The longitudinal sediment dykes are produced by folding of glaciofluvial sediment along foliation. The direction of folding is parallel to the long axis of the dyke, with the fold axis transverse to it. The width of the sediment dykes varies both longitudinally and vertically and in several places they are completely closed and the sediment fill is pinched out by the ice walls. At the upglacier and downglacier ends of the dyke the sediment fill stops abruptly although the structure can be traced further as a closed and debris-free ice fracture (Figure 2e).

Longitudinal ridge accumulations are larger ridges between 20 and 30 m long and 1 to 3 m high. They have crests of in situ sand and gravel, while the ridge flanks contain slumped debris. Two types of contact between the in situ sediment and the ice have been observed. Firstly, the sediment may lie above a flat or gently convex ice surface. No fractures are observed within the underlying ice and the sediment piles appear to be independent of the glacier surface. Secondly, the sediment may lie within a distinct semicircular channel between 10 and 50 units in diameter. Again, this is not located along an ice fracture and appears to be unrelated to any ice structure, although it is sub-parallel to the foliation. The in situ sediment located on the crest of these ice-cored ridges is between 0.2 and 0.6 m wide and may be up to 1 m in thickness. It consists of a range of pebble and granule gravel units, which may be organised into graded units, but are more commonly massive. Units of pebble gravel supported by a silt matrix are also common. The bedding is either subhorizontal or dips gently upglacier at between 30 and 50. These ice-cored sediment ridges frequently occur downstream of sediment dykes or other debris pinnacles and have been interpreted as the product of sediment reworking by englacial or supraglacial streams flowing parallel to the longitudinal foliation (Glasser et al. 1999).

4. Landform and sediment assemblages

4.1. MORAINE COMPLEXES

Modem glaciers in Svalbard are invariably associated with large end-moraine complexes that were fonned during the Neoglacial maximum around the beginning of the twentieth century. Most glaciers have receded up to several kilometres from these complexes, leaving lesser moraine systems in the intervening zone. Most authors have referred to these moraine complexes as "push moraines", even where defonnation (especially folding) is involved (e.g. Croot 1988; Hagen 1987, 1988; Boulton eta/. 1996,1999; Van der Wateren 1995). However, not all such complexes are the product ofice-push, and in many cases thrusting of sediment and ice is the primary mechanism. Hence the tenn ''thrust moraine complex" has been used where the mechanism has been clearly demonstrated, or more descriptively, "moraine-mound" complex (Hambrey and Huddart 1995; Huddart and


Hambrey 1996; Hambrey et al. 1997; Bennett et al. 1999). In practice, these moraine types probably are representatives within a broad spectrum of deformation styles.

4.2. MORAINE COMPLEXES DOMINA TED BY THRUSTING

4.2.1. Morphology Hummocky moraines, resembling those in Svalbard, have been described from several regions, e.g. Scotland (Sissons 1980), Scandinavia (Hoppe 1952; Andersson 1998) and North America (Eyles et al. 1999). A genetic connotation, that of wastage of debriscovered stagnant ice, has often been attached to hummocky moraines, but from recent work in Scotland (Benn 1992, Bennett and Boulton 1993), active ice margins are thought to have been responsible for the formation of at least some types of hummocky moraine. For this reason, Bennett et al. (1996) advocated the neutral or non-genetic term "moraine mound complex" to describe these landform assemblages. Svalbard moraine complexes commonly comprise arcuate belts of consistently aligned hummocks or mounds of a wide variety of morphological types. They include linear ridges some 100m long, short-crested ridges several metres long, and near-conical mounds; all reaching elevations of several metres. Irrespective of size, before degradation, they show glacier-facing rectilinear slopes with consistent angles of around 30°, irregular distal slopes that are commonly steeper, and stacking of different sedimentary facies (Figure 3). It has been demonstrated that the rectilinear slopes and stacking indicate thrusting in pro glacial, ice-marginal and englacial positions (Hambrey et al. 1997). Following emplacement, ice cores may be present, leading to the degradation of many mounds by mass-movement processes. However, emplacement of thick sediment wedges leads to the survival of the initial moraine morphology, especially if the sediments are not prone to flowage (e.g. well-drained gravels). On the other hand, large-scale degradation can occur if there is substantial buried ice, and if the sediment is composed of flow-prone diamicton.

NW

lliiEil diamicton IZ3 gravel

sand gravel/ g ravally sand sand

Ioe marginal

rIVer outwosh le.,aco C

100m

SE

o

Figure 3. Measured cross-section through moraine complex of Uversbreen, showing nature of thrusting and stacking of different sedimentary facies (from Hambrey and Huddart 1995, with permission of John Wiley and Sons).


4.2.2. Sedimentary facies The sedimentary facies in moraine-mound complexes is as varied as the material over which the glacier flows. Most terrestrial glaciers flow over ground occupied by extensive sheets of glaciofluvial sediment and diamicton of basal glacial origin, and therefore these facies dominate the moraine assemblages (e.g. at Uversbreen: Hambrey and Huddart 1995; Pedersenbreen: Bennett et al. 1996; and KongsvegenlKronebreen: Bennett et al. 1999) (Figure 4b, c). Tidewater glaciers, such as Comfortlesbreen (Huddart and Hambrey 1996) and KongsvegenlKronebreen (Bennett et al. 1999) tend to rework glaciomarine facies, ranging from ice-contact facies (diamicton, coarse gravel), through ice-proximal laminites (cyclopsams and cyclopels) to homogeneous mud with dispersed clasts up to boulder-size of ice-rafted origin (Figure 4d, e). Mixing with terrestrial sediments occurs where tidewater glaciers impinge on land. The wide variety offacies, however, is often organised systematically. A single ridge is commonly composed of one facies, but may be stacked as an inclined slab on another ridge of a different facies (Figure 3). Mapping of facies in ridges enables one to detennine approximately the relative proportions offacies over which the glacier flowed.

It is often possible to detennine the original mode offonnation of the individual facies from the preservation of sedimentary structures, for example cross-bedding and grading in sand and gravel, or lamination with dropstones in fme-grained sediments. Shell debris, including intact fonns are a feature of reworked glaciomarine sediment. Diamicton, on the other hand, being more susceptible to ductile defonnation because of the large mud fraction rarely preserves its original fabric. Typical basal till fabrics, with strong preferred alignment of clasts, are often (though not always) destroyed during emplacement.

4.2.3. Internal structure The internal structure of individual mounds has been documented in a number of cases, and each shows signs of defonnation. In addition to low-angle thrust-faults there are various nonnal faults and even recumbent folds with sheared-off lower limbs (Figure 4t). Such structures develop under strong longitudinal compression. Sometimes, however, the moraines show an earlier phase of longitudinal extension, such as boudinage in fme sediments, indicative perhaps of an early phase of loading prior to thrusting. The degree of defonnation is variable. Sedimentary bedding may survive intact between thrusts, as in the gravels ofUversbreen. Alternatively, bedding may be strongly modified, as in the finegrained glaciomarine sediments of Kronebreen and Comfortlessbreen.

4.3. MORAINE COMPLEXES DOMINATED BY FOLDING

One fme example of a "push moraine", dominated by folding, is that produced by a surge of Holmstrmnbreen (van der Wateren 1995; Boulton et al. 1999). The date of the surge is unknown, but it predates observations made in 1927. Continuous river sections cut through the moraine complex, revealing the internal structure in unique detail.


4.3.1. Effects of surge Holmstmmbreen is characterised by looped moraines indicative of repeated surge behaviour. The most recent surge impinged on an ice-contact scarp of pro glacial fluvial outwash. These sediments were pushed along a decollement to form the moraine complex, and deformation extended for 1.5 Ian beyond the glacier front.

4.3.2. Morphology and structure of moraine complex Van der Wateren (1995) and Boulton et al. (1999) recognised four zones in the push moraine complex:

(1) Exposed glacier, comprising looped patterns of isoclinally folded medial moraines.

(2) Ice-cored moraine zone, comprising a high-relief ridge complex 2 Ian wide underlain by up to 200 m of dead ice, grounded below sea level. Forming a continuous drape on the surface is till (much of it reworked), supraglacial sediment and supraglacial lacustrine deposits. A series of ridges parallel to foliation were interpreted as esker deposits.

(3) The glaCially-pushed sediment zone is a 1.5 Ian wide ridged terrain, forming a broad arc across the valley (Figure 5a). It is underlain by a decollement at a depth of30 m, and the overlying sediments represent a thin frozen nappe structure. This decollement emerges at the outer moraine limit as a "foreland thrust". The zone was emplaced by "pushing", although the accompanying deformation involved folding and thrusting. The ridges reflect the underlying anticlinal structure of the deformed sediments (Figure 5b). The glacially-pushed sediment zone is further subdivided on a structural basis into three (sub)zones (Figure 5c): (i) an external subzone extending glacierwards from the foreland thrust, comprising parallel folds and small-scale thrusts; (ii) an intermediate subzone with assymetric folds and overthrusts; and (iii) an internal subzone dominated by thrust and gliding nappes.

(4) Proglacial outwash zone, 1-3 Ian wide, comprising braided streams on top of coarse gravel fans passing into a fjord-head delta complex.

4.3.3. Sedimentology afmoraine complex A wide range of lithofacies make up the moraine complex. Diamicton is dominant and is interpreted as of basal till origin, though much reworking by sediment flowage is evident. Rhythmically bedded clay, silt and sand are interpreted as glaciolacustrine sediment, mud as glaciomarine sediment, and sand and gravel as glaciofluvial sediment.

4.3.4. Mode and degree of shortening Boulton et al. (1999) identified the underlying decollement as critical in moraine formation. This separates a thin nappe of frozen sediment above and non-frozen sediment beneath. In the latter high artesian water pressure reduced friction to a low value, enabling glide. The style offolding and faulting indicates incremental compressive shortening in the moraine complex of about 900 m; this equates to a net longitudinal compression of 0.58 to 0.46.


48

4b

24 Michael 1. Hambrey and Neil F. Glasser

4c

4d


Figure 4. Moraine-mound complexes with typical facies. (a) General view of moraine-mound complex on Ossian Sarsfjellet, formed by advance of Krone breen, showing rectilinear slopes facing towards glacier (to right), and stacking ofindividual thrust slices. (b) Typical diamicton lithofacies in moraine-mound of the Kronebreen complex on Ossian Sarsfjellet. (c) Glaciofluvial facies in moraine-mound at Finsterwalderbreen. (d) Glaciomarine sand-mud laminites in moraine-mound, Comfortlessbreen. (e) Glaciomarine shelly mud in moraine-mound at Comfortlessbreen. (f) Recumbent fold in moraine-mound at Comfortlessbreen.


4.4. MORAINE COMPLEX RESULTING FROM A DEFORMING BED SURGE

Next to Holmstrmnbreen is a rather different moraine complex formed on the island of Coraholmen as a result of a surge of Sefstrmnbreen across the floor of a fjord (Boulton et al. 1996). The surge took place between 1882 and 1886, and advanced on a carpet of mud, eroded from its original location, transported and smeared over the sea bed and island as deformation till. This surge was characterised by high water pressures and low friction at the bed. The underlying sediment was subjected to repeated folding and homogenisation, represented by three styles of deformation: (i) flat-lying isoclinal folds resulting from subglacial shear; (li) early folds refolded into anticlinal folds with vertical axial planes, representing intrusion of sediment into basal crevasses; and (ii) further refolding into mushroom-like structures with vertical axial planes as a result of pushing and extrusion at the glacier front.

The main landforms observed on the island resulting from the surge include: (i) a zone of basins bounded by rectilinear ridges about 6 m high, resulting from squeezing of subglacial material into crevasses, and (ii) a zone of2 m-high parallel ridges at the outer extremity of the surge, as a result of the extrusion process. The principal sediments that made up the deforming bed were clast-poor muddy and clast-rich sandy diamictons, clay, sand and gravel.

4.5. DEFORMATIONAL SIGNIFICANCE OF DIFFERENT MORAINE STYLES

From the above descriptions, it is clear that there is a continuum of forms of moraine complex, based on displacement characteristics of the glacier bed or the decollement surface. Where friction at the decollement surface is high, then thrust-dominated moraine complexes form (e.g. Comfortlessbreen, UvSrsbreen, Kronebreen).lffriction is reduced, as a result of higher water pressures, then folds are the dominant structures, although small thrusts are also present (e.g. Holmstrmnbreen). Where friction at the glacier bed is very low, as when ice moves over saturated sediment on the sea floor, the deformation in the sediment is represented by polyphase folding alone (Sefstrmnbreen). The application of these principles to Pleistocene moraine complexes will be of value in assessing the nature of terrain over which the glacier moved.

4.6. OTHER CONSTRUCTIONAL LANDFORMS

At least three other types of constructional landforms exist in the proglacial areas of Svalbard glaciers: supraglacial debris trains, foliation-parallel ridges and geometrical ridge networks.

Supraglacial debris trains are produced by the folding of supraglaciaUy-derived debris layers in the ice, as described above. These are visible both on the glacier surface, where they emerge as a result of ablation near the snout, and as regular trains of angular debris extending for considerable distances across the proglacial area (Figure 2a). Commonly, debris-trains drape moraine complexes. Individual debris trains can often be traced to their source areas in the headwall of a glacier, where they are fed by rockfall material. These debris trains are recognisable by their angular, uni-lithological nature and


lack of fine matrix. Mapping and measurement of associated ice structures shows that they form wherever debris is incorporated into the glacier by the folding of primary stratification (Glasser et af. 1999; Hambrey et af. 1999). Debris trains remain prominent following deposition because the large blocks and lack of associated fine sediment fails to support extensive vegetation.

A

B


c

Internal (sub) zone

760

N Intermediate (sub) zone s

N External (sub) zone s

:;: .

lOU

Figure 5. Moraine complex of "push" origin at Holmstrembreen (from van der Wateren 1995, with permission of the Geological Survey of the Netherlands; also published in Boulton el oJ. 1999). (a) Plan of ridge crest-lines within the "glacially pushed sediment zone". (b) Outcrop patterns of sedimentary beds and fault planes; several fold closures and fault displacements can also be seen. (c) Structural styles in the same zone, representing internal (sub )zone with thrust and gliding nappes (top); intermediate subzone with assymetric style recumbent folds and overthrust (middle); and internal subzone with parallel-style folds and small-scale thrusts (bottom).


Foliation-parallel ridges are ridges ofbasally-derived debris (Glasser et al. 1998). They are especially well-developed at Kongsvegen, where significant quantities of basal material are observed on the glacier surface parallel to longitudinal foliation. Although the ridges are commonly 1-2 m wide and up to 1.5 m high, the source debris layers in the ice below are rarely greater than 0.1 m wide. The dispersion of material associated with these features is a result of the melting of their ice core. Some ridges are composed of a clast-rich sandy diamicton, characterised by subangular and subrounded clasts, which are occasionally striated. Lithologically, these foliation-parallel ridges are highly variable. The ridges can often be traced onto the glacier forefield as low «1m high) ridges (Figure 6). The foliation-parallel ridges are important since the incorporation of basal debris along longitudinal foliation is not a universally acknowledged process, and similar ridges elsewhere may have been mistaken as flutes. The mechanism invoked to explain this process is one where lateral compression of ice leads to the development of a transposition foliation parallel to flow, combined with the incorporation of basal debris-rich ice or soft basal sediment in the fold complex (Glasser et al. 1998; Hambrey et al. 1999). A decollement zone must therefore exist beneath the deforming layer, and incorporation of debris must take place where the ice is wet-based. The incorporation of debris by this process clearly precedes most thrusting, as foliation-parallel ridges are truncated by thrust moraines in the pro glacial area. Their preservation potential appears to be low, because of destruction by mass-movement and fluvial processes.

Geometrical ridge networks are created when both longitudinal and transverse debris accumulations melt out of the glacier and become superimposed. This landformlsediment assemblage has been described in the proglacial area of Kongsvegen as a result of the 1948 surge (Bennett et af. 1996). Here, small (4-8 m high) thrust ridges intersect low « 1 m) debris trains or foliation-parallel ridges to form a complex of crosscutting ridges on the glacier forefield. The preservation potential of these networks is probably low, since they are continually modified by slope processes and glacigenic sediment flows. The landform manifestation of structurally-controlled supraglacial and englacial fluvial deposits has not yet been observed in the pro glacial areas of Svalbard glaciers. Their preservation potential is probably low.

s. Discussion

5.1. CONCEPTUAL MODELS OF DEBRIS-ENTRAINMENT AND LANDFORM GENESIS

From the above descriptions, it is apparent that there are clear linkages between the development of glacier structures and the genesis oflandformlsediment assemblages (Table I). In this section conceptual models are described which it is anticipated will fmd wide applicability in other glacierised areas.

The supply of rockfall material around the headwalls of a glacier and its subsequent modification during flow are important, though only recently recognised mechanisms, for the development of medial moraines on the glacier and trains of angular debris beyond (Hambrey et af. 1999). In this process, rockfall is first buried in a wide


accumulation area, then transported in an englacial position (Figure 7). Converging flow occurs where ice funnels into a narrow tongue, resulting in open to tight folding of the stratification (including debris). Fold axes are parallel to flow, and where debris layers reemerge at the surface as a result of ablation, they form longitudinal ridges on the glacier surface. As these ridges are deposited during glacier recession, debris trains of the same orientation are left behind on the proglacial area.

Where the glacier is sliding on its bed, debris is incorporated during the regelation process, giving rise to the thick basal debris layer that is characteristic of poly thermal glaciers. Layers are repeated by shearing and isoclinal folding. Unless modified by other processes, such as thrusting, deposition will ultimately give rise to a sheet of diamicton that has the characteristics of lodgement till or melt-out till (Boulton 1970b). This sheet may cover extensive areas of the lower parts of the proglacial area, though often it is heavily dissected by rivers.

Figure 6. Foliation parallel ridges at the snout of Kongsvegen, where they are intersected by thrust ridges.

Converging flow may also result in tight folding of the basal or subglacial debrislayer so tightly that it is transposed within longitudinal foliation in the lower part of the glacier; this requires the bed to be soft and a decollement surface to form at the bedrock/sediment interface. Through ablation, especially where folding is intense (as in association with medial moraines or at the margins), the folded basal debris layers become exposed at the glacier surface (Figure 8). On deposition, they leave behind foliationparallel ridges.


Table 1. Structurally controlled landforms in order off ormation associated with maritime hIgh-arctic glaciers.

Relaltlnl Glacier structure landform

I Folding of s1ndificaIim wi1h SlJII8glaciaUy derived debris; associated l.oogitudinal debris sIripes wi1h axial planar folialioo of 100gitudinal crienIIIion

2 Basal debris layer Sheets of diamicton 3 Basal debris ~ in lqiIudinal folialion FoJiaIioo..panlIel ridges 4 11Irustq in zone of IIlIIIsiIion ftom a wet to Iiozen bed MDiDtnowII (or "thrust moraioc")

complex 5 Glacier ovmi~ dcli:mab1e bed IIIIl iJmporaIing satimaIt by 'PushIllll11in:s'

folding; basal dtcollemeIi in IIIIlerlying sedimai 6 TbrustIfolding-derival glaciolluvial sedinent reIlIIkal imo

lOOIlitOOinal~ lAqibllinal ridge sedimai acmulaIiooI

• IAlw

•• • •• <-_.->

•••• • •••• High

Relative Importance In landform a.semblale

•••

... • •••••

•••

.

Thrusting in polythermal glaciers is one of two mechanisms that gives rise to the most prominent of all landform assemblages in the proglacial areas of polythermal glaciers (Figure 9). However, debris incorporation by this process is non-uniform, but where substantial wedges of subglacial sediment are detached from the bed, particularly at the transition from sliding to frozen-bed conditions, the resultant depositional product is a transverse ridge of debris on the ice surface and a prominent moraine mound following recession. Such landforms may originate englacially, so are not indicative of ice margins. Similar landforms result from thrust-propagation into the proglacial area, sometimes as much as -100m; in other words, they form beyond the glacier margin, so these too do not represent true ice-marginal positions. Thrusts are commonly associated with recumbent folding, and intersect earlier structures such as folds with longitudinal axes and longitudinal foliation.

Boulton et al. (1996, 1999) have described other types of moraine complex ("push moraines") which do not show the same close linkage with internal structures in the glaciers. Rather the deformation is mainly subglacial or proglacial. However, all of these moraine types seem to be part of a continuum of forms, and the full suite of types has probably yet to be defmed.

The final landform/sediment assemblage is the result of englaciallsupraglacial stream reworking of glaciofluvial material already ingested by folding and thrusting. Where this process occurs, large quantities of glaciofluvial sediment can become mixed with basal glacial debris and elevated to the ice surface, where this sediment is further mixed with supraglacial debris. A diverse range of clast and particle-size characteristics results, and facies associations are complex. Although large and robust landforms are rarely preserved by this process, the mixing of facies is an important sedimentological characteristic.


~ ,

C ,.

i~ .• ' C'

:~ ., "

, :..: .,', ,; ... ,. ]

(b) Cross-sections

CVC'

]

Rockfall from ridges bordering upper reaches of glacier. forming spreads of supraglacia/ debris; bUrial of debris

Converging flow and initiation of folding of stratification and supraglacial debris layers

Zone of strong lateral compression giving rise 10 tight open folds in • stratification and debris layers; development of axial plane foliation of longitudinal orientation

Emergence of 'medial moraines' as strongly folded layers of supraglacial debris derived from upper reaches of accumulation area

JfV ridges bounding upper glacier basin

~B rOCkfall debris derived ~v. from ridges

1 flow lines

~~ stratification

:: -'l debris concentrations

~ associated with stratification

Figure 7. Conceptual model of a glacier with multiple basins feeding a narrow tongue, showing how progressive fo1dingofsupraglacially derived debris leads to the development of medial moraines. Deposition of these moraines will produce debris-trains in the proglacial area (from Hambrey et al. 1999, with permission of the International Glaciological Society).

Landform and Sediment Assemblages

stralification with supraglacial debris

\ I I I \ I I I longitudinal \ ,I ! loliation

m,.~ .... rockfall·derived debris

basally-derived debris

c.2oom

Vertical exaggerallon approx. x 4

33

Figure 8. Inferred typical cross-section of the lower part ofa typical Svalbard valley glacier, illustrating fold styles where debris is so strongly folded that it reaches the surface. Ultimate deposition of the basal-debris zone will form a foliation-parallel ridge (from Hambrey et al. 1999, with permission of the International Glaciological Society).

FLOW

I frQUn to bed I

Cross section GlMlot '"",fIold

BT

Figure 9. Schematic three-dimensional view of the snout of a typical terrestrial valley glacier in Svalbard, showing the relationships between folded stratification, longitudinal foliation, thrusting and debris distribution, and their manifestation on the glacier forefield (from Hambrey et af. 1999, with permission of the International Glaciological Society).


5.2. COMPARISON WITH YOUNGER DRYAS MORAINE COMPLEXES IN BRITAIN

Previous models of glacial processes during the Younger Dryas in Britain have drawn heavily on alpine or Icelandic, as opposed to Arctic, analogues to explain the broad belts of "hummocky moraine" that are associated with this time period (see review in Bennett 1994). Since the pioneering work of Boulton (1972) there have been few attempts to link depositional processes at modem arctic glaciers to those in the Pleistocene landform record, and it is only recently that interest in this subject has re-emerged. Comparisons between moraine mound complexes at modem High-Arctic glaciers and those formed during the Younger Dryas in Britain have been made at Cwm Idwal in Snowdonia, North Wales (Hambrey et al. 1997; Graham and Midgley in press) and in the Valley of a Hundred Hills (Coire a' Cheud-chnoic) in the Northwest Highlands of Scotland (Bennett et al. 1998) (Figure 10). A generalised model for these features is depicted in Figure 11. These comparisons stress the similarities between moraine mound complexes formed by Svalbard glaciers and British "hummocky moraine" in terms of moraine mound morphology, sedimentology and facies variability. Other localities in Britain where Younger Dryas moraines exist remain to be reinvestigated. Understanding of the genesis of similar hummocky moraine assemblages in North America may also benefit from these comparisons.

Structural glaciological principles and their role in debris transport and deposition at modem Svalbard glacier margins can also be applied to other British Younger Dryas landforms and sediments. For example, the recognition of foliation-parallel ridges at modem glacier margins has implications for the interpretation of Pleistocene landforms, in particular areas of fluted terrain. Two broad types of fluted terrain are commonly recognised (Heikkinen and Tikkanen 1979; Rose 1987; Gordon et 01. 1992). These are:

(1) flutes, which are low (<3m), narrow (<3m) regularly spaced ridges which are usually less than 1 00 m long and are common in front of modem glacier margins, and

(2) megaflutes - which are much taller (>5m), broader and longer (> 100m). Flutes have been widely interpreted as forming in the lee of subglacially-lodged boulders (Dyson 1952; Hoppe and Schytt 1953; Schytt 1963; Paul and Evans 1974; Boulton 1976), although several authors have challenged the ability of this mode of formation to explain all the observed forms (Karlen 1981; Rose 1989; Gordon et al. 1992). The formation of megaflutes is less well understood, and for the most part these features do not appear to be associated with cavities in the lee of boulders or bedrock obstacles. Several authors have suggested a genetic link between megatlutes and drumlinoid forms (Gravenor and Meneley 1958; Prest 1968; Aario 1977; Rose 1987; Clark 1993).

The foliation-parallel ridges at Kongsvegen add to the range of mechanisms capable of producing low, linear ridges parallel to ice-flow. It is therefore possible that some of the features previously interpreted as glacial flutes may, in fact, have their origin as foliation-parallel ridges. Both types of ridge are composed of subglacial debris, and thus may be impossible to distinguish from each other, since the fabric induced by folding into basal ice may be similar to that associated with sediment flow into a lee side cavity. More observations are required concerning the range of size, length and morphology offoliation-


parallel ridges. Until these data become available it is only possible to highlight the importance of this mechanism as a possible alternative when interpreting such landfonns.

Other linear features of importance in the modem Svalbard glacier setting are debris-trains comprising angular debris, which drape the proglacial area, unless reworked by streams. These also may have been mistaken as flutes, but the present authors have observed a number of possible candidates of Younger Dryas age in cirques of the Northwest Highlands ofScotIand (e.g. in Coire Mhic Fearchair on Beinn Eighe) (Figure 12).

Geometrical ridge networks created during glacier surges, such as that of Kongsvegen are different from crevasse-fill ridges (Sharp 1985) so are important for interpretation of the landfonn record, butno Pleistocene examples have yet been identified. Crevasse-fill ridges have been reported from Svalbard, however, notable at Sefstrmnbreen (Boulton et aJ. 1996).

5.3. PALAEOCLIMATIC IMPLICATIONS OF SEDIMENTILANDFORM ASSEMBLAGES

Large moraine complexes, which fonned by thrusting and folding processes, are a characteristic feature of the polythennal glaciers of Svalbard, but are less commonly associated with temperate glaciers. It is therefore tempting to suggest that similar landfonns in Britain, associated with the Younger Dryas event, were fonned under similar climatic conditions, with the proviso that debris-bearing thrusts can occur at temperate glaciers if topographic conditions (such as movement against a reverse slope) are favourable (Hambrey et aJ. 1997). This contradicts earlier palaeoclimate inferences which assumed temperate glaciers (e.g. Payne and Sugden 1990; Thorp 1991). However, periglacial evidence and evidence of pennafrost also suggests a Svalbard type climate was more likely in the Younger Dryas (Ballantyne and Harris 1994). The dynamic nature offonnation of the Svalbard moraine complexes, and by analogy some of those in Britain (as also suggested by Benn 1992 and Bennett and Boulton 1993), also leads us to dismiss the icestagnation hypothesis, and with it the inference of abrupt climatic warming (Sissons 1980). Re-examination of similar moraine complexes elsewhere, that have been interpreted as ice stagnation phenomena, may be beneficial in the light of evidence from Svalbard.

Other landfonns and sediments are less likely to be diagnostic of palaeoclimate than the moraines, but the following points may be made: (i) diamicton of basal ice or debris-flow origin is much more abundant at polythennal glaciers than at temperate ones, as reworking by streams is less pronounced at the fonner; (ii) the basal debris load is great compared with the supragtacialload at polythennal glacierS, while the reverse is true for temperate glaciers; (iii) the fonnation of debris stripes and foliation-parallel ridges is a feature of polythennal glaciers, but they do not appear to feature in descriptions of temperate glaciers, possibly because of reworking by streams.


lOb

Figure 10. Comparison of thrust-related moraine-mound complexes, (a) at the margins ofKronebreen on Ossian Sarsfjellet, and (b) in Coire a' Cheud-chnoic, Torridon, NW Scotland (b).

Landfonn and Sediment Assemblages

o Advancing glacier at Neoglacial limit

cliff-fall thrust propagation debris

into pregtacial area

debris-bearing thrusts with recumbent folding

c e r

® Present-day receding glacier of Svalbard type

outer

t1wJl basal debris + subglacially derived englacial debris

E;;:M;\j sand and gravel 01 fluvial origin

f: .... j diamicton liHle·affected by thrusting

• diamicton thrust slices

~ bedrock

~ folded basal Ice with debris

@ sale thrust

debris-bearing thrusts with recumbent folding

37

Figure 11. Conceptual model for the development of moraine-mound complexes ("hummocky moraines") of Younger Dryas age in highland areas of Britliin. based on thrusting models developed for polythermal glaciers in Svalbard (from Hambrey et al. 1997. with permission of the Geological Society of London).


Figure 12. Debris-stripe probably fonned by folding of asupraglacial debris-layer into a medial moraine. This example is represented by white Cambrian quartzite in an area dominated by the red Proterozoic Torridonian Sandstone, Coire Mhic Fhearchair, NW Scotland.

6. Conclusions

Landfonns associated with the polythennal glaciers of Svalbard are not in themselves unique, but the particular assemblages and proportions of different sedimentary facies are, in general, different from temperate and cold polar glacier systems. The major conclusion from this study is that defonnation within glacier ice, as debris is entrained and subsequently transported, is a primary control on the nature of landfonn and sediment assemblages in the proglacial areas of valley glaciers.

Several modes of debris entrainment have been distinguished in Svalbard glaciers, that result in certain landfonns:

(i) Incorporation of angular rockfall material within the stratified sequence of snow/frrnlsuperimposed ice. This debris takes an englacial path through the glacier, becoming folded with axes parallel to flow and often in axial planar relationship with foliation. Near the snout the debris emerges at the surface on the hinges and limbs of the folds, producing medial moraines which on deposition give rise to flow-parallel debris trains.

(ii) Entrainment of debris at the bed to fonn the basal ice layer, involving primarily regelation, bulk freezing-on, folding and shearing. In High-Arctic glaciers, these processes can result in a debris layer several metres thick within the glacier and a sheet of basal till on deposition.


(iii) Incorporation of debris of basal character within longitudinal foliation. The foliation is the product of shearing or very tight folding within the ice. The resulting landforms are foliation-parallel ridges.

(iv) Thrusting, whereby basal glacial sediments (including regelation ice) and subglacial sediments are uplifted into an englacial position, sometimes emerging at the ice surface. Thrusting is a dynamic process, and in polythermal glaciers may be linked mainly to the transition from sliding to frozen bed conditions. This process is one of two that produces the large end moraine complexes that typify Svalbard glaciers.

(v) Subglacial upright folding with transverse axes and faulting of sediment as a result of a glacier overriding a deformable bed. This process also produces large moraine complexes in Svalbard.

(vi) Reworking of thrust- or fold-derived glaciofluvial material to produce longitudinal debris ridges within and on the ice. The resulting landforms are, however, ephemeral.

In terms of deposits laid down by polythermal glaciers, at least on land, the most important are diamictons, which typically would be interpreted as lodgement or meltout tills. The release of this sediment is controlled by the thermal regime of the glacier (the ice must be at the pressure melting point at the bed), the distribution of debris in the basal layer, and whether the bed beneath the glacier is deformable. Once released, much of this debris is reworked by glaciofluvial processes, but not to the same extent as with temperate glaciers.

The principal landforms resulting from ice-deformational processes are moraine mound complexes (also referred to as thrust or push moraines where their genesis is known). The deformed sediments that commonly make up these complexes are highly variable and include diamicton of subglacial derivation, sand and gravel of glaciofluvial origin, and mud with scattered gravel clasts and laminites of fjordal origin (since many Svalbard glaciers end partly in tidewater).

The application of these observations and inferences to areas where glaciers no longer exist is still at an early stage. However, there are clear similarities in terms of morphology and structure between some Younger Dryas moraines in highland Britain with those in Svalbard. Reappraisal of the ice-dynamic and climatic conditions for that cold interval are clearly necessary. Other landforms are more subtle, and are not expected to have a high preservation potential, but a handful of localities where folded supraglacial debris trains and foliation-parallel ridges occur have already been noted in Scotland.

The application of structural glaciological principles needs to be extended to the interpretation of landforms associated with glaciers elsewhere. It is clear from our own preliminary observations in temperate glaciers in the Swiss Alps and Patagonia that this will be a fruitful line of investigation.

Acknowledgements

We thank our colleagues Matthew Bennett, Julian Dowdeswell, David Huddart and Tavi Murray for several jointly productive seasons in unravelling the nature of debris transport and depositional processes associated with Svalbard glaciers. Financial support was provided by the U.K Natural Environment Research Council (Grant Nos. GR9/02185 and


GST/03/2192 the European Union Environment and Climate Programme (Grant No. EN5V-C793-0299), and Liverpool John Moores University. For logistical support we thank especially Nick Cox, Base Manager at the NERC Arctic Research Station in NyAlesund.

References

Aario, R. (1977) Associations of flutings, drumlins, hummocks and transverse ridges, Geojournall, 65-72. Andersson, G. (1998) Genesis of hummocky moraine in the Bolmen area, southwestern Sweden, Boreas 27, 55-

67. Ballantyne, C. K. and Harris, C. (1994) The periglaciation o/Great Britain, Cambridge University Press,

Cambridge. Bamber, J. L. (1989) 'Icelbed interface and englacial properties of Svalbard ice masses deduced from airborne

radio echo-sounding', Journal o/Glaciology 35, 30-37. Benn, D. I. (1992) The genesis and significance of "hummocky moraine": evidence from the Isle of Skye,

Scotland, Quaternary Science Reviews 11, 781-799. Bennett, M.R. (1994) Morphological evidence as a guide to deglaciation following the Loch Lomond

Readvance: a review of research approaches and models, Scottish Geographical Magazine, 110,24-32.

Bennett, M. R. and G. S. Boulton (1993) A reinterpretation of Scottish "hummocky moraine" and its significance for the deglaciation of the Scottish Highlands during the Younger Dryas or Loch Lomond Stadial, Geological Magazine 130,301-308.

Bennett, M.R., MJ. Hambrey, D. Huddart and J.F. Ghienne (1996a) The formation of a geometrical ridge network by the surge-type glacier Kongsvegen, Svalbard, Journal o/Quaternary Science 11, 437-449.

Bennett, M.R., M.J. Hambrey, D. Huddart and J.F. Ghienne (1996b) Moraine development at the high-arctic valley glacier Pedersenbreen, Svalbard, Geogrqfiska Anna/er 7SA, 209-222.

Bennett, M.R., Hambrey, M.J., Huddart, D. and Glasser, N.F. (1998) Glacial thrusting and moraine-mound formation in Svalbard and Britain: the example of Coire a' Cheud-chnoic (Valley of a Hundred Hills), Torridon, Scotland Quaternary Proceedings, 6, 17-34.

Bennett, M.R., Hambrey, M.J., Huddart, D., Glasser, N.F. and Crawford, K.R. (1999) The landform and sediment assemblage produced by a tidewater glacier surge in Kongsfjorden, Svalbard, Quaternary Science ReViews, 18,1213-1246.

Bjomsson, H., Y. Gjessing, S-E. Hamran, J.~. Hagen, O. Liestel, F. P8Isson and B. Erlingsson (1996) The thermal regime of sub-polar glaciers mapped by multi-frequency radio-echo sounding, Journal 0/ Glaciology 42, 23-32.

Boulton, G. S. (1970a) On the origin and transport of englacial debris in Svalbard glaciers, Journal 0/ Glaciology, 9, 213-229.

Boulton, G. S. (I 970b ) The deposition of subglacial and meltout tills at the margins of certain Svalbard glaciers, Journal o/Glaciology 9, 231-245.

Boulton, G. S. (1972) Modern Arctic glaciers as depositional models for former ice sheets, Quarterly Journal o/the Geological Society o/London 128,361-393.

Boulton, G. S. (1976) Origin of glacially fluted surfaces, Journal o/Glaciology 17, 287-301. Boulton, G. S. (1978) Boulder shape sizes and grain-size distributions of debris as indicators of transport paths

through a glacier and till genesis, Sedimentology 25, 773-799. Boulton, G. S., Van der Meer, J.J.M., Hart, J., Beets, D., Ruegg, G.H.J., Van der Wateren, F. M. and Jarvis J.

(1996) Till and moraine emplacement in a deforming bed surge - an example from a marine environment, Quaternary Science Reviews IS, 961-987.

Boulton, G. S., Van der Meer, J.J.M., Beets, D. J., Hart, J. K. and Ruegg, G.H.J. (1999) The sedimentary and structural evolution of a recent push moraine complex: HolmstrOOlbreen, Spitsbergen, Quaternary Science Reviews 18, 339-371.

Clapperton, C.M. (1975) The debris content of surging glaciers in Svalbard and Iceland, Journal o/Glaciology, 14,395-406.


Clark, C. D. (1993) Mega-scale glacial lineations and cross-cutting ice-flow landfonns, Earth Sur/ace Processes and Lan4forms 18, 1-29.

Dowdeswell, I.A., Hamilton, G.S. and Hagen, 1.0. (1991) The duration of the active phase of surge-type glaciers: contrasts between Svalbard and other regions, Journal o/Glaciology 37,388-400.

Dyson, J. L. (1952) Ice-ridged moraines and their relation to glaciers, American Journal 0/ Science 250, 204-211.

Eyles, N., Boyce, J. I. and Berendregt, R. W. (1999). Hummocky moraine: sedimentary record of stagnant laurentides Ice Sheet lobes resting on soft beds, Sedimentary Geology 123, 163-174.

Glasser, N. F., Hambrey, M. J., Crawford, K. R., Bennett, M. R. and Huddart, D. (1998) The structural glaciology ofKongsvegen, Svalbard and its role in landfonn genesis, Journal o/Glaciology 144, 136-148.

Glasser, N.F., Bennett, M.R. mll\uddart, D. (1998) Ice-owginal dwacteristics ofFridIjcMJreen(Svalbanl) during ils recenlsurge, Polar Research 17,93-100. Glasser, N.F., Bennett, M.R. and Huddart, D. (1999) Distribution of glaciofluvial sediment within and on the

surface of a high arctic valley glacier: Marthabreen, Svalbard, Earth Surface Processes and Lan4forms, 24, 303-318.

Graham, D. 1 and Midgley, N. (2000). Moraine-mound formation by englacial thrusting: the Younger Dryas moraines ofCwm Idwal, NorthWales, in Maltman, A. J., Hubbard, B. and Hambrey, M. 1. (eds.) De/ormation o/Glacial Materials, Spec. Pub!. Geological Society of London, 176,321-336.

Gordon, J. E., Whalley, W. B., Gellatly, A. F. and Vere, D. M. (1992) The fonnation of glacial flutes: assessment of models with evidence from Lyngsdalen, North Norway, Quaternary Science Reviews 11, 709-731.

Gravenor, C. P. and Meneley, W. A. (1958) Glacial flutings in central and northern Alberta, American Journal o/Science 256,715-728.

Hagen, 1.0. (1987) Glacier surge at Usherbreen, Svalbard, Polor Research 5, 239-252. Hagen, J.~. (1988) Glacier surges in Svalbard with examples from Usherbreen, Norsk Geogra/lsk Tidsskrift,

42,202-213. Hagen, 1. O. and Saetrang, A. (1991) Radio-echo soundings of sub-polar glaciers with low-frequency radar,

Polar Research 26, IS-57. Hagen, J. D., Korsen, O. M. and Vatne, G. (1991) Drainage pattern in a subpolar glacier: Brsggerbreen,

Svalbard. In Gjessing, Y., Hagen, J. D., Hassel, K. A., Sand, K. and Wold, B. (eds) Arctic hydrology. Present and Future Tasks. Oslo, Norwegian National Committee for Hydrology Report No. 23.

Hagen, J. D., Liesml, D., Roland, E. and Jorgensen, T. (1993) Glacier atlas 0/ Svalbard and Jan Moyen, Oslo, Norsk Polarinstitutt Medd. 129.

Hambrey, M. 1. and Dowdeswell, 1. A. (1997) Structural evolution of a surge-type polythennal glacier: Hessbreen, Svalbard, Annals o/Glaciology, 24, 375-381.

Hambrey, M. 1. and Huddart, D. (1995) Englacial and proglacial glaciotectonic processes at the snout of a thennally complex glacier in Svalbard, Journal oiQuaternary Science 10,313-326.

Hambrey, M. 1. and MUlIer, F. (1978) Structures and ice defonnation in the White Glacier, Axel Heiberg Island, Northwest Territories, Canada, Journal o/Glaciology 20,41-66.

Hambrey, M. J., DowdernclI,J. A., MlIIlIIY,1. mlPorter, P. R. (1996)ThrustingmJ debris-em'ainmenl ina surging glacier. BakaniMreen, SVIIbard,Annals 0/ Glaciology 22,241-248.

Hambrey, M. I., Huddart, D., Bennett, M. R. and Glasser, N. F. (1997) Genesis of 'hummocky moraines' by thrusting in glacier ice: evidence from Svalbard and Britain, Journal o/the GeolOgical Society 153, 623-632.

Hambrey, M. J., Bennett, M. R., Dowdeswell, J. A., Glasser, N. F. and Huddart, D. (1999). Debris-entrainment in polythennal valley glaciers, Svalbard, Journal o/Glaciology 45, 69-86

Hamilton, G. S. and Dowdeswell, J. A. 1996. Controls on glacier surging in Svalbard, Journal o/Glaciology, 42, 157-168.

Heikkinen, O. and Tikkanen, M. (1979). Glacial flutings in northern Finnish Lapland, Fennia 15,1-12. Hodgkins, R. (1997) Glacier hydrology in Svalbard, Norwegian High Arctic,. Quaternary Science Reviews 16,

957-973. Hooke, R. L. (1973) Flow near the margin of the Barnes Ice Cap, and the development ofice-cored moraines,

GeolOgical Society 0/ America Bulktin 84, 3929-3952. Hoppe, G. (1952) Hummocky moraine regions, with special reference to the interior ofNorbotten, Geogrqfiska

Annaler34,1-72. Hoppe, G. and Schytt, V. (1953) Some observations on fluted moraine surfaces, GeograflSka Annaler 35, 105-

111.


Huddart, D. and Hambrey, M. J. (1996) Sedimentary and tectonic development ofa high-arctic, thrust-moraine complex: ComfortJessbreen, Svalbard, Boreas, 25, 227-243.

Huddart, D., Bennett, M. R., Hambrey, M. J., Glasser, N. F. and Crawford, K. R. (1998) Origin ofweU rounded gravels in glacial deposits from Brsggerhalveya, northwest Spitsbergen: potential problems caused by sediment reworking in the glacial environment, Polar Research, 17, 61-69.

Karl~n, W. (1981) Flutes on bare bedrock, Journalo/Glaciology 27, 190-192. Knight, P.G . (1997) The basal ice layer of glaciers and ice sheets. Quaternary Science Reviews, 16, 975-993. Lawson, W. (1996). Structural evolution of Variegated Glacier, Alaska, USA, since 1948, Journal 0/

GlaCiology 42, 261-270. Meier, M. F. and Post, A. (1969) What are glacier surges? Canadian Journal 0/ Earth Science, 6,807-817. Melvold, K. and Hagen, J. O. (1998) Evolution of a surge-type glacier in its quiescent phase: Kongsvegen,

Spitsbergen, 1964-95, Journal o/Glaciology, 44, 394-404. Murray, T., Gooch, D. L. and Stuart, G. W. (1997) Structures within the surge front at Bakaninbreen, Svalbard,

using ground-penetrating radar, Annals o/Glaciology, 24, 122-129. 0degard, R. S., Hamran, S. E., Bs, P. H., EtzelmOller, B., Vatne, G. and Sollid, J. L. (1992.) Thermal regime

of a valley glacier, Erikbreen, northern Spitsbergen, Polar Research, 11, 69-79. Paul, M.A. and Evans, H. (1974) Observations on the internal structure and origin of some flutes in glaciofluvial

sediments, Blomstrandbreen, north-west Spitsbergen, Journal o/Glaciology 13,393-400. Payne, A. J. and Sugden, D. E. (1990) Topography and ice sheet growth , Earth Surface Processes and

Land/orms 16, 625-639. Prest, V.K. (1968) Nomenclature o/moraines and ice-flow features as applied to the glaCial map o/Canada,

Geological Survey of Canada Paper 67-57. Rose, J. (1987) Drumlins as part ofa glacier bedform continuum. In Menzies, 1 and J. Rose, eds. Drumlin

Symposium, Rotterdam, A A. Balkema, 103-116. Rose, J. (1989) Glacier stress patterns and sediment transfer associated with the formation of superimposed

flutes, Sedimentary Geology 62, 151-176. Schytt, V. (1963) Fluted moraine surfaces, .Journal o/Glaciology 4,825-827. Sharp, M. (1985) "Crevasse-fiU" ridges - a landform type characteristic of surging glaciers? Geograjiska

Annaler 67A, 213-220. Sissons, 1 B. (1980) Palaeoclimatic inferences from Loch Lomond Advance glaciers. In: Lowe, J. 1, Gray, 1

M. and Robinson, J. E. (eds.) Studies in the Late Glacialo/North West Europe, Pergamon Press, Oxford, 31-43.

Solheim, A and Pfirman, S. L. (1985) Sea-floor morphology outside agrounded, surging glacier: BrAsveUbreen, Svalbard, Marine Geology, 65, 127-143.

Swinzow, G. K. (1962) Investigation of shear zones in the ice sheet margin, Thule area, Greenland, Journal 0/ Glaciology 4,215-229.

Thorp, P. W. (1991) The glaciation and glacial deposits of the western Grampians. In: Ehlers, J., Gibbard, P. L. and Rose, J. (eds.) Glacial deposits o/Great Britain and Ireland, Balkema, Rotterdam, 137-149.

Van der Wateren, F. M. (1995) Structural geology and sedimentology of push moraines, Mededelingen Rijks Geologische DienstNr. 54,168pp.

Vatne, G., EtzelmQUer, B., 0degard, R. S. and Sollid, J. L. (1996) Meltwater routing in a high arctic glacier, Hannabreen, northern Spitsbergen, Norsk Geogrqfisk Tidsskrift 50,67-74.

Weertman, J. (1961) Mechanism for the formation of inner moraines found near the edge of cold ice caps and ice sheets, Journal o/Glaciology 3,965-978

PROGLACIAL AND PARAGLACIAL FLUVIAL AND LACUSTRINE ENVIRONMENTS IN TRANSITION

Peter G. Johnson Department of Geography University of Ottawa Ottawa, ON KIN6N5

ABSTRACT

Geomorphology has two interlocking paradigms; the first is a process paradigm where there is a hierarchy of knowledge through the physics of a medium, the mechanics of process, to the landscape form; the second is a temporal paradigm from the history of the landscape, to the present condition, and with a prediction capacity. These are central to concepts of landscape transitions and landform assemblages. The paper explores landscape transitions occurring in proglacial and paraglacial environments based on the horizontal link of sediment transfer rather than on the traditional vertical division into glaciology, hydrology and periglacial geomorphology. In glacierized mountain areas the superimposition of transitions at different scales can readily be demonstrated within the regional landscape. Transitions may occur slowly, rapidly or instantaneously but their frequency of change will vary depending on the underlying cause of the transition. The concept of landform assemblages should integrate the ideas of process activity at different spatial, quantitative and temporal scales. Landform assemblages are therefore infinitely variable. The ways in which the concepts of transition and landform assemblages influence geomorphological interpretation is illustrated by reference to the glacierized environments of the southwest Yukon.

I. Introduction

Traditionally the science of geomorphology has been organised according to specific environments such as glacial, fluvial, periglacial, and coastal, which can be thought of as a vertical subdivision of the discipline. A survey of the bookshelves indicates that virtually all introductory textbooks are organised this way and many advanced level texts are similarly aligned. Sediment transfer, as one basic process concept of geomorphology, provides a horizontal link uniting the vertical components of the science. Geomorphology can, in fact, be defmed by sediment transfer from bedrock to geosyncline and the temporary landscapes which are produced during this transfer. One result of the vertical division has

43 K. Hewitt et al. (eds.), Landscapes a/Transition, 43-62. ID 2002 Kluwer Academic Publishers.

44 Peter Johnson

been that some domains often receive very little consideration despite their importance in the landscape. The whole of the earth's surface is made up of slopes, yet there is often no strong focus on slope processes. Lacustrine environments often receive even less attention than slopes despite the fact that they are integral parts of the fluvial environment. Indeed, in some countries such as Canada, lakes are very critical components of the landscape and the hydrological cycle.

Sediment transfer can be conceptualised within the framework of the hydrological cycle in which both the presence and absence of water are important and the various phases of water (vapour, liquid and solid) are critical elements. In high latitude and high altitude environments ice plays a major geomorphic role. This is not only in glacial and periglacial environments but also in fluvial, lacustrine and oceanic environments. In traditional vertical approaches to geomorphology this century, glaciation has received considerable attention and periglaciaVpermafrost has received prominence in the last 30 to 40 years. Indepth treatment of ice on rivers, lakes, and oceans is however still lacking in most standard physical geography and physical geology texts.

The landform assemblages produced by erosion, transport and deposition represent different periods of evolution characterised by changes or variations in the movement and storage of sediment. Some storage forms may last only for a few minutes or hours while others may persist for centuries or millennia. Landscapes are essentially, however, always in transition and this can be demonstrated at a continuum of process scales varying both spatially and temporally. Meso-scale landform assemblages such as drumlin fields may be relatively stable at a meso-time scale, but are still subject to micro-scale processes. Transitions may be considerably slower but they are still taking place. Microscale processes operating over short time frames at the termini of glaciers are superimposed on the macro-scale processes such as periods of glacier advance and retreat. Landform assemblages associated with stages of evolution of a landscape can also be classified on different spatial scales and may be in transition at different temporal rates.

1.1. PREDICTION AND PARADIGMS

The need for prediction is implicit in the concept oflandscapes, or landform assemblages, in transition. Prediction oflandscape change is becoming more critical as various scenarios of climate change or climate variability are evaluated. Questions such as what changes might occur due to 2 X CO2 in areas with massive ground ice accumulations in high latitudes, what will be the hydrological response to changes in snowfall patterns as a result of global warming, and what will be the impact on landform forming processes such as debris flow and landslides, have become political as well as scientific questions. EI Niflo has become the popular explanation for climate variability and recent extremes, but is EI Nino related to global warming and what are the potential effects of changes in its frequency and magnitude? As an example much discussion has centred recently, even in the popular press, on the increased number of hurricanes and their geomorphic effect, particularly in Central America, which has occurred during a strong El Nino year.

Proglacial and Paraglacial Fluvial and Lacustrine Environments 45

I have emphasised in teaching for a number of years that geomorphology has two interlocking paradigms. The first is a process paradigm where there is a hierarchy of knowledge required:

Physics of Medium - Mechanics of Process - Form of Landscape

It is essential to have a comprehension of materials and mechanics before one can attempt to interpret landforms. Inference from form is not acceptable and landform, or landform assemblage, production must be demonstrated to be physically possible. A coastal geomorphologist has to understand wave theory and a glacial geomorphologist has to understand the physics and chemistry of ice in order to understand the evolution of the landscape.

An excellent example of the problems that have been caused by interpretation based on form is presented elsewhere in this volume with respect to the interpretation of diamicts in the valleys of the Karakoram and the Himalaya. Diamicts form a large number of cross valley ridges in the Karakoram. These landforms, long interpreted as moraines on the basis of their cross-valley ridge form, are now being interpreted as landslide deposits. This change in interpretation has occurred due to an understanding of the physics and mechanics of mass movements and more rigorous field investigation. It will necessitate a fundamental change in the interpretation of the landscape of a major mountainous region. The second paradigm is temporal:

The History of Landscape Development. (e.g. Quaternary Geomorphology)

• TRANSITIONS

• Contemporary Process and Landscape. (Need not be an analog for the past or the future)

• TRANSITIONS 4

Prediction of Landscape Change. (Under scenarios of climate change or variability)

This does not imply that every geomorphologist has to be concerned with all three components but it emphasises that there must be a predictive capability within the discipline, and there must be a capability to understand transitions, which develops from a comprehension of past and present environments solidly based on the first paradigm. Slaymaker and Spencer (1998) have developed this argument in their discussion on refocusing physical geography.

The next sections of this paper explore transitions occurring in proglacial and paraglacial environments using field examples drawn principally from the region of the St. Elias Mountains and its boundary ranges in the southwest of the Yukon Territory. The

46 Peter Johnson

principles developed from these observations may be applied to any glacierised or glaciated region.

In glacierised mountain areas the superimposition of transitions at different spatial and temporal scales can readily be demonstrated within the regional landscape. Transitions may occur slowly, rapidly or instantaneously but their frequency of change will vary depending on the underlying cause of the transition. Subglacial conduits may change from hour to hour, day to day, or season to season. Ice-marginal lakes may be transient or may persist in the landscape, but even in the latter case they may experience transitions in their sedimentary environments as, for example, the glacial source of sediment becomes more distant from the lake, or as temperature contrasts between river and lake environments change.

1.2. THE HYDROLOGICAL SYSTEM

Sediment transfer models (Statham 1977), and the landscape transitions which result from sediment transfer, can be organised within the broad framework of the hydrological system. The absence of water must be considered as part of the hydrological system just as the occurrence of the solid form, ice, is a very important element in many environments. Unfortunately few discussions of the hydrological cycle emphasise high latitude and high altitude environments. The stimulating new approaches to river basins by Calow and Petts (1992, 1994) and Petts and Amoros (1996) need to be expanded to include the roles of river ice and lake ice (Prowse and Gridley 1993). The hydrological system of glacial and proglacial environments is characterised by extreme variability both spatially and temporally, and frequently the evolution of the systems can be catastrophic. Arguments about uniformitarianism, catastrophism, dynamic equilibrium etc. are irrelevant in these environments because the evolution of the landscape takes place by a combination of processes which incorporate elements of all of these concepts.

Fluvial systems in glacial and proglacial environments have daily, intra-seasonal, and inter-seasonal fluctuations in dischar~e and in the sediment transported. In addition discharge and suspended sediment may not always be highly correlated. As a corollary this produces highly complex depositional sequences. Within these environments lakes may exist for periods ranging from a few hours to many centuries and the variability of the sediment input induces complex sedimentary successions. These complexities are amply illustrated in the facies approach to glacigenic sediments ofBrodzikowski and Van Loon (1991). In the transition from the glacial to the paraglacial environment it is impossible to isolate fluvial from lacustrine environments because of the range from gradual to abrupt transitions between sites. It is the variability within the glacier drainage basin which accounts for the complexity of the lacustrine environments.

The theme of the following illustrations is that proglacial environments are characterised by rapid transitions at different spatial, quantitative, and temporal scales. The landform assemblages may be similar during long periods of evolution but the details will change. The illustrations will be organised around five perspectives highlighted by field examples drawn primarily from the St. Elias Mountains in the southwest Yukon (Figure I).


1. The "nonnal" fluvial regime in glacierised basins. 2. Rapid changes in the glacier drainage system. 3. Ice-dammed lakes, drainage and sedimentation. 4. Major drainage system changes. 5. Other lacustrine environments in glacierised basins .

... 10 20 JO -40 ~

Figure 1. Location map ofsites in the Kluane region of the southwest Yukon.

1.3. THE "NORMAL" FLUVIAL REGIME IN GLACIERlSED BASINS

The "nonnal" fluvial regime within glacierised basins is in itself very variable (Young 1990). The annual regime progresses from very low flow in the winter, through snow melt contribution from the non-glacierised component of the basin, snow melt on the glacier and finally glacier-ice melt before fall shut down of the system back to winter flow levels. The

48 Peter Johnson

regime is complicated by the altitudinal progression of the different melt phases through the summer and by the very strong diurnal fluctuations which occur during the spring, summer and fall. Timing of the onset of snowmelt and ice-melt and the cessation of melt at any altitude may vary by a number of weeks from year to year. These "normal" variations in themselves cause migration of the fluvial channels and changes in the lacustrine environments within the system. The diurnal variations alone produce patterns of mobilisation and deposition of sediment which induce rapid transitions in the braided streams below the glacier terminus and may be sufficient to produce daily rythmites in lake sediments.

The rise to peak flows in the spring may be very dramatic. For a period of 1 0 days in May-June 1972 on the Donjek River at the Donjek Glacier (Figure 1) the minimum discharge on any day exceeded the maximum discharge of36 hours earlier. In the braided stream bed at this time there were still accumulations of aufeis and river ice from the winter. The diversions induced by this ice and the changes in sediment transport produced by the rise in discharge caused a complete reworking of the channel system.

The variability of mountain climate frequently induces high magnitude events. High rainfall onto a saturated basin overloads the glacier drainage system and also produces overland flow, initiating mass movement of sediments in the basin. In 1983 at Peyto Glacier, Alberta (Johnson and Power 1986), a regional storm system contributed a measured 41 mm ofrain (with winds averaging 40 kmIhr) on the night of July 11th - 12th, following nine days with a total of 47 mm of rain. On July 11th the discharge from the glacier had already contained large pieces of glacier ice which contributed to channel blockage over the next four days. The rainfall of the night of the 11th - 12th discharged as surface runoff exceeding the capacity of a channel through the ice-cored moraine at the glacier terminus. Discharge backed up, overtopped the moraine in a flood wave, and removed the till cover over an extensive area of the moraine. The flood wave washed out the Environment Canada gauging site down valley. A further local storm on the 13th July with a measured 15.5 mm of rain in two hours triggered a debris flow of all of the sediments over the ice core of the lateral moraine on the south side of the valley (Figure.2), again blocking the stream channel and drainage from the glacier. This blockage was subsequently breached and a second flood wave extensively modified the braided stream channel and completely destroyed the hydrological monitoring site.

1.4. RAPID CHANGES IN THE GLACIER DRAINAGE SYSTEM

Measurements of the outlet streams from small glaciers at a number of sites in the Kluane Ranges (Grizzly Creek, Maxwell Creek, and the Kaskawulsh Basin) (Figure I) have recorded frequent pulsing of discharge (Johnson 1991 a). These pulses have been caused both by glaciological and geomorphological changes but do not appear to have been related to the magnitude of diurnal fluctuations.

A basin containing a small annually draining ice-marginal lake in the Kaskawulsh Glacier system (Kasper 1989; Johnson and Kasper 1992) (Figure 3) illustrates these changes. The tributary basin which contributes water to the ice-dammed lake has three contrasting glacier components, one glacier is subject to periods offast glacier flow, and the other two exhibit more uniform flow conditions over time. Of the latter, one has


backwasted 2 km into its cirque leaving a valley choked with ice, glacial and avalanche debris. The other terminates down on the floor of the basin. The glacier which is subject to periods of fast glacier flow was very active in 1986 and 1987 and exhibited drainage fluctuations from almost zero to >5 m3sec·1 over periods of I - 5 mins. These pulses occurred at any time during the day or night and could not therefore be attributed to the effects of flow on the rising limb of the hydrograph. They were apparently caused by internal glaciological changes within the fast moving glacier and were not predictable. Twice during field seasons at this site pulses were of sufficient magnitude to wash out the stream gauge and the stilling well. The combination of the diurnal fluctuations and the drainage pulses caused rapid changes in the braided stream and the delta formation into the ice-dammed lake.

Figure 2. The ice-cored moraine on the south side of the Peyto Glacier terminus almost entirely stripped of its till cover by a debris flow. The flooded area at the glacier terminus had an accumulation of up to 2 m of sediment.

Of the other two glaciers the one which terminates on the lake basin floor was subject to rapid changes in the stream portal position. In 1987 three portals were formed, the first through the ice-cored moraine on the north side of the valley. This was succeeded by a hydrostatic outburst though the moraine on the south side. This second channel collapsed in two sections resulting in partial rerouting of the stream. Finally the portal migrated, again as a catastrophic outburst, to the ice-cored outwash zone in the centre of the valley.

50 Peter Johnson

Figure 3. The Kaskawulsh ice-dammed lake basin. The basin to the right of the photograph with the distinct trim line is the sub-basin with the glacier subject to periods of fast glacier flow

The third of the tributaries, where the glacier had backwasted into the cirque, was dammed on a regular basis by avalanching, mud flows, and collapse of the channel cut through ice and snow deposits which filled the valley.


In the latter two cases, all of the pulses occurred during the rising and peak flow elements of the hydro graph suggesting that they were discharge related. This basin, as with many other proglacial environments in the region, has large volumes of stagnant ice incorporated in the sediments (Figure 4). The ice core was being degraded by a combination of melt due to heat flow through the sediments, erosion of sediments by surface streams, and thermal erosion of the ice by subsurface streams. The continuing and are an illustration of an environment in transition due to processes operating at different spatial and temporal scales superimposed geomorphological changes promoted by the degradation of this stagnant ice.

Figure 4. The stagnant ice-cored area on the floor of the Kaskawulsh Lake basin. At this time in late July the lake was at its maximum recent extent.

In Grizzly Creek similar patterns of drainage fluctuation occurred over six years of observations between 1974 and 1980. Typically at the glacier terminus early season lateral meltwater drainage was diverted to a subglacial conduit in early summer, usually resulting in a high magnitude hydrostatic resurgence at the glacier terminus. The highest magnitude event observed deposited a layer of gravel up to 1 m thick over a proglacial snow patch and fluvio-glacial deposits at the glacier terminus which covered an area of approximately 500 m2 •

The proglacial environment of the Grizzly Creek glacier deglaciated since the maximum of the Little Ice Age (Figure 5) also contains large amounts of buried glacier ice. This was exposed beneath < 1 m of fluvially deposited gravel during periods of incision of the creek. Thicknesses of 2.5 m were observed without a.ny indication of reaching basal sediments. In 1986 the drainage from the glacier was diverted into a small pond on the

52 Peter Johnson

proximal side of the Little Ice Age moraine which drained frequently but irregularly through the Little Ice Age moraine and a rock glacier on the east side of the valley (Figures 6 and 7).

Figure S. The proglacial zone of the Grizzly Creek glacier at high flow. This braided stream environment is underlain by stagnant glacier ice. (Scale can be judged from three figures left of centre of the photograph.)

One very significant example of drainage change in the Kluane region occurs at the Kaskawulsh Glacier (Johnson 1991 b). The current terminus of the glacier lies in a rather unique situation at the diffluence of two major river systems, the Kaskawulsh River which drains south into the Alsek River and then directly into the Pacific Ocean, and the Slims River which drains into Kluane Lake, the Kluane, Donjek, White and Yukon Rivers and fmally into the Bering Straits, a significantly longer journey to the ocean (Figure 8). The drainage from the Kaskawulsh Glacier is primarily oriented into the Slims River. This occurs from subglacial channels under pressure, often forming geysers. On occasions however, either at the beginning of the season or during the summer the Kaskawulsh River may become the major recipient of discharge from the glacier. The effect of this is most evident at Kluane Lake (Figure 9) where the rate of filling is much slower, the maximum season water level is well below average, and the discharge of the Kluane River is greatly reduced. This drainage reversal seriously reduces the sediment transport and has a major effect on the sedimentation in Kluane Lake.

Proglacial and Paraglacial Fluvial and Lacustrine Environments

Figure 6. Pond full stage in the Grizzly Creek ice-cored moraine.

Figure 7. Pond empty stage in the Grizzly Creek ice-cored moraine. The stream can be seen draining into the ice core. Its subsurface course was through the moraine and a rock glacier on the distal side of the moraine.

53

54 Peter Johnson

Figure 8. The tenninusofthe Kaskawulsh Glacier looking south down the Kaskawulsh River. The Slims River lies to the left of the photograph. The location of the hypothesised southward course for Kluane Lake drainage lies between the bedrock knoll in the centre of the photograph and the east side (left) of the Valley.

Figure 9. View north ofKluane Lake with the Slims River delta in the lower left.


Other major drainage changes occur frequently in the proglacial and paraglacial zones. Two Kluane Region examples can be cited, one a recent occurrence at the Donjek Glacier, and the second relates to a possible drainage reversal ofKluane Lake.

1.5. MAJOR DRAINAGE CHANGES

A photographic record of the Donjek Glacier, which terminates in the Donjek Valley, exists from the mid 1930s to the present (collection of Walter Wood, Government of Canada aerial photographs, and scientists and tourists photographs). During the 60+ years of record the Donjek River has been confined mainly to the east side of the valley between the glacier terminus and the valley side (Figure 10). The Donjek Glacier is subject to periods offast glacier flow which have in the past dammed the Donjek River and formed a large ice-dammed lake upvalley (Clarke 1980; Perchanok 1980). Catastrophic drainage events of this lake may have been responsible for the incision of the rock channel of the east side course of the river. The photographic evidence indicates a few years during backwasting phases of the glacier when the lower waterfall section of the bedrock channel was bypassed as the river cut into the centre of the valley and exited downvalley through the massive Neoglacial terminal moraine system. Sometime, apparently early in 1998, the river entered a subglacial course from the heel of the terminus to a portal at the downvalley extent of the terminus lobe completely abandoning the old channel. Erosion along this subglacial channel resulted in collapse of the ice leaving a long ice-walled channel through much of the glacier terminus. Below the portal the valley floor was covered in blocks of ice (Figure 11) as far as Spring Creek. This indicates that the collapse of the subglacial channel may have blocked the rest of the conduit which finally resulted in a hydrostatic blowout. The flood wave deposited the ice blocks over an extensive area of the flood plain.

The currently accepted hypothesis on the history of Kluane Lake proposed by Hugh Bostock (1952, 1969) has recently come under scrutiny by the author. Bostock (1952, 1969) proposed that as deglaciation proceeded in the Kaskawulsh Basin at the end of the Ice Age, the glacier backwasted to a position upvalley of the diffluence of the Slims River and Kaskawulsh River Valleys. Kluane Lake then flowed south into the Alsek River (Figures 1 and 8). Bostock proposed that Kluane Lake level was 10m lower than present through the Holocene and continued to drain south until the most recent advance of the Kaskawulsh Glacier in the Little Ice Age blocked the channel. This produced a rapid rise in Kluane Lake to 10m above present level and initiated a new outlet to the north, the Kluane River, into the Donjek and White River systems. This hypothesis was based on the occurrence of drowned spruce, about 400 years old, around the shores ofKluane Lake; the form of Christmas Bay at the south end of the lake; and on limited observations of raised beaches where he found no White River volcanic ash. Clague (1981) reported other drowned deposits 50 m below current water level near Sheep Mountain and suggested that these represented even lower water levels in an early Kluane Lake. The instability of the terrain in this area however may account for the movement of this material into the lake bottom. The Bostock hypothesis of a drainage reversal about 400 years ago is considered by the author to be unlikely because of the massive amount of sediment in the Kluane Lake basin. It is proposed that Kluane Lake may have been lower until early in the Neoglacial

56 Peter 10hnson

and that a southward channel was blocked by an early Neoglacial advance of the Kaskawulsh Glacier.

The groundwater table in the glacial deposits south ofKluane Lake (Figure 12) is controlled by the water level of the lake. A small lake close to the main lake was dry until 3180 years BP which may indicate that Kluane Lake rose at this period, causing a rise in the water table and a filling of the small lake basin.

Figure 10. Terminus of the Donjek Glacier prior to the diversion of the Donjek River in 1998. The Donjek River lies along the valley side in a rock cut gorge. The subglacial diversion of the river in 1998 was through the section of the terminus in the top right quadrant.

1.6. ICE-DAMMED AND PROGLACIAL LAKES; SEDIMENTATION AND DRAINAGE

Ice-dammed lakes are a common feature of glaciated areas. There have been many studies of the hydrology and sedimentology of ice-dammed lakes in all glacierised regions of the world because of the hazard posed by sudden drainage. Ice-dammed lakes have a range of hydrological regimes from annually draining to continuous existence over hundreds of years. They represent therefore transitions at micro to meso time scales. Sedimentary evidence of the former existence of small lakes is frequently sparse however in the southwest Yukon due to melt of extensive ice cores and fluvial and glacial reworking of sediments.

An assessment of ice-dammed lake sites in the southwest Yukon was made at the time of the Foothills proposal for an Alaska Highway gas pipeline (Environment Canada 1977). This assessment found that the majority of the lakes and the lake sites in the St. Elias Mountains were small and posed no threat to the potential pipeline route. A few


larger lake sites were identified which, if they refilled to maximum extents indicated by geomorphological evidence, might pose a threat. Drainage of the Donjek Lake (Clarke 1980) was assessed as a potential hazard should it refill to its maximum historic levels, and Glacial Lake Alsek (Schmok and Clarke 1989) flooded the site of Haines Junction at the junction of the Alaska Highway and the Haines Highway in the middle of the last century. With the current downwasting and backwasting of the glaciers in the region it is considered that there is no longer a threat from these lakes to infrastructure in the region but the history of filling and drainage as a result of surges of the Donjek Glacier and the Lowell Glacier demonstrates the importance in the development of the proglacial and paraglacial landscape. If, however, there is substantial glacier backwasting as a result of global warming then there is the potential for new lake sites to be formed.

Figure 11 . Down valley of the Donjek Glacier moraines after the drainage diversion. The white cover to the braided stream plain is blocks of ice from the hydrostatic blowout after collapse of the subglacial channel.

Three small ice-dammed lakes sites in the St. Elias Mountains have been studied in detail, one on the Dusty Glacier system (Johnson 1995), a second on the Kaskwulsh Glacier system (Kasper 1989; Johnson and Kasper 1992), and a third (Hazard Lake) on the Steele Glacier system (Liverman 1987) (for locations see Figure 1). These studies have demonstrated the complexity of the hydrological regimes and the sedimentary environments of the St. Elias Mountains ice-dammed lakes (Johnson 1997). The Dusty Glacier lake was dammed in a tributary basin during periods of surge of the glacier. There is a record in an exposure close to the damming moraine of three short duration lakes. Twigs from near the base of the middle lake have been dated at 400 BP which indicates that the lakes formed at the time of the maximum of the Little Ice Age. The sedimentary record is

58 Peter Johnson

composed of daily rythmites of silt with some drop stones and coarser turbidity flow layers. Ice crystal casts suggest periods of drawdown of the lake during each of the three events. This sedimentary succession accumulated against an ice-cored moraine over glacially deposited gravels. It is now being rapidly eroded as the ice core melts at exposures along the margins of the glacier and where the basin stream has incised through the moraine. Ultimately ther will be little evidence of the existence of the lake in the deglaciated environment.

Figure 12. Stagnant ice deposits in the Shakwak Trench at the south end of Kluane Lake. The irregularly shaped lake in the centre of the photograph is Jenny Lake, one of a number of closed basin marl lakes.

The Kaskawulsh Glacier lake (Figures 3 and 4) is currently an annually draining lake but the depositional record in the basin indicates a transition from a more stable lake higher in the basin at the maximum of the Little Ice Age to the present annually draining form. A study of the hydrological balance of the basin indicates that the input to the lake is considerably greater than the storage and it is hypothesised that the lake basin is constantly leaking subglacially. The glacial and glaciolacustrine sediments in this basin are deposited on an ice core and are already rapidly degrading. Only small exposures of the upper lake sediments remain and continued erosion and ice core degradation may result in complete removal of the physical evidence ofa lake.

1.7. SEMI-PERMANENT LACUSTRINE ENVIRONMENTS IN GLACIATED AND GLACIERISED BASINS

A wide range of lacustrine environments may develop in the paraglacial region (Ashley, Shaw and Smith 1985; Brodzikowski and Van Loon 1991). They form in cirques, in


bedrock depressions or behind moraines in glaciated valleys, or in depressions in irregular glacial deposits. Hydrologically they may be open or closed systems. These lakes experience an equally wide range of sedimentary processes, often exhibiting a transition from being in a glacierised region to a non-glacierised region as deglaciation proceeds. During deglaciation a lake becomes more and more remote from the glacier bringing about changes in the supply of sediment and in the river regime. Total cut off from the glacier source fundamentally changes the hydrological regime and the sediment regime. Mush Lake and Kathleen Lake in the southern Kluane area are fed by nival regime rivers as compared with Kluane Lake which is still fed by the glacierised regime Slims River.

The glacial till deposited in the Shakwak Trench by glaciers originating in the St. Elias Mountains has a high dolomite content and as a consequence the groundwater is carbonate enriched. Closed basin lakes in the morainic deposits are fed entirely by groundwater and the sediment accumulation is dominated by autochthonous carbonates (Figure 12). The only allochthonous source of sediment is loess, which was generated along the Shakwak Trench in the period immediately after deglaciation, and from the Slims River in the Holocene. In these carbonate (marl) lakes there is an uninterrupted depositional record from the postglacial period. Loess deposition can be detected throughout this record by the magnetic susceptibility signal from lake sediment cores. Small quantities of magnetite from the loess produce a very low response superimposed on the magnetic susceptibility record of the carbonates which is normally zero or slightly negative. This indication of loess accumulation through the Holocene suggests that the Slims River valley source was active though the Holocene and is contrary to the previously suggested cycle of postglacial loess being weathered in the early Holocene to produce the Slims Soil and a renewed period ofloess formation in the Neoglacial (Denton and Stuiver 1966). It is proposed that loess accumulation was continuous in the Holocene but that weathering was occurring as deposition proceeded in the early Holocene compared to the processes occurring during the Neoglacial accumulation.

2. Discussion

These numerous examples bear testimony to the interconnectedness of the fluvial and lacustrine systems in proglacial and paraglacial areas; to their susceptibility to rapid change; and to the contrasts on the continuum between rapidly evolving and semi permanent environments. Transitions involving differences in magnitude at different spatial and temporal scales are the mode by which the landscape continues to evolve. It is the superimposition of processes at these different scales, which is evident in the glacierised and glaciated environment.

Landforms, or more generally landform assemblages producing the landscape, are snapshots of one instant in the transitions. Perhaps, in the past, landform assemblages have been taken as fixed environments because the temporal scale of transition has not been taken into account. In the modem proglacial environment, transitions are rapid, occurring on time scales from seconds to decades, compared with centuries to millennia for Quaternary environments. The temporal scale of transition of glaciated landscapes slows

60 Peter 10hnson

down as deglaciation is accomplished, and as the hydrological regimes change from glacierised to nival in character.

Slope processes which modify the landscape change as one moves from glacial to interglacial period. The inherent instability of many of the deposits as they are laid down during glaciation becomes inherent stability, but the landscape continues to evolve, albeit at a different temporal rate. We are learning however, that many of the glaciated environments which we have considered inherently stable may in fact be unstable as the effects of glacial unloading of the valley sides continue to become apparent.

The two paradigms introduced at the beginning of this paper must be seen as fundamental to geomorphology. It has been demonstrated that the concept of transition is critical to geomorphology when viewed from a process perspective. The concept of landform assemblages should integrate the ideas of process activity at different spatial, quantitative and temporal scales. Although one recognises particular environments such as proglacial areas these are constantly in flux through the activity of micro, meso, and macro processes operating over different time frames. The classification into micro, meso, and macro might in itself be problematic in that they are only points along a continuum. I have referred to fast glacier flow in relation to some of the glaciers discussed. This phrase has come to be used instead of glacier surges by some scientists because it is thought that there is a continuum of flow regimes from those with no variation over time to those which show infrequent very high velocity periods and the idea of unique flow mechanics for a surge may not be valid.

In order to fully comprehend geomorphic processes we must be aware of the physics of the material on which the process depends, for example how does polycrystalline ice behave under stress at different temperatures, and of the mechanics of the situation, for example what happens to the ice and the substrate at the zone of contact at the base of the glacier.

The most important element of the second paradigm is the predictive capacity which acknowledges that the landscape is in transition and that elements of the landform assemblages are changing at any time. The problem with prediction oflandscape change, or hydrological change, is that it depends on the correct prediction of climate change or climate variability, not just in an annual time frame but also seasonally. Landscape processes will be determined by changes in the winter accumulation, in the rate of summer melt, on the duration of ice cover etc. The validity of models of climate change will always be questioned, not just because of the basic assumption of 2xCoa but also on various questions concerning components such as proxy data and feed back mechanisms. Without prediction however, geomorphology will remain static.

In order to attempt prediction we need to be able to establish what has happened in the past and how that relates to what is happening in the present. It is from these scenarios that we can assess what may be the changes to processes and landform assemblages in the future. It is certain however, that landforms and landform assemblages will continue to be in transition.


Acknowledgements

The work referred to in this paper has been conducted over many years with funding from the Natural Sciences and Engineering Council, Environment Canada, The University of Ottawa, and the Northern Research Institute of Yukon College. The permission of Champagne Aishihik First Nation, Kluane First Nation, Kluane National Park and National Park Reserve, and the Government of Yukon to work in the region has been very much appreciated. Logistics provided by the Kluane Lake Research Station of the Arctic Institute of North America have been a mainstay of the research.

References

Ashley, G.M., Shaw, J. and Smith, N.D. (1985) Glacial Sedimentary Environments. Society of Economic Paleotologists and Mineralogists. Short Course No. 16. 246pp.

Bostock, H.S. (1952) Geology of the northwest Shakwak Valley, Yukon Territory. Geol. SUl'I'. Can. Memoir 267. 54pp.

Bostock, H.S. (1969) K1uane Lake, Yukon Territory, its drainage and allied problems (115 G, and 115F.E). GeoL SUl'I'. Can. Paper, 69-28, 19pp.

Brodzikowski, K. and van Loon, A.J. (1991) Glacigenic Sediments, Developments In Sedimentology 49, Elsevier, 674pp.

Calow, P. and Petts, G.E. (cds.) (1992) Rivers Handbook, Vol. 1, Blackwell Scientific, 526pp. Calow, P. and Petts, G.E. (eds.) (1994) Rivers Handbook, Vol. 2, Blackwell Scientific, 523pp. Clague, J.J. (1981) Landslides at the south end ofKluane Lake, Yukon Territory, Can. J. Earth Sci. 18, 959-

971. Clarke, G.K.C. (1980) An estimate of the magnitude of outburst floods from Lake Donjek, Yukon Territory,

Canada, Report for the Department of Indian and Northern Affairs, 90pp. Denton, G.H. and Stuiver, M. (1966) Neoglacial Chronology, Northeastern St. Elias Mountains, Canada, Am.

J. Sci. 264, 577-599. Environment Canada. (1977) Report on the influence of glaciers on the hydrology of streams affecting the

proposed Alcan pipeline route, Glaciology Division, Inland Waters Directorate, Fisheries and Environment Canada, 38pp.

Johnson, P.G. (1991a) Pulses in Glacier Discharge: Indicators of the Internal Drainage System of Glaciers. In Northern Hydrology. Selected Perspectives, NHRI Symposium No.6, Prowse, T.D. and Ommanney, C.S.L. (eds.), 165 - 176.

Johnson, P.G. (1991b) Discharge Regimes of a Glacierized Basin, Slims River, Yukon. In Northern Hydrology. Selected Perspectives, NHRI Symposium No.6, Prowse, T.D. and Ommanney, C.S.L. (eds.), 151 - 164.

Johnson, P.G. (1995) Ice-dammed lake history, Dusty Glacier, St.Elias Mountains, Yukon, Can. Geog. 39, 26-273.

Johnson, P.G. (1997) Spatial and temporal variability ofice-dammed lake sediments in alpine environments, Quat. Sci. Rev. 16, 635-647.

Johnson, P.G. and Kasper, J.N. (1992) The development of an ice-dammed lake: the contemporary and older sedimentary record, Arctic and Alpine Research, 24, 304-313.

Johnson, P.G. and Power, J.M. (1986) The role of high-magnitude runoff events in glacierized basins. Proceedings of the International Symposium on Glacier Mass Balance, Fluctuations and Runoff1985, Data o/Glaciological Studies No. 58, 82-86.

Kamb, B. and Engelhardt, H. (1987) Waves of accelerated motion in a glacier approaching surge: the minisurges of Variegated Glacier, Alaska, U.S.A, J. Glaciol, 33, 2-46.

Kasper, J.N. (1989) An ice-dammed lake in the St. Elias Range, southwestern Yukon Territory, Unpublished M.A. Thesis, University of Ottawa 197pp.

62 Peter Johnson

Liverman,D.G.E. (1987) Sedimentation in ice-dammed Hazard Lake, Yukon, Can. J. Earth Sci. 24, 1797-1806.

Perchanok, M.S. (1980) Improved flood estimates of the Oonjek River. Report to Northern Environmental Protecion Branch, Department ofIndian Affairs and Northern Development, 60pp.

Petts, G.E. and Amoros, C. (eds.) (1996) Fluvial Hydrosystems, Chapman and Hall, 322pp. Prowse, T.D. and Gridley, N.C. (1993) Environmental aspects of river ice. National Hydrology Research

Institute Science Report, No. S, ISSpp. Schmok, J.P. and Clarke, G.K.C. (1989) Lacustrine sedimentary record ofice-dammed Neoglacial Lake A1sek,

Can. J. Earth Sci. 26, 2092-21 OS. Slaymaker, O. and Spencer, T. (1998) Plrysical Geograplry and Global Environmental Change, Longman,

292pp. Statham,l. (1977) Earth Surface Sediment Transport, Oxford, 184pp. Young, GJ. (1990) Glacier Hydrology, Chapter 6. in Northem Hydrology; Canodlan PerspectiVes, National

Hydrology Research Institute Science ReportNo.l. Prowse, T.D. and Ommanney, C.S.L. (eds.), 135-162.

POSTGLACIAL LANDFORM AND SEDIMENT ASSOCIATIONS IN A LANDSLIDE-FRAGMENTED RIVER SYSTEM: THE TRANSHIMALAYAN INDUS STREAMS, CENTRAL ASIA

Kenneth Hewitt Cold Regions Research Centre and Department of Geography and Environmental Studies Wilfrid Laurier University 75 University Avenue West Waterloo, Ontario N2L 3C5

Abstract

The chapter concerns deposits and land form associations along high mountain rivers interrupted by landslides. The catastrophic landslides are largely due to failure of rock walls that were over steepened by glaciers and debutressed by deglaciation. Some 180 rock avalanche deposits have been identified that form(ed) cross-valley barriers on the Upper Indus streams. They provide a common explanation for widely discussed features in the fluvial zone. Aggradational sequences, built behind the landslide barriers, include mass movement, lacustrine, fluvial, and aeolian materials. Breaching has led to distinctive sets of erosional landforms, notably barrier-related and "defended" river terraces, trenched fans and fan terraces, rock gorges superimposed from valley fill, and mid-valley "isolated rocks." Landslide interruptions have been major constraints on valley fill sedimentation and the phasing of late-, post- and paraglacial sediment transfers. Along with more frequent but smaller debris flow and avalanche barriers, and glacier interference with rivers, they have helped to produce a chronically fragmented drainage system. The interactions of a range of geomorphic and sedimentation processes in interrupted river reaches are decisive for the late Quaternary landscapes of the fluvial zone. The chapter focuses on the depositional and degradational landform associations involved. Climate change and tectonically driven stream incision, hitherto used to explain the features of interest, are shown to have been masked, suppressed or redistributed in response to interruptions of the river system. However, a further type of paraglacial transition is suggested involving postglacial responses of glacially destabilised rockwalls.

63 K. Hewitt et al. (eds.), Landscapes a/Transition, 63-9J. © 2002 Kluwer Academic Publishers.

64 Kenneth Hewitt

1. Introduction

For more than 140 years, earth scientists have given special attention to the great quantities of intermontane sediments along the transHimalayan Indus streams (Strachey 1853; Thomson 1852; Cunningham 1854; Schlagintweit 1857; Drew 1875; Oestreich 1906). Throughout the Upper Indus basin there are extensive river flats underlain by deep, valleyfill sediments (Figure 1). Trunk streams and tributaries have stepped profiles; relatively gentle, more open reaches, alternate with steep narrow ones. The former are flanked by stream terraces, alluvial remnants of large lakes and many, often coalescing, sediment fans. In some places, the rivers flow in bedrock, but most of the steep, canyon-like sections are cut into coarse valley-fill and flanked by stream terraces.

The valley-fill sediments include the "fans, alluvium and lacustrine deposits" of Drew (1873), and what Norin (1925) called "thick accumulations of river conglomerates" and "moraines". They are the "diluvial" deposits of Paffen et al., (1956, 13-15); the "alluvial/colluvial terraces and valley fills" of Goudie et al. (1984), and much of the "Quaternary and Recent [sediments]" of Searle (1991). Eolian deposits are also widespread, including loess, sand sheets and occasional dune fields.

It would be difficult to over-estimate the role of these sediments and related features in past reconstructions of the Quaternary and regional geomorphology. Giotto Dainelli, in the most influential synthesis, referred to them as "characteristic morphological elements" of the Upper Indus, "within the ancient valley floor" (1922,13 and Plates V and XXVII). Owen and Derbyshire (1993, 123-4) identify them with the "main types" of processes, landforms and sediments in the Karakoram. Hitherto, the valley-fill materials have been attributed mainly to climate change, especially deglaciation, dessication and paraglacial sedimentation (Schneider 1959; Hewitt 1968; Kalvoda 1992; ed. Shroder 1993). The widespread trenching of the valley fil, and occasional rock gorges, are attributed to high but differential rates of tectonic uplift (Seeber and Gornitz 1983: Goudie et al. 1984, 400; Burbank et al. 1997). In effect, the valley floor features were seen to record the re-establishment, in the Holocene, of stream incision and high rates of sediment removal in response to tectonic uplift, following sedimentation "excesses" oflate-glacial and paraglacial transitions.

However, in recent surveys I found these phenomena to be associated, in nearly every case, with post-glacial landslide barriers, especially those created by catastrophic rock slides (Hewitt 1998a; 1998b). Large landslides, and their potential for damming streams, have certainly been recognized by others in the region (Drew 1875; Oestreich 1906; Hewitt 1968; Hughes 1984; Goudie et al. 1984; Cronin 1989; Shroder 1989, 1993; Searle,1991; Owen 1991; Owen and Derbyshire 1993, 125). But few catastrophic rock slides were actually identified, and their deposits were often misidentified as moraines (Hewitt 1999).

Of 182 catastrophic rock slides found to date, 174 were reconstructed from more or less ancient deposits in the present-day fluvial zone (Figure 2). Most descended from glacially over-steepened rock walls and were emplaced in ice-free areas. They left crossvalley deposits, the consequences of which are the main focus of this paper.

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Figure 2. Features of the landslide- interrupted river reach illustrated by the Katzarah rockslide-rock avalanche which impounds the Skardu intermontane basin (cfHewitt, 1999,223-228) : i) View of the cross-valley barrier looking westwards down the Indus. The present-day river flats and braided channels are at the level of the partially breached dam. The highest summer flows cover the flats completely. Arrows identify the undisturbed rock avalanche surface and barrier (ra). Lacustrine deposits (L) against the barrier, and the terrace (T) in the foreground, represent the original high level impoundment. The source of the rock avalanche (S), is hidden inside the hanging valley above the village of Bra gar do (8). p:: photo station for : ii) View of the same features looking upstream/eastwards. In place rock avalanche material (ra), lies in the foreground, right middle ground, and forms islands protruding through the valley fill. The original, main terrace level (T) is shown around the margins of the basin, and large sediment fans (F) developed at that level. In the foreground is an erosional terrace (E) cut in the early phase of barrier breaching. (P:: photo station of2.i)

Postglacial Landform and Sediment Associations 67

2. Rock Avalanches and Landslide-interrupted Rivers

Rock avalanches are generated when very large rock slides descend some hundreds of metres from steep mountain walls. The process crushes and pulverizes the rock mass, and gives great momentum to the resulting debris (HstlI978; Hungr 1989). This may travel many kilometres beyond the source slope, at speeds often exceeding 200 km br"l. Once in place, the morphology and composition of rock avalanche deposits promote resistance to erosion. Their impact also reflects sheer scale (Table I). Over half the Karakoram rock avalanches found, originally covered more than 10 km2 of v alley floor and were more than 50 x 106 m3 in volume. Two of the events covered over 50 km2 and exceeded 109 m3 in volume. Debris thicknesses range from 5m to over 500m (Hewitt I 998a). These dimensions show that the cross-valley barriers fully encompass the depths of valley-fill sedimentation, heights of river terraces, and other features attributed to them here (see Hewitt 1998b).

Virtually the entire fluvial zone is affected by these barriers, or erosion and sedimentation related to them. Roughly one rock avalanche occurs in every 14km of river valley surveyed (Figure 3). Between one fifth and a quarter of the length of valley floor is, or was, buried by rock avalanche debris. Lacustrine deposits were found upstream of almost every example, although most lakes are now drained or filled with sediment. However, though breached, at least 120 of the landslide barriers are not completely cut through. They persist as local base-level and steps in the river profiles.

The largest dams, such as "Gol-Ghone" (see below), would drown the valleys of today for over 100 km upstream. Barriers such as Batkor, Nomal, Sassi, Rondu-Mendi or Gol-Ghone, blocked as much as one third of the transhimalayan Indus (Figure 4). Two or more may well have coincided or overlapped in time, impounding almost all of it. This implies starving the main and lower Indus of sediment. Meanwhile, sediment yields today are among the highest in the world (Hewitt 1968). Ifrepresentative of the late Quaternary, they are quite sufficient to explain the valley fill observed, and high rates of sedimentation implied during impoundment episodes. Some of the great floods evident in the stratigraphic record may reflect catastrophic breaching of landslide dams (Desio and Orombelli 1983).

The nature of rock slide-rock avalanches is extensively discussed elsewhere, their relations to geological setting and slope stability in the Karakoram, and criteria for distinguishing their deposits (Hewitt 1998b; 1999; 2000). The present paper will focus on the sediment and landform associations that develop in interrupted river reaches. First, three examples are described to illustrate features and complexities encountered in the field.

68 Kenneth Hewitt

Table 1. Estimated dimensions of some major rock avalanches that have danuncd the Upper Indus streams

Rock River Vol. of Area of llelgbt: HeJaht: Lea&tb: Width: Lake Av. (lee Flal) depolltl depolitl orlpnal breached CI'OII down lengtb

vaHey valley

SHYOK BASIN (HAL TISTAN):

Litak Husbe 80 4+ 120 120 2000 2000+ 7

Haldi Slitoro 600 25 260 240 2000 5000 20+

Yugo- Shyok 100+ 3+ 60+ 10 1500 2000+ 32 Kurphak

Kunis- Shyok 1000+ 25+ 200+ 150 2000 11000 50+ Gwa

MIDDLE INDUS BASIN (HALTISTAN):

Gol- todus (99) 600+ 18+ 600+(1) 550 2500 8000+ 1 Ghonc 'A'

Gol-GhoDe 'B' Indus (99) 300+ 15 5SO 500 2500 6000+ 70+

Tsok Dumordo ISO 8 110 95 1800 4500 6

Sltpara- Sltpara 400 22 120 60 2000 5000+ 7 Skardu (SS)

Katzalllh todus(K2) 200+ 20 150+ 55 4500 9000 35

Roodu-Mendi \odus (RM) 1500+ 49+ 1100 950 4400 14000 180+

GILGIT-HUNZA BASIN:

Oulung Yasin 450 20 150 120 1500 2500 6 Bar

Gupis Ghizar 160 5+ 30 30 2000 2000 15

Bhort Karam- 110 8 25 20 2500 3000 10 bar

Balkor Gilgit 1000 17 110 110 2200 3000 90

Nomal Hunza (No) 10000+ 40+ 700(1) 600 5000 6000 150

Naltar Naltar 300 14 20 20 2500 6000 40 Lakea

Balti!- Hunza 400 150 ISO 3000 3000 20 Sumayar

Yashhandao- Hunza 250 21 50 50 6000 3000 Barn!

CIDTRAL BASIN:

Mir-gram- Yarlchun 600 18 120 100 3000 7000 55 Parwak (MP)

Postglacial Landfonn and Sediment Associations

Gaciers r--.. Ridges Tlsar· TONn. ~ Mounlliln pass

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Tental .... identificat ion -... lacustrine beds ,...

Figure 3. Cross-valley rock avalanche deposits identified in the Indus, Shigar and Shyok Basins of Baltistan, indicating the extent of valley floor presently or formerly covered by them, and related incidence of sediments formed in lakes they once impounded.

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69

70 Kenneth Hewitt

2.1. THE OOL-OHONE EVENTS (Figure 5)

Just before the Indus enters Skardu Basin in Baltistan, the remains of two rock slide-rock avalanches fill the valley floor over a distance of llkm. They descended more than l200m from the left/west wall. Their detachment zones appear to involve a common zone of mountain wall sagging and collapse. The two rock avalanches overlap about half way between 001 and Ohone. Both climbed the east, impact slope up to 700m above present river level (Hewitt, 2000). Subsidiary lobes were deflected far up and down the valley. The river has removed at least half the original mass.

There are abandoned spillways across the younger Ool-Ohone "B" event, on the right/east flank, and some 550m above the Indus. They record initial barrier height and a lake that overflowed it (Hewitt 1 998b, 67). Under present conditions, water this deep would reach back 70-80km. Lacustrine deposits occur along the Indus and Shyok valleys, but they are repeatedly pinched out or interfingered by debris flow, coarse river gravel and rock falls. Agricultural fields at 001 lie on alluvial flats of the former lake bed. Houses cluster on boulder-covered mounds or pressure ridges of rock avalanche debris.

Eventually, the high overflow channel was abandoned. A main breach developed along the proximal zone of the landslides. Erosional terraces were cut in rock avalanche debris through an elevation range of 500m (Figure 5 ii). However, despite its deep canyon and relatively steep fall through the landslides, the river has yet to cut down to the former valley floor. Wide terraces occur below Ohone, their surface covered by large flood-transported boulders, mainly of the rock avalanche lithologies. These record great floods that passed over and through the barrier. Above 001, river flats exposed at low flows are major sources of sand and dust that cover much of the rock avalanche and terrace surfaces.

Dainelli (1922, Plate XXV, Figure12 and XXXVII, Figurel) classified the landslide materials as moraines. He recognized that two distinct events are present, but attributed them to his "2nd" and "3rd" glacial advances, an interpretation unchallenged until now. No defmitive glacial, meltwater orthermokarst features were found in his ''moraines''. The concept of a tributary glacier from 001 valley blocking the main Indus is hard to accept. How could it deposit such huge, uniform volumes of debris and then quietly melt away. leaving intact cross valley moraines 500m thick and capable of damming a huge, long-lived lake?

Postglacial Landform and Sediment Associations

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71

Figure 5. The Gol-Ghone rock avalanches (cfHewilt, 1998,67): i) Sketch map indicating the original extent of the two rocksliderock avalanches, and present-day features in the barrier zone, ii) Longitudinal cross-section a - b (Fig, 5.i.) through Gol-Ghone 'A', reconstructing the highest part of the original cross-valley barrier, and present day features related to its erosion. The highest lake level of the Gol-Ghone '8' event is shown, as reconstructed from spillways.

72 Kenneth Hewitt

2.2. THE MlRAGRAM-PARWAK EVENT

A rock avalanche descended from rock walls in the Miragram tributary into the Yarkhun valley south west ofMastuj. It filled the valley floor for some 7 km and dammed the river. Original volume appears to have exceeded 600 x 106 m3 (Table 2). The leading part swept directly across the main valley and climbed over 400m up the opposing wall. A typical, asymmetrically thickened, cross-valley deposit occurs, with well-preserved pressure ridge or ''hrandung'' against the impact slope (Heim 1932; Hewitt, 2001). Rock avalanche debris was diverted around this stalled mass. One lobe traveled down valley, but has been largely removed by erosion. The other traveled over Skm up-valley leaving a series of longitudinal and transverse pressure ridges (Figure 6). Where the river now cuts through, rock avalanche material as much as 100 m thick is exposed. It consist of black dioritic and white granitic lithologies from the Buni Zorn pluton, which outcrop high in Miragram Gol, while the main valley is cut in metasedimentary and metamorphic rocks of the Reshun Formation (Searle 1991, 87). Typically, the angular boulders or megaclasts, and samples of the crushed and pulverized matrix materials, consist wholly of one or other of the plutonics (Hewitt 1999, Table 2).

Haserodt (1989, 21S) thought the ridges of rock avalanche materials were lateral moraines, and explained the valley-fill sediments by glacier damming. The valley has been glaciated, but I found nothing attributable to ice, meltwater or thermokarst processes in extensive exposures of Haserodt"s "Wallmorlinen" ("lateral moraines"), only rock avalanche debris. Further, one would expect a main valley glacier or tributaries here to carry and deposit debris from the underlying Reshun, or up-valley Darkot, Formations, not purely Buni Zorn materials. The depositional features show the rock avalanche descended into an ice-free valley.

Upstream, and overlapping the rock avalanche, are remnants of a complex episode with lacustrine, fluvial, deltaic, mass movement, and aeolian sedimentation, and a series of sediment fans. The original lake must have stretched back more than 2Skm. The highest cross-valley barrier, in front of Miragram Gol, is recorded in spillways, now dry, 200m above present river level. Depressions in the rock avalanche deposit are partly infilled by lake bed and aeolian sediments, and by rock falls and debris flows from the valley walls. The prograding Turi Parwak sediment fan buried the distal part of the upvalley rock avalanche lobe and eventually spanned and cut the lake in half. Immediately above it are some40m of multi-year, fme-grained, buff-coloured beds, typical lacustrine deposits of the region. They overlap and interfmger debris flow and stream sediments from the fan. Upstream they are again pinched out by coarse river gravels of the largest of these fans coming from Laspur valley. Above Mastuj, in the main and Laspur valleys, multiple smaller, coalescing debris flow fans, spread into and filled the lake.

Barrier breaching and stream incision have revealed the complexities of the fonner aggradation in river cliffs and segmented fans. The river has removed much of the fmer- grained sediment along its axis, but flows in deep narrow gorges through the coarse, debris flow fans and rock avalanche. Forced along the rim of the Turi Parwak fan, with renewed incision the river became superimposed on bedrock, and has cut a gorge some SO-6Om deep. For many centuries, this gorge has been the local base-level for the Mastuj basin, indeed, the whole upper Yarkhun. However, the degrading rock avalanche barrier controls the exit from the rock gorge.


(i)

74 Kenneth Hewitt

(iii) Figure 6. The Miragram -Parwak rock avalanche: i. Sketch map reconstructing the original rock avalanche barrier and landslide-dammed lake. ii. View looking west and downstream from the Turi Parwak fan. M = exit of Miragram valley, B = asymmetrical main accumulation of rock avalanche debris and pressure ridge or 'brandung' against the impact slope, R = ridges of original, up-valley rock avalanche deposit along the floor of the MastujIY arkhun River, X = stream cut section in rock avalanche debris, and site of: iii an eroded section in densely compacted rock avalanche, where the river cuts through a pressure ridge in the mid-section of the up valley lobe. The material consists ofdioritic (black) and granitic (white) bands of crushed and pulverised of rock.

2.3. THE TSOK-DUMORDO EVENT (Figure 7)

A post-glacial rock avalanche descended into the Dumordo River valley, six kilometres below the terminus of Panmah Glacier. It swept directly across the valley forming a huge pressure ridge against the impact slope and filling the valley to a depth of over 120m. Subsidiary lobes traveled up and down the main valley. Once more, this is an event formerly classed as moraine, specifically as the "Tsok stade" ofPanmah Glacier (Savoia-Aosta and Desio 1936, Figure 114). Yet, the material comprises fractured and powdered metamorphic rock outcropping in the west/right flank tributary, which was the source of the rock avalanche. By contrast, the moraines and outwash gravels of the Panmah are dominated by rocks of the Baltoro Plutonic Unit (Searle 1991,169-75).

The Dumordo was dammed, triggering an episode of complex deposition. The lake was filled rapidly by sediments composed mainly ofPanmah outwash gravels and sands, and of coarse sediment from a series of steep tributary valleys down which come many debris flows, snow avalanches, and torrential outwash streams from small glaciers (Hewitt 1998, 58, 69). The

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76 Kenneth Hewitt

2.4. DEVELOPMENTS IN LANDSLIDE-INTERRUPTED STREAM VALLEYS

These examples show the kinds of features that develop in impounded river reaches. While local, basin and regional contexts introduce important constraints, all examples can be said to involve two main form-generating phases. In the fIrst, sedimentation predominates, building a complex aggradational sequence and series of, mainly, constructional landforms. The rock avalanche and its mode of emplacement begins this phase. It acts as a local base level for processes and materials from contributing slopes, and the upstream drainage area. Aggradation may continue until the reservoir is infIlled by sediments. These are two or three orders of magnitude greater in volume and area than the rock avalanche itself.

Sooner or later a second, degradational phase occurs, increasingly dominated by erosional forms. The rock avalanche and other deposits are trenched and segmented, but the barrier continues to control developments according to the rate and pattern of breaching. Most Upper Indus examples record slowly degrading barriers with complicated mixes of local, back water and flank aggradation, and main channel erosion, but gradually shifting in favor of the latter. Even in the last stages, barrier remnants continue to constrain fluvial activity.

A majority of examples seem to have lasted centuries or millennia, and degraded slowly. But sudden, partial breaches have occurred, possibly due to the well-known floods from ice dams and landslide dams upvalley (Hewitt 1982). They are recorded in great boulder deposits such as those below Ghone, and may help explain the flood-transported boulders out beyond the Himalayan foothills 500 km downstream (Desio and Orombelli 1983). Rapid breaching may degrade the whole depositional complex quickly, even catastrophically.

During the degradational phase other complications and distinctive features arise. There are shifting patterns of truncation, trenching, segmentation, erosional resistance and reworking of the sediment bodies. As talus slopes, cones, fans, and deltas are segmented, secondary deposits of reworked material grow at the mouths of chutes, gullies and distributaries cut in the aggradation complex. Erosional forms influence what is preserved and revealed of the depositional assemblages, as well as creating "barrier-defended" forms (see below).

Finally, fluvial and other processes are selectively superimposed on pre-existing valleyfIll sediments or underlying bedrock, sometimes previous landslide interruption complexes. Prelandslide buried sediments or exhumed erosional features become increasingly important in the degradation phase. Streams re-excavate and begin to incise buried valleys.

3. Constructional Landforms and the Aggradational Complex

The spatial arrangement of valley-fIll can be summarised in relation to the cross-valley barrier. In addition, we can distinguish "axial" and "valley fringe" contexts (Table 2). The former involve responses of main valley drainage; the latter of mass movement or torrential deposition from surrounding slopes, tributary canyons and chutes. However, as the examples showed, progradation of valley-fringe sediments over axial ones and vice versa are locally, often generally, more important. Strong valley fringe processes may create full cross-valley forms whereas, in quiet reaches below stable slopes, lacustrine or fluvial deposits prevail to the valley wall. Aeolian


erosion and deposition superimpose yet other patterns. Some of the distinctive aspects of aggradation complexes will be highlighted.

3.1. ROCK AVALANCHE EMPLACEMENT AND BARRIER MORPHOLOGY

Pre-existing terrain and valley sediments affect the emplacement of the rock avalanche with consequences for the entire interruption epicycle. In most cases, at least three lobes developed, recording an initial cross-valley runout and divergence in up-, and down-valley directions -- the "deformed-1'" of Nicoletti and Sorriso-Valvi (1991) or "Type III" landslide dam of Costa and Schuster (1987, 9). These, and often far more complicated morphologies, represent responses of the mobile rock avalanche debris to rugged terrain. Such developments and their role in the s hape, height and stability of rock avalanche dams, are discussed elsewhere (Hewitt 1998; 2001).

Rock avalanches crossing large fans and soft or wet sediments on valley floors, may deform and entrain the substrate. In the intermontane Skardu and Shigar Basins in Baltistan, at least ten younger rock avalanches, ran out over sediments accumulated in the aggradation complex behind the Katzarah barrier. The finer-grained sediments were folded and faulted, and large quantities entrained and redeposited, creating a "chaos" of jumbled, contorted and remolded alluvial masses. Rock avalanche materials are found beneath, above, or mixed in with them (Hewitt, 2000, Figure 8). The total area of such distinctive deposits exceeds 25km2 (Hewitt 1999). Formerly they were misidentified as glacitectonic forms. Instead, they are the most extensive and diverse examples oflandslide-deformed sediment found so far, although features of this type occur in many other Upper Indus rock avalanche complexes.

3.2. LANDSLIDE-DAMMED LAKES

Cross-valley rock avalanche barriers usually formed dams impounding large lakes. The body of water could be quite complicated, drowning much of the rock avalanche, large areas up-stream and flooding two or more valleys.

The relatively young, 1911 rock avalanche and 53 km long Lake Sarez in the Pamir, suggests how many Karakoram dams looked in their early stages (Schuster 1986, 3-6). The aggradational complex would then develop according to lake geometry, the pattern ofinfdling, and conditions in contributing basins and slopes. At Satpara Lake, near Skardu, surrounding slopes and gullies contribute little sediment. A delta of coarse river gravels occupies half the former lake area (Hewitt 1998, 61). By contrast, aggradation in the Miragram-Parwak impoundment ended with a system of coalescing fans. Despite the huge Yarkhun basin above, which may well have contributed the larger volume of sediment, fans from tributaries dominate the constructional landforms, incorporating or burying lacustrine and other materials from upstream.

78 Kenneth Hewitt

Table 2. Morphological, sedimentary and erosional elements of the landslide-interruption complex, and constituents

A. Pre-emting Conditions and Constraints i) At the interruption site:

Valley morphology and geological setting Valley floor sediments (affecting rock avalanche run-out)

ii) In the interrupted basin: Topoclimatic situation - governing processes responding and contributing to the interruption epicycle Sediment delivel}' - from present-day processes (eg. glacial), and lag deposits of valley floor, valley sides or tributaries (eg. paraglacial effects)

B. Constructional landforms and the aggradation assemblage i) The Barrier zone:

Rock avalanche morphology, deposits Disturbed and redeposited substrate materials Ponds in closed depressions Spillways may include onlapping of any or all of ..

ii) The Upstream Depositional Complex: AXIAL ELEMENTS

Landslide dammed lake with fluvial inflow, mainly nival-, glacio-fluvial, sometimes glaciers Lacustrine deposits Floodplain and stream channel elements Deltas lce-contact deposition (sometimes)

V ALLEY FRINGE ELEMENTS from rock walls, canyons, chutes, hanging valleys and minor ice masses Snow and Ice Avalanche deposits Debris flow deposits Rock fall, talus and talus cones Torrent cones (Tributal}') glacier deposits N.B. widespread intertonguing or overlapping of the above; and some or all may be repeated in:

TRIBUTARIES ENTERING AN IMPOUNDED MAIN VALLEY (some or all of ii above) iii) Downstream and barrier-defended Complex:

AXIAL ELEMENTS (fluvial but sediment-starved and may be degredational in this phase) V ALLEY FRINGE ELEMENTS (as above and likely prograding in this phase)

iv) Aeolian and Dessication deposits (may oceur throughout): Dust (Loess) Sand sheets, (small) dunes, and (occasionally) large dunes and dune fields Salt pans, flats and efflorescence N.B. (Erosional) deflation features common in the aggradation phase (intertonguing of aeolian deposits with axial and valley fringe elements)

C. Erosional landforms and the degrading interruption i) The Barrier Zone:

Weathering forms (tafoni, hoodoos, earth pillars, desert varnish, frost-shattering) Gullied and reworked rock avalanche deposits Spillways, barrier breaching and boulder-armoured channels Erosional terraces Superimposed or 'epigenetic' gorges

ii) The Upstream Degrading basin sediments: Fluvial and fan terrace systems (including barrier-defended) Segmented ('Lion's Paw') fans, cones Localised deposits of reworked interruption sediments

iii) Downstream and barrier-defended complex: Terrace systems (including barrier-defended) Segmented ('Lion's Paw') fans, cones Deposits of reworked interruption sediments including Catastrophic flood deposits of rock avalanche boulders

iv) Aeolian deflatinn and depositinn (as above) v) Superimposed forms in bedrock:

Erosional straths (rock terraces) Slip-off slopes and pothole 'swarms' 'Epigenetic' gorges superimposed from interruption valley fill sediments

vi) Exhumed Erosional and Depositional Features: Pre-existing, buried glacial, fluvial and other valley fill and erosional forms Former, buried interruption complexes, (potentially) including all forms noted above, and relicts from any or all phases

of an interruption epicycle 'Isolated rocks' (combined result of superimposed and exhumed bedrock channels)

Postglacial Landfonn and Sediment Associations 79

Figure 8. The Litak event lacustrine deposits (L) behind a rock avalanche dam (arrows)on the upper Hushe River Shyok Basin, Baltistan: looking south and down river from Hushe village, showing the mUlti-year accumulation oflake beds (L). They reached a depth of about 200 m before the barrier was breached. Norin (1925) interpreted the barrier as a moraine dam, but Gansser (Burgisser et ai, 1982) realized it originated as a landslide from 8 tributary on the left/east side of the Valley.

Remnants of extensive, buff-coloured lacustrine deposits, metres or tens of metres thick, are exposed in terraces throughout the Upper Indus valleys (see Figure 2 and Owen 1996). Most record rock avalanche impoundments but were fonnerly attributed to "glacial lakes" (Burgisser et at. 1982; Zanchi and Gaetani 1996). The frne-grained, rhythmic beds are strikingly obvious remnants oflarge lakes, but rarely do they represent the bulk of material that accumulated in them (Figure 9); coarse deltaic and fluvial materials, debris flows, and sediment fans, often contributed more. Meanwhile, in the frnal stages of infilling the entire lake area may be converted to a sandur or braided flood plain, as has happened at Jut, in Kar Gah valley near Gilgit and the Naltar Lakes site (Hewitt 2001). Their rock avalanche barriers remain intact, but earlier lacustrine and other sediments are hidden under coarse fluvial materials, or vegetated islands and anastomosing channels.

Lacustrine deposits also occur in ponds on the impenneable rock avalanche, or between it and valley walls. Examples are seen at Katzarah, and on the upper surface of the Nomal Complex near Gilgit.

3.3 . SEDIMENT FANS

Drew (1875) coined the tenn "alluvial fan" for depositional features that are very widespread in the Upper Indus Basin. Today, "sediment fans" is the preferred general tenn, since debris flow

80 Kenneth Hewitt

materials are important in most cases, and many fans prove to be mixes of mass movement, lacustrine, fluvial, fluvio-glacial, and torrential fan deposits. Large fans of coarse fluvial material occur at the mouths of major tributaries. Sediment fans form a large part of most interruption complexes.

Derbyshire and Owen (1990) provided extensive, pioneering sediment logs for fans and adjacent areas along the lower Gilgit and Hunza Rivers. These are invaluable for showing the forms and complexity offormer fan sedimentation. However, these authors attribute the latter to late- glacial, including glacier-dammed conditions, and to paraglacial sedimentation at sites where I fmd them recording aggradational episodes behind rock avalanche barriers. The majority of sediment fans along the Upper Indus streams in the Karakoram, including the larger and widely discussed examples, lie in landslide-interrupted river reaches. Such forms might develop for other climageomorphic reasons, and do so elsewhere in inner Asian valleys. But their actual geometries and extent here are mainly controlled by episodic aggradation where steep tributaries entered landslide- interrupted river reaches.

The size, elevation range, lithologies and hydrology of contributing basins are important factors in individual fan geometries and composition. Hughes (1984, 263-266) showed this in an aggrading river reach of the upper Yasin valley, where fan slope and size varies considerably, as does the relative mix of debris flows and outwash gravels from small glacierised catchments.

3.4. AEOLIAN DEPOSITS

The aggradation complexes, including rock avalanches, are both major sources of wind blown material, and the main valley floor surfaces where aeolian deposition occurs. In addition, vast areas of bare rock wall and slope deposits, debris-covered glacier tongues and outwash plains, are sources of dust and sand. The more rapid mass movements, notably rock slides, rock avalanches and even certain phases of debris flows, generate large amounts of dust (Hewitt 1988). Along with a predominantly arid or semi-arid fluvial zone and powerful valley winds, conditions support intense aeolian activity (Goudie et al. 1984; Owen et al. 1992).

The bouldery, open work surfaces of most rock avalanches are infilled by dustto a greater or lesser extent. The youngest Ghoro Choh rock avalanche, at the head of Shigar valley, Baltistan, is largely buried in dunes where its boulders trap large volumes ofsand blowing off the Basha River flats, which it impounds (Hewitt 1999, Figure 10). In Shigar and Skardu Basins, dust and sand move in fierce, almost daily katabatic winds crossing terrace and fan surfaces and bare river flats behind the Katzarah barrier. There are extensive, active dune fields (Owen 1988; Searle 1991, 277).

Aeolian deposition and erosion tend to occur in close association in these valleys. Even during the aggradational phase deflationary features are widespread. Winnowing of coarse debris flow and fluvial deposits creates boulder pavements. Where there are springs, intermittent streams or rare high floods, wet deflationary features occur. Conversely, aeolian deposition continues in the degradation phase and may achieve its greatest development then.


4. Erosional Landforms oftbe Degradational Pbase

As the rock avalanche barrier is breached, impounded materials are segmented, reworked or removed, and distinctive erosional features arise.

4.1. SPILLWAYS AND BARRIER BREACHING

Most of the rock avalanche barriers appear to have survived at least until the upstream impoundments filled with water, and overflow occurred. Spillways across the lowest part of the barrier, as at Gol-Ghone and Miragram-Parwak, become well-defmed major charmels incised in rock avalanche debris. As rivers cut deeper, they usually form narrow gorges with near-vertical walls. The bed is armoured by large blocks and is relatively steep. Most of the steps in the long profiles of Upper Indus streams coincide with degrading rock avalanche barriers.

4.2. RIVER AND FAN TERRACE SYSTEMS

Extensive terrace systems are associated with degrading rock avalanches. Even where continuous, they reflect generative conditions that vary above, through and below the barriers. The vertical face or riser of a stream terrace records incision into the aggradational sequence. Trenching and segmentation of fans, or "fan terraces" (Derbyshire and Owen 1993) are part of the same developments. The terrace surface may record an aggradational level, perhaps in a former lake or flood plain, or horizontal erosion by the stream. Terraces cut in one sediment body may be veneered with deposits left by the stream. These record either the highest level of aggradation in the impounded reach, or temporary "still-stands" in the degradation phase, or a temporary reversion to aggrading conditions.

Figure 9.Superimposed ('epigenetic') gorge (EO) of the upper Yasin River, above Darkot. A roc~ aVfanche (arrows) descending from the Dulung Bar hanging valley (DB), emplaced about 280 x 10 m on the floor of the Yasin, and impounded streams from glaciers around Darkot Pass (P). A spillway across the distal (eastern) rim of the barrier cut down into a bedrock spur, draining the lake above, and the river remains incised there (EO). The cross-valley rock avalanche barrier is essentially intact, and the original river channel buried beneath it.

82 Kenneth Hewitt

In tenns of horizontal relations, divergent terraces are found above and through the upper part of the barrier zone. Parallel or convergent terraces may occur in its lower part. Convergentterraces tend to occur below the barriers. Terraces in the interruption sequence involve control, especially protection from erosion, by more resistant bodies of sediment, notably the rock avalanche. They develop above, below or between the more resistant material which deflects the stream, directs down cutting, and creates local shadow effects. Such terraces have been called "defended", but originally ascribed to lithology-controlled "accidents" within the Davisian cycle (Engeln 1942, 242; Cotton 1958, 246-8). In the present case, most relate to remnants of crossvalley rock avalanche deposits, sometimes exhumed bedrock. The Tsok-Dumordo event, described earlier, exhibits most of these features (Figure 7).

River terraces of the Upper Indus have usually been interpreted as responses to climate change and/or tectonic uplift. Terraces at particular elevations were thought to be closely related, with continuity of their development through the river system (cfLeopold, Wolman and Miller 1964,474). However, most of those investigated are not "climatic" or "tectonic" terraces (Schumm 1977,218-9; Tricart 1974). Rather, they involve rock avalanches and other barriers or related geomorphic complications. Because they are so widespread they create the illusion of linked, basin-wide developments. But no one has actually traced out the geometry, or used unambiguous morphostratigraphical, dating or tectonic evidence, to reconstruct the kinds of Quaternary terrace sequences hitherto assumed to exist.

4.3. SUPERIMPOSED OR "EPIGENETIC" GORGES AND "ISOLATED ROCKS" (Figure 9)

Streams in an interrupted reach are unlikely to become incised in valley fill directly over, and be let down into, the original river course. Rather, they can be superimposed upon bedrock flanks and spurs of the fonner valley, and may become fixed there in a bedrock gorge (Hewitt 1998,68). Thirty-five of these local bedrock gorges have been identified with rock avalanche barriers. As at Miragram-Parwak, they strongly constrain stream development. However, as degradation proceeds, the stream may also be "captured" to follow another route, abandoning the superimposed rock gorge. There are many dry epigenetic gorges associated with landslide interruptions. Active and abandoned epigenetic gorges occur frequently along the Ghizar above Gupis, along the Middle Indus upstream from Shengus to Katzarah in Skardu Basin, in the Braldu gorge, and the Hunza between Jaglot and Pasu. Similar, short superimposed rock gorge sections are associated with glaciers that enter and interrupt main stream valleys. Thirty-six of these have been identified in surveys to date clearly associated with glacier tennini or Neoglacial moraines.

There is no systematic pattern to where and when epigenetic gorges fonn. They are singularities of interruption episodes, unrelated to stream network or order, lithology, geological structures, tectonics, or patterns of glacial retreat. But their abundance makes them a characteristic feature of the fluvial zone. In tenns of bedrock incision, they mean that the river does the same work twice, or more times, in the given reach. This creates a severe, but hitherto unrecognised, complication in relations of stream incision and thalweg slopes to tectonic uplift (cf Seeber and Gornitz 1983; Burbank et al. 1997).


5. Interpretations

The descriptions identify a range of interruption-related features. At least two interpretive problems arise: how to characterize these sediment-landform associations, and their implications for high mountain stream valley development. Questions of approach are also raised.

To date, studies of Karakoram valley fill sediments have focussed on particular processes or process systems such as glacial, fluvial, or aeolian (Hewitt 1968; Goudie et al. 1984; Owen 1988; Shroder 1989). The links and comparisons emphasized are those among examples of, say, lacustrine deposits or sediment fans (Burgisser et al. 1982; Derbyshire and Owen 1990; Owen 1996). These are necessary steps, but have tended to direct attention away from the interruption phenomena and their definitive aspects.

Meanwhile, studies of large landslides and related events deal mainly with individual cases, or emphasize the singular, catastrophic aspects (Voight 1978; Heuberger et aI. 1984; Hewitt 1988). This can create an impression that they are only about extreme or "accidental" events. In the Upper Indus Basin, the numbers, scale and distribution of events suggest they are characteristic features, making their comparative and general influence more compelling.

5.1. TIME-TRANSGRESSIVE ASSEMBLAGES

In the aggradation phase a complex sedimentary architecture arises, reflecting polygeneti~, and time transgressive sedimentation constrained by the impounded basin (FigureIO). A set of forms develops causally related through the interaction of a range of processes. Impoundment geometry and history dominate the morphostratigraphic or facies "architecture" of valley fills.

Depositional sequences following the landslide interruptions involve most or all of the region's ''process systems" (Owen and Derbyshire 1993, 124). Indeed, all possible classes of fluvial, deltaic, lacustrine, glacial, aeolian and mass movement facies, may occur at different sites or times in the sequence. What is special about deposits as, say, debris flow or dune materials, is less informative than their interactions and locations within the interruption epicycle. The size or process signature of particular units is less diagnostic of the controlling environment than relationships to other sediment bodies. Sediment units or contacts that reflect critical transformations or time-transgressive developments, are keys to interpretation. Prograding sources of sediment and their interactions, overlapping, by-passing, and blocking each other, ensure complex aggradation and diachronous units. This happens without regional environmental change. Unconformities, pinching out, intertonguing or thickening of particular strata, or the relative importance of anyone type of sediment, are generally unrelated to, and tell us nothing about, larger climatic or tectonic controls.

The tendency of process-focussed classifications to look at the larger or more typical bodies of sediment of particular classes, may have contributed to the failure to identify interruption epicycles. This also fostered the view that transitional features reflect external and basin-wide controls such as tectonics, climate change, deglaciation, or post-, and paraglacial signatures.

84

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SCHEMATIC

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~ CO¥se stren go'MS o dooris fl(1N fans o aedian dust irld d<JleS

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... ~ -~ ~~ ... ~ -~ ~~ .... -:=:-- -~ <:::0. <:>. <::.. c::::. c::!Io

-------=a Figure 10. Schematic representation of the aggradational complex in a landslide-interrupted high mountain

river Valley.

Organization relative to the barrier and contributing basin(s), used in descriptions above, has some resemblance to the "non-marine" depositional systems of Miall (1984, 279-85). Proximal, medial and distal sedimentary environments relate to the contributing trunk stream(s). But accretion or progradation from valley walls, tributary fans and deltas, can be equally, or more important in the evolving sequence and constructional landforms.

However, unlike Miall's examples, these are not enclosed basins due to tectonic subsidence, or basins of interior drainage. They represent temporary impoundments by geomorphic processes in otherwise open river systems. Most are also an order of magnitude smaller in size and duration than those usually considered by sedimentologists. They have little

Postglacial Landfonn and Sediment Associations 85

chance of long-tenn geological survival, being almost certain to undergo partial or complete removal after a few centuries or millennia in this rapidly uplifting region. Yet, tectonic movements, per se, are unlikely to have a direct influence on the sedimentation. Thus, we observe the sediments and related landfonn associations at some transitional stage specific to the, geologically rather short-lived, interruption epicycles. This, and the many pennutations involved, are as important for field recognition and the diversity of events, as reconstructing a "model" depositional assemblage.

5.2. IMPLICATIONS: AN EXTENDED PARAGLACIAL?

The landslide interruptions have two distinctive effects on landscape development. They begin with catastrophic events. Once in place, however, they constrain the fluvial system and local geomorphic processes for some millennia. Stream slope and energy are reduced. Individual interruptions and the post-glacial set of them introduce a transitional, geologically temporary, but extensive disturbance regime in the river system and related geomorphic activity. Continuity in stream development and the sediment delivery system are severely and repeatedly disrupted. Responses to climate or tectonic change are, at different times, moderated, reversed or exaggerated.

The place of paraglacial sedimentation in the Upper Indus valleys is specially important but problematic. Unstable glacial materials have, and do, contribute to intennontane sedimentation-in places the larger share. Sediments and landfonns in the interruption sequences often resemble, in type and range, those identified in the classic paraglacial situations (Church and Ryder 1972). They have been explained as such (Jijun et al. 1984). Yet, it is difficult to decide exactly how paraglacial sedimentation, as nonnally understood, is present or develops. The fans and other valley-fill materials are rarely, or not primarily, emplaced as a function of a paraglacial sedimentation regime, but of the landslides and other interruptions. Valley fill excavation and terrace systems, rather than recording later phases of diminishing paraglacial sediment delivery, or the influence of changing hydrological or tectonic conditions, are most directly controlled by barrier erosion. A paraglacial effect, though present, is masked or "overprinted" by the fragmentation of the drainage system.

This applies to many sites notorious for large debris flows attributed to a paraglacial regime, but which turn out to involve remobilised rock avalanche materials (cf Owen and Derbyshire 1993, 126-127). Examples include the Upper Henzul, Nomal, Jutial and Serat landslide complexes of Gilgit and Hunza; the Gomboro and Chongo Complexes in the Braldu gorge, and a host of individual rock avalanches along the Yasin, Karambar, Bagrot, Naltar, Chapursan, Stak, Tonnik, Satpara, Shyok and Hushe valleys (Figure I). In several cases, degrading rock avalanche deposits provide most of the sediments in downstream sediment fans. Huge quantities of coarse deposition in the Bagrot fan, near Gilgit, rather than representing reworked glacial sediments, derive mainly from erosion of at least five partly or wholly breached rock avalanche barriers inside Bagrot valley (Derbyshire and Owen 1990,42-45; Hewitt, 2001). As here, remobilised impoundment debris has a regional role in generating coarse sedimentation around valley junctions, resembling paraglacial sedimentation but clearly of a different origin.

86 Kenne1h Hewitt

However, one could argue that most of the cross-valley rock avalanches are themselves in a class of events due to:

... non glacial processes that are directly conditioned by glaciation ... [and] that are the direct result of the earlier presence of the ice.

(Church and Ryder,1972, 3059)

Most of the rock walls that fail were over steepened by glacial action, and seem to have been rendered unstable by debuttressing during and after deglaciation (Hewitt 1998a; 1999; 200 1; cf Bovis and Stewart 1998). Most originated around the upper limit of the last major glaciation, and descended to ice-free valley floors. As such they, and the interruption epicycles following from them, are part of a range of responses in valleys recovering from glaciation.

The spatial and temporal form of this kind of transition, as it affects a whole basin, is not known. No systematic sequence of interruptions has been discerned by stream order or elevation in particular valleys (Hewitt 1999). Barrier-related aggradation does tend to be more prevalent at higher altitudes, closer to present glaciers, and degradation more advanced on the main Indus streams. But it is not entirely clear why this should be so, and cannot be taken at face value in the absence of dating of the events. Cruden and Hu"s (1993) "exhaustion" model for post-glacial rock avalanches could apply. That is: a relaxation process in which catastrophic landsliding tends to diminish with time since glaciation. Likely sites are progressively "used up" and rendered more stable, in a process whose temporal shape is not unlike the classic paraglacial (Church, this volume). Of course, it is complicated by continuing slope steepening and ice fluctuations in the zones of neoglacial advance and present-day glaciers. It remains to be seen whether new dating methods will allow us to test the hypothesis.

6. Conclusion

Extensive valley-fill sedimentation and related landforms characterize the trans-Himalayan Indus stream valleys and have been the focus of modern geomorphological and Quaternary studies. Here, it is suggested they mainly reflect fragmentation of the drainage system by post-glacial, geomomhic events. The more influential interruptions are rock avalanche barriers that last a few millennia, in some cases perhaps, tens of millennia. They have constrained and modified sediment yield, sedimentation and stream development over the past 10,000-30,000 years, at least, and in river systems draining more than 100,000 km2 of the most rugged terrain on Earth.

There is also a long history of stream interruption by other processes, notably avalanches, glaciers and debris flows, or the growth of sediment fans (Mason 1935; Hewitt 1982, 1998a; Hughes 1984; Hughes and Nash 1986). Temporary impoundments occur in most years. But the more frequent events generate smaller, lower and shorter -lived impoundments, and generally occur in river reaches already affected by rock avalanche barriers. This suggests a hierarchy of interruptions by size and duration, and river systems chronically affected by non-fluvial processes.

Important features in the fluvial zone do not derive, in a direct causal sense, from the typical fluvial factors of network position, discharge, or sediment yield. Instead, they arise from


the local interaction of fluvial with other geomorphic processes. Rather than recording the late glacial legacy along the stream valleys, as is widely believed, these developments have generally buried, modified or removed it. Nor can we any longer say that ''the glacial system" is dominant here (Owen and Derbyshire 1993). The interruptions modify stream flows and redistribute the large sediment loads coming from contemporary, heavily glacierised headwaters throughout the fluvial zone. Wash loads, which dominate sediment delivery from the mountains, may be swept right through in high summer flows. But great quantities of sand and, especially, coarser loads, move step-wise through the dozens of existing, incompletely removed interruptions. They have and do enter long term storage or 'queues' of sediment behind the barriers

The number of interruptions is sufficient to affect development of the entire fluvial zone and to create forms characteristic of the region. Past work had not recognized the role of landslide interruptions in valley fill sedimentation, or their widespread occurrence. Hence, revisions of the late-, post- and paraglacial glacial chronology also seem necessary.

References

Abele, G.O. (1974) Vom Eis gefonnte Bergstwz\andschaften: zur Frage der glazialen Umgestaltung der BergstOrze vom Sierre, Flims, Ems und vom Fernpass, Zeischriftf. Geomorphologle, 120-147.

Adams, J. (1981) Earthquake-dammed lakes in New Zealand, Geology, 9, 215-219. Bovis, MJ. and Stewart, T.W. (1998) Long-term deformation of a glaeially undercut rock slope, southwest British

Columbia, in Moor, D. and Hunge, O. (eds.) Proa:edings Eighth International Congress, International Association for Engineering Geology and the Environment AA Balkema, Rotterdam, 1267-1276

Burbank, D.W., Leland, J., Fielding, E., Anderson, R.S., Brozovic, N., Reid, M.R. and Duncan, C. (1996) Bedrock incision, rock uplift and threshold hillslopes in the northwestern Himalayas, Nature (London) 379, SOS-5 I O.

BOrgisser, H.M., Gansser, A, and Pika, J. (1982) Late Glaeia1lake sediments of the Indus valley area, northwestern Himalayas, Eclogae geologUche Helvetica 75/1, 51-63.

Brunsden, D., and Jones, D.K.C. (1984) The geomorphology of high magnitude-low frequency events in the Karakoram mountains. In Miller KJ. (ed.) The International Karakoram Project" Cambridge University Press, Cambridge. I, 383-388.

Church, M. and Ryder, J. (1972) Coates, D.R. (ed.) (1977) Landslides, Reviews in Engineering Geology, Geological Society of America, Boulder,

Colorado. Costa, J.E. and Schuster, R.L. (1987) The Formation and Failure of Natural dams, Open-File Report, US Geological

Survey, Vancouver, Washington, 39. Cotton, CA, (1958) Geomorphology 7th edition, Whitcomb and Tombs, Christchurch, NZ. Cronin, V.S. (1989) Structural setting of the Skardu intermontane basin, Karakoram Himalaya, Pakistan. In L.L.

MaIinconico, Jr. and RJ. Lillie, (eds.) Tectonics of the Western Himalayas Geological Society of America Special Paper, 232, 203-216.

Cruden, D.M. (1985) Rock slope movements in the Canadian Cordillera. Canadian Geotechnical Journal, 14, 528-S40. Cruden, D.M. and Hn, X-Q (1993) Exhaustion and steady-state models for predicting landslide hazard in the Canadian

Rockies, Geomorphology 21, 18-30. Cunningham, A. (1854) LadaIc, physical. statistical and historical, with notices of su"oundlng countries, London

(Reprint: Sagar Publications, New Delhi, 1970). Dainelli, G. (1922) Studi sui Glaciate. In "G. Dainclli, Ed.) Relazione Sclentijiche della Spedizlone Italiano de Filippi,

nell'Himalaia, Caracorum e Turchestan Cinese (1913-14), Serei II. Resultati geologici e geografici, vol. 3 + volume of plates, Zanichelli, Bologna.

Derbyshire, E. (1996) Quaternary glaeial sediments, glaeiation style, climate and uplift in the Karakoram and northwest Himalaya: review and speculations, Palaeogeography, Palaeoclimatology, Palaeoecology, 120, 147-1S7.

Derbyshire, E., Jijun, L., Perrott, FA, Shuying, X. and Waters, R.S. (1984) Quaternary glaeial history of the Hunza Valley, Karakoram mountains, Pakistan, In K.J. Miller, (ed.), The International Karakoram Project, 2, 456-495.

88 Kenneth Hewitt

Derbyshire, E., and Owen, L. (1990) Quaternary alluvial fans in the Karakoram Mountains. In. A. H. Rachocki, and M. Church, (eds.) Alluvial Fans: afield approach, Wiley, New York, 27·53 ..

De Terra, H. and Patterson, T.T. (1939) Studies on the Ice Age in India, Carnegie Institute, Washington, D.C. Desio, A., and Orombelli, G. (1983) The "Punjab erratics" and the maximum extent of the glaciers in the Middle Indus

Valley (pakistan) during the Pleistocene, In Accademia nazionale dei Lincei, Classe di Scienze fisiche, matematiche e naturali, Ser. VII, XVI, Memorie, Sezione.lla, 135·180.

Drew, F. (1873) Alluvial and lacustrine deposits and glacial records of the Upper Indus Basin, in Quarterly Journal GeolOgical Society, London, 29, 441-471.

............... 1875, The Jarnmoo and Kashmir Territories: a geographical account London, Standford. Eisbacher, G.H. (1979) Cliff collapse and rock avalanches (sturzstroms) in the Mackenzie Mountains, northwestern

Canada, Canadian Geotechnical Journal, 16 309·334. Eisbacher, G.H. and Clague, J.J. (1985) Destructive Mass Movements inHigh Mountains, Geological Survey of Canada,

Paper84·16,~a

von Engeln, 0.0. (1942) Geomorphology: SystematiC and RegiOllQI, Macmillan, New York. Evans, S.G. (1986) Landslide damming in the Cordillera ofwestem Canada, in Schuster R.L. (ed.) Lmulsllde Dams:

Processes, Risk, and Mitigation Geotechnical Special Publication No.3, American Society of Civil Engineers, New York, 111·130.

Ferguson, R.I. (1984) Sediment load of the Hunza River. in K.J. Miller (ed.) The International Karakoram Project, 2, Cambridge University Press, Cambridge, 1984, 581·598.

Fort, M. (2000) Glaciers and mass wasting processes: their influence on the shaping of Kali Gandaki Valley (higher himalaya ofNepa), Quaternary International.

Fort, M, Burban, D.W. and Freytat, P. (1989) Lacustrine Sedimentation in a Semiarid Alpine Setting: An Examination from Ladakh, Northwestern Hamilayas, Quaternary Research, 31, 332·350.

French, H. (1987) Periglacial Geomorphology in North America, Progress in PhySical Geography, 11/4, 533·547. Goguel, 1. and Pachoud, A. (1972) Geoogie et dynarnique de I'ecroulement du Mont Granier, dans Ie massif de la

Chartreuse, en novembre 1248, Bulletin du Bureau Recherche GeolOgie Min;eres (2nd Ser.), Sect. III. 1,29·38. Goudie, A.S .. , Brunsden, D., Collins, D.N., Derbyshire, E., Ferguson, R.J., Jones, D.K.C., Perrott, F.A., Said, M., Waters,

R.S., and Whalley, W.B., (1984) The geomorphology of the Hunza Valley, Karakoram Mountains, Pakistan in K.J. Miller (ed.) The International Karakoram Project, 2, Cambridge University Press, Cambridge, 359-410.

Haserodt, K. (1989) Zur pleistoZ8nen und postglazialen Vergletscherung zwischen Hindukusch, Karakorum und Westhimalaya In: Haserodt, K. (Ed.) "Hochgebirgsrlume Nordpakistans im Hindukusch, Karakorum und Westhimalaya" Beittage und Materialen zur Regionalen Geographie, 2, 182·233, Institut ftlr Geographie der Technischen Universitat, Berlin.

Heim, A. 1932. Bergsturz und Menschenleben. ZOrich, Fretz and Wasmuth. Heuberger, H., Masch, L., Preuss, E., and ScluOcker, A. (1984) QuaternlUy landslides and rock fusion in Centtal Nepal

and the Tyrolean Alps, in MOUIIIain Research and Development, 4:4, 345·362. Hewitt, K. (1968) Studies in the Geomorphology of the Mountain regions of the Upper Indus Basin. Unpublished Ph.D.

Dissertation, King's College, University of London, 2 vols . ................... (1982) Natural dams and outburst floods of the Karakoram Himalaya.,in Hydrological Aspects of Alpine

and High Mountain Areas Glen, J. (ed) IAHSIAISH Publication, 138,259·269 . .................. (1988) Catastrophic landslide deposits in the Karakoram Himalaya, Science, 141, 64-67 . ............. (1998a) Himalayan Indus stteams in the Holocene: glacier·, and landslide·'interrupted' fluvial systems, in

Stellrecht, I. Karalroram·Hindukush·Himalaya: Dynamics ojChange, 4/1, Rodgers KOppe Verlag, KOln, 3·28 . .. ...... ..... (1998b) Catastrophic landslides and the Upper Indus stteams, Karakoram Himalayan, northern Pakistan jp ed.

Shroder, J.T. jr Mass Movement in the Himalaya, Geomorphology: International Journal oj Pure and Applied Geomorphology, 26/1·3, Special Issue, "Mass Movement in the HImalaya", 47·80

............. (1999) Quaternary moraines vs catastrophic rock avalanches in the Karakoram Himalaya, Northern Pakistan. Quaternary Research, 5113, 220·237

...... ....... (2000) Styles of rock avalanche depositional complex in very rugged terrain, Karakoram Himalaya, Pakistan. in Evans S.G. (ed.) Catastrophic Lmulslldes: processes, events, environments, Reviews in Engineering Geology, Geological Society of America, Boulder, Colorado.

.. ........... (2001) Catasttophic rock slides and the geomorphology of the Hunza and Gilgit Basins, Karakoram Himalaya, Erdkunde 55, 72·93.

HsO, KJ. (1975) Catastrophic debris streams (sturzstrorns) generated by rock falls. GeolOgical Society oj America, Bulletin,86, 129·140.


............... (1978) Albert Heim : observations oflandslides and relevance to modem interpretations. in B. Voight. (cd.) Roclc3lldes and Avalanches, I Natural Phenomena, Elsevier, New York, 70-93.

Hughes, R.E. (1984) Yasin Valley: the analysis of geomorphology and building types, in Miller, K.n (cd) 1984, The I1Iternatl01lai Karakoram p,.oject, v. 2, Cambridge University Press, Cambridge, 253-288.

Hughes, R. and Nash, D. (1986) The Gupis debris flow and naturaJ dam, July 1980, Disasters, 10, 8-14. Hungr, O. (1989) Mobility of rock avalanches. National Research Ce1ltre/or Disaster Preve1ltlon, Bulleti1l. Tsukuba,

Japan, I-B. Hutchinson, J.N., (1988) General report: morphological and geotechnical parameters oflandslides in relation to geology

and hydrology, in Landslides. Ch. Bonnard (ed.) Proceedings, 5th International Symposium on Landslides, Lausanne, Switzerland, 10-15 July, I, 3-35.

Jijun, L., Derbyshire, E., and Shuying, X. (1984) Glacial and paraglacial sediments of the Hunza Valley, North-West Karakoram, Pakistan: a preliminary analysis. 111 K.J. Miller, (ed.) The I1Iternationai Karakoram Project. ,2,496-535.

Jones, D.K.C., Brunsden. D, and Goudie, A.S. (1983) A preliminary geomorphological assessment of part of the Karakoram Highway, Quarterly Journal 0/E1Igineer/1lgGeology, London 16, 331-355.

Kalvoda, J. (1992) GeomorpholOgical Record 0/ the Quaternary Orogeny 111 the Hami/aya and the Karakoram, Developments in Earth Surface Processes 3, Elsevier, Amsterdam.

Krieger, M.H., (1977) Large landslides, composed 0/ Megabreccia, Interbedded 111 Miocene Basi1l Deposits, SoutheastemArlzona, United States Geological Survey, Professional Paper 1008, Washington, D.C.

Laznicka, P. (1988) Breccias and coarse fragme1ltltes: petrology, environments, associations, ores, Developme1lts i1l Ec01l0mic Geology, 25, Elsevier, New York.

Leopold, L.B., Wolman, M.J. and Miller).P. (1964)f/Ullial Processes /11 Geomorphology, W.H. Freeman, San Francisco. Mason, K. 1929 Indus floods and Shyok Glaciers, Himalayan Journal 1, 10-29. Miall, A.D. (1990) Prl1lciples 0/Sedime1ltary Basi1l Analysis, 2nd edition, Springer-verlag, New York. ed. Miller, K. 1. (1984) The International Karakoram Project, 2 vols., Cambridge U.P. Cambridge. Mollard, J.D. (1977) Regional landslide types in Canada in D.R Coates, (ed.) Landslides, Reviews in Engineering

Geology Geological Society of America, Boulder, Colorado, 29-56. Mudge, M.R. (1965) Rock avalanche and rockslide avalanche deposits at Sawtooth Range, Montana, GeolOgical Society

0/ America Bulleti1l, 76, 1003-1014. Nicoletti, P.G. and Sorriso-Valvo, M. (1991) Geomorphic controls of the shape and mobility of rock avalanches.

GeolOgical Society 0/ America Bulleti1l, 103, 1365-1373. Norin, E. (1925) Preliminary notes on the late Quaternary glaciation of the North-Western Himalaya. Geografiska Anna1er

7, 165-194. Oestreich, K, (1906) Die Taler des nordwestlichen Himalaya. In: Petermanns Geographische Mitteilungen, 33,

Erganzungsheft 155. Owen, L. (1988) Terraces, Uplift and Climatic Change in the Karakoram Mountains, Northern Pakistan, Unpublished

Ph.D. Dissertation, University of Leicester . ............... (1991) Mass movement deposits in the Karakoram Mountains: their sedimentary characteristics, recognition

and role in Karakoram landform evolution. 'kttscbrfftftJr Geomorpholog/e, N.F., 35, 401-424. Owen, L. (1996) Quaternary lactustrine deposits in ahigh-energy semi-arid mountain environment, Larakoram Mountains,

Northern Pakistan, Journal o/Quaternary SCience, 11(6),468-483. Owen, L. and Derbyshire, E. (1993) Quaternary and Holocene Intermontane Basin Sedimentation in the Karakoram

Mountains, in J.F. Shroder, Jr., ed. Himalaya to the Sea: Geology, Geomorphology and the Quaternary. Owen, L., White, B.J., Rendell, H. and Derbyshire, E. (1992) Loessic Silt Deposits in the Western Himalaya: Their

Sedimentology, Genesis and Age, Catena, 19,493-509. Paffen, K.H., Pillewizer, W. and Schneider, H-J. (1956) Forschungen im Hunza Karakorum, Erdlamde, 10, 1-33. Porter, S.C., and Orombelli, G. (1980) Catastrophic rockfall of September 12, 1717 on the Italian flank of the Mont Blanc

massif. 'kitscbrfftf Geomorphologie N.F., 24:2, 200-218. Savoia-Aosta, A. and Desio, A. (1935) La Spedlzi01lfl Geograftca Itallana nel Caracorum: Storia del Viaggio e Remltati

Geograftcl, Bertarelli, Milano-Roma. Scheidegger, A.E. (1972) On the prediction of the reach and velocity of catastrophic landslides. In: Rock Mechanics 5,

231-236. Schlagintweit, A. (1857) Report on the progress of the Magnetic Survey oflndia and researches connected with it ftom

May to November 1856, in Survey 0/ India Report No 9 Debra Dun.

90 Kenneth Hewitt

Schneider, H-J. (1959) Zur Diluvial Geschichtc des NW-Karakorum, Mittelirmgen Geographla Gessebchaft. MfJnchen, 44,2012-216.

Schumm, SA (1977) The Fluvial System, Wiley-Interscience, New York. Schuster, R.L. (1986) A perspective on landslide dams. in Schuster R.L. (ed.) Ltmtlslide Dams: Processes, Risk, and

Mitigation, Geotechnical Special Publication No.3, American Society of Civil Engineers, New York, 1-20. Selby, MJ. (1993) Hlllslope MaterilIls and Processes 2nd edition, Oxford University Press, Oxford. Searle, M.P. (1991) Geology and Tectonics o/the Karakoram Mountains, New York, John Wiley. Seeber, L. and Gomitz, V. (1983) River profiles along the Himalayan arc as indicators of active tectonics, in

Tectonophysics 92, 335-367. Shrader, J.F. Jr. (1989) Slope failure: extent and economic significance in Afghanistan and Pakistan in Brabb, E.E. and

Harrod, B.L. Landslides: &tent and Economic Significance, A.A. Balkema, Rooterdam, 325-341. Shrader, J.F.jr. (ed.) (1993) Himalaya to the Sea: geology, geomorphology and the Quaternary, Routledge, New York,

159-183. Shroder, J.F. Jr., Saqib Khan, M., Lawrence, R.D., Madin, I.P. and Higgins, S.M. (1989) Quaternary glacial chronology

and neotectonics in the Himalaya of Northern Pakistan, in Tectonics o/the Western Himalayas, Geological Society of America Special Paper. 232, 275-294.

Shrader, J.P Jr., Owen, L,. andDcrbyshire, E. (1993) Quaternary glaciation of the Karakoram and Nanga Parbat Himalaya, in Himalaya to the Sea: geology, geomorphology and the Quaternary" (J.P. Shrader, Jr. cd.), Routledge, New York, 132-158.

Strachey, R. 1853 A Physical Geography of Western Tibet, JOUI7I(llo/the Royal Geographical Soctety, 23, 1-69. Thomson, T. (1852) Western Himalaya and Tibet, London. Tricart, (1974) Structural Geomorphology (translated: S.H. Beaver and E. Derbyshire), Longmans, London. von Engeln, 0.0. (1942) Geomorphology: SystematiC and Regionol, Wiley, New York. ed. Voight. B. (1978) Rockslides and Avalanches, J Natural Phenomena, Elsevier, New York. Yarnold, lC., (1993) Rock-avalanche characteristics in dry climates and the effect oftlow into lakes: Insights from mid

Tertiary sedimentary breccias near Artillery Peak, Arizona, Geological Society 0/ America Bulletin, lOS, 345-360. Yarnold, J.C., and Lombard, J.P. (1989) A facies model for large rock-avalanche deposits formed in dry climates, in

Conglomerates in Basin Analysis: a Symposium dedicated to A.O. Woo4ford (I.P. Colburn, P.L. Abbott and J. Minch, eds.,) Pacific Section S.E.P.M. 62, 9-3.

Whitehouse,l. E. (1983) Distribution ofJarge rock avalanche deposits in the central Southern Alps, New ZealandJOUI7Kll o/Geology and Geophysics. 26, 271-79.

Zanchi, A. and Gaetani, M. (1996) Introduction to the geological map of the North Karakoram terrain from Chapursan Valley to the Shumsha\ Pass, 1:150,000 scale, Rlvista ltaliana Paleontologta, Stratigraphia 10011, 125-136.

Zietler, P.K. 1985, Cooling history of the NW Himalaya, Pakistan, Tectonicif. 4, 127-151.


Acknowledgements

I thank Dr. Ian Brookes for reading and helpful comments on a draft of the chapter. Parts of the research on which this paper is based were funded by the International Development Research Centre, Ottawa, and the OfficeofResearch, Wilfrid Laurier University, Waterloo. I thank officers of the Water and Power Development Authority, Pakistan and local guides from the Northern Areas for assistance in the field, Ms. Pam Schaus for preparing the figures, Ms. Jo-Anne Horton for preparing the text.

FLUVIAL SEDIMENT TRANSFER IN COLD REGIONS

Michael Church Department of Geography The University of British Columbia Vancouver, British Columbia, Canada V6T IZ2 [email protected]

Abstract

The fluvial signature of sedimentary disturbance can be likened to the response of a cascade of linear reservoirs, at least qualitatively. The reservoirs are fluviallandfonns, including channel deposits. The signature includes contemporaneous degradation and aggradation at different places in the landscape. The most significant fluvial disturbance to have affected terrestrial landscapes in recent earth history is recurrent, unstable continental glaciation in the Northern Hemisphere. Successive glaciations yielded major fluvial disturbances, the residual effects of which are still detectable in undisturbed landscapes. However, the fluvial response to glaciation, which may be termed the 'paraglacialresponse' is similar to the response to other sedimentary disturbances. In most places, residual paraglacial effects are now overwhelmed by other, more recent fluvial disturbances, in particular the effects of human land use. Complications of the conceptual model include the variable mobility of sediments of different calibre and the superimposition of various disturbances. There are few data to allow quantitative analysis of the model, but it presently serves to indicate aspects of fluvial sediment transport and storage that require attention in order to improve our understanding of fluvial sediment transfer in the landscape.

1. Introduction

Weare used to thinking of fluvial sediment transport as a hydraulically driven, mechanical process that is forced by stream flows. In fact, the transfer of sediment through a drainage basin on geomorphologically significant time scales (say 1()2 to l06 a) is mainly a matter of sediment storage. Sediment remains immobile for most of the time because, at most places in subaerial landscapes, the forces necessary to mobilise material and transport it

93

K. Hewitt et al. (eds.), Landscapes a/Transition, 93-117. © 2002 Kluwer Academic Publishers.

94 Michael Church

are experienced only occasionally. This circumstance creates most valley bottom fluvial landforms. The process remains constrained, of course, by continuity, so that degradation and aggradation are simultaneous processes occurring at different (but connected) places in the landscape. Over significant periods, these connections create characteristic signatures of geomorphic response to primary disturbances of the land surface. Prominent disturbances include wildfire, local or regional human manipulation of the land surface, regional climatic change, and tectonics. The most dramatic primary sedimentary disturbances to visit Earth's surface within the past 2 Ma have been the recurrent, unstable Pleistocene glaciations of the northern hemisphere.

The pattern of subaerial sedimentary response exhibits some general characteristics that are susceptible to model representation. The pulse of sediment that moves downstream following initial delivery of a sediment charge to a stream channel decays monotonically over time (Figure la). The more sudden is the initial delivery, the more rapid is the initial rise in sediment transport. The decline is approximately exponential. For an impulsively delivered sediment charge -- such as a bank collapse or landslide into a watercourse -- that is readily entrained into the streamflow, the sediment transport away from the point of entry behaves like the drainage from a linear reservoir (Figures 1 b, 2), which is strictly exponential. Lisle (2000) has investigated sediment evacuation from such simple sources and found that, indeed, the process is approximately exponential. But trapping points occur downstream, so that sediment is intercepted and onwardly remobilised at a later time. Sediment in a drainage system moves through a cascade of reservoirs. In general, the local rate of remobilisation will depend upon the amount of sediment that is stored in a particular sediment trap, or reservoir. Ifwe think of this cascading process as a sequence of identical linear reservoirs, the downstream response to a slug input of sediment behaves as a Gamma variate in time (Figure 2). At successive observing points downstream, the sediment "wave" becomes progressively attenuated in magnitude and more extended in time. It behaves dispersively.

This simple picture is distorted by the circumstance that successive sedimentary ''reservoirs'' may not act in a similar way and, even more so, by the fact that the readiness with which sediments are transported depends strongly on the size and density of the individual particles, so that sediment transport is also a sorting process. Mixtures of sediment segregate in the cascade into sub-populations with different transit times through the landscape, and so fluvial sediment deposits exhibit sorting.

It appears as if this simple model holds tolerably well not just for point infusions of sediment into the fluvial system, but also for areally extended sources. Figure 3 illustrates the response to wildfIre. There is now, however, a spatial complication. Part of the sedimentary signal originating at the headmost point of disturbance must traverse downstream through a reach subjected to a similar disturbance, so the signal may be reinforced for some distance. At downstream points, a response time will be observed after the initial disturbance that is not merely the consequence of the input sediment traversing a series of reservoirs. If the disturbance is visited on the area with varying intensity, the reinforcement may vary spatially. A further complication is that the initial disturbance may continue for some fmite period, so temporal reinforcement occurs as well. Simple mathematical treatment of the model slips away.

30

~

";"(1) 20 '" .s o

10

o

Fluvial Sediment Transfer in Cold Regions

(a) z 75 o

~ ~ w o z --50 8-1-';0) z.s w :e o w en 25 o w o z W Q. en ::>

Q

en o~----------~----------------~~ Nov. 10 Nov. 11

1983

(b)

time Figure 1. Simple fluvial sediment response: (a) Suspended sediment transport in Carnation Creek, British Columbia, in response to storm runoff. Carnation Creek is a partially logged drainage basin with area A.J - 10.1 1an2 (Water Survey of Canada Stn. 08HB048; data from Tassone, 1987). Rapid sedimentary response of this type indicates sediment mobilisation from flushing gullies or from strearnbank collapse relatively close to the gauge; (b) response of a simple linear reservoir to an impulsive load.

95

96

n - step reservoir

Michael Church

1 - step reservoir after ~ ~-mpUISive disturbance. V

Q = kVe -k.

k (t-Il is the reservoir constant

time

" = V knt n.1 e -k • Rn Vnl (n _ 1)!

This is Gamma variate

Figure 2. Mass transfer through a sequence of linear reservoirs. For simple representation, the reservoirs are assumed to be similar.

But what these mechanistic complications imply for the way we observe sedimentary disturbance depends upon the spatial scale and temporal resolution of our observations. At more extended spatial scales, and over longer periods of time, observations lose local resolution so that it may remain appropriate to specify a simple model of sedimentary disturbance even though the actual local history may have been quite complex. This is particularly true of observations at geomorphological or geological time scales, for which the only observations are the sedimentary deposits (reservoirs) themselves, as in Figure 3b.

2. Some characteristics of fluvial sedimentary response to disturbance

To examine the signature of sedimentary disturbance in a systematic way, we need some defInitions. Figure 4 presents a set of defInitions for temporal response. It is not an accident that these defInitions bear some resemblance to those used for hydrograph analysis. A fluvial sediment transfer episode is conceptually similar in many ways. An important difference, however, is the very much longer time scale for a sedimentary disturbance to traverse the landscape because of the discontinuous transport of most of the material. It is also necessary to have some measure of the scale of the initial disturbance. The area of the disturbance will be adopted for that purpose in this discussion because it is an accessible measure, although area x duration might be a more complete measure.

-'5-

~ "'g

~ >=

~ ~


140

120 (a)

100

80

60

40

20

0 1965 1970 1975 1980

YEAR

(b) _floodPlain afll2rac1alll0(1

___ T2 terrace ____ T3 terrace T4 terrace

---~--~

• fire ·related debris flows

D probable fire· related sedimentation

CALIBRATED YEARS CALIBRATED YEARS AD

1800

Figure 3. Sediment yield response to wildfire: (a) simple response of a S lcm2 drainage basin in Washington state, U.S.A. Qualitatively similar responses more than 2Sx greater were observed in two other small basins that experienced debris flows; (b) Summary of fluvial response to fire and other alluvial activity in small drainage basins in Yellowstone National Park, Montana-Wyoming border, U.S.A. The graph of sediment yield is constructed from the distribution of 3S calibrated radiocarbon dates on charcoal and other organic materials enclosed in the fluvial sediment deposits. The graph consists of the summed probability densities associated with the individual date determinations. Graphical smoothing interpolated by the writer. Association of the dates with aggradational phases is shown above the ''yield curves". (Adapted from Meyer et 01., 1992, their Figure 2).

97

98 Michael Chmch

It is difficult to assemble data on the characteristics of sedimentary disturbance. In particular, most reported sediment budgets do not distinguish production from yield. This problem is evident in the sediment mass balance equation, 0 = I - 4S, in which I is production (specifically, input to a well-delimited system), 0 is yield (output), and 4S is the storage change in the system. Transient sediment storage is, of course, what creates long relaxation times after a disturbance event, but part of that storage may be very long term; that is, for a bounded sedimentary system, IIdt p f 0dt, necessarily, at least not for a very long time. In this circumstance, persistent aggradation or persistent degradation occur within the bounds of the system. James (1989) gives a striking example of this complication in the context of reintetpreting the classically described (Gilbert 1917) effects of hydraulic mining in the Sierra Nevada ofCalifomia.

Anatomy of sedimentary disturbance

weathering + response .~.. th . = magnll1catlon

I wea ermg

subaerial weathering

time

I ~F~~~ 1-... ..---- relaxation time ---~.-I Figure 4. Definitions for a sedimentary disturbance. Sediment yield may be considered in volumetric or in mass terms, the difference between them being mineral density. "Subaerial weathering" indicates the fluvial sediment yield derived from normal weathering processes acting in equilibrium with erosive agents in the landscape. It is introduced here as a normative reference rate of sediment yield.

Furthermore, different specific sources may behave in characteristically different ways in a drainage basin, and so may different sediment size grades. Landforms are, of course, inherently integral measures of a sediment budget, or a partial sediment budget, over an extended period of time, but absolute chronology is necessary in order to intetpret them.

Data of fluvial sediment response are given in Table 1. Response time (cf. Figure 4), the characteristic about which we appear to have most information, is graphed for several drainage basins in Figure 5. For the widest range of data, the response time of a sedimentary disturbance appears to scale approximately with drainage area, that is with V, L being a characteristic length scale of a drainage basin upstream from a measurement

Fluvial Sediment Transfer in Cold Regions 99

point. and also a measure of the distance of travel of the sediment wave in the landscape. Viewed conversely, the distance of propagation ofa sedimentary disturbance appears to grow very nearly as tll2. Sediment wave propagation appears to be different, then, than sediment virtual velocity, defined as the speed of travel of labelled sediment grains in the system, rest periods included (Hassan et al. 1992). We expect virtual velocity to scale directly with t (i.e., L oc t, hence drainage area A.,. cc f) so long as the grains remain active. (It must be borne in mind that the sediment wave speed, or celerity, is distinct from the speed of travel of individual sediment particles.) This is consistent with the dispersive propagation of the disturbance through a series of reservoirs. It results from the propensity of grains sooner or later to drop into long-term storage, in an inactive channel bar or in a floodplain.

DISTANCE (km)

10° 101 102 105r-------~------,-------~------_.------~ ....

Sedl",enf response tI",. .. .. .... .... .... ,,',\...... , .

...... 0-.. /,. 6

.. "/' Yello"".tone ".-"L" •

.... .... ....

......... ....... .lfenlehe

...... Coon Creek

" "

.... .. ~ ....

DRAINAGE AREA (km2 )

Figure S. Sediment response time plotted versus drainage area, the latter being used as a measure of the magnitude of the disturbance. A distance scale (for travel distance of the disturbance) is also given as L = "'A.s(upper abscissa).

Three points in Figure 5 that do not scale with t derive from studies of the impact of hydraulic mining on stream channels in California. The study of this case by G.K.Gilbert (1917) originally introduced the concept of sediment movement as a wave through river channels. Gilbert's data of Yuba and Sacramento rivers (extended by Graves and Eliab 1977) were based on channel bed elevations, not upon sediment transport measurements. James (1989,1991) critically reexamined the observations and concluded that the history of channel bed elevations (Figured also in Meade et al. 1990) does not reasonably reflect the history of sediment movement since other factors, including dyking,

100 Michael Church

dam construction and flood history, influence the evolution of channel morphology. Nonetheless, the apparent response time in these drainage basins -- the most reliably observed parameter of sediment propagation -- remains anomalously short in comparison with t-scaling. Aggradation ended only shortly after the cessation of direct delivery of mine spoil to stream channels in 1884. An important reason for this may be that direct delivery occurred at a limited number of points over a relatively short period. The sources were essentially point sources at streambank, whereas the main sequence of results in Figure 5 represent areally extensive sources within the drainage basin. This opens the distinct possibility that different scaling laws may describe sediment movement off the land surface, and sediment movement through channels. In fact, 2 of the California points and two others scale reasonably with f, which is the expected scaling by virtual velocity of channel transport. But, by any standard, the Sacramento datum remains anomalous.

The summary picture that we can draw from this investigation is that space-time response for highly mobile sediments, or for sediments directly entrained into and remaining in stream channels, follows the expected relation for some defined sediment virtual velocity, but that sediment mobilisation from the land surface, with abundant opportunity for intervening storage at locations such as slope base (cf. Costa 1975) or in floodplains departs toward a scaling in which A oc t, or L oc f·5. Individual cases may fall between these limits.

Response magnitude depends upon how much sedimentary material is stored in the region of disturbance and upon how readily it may be mobilised, hence upon regional history and material properties. A major problem associated with quantification of response magnitude to an event disturbance is knowing what the reference rate of subaerial dendudation (presumably, approximately equivalent to rock weathering rate under zonal climate) would be. The remaining disturbance parameter, relaxation time, also is difficult to estimate in most cases because of the serial occurrence of overlapping disturbances, or because observations do not yet permit an estimate to be made.

The temporal view developed above gives rise to a corresponding spatial signature. The spatial signature of fluvial sedimentary response has not been much investigated because it requires simultaneous measurements of sediment transfers at a large number of places in the regional landscape. These measurements, furthermore, must be averaged over a sufficient period to be representative of the long term movement of sediment through the fluvial system. The writer knows of no such measurements. Shortterm variability in fluvial sediment transport is relatively large because sediment transport (G, understood to be on a mass basis) is a positively non-linear function of streamflow (Q); i.e., G oc Qb, in which b '" 1.5 to 2.5 for suspended sediment and systems with sandy bed material, and may be much higher in competence-limited gravel-transporting systems. These conditions magnify variations in flow and make it difficult to discern long-term trends in relatively short records.

Figure 6 presents an approximation of the spatial response to sedimentary disturbance. It is an approximation only because the data are derived from contemporary measurements of suspended sediment transport at stream-gauging stations, and individual records vary from 1 year to about two decades in length. Hence, they are not strictly contemporaneous records, and some temporal variability must be aliased into the displayed


spatial variation. This Figure is the signature of fluvial sedimentary response to the Fraser (late Wisconsinan) glaciation of British Columbia, displayed for t '" 12ka (i.e., the present day) if we designate the initiation of the disturbance to be the approximate time of deglaciation. Of course, deglaciation did not happen everywhere in the region at once but, because of the long time scale of the disturbance, the measurements are, in fact, acceptably "simultaneous" (but not equally representative). An interesting feature of the signature is the positive relation between sediment yield and drainage area for areas 102 km2 < A.t < 3xl04 km2• In this range, g oc Al6, approximately, g being yield per unit area, so G oc A/6•

Beyond about 3xl04 km2, the exponent becomes negative and continues the upper envelope of response in presently glacierised basins.

104r-------~--------~------_,--------_r--------r_------_,

•

------~--~--~~

o~

'b • o

o o

• glacial

Stlklne, Liard, Peace (north) 0 JI Columbia, Coast, Fraser (south). "

DRAINAGE AREA (km 2 )

Figure 6. Contemporary fluvial suspended sediment yield, represented as specific sediment yield (GIN, in British Columbia, Canada, based mainly on data of the Water Survey of Canada. Stations with recently initiated land disturbance, and ones with significant lakes have been deleted. Data are divided into two groups, the northern group (Peace, Liard, Stikine drainages) having generally higher yields because of the presence of more erodible rocks. "Glacial" indicates contemporary glaciers in the drainage basin. (Adapted from Church et 01., 1989: Figure 7).

What do these results mean? An increase of g with Ad indicates that specific sediment yield (yield per unit area) increases downstream. This signifies either that downstream tributaries contribute systematically greater sediment yields than do headwater ones, or that sediment is being entrained from along the channels themselves. In the wildland landscape of British Columbia, the former possibility seems rather unlikely, as a generality. We draw the conclusion that headward drainage systems -- out to about 175 km flow distance from the headwaters -- are degrading. (This may not be an appropriate inference everywhere, depending upon regional patterns of land use.) Larger drainage basins exhibit a systematic reduction in specific sediment yield with distance downstream, indicating aggradation. This is the usually observed signature of sedimentary disturbance

102 Michael Church

in agriculturally disturbed landscapes. The actual, more complex condition illustrates systematic changes in the storage term in reach-delimited sediment mass balance and the direct connection of degradation (upstream) with aggradation (downstream) in the landscape as the outcome of contemporary sediment transfers. It also reflects the propagation of the glacial sedimentary disturbance through the landscape. Headwaters have been exhausted of mobilisable sediment, while more distal sites are still passing material onward through the cascade of reservoirs associated with the fluvial system. (In fact, drainage basins with A.! < 102 Ion2 do not, in general, conform with the pattern since contemporary sedimentary disturbances may easily overwhelm the residual glacial signal there.) For drainage basins of intermediate size, degradation is widespread, but not universal in glaciated Canada (cf. Church et al. 1999).

An interesting feature of the regional example given in Figure 6 is that the observed sedimentary signature lies entirely within the region of the former primary disturbance. The functional space within which the disturbance is propagating is not the compact geographical space of the region, but a near-linear topological space distributed along the stream system. The direction and geographical pattern of propagation of the disturbance are determined by topography, which controls the structure of the drainage network. Within the example, regional topography remains essentially constant (even though the complete evolution of the signal evidently will require much more than 10 lea), but fluvial landforms -- representing transiently stored sediments -- will develop and disappear. These are general features of fluvial systems that defme a domain of coherent behaviour of fluvial sediment yield in the landscape. The domain is space-delimited, but not in the usual manner of outer boundaries of an otherwise undifferentiated geographic space (such as, for example, drainage divides). To encompass the entire landscape it would be necessary to make mass wasting processes on hillside slopes an integral part of the analysis, something that is not pursued in this paper.

3. The signature of glacial disturbance

Glacial sediment yield is relatively extreme. Glaciers themselves are effective abraders of subjacent rock when basal slip occurs, as the result of both thaw-freeze mechanisms and the effect of entrained clasts grinding across the surface. Soil materials may be frozen onto the glacier sole or incorporated into the ice body along failure planes in compressive flow, so they are also an effective sediment transporting agent. At the ice margin, material that melts out from the ice, or is exposed from beneath ice, is immediately susceptible to transport by wind and water in an environment where both are seasonally strongly effective. The freeze-thaw environment and lack of vegetative cover magnify the effectiveness of these agents. As the result, specific rates of sediment yield are high (Figure 7a). Observations vary by as much as two orders of magnitude as the result of variable rock erodibility and variable glacier energy. "Maritime glaciation" on Neogene rocks in southeastern Alaska, as Figured in the diagram, must yield some of the highest glacial sediment yield rates on Earth.


As one moves away from the glacier margin, specific yields are observed to decline (Figure 7b). One may suppose that this is the consequence of reduced regional yield where an increasing proportion of non-glacierised terrain -- with presumably much reduced unit area yield -- is included in the regional average result. This simple picture hides a more complex reality. As one moves downstream in a fluvial system, areal sediment yield may appear to change because material is moving into or out of storage in reservoirs along the stream system. Scale-related change in specific yield may indicate the occurrence of restricted sources, as described here, or it may indicate the change of sediment storage, as described in the last section. Persistently declining yields indicate either a concentration of sediment sources at the headwater of the system, or aggradation downstream. In the case of a glacial system, both effects are apt to occur, as we see by the common juxtaposition of ice margins and outwash. The data of Figure 7b also suggest that response magnitude, or yield magnification, due to glaciation is in the range 101.5 to 1 Q2x. Such other data as are available (Table 1) suggest that this is a relatively robust estimate of the range for major surface disturbances.

(a) ~.; ..........

of' 10· . -.............

~ e

" • • ' .. ! • e, -. ' ..

• ~ 10·

0 • 0 0 ¥ 10' 8

~ .0 (b) 0

.. 0 ! .~ i .,. • • • " ~ 10' • 10' 0 ,. ~o 0 00. ..

I '~i-0 _.0 0

0

• e 0 • o • • .,.- .... .0: 0 • e. 0 e

10" .. NoTJW8Y 0 • • _.,'alld ." e AI_ka o 0",., I ",' •

10' 10" 10' 10' 10' 10· 10" 10' 10"

AREA (lanz) '" GLACIAL COVER

Figure 7. Measured glacial sediment yields, represented as specific yield rate: (a) Data compiled in Hallet et al. (1996). The three distinguished regions exhibit distinct full-glacial sediment yield rates because of different bedrock types; (b) variation of fluvial suspended sediment yield with percent contemporary glacierization, from compilations summarised in Hattet et 01. (1996).

104 Michael Church

Table 1. Data of fluvial sediment response to disturbance

SitelEvtnt A,,' t' t..' lateaslty lIlapifk:atl loa' tonDel kDI.J.-1 .. '

Contemporary G1aciar

Nigardsbreen, Norway 48 400 22x

Engabreen, Norway 38 1050 66

Norway (average) 1000 70

Iceland (summary) 50-100

Alaska (summary)' 1000 100

Contemporary Non-Glacial

Washington fires 5 2 8? 120+ 100

Yellowstone fires8 256-570 50 >1000

Western Run, Md.9 ISS 140

Coon Cr., WisconsinlO 360 80 2800

Lloyd Shoalsll 3600 110 220 >7

Yuba R, Californial2 3500 42

Sacramento R, Californial2 55000 35 10

Bear R, California 13 756 28 >100

Pleistocene Glacial

British Columbia 104 104 _102 102

Notes and references: I.At measurement point for sediment transport, or at area of peak response (B.C.); 2. response time. 3. relaxation time. 4. sediment yield at peak response. 5. Intensity/weathering rate (the denominator generally being estimated). 6. Mainly after Hallet el al. (1996) 7. From Guymon (1974) 8. Meyer el al. (J 992) 9. Costa (1975), not precisely determined. 10. Trimble (1983) II. Altamaha R., southeastern U.S.A. (Meade and Trimble, 1974) 12. Gilbert, 1917; Graves and Eliab, 1977; James, 1989, 1991.

Fluvial Sediment transfer in Cold Regions 105

Away from an ice margin, or immediately following deglacierisation, the entire fluvial landscape is apt to be aggrading, reflecting the remobilisation of glacial sediments by running water and its deposition along stream courses (Figure 8a). Hillslopes and glacial deposits experience degradation. After the lapse of some time, mobilisable sediment is exhausted in these places (in part, usually, because of stabilisation of part of the glacial mantle under developed tundra, forest or prairie vegetation). Fonnerly aggraded outwash deposits along the valleys begin to be degraded by sediment-starved runoffwaters and resedimented farther down the system. This gives rise to a characteristic signature of coupled degradation-aggradation (Figures 8b,c), which looks very like the contemporary pattern in British Columbia and like the pattern of material staged through a sequence of reservoirs (Figure 1). Over a very long time, the spatial signature of fluvial sediment transfer following a major glacial disturbance appears to evolve in approximately the manner shown in Figure 9.

The signature of disturbance in time is similar. The signature of "paraglacial sediment yield" (Church and Ryder 1972; Church and Slaymaker 1989) is already depicted as such (Figure 10: inset). For each position in the drainage system, there is a distinctive signature in time. Figure 5, which represents the space-time base for full representation of sedimentary disturbance, shows that propagation in time and space (as area) are very nearly similar, so far as the data allow us to judge. In fact, the upper envelope of data in Figure 5 is not inconsistent with A.J oc t for sediment response time, implying that L oc f·s for sediment wave propagation through the fluvial system. Figure 10 shows the conjectured pattern of sediment yield in both distance and time following a disturbance. Distance is measured from the centre of the disturbance, time from the inception of significant impact on the fluvial system. For a major disturbance, such as the Pleistocene Northern Hemisphere glaciations, the time scale is very long, which reduces the significance of the practical difficulty to determine when to should be declared. Much of the distance in Figure 10 occurs within the boundaries of the disturbance, but it is evident that the model applies, as well, to extraglacial areas influenced by glacial runoff or by coeval severe climates (e.g., Ashley and Hamilton 1993). Figure 10 represents the full regional signature of paraglacial fluvial sedimentary response, which is repeated for each major glacial episode (Clague 1986).

4. Complications of the signature of glacial disturbance

The simple conceptual picture described above is complicated in nature by the fact that sediments of different sizes may propagate at different rates through the fluvial system, so that it is actually a family of waves that moves through the landscape. Figure 11 a gives a conceptual view of this circumstance, but there are no empirical data to test it, beyond the casual observation that certain data in Figure 5 that deviate from the main trend (Coon

106 Michael Church

Creek; Lloyd Shoals of upper Altamaha R) are those derived from fme-grain sedimentary systems, and the deviance is in the direction of a more sensitive relation in time. In these systems, sediments, once eroded from the land surface, may be reentrained more frequently, so the system behaviour is displaced toward virtual velocity scaling. This picture is also consistent with the common observation that coarser materials remain instorage for lengthy periods close to drainage headwaters, or to the source of a sedimentary disturbance.


Figure 8. Stages in glaciofluvial sedimentary response: (a) ice-marginal outwash aggradation in a British Columbia mountain valley at ca. 13 000 BP (a.k.a. Tasman Glacier outwash, Mt.Cook, New Zealand, 1972AD); (b) British Columbia mountain valley at the inception of postglacial degradation, ca. 9 000 BP (actually, Iskut River, British Columbia, ca.1982AD. Iskut River drains contemporary alpine glaciers.); (c) contemporary British Columbia mountain valley (Fraser Valley near LiJlooet, British Columbia, really).

A significant further complication that is reflected in the longevity of the fluvial landforms derives from the interaction of sediment sorting and changing stream competence as hydrology was modified in the postglacial period. Sediment sorting derives from competence limitations of streamflow. Competence is determined by stream gradient and flow magnitude. Through a drainage basin, these change systematically (Figure 11 b) in such a way that the ability of the stream system to move sediments onward is steadily reduced to smaller sizes. The large influx of sediment early in the disturbance history creates substantial aggradation near the source of disturbance (for example, within the region of glacial disturbance). Later on, after sediment influx to the fluvial system has declined, some of the initially deposited sediment is reentrained and moved farther on. What sediments are reentrained and moved on depends upon two factors: (i) continuing competence of the flows, which may decline as late-glacial, meltwater-augmented runoff gives way to hydro-meteorologically determined runoff in the postglacial period, and (ii) the propensity of the earlier deposits to become progressively armoured at the surface by the largest and least mobile materials in the deposit as reentrainment of fmer material occurs (this phenomenon provides the basis for James's reinterpretation of the California hydraulic mining data). The result may be substantial degradation and fluvial terrace production in headwater reaches, where the high topographic slope maintains stream competence to evacuate sediments once the initial aggradation, prompted by the high

108 Michael Church

volume of the late-glacial sediment influx, has ended. Similarly, in regions far downstream and reached only by sands, degradation may eventually set in -- but only after a relatively long lapse of time -- as rivers remain competent to entrain these fme materials. In intermediate reaches, however, hydrologically imposed reductions in stream competence and surface armouring of gravel deposits may combine to limit the degradation that may occur, especially along smaller river systems. Terrace production may be limited in this zone (Figure llc). Furthermore, this zone may remain, for a long time, the source of slowly released sediment that moves on into the distal reaches oflarger river systems, so that floodplains there continue to aggrade at rates in excess of any that would be supported by contemporary weathering.

)

subaerial weatheringJt )

(log) area o --Figure 9. Evolution of the signature of glacial disturbance over time. The

contemporary stage of postglacial sediment transfer in British Columbia (cf. Figure 6) is emphasised, and the stages supposed to be represented in Figure 8 are indicated.


The eventual breaching of channel armour as the result of major floods or hydrological regime change may also set off renewed epicycles of upstream degradation and downstream aggradation, so that the eventual response of the river system to sedimentary disturbance may become complex in Schumm's (1973) sense.

The features sketched in the foregoing paragraphs and illustrated in Figure 11 are evident in the postglacial fluvial landscape of British Columbia, indicating that, for glacial sedimentary disturbance, the time-scale for subsequent development of these effects can be of order 104 a. At the far distal end of the largest river system contained wholly within the province, sedimentary evidence indicates that Fraser delta -- a significant sand-body delta -- has been developing at rates similar to the present-day rate throughout the Holocene Epoch (Mathews and Shepard 1962). At this distance from the fluvial sediment sources (between 102 and 1 ()3 km), temporal variability of postglacial sediment yield is almost entirely damped.

~ 3.5 ... ..!:l! 3

~ 25

05

o 50

. ' - ...:_!'- ---- -~ -.... _ ............ -_ ....

: prlrrtGrf .$!bOWQt dmiJdctJ0r2 fr;)tr

tnJr(grurph~.kfltJ(1J 10000,.,,)

50

Figure 10. Paraglacial fluvial sedimentary response in space and time. For illustration, it is assumed that peak response follows A.J ~ t. Inset: paraglacial fluvial response as Figured by Church and Slaymaker (1989; their Figure 2).

110 Michael Church

Another significant complication of the signature of glacial disturbance is that disturbances of the landscape are never isolated in time or space. Disturbances of different kind and intensity occur successively at the same place in the landscape. These might be serial recurrences of a similar kind, as illustrated in Figure 3b, or they may be disturbances of different kinds, as illustrated qualitatively in Figure 12. When disturbances are not colocated, it is reasonable to suppose that their propagated effect on the downstream fluvial system, if they converge, is additive. Where they are co-located, it is possible that they will be non-additive, because the effect of the frrst disturbance on land surface condition may influence the magnitude of a following disturbance. Since the effect of late Pleistocene glaciation was widespread, it is reasonable to suppose that the progress of that sedimentary disturbance has influenced more localised disturbances in the Holocene landscape. The influence may be to magnify the effect of later disturbance, when substantial volumes of glacial sediment remain to be mobilised once surface cover is destroyed, or it may be to dampen the effect, when the earlier disturbance has already removed the readily mobilised material. This sort of interaction is present at some sites where ostensibly long-delayed 'paraglacial' effects have been identifed (e.g., Forbes and Taylor 1987). Figure 13 illustrates the palimpsestic nature of surface disturbance and a generalisation of the regional signature of sedimentary disturbance to which it gives rise.

Figure lOin fact represents the regional signature of any sedimentary disturbance. Specification of the space and time scales depends upon the magnitude of the disturbance. Figure 13b is simply the sum of many such disturbances with varying magnitudes. Today, the overwhelmingly dominant sedimentary disturbance in most of the world is the effect of human activity. Agricultural activity, in particular, gives rise to continuous and severe land surface disturbance, so that fluvial sediment transfers are elevated well above those that would occur in natural landscapes. The characteristic signature of human disturbance is similar to that offutl glacial response: declining specific yields at all downstream distances from source (cf. Walling 1983). This indicates pervasive aggradation in stream systems and is the consequence of the human-induced erosion of the land surface. In comparison, the residual signature of glacial disturbance is trivial and goes undetected. Indeed, Canada may offer the only glaciated landscape within which the residual Pleistocene glacial sedimentary signature remains clearly identifiable. Even here, other events have quite overtaken residual paraglacial effects in many basins of intermediate size in southern Canada (cf. Jordan and Slaymaker 1991; Slaymaker 1993; Brooks 1994).

Neoglacial sedimentary signatures are local. The fluvial signature of Neoglacial disturbance has been detected outside the immediate area of alpine glaciation (e.g., Gottesfeld and Gottesfeld 1990). More generally, a fluvial response to episodic Holocene deteriorations of climate has been detected in a number of studies (e.g., Knox 1985; Chatters and Hoover 1988; Schirmer 1988; Macklin and Lewin 1993; ). These effects are interesting because they reflect regional hydrological change, as well. But the most recent


Neoglacial effects (17th to 19th centuries) are often difficult to separate from the impacts of societies that were developing rapidly, both agriculturally and industrially, in the same period . The effects of earlier Holocene environmental fluctuations are preserved in the the fluvial record in favourable circumstances (e.g., Costa 1978; Baker et al. 1983), but specifically glacial signatures often have been destroyed or rendered uninterpretable by later epicycles of fluvial activity.

5. Discussion and Conclusions

In this paper, I have argued that there is a coherent, recognisable response to sediment input to the fluvial system. Furthermore, the characteristic time scale of disturbancerelated sediment inputs is short in comparison with the response time away from the source, at least for sediments that are apt to be deposited along the stream system. With respect to these sediments, the fluvial system can be likened to a sequence of reservoirs. Sediment storage and onward movement through the fluvial system is at least qualitatively like the movement of material through a sequence of reservoirs.

The most significant sedimentary disturbance in geologically recent time has been that associated with the major, unstable Northern Hemisphere glaciations of the Pleistocene Epoch. The sedimentary disturbance associated with Pleistocene glaciation is still present and detectable in landscapes that have not more recently been extensively disturbed by human activities. In the present day, degradation is occurring of proximal glacial deposits, including outwash, and aggradation of more distal floodplain environments. From an example in British Columbia, Canada, a mountainous, largely unsettled environment, it appears that the length scale of degradation is now about 175 km from headward fluvial sources. The time since inception of glacial disturbance of the fluvial system (inside the limits of the Cordilleran ice) is about 12 000 years. This datum, in comparison with data of more recent disturbances, suggests that L oc fl·s approximately, in which t is response time, and Ad ex tl.O approximately, in which Ad denotes the area having experienced peak response by time t. However, the available data are consistently biased away from a simple linear relation for area. At the "fast response" extreme, L oc tl.O

and Ad oc tl·o, consistent with the virtual velocity of sediments which consistently bypass significant reservoirs. Sediment size-specific response analyses would yield a variable pattern of results between these extremes, according to the relative mobility of the sediments.

What is more striking than these appearances, however, is how little we know about the pattern of regional sedimentary response to land disturbance. Response occurs on a wide range of scales, varying from the effect of bank collapse on a river during an individual storm, or the influx of sediment from the land surface at the same scale, through responses to varying degrees of land disturbance, including mining, urban land

112 Michael Church

conversion, agriculture, forestry, and wildftre, to extended regional responses created by hydroclimatic change, tectonics, and glaciation. It appears as if the magnitude of sedimentary disturbance, in comparison with rates of subaerial weathering, characeristically varies from 2x to perhaps 102x, with glaciation representing a relatively extreme response. But there are too few data of undisturbed weathering rates to be very certain of this.

More generally, attention to the question of sedimentary disturbance and response makes us realise that we know virtually nothing about the following basic questions concerning sediment movement in the landscape: • What is the relative response time after a mobilising disturbance for sediments of

varying calibre, from cobbles to silts? • What are characteristic storage times for sediments in particular fluvial landforms?

(Alternately, how does one establish k, the reservoir characteristic specifted in Figure 2?)

• What roles do selective fluvial transport and abrasion play in determining these response times? ( "Abrasion" associated with the weathering of fluvial sediments in storage along the channel orin floodplains may be important even though direct abrasion associated with transport may not; clJones and Humphrey 1997.)

• Are the relations given above for L, Ad and t general, or do they depend speciftcally upon climate and lithology? (As usual, the data derive almost entirely from the northern temperate to subarctic regions.)

• Is it generally possible to separate the effects of distinct sedimentary disturbances in the fluvial system, thence to reconstruct the actual erosional history of the landscape?

It may be possible to tackle some of these questions with recently developed methods, including techniques for sediment tracing and new luminescent and radiogenic methods for absolute dating of mineral sediments. Regardless, we frrst need a more adequate conceptual understanding of sediment transfer in the environment. It has been the purpose of this paper to contribute toward this understanding

Acknowledgments

Paul Villard constructed Figures 10 and 13(b), while Paul Jance drafted all the other diagrams. I am grateful to them both for their technical skill. Olav Slaymaker found time to review and comment upon the paper. I appreciate his interest: one requires some reinforcement to put forward ideas as unsupported as these. I dedicate this paper to Stan Schumm, whose own conceptual flights of some twenty-five years ago form the foundation for some of the ideas presented here.


(c)

degradatlofk-l teTrace production

area/distance

area/distance

_ _ aggradation -- ..... ------. area/distance

113

Figure 11. Complications of the fluvial sedimentary response to disturbance: (a) the sediment wave at a ftxed time for various sedimentary grain sizes; (b) controls of sediment mobility downstream in fluvial systems; (c) conjectural distribution of fluvial sediments after disturbance, showing zones of degradation, delayed degradation, and persistent aggradation of the fluvial system.

114

CD

-= CD

U

0 -U)

CD

CL

U)

Michael Church

volcanism volcanism

I · f' human eog aCla lonimpacf

time Figure 12. Conceptual model of successive major sedimentary disturbances in a Holocene landscape, based on Lillooet Valley, British Columbia (Jordan and Slaymaker, 1991; their Figure 8). Lillooet River drains 3150 km2 and still retains 500 km2 (16%) glacier cover.


(b)

~ 3.5

'" ! 2.5

-!

1 1.5

0.5

0 50

50

Figure 13. Cartoons to illustrate the recurrent nature of sedimentary disturbance in the land-scape: (a) spatial pattern of disturbances. The varying circle size denotes varying magnitude, and the arrows (shown for only a few disturbances) indicate the topographically determined direction of propagation. Full circle and darkly shaded arrow represents the regional glacial disturbance; light shading represents wildfire effects; open circles represent human disturbance consequent upon land usc. Not all disturbances are co-located, and they do not all propagate into each other; (b) a generalisation of Figure 10 to show the pattern of regional sediment yield that would result from (a).

116 Michael Church

References

Ashley, G.M. and Hamilton, T.D. (1993) Fluvial response to late Quaternary climatic fluctuations, central Kobuk Valley, ncrthwestern Alaska, Journal of Sedimentary Petrology 63: 814-827.

Baker, V.R., Kochel, R.C., Patton, P.C. and Pickup, G. (1983) Palaeohydrologic analysis of Holocene flood slack-water sediments, International Association ofSedimentologists Special Publication 6: 229-239.

Brooks, G.R. (1994) The fluvial reworking of late Pleistocene drift, Squamish River drainage basin, southwestern British Columbia, Geographie physique et Quaternaire 48: 51-68.

Chatters, J.C. and Hoover, K.A. (1988) Response of the Columbia River fluvial system to Holocene climatic change, Quaternary Research 37: 42-59.

Church, M. and Ryder, J.M. (1972) Paraglacial sedimentation: a consideration offluvial processes conditioned by glaciation, Geological Society of America Bulletin 83: 3072-3095.

Church, M. and Slaymaker, O. (1989) Disequilibrium of Holocene sediment yield in glaciated British Columbia, Nature 337: 452-454.

Church, M., Kellerhals, R. and Day, TJ. (1989) Regional clastic sediment yield in British Columbia, Canadian Journal of Earth Sciences 26: 31-45.

Church, M., Ham, D., Hassan, M.S. and Slaymaker, O. (1999) Fluvial clastic sediment yield in Canada: scaled analysis, Canadian Journal of Earth Sciences 36: 1267-1280.

Clague, J.J. (1986) The Quaternary stratigraphic record of British Columbia -- evidence for episodic sedimentation and erosion controlled by glaciation, Canadian Journal of Earth Sciences 23: 885-894.

Costa, J.E. (1975) Effects of agriculture on erosion and sedimentation in the Piedmont Province, Maryland, Geological Society of America Bulletin 86: 1281-1286.

Costa, J.E. (1978) Holocene stratigraphy in flood frequency analysis, Water Resources Research 14: 626-632. Forbes, D.F. and Taylor, R.B. (1987) Coarse-grained beach sedimentation under paraglacial conditions,

Canadian Atlantic coast, in D.M.Fitzgerald and P.S.Rosen (eds.), Glaciated Coasts, Academic Press, Orlando, Florida: 51-86.

Gilbert, G.K. (1917) Hydraulic mining debris in the Sierra Nevada, United States Geological Survey, Professional Paper IDS: 154pp.

Gottesfeld, A. and Johnson Gottesfeld, L.M. (1990. Floodplain dynamics of a wandering river, dendrochronology of the Morice River, British Columbia, Canada, Geomorphology 3: 159-179.

Graves, W.P. and Eliab, P.L. (1977) Sediment study: alternative delta water facilities -peripheral canal plan, California Department of Water Resources, Central District, Sacramento: 117pp, (Quoted in Meade and Yuzyk 1990, and in James 19891991).

Guymon, G.L. (1974) Regional sediment yield analysis of Alaska streams, American Society of Civil Engineers Proceedings: Journal of the Hydraulics Division 100: 41-51.

Hallet, B., Hunter, L. and Bogen, J. (1996. Rates of erosion and sediment evacuation by glaciers: a review of field data and their implications, Global and Planetary Change 12: 213-235.

Hassan, M.A., Church, M. and Ashworth, PJ. (1992) Virtual rate and mean distance of travel of individual clasts in gravel-bed channels, Earth Surface Processes and Landforms 17: 617-627.

James, L.A. (1989) Sustained storage and transport of hydraulic gold mining sediment in the Bear River, California, Association of American Geographers Annals 79: 570-592.

James, L.A. (1991) Incision and morphologic evolution of an alluvial channel recovering from hydraulic mining sediment, Geological Society of America Bulletin 103: 723-736.

Jones, L.S. and Humphrey, N.F. (1997) Weathering-controlled abrasion in a coarse-grained, meandering reach of the Rio Grande: implications for the rock record, Geological Society of America Bulletin 109: 1080-1088.

Jordan, R.P. and Slaymaker, O. (1991) Holocene sediment production in Lillooet River basin, British Columbia: a sediment budget approach, Geographie physique et Quaternaire 45: 45-57.

Knox, J.C. (1985) Responses of floods to Holocene climatic change in the Upper Mississippi Valley, Quaternary Research 23: 287-300.


Macklin, M.G. and Lewin, J. (1993) Holocene river alluviation in Britain, Zeitschriftfur Geomorphologie Supp. 88: 109-122.

Mathews, W.H. and Shepard, F.P. (1962) Sedimentation of Fraser River delta, British Columbia, American Association o/Petroleum Geologists Bulletin 46: 1416-1438.

Meade, R.H. and Trimble, S.W. (1974) Changes in sediment loads in rivers of the Atlantic drainage of the United States since (1900) International Association of Hydrological Science, Publication 113: 99-104.

Meade, R.H., Yuzyk, T.R. and Day, T.J. (1990) Movement and storage of sediment in rivers of the United States and Canada, in Wolman, M.G. and Riggs, H.C., editors, Surface water hydrology. The Geology 0/ North America, vol. 0-1) Geological Society of America, Boulder, CO: 255-280.

Meyer, G.A., Wells, S.G., Balling, Jr., R.C. and Jull, A.J.T. (1992) Response of alluvial systems to fire and climate change in Yellowstone National Park, Nature 357: 147-150.

Schirmer, W. (1988) Holocene valley development on the Upper Rhine and Main, in Lang, G. and Schluchter, C., editors, Lake. Mire and River Environments. Rotterdam, Balkema: 153-160.

Schumm, S.A. (1973) Geomorphic thresholds and complex response of drainage systems, in Morisawa, M.E., editor, Fluvial geomorphology. Proceedings volume of the 4th Annual Binghamton Geomorphology Symposium, State University of New York, Binghamton, NY. Publications in Geomorphology: 299-310.

Slaymaker, O. (1993) The sediment budget of the Lillooet River basin, British Columbia, Physical Geography 14: 304-320.

Tassone, B.L. (1987) Sediment loads from 1973 to 1984: 08HB048 Carnation Creek at the mouth, British Columbia, in Chamberlin, T.W., editor, Proceedings o/the Workshop: Applying 15 Years o/Carnation Creek Results, Carnation Creek Steering Committee, Pacific Biological Station, Nanaimo, British Columbia: 46-58.

Trimble, S. W. (1983) A sediment budget for Coon Creek basin in the Driftless Area, Wisconsin, 1853-1977, American Journal o/Science 283: 454-474.

Walling, D.E. (1983) The sediment delivery problem, Journal 0/ Hydrology 65: 209-237.

PART II

Cold Lowland and Coastal Environments

WHERE ON EARTH IS PERMAFROST? BOUNDARIES AND TRANSITIONS

Michael W. Smith and Dan W. Riseborough Department o/Geography Carleton University Ottawa, Canada

Abstract

Geomorphology has two interlocking paradigms; the fITst is a process paradigm where there is a hierarchy of knowledge through the physics of a medium, the mechanics of process, to the landscape form; the second is a temporal paradigm from the history of the landscape, to the present condition, and with a prediction capacity. These are central to concepts of landscape transitions and landform assemblages. The paper explores landscape transitions occurring in pro glacial and paraglacial environments based on the horizontal link of sediment transfer rather than on the traditional vertical division into glaciology, hydrology and periglacial geomorphology. In glacierized mountain areas the superimposition of transitions at different scales can readily be demonstrated within the regional landscape. Transitions may occur slowly, rapidly or instantaneously but their frequency of change will vary depending on the underlying cause of the transition. The concept of landform assemblages should integrate the ideas of process activity at different spatial, quantitative and temporal scales. Landform assemblages are therefore infmitely variable. The ways in which the concepts of transition and landform assemblages influence geomorphological interpretation is illustrated by reference to the glacierized environments of the southwest Yukon.

1. Introduction

Permafrost, a thermal condition of the ground, is, ultimately, a climatic phenomenon. An enduring question in permafrost science has been the precise nature of the relationship between mean air temperature and permafrost temperature. Apart from its intrinsic interest, this question has gained new importance with the increasing concern that permafrost environments will be particularly affected by global warming widely expected over the next century. The IPee (1990) suggested that research be directed toward an improved defming of the permafrost-climate system, including the effects of temperature forcing due to climatic variation, local environmental factors such as snow and vegetation, and lithologic conditions.

121

K. Hewitt et al. (eds.), Landscapes a/Transition, 121-139. rg 2002 Kluwer Academic Publishers.

122 Michael W. Smith and Dan W. Riseborough

This chapter analyses the climatic and environmental conditions that define the limits and continuity of permafrost occurrence in Canada. We believe the analysis could be extended to any geographic region. Historically, the extent of permafrost zones in Canada has been interpreted to correspond broadly to the continental-scale pattern of air temperature isotherms (French 1993). However, the actual relationships have not been clearly determined. Furthermore, while the southern limit of( discontinuous) permafrost has been the subject of considerable field study, discussion and speculation, there has been little analysis of the conditions favouring permafrost-free conditions at the northern extent of the discontinuous zone.

In this chapter, we present an explicit analysis of the question "What is the relation between mean annual air temperature (MAA T) and mean annual ground temperature (MAGT)"?

Following from this, we ask: What is the climatic sensitivity of permafrost? What is the maximum MAA T under which (discontinuous) permafrost can exist? Is permafrost at its extreme southern margin a contemporary feature or a relict associated with the Neoglacial period? What are the limiting conditions for continuous permafrost?

2. Permafrost Distribution in Canada

Permafrost is a significant feature of Canada's natural environment. According to the Permafrost Map of Canada (Figure 1), permafrost regions encompass about 50% of the nation's landmass. At the national scale, maps of permafrost in Canada show a broad latitudinal zonation of continuous, widespread (discontinuous) and scattered (sporadic) permafrost occurrence. The different zonal bands are intended to convey the relative areal dominance of permafrost and permafrost-free conditions in a region. Most of the published maps show a similar continental scale pattern (Nelson 1989). However, knowledge about permafrost conditions within any particular zone is limited by the scarcity of field measurements of ground thermal conditions. This is not surprising given the vast extent of the Canadian north, and the difficulty and expense of sampling the landscape at appropriate intervals.

In Figure 1, latitudinal bands of permafrost occurrence are seen to extend across the northern part of the continent, reaching furthest south near James Bay, with boundaries trending northwest towards the Alaska border, and east to northeast toward the Atlantic coast of QuebeC-Labrador. At this scale, permafrost distribution relates broadly to the isotherms of mean annual air temperature. The pattern of MAA T itself relates to the influence of the Western Cordillera on the upper airflow, and the subsequent persistence of arctic air masses for longer periods in eastern Canada (Rouse 1993). The discontinuity shown in the pattern of permafrost occurrence west and east of Hudson Bay seems to be unrelated to mean annual air temperature. This is discussed later in the paper.

\

Where on Earth is Permafrost?

= Continuous pennafrosl Widespread discontinuous permafrost

Sporadic discontinuous penna f rost

Figure 1. Pennafrost Map of Canada (Heginbottom et 01. 1995)

123

French (1993) states that the southern limit of continuous permafrost corresponds with a MAA T of -6 °e to -8 °e, with southern discontinuous limit at about -I °e. Brown (1970) suggested that permafrost is I to 5.5e wanner than air temperature, averaging about 3°e wanner. Assuming this, the southern limit of discontinuous permafrost would correspond with the -I °e mean annual air isotherm, and the southern limit of continuous permafrost with the -5.5°e isotherm. The southern boundary of discontinuous permafrost corresponds reasonably well with most temperature limits in the literature, while the value for the southern boundary of continuous permafrost is wanner than expected. Brown (1970) stated that the southern limit of permafrost "corresponds roughly with the 30°F (-1°C) mean annual air isotherm"; the continuous boundary is based on the 17°F (-8°C) isotherm, except that the band along the coast of Hudson Bay west of James Bay is anomalous.

2.1. PERMAFROST TRANSITIONS

The concept of permafrost zones tends to suggest distinct kinds of terrain. In reality the zones represent the transition from the seasonally frozen ground of temperate regions of the south to the extensive perennially frozen ground of the far north. As one moves progressively southward from the far north, where permafrost is present almost everywhere (continuous), the climate ameliorates and permafrost gradually becomes thinner and less


widespread (discontinuous). Eventually, permafrost exists only in scattered islands, such as in the peatlands of northern Alberta (Lindsay and Odynsky 1965, Zoltai 1971). But, ultimately, it disappears altogether.

While maps of permafrost distribution portray a broad zonal distribution, it is well

,-----

I ,:..""=/ • • • • • • • .. • • , . • :

/

o To "...,.,tura Ie)

Figure 2. Local Ground Temperature Distributions.

100%

9M!.

.80%

·709\

60"- ~ ! l

.509\ IL

• At 1he DiScoI1iroIous t PennafrOst Boundary 400/0 §

U

30%

·209\

·10%

O'M. 6

known that the actual relationship between permafrost and climate is not simply a function oflatitude. Local variations in vegetation, topography, snow cover and soil conditions can produce variations of several degrees in mean ground temperatures over a small area. Within a region of uniform climate, there will be a range of local ground temperatures, reflecting the variation of terrain, vegetation and lithologic conditions. In particular, these variations result in a patchwork of discontinuously frozen ground where mean annual air temperatures are within a few degrees of O°C. The zone of discontinuous permafrost comprises the band of transition between conditions where permafrost is ubiquitous (spatially continuous) to where it is completely absent. At the southern boundary of this zone, the coldest temperature anywhere in the given area is DoC, while at the northern boundary the warmest temperature is O°C. This is shown schematically in Figure 2. Each curve represents a local or regional distribution of mean annual ground. Within the discontinuous permafrost zone, the distribution will extend to both sides ofO°C.

The geographical extent of the discontinuous zone is the spatial consequence of the range oflocal ground thermal conditions that develops in response to regional climate. According to the permafrost map of Brown (1978), the width of the discontinuous zone ranges from about 500 km west of James Bay to about 1200 km in northern Quebec. Ives (1974) stated that the latitudinal gradient in MAAT is on the order of leo per 200 km.

Where on Earth is Pennafrost? 125

This gradient suggests typical local variations in ground temperature relationship ranging from 2.5°C where the discontinuous zone is narrowest, up to 6°C at the widest. This is in reasonable agreement with Brown's (1970, p. 20) observation that permafrost temperatures are -1 ° to 6°C warmer than mean annual air temperatures, giving a range in permafrost temperature of 5°C.

3. Permafrost and Climate

Field observations over the years have indicated a broad relationship between mean annual air temperature and pennafrost distribution. However, many discussions over the same period (Brown 1960, Williams and Smith 1989) have recognized that the relationship is not so simple.

According to Gold and Lachenbruch (1973), in northern latitudes the mean annual ground surface temperature (MAGST) is variably 1° to 6°C warmer than the corresponding MAA T, largely as a result of variations in snow cover. For example, temperature data for the Fort Simpson region (Judge 1973) show that while the mean annual air temperature is about _4°C there, the approximate range ofMAGST is between 0.7°C and l.3°C, largely as the result of snow cover. On this basis, one might infer that regions with MAA T greater than about -5°C would be pennafrost-free. This is not the case, however. In fact, pennafrost is believed to underlie about 30% of the Fort Simpson region (Wright 1995). This circumstance can be explained by the seasonal variation in ground thennal properties, which can lead to mean ground temperatures that are (considerably) lower than the mean ground surface temperature (Goodrich 1978).

Within a region of unifonn climate, we can visualize a range of local ground temperatures, reflecting variation in snow cover, vegetation and lithologic conditions. The effect of snow cover and the ground thennal properties, in particular, significantly modulate the relation between air temperature and pennafrost temperature. Whether pennafrost temperatures exists in a particular place or not, depends on the interplay of regional climatic conditions and various local factors (Smith and Riseborough 1983).

3.1. THE CLIMATE-PERMAFROST RELATIONSHIP

Much of the climate-pennafrost relationship is understood in a qualitative way, and there has been little success in defming it in an explicitly functional and exact manner. Models and analyses based upon surface energy exchange have been infonnative in understanding the detailed physics of the pennafrost-climate relationship. However, they are impractical beyond the site scale, because of the limited database for characterizing the microclimates of a broad range of vegetation and terrain conditions. A necessity for regional pennafrost analysis is the simplification of surface energy balance computations in order to reduce the number oflocal variables needed. Index methods have the virtue of using minimal climatic and soils infonnation to delineate the broad regional features of pennafrost distribution (Nelson 1986, Nelson and Outcalt 1987, Jorgensen and Kreig 1988), but they do not contain explicit relations between climate and the temperature conditions of pennafrost.


Here, we understand that the mean annual temperature of permafrost is related systematically to climate, surface conditions (principally snow cover and vegetation) and lithology. Following Lachenbruch et al. (1988), the permafrost-climate relationship can be represented by the temperature regime at a minimum of three levels: 1. the air temperature (as measured at standard screen height); 2. the temperature at the ground surface; 3. the temperature at the top of permafrost. The temperatures at each of these levels differ on a mean annual basis (Figure 3).

Mean Annual Temperature Profile

I Lapse Rate 1-

SUrface Bmtwlary Layer Offset MAAT -

0 --. MAGST S~/Wgdation ------

E T~y Offset .Aaive lAyer

:E -,-.!;!) TTOP G) J:

I

§. -2 ~ Geothennal Pemuljrost Gradient

·3

~+-~--~~~~--r--9 -8 -7 -6 -5 -4

Mean Amual Temperature

Figure 3. Schematic Model of the Climate-Pennafrost System.

The difference between the air and the ground surface is influenced most by the presence or absence of snow cover, which determines whether heat transfer immediately above the ground surface is predominantly conductive or convective. The primary influences of vegetation are a reduction of solar radiation reaching the ground surface in summer and the effects on the accumulation and persistence of snow cover (Luthin and Guymon 1974, Rouse 1984). Smith (1975) concluded that the direct effect of vegetation was less important than its role in snow accumulation. Nicholson and Granberg (1973) found the variation in mean annual ground temperatures (MAGT) in the Schefferville area was determined primarily by the snow depth, with variations in summertime conditions being less important. Since the effect due to winter snow cover is generally greater than that due to vegetation in summer, the mean annual ground surface temperature (MAGST) is warmer than the mean annual air temperature (MAA T) in most cases. Subsequently, we refer to this effect as the nival offset.

Where on Earth is Permafrost? 127

Between the ground surface and the top of permafrost, heat transfer by conduction varies seasonally between frozen and thawed states, since the thermal conductivity of ice is four times that of water. This produces a difference between MAGST and MAGT (Goodrich 1978), with the mean annual ground temperature becoming progressively colder through the active layer (Burn and Smith 1988, Romanovsky and Osterkamp 1995). Thus the lowest MAGT occurs at the permafrost table; we call this TTOP (mean temperature at the top of permafrost). As a result of this so-called thermal offset, permafrost can exist even where the MAGST is above O°C.

The nival offset and the thermal offset together defme the relationship between air temperature and permafrost temperature. The presence of snow cover results in a mean ground surface temperature (MAGST) that is higher than the mean annual air temperature (MAA T), while the thermal offset results in permafrost temperatures (TTOP) colder than the ground surface. The interplay between these two effects is responsible for the conditions of permafrost occurrence at any location.

3.2. TTOP MODEL

Smith and Riseborough (1996) presented a formulation of the permafrost-climate relation for analyzing the influence of climate, terrain and lithologic factors on the temperature condition and distribution of permafrost. The TTOP model links permafrost temperature conditions with the surface climatology through seasonal surface transfer functions and subsurface thermal properties. The model is exact for equilibrium conditions and provides a reasonably accurate estimate of subsurface temperatures under transient conditions. While part of the TTOP relationship was known previously (Kudriatsev 1958, cited in Romanovsky and Osterkamp 1995), it seems that its general utility was not recognised prior to the present work. The TTOP model can provide a general formulation of the climatepermafrost relationship, linking ground temperatures to the annual cycle of atmospheric temperatures, with the modulating effects of local surface and lithologic conditions.

The TTOP equation is (Smith and Riseborough 1996):

which can be expressed as:

ITOP = (laf kj)ntJt - nt·If P

ITOP = rk.ntJt - nt·If P

(1)

(2)


TrOP = Temperature at Top Of Permafrost (~ ODe) It = thawing index for air temperature If= freezing index for air temperature (absolute value) let = thermal conductivity of ground (thawed) kl= thermal conductivity of ground (frozen) rk = ktlkl(thermal conductivity ratio) nt = scaling factor between summer air and surface temperatures (vegetation effect) nf= scaling factor between winter air and surface temperatures (snow cover effect) P = Period (365 days)

From equation (2), we can deduce the following air temperature (climatic) controls on TrOP:

• TrOP is negative (permafrost condition) wherever (nf.lj) > (rk.nL/t) • TTOP decreases with an increase in Ifand the corresponding decrease in It. • Where If » It, TTOP is likely to be negative everywhere in an area (regardless

of nt, n/and rk). • When It = If(Le. MAA T = ODC), TTOP will be negative wherever nl > (nt.rk). • The sum of It and If increases with latitude, and with continentality. If nl >

(rk.nt), TrOP will be lower as the sum of It and Ifincreases. Thus, the latitudinal decrease of TTOP with decreasing MAA T is reinforced by the associated increase in the annual total of degree-days. This effect increases as the difference between nfand (rk.nt) gets larger.

From equation (2) we can also examine the localised controls on TTOP. Because nf, nt and rk are multipliers, their effect on TTOP increases as It and Ifincrease. We consider the possible range of variation in: • rk is on the order of3-4 times (Riseborough and Smith 1998) • nt is on the order of 1.5-times (Jorgensen and Krieg 1988) • nl is on the order of 5-times (Riseborough and Smith 1998) From this, it is apparent that lithologic conditions and snow cover comprise the primary localised influences on permafrost temperatures, while variations in nt (vegetation effect) appear to be a second order effect. (To simplify the analyses below, geographical variations in nt have been ignored.) Close to the southern limit of permafrost, It and If are sufficiently similar that small differences in the surface and subsurface parameters become critical in determining the conditions in which permafrost can exist (that is TTOP negative). For permafrost to exist when It > If, nlmust exceed (nt.rk).

Equation (2) applies to the temperature at the top of permafrost (base of the active layer) - i.e. for TTOP ~ ODC. Where there is no permafrost but only a seasonally frozen layer overlaying permanently unfrozen ground, then the equation becomes:

MAGT= ntJt-(nJJ.f}/rk p (3)


where MAGT is the mean annual temperature at the base of the seasonal layer (~ O°C).

3.3. OFFSET FORMULATIONS

From equation (2), we can derive the expressions for the nival and thermal offsets for permafrost conditions. In the formulations below, the following equalities are implicit:

MAAT =(It-If)IP MAGST = (It - nf.1j)1P

with nt being taken as unity.

3.3.1. Nivalofftet

MAGST-MAAT= .If(l-nj) P

(4)

This is defmed as the difference between MAGST and MAA T:For any snow cover condition (n/value), the nival offset increases with the value oflJ. When there is no snow, n/equals one and the nival offset is zero. The value ofn/otherwise depends on snow cover conditions. By assuming that the freezing season coincides exactly with the presence of a snow cover, Riseborough and Smith (1998) determined that the value of n/ varies systematically with mean snow depth (z) and MAA T, using results from a numerical ground thermal simulator (based on Goodrich 1982).

A trend surface for nf= f(MAAT, z) was created from this analysis (Figure 4), allowing calculation of n/ for climate stations with air temperature and snow depth or snowfall data (snow depth was be estimated from snowfall data using the procedure presented in Nelson and Outcalt 1987). Where the snow cover is less than 30 cm, n/ is little affected by MAA T. Above this, there is a sharp transition to a strong dependence on MAAT.

3.3.2. Thermalofftet This is defined as the difference between TIOP and MAGST:

ITOP - MAGST= It(rk-l) p (5)

While the magnitude of the thermal offset depends directly on It, it is also determined by the ratio of thermal conductivity in the thawed and frozen states. Since rk is less than or equal to one (see Riseborough and Smith 1998), the thermal offset is negative and the temperature at the top of permafrost (TTOP) will be colder than the MAGST. This brings explicit recognition of the importance of the ground materials in the climate-permafrost relationship. For example, TTOP is reduced by 3"C between rk=1 and rk=0.3. This is equivalent to a distance of 300 kIn or more in terms of an equivalent change in MAA T.


In the case of bedrock, ,k approaches unity, and there is no offset effect. For example, Brown (1973) found no permafrost in bedrock outcrops in the Thompson, Manitoba, area, where the MAA Twas -3.3°C (recall that MAGST will be a few degrees warmer than MAA T because of the nival offset). Similarly, at Yellowknife, NWT, with a MAA T of -5.5°C, permafrost does not occur in exposed Precambrian bedrock but it does occur at other sites (Brown 1973).

For mineral soils, ,k will be in the range 0.6 to 0.9, depending on the water content but being largely insensitive to mineralogy (see Riseborough and Smith 1998). Thus, in mineral soil the thermal offset will range from -10 to -40% of (ItIP). The greatest range in rk values will be found in organic soils and may be from about 0.3 or even less under saturated conditions to about 0.9 for very dry conditions.

1.D_------------------------

:·:r--\W~-··-...-·-======-_={f-=-0.7 I--~\-,<: ----.. ------'-o-DC

-0-+2 0.6 ·1-----\\cT-'~..-~.:>....;;:----------------j __ .5C

00 02 0.8 O.S 1.0 1.?

Snow Depth 1m)

Figure 4. NfFunction

4. Application ofTIOP model to permafrost limits and transitions

We can now visualize the temperature of permafrost, TTOP, as the result of the interplay between the air temperature, the nival offset and the thermal offset. The net effect will determine whether permafrost exists at a location or not. This will depend on the values of ,k and nf, as well as the values ofIt and If, all of which are geographically and temporally variable.

Figure 5 shows the trend in It and If with MAAT, based on all geographically relevant Canadian climate stations. As expected, If increases with latitude while It decreases. Consequently, the importance of rk and nlto TTOP changes with latitude. We note further that:


Since q changes more rapidly than It, the relative importance of rk and nl also changes with latitude. The annual total of q and It increases with the latitude (as the annual range of temperature increases northwards), and this will contribute to lowered values of TTOP, as discussed previously (depending on the values of nJ, rk and nt). We term this the annual range effect.

As one moves southward towards regions of higher MAAT and increasing values of It, ground temperatures will inevitably rise. However, we note two limitations from equations (2), (4) and (5): • The increasing magnitude of the thermal offset with It reduces the impact of the

warmer summers on TTOP, the amount depending on the value of rk. • Since q becomes smaller with the increase in MAA T, the effect of nlon TTOP

is correspondingly reduced. Thus the thermal conductivity ratio, via the thermal offset, is the critical local factor in determining the southernmost limit of (discontinuous) permafrost occurrence.

f

OOOO~nr--------------------------------------------------<

.--241.8'MAAT+ 2142 R'.'.M

!4000r---------------~~~~_r----------------------------~ Ii' B M

i3000~-------------------------A~~~--------------------~ I ,

-I

~~~-----v---~-------··-----------···-··---·--·--·-··---.=~~ .• ~-~

O~--__ ----__ --__ ----__ --__ ----~------------__ ----~--~ .16.0 .14.0 ·12.0 -10.0 -8.0 -6.0 -'1.0 -2.0 0.0 2.0 4.0 6.0

MAATIC)

Figure 5. Trend of It and If with Latitude.


As one moves northward, If gets increasingly large and ground temperatures decrease. Again, we note: • The effect of the increase in If on TrOP is limited by the corresponding increase

in the value of the nival offset, depending on the value ofnt • Meanwhile, as It gets correspondingly smaller, the thermal offset decreases in

absolute magnitude and,k becomes progressively less important to TTOP. Thus we conclude that snow cover, via the nival offset, is the critical local factor in determining the northern limit of discontinuous permafrost (i.e. southern limit of continuous permafrost).

Figure 6 shows the climate data in Figure 5 plotted in degree-day index space. While there is a clear relationship with air temperature, it is interesting to note that some stations depart from this between _5° and -15"C. These stations fall largely within two geographic regions: QuebeclLabrador and Baffm Island. The climate regime in the eastern Arctic differs from that in the west, being less continental. For a given value ofMAA T, the annual temperature regime exhibits a smaller range, as indicated by a smaller sum ofIt and If. Consequently, TrOP values will be higher for similar values ofMAAT, as discussed above. For example, Fort Reliance (109.17 DW) has the same MAAT, -6.8DC, as Schefferville (66.82 oW), but an annual degree-day total of 54 1 0 compared to 4300. In this case, with other factors being equal, TrOP is warmer in Schefferville by O.SOC or so. This effect of reduced annual range may help to explain to the reported displacement of permafrost zones poleward east of Hudson Bay.

2000~~------~~~~~~~~--~--------------------~ g .. ! 1500·r--------.~------~~~--~----~~--------~ ·r & 1000 .------/-----------.~-------__J~--

+

5OOr--7L-----------~------------~----~~~--~----__;

OL-____________ ,~ __________ ~~ __________ _L ________ ~

a 1000 2000 4000 6000 7000

Freezing Index (If)

Figure 6. It and /fin Index Space

Figure 7 shows the trends in thermal offset and nival offset with the latitudinal change in MAAT, using temperature and snowfall data for the same climate stations as in Figure 5. The thermal offset is plotted for two values of ,k. In both cases, the trend is linear, with the slope determined by the value of Tk. One notes that the thermal offset

Where on Earth is Permafrost? l33

becomes increasingly negative as It (and MAA T) increases - i.e. as one moves south. In contrast, the nival offset trend is curvilinear, although the data are more scattered. The offset initially increases northward with the increase in .lfbut then shows a decreasing trend with the latitudinal decline in snowfall.

• • • • • • ..... *. . . . • •

• • • + •• • R'I. 0 .93

y.·O.14x ·2.3e;

•• +. •• • •• +

-16.0 -14.0 -12.0 -10.0 -8.0 -6.0 -4.0 -2.0 0.0 2.0 4.0 6.0

MAAT(C)

Figure 7. Trend of It and .lfwith Latitude

Figure 8 shows the variation of the nival offset at national scale. The pattern of values is related to the geographical distribution of snowfall. One notes that the offset increases and remains steady throughout much of the region of discontinuous permafrost but decreases abruptly at higher latitude. A continental effect of reduced snowfall is also evident in lower values for the nival offset in the central Arctic. A further effect, not represented here, is the reduction in nival offset associated with the denser snow covers of higher latitudes.

The implication of Figures 7 and 8 is that there is a gradual transition in TTOP southwards towards the ultimate limit of permafrost occurrence, as the effect of rising MAA T is offset by the increasing value of the thermal offset and the decreasing nival offset. In contrast, there is a much more abrupt transition from discontinuous to continuous permafrost associated with geographical changes in snow cover conditions northwards.

4.1. SOUTHERNMOST EXTENT OF DISCONTINUOUS PERMAFROST

Brown (1970) wrote that peatlands and permafrost are very closely related, especially in the southern fringe of the permafrost region. He felt that the occurrence of permafrost in this fringe area is governed to a considerable extent by the thermal properties of the peat, particularly the seasonal changes in the thermal conductivity between the thawed and frozen states. He presented descriptive reasoning for why ground temperatures would be


colder under peat than in adjacent areas without peat (ibid p. 176). Here, we present a more formal reasoning.

500 500 Kilometers -==--o

Figure 8. Nival Offset Map for Canada.

Nival Offset 0-0.9 1 - 1.9 2 - 2 .9 3 - 3.9 4 - 4 .9 5 - 5.9 6 - 6.9 7-8 No Data

According to French (1993) the southern limit of discontinuous permafrost coincides generally with a MAAT of -1°C. The _1°C MAAT isotherm reaches to approximately 58"N in Alberta, and 54-55"N in Saskatchewan, based on climate normal data. Nonetheless, Lindsay and Odynsky (1965) reported occurrences of frozen ground (permafrost?) in peatlands in northwestern Alberta, as far south as 55"N or so. Zoltai (1971) reported occurrences of permafrost in northern Saskatchewan as far south as 54"N. He gave the name "localized permafrost zone" to such areas, and suggested that the southern boundary of the local permafrost zone corresponded to the O°C mean annual air isotherm.

Brown (1967) considered permafrost in peatlands to be a special case; his map includes special symbols for "patches of permafrost ... south of the permafrost limit." Brown (1979, Map # I 0) includes a zone which " ... relict permafrost patches formed under previous climate regime." Halsey el a{ (1995) reported that since the Little Ice Age, MAA T isotherms have shifted a greater distance north than has permafrost. Thus, where


mean annual temperatures increase, large amounts of permafrost will persist as relict permafrost. They imply that it should eventually disappear, under the present climate. Finally, Camill and Clark (1998) state that the existence of permafrost in regions beyond the broad limit of distribution may reflect cool conditions from the Little Ice Age (from A.D. 1400 to 1850), and suggest that permafrost can be out of equilibrium with regional climate for a century or more. They conclude that during the transition to a future warmer climate, local processes will likely buffer permafrost peatlands from increasing temperature.

From this, we conclude that the thermal conditions in marginal permafrost are not precisely understood and that the climatic relations of marginal permafrost have not been clearly defmed. In particular, there is some debate whether permafrost at the extreme margins is relict or contemporary. We recall that, even under equilibrium conditions, TIOP can be less than MAA T, perhaps by several degrees, because of the thermal offset effect.

On the Map of Permafrost in Canada (Figure 1) Fort Vermillion, Alberta (58° 23'N), is shown to be close to the southern boundary of permafrost. For the climatic conditions at Fort Vermillion, It is about 1740 degree-days. Thus the thermal offset in mineral soil will range from -0.5 to -2.0Co; in organic soils the thermal offset could range from about-O.5 to -3.5Co. The MAAT at Fort Vermillion is -1.4°C, and the MAGST is calculated as +2.7°C. Thus the coldest likely MAGT in mineral soil will be +0.7°C (2.7°-2.0°); in organic terrain, the coldest MAGT will be -0.8°C (2.7°-3S).

The ultimate limiting condition for permafrost occurrence is when TIOP (MAGT) = O°C. From equation (2) this is given by:

rk.lt=nfJf

(ignoring the effect of nt). The same limit is derived from equation (3). Clearly, the solution to this is not single-valued. However, for various climate and snow cover conditions (It, q; nj), one can calculate the value of rk necessary to sustain permafrost at a location. In the case of Fort Vermillion, Fort McMurray and Peace River, these values are shown in the following Table.

Table 1.

It If MAAT MAGST MAGT MAGTo rk m crit

Fort Vermillion 2040 2505 -1.4 2.7 +0.7 -0.8 0.43 Fort McMurray 2050 2200 -0.4 3.0 +1.1 -0.4 0.38 Peace River 2235 1970 +0.7 3.6 +1.5 -0.0 0.31

Since a reasonable minimum value for rk is ....().3, Peace River represents the southern latitudinal limit for permafrost in northern Alberta under the present climate (ignoring the factor of mountain permafrost). According to St-Onge (pers. comrn.), permafrost occurrence in bogs is well known by local inhabitants in the Swan Hills area just south of


Peace River (see also, Lindsay and Odynsky 1965). Similar analyses could be repeated for other areas, such as northern Saskatchewan, where corroborative data can be compiled.

As stated Halsey et 01. (1995) report that since the Little Ice Age, MAAT isotherms have shifted a greater distance north than has permafrost in northern Alberta. Thus, where mean annual temperatures have increased, permafrost has persisted as relict permafrost. This "disequilibrium degradation" response, as they term it, is indicated by the presence of relict permafrost where current (air?) temperatures are between 0.5 and -3 .soC. They imply that it should eventually disappear, under the present climate. We conclude, however, that the thermal offset effect accounts for the occurrence of permafrost in peatlands at these latitudes and that permafrost can exist under the present climate. It is possible that with climate warming, summers have become drier resulting in a decrease in the thermal conductivity ratio. When this effect is combined with higher air temperatures, the magnitude of thermal forcing can be significantly reduced (see Smith and Riseborough 1996). The observations of Halsey et 01. (1995) can be explained by the buffering capacity of peat soils, which allows permafrostto persist under warming conditions, especially when enhanced by drier summer conditions. This implies that where organic material is present, the ground thermal response to climatic warming may be greatly reduced and permafrost may be quite persistent.

4.2. NORTHERN LIMIT OF DISCONTINUOUS PERMAFROST

4.2.1. Formulation o/n/crll

At the northern boundary of discontinuous permafrost, the warmest value ofTTOP in the area must be OQC. This will occur at sites where the influence of It is maximized - i.e. where rk is equal to one (a value of rk=l is found in bedrock locations). Therefore, the ultimate limiting condition for discontinuous permafrost is, from equation (2):

OQC = (It - nf.lj)IP

For any combination of It and If, we can derme the critical (minimum) value of nf (snow cover) sufficient to maintain unfrozen ground at the location:

nfcrit = Itlq

In other words, the northernmost occurrence of unfrozen ground will be found in bedrock (or very dry soil) wherever snow cover conditions exceed the equivalent value of nf crit.

For decreasing values of MAA T, paired values of It and qwere taken from the best-fit lines in Figure 5. Inserting these values in the equation above allowed calculation of the corresponding value of nfcrit. Using the relationship shown in Figure 4, these n/ values were converted into values of equivalent snow depth. The results are plotted in Figure 9 for both bedrock and mineral soil.

The figure shows the minimum snow depth necessary to preserve unfrozen ground as MAA T declines. In both cases (soil and bedrock), there is an abrupt rise in the critical snow depth to unrealistic values between -6 and -SoC. Beyond this point, permafrost becomes continuous. As expected, because of the thermal offset effect, the temperature


limit for mineral soil is reached at a higher MAAT (i.e. further south) than for bedrock. Finally, it is interesting to note that the continuous/discontinuous transition occurs at MAA T's between -6 and _8° C. This corresponds to, and may properly explain, the air temperature limit for continuous permafrost reported by French (1993).

Ongoing explorations with this approach indicate that the thermal properties of the substrate (soil latent heat and thermal conductivity) influence the relationship of nlto MAA T and snow cover. It appears that sites with a low thermal conductivity (peat) or low latent heat (bedrock) will have a relatively low nl compared to a site with a moderate conductivity and latent heat content (mineral soil).

Critical Snow Dep1h (no p81'1111frolt II actual Inow >critlcll)

1. 2$ ~---.:::-r, ---r--------,---- - -------, .....

"-

"

I i

1.00 '

0.761---.....j---lt-,----+--

_J~ . • Q & 0.50 - -

~

025 ---i

0.00 ·10 ·8

I, " ..... .....t..

,-~ · 4

MMTIc)

Figure 9, Snow Depth Necessary to Prevent Pennafrost

5. Conclusions

--~,----

-..... , ·2 o 2

Application of the TrOP model has elucidated the environmental conditions that limit the occurrence of permafrost at both the southern and northern margins of the discontinuous zone. While permafrost is ultimately a climatic phenomenon, the ground thermal conductivity ratio, via the thermal offset, is the critical factor in determining the southernmost limit of (discontinuous) permafrost occurrence. In contrast, we conclude that snow cover, via the nival offset, is the critical factor in determining the northern limit of discontinuous permafrost (i.e. southern limit of continuous permafrost).

Furthermore, we fmd there is a gradual decrease in MAGT southwards towards the limit of permafrost occurrence, as the effect of rising MAA T is offset by the increasing value of the thermal offset and the decreasing nival offset. This results in a diffuse geographical transition in the southward occurrence of permafrost. While local conditions


may permit the scattered occurrence of permafrost, it becomes gradually less common towards the southernmost margin. Eventually, permafrost exists only in scattered islands, such as in the peatlands of north em Alberta. But, ultimately, it disappears altogether. In contrast, northwards there is a much more abrupt transition from discontinuous to continuous permafrost associated with geographical changes in snow cover conditions northwards and the combined effect of air temperature and snow depth on the niva! offset.

Since the TTOP model uses only a few variables and requires limited types of data, it is particularly suited for application over a range of scales. It offers a rational method for regional and national scale mapping of permafrost distribution under current and possible future climate conditions. Using current studies of projected climate warming, it is also possible to estimate the future extent of permafrost.

References

Brown, R.J.E. (1960) The distribution of permafrost and its relation to air temperature in Canada and the USSR, Arctic, 13:163·177.

Brown, R.J.E. (1967) Permafrost in Canada, Geological Survey of Canada Map 1246a and National Research council of Canada, Division of Building Research Map NRC·9769.

Brown, R.lE. (1970) Permafrost in Canada, University of Toronto Press, Toronto, 234 p. Brown, R.J.E. (1973) Influence of climatic and terrain factors on ground temperatures at three locations in the

permaftostregion of Canada, Secondlnternationai Conference onPel71l4frost, Yakutsk, USSR. North American Contribution: 27·34.

Brown, R.J.E. (1978) Permafrost map 0/ Canada, Plate 32 in Hydrological Atlas of Canada, Ottawa: Department of Fisheries and Environment

Brown, R. lE. (1979) Permafrost distribution in the Southern Part of the discontinuous permafrost zone in Quebec and Labrador, Geographie Physique et Quaternaire 33: 279·289.

Bum, C.R. and C.A.S.Smith (1988) Observations of the 'thermal offset' in near·surface mean annual ground temperatures at several sites near Mayo, Yukon Territory, Canada, Arctic 41(2): 99·104.

Camill, P and J.S.Clarke. (1998). Climate change disequilibrium of boreal permafrost peatlands cause by local processes. The American Naturalist, 151: 207·222.

French, H.M. (1993) Cold climate processes and landforms, in Canada's ColdEnvironments, H. M. French and o Slaymaker (eds.), McOiII-Queen's University Press. p. 143·167.

Gold, L.W. and A.H. Lachenbruch. (1973) Thermal conditions in permafrost: a review of North American literature. Proceedings Second international Conference on Permqfrost, Yakutsk, USSR,3·23.

Goodrich, L.E. (1978) Some results of a numerical study of ground thermal regimes, Third Internternatlonal Canference on Pel71l4frost, Edmonton. Canada: 30·34.

Goodrich, L.E. (1982) The influence of snow cover on the ground thermal regime. Canadian Geotechtechnical Journal 19: 421-432.

Halsey, L.A., Vitt, D.H. and Zoltai, S.C. (1995) Disequilibrium Response of Permafrost in Boreal Continental Western Canada to Climate Change, Climatic Change 30: 57·73.

Heginbottom, J.A., Dubreuil, M.A. and Harker, P.A. (1995) Canada· Permafrost, in National Atlas o/Canada, 5th Edition, National Atlas Information Service, Natural Resources Canada, MeR 4177.

IPCC (1990) Seasonal snow cover, ice and permafrost, Potential impacts o/climate change, Report of Working Group 2: 7.1·7.45.

Ives, Jack D. (1974) Permafrost, in Arctic and Alpine Environments, J.D. Ives and R.G. Barry (cds.), 159·94, London: Methuen.

Jorgenson, M.T. and R.A.Kreig. (1988) A model for mapping permafrost distribution based on landscape component maps and climatic variables, Fifth Internatianal Co1lference on Pel71l4frost, Trondheim, Norway, 1: 176·182.

Judge, A. (1973) The prediction ofpermaftost thicknesses. Canadian Geotechnical Journal., 101: 1·11.

Where on Earth is Pennafrost? 139

Lachenbruch, A.H., T.T. Cladouhos and R. W. Saltus, (1988). Permafrost temperature and the changing climate.,Fifth International Conference on Permqfrost, Trondheim, Norway, 3: 9-17.

Lindsay, J.D. and Odynsky, W.( 1965) Permafrost in organic soils of northern Alberta. Canadian Journal of Soil Science, 45: 265-269.

Luthin, J.N. and G.L.Guymon. (1974) Soil moisture-vegetation-temperature relationships in central Alaska, Journal of Hydrology 23: 233-246.

Nelson, F .E. (1986) Permafrost distribution in central Canada: application ofa climate-based predictive model, Annals of the Association of American Geographers 76(4): 550-569.

Nelson, F. E. (1989) Permafrost zonation in Eastern Canada: A review of published maps, Physical Geography 10: 233-248.

Nelson, F .E. and S.1.0utcalt (1987) A frost index number for spatial prediction of ground frost zones, Arctic and Alpine 19: 279-288.

Nicholson, F.H. and Granberg, H.B. (1973) Permafrost and snowcover relationships near Schefferville, 2nd International Con! on Permafrost. Yakutsk, North American Volume: 151-158.

Riseborough, D.W. and M.W. Smith, (1998) Exploring the limits of permafrost, Proceedings, Seventh International Conference on Permafrost, Yellowknife, 935-942.

Romanovsky, V.E. and T.E.Osterkamp. (1995) Interannual variations of the thermal regime of the active layer and near surface permafrost in Northern Alaska, Permafrost and Periglacial Processes 6 (3): 313-335.

Rouse, W. (1984) Microclimate of Arctic tree line, 2: Soil microclimate of tundra and forest, Water Resources Research 20(1): 67-73.

Rouse, W. (\993) Northern climates, in Canada's Cold Environments, H. M. French and O. Slaymaker (eds). McGill-Queen's University Press, 143-167.

Smith, M.W. (1975) Microclimatic influences on ground temperatures and permafrost distribution, Mackenzie Delta, Northwest Territories, Canadian Journal Earth Sci. 12: 1421-1438.

Smith, M.W. and Riseborough, D.W. (1983) Permafrost sensitivity to climate change, 4th International Permafrost Con!, Fairbanks. AK, Nat. Academy Press, Washington, DC: 1178-1183.

Smith, M.W. and D.W. Riseborough (1996) Ground temperature monitoring and detection of climate change, Permafrost and Periglacial Processes 7(4): 301-310.

Williams, PJ. and M.W. Smith (1989) The Frozen Earth: Fundamentals of Geocryology, Cambridge University Press, Cambridge, England, 306 p.

Wright, J.F. (1995) A Hybrid Modelfor Predicting Permafrost Occurrence and Thickness, Unpublished MA thesis, Carleton University, Ottawa, Canada, 80 p.

Zoltai, S.C. (1971) Southern limit of permafrost features in peat landforms, Manitoba and Saskatchewan, GeolOgical Association of Canada, Special Paper #9: 305-310.

TYPICAL ASPECTS OF COLD REGIONS SHORELINES

Mary-Louise Byrne Geography and Environmental Studies Wilfrid Laurier University Waterloo, ON N2L 3C5

Jean-Claude Dionne Dept. Geographie et Centre d'Etudes nordiques Universite Laval, Quebec, QC GIK 7P4

Abstract

Cold regions shorelines (about one-third of the world's coastline) do not differ much from those of mid- and low-latitude regions because the fundamental factors governing their development (lithology, structure, tectonics) are nearly the same. Major features (cliffs, platforms, drowned valleys, dunes and beaches) are also found in the three major morphoclimatic regions (warm, temperate, cold). Many differences, however, do exist in the details and in the processes involved in shoreline development. There are two main categories of factors to consider. Two climatic factors (ice and frost) play an important role in shaping cold regions coasts, whereas the geological background, particularly the Quaternary heritage (glaciations and marine submergence) is of major importance for an adequate knowledge of cold regions coastal features (paraglacial vs. periglacial shorelines).

The role of ice and frost in the different coastal environments in cold regions is generally relatively well known. Detailed studies, however, when compared to mid- and low-latitude regions, are much less common and specialized. Large areas have not been surveyed adequately, and data on processes, particularly over the long-term, are missing almost everywhere. This is especially true in the Southern Hemisphere. The coastal environments for which data exist are mainly deltas, beaches, tidal flats and margins, cliffs and platforms, and the shallow inner portion of the shelf.

First, three typical aspects will be examined and discussed briefly: ice action on rocky shorelines and in intertidal marshes, frost action on rocky shorelines and in intertidal marshes, and the debris content of the ice cover - a prerequisite for ice rafting. Finally, cold coastal dunes and aeolian processes will be summarized.

141

K. Hewitt et al. (eds.), Landscapes of Transition, 141-158. © 2002 Kluwer Academic Publishers.

142 Mary-Louise Byrne and Jean-Claude Dionne

A better knowledge of cold coastal features and processes is necessary not only to preserve these fragile environments, but also for planning adequate economic development and a better use of this natural resource.

1. Introduction

In general terms, cold regions coasts may be dermed geographically as those areas where frost and ice processes are active during a period of the year which is sufficient to have a significant, if not permanent, impact on the near terrestrial, coastal and marine environments. This implies a minimum intensity and duration of frost and ice action according to latitude and exposure. Regions where frost and ice are briefly or rarely present should not necessarily be included in cold regions. However, ice and frost that are active for only a short period may form quasi-permanent features in areas of low wave action. Conversely, the effect of ice and frost action over prolonged periods may be removed quickly by strong wave action in a short period of time. Koster (1988) de:fmed a cold climate region as an area where the mean annual air temperature was less than 3°C, or where the coldest mean monthly temperature is below -3°C. But derming a cold coast is a complex task requiring a balance between coastal processes (waves and tides) and the influence of ice processes at anyone site (Viles and Spencer 1995).

Nichols (1961) dermed cold regions coasts as those where there is the presence, or a history of abundant sea ice, lake-ice, deeply frozen ground, or glaciers that terminate in water. Figure 1 depicts the extent of cold coasts in the northern hemisphere and Figure 2, the glacier cover, ice shelves and cold coasts of the southern hemisphere. The ocean margins, bays, seas and lakeshores in this area contain almost every type of possible cold coastal configuration. There is a great variety of processes and marked morphological differences, according to latitude, the extent and duration of the ice, and length of the frost season.

In North America (Figure 1), cold coastal environments include fresh and salt water shorelines from the Great Lakes to Nova Scotia, and from New England (Maine) to the Arctic. The ice and frost seasons range from 2 to 12 months. At lower latitudes, where the duration ofice cover is shorter, cold processes are important because they affect the seasonal distribution of sediments and can have a lasting impact on the shoreline in general. In Canada, only the Pacific coastline of British Columbia can be excluded from cold regions, although there are some effects of ice and frost locally on the northern B.C. shoreline.

Elsewhere in the world, cold regions coastlines occur in Alaska, in northernmost Europe (Norway, Sweden, Finland), in Russia and Siberia (Barents Sea, Kara Sea, Laptev Sea, East Siberian Sea and Chukchi Sea), in Greenland and Spitsbergen. In the Southern Hemisphere (Figure 2), they characterize Antarctica, most sub-Antarctic Islands, and some of the southernmost areas of Chile and Argentina.

Typical Aspects of Cold Regions Shorelines

- Cold coast o Ice sheet Permanent ice shelf

• Limit of winter ice pack or complete cover

143

Figure I. The extent of cold coasts in the northern hemisphere. Coasts were categorized as cold when they had evidence of subsea pennafrost, subaerial continuous and discontinuous pennafrost, or were at the approximate limit of seasonal pack ice. As well, the Great Lakes were included because they are covered with ice for many months of the year and have a significant contribution of sediment movment that results from ice fonnation.

This paper will focus on coasts that experience ice and frost seasons, exclusive of currently glaciated coastlines (tidewater glaciers). The purpose is:

1) to outline the importance of cold coasts in Canada; 2) to compare the processes and morphology oflow and high latitude coastlines, concentrating on the typical morpho-sedimentologic aspects of shore platforms, tidal marshes, the debris content of ice cover, and niveo-aeolian processes in coastal dunes; and, 3) to identify the more important cold regions coastal processes. In particular, examples from regions where ice and frost action on shorelines occur annually during a 4 to 6 month period, are presented. This is to help contest the traditional


concept that ice infested coasts are generally protected from wave action and are only slightly modified by ice and frost action (Zenkovich 1967).

.<1 20 / ___ ....aaGl~rtblt"'~

eo

'00'

.... Cold coast o lcesheet

/ 1'--

/ r '" - , '?' ,.. ....

<G. / v I'

o /

'

FIU.land \ . ~ ", , , .. ··"· .. r-

I , , ,,' ...... '-----

"U

J> ()

()

\ \ I I AAtarclica \ Po"",,",,,

\ I \ ~) I \ I I \ \ \ , \ )

o ( " \ ~ \

"<J \ -t- I

I

J lM3ddel/

Sea

• Permanent ice shelf

,..,.

/'

,." ..... :---- ... ,

• South Pol.

East AAtarclica

. ...... , ,'- .... ,

" .. _- --

.w Figure 2. The glacier cover, ice shelves and cold coasts of the southern hemisphere .

2. Importance of cold regions coastlines

· • I

• f) I , , , ,

I , .. J

'/ • J

I

,

.,,'

"'\

'f

ed'

It"~

Although there have been no precise measurements, the total length of cold regions shorelines amounts to at least one-third of the world total. On this basis alone, the continued neglect of this environment can hardly be justified. Meanwhile, the expanding interest in the natural resources of cold regions, and advances in cryo-technologies made during the last few decades, have provided opportunities to study the coasts of cold regions, particularly in the northern Hemisphere. In fact, over 1000 papers have been written in English, and there are some relevant reviews of specific aspects of cold coasts (John and Sugden 1975: Taylor and McCann 1983; Dionne 1988; 1989; Forbes and Taylor 1994; Trenhaile 1987, 1997). However, compared to mid-and low-latitude regions, the

Typical Aspects of Cold Regions Shorelines 145

processes and morphological features of coastlines in cold regions need to be better documented. The processes involved are well-known, but quantitative data are rare or missing almost everywhere.

Canada may well have the longest coastline of any country (Trenhaile 1998, p. 268) and most of the coastal zones are located in cold regions (Taylor and McCann, 1983). Forbes and Taylor (1994) estimated Canada's shoreline to be about 250,000 km long.

However, because such measurements have been made on relatively small scale maps, the length could be as much as 400,000 kilometers. This enormous length of shoreline, and commensurate variations in geology (lithology, physiography, structure and tectonics), wave conditions, tidal range and other aspects, produce coastal environments that are exceptionally diverse, and of major scientific and economic interest. It is important to understand not only the current state of coastal systems, but also to comprehend that alterations to these systems may result from impacts of projected climate change.

3. Distinctive Features of Cold Coastlines

In general terms, many aspects of cold shorelines are similar to those of warmer areas. Cold coasts are affected and shaped by waves, currents and tides. They have the same large morphological features as coasts in warmer climates; that is, embayments, cliffs and platforms and associated stacks and caves, terraces, deltas and tidal flats, beaches, barriers, spits and dunes. There are, however, some major differences, which can be attributed to climate and related processes. Forbes and Taylor (1994) suggested that the unique characteristics of cold coastal processes are due to perennial or seasonal sea ice, permafrost, ground ice, frost action, isostasy, glacial history including isostatic rebound, and other zonal factors.

Rocky shorelines are widespread in coastal regions in both hemispheres, averaging between 80 and 85 percent of the total length of the coastline. In high latitude regions, there are only two broad geographical areas where the marine shoreline is mostly sculptured into unconsolidated deposits: in northernmost Alaska and along the Beaufort Sea coastline in Canada; and, along the Arctic coastline of Russia and Siberia. Here thermo-abrasion is the dominant process of evolution, particularly in the ice-rich permanently frozen, fme-grained sediments.

Ice and frost action are two of the more characteristic agents at work in contemporary cold environments, whereas the legacy of glaciations during the Quaternary often governs and dominates the recent and present evolution of cold regions shorelines. Thus, Taylor and McCann (1983) studied shorelines in northern Canada within the general framework of shoreline dynamics with ice being an additional variable in the system. As on all coasts we also recognise a major distinction between cold rocky and depositional (beach, marsh, estuarine, deltaic, and dune ... ) types of shorelines. This paper will focus on non-glacerized shorelines, as glaciated coasts and tidewater glaciers have been covered elsewhere.


3.1. ICE AND FROST ACTION ON ROCKY SHORELINES: CANADIAN EXAMPLES

According to Trenhaile (1983 p. 77), "the rate and mode of development of rock coasts in cool environments need to be understood, not only because they constitute a significant portion of the world's coasts, but also because much of the temperate zone experienced, and often still bears the imprint of, erosion which occurred under cool climatic conditions during the Pleistocene." Yet he adds, "there is a lack of even basic agreement on the efficiency of coastal processes in high latitudes" (Trenhaile 1983 p. 77). However, in the last ten years a number of studies have begun to examine these processes. Again according to Trenhaile (1997, p. 307) "Shore ice, pack ice, and rock-shod glacial ice are thought to abrade, quarry, and transport debris across well-developed shore platforms in the Antarctic Peninsula" . And there is "clear evidence in this area of the erosive efficacy of shore ice, and probable evidence of its ability to planate rocky intertidal zones" (ibid). Trenhaile expressed the caveat that the lithology, structure, shore processes and sediment availability in the particular study areas summarized, made it easy for shore ice to perform its work. Curiously, few studies have been concerned with the evolution of Canada's cold shorelines under present climatic conditions (Dionne 1988, 1993, 1993b). Dionne and Brodeur (1988) identified and documented mechanisms of frost weathering and ice action in shore platform development, concluding that they are important agents in shaping rocky shorelines, at least locally. Of course, sedimentary rocks are prone to, and more affected by, frost and ice action than igneous and metamorphic rocks, particularly in areas where the bedrock has been eroded and shaped by glaciers and exposed to shore erosion only recently. Coastal scientists still need to determine if these cold region mechanisms operate as efficiently in areas less favourable to the work of shore ice.

The opinions of Bird (1967) and Zenkovich (1967) among others, concerning the minor role frost weathering and ice play in the evolution of rock shorelines, is apparently based on insufficient field surveys. They depart from statements made long ago by Chalmers (1882) and Goldthwait (1933), and by other workers concerning the strandflat, ahigh-Iatitude feature for which various interpretations have been suggested, including that of Nan sen (1922), who emphasized the role of frost shattering and waves in removing the debris. More recent studies generally support the older views (Dawson et al. 1987). Improved knowledge of shore platform development in cold regions should contribute to a better understanding of raised coastal platforms cut during the Quaternary or earlier (Gray 1978; Dawson 1980) and also to the complex, high-latitude strandflat.

3.1.1. Frost Action There are two important frost weathering processes active in the shore zone contributing to shore platform development: frost shattering and frost wedging. Frost shattering is the breakdown of consolidated rock under pressure related to freezing and thawing of water in the micropores. Frost wedging is the opening of fissures and prying apart of large rock fragments by ice formed in the fissures. It is thought that both mechanisms operate together and playa relatively significant role in physical weathering of bedrock exposed to the shore zone, even though there are few quantitative data available and the role offrost is not proven (Allard and Tremblay 1983; Hansom 1983; Fournier and Allard 1992; Dionne 1993 1993b).


Frost shattering has been observed on the eastern coast of Hudson Bay in various lithologies including basalt, sandstone, quartzite, shale and dolostone (Michaud et a/. 1989). In Ungava Bay (Fournier and Allard 1992), frost shattering occurs in igneous and metamorphic rocks; whereas, in the St. Lawrence Estuary, frost shattering is common in shale, slate and limestone (Dionne and Brodeur 1988). Along the same shorelines, frost wedging is also a common process for disjointing large rock fragments, a prerequisite for their subsequent removal by ice.

The third mechanism observed in cold regions is frost heaving - the predominantly upward movement of material during freezing caused by the movement of water toward the freezing plane and then its expansion upon freezing (Washburn 1979). Typical features of frost heaving in bedrock, however, are rare or absent in the shore zone although they are relatively abundant in bare bedrock above the 10 m level in areas with permafrost. It is not known if similar features observed on shore platforms of the St. Lawrence Estuary are produced by the same mechanism.

3.1.2. Ice Action Erosion mechanisms related to shore ice are more varied and complex than those related to frost action. The first is abrasion. Minor abrasion features made by shore ice, such as scratching, striations, and small grooves, are relatively common on rock shorelines, and are occasionally confused with glacial abrasion marks. Along the Hudson Bay shoreline, for example, these marks are often superimposed on glacially polished and striated surfaces. The amount of debris removed by abrasion is, however, of little significance for overall shore platform erosion.

Percussion marks which resemble glacial grooves, are made directly by ice or by rock fragments frozen at the base of ice floes. They have been observed on limestone strata interbedded with shales on wide, nearly horizontal shore platforms along the St. Lawrence Estuary (Dionne 1993). It is a new erosion mechanism whose occurrence has not been widely documented elsewhere. However, locally in this area, it is important for the formation of shore platforms cut into soft bedrock.

"Ablation", that is dislodgement and removing of rock material, is a common and significant process observed everywhere on rocky shorelines in various lithologies. Large fragments produced by frost shattering and frost wedging are dislodged and removed by ice pushing and rafting. This leaves cavities of various sizes, and produces small benches or terracettes. The process is particularly evident on glacially polished and striated rock surfaces out-cropping on the east shore of Hudson Bay, and the northern shore of the Gulf of St. Lawrence. Fournier and Allard (1992) considered this process to be of major importance on the rocky shores ofUngava Bay.

"Ice push" is the movement, offshore or shoreward, of rock fragments over the bedrock surface by ice. Rock fragments are either scattered on the platform or concentrated in shallow depressions and along small scarps. They may also form large block ridges at the high tidal level, in which flat fragments are often in a nearly vertical position (Dionne 1988, 1988b).

Dionne and Brodeur (1988) stressed the importance of ice on shore platforms in the St. Lawrence Estuary, which cuts into shale and slate. The platforms, particularly where they are folded, are often scarped and leveled by ice. Ice floes grounded along the bottom can scrape, grind, and level the surface producing remarkably horizontal platforms.


At breakup, the platfonn surface shows a series of elongated depressions perpendicular to shoreline. The small fragments removed by ice are subsequently carried out by waves and tidal currents leaving a bare rock surface.

Plucking also occurs on shore platfonns in sedimentary rocks. Scoured rockpools up to a few meters across and about 20 cm deep, are produced directly by ice floes moving up and down in response to tidal fluctuations. Rock fragments removed from the substrate are usually left on the seaward side of the cavities. Occasionally, single rock fragments are removed and transported over some distance leaving a shallow angular depression (Dionne and Brodeur 1988).

Isolated erratic boulders commonly occur on shore platfonns, and they are frequently displaced various distances by ice floes at breakup. In the St. Lawrence Estuary, the platfonns cut into shale are locally scoured by boulders pushed by ice floes. Under pressure exerted by ice floes when moved by tidal currents, boulders plow and scour the. substrate. At Neuville, Quebec, Dionne and Brodeur (1988) documented boulders up to 50 tonnes that had scoured the underlying bedrock surface and left furrows up to 25 meters long and 40 cm deep.

Another important ice process is rafting (Dionne 1993 1993b; Drake and McCann 1982). Physical weathering produces fragments of various sizes, fonning a litter that may progressively cover the shore platfonn and halt its evolution. For continuous erosion, it is necessary that the debris be removed by waves and currents, or by another agent. In cold regions, shore ice action, particularly in protected sites, is often more important than waves and tides in removing debris from the surface of the shore platfonns. Debris produced by frost weathering and other mechanisms are incorporated into the ice cover during the winter or frozen at its base. A large volume of debris is removed and rafted away at breakup. The surface of the platfonn is thus cleared of protective debris, and the erosional processes remain active over a longer period, thereby allowing a wide platfonn to develop.

Frost and ice action in shore platfonn development vary with latitude, geographical setting, and hydrological factors. Higher latitude regions are not necessarily the most affected by their combined action. Frost action, for example, works best in areas where there are many cycles of freezing and thawing. But, higher latitudes have fewer freeze-thaw cycles than lower ones. Additional field work and quantitative data are necessary to evaluate correctly the exact role of frost and shore ice in the development of shore platfonns in cold regions.

3.2. FROST AND ICE ACTION IN TIDAL MARSH DEVELOPMENT

Intertidal marshes in cold regions are far less well-known than those of mid- and lowlatitudes, and the literature is mostly concerned with the botanical and ecological aspects (Frey and Basan 1978). Yet, there are salt marshes throughout cold regions of the northern and southern hemispheres (Chapman 1977). In North America, coastal marshes occur in eastern Maritime Canada, Labrador, Ungava Bay, James and Hudson Bays, in many protected sites in the Arctic Archipelago, and along the Beaufort Sea shoreline. The largest tidal marshes occur in Nova Scotia and New Brunswick, the St. Lawrence Estuary and Gulf, and in James and Hudson Bays. Most marshes are small fringing fonns at the head of shallow embayments, and no large intertidal marshes have been documented in higher latitudes in Canada. Frost and ice action influence the development of tidal marshes in all


these regions. The role of frost increases with latitude, whereas ice action is more important in meso and macro tidal environments in lower latitude regions.

3.2.1. Frost Action Frost operates seasonally and in association with permafrost. It is of course, more pronounced in higher latitudes, but minor features directly related to seasonal frost also occur on mid-latitude intertidal marshes. The role of frost action in coastal marsh development is still not well-documented. The occurrence of permafrost, for example, has been reported only recently (Dyke 1988; Allard et al. 1992), and the effect of seasonal frost has not been adequately studied.

Tidal marshes in James, Hudson and Ungava Bays, and the Beaufort Sea, are influenced by permafrost and seasonal frost. Segregated ice, formed by the movement of water to the freezing plane and resulting in lenses or layers of ice in the soil (Mackay 1972), raises the surface of the marsh above the higher water level and modifies the normal conditions of their development. Small mounds with polygonal patterns characterize the lower marsh surface, whereas the high marshes are covered by larger and higher mineral mounds (palsas) and thermokarst depressions. Frost heaving is another mechanism observed in tidal marshes influenced by seasonal frost. Boulder-sized clasts, interbedded in the underlying muddy or silty-clay substrate, are progressively ejected at the surface. This process, along with ice rafting, plays a role in concentrating coarse debris at the surface of tidal marshes, a feature unique to cold regions. In addition, thawing of ground ice often results in shallow depressions that modify the topography of the tidal marsh (Fournier et al. 1987).

In summary, the range offeatures produced by frost is relatively small, and their role has not yet been quantified or adequately evaluated. However, these are, evidently features unique to cold coasts. Depending upon the geographic location, frost and freezing indirectly aid in tidal marsh development through the formation of an ice cover at the surface of the tidal marsh and in the offshore zone. This protective role of ice will now be discussed.

3.2.2. Ice Action Shore ice action in tidal marshes was reported as long ago as 1875 by Hind. Bancroft (1902) and Massart (1907) also discussed ice action, but its role has been defined and quantified adequately only recently. Dionne (1989) summarized typical shore ice effects, taking account of protection, erosion, and sedimentation.

First, it is clear that ice cover protects tidal marshes from erosion by waves during winter. However, since this is a coastal environment characterized by low-wave energy, the protection offered in most areas is of limited significance. Meanwhile, it now seems the protective role of the ice cover is often less efficient than many past workers assumed. In meso- and macro-tidal environments of lower latitude cold regions, the ice cover is generally not frozen to the bottom. At high tide, the flood penetrates under the ice carpet and causes deposition, whereas erosion is usually associated with ebb currents. This aspect is still not well-documented. At breakup, intertidal marshes, particularly at their seaward margins, suffer much mechanical erosion by ice floes.

Second, the general physiography or morphology of most tidal marshes in cold regions does not differ much from temperate region marshes (Dionne 1989b). Both are


often characterized by shallow depressions or "salt pans" of various shapes and sizes. The difference is in the origin of the depressions (Fournier et aJ. 1987). In cold regions, most pools or basins are related to ice and frost action (Dionne 1989). Thermokarst depressions are more common in the marshes of higher latitudes (Fournier et al. 1987). However, icemade depressions resulting from the removal of pieces of turf from the marsh surface are common in lower latitudes in meso-and macro-tidal environments (Dionne 1989). Dionne (1989) concluded that a large volume of debris is removed by ice floes in tidal marshes over the lower St. Lawrence Estuary. He found that marsh clumps released in the tidal zone modify the morphology of the marsh surface, and create a micro-relief which influences sedimentation processes during the ice free season, conversely, the scars left by the clumps removed by ice floes create a pattern of depressions in the marsh surface.

Another typical effect of shore ice is the 'mowing' of the grass cover. Although this process does not disturb the substrate, it plays a role in sedimentation and is the main factor involved in the low organic content of most tidal marshes in cold regions. It may also explain why, in contrast to the vast quantities of peat in fresh water cold environments, few organic or peaty tidal marshes have been documented. When moved by the tide, ice floes not only remove the vegetation cover, but also scratch the marsh surface and even scour it leaving shallow elongated depressions. Similar furrows are also made by ploughing boulders pushed by ice floes (Dionne 1988). Another morphologic characteristic of cold regions tidal marshes, especially at higher latitudes, is the absence, or rarity, of a well-developed tidal creek pattern. No satisfying explanation has yet been suggested.

There are also marked differences between sedimentation in salt marshes of cold and temperate regions. Shore ice is the main agent responsible for this difference, and the principal characteristic is the abundance of coarse debris (Dionne 1972). Intertidal marshes elsewhere are typically a fme-grained sedimentary environment, but in cold regions, most tidal marshes are boulder strewn (Dionne 1989). The abundance of coarse debris, of course, depends mainly on the presence of a nearby source, but also results from ice rafting. Boulders may have been rafted, pushed, or rolled. Commonly, however, a large percentage of coarse debris is a lag from eroded glacial till or stony marine deposits. In James Bay, UngavaBay and the St. Lawrence Estuary, for example, the substrate of most tidal marshes is an erosion surface cut into marine clay containing ice-rafted clasts. Removed by shore ice, these clasts may fonn boulder pavements. Locally, large to small boulders may also have been reworked and deposited at the marsh surface. Once again, there are few quantitative data on the annual input and output by ice rafting in coastal marshes.

In tidal marshes where the offshore water is turbid, as in the middle St. Lawrence Estuary or the eastern Foxe Basin, the ice cover contains a large volume of mud (Dionne 1984). In these marshes, a problem is to evaluate the volume of sediment brought in or moved out by ice annually. In the St. Lawrence Estuary, the output is more important than the input (Dionne 1984, 1997). Of a few million tonnes of mud trapped in the ice cover in the tidal flats of the middle st. Lawrence Estuary, only a small percentage (5 to 10 percent on average) is released in the tidal marshes. This may largely explain the low rate of lateral and vertical accretion of most tidal marshes in cold regions, even where offshore water is very turbid.


Geologists often question the potential of preservation of ice-made features in tidal flats and associated marshes. However, if very few papers report ice-made features, they do exist even on emerging shorelines. Dionne (1998) documented typically ice-rafted clasts dispersed in fine-grained, thinly stratified and laminated sediments, with various types of deformations and bowl-shaped circular and elongated sedimentary structures on tidal flats. These features have been observed in emerged tidal marsh deposits in the middle st. Lawrence Estuary, providing evidence that ice-made features can be preserved and be useful tools for identifying former environmental and sedimentary conditions.

Another agent plays a role in the evolution of many cold regions tidal marshes. In the Northern Hemisphere, geese feeding in tidal marshes, commonly cause local erosion. In the bulrush tidal marshes of the middle St. Lawrence Estuary, erosion by the Greater Snow Geese is an important factor to consider in tidal marsh development. The muddy sediment disturbed by geese feeding in the low tidal marsh, is moved easily by waves and tidal currents, and exported offshore causing vertical erosion of the low marsh surface, as evidenced by perched boulders. In comparison with the tidal marshes of temperate and warmer regions, the activity oflower invertebrates, worms, crustaceans, and other animals in cold regions, is usually not significant, bioturbation being caused mainly by plant roots (Martini 1981, 1991).

In summary, cold regions tidal marshes differ in several respects from those of lower latitudes. Frost and ice are the main agents involved. Presently, only a few detailed morpho-sedimentologic studies are available for tidal marshes in the Bay of Fundy, St. Lawrence Estuary, and James and Hudson Bays. There is thus the need for additional information, particularly in the higher latitudes of both hemispheres.

3.3. DEBRIS CONTENT OF THE ICE COVER AND THE POTENTIAL OF ICE RAFTING

Most specialists recognize that ice rafting is a common process almost everywhere in cold region coastal environments, although few quantitative data have been published. Recent work by Reimnitz et al. (1993) in the Beaufort Sea and in the Arctic Ocean, and by Barnes et al. (1993) in Lake Michigan, for example, have pointed out the importance of sea and lake ice rafting and called for a better examination of the role of shore ice in sedimentation. In particular, the traditional concept that ice rafting is essentially performed by icebergs needs to be re-examined. Shore ice and river ice contribute many millions of tonnes of sediment of all grain sizes to inland seas, continental shelves and the deep oceans. In the field, debris is usually observed at the surface of the ice cover, at the base, and at different levels, particularly when the ice carpet is one metre or more thick. The content varies greatly from site to site and from year-to-year, but a huge volume of debris is commonly observed in shore ice where sediment is available and where ice is in contact with the bottom. A prerequisite for ice rafting, however, is the debris content of the ice cover, and we lack data on this. Consequently, it is difficult to evaluate adequately the over-all morpho-sedimentologic role of coastal ice. Yet, many processes incorporating debris in the ice cover are known (Dionne 1968, 1970, 1993b), and there is no doubt that shore ice offers better conditions for incorporating sediment into the ice cover than offshore ice.

If offshore ice is usually debris free, many million tonnes of sediment are effectively rafted annually by coastal ice and we must remember the huge potential that the


enormous length of cold regions shorelines represents. Ifwe add to this volume of material rafted by icebergs, floating ice is certainly the main agent of coarse debris transfer from the land to the continental shelves and to the surrounding ocean basins in cold regions. The opinion that coastal ice is generally poor in debris, and its morpho-sedimentologic role is insignificant, is apparently based on inadequate surveys and a lack offield work. In the St Lawrence Estuary for example, an area where the ice season extends from December to April, coarse debris in the ice cover amounts to between 1 and 2 million tons annually (Dionne 1993, 1993b). Fine-grained sediment most likely represents a volume, exceeding 3 to 4 million tons. Is this estimate representative of what happens elsewhere? It is difficult to answer this question. However, we should remember that the estimates made by, for example, McCann and Dale (1986) for Winter Bay and by Gilbert (1983) for Pangnirtung Fiord, are larger than the estimate for the maximum turbidity zone of the middle St Lawrence Estuary (Dionne 1984, 1997).

3.4. NIVEO-AEOLIAN PROCESSES IN COLD COASTAL DUNES

Sandy beaches and coastal sand dunes represent a transitional area between the marine or lacustrine and the terrigenous environments. At higher latitudes, beaches are generally made of coarser sediments such as cobbles and boulders (Taylor and McCann 1983), sand dunes are rare. Moving southward into the subarctic, sandy beaches and coastal dunes are more common, especially at the mouths of rivers (Ruz and Allard 1994). This reflects an increase in the supply of fmer sediment to the shoreline. Together with this increase in supply, there is increased importance in the presence of vegetation on the shoreline. Coastal dunes in Canada are found more frequently in the Great Lakes Basin and on the east coast of Canada, especially in Prince Edward Island (McKenna Neuman 1993).

Koster and Dijkmans (1988, p. 153) quoted Cailleux (1978), who defmed niveoaeolian deposits as mixed deposits of wind driven snow and sand, silt, vegetal debris, or other detritus. The (micro-) forms generated through melting and/or sublimation of snow are called denivation forms. Their conclusions included that niveo-aeolian and denivation processes play an important role in modern periglacial or cold-climate dune environments and that denivation features will be preserved as paleoclimatic indicators in aeolian strata under conditions such as a distinctly unimodal wind regime and/or high migration rates of the dunes. Fossil denivational structures will occur as various local disturbances of the aeolian stratification. They also remarked that niveo-aeolian transportation and deposition of dune sands is favoured by cold or semi-arid conditions, rather than by snow-rich climates.

Ruz and Allard (1994) investigated foredune development along a subarctic emerging coastline in eastern Hudson Bay. They discovered that niveo-aeolian processes were significant, that rates of sedimentation in the back shore area were lower than in temperate areas, and that there were two sources of sand. The frrst was relatively fine sand supplied by erosion of the fall storm berm. The second was relatively coarse sand which was supplied by the sand laden winter ice-foot These left distinctive, alternating layers of coarse and fine sand in the dunes. The primary depositional processes occur in winter rather than in summer because winter niveo-aeolian deposition is trapped by dormant vegetation slowing the wind. In the spring, the ice-foot disappears, cutting off that source of sediment. Vegetation grows through summer and is able to trap the sand carried from


the fall storm berm to the backshore. Because the rates of sedimentation are lower, it takes longer for the established foredune to develop in this subarctic environment.

To the south, in the Great Lakes Basin, the supply of sand is greater than in the more northerly latitudes and the importance of vegetation increases. Although summer temperatures are very warm, sediment transport can be greater in winter due to stronger storms winds. McKenna Neuman (1990) presented laboratory experiments that illustrated the importance of sublimation in freeing frozen beach sand for transport in winter. Van Dijk and Law (1995) reported from field experimentation that sediment freed from a frozen beach through sublimation, allows a ready supply for transport into the dunes at Presquile Provincial Park on Lake Ontario. Byrne (1997) documented the relative rates of transport of sand through a trough blowout at Pinery Provincial Park, Ontario where the transport of sand shows a marked seasonality with greater transport in the winter months. Coarser sediment is moved during winter when niveo-aeolian processes are active (Figure 3). The coarser layers of sediment transported by niveo-aeolian processes are distinct cold region features in coastal dunes. The volumes of sediment transported into the system from the beach during the cold periods are only recently beginning to be documented. Examination of niveo-aeolian deposits during a field season in winter 2001 at Pinery Provincial Park revealed that, even in a snow-rich climate, snow and sand deposits were remarkable (Figures 4 to 7). The winter supply of sand is significant, providing sediment from the beach and foredune to the established ridge. This supply is cut off in the summer when vegetation growth on the foredune traps the beach sand, preventing it from reaching the established ridge. It is important to emphasize that Pinery is an Oak-Pine Savannah and is noted as a northern extension of the Carolinian forest - indications of warm and temperate environments. Yet, cold coastal processes operate and contribute significantly to the overall budget of sand in the dune system. Further research needs to be undertaken to determine the importance of this supply of sand to the growth and development of the dunes.

4. Conclusion

In general, the major landforms of coastal landscapes in cold regions are similar to those found elsewhere; in detail, there are many typical features of great interest. The distinctiveness is mainly related to frost and ice processes. Because of exploration for natural resources (e.g. petroleum, diamonds), cold regions will most probably experience accelerated development during the next decade. It is thus likely and fundamental that we have a better knowledge of the physical characteristics of the land and processes involved especially in the coastal zone. Earth scientists should contribute positively to a field that has been too often neglected in the past.


Figure 3. Coarse sand layers exposed in a pit dug in a snow bank at Pinery Provincial Park.

Figure 4. Snow core showing the layers of sand that were deposited with blowing snow in a snowbank on the lee of a dune at Pinery Provincial Park.

Typical Aspects of Cold Regions Shorelines

Figure S. Exposed sediment on the beach is available for transport over the low foredune and into the established dunes by the strong winter winds. Sediment-rich ice ridges are clearly visible on Lake Huron in the background.

Figure 6. Looking up the trough blowout from the foredune ridge, patches of snow, sand and snow and sand are clearly visible.

155


Figure 7. A view looking down the trough blowout toward the lake. Surface sediment is exposed and available for transport. There are also several snow patches dotting the surface and mixtures of sand and snow that have resulted from niveo-aeolian processes in the foreground. The ice ridges on Lake Huron are clearly visible in the photograph. All are clear indications of cold processes in what is usually thought of as a temperate landscape.

References

Allard, M., Fortier, R., and Seguin, M.K. (1992) The thermal regime 0/ intertidal permafrost, George River estuary, Ungava Bay, Quebec, Canadian Journal 0/ Earth Sciences 29 (2): 249-259.

Allard, M. and Tremblay, G., (1983) Les processus d'erosion littorale periglaciaire de la region de Poste-de-IaBaleine et des i1es Manitounuk sur la cote est de la mer d'Hudson, Canada, Zeitschrift fur Geomorphologie, Supplement Band 47, 27-60.

Bancroft, J.A., (1902) Ice-borne sediments in Minas Basin, N.S. Transactions Nova Scotia Institute o/Science 11,158-162.

Barnes, P.W., Kempema, E.W., Reimnitz, E., McCormick, M., Weber, W.S., and Hayden, E.C., (1993) Beach profile modification and sediment transport by ice: an overlooked process on Lake Michigan, Journal o/Coastal Research 9 (I), 65-86.

Bird, J.B., (1967) The Physiography of Arctic Canada, Johns Hopkins Press, Baltimore. Byrne, M-L., (1997) Seasonal Sand Transport through a Trough Blowout at Pinery Provincial Park, Ontario.

Canadian Journal 0/ Earth Science 34, 1460-1466. Chalmers, R., (1882) On erosion from coast-ice and floating ice in the Baie des Chaleurs., Trans. R. Soc. Can.

1(4),285-286. Chapman, V.J. (ed.), (1977) Wet Coastal Ecosystems, Elsevier, Amsterdam, 428 pp. Dawson, A.G., (1980) Shore erosion by frost: an example from the Scottish late-glacial, In J.1. Lowe, J.M. Gray

and lE. Robinson (eds.) Studies in the Late-Glacial o/North-West Europe, Pergamon, Oxford 45-53. Dionne, J-C. (1968) Morphologie et sedimentologie glacielles, littoral sud du Saint-Laurent, Zeitschrift fur

Geomorphologie Suppl.Bd. 7, 56-84.


Dionne, J-C. (1970) Aspects morpho-sedimentologiques du glaciel, en particulier des cotes du Saint-Laurent, Quebec Lab. Reck Forestieres, Rapp. Infor. Q-F-X-9, 324 p.

Dionne, J-C. (1972) Caracteristiques des schorres des regions froides, en particulier de I'estuaire du SaintLaurent Zietschrift for Geomorphologie 13, 131-162.

Dionne, J-C. (1984) An estimate of ice-drifted sediments based on the mud content of the ice cover at Montmagny, middle St. Lawrence Estuary Marine Geology 57(1-4): 149-166.

Dionne, J.-C., (1988) Characteristic features of modern tidal tIats in cold regions. In de Boer, P.L. et al., editors, Tide-influenced sedimentary environments and facies, Dordrecht D.Reidel, 301-32.

Dionne, J.-C.(1988b) Ploughing boulders along shorelines, with particular reference to the St. Lawrence Estuary Geomorphology 1(4): 297-308.

Dionne, J-C. (1989) An estimate of shore ice action in a Spartina tidal marsh, St.Lawrence Estuary, Quebec, Canada Journal of Coastal Research 5(2): 281-293.

Dionne, J-C. (1989b) The role of ice and frost in tidal march development. A review with particular reference to Quebec, Canada Earth Science Re1Iiews 26.

Dionne, J-C. (1993) Quantitative data on coarse debris on the surface of the ice cover, Saint-Lawrence Estuary. [Donnees quantitatives sur les cailloux a la surface du couvert glacial, estuaire du

Saint-Laurent.] Geographie physique et Quaternaire 47(2): 181-192. Dionne, J-C., (1993b) Sediment load of shore ice and ice rafting potential, Upper St. Lawrence Estuary, Quebec,

Canada Journal of Coastal Research 9(3),628-646. Dionne, J.-C., and D. Brodeur, (1988) Erosion des plate-formes rocheuses littorales par affouillement glaciel

Zeitschriftfur Geomorphologie 32,101-115. Dionne, J-C., (1997) Nouvelles donnees sur la transgression Laurentienne, cote sud moyen estuaire du

Saint-Laurent, Quebec (Additional evidence for the Laurentide transgression on the south shore of the middle St. Lawrence estuary) Geographle physique et Quaternaire 51(2), 201-210.

Dionne, J-C., (1998) Sedimentary structues made by shore ice in muddy tidal-tIat deposits, St. Lawrence Estuary, Quebec Sedimentary Geology 116, 261~274.

Drake, 1.1. and McCann, S.B., (1982) Canadian Journal of Earth Sciences 19, 748-754. Dyke, L., (1988) Permafrost aggradation along an emergent coast, Churchill, Manitoba. In Senneset, K (ed.),

Proceedings, 5th International Conference on Permafrost, Trondheim, Tapir, vol. I, 138-142. Forbes, D.L.; and Taylor, R.B. (1994) Ice in the shore zone and the geomorphology of cold coasts Progress in

Physical Geography 18(1): 59-96 Fournier, A. and Allard, M, (1992) Periglacial shoreline erosion ofa rocky coast: George River Estuary, Northern Quebec Journal of Coastal Research 8(4), 926-942. Fournier, A., Allard, M., and Seguin, M.K., (1987) Typologie morpho-genetique des marelles du marais littoral

de la baie de kangiqsualujjuaq, Estuaire du George, Quebec Nordique Geographie physique et Quaternaire 41, 47-64.

Frey, R.W.. and Basan, P.P., (1978) Coastal salt marshes, in R.A. Davis (Ed.), Coastal Sedimentary Environments, New York, Spinger-Verlag, 101-169.

Gilbert, R. (1983) Sedimentary processes of Canadian arctic fiords, Sedimentary Geology 36(2-4), 147-175. Goldthwait, J. W., (1933) The sawing of cliffs and platforms by batture ice, Originally unpublished Ms, reprinted

in 1971 in Pleistocene Geology of the Central St. Lawrence,Geol. Surv. Canada. Memoir 359, 144-5. Gray, J.M., (1978) Low-level shore platforms in the wouthwest Scottish highlands: altitude, age, and correlation

Trans. Inst. Brit. Geogr.3, 151-64. Hansom,J .D., (1983) Shore platform development in the South Shetland Islands, Antarctica Marine Geology 53,

211-229. Hind, H.Y., (1875) The ice phonomena and the tides ofthe Bay of Fundy, Canadian Monthly Review 8, 189-

203. John, B.S. and Sugden, D.E., (1975) Coastal geomorphology of high latitudes. Prog. Phys Geogr. 7,53-132. Koster, E.A., 1988; ancient and modem cold-climate aeolian sand deposition: a review. Journal of Quaternary

Science 3( I), 69-83. Koster, E.A. and Dijkmans, J.W.A., (1988) Niveo-aeolian deposits and denivation forms, with speical reference

to the Great Kobuk sand dunes, Northwestern Alaska, Earth Surface Processes and Landforms 13, 153-170.

Mackay, J.R. (1972) The world of underground ice, Annals of the Association of American Geographers 62, 1-22.

Martini, LP., (1981 Ice effect on erosion and sedimentation on the Ontario shores of James Bay, Canada. Zeitschriftfor Geomorphologie, 25, 1-16.


Martini, I.P., (1991) Sedimentology of subarctic tidal tlats of western James Bay and Hudson Bay, Ontario, Canada. In D.G. Smith, G.E. Reinson, BA Zaitlin and R.A. Rahmani (Eds.) Clastic Tidal Sedimentology, Canadian Society o/Petroleum Geologists Memoir 16, 301-312.

Massart, J. (1907) Essai de geographie botanique des districts littoraux et anuviaux de la Belgique, Bulletin de la Societe Royale de Botan/que beige 44, 192-269.

McCann and Taylor McCann, S.B., and Dale, J.E., (1986) Sea ice breakup and tidal tlat processes, Frobisher Bay, Baffin Island,

PhysicaIGeography7,168-180. McKenna Neuman, C., (1990) Role of sublimation in particle supply for aeolian transport in cold environments,

Geograpfiska Anna/er 7lA ( 3-4), 329-335. McKenna Neuman, C., (1993) a review of aeolian transport processes in cold environments, Progress in

Physical Geography 17(2}, 137-155. Michaud, Y.; Dionne, J-C.; and Dyke, L.D. (1989) Frost bursting: a violent expression offi'ost action in rock,

Canadian Journal 0/ Earth Sciences 26(10}, 2075-2080. Nansen, F. (1922) The stran4flat and isostasy, Videnskapssel skapets Skrifter I, Mat.-Naturv. Klasse 1921, no.

II (Kristiana). Nichols, R.L., (1961) American Journal o/Science 259, 694-708. Reimnitz, E., Barnes, P.W., and Weber, W.S., (1993) Particulate matter in pack ice of the Beaufort Gyre, Journal

o/Glaciology 36, 186-198. Ruz, M.-H. and Allard, M. (1994) Coastal dune development in cold-climate environments, Physical Geography

15(4), 372-380. Taylor, R.B., and McCann, S.B. (1983) Coastal depositional landforms in northern Canada, in J.T. Lowe and

A Dawson (eds.), Shorelines andIsostasy (Academic Press, London), Institute of British Geographers, Special Publication 16, 53-75.

Trenhaile, A.S., (1983) The development of shore platforms in high latitudes, in Shorelines and Isostasy, ed. D.E.Smith and A.G.Dawson (Academic Press, London), Inst Brit Geogr. Special Publ. 16,77-93.

Trenhaile,AS., (1987) The Geomorphology 0/ Rock Coasts, Oxford: Clarendon Press, 384 p. Trenhaile, AS., (1997) Coastal Dynamics and Lan4forms, Oxford University Press, Oxford, U.K., 366 p. Trenhaile, A.S., (1998) Geomorphology: a Canadian Perspective, Oxford University Press, Toronto 330 p. van Dijk, D. and Law, J. (1995) Sublimation and aeolian sand movement from a frozen surface: experimental

results from Presqu'i1e Beach, Ontario Geomorphology 11, 177-187. Viles, H. A and Spencer, T. (1995) Coastal problems: geomorphology, ecology, and society at the coast,

London: Edward Arnold, 350 p. Washburn, AL., (1979) Geocryology: a survey o/periglacial processes and environments, New York:

Wiley. Zenkovich, V.P., (1967) Processes o/Coastal Development, Oliver and Boyd, Edinburgh.

LANDFORM DEVELOPMENT IN AN ARCTIC DELTA: THE ROLES OF SNOW, ICE AND PERMAFROST

H. JESSE WALKER Department o/Geography and Anthropology Louisiana State University Baton Rouge, LA 70803-4105

Abstract

Among the distinctive characteristics of arctic deltas is the presence of snow, ice and permafrost. In many other aspects arctic deltas differ little from deltas found elsewhere, i.e. they have distributaries, sand bars, mudflats, sand dunes, subaerial and subaqueous portions and various kinds oflakes. Included among the unique forms in arctic deltas are ice wedges, ice-wedge polygons, pingos and thermokarst lakes.

The 600 km2 Colville River delta in northern Alaska illustrates well the forms and processes typical of arctic deltas and also illustrates the relationship between an arctic delta and its drainage basin, the ocean, the atmosphere and the biosphere. The drainage basin, which occupies about 113 of the North Slope of Alaska, is confined entirely to the continuous permafrost zone north of the Brooks Range. It is a region where snow cover (albeit generally thin) and river, lake and sea ice last for eight to nine months. During that period of time, geomorphic activity is minimal. Hydrologically, the sea is connected with the river in the deeper parts of the river's channels so that sea water penetrates beneath river ice some distance upstream during winter. In contrast fresh water flow during winter from the basin is virtually nil because of the frozen surface of water bodies and a frozen active layer which eliminates ground water as a source.

Snow melt, which generally begins in May, gradually leads to breakup. A prebreakup flood period is followed by breakup and post-break flooding; combined they last about three weeks. During that period, flood water flowing against riverbanks can cause thermoerosional niching which often leads to bank collapse. The character of the bank collapse depends to a large extent on the texture and type of frozen bank material.

During breakup flooding, water flows over the bottom-fast ice of the river, into connected lakes and out to the sea. Much of the river's suspended load is deposited in these overflow areas. As floodwater progresses out over the sea ice it eventually flows down through pressure ridge cracks and scour holes to progress seaward beneath the floating ice.

159

K. Hewitt et al. (eds.), Landscapes o/Transition, 159-183. © 2002 Kluwer Academic Publishers.

160 H. Jesse Walker

Being a low-lying area and being situated in the Arctic, arctic deltas are vulnerable to climatic and sea-level changes. Changes in temperature, snow fall, duration and extent of sea ice, and active layer thickness are among the many changes that would have major impacts on deltaic processes and forms.

1. Introduction

Deltas, among the most changeable oflandscapes, occupy only about two percent of the earth's coastline. Some of the most distinctive of these deltas are those that have formed in the Arctic. They, like their counterparts in other environments, range in size from very small to that of the Lena River which, with an area of 32,000 kIn2, is the world's third largest. Arctic deltas also range in age from those just beginning to form (e.g. in lakes that have recently been tapped) to those continuing their formation since sea level reached its nearly still-stand position 5000 years ago.

Geologically speaking arctic deltas are young, but most of them have had time to develop forms, such as distributary channels, sand bars, mudflats, dunes and lakes, typical of deltas elsewhere. In addition, arctic deltas also possess forms uncommon elsewhere including ice wedges, ice-wedge polygons, frost mounds, pingos and thermokarst features.

Among the many natural features associated with arctic deltas, those of snow, ice and permafrost are especially important. Snow and subaerial forms of ice are seasonal whereas permafrost and most types of ground ice are perennial. Although snow, ice and permafrost are distinctive and highly varied entities, their impacts on the landscape are often modified by the interactions among them.

The Colville River delta is used here for the purpose of examining the many ways in which snow, ice and permafrost occur, modify and interact in an arctic delta setting. If only about 5% the size of the Mackenzie delta, the Colville delta possesses all ofthe forms and is affected by all of the processes shared by the large deltas of the Arctic.

2. The Colville River Delta: Its Setting and Physical Characteristics

The Colville River delta is located 250 kIn east of Barrow , Alaska where the Colville River enters the Beaufort Sea (Figure 1). This river, some 600 kIn long, drains about one-third of the North Slope of Alaska. Most of the river's tributaries originate in the Brooks Range. In their upper portions they flow over terrain that was formerly glaciated. During those periods when most of Canada and Alaska were under a glacial cover, however most of the North Slope was ice free. Nonetheless, it was subjected to continuous periglacial processes-processes that also prevail today. Thus, the Colville River basin lies entirely within the zone of continuous permafrost and displays those characteristics typical of such regions including: 1) a long period of subfreezing temperature, 2) low annual precipitation most of which is snow, 3) a long period of snow cover over all surfaces, 4) a thin active layer, 5) little or no river flow during winter, 6) river-ice breakup as a major hydrologic phenomenon and 7) low, tundra-type vegetation.

Landform Development in an Arctic Delta 161

The delta, with an area of about 600 km2, is roughly triangular in shape with sides that are 32, 37 and 42 km long (Figure 2). The main route of water to the sea is along its eastern flank where the channel reaches depths of 12 m in some locations. A number of distributaries branch off toward the northwest from this main channel. The most important of these is Nechelic Channel which forms the western border of the delta and flows past Nuiqsut, the delta's only town. The actual number of distributaries and other deltaic waterways varies with river stage and with the trend of erosion and deposition through time. The numerous bifurcations and rejoinings in the delta provide more than 5000 routes to the ocean once water passes the head of the delta (Walker 1983). Although the relative discharge among distributaries varies with season, the main channel, which is nearly 1 km wide near its head, carries some 70% of the annual discharge. Most of the rest of the water flows through Nechelic Channel (Amborg et al. 1966).

Figure I. The North Slope of Alaska showing the Colville drainage basin and the Colville delta.

Distributary banks in the delta, which range in height up to 9 m, are composed of a variety of materials including silt, sand, gravel and peat. In 1971 about 59% of the riverbanks were erosional, 35% depositional and 6% neutral (Ritchie and Walker 1974).

In addition to distributaries and their banks, the delta surface is characterized by a variety of other forms. Subaerial forms with the maximum relief are sand dunes which occur in many parts of the delta. Dunes are especially well developed on the lee side (left bank) of the major channels. Both stable and active dunes are present with active dune belts, one of which is more than 4 km long, occurring down-drift from the extensive bars and flats that are associated with present-day channels.

Sand bars and mudflats are found along all of the delta's distributaries as well as across nearly all of the front of the delta. Most of these bars and flats are submerged only

162 H. Jesse Walker

during breakup flooding. For the rest of the year, they are either snow covered (approximately 8 months) or exposed to the atmosphere and thus subject to deflation.

The lakes of the delta are highly varied in size, shape, depth and origin. They are present in terrace flank depressions, abandoned river channels, swales on point bars, interand intra-dune depressions, low-centered polygons (Figure 3) and the troughs between icewedge polygons (Figure 4) (Dawson 1975). They range in depth from a fraction of a meter to more than 10m, in size from a few square meters to more than 2 km2, and in shape from semi-round to elongated and sinuous as in the case of abandoned river channels (Figure2).

10

ALBEDO FLOODED 10% ~ River, Lokes, Tundro ond Seo ke

55% c::::=::J Ice flooting on Floodwoter

'0 HOODED 15% k' ~:"~n Tundro 75% _ loke. (tcc covered)

65'1'. c=:J Sea Ice

Figure 2, The Colville delta during breakup,

Landfonn Development in an Arctic Delta 163

6. Snow, Ice and Permafrost: Characteristics and Distribution

3.1. SNOW: SEASONALITY, PROPERTIES AND DISTRIBUTION

The most ephemeral of the trinity of snow, ice and pennafrost is snow. Nonetheless, it is very important in an arctic deltaic environment. It is the dominant surface material for most of the year (Figure 5), it is the source of most of the water that flows to the sea and it affects virtually all natural and human activities in the delta.

Figure 3. Low-centered ice wedge polygons.

164 H. Jesse Walker

Figure 4. Linear pond on top of ice wedge.

Although snow may fall during any month of the year, it is relatively rare in midsummer when temperature is at its maximum and in mid-winter when temperature is at its minimum. Whereas minimal amounts of snow are added to the pack during mid-winter, snow redistribution by wind is of common occurrence. Irregularities in the delta surface, especially sand dunes, along riverbanks and along pressure ridges, are the loci for the accumulation of snow drifts (Figure 6). In contrast, flat surfaces such as sand bars, mudflats and frozen water-bodies usually possess thin layers of snow. Occasionally, after wind stonns, parts of such flat surfaces may be swept clean of snow. Wind also packs the snow, helps increase its density and produces sastrugi.


3.2. ICE: TYPES, SEASONALITY AND DISTRIBUTION

Ice, unlike snow, can also fonn at the surface and underground. Like snow it is temperature dependent, although the actual temperature at which it fonns is affected by impurities in the water and by pressure.

In the case of the Colville River delta, ice is present during most (7 to 9 months) of the year on all surface water bodies as well as in the active layer. In the ground below the top of the pennafrost (usually the same as the base of the active layer in the delta) ice is perennial and occurs in a variety of fonns and sizes from ice wedges, which may grow to more than 3 m wide at the top, to pore ice of microscopic proportions.

Figure 5. Snow covered landscape.

166 H. Jesse Walker

Surface ice on the delta's distributaries, ponds and lakes and on the ocean, usually begins to develop in October and lasts until Mayor later. Even under the coldest winter conditions it rarely grows to be more than about 2 m thick. In water bodies less than about 2 m deep, ice may become bottom-fast whereas in deeper bodies it floats.

Figure 6. Pressure ridge at delta front with snow drifts.

3.3. PERMAFROST AND THE ACTIVE LAYER

The Colville River delta is within the zone of continuous permafrost which means that permafrost is present under all surfaces except beneath the deeper portions of the delta's


distributaries, the deeper lakes and under the sea out from the delta front but even there relict permafrost may be present. On the North Slope permafrost reaches depths of more than 650 m. Because the delta, like much of the North Slope, has many lakes that are more than 2 m deep, taliks beneath them are common. The delta is a water laden environment so most of the permafrost is of the moist variety; included ice comprises much of the total volume of the top several meters.

The active layer within the delta varies in thickness from only a few 1 O·of centimeters in areas covered by grass tussock vegetation to more than 2 m in nonvegetated sand dunes.

4. Snow, Ice and Permafrost: Examples ofImpacts and Interactions

4.1. SNOW

4.1.1. Snow: Impact on Albedo and Active Layer Development Although the total relief in the subaerial portions of the delta is less than 12 m, localized relief, especially in sand dunes provides an irregular surface. The entire surface, with rare exception is snow covered during most of the year. Once the sun rises in late winter, the albedo is high because of the reflective nature of the snow cover. Its decline begins with snowmelt and at anyone point in the delta may be rapid. It can drop from more than 80% over a clean snow surface to less than 20% within a few days (peake and Walker 1976). For the delta as a whole the decline takes much longer. The first locations showing reduced albedo are dune tops and river bars and flats. In terms of the delta as a whole the most rapid reductions occur just prior to and during the period of river breakup flooding (Figure 2) when snow melt is rapid. The relatively warm river water causes air temperatures over the delta (and therefore snow-melt rates) to be higher than over surrounding tundra surfaces, increasing the rate of snow melt. Further, at highest flows as much as two-thirds of the delta is flooded (Figure 2). Nonetheless, because of remnant snow drifts in some areas and longlasting lake ice cover on large lakes, it takes "about two months for the albedo decline to occur over the entire delta.

By September, the active layer reaches its maximum thickness and begins to freeze back up from below (see Figure 23). Soon thereafter air temperatures dip below ODe and the active layer begins to freeze from the top. About the same time snow begins to accumulate and, by providing insulation for the ground surface beneath, reduces the rate of active layer freezeup. During the snow-melt period in spring, the reverse is true. As long as a snow cover remains, active layer development is delayed (Figure 7). Thus, large drifts against riverbanks and dunes often delay thaw in some protected locations until as late as July.

4.1.2. Snow: Ice Thickness and River Ice Depression Once ice has begun to form on the water bodies of the delta snow accumulates on it just as it does on other surfaces. Because snow is an insulator, the timing of its occurrence and its depth will affect the rate of growth ofice. If thick snow cover develops early in the winter, it reduces the rate of growth of ice and results in a thinner layer than might otherwise occur. Although river and lake ice varies in thickness from year to year, it normally develops to

168 H. Jesse Walker

about 2 m thick. The insulation properties of snow are also important over other surfaces in the delta. For example, ground squirrels which are year-round residents, spend the winter in the sand dunes that are insulated from atmospheric temperatures to some extent by snow drifts.

o 2S 12

o 26 12

Figure 7. Snow cover and active layer development at Putu.

Snow drifts, which are often as much as 8 m deep along some of the delta's riverbanks, cause river ice to be depressed because of their weight (Figure 8). In addition, in the process of depressing the river ice, cracks occur allowing water to be forced to the surface and to freeze at the base of the snow pack.

4.2. ICE

In the delta, as noted above, ice is almost ubiquitous during the long winter. It is present on lakes, rivers, the sea and in the ground. Of these many forms ofice, river ice is the most important within the delta proper because during breakup it becomes an erosional, transportational and depositional agent. In the subaqueous portion of the delta, sea ice serves the same purposes. Lake ice, on the other hand, tends to be more passive although, in large lakes, under the influence of wind (Figure 9), it can also serve as an erosional agent. Ground ice, including ice wedges, becomes quite significant as thermokarst situations develop in the delta.


4.2.1. River Ice: Breakup Flooding, Erosion and Deposition River ice, which grows to thicknesses of2 m, becomes bottomfast in most of the delta's distributaries. Only in the deeper portions of channels, especially the main and Nechelic Channels, does unfrozen waterremain through the winter. Even in these channels, however, the margins of the ice cover are bottomfast. Because the thalweg portion of the channels remain unfrozen, seawater penetrates upstream beneath the ice. In the case of the Colville seawater reaches more than 60 km inland (Figure 10).

Such is the situation until snowmelt water begins to flow from the tundra to the river channels. It first begins to accumulate on the river ice but soon flows downstream on top of the river ice. At some point the floating ice breaks loose from its bottomfast portion, rising and falling with floodwater (Figure 11). This river and under-ice flow lasts until ice breakup occurs. Sediment is deposited on top of the bottomfast ice as well as on adjacent river bars and mudflats. Once buoyancy overcomes the adhesion of bottomfast ice to the bed plus any added weight due to its load of deposited sediment, the bottomfast ice rises to the surface and, when there is ice movement downstream, joins the floating mass.

Station Locations

-4

-5

-6

o 100 200 300 meters

E:;::! Snow R River ice

6 Salinity %0

400 500

Figure 8. Depression of river ice on the Nechelic Channel.

Some of the ice that moves downstream during breakup picks up sediment (including at times large rocks) and vegetal material (such as willow trees as much as 3 m tall) and transports them downstream. Some of the ice floating in floodwater is carried onto

170 H. Jesse Walker

sandbars, mudflats and even the tundra surface, in places often becoming stranded (Figure 12). Once it melts, the material carried settles in place as erratics (Walker 1969).

Temo ·C Wind JUNE W A V

15 3 10 5.5 19 3 8 5.5 25 4 6 12.5 26 4 4 14.0 27 6 4 13.0

W- Water temp. ·C A -Air temp. · C

01 S N

NE NE NE

V -Wind velocity M.P.H. Ol- Wind direction

~11.5-Wlnd direction

Figure 9. Wind and lake-icc decay.

River

iIlllIIIIII Prebreokup ~Breokup

.040 30 20

_Londvoord 10

Figure 10. River and ocean salinity in the Colville delta.

o 500 meters ''--...L.-....I'

Ocean

10 20 30

Km Seoward-


m

Figure II . River cross-section during breakup.

During breakup, ice jams frequently occur, especially if breakup occurs on a falling stage. Such jams are accompanied by ice shove as the ice butts into or overtops riverbanks. Further, as the ice, much of it 1.5 to 2 m thick, rides over bars it leaves gouge marks in the surface (Figure 13).

The ice that becomes stranded melts in place. Because it is freshwater ice it produces "candles" that drop to the bar or flat beneath creating slight impressions (Figure 14). This is one of the many microfonns that are unique to arctic deltas.

4.2.2. Sea Ice and River Flooding During winter sea ice is bottomfast across the delta front except over the deeper channels. It is through these channels that tidal water penetrates upstream. During flooding river water with its heavy sediment load moves seaward through the river-mouth channels and over the sea ice. As the floodwater spreads out in front of the delta, sediment is deposited on top of the sea ice just as it is on bottomfast ice in the river channel (Walker 1974). In the case of the Colville, the floodwater progresses under as well as over the ice. The water flow beneath the ice produces a fresh-water wedge that expands rapidly once the floodwaters reach the sea (Figure 15). This wedge has been tracked to distances of more than 40 km from the front of the delta. Surface flow progresses to some 8-10 km from the delta front. This outer limit, which is arc-shaped (Figure 16), tends to follow pressure ridges in 3 to 6 m water depths. The added weight probably aids in the development of drainage loci (Reimnitz et 01. 1974) which often fonn circular eddies. The relative proportions of drainage through these localized whirlpools or strudel and along pressure ridges is unknown. In 1971, 140 strudel were counted with about two-thirds located off the

172 H. Jesse Walker

eastern channel mouths. Drainage patterns develop rapidly on the ice with many curving at the juncture of the whirlpool.

By the time the sea-ice surface is drained of its floodwater, there is left behind a covering of sediment and vegetal debris. In 1971 sediment in some locations reached thicknesses of more than 10 cm. The average, however, was less than 1 cm, with the thickest amounts being off the east channel.

Once floodwaters reach the ocean, thermal erosion and melting begin to destroy the sea ice. The sediment on the sea ice, with its reduced albedo, aids in increasing the melting rate of the ice beneath it. In the case of the Colville most of the nearshore ice melts in place dropping its sediment load to the floor below. It appears that little if any of it is transported seaward to be deposited elsewhere (Walker 1974).

4.2.3. Sea Ice: The Subaqueous Delta As noted above, the sea just off the delta intercepts sediment during river flooding only to deposit it on the sea bottom as the ice melts. The bottom also has numerous pits that result from surface drainage-pits caused by strudel scour. These forms are associated with breakup flooding. Other forms that develop on the sea floor are the result of sea ice movement. During the ice-free portion of the year sea-ice drift in the Beaufort Sea creates


gouges as ice floes move along the bottom. These gouges are numerous in the deeper portions of the subaqueous delta.

Figure 13. Gouge mark from river ice on gravel bar.

4.2.4. Ice Wedges: Inversion of Relief and Lake Tapping

Large lakes are numerous in the Colville delta. Twenty-four have a surface area of more than 1 km2• Most of the shoreline of these lakes, except those portions bordered by sand dunes, is serrated reflecting the type of erosion that occurs along ice-wedge polygons. The enlargement that occurs is the result mainly of wave action in summer and, to some extent, ice shove during breakup. These active processes are accompanied by the thaw of ice wedges which accounts for their irregular shoreline. As the deeper lakes expand, abrasion platforms develop. These platforms often possess relatively deep, submerged troughs which represent an inversion of relief (Figure 17) due to the melting of the ice wedges that formerly bordered the lake.

174 H. Jesse Walker

Lake expansion also plays a role in a process known as lake tapping (Walker 1978). Much more important, however, is the migration of the delta's distributaries in the direction of a lake (Figure 18). As the tundra bands between the lake and the channel become narrow, ice wedges in the cut bank thaw faster than the surrounding permafrost material and eventually provide a conduit for drainage. Because the lakes are 2+ m higher than the river channels, drainage is rapid and the trench left by the melting ice wedges rapidly enlarges. Once tapping has occurred however, lakes are subject to flooding by the river and deposition within the lake begins. Further, some of the former lake bed becomes subaerial and subject to the initiation of new ice-wedge growth.

Figure 14. "Candle" impression in mud after river ice melt.

Landfonn Development in an Arctic Delta

A. Floodwater- Seawater Interface Advance. 3-14 June 1973

o

Jun ••

J..,ne f:·:·:.:,,:::::.1 9 Nne Bl

___ 1.4 )un.~

10 IS 20 15 30

10

B. Salinity - Station 73 c . Temperature - Station 73

Oy-----------------~

I i

ic.

, ! \ Ii I ( .. ~ • \ '.- . -.-.-.-. - :::::.::,<~ ~ , o ----___ _

10 20

Salini Iy· 0/00

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-!. 0. . o

6

Figure 15. Flood water/sea water interface during breakup.

Figure 16. Flooding on top of sea ice during breakup.

\ ,; \ ! \ \..... . ....... . \, ._.-.-.-.-._.-.,

.1

......_ 1

- ......... i '\.i

\i

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__ • SurY.y linn

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175

176 H. Jesse Walker

Figure 17. Inversion of relief in an enlarging lake.


Because riverbanks are prime locations for the development of snow drifts, floodwater first undercuts these drifts before reaching the frozen banks where it begins to fonn a thennoerosional niche (Walker and Arnborg 1966). Niches, which may be more than 8 m deep, undercut both pennafrost and included ice wedges (Figure 19).

Figure 18. Serrated lake and river banks. The river (right) tapped the lake in 1972.

178 H. Jesse Walker

4.2.5 . Ice Wedges and Bank Erosion In the Colville River delta, ice wedges vary greatly in location, distribution and size. They are best developed in zones of dense peat, such as that present along the right bank of the main channel. Bank erosion is initiated with the flooding that precedes river-ice breakup and continues through post-breakup flooding. Only rarely during summer do rain-generated floods impact riverbanks.

Figure 19. Double thennoerosional niche under ice wedge and pennafrost bank.


What happens to the bank depends to a large extent on the material in the bank. Peat banks tend to be more resistant to the floodwaters so that ice wedges are thawed back leaving peat blocks isolated. In more mineral-laden banks, such as the Gubik Fonnation along the left bank of the Nechelic Channel, block collapse frequently occurs. Usually, this collapse is along ice wedges which are zones of weakness in comparison to the surrounding pennafrost (Figure 20). The melting of ice wedges, once collapse has occurred, is rapid leaving behind the slumping blocks as the pennafrost thaws (Figure 21).

Figure 20. Collapsed blocks on Nechelic Channel.

180

m ~ I .. Wod"o

o Snow

Orgonlc Mat

B Permalrost

SIumpod

Malerial

H. Jesse Walker

Figure 21. Undercutting, block collapse and bank retreat along the Nechelic Channel

4.3. PERMAFROST

4.3.1. Permafrost and Perched Lakes A perched lake is one" ... which has a surface level lying at a considerably high elevation than that of other water bodies directly adjacent or closely associated" (Veatch and Humphreys 1964). Lake perching is only possible where an aquaclude exists between the lake's water table and the regional water table. In the case of perched lakes in the Arctic, the aquaclude, which is temperature dependent, conforms to the top level of the permafrost. In the Colville River delta perching is best developed in sand dunes many of which have a local relief of more than 5 m. Both active and stabilized dunes are present. Stabilized dunes, which are vegetated and generally smooth, have relatively few depressions within which water accumulates. Active dunes, on the other hand, provide two types of closed basin; one, between dune bands, the other, blowouts within dunes (Walker and Harris 1975). Most dune systems within the delta possess both types of basin.

Landfonn Development in an Arctic Delta

o 5 10

Vert. Ex. 5x

15 20 meters

25 30 35

Figure 22. Putu pond, a perched lake near the head of the Colville River delta.

a

'" A 0

'" " Pond B

~:..Y'Bottom 25 ':. I 25 ~"

'\. .. '\ i ,'. i \ \~" i e- 50 \ \~

50 , \

~loWWiIIOW o!!- \ i

== \ LowWll ow \

~ 75 1H~ 75 a i/Sand \ ~ «W1Jow \ ~ . ~HighWiIlO"

" \ ,: ..... ~ 100 \. I """"""'Moss 100

\ ' =-:-'-::'~:;-MoS$ ". \ Pond';" ". \. .... .. ..

\-. Sand ...... \. Bottom ". ". '. 125 f--- 125 ,

'''''", " \

" '. .-.-.- v ' ... /' 150 150 ...

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May June July AUg. Sept. May June July AUg. Sept.

Figure 23. Active layer development under different surface materials in Putu pond.

181

182 H. Jesse Walker

An example of a perched lake is Putu pond, located near the head of the delta. The basin, 2060 m2 in area, is separated by only 15 m of dune sand from a mudflat that lies more than 6 m below the lowest edge of the basin (Figure 22). The basin has a variety of vegetation types including willows as much as 2 m tall. During winter snow accumulates to more than 2 m in depth, some of which does not melt until early June. As the snow melts, the active layer begins to develop moving the aquaclude down into the dune. Snow melt results in overflow which in turn causes erosion in the dune. The eroded material builds a fan on the mudflat at the base of the dune.

Active layer development progresses at different rates and reaches different thicknesses beneath the different surfaces within the perched pond basin (Figure 23). Beneath exposed sand surfaces it reaches depths of more than 150 cm whereas under willows it develops to only about one-half that amount. By the end of the thaw season the pond, which, at frrst, lost snow-melt water via the drainage channel, remains perched within the thaw bulb created above the permafrost table. Although many factors are involved in pond perching within the delta, the most important is permafrost.

This pond, like other forms in the delta, must be considered ephemeral. The overflow channel is generally cutting into the dune sand, while the pond is filling with wind blown sand. Meanwhile decaying. vegetation is accumulating at the bottom of the pond.

5. Conclusion

The various processes and impacts discussed above have been grouped under one or the other of the triad of snow, ice and permafrost. The objective was to pick the one that seems to be the most important in shaping the particular event described. However, it is recognized that in virtually every case other factors or conditions are also important; factors such as river, air and soil temperature, river discharge and wind. The combination of all of the agents and processes operating in an arctic delta helps account for the great variety of fonns present and for the rapidity of change that occurs.

References

Amborg, L., Walker, H. J. and Peippo, J. (1966) Water discharge in the Colville River, 1962, Geografuka Annaler,48A,195-210.

Dawson, A. (1975) Landforms of the Colville River Delta, Alaska, as Interpreted from Aerial Photographs, Master's Thesis, Louisiana State University, Baton Rouge, LA.

Peake, J. S. and Walker, H. J. (1976) Albedo decline in an arctic delta during spring snowmelt, Proceedings, Association of American Geographers, 8, 8-12.

Reimnitz, E., Rodrick, C. A. and Wolf, S. C. (1974) Strudel scour: A unique arctic marine geologic phenomenon, Journal of Sedimentary Petrology, 44(2), 409-420.

Ritchie, W. and Walker, H. J. (1974) Riverbank forms of the Colville River delta, In: Reed, J. E. and Sater, J. C. (eds.) The Coast and Shelf of the Beaufort Sea, The Arctic Institute ofNorth America, Arlington, VA, 545-562.

Veatch, J. O. and Humphreys, C. R. (1964) Lake Terminology, Water Bulletin No. 14, Michigan State University, East Lansing, Michigan.

Walker, H. J. (1983) The Colville River Delta, Seciton G: The Delta's Distributaries, Report prepared for the North Slope Borough, 41 pp.


Walker, H. J. (1978) Lake tapping in the Colville River delta, Alaska, Third International Conference on Permafrost, Proceedings, I, 233-238.

Walker, H. J. (1974) The Colville River and the Beaufort Sea: some interactions, in Reed, J. E. and Sater, J. C. (eds.) The Coast and Shelf of the Beaufort Sea, The Arctic Institute of North America, Arlington, VA. 513-540.

Walker, H. J. (1969) Some aspects of erosion and sedimentation in an arctic delta during breakup, Actes due Col/oque de Bucharest, Association Intemationale d'Hydrologie Scientifique, 209-219.

Walker, H. J. and Amborg, L. (1966) Permafrost and ice-wedge effect on riverbank erosion, First International PermafrostCo1iference, Proceedings, Natiorial Academy of Science, Washington, D.C., PubI.1287,164-171.

Walker, H. J. and Harris, M. K. (1975) Perched ponds: an arctic variety, Arctic, 29(4):223-238.

THE SEARCH FOR AN ARCTIC COASTAL KARREN MODEL IN NORWAY AND SPITZBERGEN

Joyce Lundberg Department o/Geography Carleton University Ottawa, Canada K1S 5B6

Stein-Erik Lauritzen Geological Institute University 0/ Bergen Allegaten 41 Bergen,JVorw~ JV-5007

Abstract

A search was conducted, in Norway and Svalbard, for an assemblage of littoral karren features characteristic of cold regions, comparable to other coastal karsts of the world, that could be fitted into a general coastal karst model. In fact, these karren do not fit into a world model oflittorallcoastal karren The search showed that they represent a landform in transition rather than a morphology in eqUilibrium with modem processes. The suite of mainly bowl- or basin-shaped karren forms is best developed in the supra-littoral zone, rather than in the inter-tidal zone as occurs in other coastal karsts of the world. Further the forms are often simply an overprint on a glacial legacy; and they are governed more by isostatic uplift than by inter-tidal and coastal processes. The form will not remain constant over time but will vary as the rate of isostatic uplift decreases. In this chapter, the forms are documented, and the dominant processes are discussed (including a surprisingly high input from salt weathering arid a low input from direct biological weathering). They are then compared with coastal karsts from other parts of the world to demonstrate their similarities (a common dominance of roughly circular negative forms and an absence of joint control), and their differences (almost negligible erosion in the inter-tidal zones here compared with maximum low- to inter-tidal erosion in temperate coasts).

185 K. Hewitt et al. (eds.), Landscapes a/Transition, 185-203. © 2002 Kluwer Academic Publishers.

186 Joyce Lundberg and Stein-Erik Lauritzen

1. Introduction

Coastal or littoral karren is an assemblage of small- (down to millimetres) to large-scale (up to several metres) dissolutional features developed in the eu-littoral (inter-tidal, supra-tidal) and the supra-littoral zones of carbonate coasts (Trenhaile 1987; Ford and Williams 1989). The term "coastal karst" is also often used but this can mean the large-scale hydrogeological karstification of coastal regions (eg, Back et 01. 1986; Whitaker and Smart 1994). The terms "phytokarst" or "phytokarren" or "biokarst" have been applied to coastal karren features (eg, Folk et 01. 1973; Schneider and Torunski 1983; Viles 1984; Jones 1989) but these process-indicative terms are limited since much of the geomorphic work may not be caused directly by living things.

A review of the literature teveals that coastal karren is well developed both in temperate and tropical coasts. Characteristic forms can be related to tide levels and to exposure (see below). However, we know of only a few papers describing littoral-coastal forms of a cold region: Moe and Johannessen (1980) and Holbye (1989) describe features from northern Norway at 67.2"N; Malis and Ford (199S) and Malis (1997) describe features from Newfoundland, Canada at SO-S2"N. In response to this gap we undertook to survey a selection of the carbonate outcrops along the coast of Norway and in addition to look for coastal karren in two carbonate sites in Svalbard (Figure 1).

The aim of this research was to document the characteristic landform assemblage (morphology and suite offorms) on karren of cold region coasts and to add this to a general model of coastal karren development. Two fundamental questions in geomorphology were asked: (i) is the landform in equilibrium with the conditions in operation today? and (ii) does it have a characteristic form that is self-sustaining. Most coasts have had a relatively short time to develop an equilibrium with present processes, sea level having reached its present height only about 6,000 years ago (Fairbanks 1989; Bard et al. 1990). The approach to equilibrium will depend on the processes in operation, the rates of development of the features, and the rock types. Carbonates in coastal settings, have high denudation rates (typically -1 rom a') but up to -4 mm a· l : Trudgill 1983; Viles and Trudgilll984; Spencer 1985a,b; Trudgill and Crabtree 1987). Therefore they have a high chance of developing a characteristic form quickly. Indeed carbonates in cold regions might theoretically have an even better chance than in warm regions of dissolving rapidly because of the greater solubility of CO2 in cold water (Alexandersson 1976).

Normal coastal processes include abrasion, attrition, hydraulic action, wave action/quarrying, corrosion or water-layer weathering, dissolution of soluble components, and bio-erosion. In cold regions, frost action, focused mainly on the upper inter-tidal and splash zones, must also be considered, although fine-grained carbonate rocks are not very susceptible to it (Trenhaile 1987). Winter shore ice often protects the coast from wave action, though Nielsen (1979) observed significant rock destruction when ice breaks off in spring. For karren development, mechanical action must be limited and dissolution and/or bio-erosion must predominate. Holbye (1989) found little mechanical action but also little direct bio-erosion in the supra-littoral bowls of Northern Norway. Moe and Johannessen (1980) did fmd evidence, in their very sheltered site close to the same area, of direct bio-erosion in the form of cavities inhabited by Littorina snails. Malis (1997)

The Search for an Arctic Coastal Karren Model 187

observed much dissolution, but little direct bio-erosion, and little mechanical action in the inter- and supra-tidal karren of Newfoundland.

N

i

.. ............. ,'

\ 00·-\ 60·

I)

I 50 km I

Figure 1. The field sites in Norway were from three general locations: -65.2° to 65. 7°N: Site A - the eastern extremity of the island of Vega, a relatively exposed situation; Site B - Bnmngysund on the mainland to the southeast of Vega, a moderately exposed situation; Site C - Vennesund, on the protected side of Bindalstjorden about SO km south of Vega, a sheltered site; site D - Ssroffersgy, an island - 40 km north of Vega, a very sheltered location although the island is exposed to the open ocean; -67.0° to 67.6°N: Site E - I1stad, to the east of the town of Bods, facing the ocean but IS km inside Skjerstadtjorden; Site F - a couple of small sites on the large island of Sandhorngy, 20 km to the south west of Bods, in particular Mimes, an exposed site in the north of Sandhorngy; Site G - Nord Arngy, an exposed island offshore of Sandhornsy; - 68.soN: Site H - Evenes and GAsbakken, 20 km west of the town ofNarvik, both on the northern shore ofOfottjorden, but relatively open to the prevailing wind and waves. The sites in Svalbard (inset) were Site J - Trolloson Point, a very exposed site just south of Hornsund, at -77.soN; and Site KBlomstrand Island just north of Ny Alesund, a sheltered site inside Kongstjorden at -78°N.


The west coast of Norway is an emergent coast made up of resistant rocks, and subject to generally westerly storm wave action. Isostatic uplift has raised relative land level throughout the Holocene and is still ongoing. Modelling of land uplift rates for the relevant areas gave 2-3 mm a-I (Gr0D.lie 1981; Meller 1989; Fjeldskaar 1997). If the rate of eustatic sea level rise (1.15 mm a-I: Fjeldskaar 1997) is included, then the rate of coastal emergence is 1-2 mm per year. The isostatic uplift means (a) that coastal processes have not operated for long on the recently emergent rock closest to modem sea level; and (b) that each outcrop has gone through each area of the littoral, leading to inherited forms.

Glaciation has left a legacy on the coastal rocks. Sea level during glaciation was considerably lower than today, and ice scour or sub-glacial fluvial scour affected land below modem sea level. In many rock types, particularly the insoluble rocks, the inter-tidal zone is unmodified by modem processes: rocks show pristine striae, p-forms, and whalebacks. The glacial legacy applies to the carbonate outcrops but these generally do show modifications to varying degrees related to littoral dissolution.

The strandflat (Holtedahl 1998) is an uneven and partly submerged rock platform extending seawards from the coastal mountains, made up of numerous low rock skerries, shallow sea areas, sometimes with sharp breaks of slope to inselberg-type steep hills. It is 50-60 km wide in mid-Norway. It occurs only in areas with a history of glaciation, and is the result of long-term polygenetic and/or polycyclical erosion over the whole glacial period since at least the Tertiary. The effect of the strandflat on coastal karst is that exposure levels to high energy storms are not as great as might be expected for an open Atlantic coast.

The carbonate rock type is coarsely crystalline calcite marble (Lauritzen 1984) with varying amounts of impurities in the form of narrow, elongate mica schist bands disseminated to a greater or lesser degree throughout the marble. Bedding planes are absent. The coastal outcrops are a continuation of the stripe karst of inland Northern Norway. This outcrop pattern results in rather limited coastal exposures and in places very obvious structural alignments of the micaceous impurities.

2. Methods

In order to model karren development in different situations, sites were sought from a range of exposure levels; however. very exposed sites are rare. The field sites are shown in Figure 1. The tidal range is approximately 1.7 m (Holbye 1989) but the effect of wetting is extended greatly by supra-tidal swash and splash, and supra-tidal or supra-littoral spray. Only a few of the sites were of sufficiently pure carbonate to show any significant karren development; where micaceous impurities are dominant dissolution simply removes the carbonate laminae leaving the insolubles emergent.

The research method was to survey the coastal karren from low water mark up to the supra-littoral:terrestrial transition zone. The surveys documented the sizes and shapes of karren features and the processes that appeared to be involved, focussing on any biological action involved, since bio-erosion is a conspicuous feature of most other coastal karsts.

The data analysis included a comparison of all the sites to see if they conform to a pattern, and a comparison of these cold region sites with other coastal karsts of the world,


in particular with the coastal karren of the west of Ireland, which is one of the best documented of the temperate coastal karsts (Lundberg 1974, 1977; Trudgill1987; Trudgill and Crabtree 1987; Trudgill et al. 1987).

(a)


Figure 2: The coast at I1stad. a) glacial p-fonns in the supraJittoral zone being modified into dissolutional potholes; b) intertidal and supratidal etching, the large-scale fonn is the glacial inheritance.

3. Results

Of the many sites documented, two were chosen as type examples and are described in detail below. In general, karren forms were poorly developed, especially in the inter-tidal zone. The more sheltered sites still showed the glacial legacy clearly; eg, whalebacks and p-forms in pure white marbles at Smoffers0}' are very nicely preserved, with only very minor pitting (sharp-edged, semi-spherical pits 1-2 cm wide) in the depressions and striae.

The Ilstad site is a little more exposed and shows greater modification of the glacial inheritance. In the supra-littoral the intricate meandering p-forms are deepened into dissolutional potholes (Figure 2a). Closer to high tide level the smoothed glacial surface is almost obscured by overall surface pitting. These pits have no clear organisation and vary in size from shallow depressions 10-50 cm wide and bounded by micaceous laminae, to 2-4 cm wide semi-spherical pits with sharp edges covering all surfaces, including the larger depressions. The p-forms have sometimes been deepened into potholes 30-50 cm wide and deep. In the inter-tidal zone dissolution has etched out the structural features of the rock and only the larger glacial forms are apparent (Figure 2b).

The Evenes site shows the differential effect of exposure of the headland. The sheltered side shows virtually no denudation of the fme p-forms in the inter-tidal zone (a carving on the rock from 1866, 131 years ago, is still pristine) while the supra-littoral basins are up to 10 cm deep. However, supra-littoral bowls on the windy side are up to 50 cm deep. This site was not surveyed in detail because structural control by the parallel, steeply dipping, micaceous partings dominates the form.


The Mrunes site is one of the most exposed and shows only a trace of the glacial legacy, at least in the upper littoral where greater post-emergence denudation has occurred. This site was surveyed in detail. The marble is sugary in texture, with few insoluble partings. This karren assemblage shows no real zonation of forms, only a zonation of size and complexity. The sequence from low tide up to the supra-littoral is shown in Figure 3. The rock of the inter-tidal zone (Figure 4a), is covered with a dense mat offucoid algae, barnacles, mussels, and green algae. This mat protects the rock from differential wetting and erosion but tiny cavities at a scale of 1-2 mm occur (see Moe and Johannessen 1980). Pitting at larger and various scales is apparent in the upper zones. Towards high water mark the rock has a barnacle cover and shows poorly organised undulations at a scale of 15-30 cm, some of which hold water.

At high water mark and above it the shallow rough basins start to form coherent units with a hierarchy of pits. The basins are 5-15 cm wide and 2-8 cm deep; the surfaces of these basins are eaten into by semi-spherical pits 1-2 cm wide, 1-1.5 cm deep similar to Hiatella arctica borings in Newfoundland (Brookes and Stevens 1985) but in this case probably not from such a cause. Endolithic and epilithic colonisation is apparent in the brown coloration inside the basins. Cyanophyte colonisation and dissolutional etching continue inside pits and basins, but abrasion/fracture marks are common on upstanding basin rims contributing to the less pitted and more rounded nature of the rims.

The supra-littoral shows the best developed and largest bowls (Figure 4b). These are also made up of hierarchies of pits, with the bowls typically around 30 cm in diameter and 20 cm in depth. Many of these show considerable salt weathering with large salt crystals in the remnant pools. The biggest of the supra-littoral bowls are structurally controlled, dissolutional etching between insoluble quartz veins giving an elongate shape. Many basins have itinerant rounded beach boulders which have been thrown up during storms causing some fracture of upstanding rims but probably having no further geomorphic effect (ie, these boulders do not act as rock mills). Holbye (1989) described supra-littoral bowls like these have been described from nearby Nord Al1l.0Y as having a strict hierarchy of four bowl/pitsizes (2-6 cm, 10-15 cm, 25-40 cm, 100-300 cm). This is a very useful conceptual tool which can be applied also to the Mrunes site, although much more loosely. Arguably this concept can be applied to littoral karren world-wide since all pitted, scoriated surfaces seem to have something of a pit-in-pit fractal nature (Torunski 1979).

The Gasbakken site has pure marble of sugary texture with few impurities and few joints. A continuous single bed sloping from the supra-littoral down to below water level was surveyed (Figure 5). It creates the ideal situation to examine zonation of forms in relation to elevation without any interference from geological variation. The glacial legacy of striae and p-forms is well preserved at, and just above, low water. Again the greatest development ofkarren is in the supra-littoral zone, where there is also zonation mainly of size rather than form. The variety of bowl sizes here is more restricted and related to the topographic slope. Two features, which do occur in other sites, are particularly well displayed here. They are elevated rims of pioneer pits in the mid- to high tide region, and grusification in the supra-littoral zone.


I Zone D: Bowls degenerating White,grusified surface

L--.J

lOem

I Zone C: Basinslbowls

L--.J

10 em MiclO-ftettiog allover

t

I Brownlpink swfac:e

I Zone B: Rudimentary depressions I '-------', ~\ .... 10 em Fmcture scar I

GIey swfac:e on tops BIack-bIOWD swfac:e inside pits

I Zone A: Undulations I Small-scale ftettiog '--------' Waler4iDed deplCSSiOIlS ~ 'I~ 10 em

1 6 II s ; 4

j 3

l 2 i I

I 0 :2 o 10

I Mimes

15 20 25 30

Figure 3: The profile and sequence offorms from low tide up to the supralittoral at Mimes.

The sequence ofkarren development is as follows (Figures 5 and 6a). Smooth rock emerges from the water, coloured brownish-yellow from epi- and/or endolithic cyanobacteria and algae. Just above mid-tide level a threshold is reached with pitting above and virtually no pitting below (Figure 6b). The non-pitted, apparently smooth surface is actually very shallowly fretted allover. The isolated pits which breach this surface are water-filled and show edges raised above the general surface. This could either be caused by deposition of, presumably, calcite on the rims through evaporative wicking from the wet pit at low tide, or by protection of the rims from the fine fretting of the rest of the surface. Once the surface is breached, then pitting becomes general, leaving the surface rock rough and sharply fretted to a depth of about 5 mm.


Figure 4: M4mes. a) Poorly developed karren in the inter-tidal zone (right) and supra-tidal zone (left); b) Bird's eye view of supra-littoral bowls, characterised by hierarchies of pits and basins, colonised by brown and pink endolithic cyanobacteria and algae, and coated with dried salt crystals (the pencil is IS cm long).

Blomstrand Island has moderately pure marbles, but is in a relatively sheltered location and the fresher water freezes easily. The coastal profile is steep but not cliffed. Rock surfaces up to 5 m ASL are frost shattered and swept clean of debris by wave action. Above this, well into the supra-littoral, dissolution produces micro-fretting a couple ofmm deep. By 8 m ASL dissolutional cups around 20 mm wide, 10 mm deep appear on sheltered surfaces (Figure 7); other surfaces are generally frost shattered.

Both of the Svalbard sites demonstrate that dissolutional karren do form in this very high latitude coastal site, but they are apparent only where frost action is limited.

4. Discussion

4.1 . CHARACTERISTICS OF NORWEGIAN LITIORAL KARREN

The karren are basically superficial modifications of glacially scoured surfaces being raised above modem sea level by isostatic uplift. The following generalisations can be made.


I Zone D: Bowls degenerating

10cm

I Zone C: Basinslbowls ..........

White surface ~ )

10 em Salt etystals

Fracture scar ~ \ ~ Pitting on baekwall

Black-brown surface

I Zone B: Rudimentary depressions I _ ./ ~ r~ 10 em White surface on tops '"

Blaek-brown surface inside pits

Rimmed pits . Unpitted surface ~ ~ ~

r------------------------, ~: -- ~ -I Zone A: Smooth rock plus pits I Small-scale fretting

10em

I Glsbakken I

o S 10 15 20 2S 30 3S Metres from low water mark

Figure 5: The profile down a single bed of pure marble at Omfllget, Glsbakken.


Figure 6. Omflaget, Gbbakken. a) Supra-littoral bowls in foreground, non-pitted supra-tidal surface and algal mat oflow-tidal zone in background; b) Pitting at mid-tide level on the pure marbles at OmfUget, Gasbakken. Isolated pits penetrating through the smooth surface show raised rims. Once the surface is breached, pitting is general and of unremarkable form; c) Supra-littoral basins at Gbbakken, showing the effect of slope (shallow lips and deep backwalls), secondary 1-2 cm deep semi-spherical pits over all surfaces, and rounded basin rims (notebook is 15 cm long).


The form: The karren fonn is always some kind of pitlbasinlbowl, always round bottomed and with no sharp breaks of slope to the side walls. They could never be described as panshaped. The basins vary mainly in size and width:depth ratio. Smaller pits have a ratio closer to 1 and larger basins closer to 6. In each locality the depressions confonn roughly to a two- or three-level hierarchy of pit-in-pit morphology (see Holbye 1989). However, there is no evidence of a real separation of the three populations, suggesting the sizes fit into a continuum. We can only say that the smaller pits are approximately 1-3 cm wide and deep, the middle-sized ones roughly 5-12 cm wide and 3-6 cm deep, the large ones roughly 10-50 cm wide and 5-12 cm deep. We found little evidence that the largest of Holbye's categories - the 1-3 m wide shallow depressions ~ are truly karren forms, rather than a glacial legacy or structurally controlled.

The distribution: Where pitting is just beginning, pits are isolated and their distribution related to depressions, striae, or joints. Otherwise the pits and basins cover all surfaces and fill the space completely in what appears to be a space-competitive manner giving a roughly polygonal network of basin rims. Where impurities are present the distribution is more linear and bowl rims are coincident with micaceous partings.

The zonation: The karren are virtually non-existent in the inter-tidal zone, appear in the supra-tidal, and increase in size with elevation, reaching their maximum in the supralittoral before destruction by terrestrial weathering processes. Their zonation is only of size and not of basic fonn.

The history: Many sites show smooth surfaces emerging from the inter-tidal zone and pitting being initiated only above this, after which the forms are inherited from below.

4.2. MECHANISM OF FORMATION

While this research was not designed to test hypotheses about geomorphic processes, the observations lead to several conclusions and speculations about mechanisms of formation.

Bio-erosion and dissolution: 1. Cyanobacterial colonisation is evident in all but the highest bowls. It is likely that bioerosion on a microscopic scale is active in all pits and basins. The macro-scale bioeroders which produce identifiable borings or home scars are absent. Thus bioerosion is active, but less significant than in temperate or tropical sites. 2. As in all situations where bodies of water are isolated in basins, respiratory CO2

production at night, concurrent with low tide and frequent replacement of the water, will cause bio-chemically enhanced dissolution. This is most likely to occur in the lower zones. However, in Norway the rapid isostatic uplift rate prevents the fonnation of many basins in lower zones so this process must be minor. 3. The other most common situation for rapid dissolution is the supra-tidal splash zone where high fluid velocity enhances dissolution rates, the pattern observed in Newfoundland by Malis (1997). In Norway this probably explains the initiation and


expansion of the basin forms in the supra-tidal zone. However, the uplift of this zone out of the influence of the splash limits the extent to which this process can act. 4. Dissolution is probably expressed most strongly where fresh rain water can mix with salt spray in the supra-tidal to supra-littoral zones. 5. Turbulent eddies rarely have any effect on coastal karren, except where the water carries significant abrasive material, and the forms cannot be called "karren". Hence, the mechanism proposed by Holbye (1989), involving inheritance of the effects of turbulence from the time when the rock was in the inter-tidal zone, is unlikely to exist. 6. The regular distribution of both the small pits and the larger bowls cannot be explained by boring organisms, whereas in other karren sites, many pits can be directly related to the colonising organisms such as the characteristic home scars of the echinoid Paracentrotus lividus in Ireland. The roughly equidistant distribution of the pits and basins might be the result of competition for dampness, versus evaporation and capillary action. Moe and Johannessen (1980) explained the 20-30 mm wide pits above high tide level by weathering reminants of the 1-2 mm wide Littorina pits of the intertidal zone. However, the concept of the inheritance of the locus of pits from intertidal processes, does not account for the changing basin size and density and the overall greater surface denudation of the upper zones.

Mechanical action: The inter-bowl rims are upstanding, dry out quickly, and are open to mechanical action from various sources. They show little biological colonisation or biochemical activity and only minor dissolution. 1. Abrasion and fracture are displayed in the frequency of many small fresh-rock surfaces. Projectiles carried in the sea water and thrown up by storm waves cause fractures and produce rock shards and fragments. Sand-sized projectiles carried in the wind cause abrasion. 2. Spalling and grusification are displayed in the rounded shapes of rims, the exposure of the grain texture, and the build up of loose grains. Frost action probably causes shattering and spalling, although modelling by Trenhaile and Mercan (1984) suggests some problems in freezing inter-tidal and supra-tidal rock. Salt action causes spalling and grusification. Mottershead and Pye (1994), studying coastal tafoni in southern England, found that salt weathering results from a combination of dry deposition of salts and wetting by coastal aerosol.

4.3. COMPARISON WITH OTHER REGIONS AND IRELAND IN PARTICULAR

Coastal karren are usually delicately etched, yet spectacularly jagged. They usually show a hierarchy of pitlbasin sizes up to several metres in scale. Unlike terrestrial karren, hydraulic gradient and gravity are not significant controls on the form. Littoral zone processes do not preferentially etch joints and bedding planes. The characteristic form is some type of round basin separated by intervening positive forms of varying relative relief.


They are controlled both on the large and small scales by wetting regime, biological activity, and energy levels. The overall form, or the way in which the components are arranged, will be modified according to organisms present (which may cause, or protect from, erosion); exposure levels (wave energy, wind energy, wave, and wind orientation); lithological variations (chemistry, crystallography, depositional fabric); structural variations (sedimentary structure, bed thickness, orientation, dip, joint frequency, and orientation); tidal range and regime; input of fresh rain-water. Tropical karren, because of the effect of encrusters, show some unique inter-tidal features (see Focke 1978). Otherwise the erosional ramp is the typical coastal/littoral karren form that occurs world-wide. The karren show downward inheritance of forms, in which the denudation of the upper zones will eventually transform each surface into the form seaward of it.

Figure 7: Karren at Blomstrand Island. Svalbard. on a sheltered surface in the supra-littoral zone.

The temperate form from the west of Ireland (Figure 8)rnight be expected to be comparable to the cold regions form, perhaps modified by frost/ice action. Characteristically it displays the erosional ramp with basins and pinnacles but with an


obvious zonation of basin shape and size. The greatest relief is in the low- to mid-intertidal zone. The zonation is as follows: a) The supra-littoral/spray zone is characterised by perfectly flat-bottomed, steep-sided,

isolated pans,-20-30 cm wide and 10 cm deep, cut into the rock ''plateau'' which is scoriated all over with small-scale fretting. Lichens and cyanobacteria are obvious, and some minor Littorina colonisation, but the rock is relatively bare ofliving things. Salt-fresh water mixing is an important process in this zone. The pans grade downshore to shallow, roughly rounded basins.

b) The upper inter-tidal zone (barnacle zone) shows relatively flat-bottomed basins, -20-30 em wide and 10-15 em deep, with overhanging rounded edges and barnacleencrusted, rounded inter-basin plateau remnants.

c) The lower inter-tidal zone, or the mussel/echinoid zone, shows sharp,jagged, narrow pinnacles separating deep, rounded to conical basins, with widths, depths, and diameters all 30-80 cm. These basins are lined with echinoderm pits. This is the zone of greatest relief.

d) The sub-tidal zone shows subdued relief with no basin or pit karren. TrudgiU (1987) shows some proftles from Ireland with cliffed edges at low tide level or slightly overhanging, representing what might be a slight sub-tidal notch. It is obvious from this description, and a comparison of Figures 3 and 5 with Figure

8, that the Norwegian coastal karren and the Irish coastal karren have little in common. The only other cold region coastal karren described in the literature are from the west

coast of Newfoundland which has a much lower rate of isostatic uplift than Norway. Malis (1997) found the maximum relief in the high-tide swash zone of maximum energy level. Again the Norwegian coastal karren do not conform to this pattern.

5. Summary and Conclusion

To be classified as "characteristic" a landform must be in equilibrium with the conditions in operation today. This is not true for the Norwegian littoral karren which is strongly influenced by the glacial legacy and by isostatic uplift rates. So, this cold region coast is clearly a landscape in transition. The second criterion for a characteristic landform, that it must have a form that will remain constant over time through negative feedback control, is also not true for this cold region coast. The form will change as the rates of isostatic uplift change. However, there is a unity of form-namely, the greatest development in the supra-littoral zone-that could be considered as a model for cold regions karsts with isostatic uplift.


I Supra-littoral pans I Small-scale fretting and lichen

/' --~ 30cm

I Littorina zone I '" White surface on pan floors

Greysurfaces~

30cm / Black-brown surface inside basins

I Barnacle zone Micro-fretted surfaces ........ 30cm Basin surfaces smooth ? ~

Grazing gastropods, algae,etc

Mussels colonize old echinoid pits

I MussellEchinoid zone I ~

.,,/' Echinoid pits line basins

100cm

o 5 10 15

Mettes from low water mark

20

Terrestrial karren Supra-littoral pans

I Ireland

25 30 35

Figure 8: Profile oflittoral karren zones in the Burren District, west coast of Ireland (from field observations combined with Lundberg 1976, 1977 and TrudgiIl1987).


The features common to all littoral-coastal karsts are an absence of joint control, a dominance of pit-basin-bowl negative forms on an erosional ramp, and a dominance of biological (erosional or encrustational) processes. The Norwegian littoral karren show an absence of joint control and a ramp with pit-basin-bowl forms. However, this ramp is not a product of erosion; rather it is a legacy from glacial scour emerging through isostatic uplift into the littoral zone. Thus the ramp form in the littoral zone of different climates may not represent convergent evolution rather than a common developmental stage.

The Norwegian karren do not show a strong biological imprint. Bioerosion is expressed mainly through biochemically-enhanced dissolution and micro-scale bioerosion, rather than through obvious bio-erosional rock carving. Other weathering processes such as frost action, salt action, and grusification are important. Mechanical modification of the positive residuals is stronger than in other coastal karst areas, although large-scale storm removal of blocks is less common. This is partly because the rock is not regularly jointed and bedded, partly because the glacially smoothed surfaces offer few sites of weakness for hydraulic action, and because storm action is mediated by the wide strandflat.

The imprint of isostatic uplift on the Norwegian littoral-coastal karren causes development completely unlike what is known from other coasts which show a gradual downcutting of the surface so that the lower morphological zones inherit the forms from the upper zones. In Norway the land uplift causes, instead, the upper morphological zones to inherit forms from the lower zones. Thus, while the greatest rate of erosion in Ireland is in the deep MussellEchinoid bowls of the lower eu-littoral and the basins get shallower up the profile, the deepest basins in Norway are in the supra-littoral and are only about one quarter the size of the biggest Irish basins.

As in most coasts, increased exposure levels cause increased erosion rates. Sheltered shores may show virtually no post-glacial modification while exposed shores may show several decimetres of denudation. Frost action is relatively minor as ice appears to modify the forms only in the relatively sheltered and fresher waters of the Svalbard fiords. In general, inter-tidal erosion is not as active as might be expected.

This research suggests that dissolution is active in cold regions, even in Svalbard, but that bio-erosion is not so apparent as in other coastal karsts. However, the fast rate of isostatic uplift masks the fundamental processes in the cold region inter-tidal zone. No evidence was found on which to judge the probable characteristic karren morphology of a stable littoral zone.

References

Alexandersson, T. (1976) Actual and anticipated petrographic effects of carbonate undersaturation in shallow water, Nature 262: 653 - 658.

Back, W., HanshaW, B.B., Hennan, J.S, and Van Oriel,J.N. (1986) Differential dissolution ofa Pleistocene reef in the ground-water mixing zone of coastal Yucatan, Mexico, Geology 14,137-140

Bard, E., Hamelin, B., Fairbanks, R.G., and Zindler, A. (1990) Calibration of the 14C timescale over the past 30,000 years using mass spectrometric U-Tb ages from Barbados coral, Nature, 345,405-410.

Brookes, I.A and Stevens, R.K. (1985) Radiocarbon ageofrock-boring,Hiate1la arctica (Linne) and postglacial sea-level change at Cow Head, Newfoundland, Canadian Journal of Earth Sciences 22, 136-140.


Fairbanks, R.G. (1989) A 17,OOO-year glacio-eustatic sea level record: influence of glacial melting rates on the Younger Dryas event and deep ocean circulation, Nature, 342, 637-642.

Fjeldskaar, W. (1997) Flexural rigidity of Fen no scandia inferred from the postglacial uplift, Tectonics 16(4), 596-608

Focke, I.W. (1978) Limestone cliffmorphology on Cura~ao (Netherlands Anti1\es)with special attention to the origin of notches and vermetidlcoralline algal surfbenches ("cornices", "trottoirs"), Zeitschriftfuer Geomorphologie 22: 329 - 349.

Folk, R.L., Roberts, H.H. and Moore, C.H. (1973) Black Phytokarst from Hell, Cayman Islands, British West Indies, Geol. Soc. of America Bull. 84: 2351 - 2360.

Ford, D.C. and Williams, P. (1989) Karst Geomorphology and Hydrology Unwin Hyman, London, 601 p. Gnmlie, A. (1981) The Late and Postglacial Isostatic rebound, the eustatic rise of the sea level and the

uncompensated depression in the area of the Blue Road Geotraverse, Earth Evolution Sciences 1, 50-57

Holbye, U. (1989) Bowl-karren in the littoral karst ofNord-Amsy, Norway, Cave Science 16, 19-25. Holtedahl, H. 1998 The Norwegian strandflat - a geomorphological puzzle, Norsk Geologisk 1idsskrift 78, 47-

66. Jones, B. (1989) The role of microorganisms in phytokarst development on dolostones and limestones, Grand

Cayman, British West Indies, Can. J. Earth Sci. 26: 2204 - 2213. Lauritzen S.-E. (1991) Karstformer I Norge, 1:500 000, Geologisk Institutt, Avdeling B. Universitetet I Bergen. Lauritzen, S.-E. (1984) A symposium: arctic and alpine karst, Nonk Geogr. 1idsskr. 38: 139-143 Lundberg, J. (1974) The /carren of the littoral zone, Burren District, Co. Clare, Ireland, Unpublished Bachelor's

thesis, Trinity Co1\ege, Dublin. Lundberg, I. (1977) Karren of the littoral zone, Burren District, Co. Clare, Ireland, Proceedings of the 7th

International Speleological Congress, 291-293, Sheffield. Malis, C.P. (1997) Littoral karren along the western shore of Newfoundland, Unpublished M.Sc. thesis,

McMaster University, Hamilton, Ontario. Malis, C.P. and Ford, D.C. (1995) Littoral karren along the western shore of Newfoundland, Geological Society

of America, Abstracts with Program 27(6), A-56 Moe, D. and Iohannessen, P.I. (1980) Formation of cavities in calcareous rocks in the littoral zone of northern

Norway, Sarsia 65, 227-232 Meller, 1.1. (1989) Geometric simulation and mapping of Holocene relative sea-level changes in Northern

Norway, Journal of Coastal Research 5 (3), 403-417. Mottershead, D.N. and Pye, K. (1994) Tafoni on coastal slopes, South Devon, u.K, Earth SUiface Processes

and Landforms 19,543-563 Nielson, N. (1979) Ice-foot processes. Observations of erosion on a rocky coast, Disko, West Greenland,

Zeitschrift fuer Geomorphologie, 23 (3), 321-331 Schneider, I. and Torunski, H. (1983) Biokarst on limestone coasts, morphogenesis and sediment production,

Marine Ecology 4: 45 - 63. Spencer, T. (I 985a) Marine erosion rates and coastal morphology ofreeflimestones on Grand Cayman Island,

West Indies. Coral Reefs 4: 59 - 70. Spencer, T. (1985b) Weathering rates on a Caribbean Reeflimestone: results and implications, Marine Geology

69: 195 - 201. Torunski, H. (1979) Biological erosion and its significance for the morphogenesis of limestone coasts and for

nearshore sedimentation (northern Adriatic). Senckenbergiana Marit. 11: 193 - 265. Trenhaile, A.S. (1987) The Geomorphology of Rocky coasts Clarendon press, Oxford., 384 pp. Trenhaile, A.S. and Mercan, D. W. 1984 Frost weathering and the saturation of coastal rocks. Earth Sur/ace

Processes and Landforms 9,321-331. Trudgill, S.T. (1983) Preliminary estimates of intertidal limestone ernsion, one Tree island, Southern Great

Barrier Reef, Australia. Earth Sur/ace Processes and Landforms 8: 189 - 193. Trudgill, S. T. (1987) Bioerosion of intertidal limestone, Co. Clare, Eire - 3: zonation, process and form. Marine

Geology 74, 111-121 Trudgill, S.T. and Crabtree, R. W. (1987) Bioerosion of intertidal limestone, Co. Clare, Eire - 2: Hiatella

arctica. Marine Geology 74, 99-109. Trudgill, S.T., Smart, P.L., Friederich, H. and Crabtree, R.W. (1987) Bioerosion of intertidal limestone, Co.

Clare, Eire - 1: Paracentrotus lividus, Marine Geology 74, 85-98. Viles, H.A. (1984) Biokarst: review and prospect, Progress in Physical Geography 8: 523 - 542


Viles, H.A. and Trudgill, S. T. (1984) Long term remeasurements of micro -erosion meter rates, Aldabra Atoll, Indian Ocean, Earth Surface Processes and landforms 9: 89 - 94.

Whitaker, F.F. and Smart, P.L. (1994) Bacterially-mediated oxidation of organic matter: a major control on groundwater geochemistry and porosity generation in oceanic carbonate terrains, in Breakthroughs in Karst Geomicrobiology and Redox Geochemistry, Sasowsky, I.D. and Palmer, M.V. (cds). 72-74, Karst Waters Institute, Virginia.

SEDIMENTARY CHARACTERISTICS, BIOLOGICAL WNATION AND PHYSICAL PROCESSES OF THE TIDAL FLATS OF IQALUlT, NUNA VUT

JANIS E. DALE, SHANNON LEECH, S. BRIAN McCANN* AND GLENDA SAMUELSON** Department of Geography University of Regina Regina, Saskatchewan S4S 6S6

*Department of Geography and Geology McMaster University Hamilton, Ontario

**present address Department of Geography University of Calgary Calgary, Alberta T2N 1 N4

Abstract

Tidal flats in temperate areas tend to be dominated by tidal processes, resulting in distinct sedimentary characteristics and associated flora and fauna. Fewer studies exist in arctic regions where the classic form of temperate tidal flats is less obvious. This paper discusses the sedimentary characteristics, biological zonation, and dominant physical processes affecting the subarctic, macrotidal tidal flats near Jqaluit, Nunavut. Tidal, wave, and ice processes dictate the development of tidal flat morphology and the distribution of organisms.

The semi-diurnal tides average 7.8 m above low, low tide (ALLT), with large tides of 11.6 m ALLT. The two high and two low tides are experienced daily, even during ice cover that averages 8 to 9 months of the year. Tidal currents move sediment, biota, raft ice floes with their sediment or boulder-rich load, and affect water temperature and salinity at the bed. All of these factors limit the distribution of flora and fauna. Exposure indices generated from tidal data, reveal two critical tidal heights around 4.0 m and 7.5 m ALLT. The boundary between motile and less motile (sedentary) fauna occurs around 4.0 m ALLT, and 7.5 m ALLT marks the limit of most marine flora and fauna, with the rare exception being Fucus evanescens. The effectiveness of wave action is restricted to ice free periods. Thus the sorting of sediment by wave action is limited to a few months each

205

K. Hewitt et af. (eds.), Landscapes o/Transition, 205-234. © 2002 Kfuwer Academic Publishers.

206 Janis E. Dale, Shannon Leech, S. Brian McCann, Glenda Samuelson

year. Ice floes also dampen wave action. The longer the ice-free period the greater the degree of sorting of the intertidal deposits after the disruptive effects of ice formation.

The tidal flats are divided into six motphological zones and three closely associated biological zones. Fauna of the upper flat and beach areas are hardy and freshwater tolerant. Many are highly motile, opportunistic species that recolonize the area after ice breakup. These zones are the most intensively affected by ice action. There is a gradual change in fauna from highly motile species to more sedentary ones towards the lower end of the middle flat. Below 2.2 m ALLT, the tidal flat has smaller sizes and decreasing numbers of boulders. In addition, this area has the highest diversity and species richness due to increasing numbers of sedentary infauna such as Mya truncata, larger tubiculous polychaetes and the anemone Tealia sp.

Anthropogenic effects are most apparent on the flats nearest the tOWD-centre, at sites proximal to the sewage lagoon, dump sites and areas graded by the hamlet for vehicular traffic. Reduced species richness, diversity and evenness values characterize the invertebrate population at these sites. Elsewhere communities appear stable and show no statistically significant difference from studies in the 1980s.

1. Introduction

Tidal flats are low gradient, low relief, poorly vegetated intertidal coastal features, built largely of unconsolidated sediments. They occupy the area between high and low water marks of the intertidal zone, typically in sheltered coastal areas under mesotidal or macrotidal conditions (Davis 1992). While tidal flats occur at all latitudes, they are better known, more widely studied and best developed in the low and mid-latitude zones (Dionne 1988). Since the early 1900s, most tidal flat research was conducted at temperate sites in Europe and North America, where the sedimentary environment and associated processes of tidal flats have been well documented (e.g. VanStratten 1954; Evans 1965; Kellerhals and Murray 1969; Reineck 1976; Kooistra 1983; Dionne 1984; Dalrymple 1992). Tidal processes prevail over wave processes in most temperate situations. Periods of exposure and inundation are also important factors in tidal flat development. Typically, these flats are characterized by three distinct sedimentary zones, fining shoreward sequences and associated flora and fauna. Fewer studies exist in arctic regions where the classic form of temperate tidal flat is less obvious (Gilbert and Aitken 1981; McCann et al. 1981a; Dale 1982; Martini 1991). Along with tidal and wave processes, ice action also plays an important role in the development of tidal flat zones and the distribution of organisms. Studies of the sediment distribution, biological zonation and dominant processes have been ongoing since the early 1980s near Iqaluit, Nunavut. This paper discusses the sedimentary characteristics and biological zonation of these tidal flats in terms of the arctic processes that shape them.

Tidal Flats ofIqaluit, Nunavut 207

2. Site

The field area is located at the head of Frobisher Bay on the southeast comer of Baffin Island, in the territory ofNunavut (Figure 1). Some 250 kilometres of coastline have been examined during aerial studies of ice breakup along the shoreline, with detailed studies at Rogers Island and Koojesse Inlet by the Hamlet of Iqaluit (Figure 2). A good portion of this coastal area consists of tidal flats, some reaching over 3 Ion in width (Figure 3). The area experiences macrotidal conditions with semi-diurnal tides averaging 7.8 m ALLT large tides of 11.3 m and highest large tides recorded at 11.6 m (Dept. of Fisheries and Oceans 1995). Two high and two low tides are experienced daily even during periods of ice cover, which average eight to nine months of the year.

Iqaluit has a subarctic climate with a mean annual temperature of - 9.0°C and precipitation of 409.3 mm, with most occurring in the summer as rain (Environment Canada 1993). Frobisher Bay has subarctic marine conditions, receiving waters from polar and non-polar origins (Wilce 1959) and identified by the presence of subarctic species, Mytilus edulis, Littorina saxatilis and Balanus balanoides (Dunbar 1951).

The entire area was subject to numerous glacial advances during the Pleistocene (Bird and Schwartz 1985). The most recent glacier likely retreated some 6,750 +1- 170 years B.P. (Blake 1966) leaving an extensive moraine at the head of the bay (Miller et al. 1980). Local sea level fluctuated towards the end of the Late Wisconsinan, reaching its present level by 8 kaBP (Dyke and Prest 1987), the result of a combination of isostatic rebound and eustatic sea level rise. Raised shorelines and deltas indicate that the area of present tidal flats was once under fairly deep water. It is likely that it was at this time that a grey sedimentary layer was deposited and which underlies the tidal flats ofKoojesse Inlet and Rogers Island. This layer likely contributes much of the silt, sand and gravel now present on the flats. The local bedrock consists of Precambrian gneiss with outcrops of paleozoic sediments (Miller et al. 1980), and is the source of many of the boulders on the flats today. It forms the cliffed coastline common to the area.

3. Methods

The earliest studies of the tidal flats in Frobisher Bay were undertaken by Ellis (1955; 1960) and Ellis and Wilce (1961). Their work formed the basis of many of the subsequent studies by other individuals including Grainger (1954) and Wacasey and Atkinson (1975). Much of this paper is based upon studies at the head of Frobisher Bay and around the Iqaluittownsite in the early 1980s (McCann et al. 1981a,b; Dale 1982; McCann and Dale 1986) and more recent work in the 1990s (Samuelson 1997; Leech 1998) (Figure 2). These studies concern the physical characteristics of the tidal flats in terms of sediment and boulder distribution, biological zonation and the modern processes, tidal, wave and ice action that dominate the tidal flats today.


Baffin Bay

100 0 100 200 Kilometers ,......

Figure 1 : Map of the study area -- Iqaluit, N.W.T.

Figure I. Map of the study area-Iqaluit, NWT.

Tidal Flats of Iqaluit, Nunavut 209

Figure 2a. View across the tidal flat in Koojesse Inlet, total width approximately 900 m.

Figure 2b. Upper Flat zone with beach and fines flats with pooled water (5.5-10 mALL T).


.., l!! ::I .., 0 g,

~ -0 ... II .a E ::I Z

70·

60

50

40

30

20

10

0 0

Number of Exposures at Varying Tidal Heights on the Tidal Flats of Koojesse Inlet

Level above Datum (m)

Log of Maximum Duration of Exposure and Inundation in Hours for Varying Tidal Heights in Koojesse Inlet

3r------------------------------------------------, f 2.5+-~ .. ~~--------------------------------------~~ ::I o ::I: ~ 2+---------------~------------------------_.~----------~ c o ! 1.5 t----------'t---------/---------j ::I o E E i "" 0.5 '0 ! 0 .... , ... ·····t·····t·-+· .. +·· .. I··+··+-+··+···,··· .. ~··+·-i····+ .. ·t .. ·+··+·+·+·+····t····+-,·· .. ··t····I-+··+·+··+····t····+···-I···+··+····t·····,······I· .. ·I-·

-0.5 -'-----------------------------'

Level Above Datum (m)

Figure 3. a) The number oftidal exposures plotted against tidal height for July and August; b) The log of maximum duration of exposure and inundation in hours plotted against tidal height.


4. Modern Processes

4.1. TIDAL CONDITIONS

Tidal action plays a key role in the development of these subarctic tidal flats. Flood and ebb tides sweep across them at impressive speeds, particularly during spring tide conditions, when it is not uncommon to observe standing waves develop offshore between rock outcrops. Currents over the tidal flats during flood tides have been measured at .26 ms-\ (Dale 1982). Initially, water floods up channels and low areas on the flats. As the depth increases the channels are breached and water floods adjacent areas. Small current ripples and tidal wedges are created at this time. At high water, gyres of water rotating in great circles over the flats are mirrored by the movement of ice floes. During ebb tides the lowering of the water surface gradually exposes sediment mounds and boulders, until the water follows the channels on the flats, where it continues to drain throughout much of the low tide period.

Calculations of the exposure curve for the tidal flats are based on predicted tidal values using a program developed by Smart and Hale (1987). Comparisons with predicted tidal values from other years show no significant difference in the exposure curve over the summer months in Koojesse Inlet (Figure 4). The number of tidal exposures is plotted against tidal height (Figure 5) and again for the log of maximum duration of exposure and inundation in hours with tidal height (Figure 5). Values show prominent breaks in slope at 4.25 mALLT and 7.5 mALLTinJuly and 7.75 minAugust. Additional breaks occur at 3.75 m ALLT on the inundation line and at 8.5 m in July and 8.0 m in August on the exposure line. These delimit the boundary between tidal heights undergoing regular oscillations versus those experienced during spring tides. Minor monthly variations are reflected in the differences between July and August. The prominent heights around 4.0 and 7.5 m ALLT are also reflected in the morphological and biological zonation of the tidal flats (Dale 1982).

Calculations of instantaneous vertical velocities using the tidal curves, at .25 m intervals, show the highest velocities (0.14 ms-I ) occur at sites around 6.0 m ALLT, or roughly midway between the 4.0 and 7.5 m heights (Smart and Hale 1987). While ebb and flood velocities were similar in August, ebb velocities were noticeably higher in July, probably the result of variations in global tidal fluctuations as opposed to local controls. Horizontal velocities were calculated from the vertical velocities using gradient data from each transect survey at 50 m intervals down the flats. On all transects, results show that steeper slopes experience lower velocities and shallower slopes the fastest. Areas where the highest velocities were predicted to occur did correspond to coarser sediments and lower faunal densities. Average flood velocities were higher than average ebb velocities; however, the highest instantaneous velocities always occurred during ebb tides. Consequently the orientation of ripples and small tidal wedges exhibit a mixture of ebband flood-dominant directions on the flats.


CUmtnt Velocities Over the Tidal Flats With Ice Cover

0.9r----------------------------------------------------,

0.8

0.7

i i 0.6

I::: I 0.3

B 0.2

0.1

-+-HghTide

......... EbbTKIe

-"-Low Tide

........ FIoodTide

o 0.5 1.5 2 2.5 3 3.5 4 4.5 5 5.5 6 789

0.16

0.14

_ 0.12

~ I 0.1 >-:= ~ 0.08 , I 0.06

B 0.04

0.02

0 0.1

Water Depth 1m,

Current Velocities Over the Tidal Flats With Open Water in August

0.5

-+-HghTide

......... Ebb Tide

-"-Low Tide

........ FIoodTlde

1.5 2 2.5 3 3.5 4 4.5 5 5.5 6 6.5 7

Water Depth 1m,

Figure 4. Tidal current velocities over the tidal flats with a) ice cover; b) open water.

Tidal FIats ofIqaluit, Nunavut

Water Temperatures Over·the TIdal Flats During Ice Cover

3

2.5

2 -+- High Tide

-+-EbbTlde

1.5 -*-Low Tide

!:!. ____ Flood Tide

~ ::s 0.5 i ! 0

J -0.5 7 8 9

-1

-1.5

-2

Water Depth 1m,

Water Temperatures Over the TIdal Flats with Open Water in August

9

8

7

6

§: 5 I!! i 4 ! E 3 {!!.

2

1 .

-+-HghTlde __ Ebb Tide

-*-Low Tide ____ Flood Tide

o~~~~~~~~=±~~~~~~~ 01 0.5

-1 1.5 5.5 6 6.5 7

Water Depth 1m,

Figure 5. Water temperatures over the tidal flats with a) ice cover; b) open water.

213


4.2. MARINE CONDITIONS

4.2.1. Tidal Current Measurements Measurements of water temperature, salinity and tidal current velocity were taken at two sites, on both the north and south side of the inlet, over four tidal cycles, once with ice cover and once ice free. The greatest current velocities always occur in the upper 3 metres of water regardless oftotal depth (Figure 6). The higher velocities only approach the bed when the water depth is less than 3 metres, at the early stages of flood tide and later stages of ebb tide. Ebb tidal currents generate the highest velocities, supporting the findings from the exposure curve calculations. However, these currents do not increase regularly, but appear to pulse with higher values at 1.5 hour intervals during ebb tide. This may relate more to the shape and bathymetry of Frobisher Bay than to more local effects. The presence or absence of ice did not influence current velocities at the bed, although strong winds and waves set up in the ice free period, result in an increase in velocity at the water surface.

4.2.2. Marine Temperatures Marine temperatures vary throughout the tidal cycle, although the highest temperatures always occur at the surface, ranging between 0° to 3°C during ice cover and 5° to 8°C when ice free (Figure 7). The coldest temperatures occur at depths below 3 m to the bed during spring ice cover (-1.5°C) and at similar depths in the summer (2°C). The variations in water temperature during the tidal cycle likely do not adversely affect the intertidal biota. However, exposure to sunlight did raise the surface sediment temperatures from 2° to 10°C in just a few hours, although at depths below 2.5 em little heating was noted. This could have some effect on sedentary organisms especially those inhabiting the upper few centimeters.

4.2.3. Salinity More important to biological zonation is the variation in salinity, particularly evident during sea ice ablation. Salinity increases with depth while the fresher, less dense water remains at the surface (Figure 8a,b). During ebb tides, the lens of fresh water lowers towards the bed, with significant differences in salinity occurring in water depths less than 3 m. During ice breakup at low tide, salinity measurements range from 0 to 8 ppt at the bed. This could affect the freshwater intolerant species directly or indirectly by discouraging planktonic settlement in those areas affected by the freshwater lens.

The average salinity during ice free periods ranges between 29 and 32 ppt, with considerably less variability than during breakup. The lowest salinity occurs at the top of the water column, in the upper half metre. Twelve ppt is the lowest summer salinity value recorded on the tidal flats at sites distal to river influences. This value could conceivably cause duress for intertidal organisms over extended periods of time. However, wind and wave action in the summer months, plus the cyclic nature of the tide helps to mix the water rapidly at shallow depths, reducing the effectiveness of this variability. In the summer, freshwater sources include terrestrial run-off and occasionally ice rafted from down bay (Dale 1982).

Tidal Flats of Iqaluit, Nunavut

Salinity Values Over the Tidal Flats During Ice Cover

~r----------------------------------------------'

30

25

~20 ~ .5 i 15

5

~

30

25

~20 >-

;I::

.5 15 i 10

5

0

-+-HghTlde

-e-81bTide

""*- Low Tide __ Flood Tide

0.5 1.5 2 2.5 3 3.5 4 4.5 5 5.5 6 7 8 9

Water Depth 1m)

Salinity Values Over the Tidal Flats with Open Water in August

r ~

0.1 0.5

-'" ....

-+-HghTlde __ 81biide

""*- Low Tide __ Flood iide

U 2 ~ 3 ~ 4 U 5 U 6 M 7

Water Depth 1m)

Figure 6. Salinity values over the tidal flats with a) ice cover; b) open water.

215


1- Penn-m Ice Complete Freeze ev. FIm Ice DelAlrfondion Ice free

• • • • I 1990 1991

110 __

1885 .. ..

19 74 No_available 1975

.. .. 1968 1869

I 196 3 1S1114

s o N o J F M A M J J A

Figure 7. Formation and deterioration of ice on Koojesse Inlet from 1963-1990.

4.3. WAVE ACTION

The effectiveness of wave action is restricted to ice free periods, limiting wave sorting to a few months each year. Ice floes dampen wave development, thus the longer the ice free period, the greater the influence of wave action. Over two summers of study, fewer than 10 episodes of waves greater than 10 cm in height occurred on the Koojesse flats. Wave action did produce sinuous crested ripples on isolated sand patches, and modified the sediment ridges and crescentic sand waves at Apex River. During one stonn, the sand beaches at the head of the bay were pounded by waves 1 m in amplitude. These waves removed substantial amounts of sediment from the beach and redistributed it along shore and in a seaward direction. Wave ripples 2.5 cm in amplitude were observed on sand patches after the storm.

Tidal Flats ofIqaluit, Nunavut

Figure 8a. Jumbled intertidal ice floes pushed onshore and on the ice foot during breakup.

Figure 8b. Ballycatter with a large boulder in the centre, that appeared to have been lifted up from the tidal flat surface by the ice.

217


On average the most frequent wind direction from June to August is from the south-east (Crane 1978; Environment Canada 1993), the direction of maximum fetch length. However, the presence of Long Island and the orientation of the tidal flats probably helps to limit high energy wave action in Koojesse Inlet White caps are frequently observed in the open stretches outside of Koojesse Inlet, and around Rogers Island, that does experience more frequent high energy wave conditions, resulting in medium well sorted sands and the formation of2 m high sand waves.

Field observations over four summer seasons suggest that, infrequent, but high intensity wave action is more significant than frequent but low intensity wave action on the Koojesse flats. The more infrequent high energy events shape the sediment beaches, form the large sand waves found in Tan Inlet and at the Apex River.

4.4. ICE ACTION

Ice plays a significant role in the tidal flat environment throughout the year. Using data from 1963 to 1990 (Canadian Ice Service 1998), the "ice year" can be divided into 4 periods based on average dates. It begins with the freezeup period, typically lasting from October 20 - November 15; complete ice cover from November 14 - June 11; breakup from June 10- July 18, and open water from July 19 - October 19 (Figure 9). On average it takes 19.3 days between the first permanent ice conditions and complete freeze over, with full ice cover lasting 158 days. Breakup takes approximately 38.1 days between the frrst deterioration and ice free conditions. Early ice formation does not necessarily result in early ice-free conditions, nor do later ice-free conditions result from late ice formation. Similarly the earlier onset of ice deterioration may not result in earlier ice-free conditions. Local variations in temperature, precipitation, wind speed and direction dictate the rate of ice formation and breakup.

Generally the ice surface is relatively smooth and clean over deep subtidal waters and chaotic and dirty over the intertidal zone (Figure 10). Typically the intertidal ice in Koojesse Inlet averages 1.4 to 2.0 m in thickness (McCann et al. 1981a,b; NIC 1995). An icefoot develops around the high tide portion of the flats, ranging in width from 16 m to 74 m, with wider sections present on flats oflower gradients. Seaward of the ice foot is a jumbled surface with ballycatters, circular, cone-shaped ice deformation structures created by the rise and fall of the intertidal ice over large boulders. These range from 1.5 to over 3 m in height (Figure 11) There are also snow mounds and snow ridges with heights over 1 m, separated by broad depressions called craters (Leech 1998).

During freezeup, materials are incorporated in the developing intertidal sea ice, involving considerable disturbance to both sediments and associated organisms. A layer of sediment can freeze onto the base of the ice as it is lowered to the bed at low tide. During high tide a layer of seawater can freeze to the base of the ice, resulting in a series of alternating layers of sediment-rich and sediment-poor ice. The presence of large boulders restricts ice settlement to the bed and interrupts sediment incorporation, reducing the sediment load and disrupting the stratigraphy of the sediment bands within the ice.

e

long fl.

d


a .,. lcelool

EJ Open Wa'e,

I11III "-do. fce

a fractured Sea

D~ICfs...lce - .... )0< LAaCfl

219

figure 9. Typical sequence of ice breakup in Koojesse Inlet a) June 19/80, b) June 25/80, c) June 28/80, d) July 8/80.


C'

! 11 (I)

.5 I! CJ

3.5T

3.01 251

i' 1-e.. 2.0 ••••••••• • •• - •••• = t, -. w- '.. ........

i~!+------,+~-----'-:.+:-··-··-·~---3D+,------~_\~~;-----SOO~.-.. -.. -~.-.-.. -~~.-.. -.. -.. -.-.. ~~~----~~ Clllanceflam ...........

Mean - - - - - -Sorting Mean Gr.Sz. Regres.lon I 3.5 .

3.0

2.5

2.0

1.5

1.0

0.5

0.0 800 700

-0.6

-1.0 Distance from ahore reference (m)

Figure 10. Graphs showing the variability of mean grain size, sorting and regression lines along two transects indicating both fining and coarsening trends in a shoreward direction

Tidal Flats ofIqaluit, Nunawt 221

With increasing depth across the flats, the length and number of times that the ice settles to the bed is reduced. Measurements of sediment in the intertidal ice shows that the ice formed over the upper and middle flats (>3 mALL n contains the greatest amount of sediment, with decreasing amounts in a seaward direction. On average the thickness of sediments contained in the ice at the head ofKoojesse Inlet is 8.17 em . Estimates show that between 63,750 t km·2 (Dale 1982) and 68,000 t km·2 (Leech 1998) of sediment may be incorporated in the intertidal ice annually in the upper flat reaches ofKoojesse Inlet. These values are compatible with those obtained at other arctic (Gilbert 1983) and temperate tidal flats (Dionne 1984). Differences arise, however, in the redistribution of this material. At arctic sites the intertidal ice melts out first, and the sediment is redistributed within the intertidal zone. Whereas at temperate sites, sediment can be rafted into deeper waters since the intertidal ice breaks up last (Dionne 1984).

The inclusion of coarser materials is not well understood and may be due to a number of mechanisms. Small stones attached to algae can be incorporated when the algal fronds become frozen in the developing sea ice (Martini 1980). Coarser debris can also become entrained during freezeup when an expanding ice cover buckles under pressure, creating a ridge through which cobbles and boulders are brought to the ice surface. Boulders can be entrained through the formation ofballycatters (Dionne 1973; Rosen 1979). Coarse material becomes incorporated within the ice and then thrust up within the structure (Figure 11). Frost riving and plucking, are other means of incorporation (Drake and McCann 1982). Ice push by pack ice can also thrust boulders onto stranded intertidal ice which is subsequently refloated at high tide. Boulders are most likely to be incorporated when small, isolated and lying loosely on the substrate surface. The length of the longest axis ofrocks found on the ice surface averaged between 70 and 90 mm, in the zone between 3 and 6 m ALLT, but boulders up to 1 m in diameter are regularly transported. Of36 boulders embedded in the ice on five different transects, 22 were moved more than 2 m. Seven moved distances between 6 and 17 m, predominantly in a shoreward direction (Leech 1998). the direction, length and fraction of boulder movements are highly variable each season as is found in other tidal flats subject to ice action (Gilbert and Aitken 1981; Dionne 1988).

The sequence of ice breakup on the tidal flats at the head of Frobisher Bay has been monitored over several field seasons (Figure 12). Melt begins on the upper ice surface with the arrival of warmer temperatures in the spring. Sediment on the ice surface also contributes to a reduction in albedo and increased heating. The cyclic vertical movement of the ice over the intertidal zone results in cracks and leads developing at the ice foot, tidal flat boundary. Ice breakup is initiated around stream mouths by terrestrial runoff, which occasionally flows across the ice surface, draining through leads, some distance from shore. Breakup expands to the intertidal ice beside the ice foot and continues in a seaward direction. Ice depressions or craters become sites of water and sediment accumulation, enhancing melt, until melt holes develop permitting drainage of the upper surface during low tide and percolation of marine waters to the surface during high tide.


Upper Flat MIddle Flat I.owerAat

'------/-7.5 5.0 4.5 4.0 2.2 0.0

lOO}, I finer Flat 1 Boulder( FIOt I Boudef I Very BouIdeIy Flat 80 - Beach Ac:tge

% 60 ~,

I Gladed Flat I

ExposuI9 40j ______________________ ___ 20 _______ ._

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12 10

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Banen ,-Flat I ~FIat 81.4 =::: 6.2 Molle 4.6 leA Molle 2.2 0.0 ~,~ ____________________________ ~. ~ I

4 species 30 apeeIes MIddle flat 11 spec:Ie8

Figure 11. Morphological zones and biological zones graphed with exposure ranges along a transect.

Ice that forms in the upper flats experiences less total movement and is impeded by the still solid offshore ice. As the area of open water increases, ice floes are subject to tidal currents and increasing wind and wave action. The net direction of ice floe movement is shoreward with rising tides and wind action. Floes can pile onshore three or four deep on the ice foot and upper flats (Figure 10). Inshore floes ablate until the weight of sediment reduces their buoyancy. They settle to the bed and melt in situ, releasing the sediment in the upper intertidal zone.

Experiments monitoring the movement of floes show that the amount of open water, tidal conditions and wind direction dictate the rate and direction of ice floe movement. Floes monitored over one tidal cycle, during peak ice floe coverage, move on average between 225 and 500 m to the west under prevailing south-east winds and flood tide conditions. During periods oflow wind, floes move by tidal currents and follow local tidal circulation patterns over the tidal flats. Floes move at rates of 18 to 51 mIhr during flood tides and increase to rates of 1 to 3 km/hr with the addition of wind. Experiments conducted over 5 tidal oscillations in 1995 in Koojesse Inlet, show ice floes moved between 450 and 920 m in a northeastward direction due to the prevailing wind and tidal conditions at that time. Some ice from the head of the bay moved northeastward over a distance of 1.5 Ian until stranded on the upper flats.

Tidal Flats ofIqaluit, Nunavut

Figure l2a. Middle Flat zone showing bouldery flats with sediment and boulder mounds (5.5-4.5 m ALLT).

Figure l2b. Middle Flat zone showing boulder ridges (4.5- 4.0 m ALLn.

223


Figure 12c. Middle Flat zone showing very bouldery flats (4.0-2.2 m ALLT).

Figure l2d. Lower Flat and the graded flat zone (<2.2 mALL T).


Gouging by drifting floes disturbs the tidal flat surface. Gouges are most prominent during the peak breakup period, when the maximum coverage of ice floes exist along with some open water. They are best developed early in breakup when large keeled ice floes from the ice foot and intertidal ice are available. Smaller and fewer gouges occur from the smoother subtidal ice floes rafted onshore later in the season. Gouges measured on the north side ofKoojesse Inlet show that the greatest number of gouges occur oriented parallel to shore in a northeast-southwest direction. This is due to a combination of prevailing wind directions from the southeast and large rotating gyres set up by ebb and flood tidal currents. Fifty percent of ice gouges are less than 5 m in length and 89% less than 30 m (Figure 13a). These features are generally short-lived and soon reworked by wave and current action, although deeper gouges were still visible by the end of August. Most of the gouges occur within 200 m of shore, but at varying tidal heights (Figure 13b). Subtidal fast ice driven onshore during the measuring period, likely accounts for the large number of gouges between 5 and 7.0 m ALLT on shallow gradient sites. Tidal flats with steeper gradients had the greatest number of gouges at tidal depths between 2.4 and 3.5 m ALLT.

Boulders and sediment incorporated in the ice can be transported far from their original sites. Each year after breakup, hundreds of boulders are deposited on the flats at the head of Koojesse Inlet after removal the previous year by the local excavator to expedite sea lift. Boulders marked for retrieval over two successive seasons reveal that some move 50 m or more annually. The average longest axis of these boulders is 21 cm. Many of these marked boulders are subsequently deposited at tidal heights around 5 m ALLT. Observations suggest that freezeup might also be a period of considerable movement, particularly for the smaller boulders, since many of the boulders had melted out of the ice before any large scale ice movement had occurred. Many boulders appear to move onshore or perpendicular to shore dependent upon prevailing winds and tidal currents.

5. Sedimentary Zonation

The erosion, transportation and deposition of materials by ice action creates a large turnover in sediments and boulders. Boulders can be rocked, rolled and transported by the ice creating boulder pavements, boulder ridges, boulder mounds and sediment mounds which characterize the sedimentary features of these flats. However, tidal flats in Koojesse Inlet and around Rodgers Island appear to be largely erosional in nature. The relative rarity of large scale depositional forms such as those observed in more temperate flats, the thin oxidized active layer (average 6.5 cm) and net erosion measured over two summers of study supports this view.

Ice action disturbs wave- and current-induced sediment zonation, typical of temperate tidal flats and results in very weak trends in grain size. Measurements across the tidal flats shows that the average grain size is on the coarse side, with just a few sites exhibiting weak shoreward fming similar to that observed in temperate flats. The fining


shoreward trend strengthens as the ice free season progresses from wave and tidal action. Over time sheltered intertidal sites can develop deposits of mud in the upper reaches. Areas such as Tarr Inlet, and other inlets at the head of Frobisher Bay have muddy zones now forming into mud flats.

The flats are separated into three units based on tidal position: upper, middle and lower flats. These, in turn, can be divided into six morphological zones: beach, fines flat, bouldery flat, very bouldery flat and boulder ridges and graded flat. Differences in the width and gradient of the intertidal zone throughout the study area results in a wide variation in the number of zones found along anyone transect.

The upper flat can include beach and fines flat zones. Beaches occur in small embayments as pocket beaches and in areas of low intertidal gradients such as the head of Koojesse Inlet. During freezeup the beaches are protected by the presence of the ice foot. Leads in the ice foot often result in the formation of sediment ridges that soon disappear once the ice is removed. Wave action helps to sort the beach sediments in the ice free season. The fmes flat zone can form just seaward of the beach under the right wave and tidal conditions, and is best developed at the end of the ice free period. When formed this zone contains the finest sediments along a given transect. It contains a few small boulders that are almost always incorporated in the ice during freezeup and settle on the top of the sediment. The zone can have low sediment mounds with a few boulders that may enhance further sedimentation. Typically, drainage channels are shallow and indistinct, becoming more prominent with increasing depth. During high wave events wave formed ripples are common.

Along steep intertidal gradients backed by cliffed shorelines, the boundary between the ice foot and fast ice marks the position of ice push features such as boulder barricades and boulder pavements, usually at tidal heights above 5.5 m ALLT. The beach zone is usually absent and the fines flat usually poorly developed unless the zone is a depression or surrounded by boulder ridges to form shallow tidal pools that never drain completely.

The middle flat contains the bouldery to very bouldery flat zones and boulder ridges where formed. The highest concentration of the largest boulders occurs in the midtidal zone, The bouldery flat consists of both sediment and boulder mounds, with boulders appearing perched on the surface. Gravel lines the meandering drainage channels, tidal wedges form in the drainage channels and small depressions on the flat surface. Although both directions are evident, flood dominant bedforms appear most dominant.

Boulder ridges where present form around the 4.5 m to 4.0 m ALLT. This height is conspicuous in terms of the exposure curves and appears to undergo heavy ice action. Ice gouging and ice push features are common during breakup. Large ballycatters form over the ridges in the winter, with boulders thrust up onto the surface, ice push and ice rafted floes are often snagged by the higher elevation of the ridge. The large number of boulders helps protect the underlying grey silt layer on which the ridge is forming. With the absence of boulder cover the grey layer was often exposed in this area from ice erosion. The boulder ridge also helps protect the inner flats from some wave and ice action. Boulder pavements appear to be forming along the slopes of these ridges.

100

90

80

70

~ 60

~ 50

~

~ 40

~ 30

20

10


MHT

12

12 Fucus evanescence 13 Fucus veslculosus 11

13

1 Cyrtodarto klJllana 2 HateIIa arc1Ica 3 ~ dscors and M. niger 4 MIa 1Mlcato 5 Jhracla myopsIs 6 Margar!Ies sp 7 sea anemone 8 Prq:luIId 9 EChlIJIO

1 0 BaIalus baIaloIde& 11 LJtta1na saxallis

04-~~+-.-+L+-~rL~r-~~-'~~~-r-r-r-.-..-.-.-,

0.0 1.0 2.0 4.0 5.0 6.0 7.0 8.0 9.0 10.0 11.0 12.0 13.0

lEVEL PJJCNE DAlUM (MEmES)

Figure 13a. Distribution of important macroalgae and macrofauna with tidal exposure.

100

90

80

Q 70 w

~ 60

w

~ 50 u.. 0 40 2: ~ 30 If

20

10

0 0.0

MHT

1 C.CXIIlIIaIa 2 Mociorna5tu< 3C._ 4 P.afIInIa 5 P. ,,","1eImiIIa

61._ 7 SpIo a>utuo BC._ 9 E. oneil, 10~~ llP.~o 12 SpIolt.eeI 13Sp1o_ 14Sp1o~ 15 A. acuItIIcno 16 H. ImbIeo1a 17 E.11Qvo 18 E. fonga 19 P. goOenialdcus

1.0 2.0 3.0 4.0 5.0 6.0 7.0 8.0 9.0 10.0 11.0 12.0 13.0

LEVEL N!ICNE DATUM (METRESl

Figure 13b. Distribution of important meiofauna with tidal exposure.

227


The greatest concentration of boulders occurs from 4.0 m to 2.2 m ALLT in the very bouldery flat zone. On gradual slopes it is common to observe discrete boulder mounds, often with the boulders embedded in the sediment. Along steeper gradients, these boulder mounds became elongated into ridges perpendicular to shore, which channel tidal runoff. Tidal drainage rather than ice micro-relief likely dictates the pattern. Boulder pavements appear frequently around boulder mounds and boulder ridges in this zone. Both flood and ebb dominant tidal wedges are common, as are ripples on patches of sandy substrate between the mounds.

The lower flat is graded and forms below 2.2 m ALLT ( <25% exposure). There is little relief, with fewer and smaller boulders, typically well covered with algae. The unconsolidated substrate appears well sorted although tidal channels do contain gravels.

The tidal flats themselves have undergone some physical changes due to human activities. Grading of the tidal flat for sealift has been ongoing for many years at the head ofKoojesse Inlet. Two rows of boulder accumulations have developed on either side of the scraped area. A series of breakwaters have been built on the north and south shores of the inlet. The breakwater on the south side near the sewage lagoon has probably restricted the normal movement of water, resulting in a greater accumulation of fmer sediments north of the causeway than that measured seaward.

6. Biological Zonation

Most of the 46 species of invertebrates and macroalgae identified at the head of Frobisher Bay exhibit a zonal preference from high to low tide. The reasons are varied but include length of aerial exposure, tolerance to freshwater, temperature variations, texture and water content of the substrate. In addition, wave action, current velocities and ice action may also restrict the distribution of flora and fauna. Macrofaunal counts were taken using 2 by 2 m quadrats, while meiofauna were sampled from cores covering areas of approximately 100 sq cm. Macroalgae were sampled throughout the intertidal zone.

Of the 14 macroalgae species collected, Fucus evanescens was the most common. It lives attached to cobbles and boulders or in the crevices of the bedrock from low tide to 8.0 mALLT. Fucus vesiculosus was only collected in one tidal pool (6.5 mALL T) under the influence of the ice foot, where specimens were badly tom after ice breakup. Remaining algal species including those from the classes of Chlorophyceae and Rhodophyceae appear to require less than 10% subaerial exposure and live on rock surfaces below 2.5 m ALLT.

Results from the quadrat analyses reveal that the highest densities and diversities of macrofauna occur between 1.0 and 2.0 m ALLT. The highest densities (110.2 per 4 m2)

occur at Rodger's Island, a flat just outside of Koojesse Inlet, due to large numbers of Cyrtodaria kurriana. Sea anemones (Tealia spp.) are the most common macrofauna, accounting for 47.8% of the total organisms counted in the quadrat analyses. Their upper limit appears to be around 4.2 m ALLT, although their range is extended by their presence in tidal pools in the upper flats. Bivalves account for 32.1 % of the fauna examined in the quadrats, the most common are Cyrtodaria kurriana and Mya truncata. Cyrtodaria


kurriana exhibit densities of92.8 per 4 m2 at 1.6 m ALLT at Rodgers Island. They require less than 15% exposure and are very susceptible to fresh water and ice gouging. Although shells are found throughout Koojesse Inlet, they were only found alive on a rocky transect on the south side of the inlet.

Mya truncata are found throughout the inlet, although their densities appear to be decreasing, a result of local harvesting. M truncata are noticeably absent from some sites in the 1990s study which they had inhabited in the 1980s. In addition the size of the harvested clams is smaller than those measured in the 1980s. M truncata live at sites below 15% exposure, although a few survive in tidal pools in the middle tidal zone. They are long-lived and have been recorded to reach the age of 42 years (Hewitt and Dale 1985). They require cohesive but fme sediments to sustain their burrows, that have been measured to depths of21 cm. Other common molluscs include Thracia myopsis, Hiatella arctica, Musculus discors and Musculus niger all requiring less than 10% exposure.

Two other common species in the intertidal zone include Littorina saxatilis and Balanus balanoides. Littorina saxatilis are found covering the algal covered boulder surfaces from 1 to 7 mALL T, with peak densities at 4.0 mALL T. Populations of juveniles and adults appear after ice breakup, and tend to stay on the same boulder for the duration of the summer, seeking refuge under algal fronds during high wave or wind action. Balanus balanoides are found from 5.0 m ALLT to the lowest part of the tidal flat. The barnacle distribution appears to be largely controlled by the availability of appropriate boulder habitat, with highest densities occurring between 4.2 m to 4.5 mALL T (164 per 100 sq cm), mainly on east and south facing aspects.

Meiofaunal samples include molluscs, sipunculids and priapulids, although the polychaetes are the most abundant and diverse. Of the polychaetes, Microspeeli theeli (formerly Scolelepis sp.) of the Spionidae family are the most abundant with densities as high as 727.5 per 100 sq. cm. They dominate the upper flat areas and are often the only species present at sites above 5.2 mALLT. They are tubiculous, opportunistic species only surviving a single season. They seem to appear literally over night, recolonizing the upper flats after the ice is gone. The Capitellidae are the next most abundant polychaete. Although motile, they maintain a network of tubes with access to oxygen-rich waters above the sediment. They tolerate muddier substrates and have been linked to enriched and polluted sites (Rygg 1985; Dale 1992; Smith et al. 1995).

The tidal flats are divided into six morphological zones and three closely associated biological zones, upper flats, midflats and lower flats. Fauna of the upper flat and beach areas are hardy and exhibit some freshwater tolerance. Many are highly motile, opportunistic species that recolonize the area after ice breakup, since these zones are the most intensively affected by ice action. Littorina saxatilis and Microspio theeli are the most common of the upper flat species, along with Fucus evanescens.

Errant polychaetes dominate the middle flat region which ranges from 5.0 m to 2.0 m ALLT, although they are found to the low tide datum. These species migrate following food sources, many are carnivorous including Eteone spp. that have been observed feeding on the spionids. There is a gradual change in fauna from highly motile species to more sedentary ones towards the lower end of the middle flat. A variety of polychaetes are commonly found, including Microspio theeli, Spio jilicomis, Spio


goniocephala, Eteone flava and Eteone longa and Capitella spp. The area experiences 25% to 45% exposure, with meiofaunal densities ranging from 15 to over 700 organisms per 100 sq cm, the high densities resulting from the presence ofMicrospio.

The lower flat zone below 2.2 m ALLT has the highest species diversity and species richness of the tidal flats. Increasing numbers of sedentary infauna such as Mya truncata, larger sedentary and tubiculous polychaetes like Euchone analis, Chone infundibuliformis and Laonome kroyeri and the anemone Tealia sp. contribute to the increasing biological richness. These organisms are all filter feeders, that appear to require at least 90% marine inundation in order to survive. Motile species are also common at lower sites including the carnivorous polychaetes Harmothoe imbricata and Phyllodoce groenlandicus and deposit feeders such as Travisiaforbesii and Chaetozone setosa.

Exposure indices reveal 2 critical tidal heights around 4.0 m and 7.5 m ALLT that coincide with some biological boundaries. The boundary between motile and less motile (sedentary) fauna occurs around 4.0 m ALLT and 7.5 m ALLT marks the limit of most marine flora and fauna, with the rare exception of Fucus evanescens. In general, it appears that low tidal species live over a number of seasons, whereas in the upper flats recolonization occurs each year. Frequent exposure to freezing temperatures and the removal of upper layers of sediment in which many live preclude their survival over the winter. In addition, the influx of freshwater during melt restricts the movement oflarvae and the distribution of marine invertebrates particularly in the upper flat regions, subjected to lengthy periods of freshwater influence. Some sedentary molluscs and polychaetes are particularly sensitive to changes in salinity. This was observed in early August in 1980, when significant numbers of Cyrtodaria kurriana and Hiatella arctica were killed after an influx of drift ice settled on the narrow fringe of tidal flat on the south side ofKoojesse Inlet. Many of the dead molluscs were found in situ, suggesting the influence of freshwater. A fewer number of shells showed evidence of crushing by the ice.

Anthropogenic effects are most apparent on the flats nearest the town centre, at sites proximal to the sewage lagoon, dump sites and areas graded by the hamlet for vehicular traffic. Reduced species richness, diversity and evenness values characterize the invertebrate population at these sites. Elsewhere, communities appear relatively stable with no statistical significant difference in the invertebrates from the 1980s (Samuelson 1997). The most heavily disturbed zone occurs near the sewage lagoon outlet and is devoid of invertebrates. A less disturbed zone, some 400 m from the outflow is dominated by increasing numbers of species known to be pollution tolerant, including the polychaetes Capitella 'capitata' sp. and Polydora quadrilobata.

7. Discussion

The tidal flats at the head of Frobisher Bay are similar to those found in other subarctic and arctic locales. Northern tidal flats exhibit many of the classic features found on temperate tidal flats. But the combination of ice action and the considerable duration of ice cover results in some notable differences. Ice cover for up to 9 months of the year reduces the role of wave and current processes on sedimentary zonation, such as the

Tidal Flats of Iqaluit, Nunavut 231

shoreward fining sequence familiar in temperate locales. The intertidal ice cover protects underlying surfaces from waves, tidal currents and offshore ice action. But it also can contribute to substantial erosion, transpo~ and deposition of materials throughout the tidal flats, disrupting the sorting by wave and current action. Ice freezeup results in erosion of the flat surface and incorporation of a wide range of sediment sizes as large as boulders. Movement of the ice results in the transport and deposition of the sediments throughout the intertidal zone. Subarctic tidal flats subject to ice action are characterized by weak sorting trends, especially after ice breakup when the sorting is very poor. As the open water period progresses, shoreward flDing and sorting of the surficial sediment becomes more apparent. The ability of the ice to transport boulders has resulted in the formation of boulder ridges, boulder garlands, boulder pavements, boulder mounds, boulder barricades and sediment mounds in specific zones that characterize subarctic tidal flats.

The large tidal range in Koojesse Inlet also affects the flats, specifically in the generation of high current velocities. Tidal currents erode and transport coarse sands to fine clays, as well as rafting ice floes that can contain sediments, further disrupting normal sediment trends. Overall there appears to be net erosion across the flats as fme sediments are gradually deposited subtidally or in protected, isolated zones in the upper flats.

Similar to temperate areas, arctic tidal flats also exhibit zonation in the distribution of species due to the varied conditions encountered from high to low tide. The lower species richness and diversity are typical of arctic locales, where variable seasonal conditions in marine temperatures and salinities result in a harsh and dynamic environment that precludes more temperate organisms.

The flats in Koojesse Inlet exlnbit evidence of anthropogenic effects. The construction of breakwaters is changing current and sedimentation patterns. Restrictions to normal water flow results in an increase in fme deposition particularly notable in the lee of the breakwaters on the south side near the sewage lagoon outflow. Scraping of the flats in the construction of the breakwaters, as well as to expedite sea lift further disrupts the morphology, sediment sorting and biota present on the flats. Harvesting of clams such as Mya truncata that are long-lived and slow to mature has affected the distribution of these species around the inlet, to the point that they are no longer found at some former sites.

Weather plays an important role in the marine conditions and processes of the flats. Thus the prospect of climate change could have significant consequences to the processes and morphology of the tidal flats. Climate warming could result in thinner and later sea ice development and a reduction in the duration of ice cover. As a result, the influence of wave action and current action will increase and shoreward flDing should be enhanced. It is likely that sediment and boulder mounds and ice micro-relief features will be much reduced over time. The size of boulders capable of transport will decrease and overall boulders will be much less mobile. Present morphological features such as bouldery flats, boulder ridges and barricades will likely remain as fairly permanent features, with little additional enhancement. The composition of biota may also change as more temperate species inhabit the flats, and inhibit the distribution of colder species whose range will retreat northward. Changes in invertebrate species diversity and richness on arctic tidal flats could have important ramifications for organisms higher on the food


chain including present arctic fish and mammal species, migrating birds and ultimately the human population who so relies on present marine resources.

Tidal currents move sediment, biota, raft ice floes with their sediment or boulderrich load, and affect water temperature and salinity at the bed. All of these factors limit the distribution of flora and fauna. Exposure indices generated from tidal data, reveal 2 critical tidal heights around 4.0 m and 7.5 mALLT. The boundary between motile and less motile (sedentary) fauna occurs around 4.0 m ALLT and 7.5 m ALLT marks the limit of most marine flora and fauna, with the rare exception of Fucus evanescens. The effectiveness of wave action is restricted to ice free periods. Thus the sorting of sediment by wave action is limited to a few months each year. In addition, the presence of ice floes also dampens wave action. The longer the ice free period the greater the degree of sorting of the intertidal deposits after the disruptive effects of ice cover. Field observations over four summer seasons suggest that, infrequent, but high intensity wave action is more significant than frequent but low intensity wave action on the Koojesse flats.

The tidal flats are divided into six morphological zones and three closely associated biological zones. Fauna of the upper flat and beach areas are hardy and freshwater tolerant. Many are highly motile, opportunistic species that recolonize the area after ice breakup. These zones are the most intensively affected by ice action. There is a gradual change in fauna from highly motile species to more sedentary ones towards the lower end of the middle flat. Boulders mounds, boulder barricades and boulder pavements are found within this zone. Below 2.2 m ALLT the tidal flat has smaller and decreasing numbers of boulders to form the graded flat. It also has the highest diversity and species richness of the flats due to increasing numbers of sedentary infauna such as Mya truncata, larger tubiculous polychaetes and the anemone Tealia sp.

Intertidal communities appear relatively stable, except at sites close to the sewage lagoon outflow, graded areas at the head of the bay and beaches near the town centre with vehicular traffic Elsewhere statistical tests show no significant difference in the invertebrates from the 1980s.

References

Bird, E.C.F. and Schwartz, M.L. (eds.) (1985) Arctic Canada. In: The World'S Coastline. Van Nostrand Reinhold Co. Inc., New York: 241-251.

Blake, W. Jr. (1966) End moraines and deglaciation chronology in Northern Canada. with special reference to Southern Baffin Island. Geological Survey of Canada. Paper 66-26; 31 p.

Canadian Ice Services (CIS) (1998) Ice formation and deterioration records for Koojesse Inlet and ice thickness records. Environment Canada Client Services, Ottawa.

Crane, R.G. (1978) Seasonal variations of sea ice extent in the Davis Strait-Labrador Sea area and relationships with synoptic-scale atmospheric circulation. Arctic, 31(4): 434-447.

Dale, J.E. (1982) Physical and biological zonation ofsubarctic tidal flats at Frobisher Bay, southeast Baffin Island. Unpublished M.Sc. thesis, McMaster University, Hamilton, Ontario: 284 p.

Dale, J.E. (1992) The relationship between the physical environment and benthic faunal communities in Pangnirtung Fiord, Baffin Island, N.W. T. Unpublished Ph.D thesis, Queen's University, Kingston, Ontario: 424 p.

Dalrymple, R.W. (1992) Tidal depositional systems. In Facies Models: Response to Sea Level Change, Walker, R.G. and James, N.P. (eds.). Geological Association of Canada, St.John's Newfoundland: 195-218.

Davis, R. (1992) Depositional Systems. 2/1d Ed. Prentice Hall, New Jersey: 604 p.


Department of Fisheries and Oceans (1995) Canadian Tide and Current Tables 1996. Canadian Communication Group Publishing, Ottawa, Ontario: 45 p.

Dionne, J-C. (1973) La notion de pied de glace, en particulier dans I'estuaire du Saint Laurent. Cahiers de Geographie du Quebec, 41:221-250.

Dionne, J-C. (1984) An estimate of ice-drifted sediments based on the mud content of the ice cover at Montmagny, middle St. Lawrence Estuary. Marine Geology, 57: 149-166.

Dionne, J-C. (1988) Characteristic features of modern tidal flats in cold regions. In Tide-influenced Sedimentary Environments and Facies, DeBoer, P.L., VanGelder, A. and Nio, S.D. (eds.) D. Reidel Publishing Company, Dordrecht, Holland: 301-332.

Dionne, J-C. (1993) Sediment load of shore ice and ice rafting potential, Upper Sl Lawrence Estuary, Quebec, Canada. Journal o/Coastal Research, 9(3): 628-646.

Drake, U. and McCann, S.B. (1982) The movement of isolated boulders on tidal flats by ice floes. Canadian Journal 0/ Earth Sciences, 19: 748-754.

Dunbar, M.J. (1951) Eastern Arctic Waters. Bulletin o/Fisheries Resource Board o/Canada, 88: 131 p. Dyke, A.S. and Prest, V.K. (1987) Late Wisconsinan and Holocene History of the Laurentide Ice Sheet.

Geographie Physique et Quaternaire, XLI (2): 237-263. Ellis, D.V. (1955) Some observations on the shore fauna of Baffin Island. Arctic, 8: 224-235. Ellis, D.V. (1960) Marine infaunal benthos in North America. Arctic Institute 0/ North America, Technical

Paper 5: 53 p. Ellis, D.V. and Wilce, R. T. (1961) Arctic and subarctic examples of intertidal zonation. Arctic, 14: 224-235. Environment Canada (1993) Canadian Climate Normals 1961-1990: Yukon and Northwest Territories.

Canadian Climate Program, Ottawa, Ontario: 31 p. Environment Canada (1980, 1981, 1995, 1996) Monthly meteorological summary, Iqaluit, Northwest

Territories. Evans, G. (1965) Intertidal flat sediments and their environments of deposition on the Wash. In Holocene Tidal

Sedimentation, BenchmarkPapers in Geology, vol. 30, DeVries-Klein (ed.) Dowden, Hutchison and Ross Inc., Stroudsburg, Pennsylvania: 80-117.

Gilbert, R. (1983) Sedimentary processes of Canadian Arctic fjords., 36: 147-175. Gilbert, R. and Aitken, A.E. (1981) The role of sea ice in biophysical processes on intertidal flats at Pangnirtung

(Baffin Island) NWT. Proceedings Workshop on Ice Action in Shoreline Evaluation (Rimouski) National Research Council, Ottawa:35-52.

Grainger, E.H. (1954) Polychaetous annelids of Ungava Bay, Hudson Strait, Frobisher Bay and Cumberland Sound. Bulletin 0/ Fisheries Resource Board 0/ Canada, 11(5): 507-528.

Hewitt, R. and Dale, J.E. (1985) Growth increments of modern Mya truncata_L. from the Canadain Arctic, Greenland and Scotland, In: Current Research, Part B, Geological Survey o/Canada, Paper 84-lB: 179-186.

Kellerhals, P. and Murray, J.W. (1969) Tidal flats at Boundary Bay, Fraser River Delta, British Columbia. Bulletin o/Canadian Petroleum Geology, 17(1): 67-91.

Kooistra, M.J. (1983) Geomorphology of subtidal and intertidal areas in the southwest of the Netherlands. Geologie en Mijnbouw, 62: 637-641.

Leech, S.R. (1998) The transport of materials by ice in a subarctic macrotidal environment, Koojesse Inlet, Southeast Baffin Island. Unpublished M.Sc. thesis, University of Regina, Regina, Saskatchewan: 196p.

Martini, LP. (1980) Ice effect on erosion and sedimentation on the Ontario shores of James Bay, Canada. Zeitschriji Fur Geomorphologie, 25(1): 1-16.

Martini, LP. (1991) Sedimentology of subarctic tidal flats of western James Bay and Hudson Bay, Ontario, Canada. In Clastic Tidal Sedimentology, Smith, D.G. Reinson, G.E. Zaitlin, B.A. and Rahmani R.A. (eds.) Canadian Society of Petroleum Geologists, Memoir 1(i: 301-312.

McCann, S.B., Dale, J.E. and Hale, P.B. (1981a) Subarctic tidal flats in areas oflarge tidal range, southern Baffin Island, Eastern Canada. Geographie Physique et Quaternaire, XXXV (2): 183-204.

McCann, S.B., Dale, J.E. and Hale, P.B. (1981b) Ice conditions and effects in a macrotidal subarctic environment, S.E. Baffin Island. Proceedings Workshop on Ice Action in Shoreline Evaluation (Rimouski) National Research Council, Ottawa: 105-1 15.

McCann, S.B. and Dale, J.E. (1986) Sea ice breakup and tidal flat processes, Frobisher Bay, Baffin Island. Physical Geography, 7(2): 168-180.


Miller, G.H., Locke, W.W. ill and Locke,G.W. (1980) Physical characteristics of the southeastem Baffin Island coastal zone. In The CotlStline o/Canada, McCann, S.B., (ed.) Geological Survey of Canada, Paper 80-10: 251-265.

NunavutImplementation Commission (NIC) (1995) Choosing a capital: asupplementary report o/theNunavut Implementation Commission, NIC, Iqaluit: 93 p.

Reineck, H-E. (1976) Drift ice action on tidal flats, North Sea. Revue de Geographie de Montreal, XXX (1-2): 197-200.

Rosen, P.S. (1979) Boulder barricades in Central Labrador. Journal 0/ Sedimentary Petrology, 49(4): 1113-1123.

Rygg, B. (1985) Distribution of species along pollution-induced diversity gradients in benthic communities in Norwegian fjords. Marine Pollution Bulletin, 16(12): 469-474.

Samuelson, G.M. (1997) Benthic organisms as indicators of environmental change on subarctic tidal flats, Iqaluit, Baffin Island, N.W.T. Unpublished M.Sc. thesis. University of Regina, Regina, Saskatchewan: 197 p.

Smart, C.C. and Hale, P.B. (1987) Exposure and inundation statistics from published tide tables. Computers & GeOSCiences, 13(4): 357-368.

Smith, J.A., Millward, G.E., Babbedge, N.H., Attrill, M.J. and Jones, M.B. (1995) Changes in benthic community structure following construction of a harbour impoundment scheme. Netherlands Journal 0/ Aquatic Ecology, 19:449-457.

VanStratten, L.M.J.U. (1954) Sedimentology of recent tidal flat deposits and the Psammites Dir Condroz (Devonian). Geologie en Mijnbouw, 15: 25-47.

Wacasey, J. W. and Atkinson, E.G. (1975)Benthic invertebrates collected/rom Ungava Bay, Canada. Fisheries and Oceans, Ottawa: 68 p.

Wilce, R. T. (1959). The Marine Algae o/the Labrador Peninsula and Northwest Newfoundland (Ecology and Distribution). National Museum of Canada, Bulletin 158: 103 p.

Index

Abandoned river channels, 162 Ablation, 27,31,147,215 Above low, low tide (ALLT), 206,207,210,212,222,224-227,229-231,233 Abrasion, 145,147,174,187,192,198 Abrasion platforms, 174 Aeolian, 3, 10,64,73,77,79,81,84, 141, 143, 152, 153, 156-158 Aeolian deposits, 79,81 Aggradation, 73,77-79,81,82,84,87,93,94,99, 101,102,104, 106,108,110-112,114,

157 Aggradational sequences, 64 Alaska, 43,56,57,61,103,105,117,122,139,142,145,157,160-162, 184, 185 Albedo, 168,173,184,222 ALLT, see Above low, low tide Alluvial fan, 80 Anemone Tealia sp, 206,231,233 Antarctica, 142, 157 Anthropogenic effects, 206,231,232 Apex River, 217,219 Aquac1ude, 182,184 Arctic deltas, 159,160, 172 Argentina, 142 Attrition, 187 Avalanche boulder tongues, 6 Baffin Island, 132, 158,207,233-235 Balanus balanoides, 207,230 Ballycatter, 218 Bank collapse, 94, 112, 160 Barnacles, 192 Barriers, 63-65,68,70,77,78,80-83,86-88, 145 Barrow, Alaska, 161 Basal debris load, 16, 36 Beach, 117,145,153,155,156,158,192,206,210,217,227,230,233 Beaufort Sea, 145, 148, 149, 151, 161, 173, 184, 185 Bedrock, 31,35,43,43,55,58,65,73,77,79,82,83,88,104,130, 136, 137, 146-148,

208,229 Bifurcations, 161 Bio-erosion, 187-189, 197,202 Biokarst, 187,203 Biological zonation, 206-208, 212, 215, 229, 233 Boreal, 2, 5, 138 Bottom-fast ice, 160 Bottomfast, 170, 172

236 Index

Boudinage, 22 Boulder pavements, 81,150,226,227,229,232,233 Boulder ridges, 224,226,227,229,232 Boulders, 35,71,73,77,79,81148,150-152, 157, 192,206,208,212,219,222,226,

227,229,232-234 Bouldery flat, 227, 229 Brandung, 73 Breakup, 148,149,158,160-163,168-170,172-174,176,180,185206, 207, 215,

218-220,222,226,227,229,230,232-234 Breakup flooding, 160,162,168,170,173,180 British Columbia, 10,88,93,94,101,102,105,106,108-110,112,115,117,118,142,

234 Brooks Range, 160, 161 Candles, 172 Carbonate, 59, 187, 189,202,204 Carolinian forest, 153 Caves, 145 Chile, 142 Chlorophyceae, 229 Circular eddies, 172 Cirque glaciers, 14, 15 Cirques, 6, 8, 36, 58 Cliffed shorelines, 227 Cliffs, 9, 73, 141, 145, 157 Climate change, 9,44,45,60,64,65, 83, 84, 118, 138, 139, 145,232 Climate warming, 136, 138, 232 Climatic sensitivity, 122 Coastal dunes, 141, 143, 152, 153 Coastal karsts, 186, 189, 190, 202 Coastal or littoral karren, 187 Coastal karst model, 186 Coastal profile, 194 Cold regions shorelines, 141,144, 145, 152 Cold coastal, 2, 7, 141, 142, 145, 152, 153 Cold coastal dunes, 141,152 Cold coasts, 142-145,149,157 Cold regions coasts, 141, 142 Cold coasts, 5 Colville River, 160, 161, 166, 167, 180, 182-185 Complete freeze over, 219 Conductive, 126 Constructional landforms, 10,12,27,77-79,85 Continuous/discontinuous transition, 137 Continuous permafrost, 122, 123, 132, 133, 137, 138, 160, 161, 167 Convective, 126

Index

Corrosion or water-layer weathering, 187 Craters, 219,222 Crescentic sand waves, 217 Current velocities, 213, 215, 229, 232 Currents, 5, 145, 148, 149, 151, 206,212,215,223,226,232,233 Cyrtodaria kurriana, 229-231 Debris content of the ice cover, 141,151 Debuttressing, 87 Deflationary features, 81 Degree-day index, 132 Deltas, 5,77,79,85, 159-161, 172,207 Denivation forms, 141, 145, 152, 157 Dense peat, 180 Denudation, 187,191,192,198,199,202 Deposit feeders, 231

237

Deposition, 12,28,31,33-35,39,41,44,48,59,75,77-79,81,86,106, 149, 152, 157, 161,170,175,193,198,226,232,234

Diamicton, 12,21-23,26,30-32,36,40 Diffuse geographical transition, 137 Discontinuous permafrost, 123,124,132-134, 136-138, 143 Dissolution of soluble components, 187 Dissolutional potholes, 191 Distributaries, 77, 159, 161, 162, 167, 168, 170, 175, 184 Distributary banks, 162 Diversity, 7,86,206,231-233,235 Diversity and evenness, 206, 231 Drainage loci, 172 Drumlins, 6,41,43 Dump sites, 206, 231 Dunes, 8, 79, 81, 141, 143, 145, 152, 153, 155, 157 159, 160, 162, 165, 168, 169, 174,

182 Eastern Foxe Basin, 150 Eastern Hudson Bay, 152 Ebb tides, 212,215 Ebb velocities, 212 Ebb tidal currents, 215 El Nino, 44 Embayments, 145,148,227 Emerging shorelines, 151 Englacial channels, 19 Ephemeral, 40,164,184 Epigenetic gorges, 83 Equilibrium, 9,46,99, 127, 135, 186, 187,200 Erratics, 89, 170 Eskers, 6

238

Eu-littoral, 187,202 Europe, 43, 142, 156,207 Eustatic sea level rise, 189,207 Fan terrace, 79, 82 Fetch length, 219 Filter feeders, 231 Fines fiat, 227 Fining shoreward sequences, 207 Floating ice, 5,152, 156, 160, 170 Fluted terrain, 35

Index

Foliation-parallel ridges, 12, 27, 30-32, 35, 36,40 Fort Vermillion, 135 Fort Reliance, 132 Fort McMurray, 135 Frobisher Bay, 158,207,208,215,222,227,229,231,233,234 Frost riving and plucking, 222 Frostmounds, 160 Frost action, 141-150, 158, 187, 194, 198,202 Frost heaving, 147, 149 Frost weathering, 146, 148 Frost shattering, 146, 147 Frost wedging, 146, 147 Frost weathering, 203 Fucoid algae, 192 Fucus evanescens, 206,229-231,233 G.K.Gilbert, 100 Geometrical ridge networks, 27, 30, 36 Gilber4 99,100,105,117,152,157,207,222,234 Glacial disturbance, 103,106,108,109,111,112,116 Glaciallegacy, 88, 186, 189, 191, 192, 197,200 Glacier drainage system, 46, 48 Glaciofluvial debris, 15 Glaciofluvial sediment, 19,20,22,23,32,42 Gouge marks, 172 Gouging, 226, 227, 230 Graded fiat, 225,227,233 Grass tussock vegetation, 168 Great Lakes Basin, 152, 153 Great Lakes, 142, 143, 152, 153 Green algae, 192 Greenland,6, 10, 19,43, 142,203,234 Greenland Ice Sheet, 6, 19 Ground ice, 44 Ground thermal properties, 125 Groundice,3,6,145,149,160,169

Ground thennal conditions, 122,124 Gubik Fonnation, 181 Gyres, 212,226 Heat transfer, 126, 127 High energy events, 219

Index

High-AIctic, 11,12,15,16,32,35,39,41,43 HinuUaya,10,45,88-91 Hudson Bay, 122,123, 132147, 152, 158,234 Hummocky moraines, 21, 38 Hydraulic action, 187,202

239

Ice, 2,3,5,6,8,9, 12-16, 19-23,27,28,30-32,35,36,39,44-53,55-58,60-6265,73, 77,79,87,89,103,104,106,108,112,127,134-136, 138, 141-153, 155-158, 159-162,164-170,172-176,179-181,184,185,187, 189, 199,202,203, 206-208,212-220,222,223,226,227,229-235

Ice action, 141, 142, 146-149, 157, 199,206-208,219,222,226,227,229-235 Ice breakup, 158, 161, 170, 180,206,207,215,220,222,229,230,232-234 Ice butts, 172 Ice cover, 6, 60, 141-143, 148-152, 157, 168, 170,206,207,213-216,219,222,231-

234 Ice floes, 147-150, 174,206,212,217,218,223,226,232-234 Ice foot, 218,219,222,223,226,227,229 Icejams, 5,172 Ice micro-relief, 229, 232 Ice push, 147,222,227 Ice rafting, 141, 149-151, 157,234 Ice wedges, 160,166,169,174,175,179-181 Ice-cored moraine, 23,48, 49, 53, 58 Ice-dammed lake, 48-50,55,56,61 Ice-free period, 206 Ice-made depressions, 150 Ice-marginal, 12, 16, 21, 32,4246,48, 108 Ice-rafted, 22, 150, 151 If = freezing index for air temperature, 128 Impurities, 166, 189, 192, 197 Inselberg-type steep hills, 189 Inshore floes, 223 Intertidal communities, 233 Intertidal coastal features, 206 Intertidal sea ice, 219 Inter-tidal zone, 186, 189, 191, 192, 194, 197, 198,200,202 Intra-dune depressions, 162 IPCC, 121, 138 Iqaluit, Nunavut, 205-207 Isostatic rebound, 8, 10, 145, 203, 207 Isostatic uplift, 186,189,194,197,200,202

240 Index

Isotherms, 122, 134, 136 It = thawing index for air temperature, 128 James Bay, 122-124, 150, 157, 158,234 Karakoram, 10,45,65,68,78,81,84,88-91 kf = thermal conductivity of ground (frozen), 128 Koojesse Inlet, 207,208,210,212,217,219,220,222,223,226,227,229-234 Kropotkin, 8 kt = thermal conductivity of ground (thawed), 128 Lacustrine deposits, 23,65,67,68,71,73,79,80,84,89 Lacustrine environments, 44, 46-48, 58 Lake Age, 8 Lake ice, 46, 151, 168,169 Lake tapping, 174, 175, 185 Landform assemblages, 3,6,7, 12,21,32,36,43,44,46,59,60, 121 Landform associations, 64, 68, 84, 86 Landform in equilibrium, 187 Landforms, 7,9,10, 11-13,27,32,35,36,39,40,42,43,45,59,60,64,65, 77-79,82,

85-87,93,94,99,103,108,113,117,138,139,153,157,158,184, 203, 204 Landslide barriers, 64,65,68 Landslide-fragmented river system, 63 Landsystems, 7,10 Large rotating gyres, 226 Late Wisconsinan, 10 1, 207, 234 Late-glacial, 65, 156, 108, 109 Lateral and vertical accretion, 150 Latitudinal gradient in MAAT, 124 Lena River, 160 Linear pond, 165 Lithologic conditions, 121, 124, 125, 127, 128 Little Ice Age, 51,52,55,57,58, 134-36 Little Ice Age moraine, 52 Littoral karren, 186, 187, 192, 194, 199-203 Littorina saxatilis, 207,230 Local ground temperature distributions, 124 Loess, 6,8,59,65,79 Longitudinal foliation, 12, 16, 18,20,30-32,34,40 Lower flat zone, 231 Mackenzie delta, 139, 161 Macroalgae, 228, 229 Macrofauna, 228, 229 Macrotidal, 206, 207, 234 Marine temperatures, 215,232 Mean annual air isotherm, 123, 134 Mean annual air temperature (MAAT), 122,126,127 Mean annual ground surface temperature (MAGST), 125,126

Index

Mean annual ground temperature (MAGT), 122 Mechanical action, 187,188, 198 Medial moraines, 12, 16,23,30,31,33,39 Medium well sorted sands, 219 Mesotidal, 206 Middle flat, 206,224,225,227,230,233 Middle St. Lawrence Estuary, 150-152,157,234 Mineralogy, 130 Mires, 5,10 Molluscs, 230, 231 Montane, 2, 5 Moraine complexes, 12,20-22,27,35,36,40 Morphological and biological zonation, 212 Motile and less motile (sedentary) fauna, 206,231,233 Mounds, 12,21,22,71,149,160,212,219,224,226,227,229,232,233 Mudflats, 159, 160, 162, 165, 170 Mussels, 192 Mya truncata, 206, 229-234 Mytilus edulis, 207 Nechelic Channel, 161, 170, 181, 182 Neoglacial, 20,55,56,59,61,62,83,87,111,122 New England (Maine), 142 Nf function, 130 Nival, 2,6,79,168, 126, 127, 129, 130. 132-134. 137. 138 Nival offset map. 134 Nonvegetated sand dunes. 168 Northern Alaska. 139, 160 Northern Saskatchewan, 134.136 North Slope of Alaska. 160-162 Nova Scotia. 142, 148, 156 Nuiqsut, 161 Offset effect. 130, 135. 136 Offset formulations. 129 Pack ice, 143, 146, 158, 222 Paleoclimatic indicators, 152 Palsas. 6, 149 Paradigms. 10,43-45. 60. 121 Paraglacial. 3, 10,43.45,46.55.57-59.64,65, 79, 81. 84. 86-88, 90. 93. 106, 110,

111. 117, 121. 141 Paraglacial transitions, 65 Peace River, 135, 136 Peatlands. 5, 124. 133-136, 138 Perched lake, 182-184 Perched pond. 184 Perennially frozen ground. 123

241

242 Index

Periglaciru, 2,6,8-10,36,43,44,89,121,139,141,152,157,158,161 Periglaciru shorelines, 141 Permafrost, 3, 6, 8,13,36,44, 121-139, 143, 145, 147, 149, 156, 157, 159-161, 164,

166-168,175,179-182,184,185 Permafrost and perched lakes, 182 Permafrost-climate, 121, 125-127 Permafrost-free, 122, 125 Peyto Glacier, 48, 49 P-forms, 189, 191, 192 Phytokarren, 187 Phytokarst, 187,203 Pinery Provinciru Park, 153, 154, 156 Pingos, 160 Planktonic settlement, 215 Platforms, 141,143, 145-148, 157, 158,174 Pleistocene, 12,27,35,36,89,94,105,106,111,112,117,146,157,202,207 Polar-maritime environment, 13 Pollution tolerant, 231 Polychaetes, 206,230,231,233 Polycyclicru erosion, 189 Polygonru patterns, 149 Polythermal type, 13 Pore ice, 166 Post-break flooding, 160 Post-emergence denudation, 192 Pre-breakup flood period, 160 Presquile Provincial Park, 153 Pressure ridge cracks, 160 Pressure ridge, 73, 75, 160, 167 Priapulids, 230 Prince Edward Island, 152 Pristine striae, 189 Profile down, 195 Proglacial, 12, 18,21,23,27,30-33,36,39,42,43,45,46,51,52,55-57,59,60, 121 Proglaciallakes, 56 Proglaciru outwash, 23 Push moraines, 20, 32, 40, 43 Putu Pond, 183, 184 Quaternary heritage, 141 QuebeclLabrador, 132 Raft ice floes, 206,233 Rain-generated floods, 180 Raised shorelines, 207 Rejoinings, 161 Relict permafrost, 134-136,168

Rbodophyceae, 229 River bars, 168, 170 River flooding, 172, 173 River-ice breakup, 161, 180 River ice depression, 168 River terraces, 64, 68, 83

Index

Rock avalanche, 63,67,68,70, 71, 73, 75-83,86,87,89-91 Rockfall, 12,15,27,30,39,90 Rock skerries, 189 Rock slide-rock avalanches, 68, 71 Rogers Island, 207,208,219 Russia, 142, 145 Salinity, 171,206,215,216,231,233 Salt weathering, 186, 192, 198 Sand bars, 159, 160, 162, 165 Sand dunes, 79, 150, 152, 157, 159, 162, 165, 168, 169, 174, 182 Sastrugi, 165 Scaling factor, 128 Schefferville, 126,132,139 Scour holes, 160 Sea anemones, 229 Sea ice, 142, 145, 158, 160, 169, 172,173, 176,215,219,222,232-234 Sea level rise, 8, 189,207 Seasonally frozen ground, 123 Seawater, 170,219 Sediment assemblages, 3, 7, 10, 11, 12,20,30,39 Sediment bands, 219 Sediment fans, 65, 67, 73, 80, 81, 84, 86, 87 Sediment in the intertidal ice, 222 Sediment load, 89, 157, 172, 173,219,234 Sediment response time, 100, 106 Sediment yields, 68, 102, 104 Sedimentary characteristics, 90, 205-207 Sedimentary zonation, 226, 231 Segregated ice, 149 Semi-diurnal tides, 206, 207

243

Sequence, 12,15,39, 77,82-85,87? 94, 97, 101, 106, 112, 192, 193,220,222,232 Sewage lagoon, 206,229,231-233 Shore erosion, 146, 156 Shore ice, 146-151,157,187,234 Shore ice action in tidal marshes, 149 Siberia, 142,145 Sinuous crested ripples, 217 Sipunculids, 230 Slumping blocks, 181

244

Small current ripples, 212 Small mounds, 149

Index

Snow, 3,5,6,12,15,16,39,47,50,51,75,79,121,124-129,132, 133, 135-138, 151-156,159-162,164-169,179,184,219

Snow core, 154 Snow cover, 6, 124-129, 132, 133, 135-138, 160, 161, 168, 169 Snow drifts, 165, 167-169, 179 Snow patches, 156 Soil conditions, 124 Soil latent heat, 137 Sorting, 94,108,206,217,221,232,233 Species richness, 206,231-233 Spillways, 71-73,79,82 Spits, 145 Spitsbergen, 12,14,41,43,142 St. Elias Mountains, 45,46,56,57,59,61 St. Lawrence Estuary, 147,148, 150-152, 157,234 St-Onge, 135 Stacks, 145 Stagnant ice, 21,51,58 Standing waves, 212 Steep intertidal gradients, 227 Stranded intertidal ice, 222 Strandflat, 146, 158, 189,202,203 Strudel, 172, 173, 184 Strudel scour, 173, 184 Subaerial forms, 160,162 Sub-Antarctic Islands, 142 Subaqueous delta, 173, 174 Subarctic, 5, 113, 152, 153, 158,206,207,212,231-235 Subarctic climate, 207 Subarctic marine conditions, 207 Subarctic tidal flats, 158,212,232-235 Sublimation, 152, 153, 158 Submerged troughs, 174 Sub-tidal zone, 200 Superimposition, 43,46,59, 93, 121 Supply of sand, 153 Supraglacial, 15,16,18,20,23,27,30,32,36,39,40 Supraglacial debris trains, 27, 40 Supra-littoral zone, 186,192,199 Surface pitting, 191 Surges, 8, 15, 16,36,42,43,57,60,61 Surge-type, 15,41-43 Suspended load, 160

Index

Svalbard, 11-16, 18-21,27,30,34-36,38-43, 186-188, 194, 199,202 Swales, 162 Temperate tidal flats, 206, 222, 226, 231 Temperature regime, 126, 132 Terrace flank depressions, 162 Terraces, 64,65,68,71, 76, 79, 80, 82, 83,90, 145 Thalweg, 83, 170 Thennal conductivity ratio, 128, 131, 136, 137 Thennal erosion, 51,173 Thennal offset, 127, 129-133, 135-137 Thennal offset results in pennafrost temperatures (TTOP), 127 Thennalregime, 13,40,41,43,138,139,156 Thenno-abrasion, 145 Thennoerosional niche, 179,180 Thennoerosional niching, 160 Thennokarst depressions, 149, 150 Thennokarst lakes, 160 Thrust moraine complex, 20 Thrusting, 12, 15, 18-23,30-32,34,36,38,40-42 Tidal currents, 148, 151,206,215,223,226,232,233 Tidal current velocity, 215 Tidal drainage, 229 Tidal flat boundary, 222 Tidal flat morphology, 206

245

Tidal flats, 141, 145, 150, 151, 157, 158205-208,212-216,219,222,223,226,229-235

Tidal height, 211, 212 Tidal marsh development, 148, 149, 151 Tidal oscillations, 223 Tidal processes, 206, 207 Tidal wedges, 212,227,229 Tides, 142,145,148,157,206,207,212,215,223 Tidewater glaciers, 22, 143, 145 Time scales, 56,59,93,97,111 Time transgressive sedimentation, 84 Topography, 43, 103, 124, 149 Transition, 3,5,8,9, 19,32,40,43,44,46,51,58-60,64,87, 121, 123, 124, 129, 133,

135,137,138,186,189,200 Transitions, 2,3,5,8,9,43-46,48,56,59,65, 121, 123, 130 TTOP model, 127, 130, 137, 138 Tubiculous polychaetes, 206,231,233 Tundra-type vegetation, 161 Unconsolidated sediments, 5, 206 Under-ice flow, 170 Ungava Bay, 147, 148, 150, 156,234,235

246 Index

Vegetation, 5,8,9,21,28, 106, 124-126, 128, 139, 150, 152, 153, 161, 168, 184 Very bouldery flat, 227,229 Water temperature, 206, 215, 233 Wave action, 6, 9,142,144,174,187, 189, 194,206,215,217,219,223,227,229,

232,233 Waveactionlq~ng, 187 Wave processes, 207 Waves, 5,61, 106, 142, 145, 146, 148, 149, 151, 188, 198,212,215,217,219,232 Weather, 232 West ofIreland, 190, 199 VVluUebacks, 189, 191 Wind, 6,81,103,152,165,169,184,188, 198, 199,215,219,223,226,230 Younger Dryas, 35, 36, 38, 40-42, 203 Yukon, 19,43,45-47,52,56,61, 121, 138,234

landscapes of transition: landform assemblages and transformations in cold regions

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