Glacigenic Sediments

DEVELOPMENTS IN SEDIMENTOLOGY 49 Glacigenic Sediments FURTHER TITLES IN THIS SERIES VOLUMES 1-1 1, 13-15, 17, 21-25A...

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DEVELOPMENTS IN SEDIMENTOLOGY 49

Glacigenic Sediments

FURTHER TITLES IN THIS SERIES VOLUMES 1-1 1, 13-15, 17, 21-25A4,27,28,31,32 and 39areoutof print 12 R. C.G. BA THURST CARBONATE SEDIMENTS AND THEIR DlAGNESlS 16 H.H. RIEKE Ill and G. V. CHILINGARIAN COMPACTION OF ARGILLACEOUS SEDIMENTS 18 G. V. CHlLlNGARIAN and K.H. WOLF, Edidors COMPACTION OF COARSE-GRAINED SEDIMENTS 19 W. SCH WARZACHER SEDIMENTATION MODELS AND QUANTITATIVE STRATIGRAPHY 20 M.R. WALTER, Editor STROMATOLITES 258 G. LARSENand G. V. CHILINGARIAN, Editors DIAGENESISIN SEDIMENTS AND SEDIMENTARY ROCKS 26 T. SUDA and S. SHIMODA, Editors CLAYS AND CLAY MINERALS OF JAPAN 29 P. TURNER CONTINENTAL RED BEDS 30 J.R.L. ALLEN SEDIMENTARY STRUCTURES 33 G.N. BATURIN PHOSPHORITESON THE SEA FLOOR 34 J.J. FRIPIAT, Editor ADVANCED TECHNIQUES FOR CLAY MINERAL ANALYSIS 35 H. VAN OLPHEN and F. VENIALE, Editors INTERNATIONAL CLAY CONFERENCE 198 1 36 A . IJIMA, J.R. HElNand R. SIEVER, Editors SILICEOUS DEPOSITS IN THE PACIFIC REGION 37 A . SlNGERand E. GALAN, Editors PALYGORSKITE-SEPIOLITE:OCCURRENCES, GENESIS AND USES 38 M.E. BROOKFIELD and T.S. AHLBRANDT, Editors EOLIAN SEDIMENTS AND PROCESSES 40 B. VELDE CLAY MINERALS-A PHYSICO-CHEMICALEXPLANATION OF THEIR OCCURRENCE 4 1 G. V. CHlLlNGARIA N and K. H. WOLF, Editors DIAGENESIS, I 42 L.J. DOYLE and H.H. ROBERTS, Editors CARBONATE-CLASTICTRANSITIONS 43 G. V. CHILINGARIAN and K.H. WOLF, Editors DIAGENESIS, II 44 C.E. WEAVER CLAYS, MUDS, AND SHALES 45 G.S. ODIN, Editor GREEN MARINE CLAYS 46 C.H. MOORE CARBONATE DIAGENESIS AND POROSITY 47 K.H. WOLFand G. V. CHILINGAR, Editors DIAGENEIS, 111 48 J. W. MORSE and F. F. MACKENZIE GEOCHEMISTRY OF SEDIMENTARY CARBONATES

DEVELOPMENTS IN SEDIMENTOLOGY 49

Glacigenic Sediments K. BRODZIKOWSKI Uniwersytet tbdzki, lnstytut Geografii Fizycznej i Ksztalrowania Srodowiska, Zakiad Geologii, Al. KoSiuszki 2 1, 90-4 18 t b d i , Poland and

A.J. VAN LOON P. 0. Box 1254, 680 1 BG Arnhem, The Netherlands

ELSEVIER Amsterdam

-

Oxford - New York -Tokyo

1991

ELSEVIER SCIENCE PUBLISHERS B.V. Sara Burgerhartstraat 25 P.O. Box 2 1 1, 1000 AE Amsterdam, The Netherlands Distributors for the United States and Canada: ELSEVIER SCIENCE PUBLISHING COMPANY INC. 655, Avenue of the Americas New York, N.Y. 10010, U.S.A.

ISBN 0-444-88307 (VOI. 49)

0 Elsevier Science Publishers BV., 1991 All rights reserved. No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publisher, Elsevier Science Publishers B.V./ Physical Sciences 2 3 Engineering Division, P.O. Box 330, 1000 AH Amsterdam, The Netherlands. Special regulations for readers in the USA -This publication has been registered with the Copyright Clearance Center Inc. (CCC), Salem, Massachusetts. Information can be obtained from the CCC about conditions under which photocopies of parts of this publication may be made in the USA. All other copyright questions, including photocopying outside of the USA, should be referred t o the publisher. No responsibility is assumed by the Publisher for any injury and/or damage t o persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. This book is printed on acid-free paper. Printed in The Netherlands

I dedicate this book to Jan D . de Jong, who was not only a very stimulating teacher i n both sedimentology a n d Quaternary geology, but who is also - with his wife Roni - a true friend o f m y family. T o m van Loon

I dedicate this book to the memory o f m y teacher and friend J e r z y Cegta who, during my first steps i n glacial geology, showed how passionate sedimentology can be. Krzysztof Brodzikowski

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VII

PREFACE Large ice sheets have covered the Earth's continents several times in the geological past. The most recent series of glaciations occurred during the last 2.5 million years. Huge areas, particularly on the northern hemisphere, were covered by ice during one or more phases of this period of glaciations. The ice and its meltwaters shaped the landscape, eroding the substratum and depositing the transported debris elsewhere in a complex system of interactive processes that were also influenced by parameters such as solar irradiation and the Earth's heat flux. The last ice age ended some 10,000 years ago, but even contemporary society in several countries still depends on the morphology and mineral resources left by the ice o r by the meltwater. The economy - agriculture, water supply, civil engineering, mining and various other branches - therefore profits from any knowledge that exists about the ice ages and their influence on the present-day landscape and surface layers. Various branches of science (geology, geomorphology and hydrology to mention only the most important) are also interested in the deposits that resulted from the ice covers. Both (fundamental and applied) science and economy-oriented disciplines have long carried out research in the field of glaciations; the results of this research were published mainly for the benefit of colleagues, and data available from other fields of research were not commonly referred to. This led to a somewhat chaotic terminology, especially with respect to the various processes that play a role during glaciation and deglaciation, but also concerning the types of deposits that were formed during such stages. This book aims primarily at providing those involved in fundamental or applied research in these fields with an overview of the various aspects concerned. A generally applicable terminology is proposed which should facilitate communication between scientists from several fields. Because the book is mainly devoted t o depositional processes and the resulting deposits, the approach and the terminology followed in this book are obviously founded strongly on sedimentology, the geological discipline that deals specifically with these phenomena. The book should be of help in describing the sediments involved, interpreting their genesis, establishing their extent and their mutual relationships, and thus in the reconstruction of the palaeogeographic development. It is the result both of research carried out by the authors

VIII together in Poland and in The Netherlands during the period 1979-1990, and of the investigations carried out by the authors individually in the course of many more years and in many more countries. Field work in a large open-cast mine in central Poland with a Quaternary overburden containing deposits from seven glaciations was of enormous help for obtaining a 3-dimensional picture of the geological structure of a glaciated area; the several square kilometres of sections provided by the quarry were (and still are) an almost inexhaustible source of data. Ongoing mining exposes ever new walls, allowing the authors to check the validity of models based on earlier observations. Other important study areas are situated in Canada, where one of the authors (K.B.) had the possibility to study the bluffs of Lake Ontario and the Fraser River canyon in British Columbia, thanks to the cooperation of Prof. Nick Eyles and Prof. A. Miall (University of Toronto); the other author (A.J.v.L.) could study glaciomarine deposits of Huronian age in the Cobalt area and similar deposits of Pleistocene age formed in the former Champlain Sea near Ottawa, thanks t o the help provided by Dr. D.R. Sharpe (Geological Survey of Canada). These areas provided important additional information on glacial sequences and facies associations. The field data form an important base for this book that is therefore a practical, rather than a theoretical treatise, although it also takes into account data from several hundreds of publications by numerous authors. The authors are much indebted to all those who contributed t o the book in one way or another. Special thanks are due to Mr. S. Drozdowski and Mr. J. Kowalski who allowed systematical investigations in the open-cast mine of Befchatcjw (near L6di, central Poland). Prof. S. Kozarski (Poznaii University), Prof. H. Klatkowa (Lddd University), Dr. L. Kasza (Wroctaw University), Dr. Ruszczyiiska-Szenajch (Warsaw university), Prof. N. Eyles and Prof. A. Miall (University of Toronto) and Dr. L. Eissmann (Leipzig University) helped solving problems in the course of several discussions; Dr. D. Krzyszkowski (Wroctaw University) and Dr. T. Zieliiiski (&ask University at Sosnowiec) helped the authors with the literature search concerning fluvioglacial and melt-out deposits. Mr Marek Ciennik (Lodi University) helped much with checking manuscript and references for spelling errors, omissions, etc. Dr. M.-L. Desbarats-Schonbaum was a stimulating corrector of the English language, who contributed probably more t o the book than she is aware of. Mr. D. Benn (University of St. Andrews, Fife), M.Sc., made helpful suggestions regarding the classification of glacigenic sediments. July 1990

Krzysztof Brodzikowski Tom van Loon

IX

CONTENTS PREFACE INTRODUCTION

1

EARLY LITERATURE AND PROBLEMS IN TERMINOLOGY From Diluvium to diamicton Landforms versus sediments Ambiguous descriptions Nomenclature SOURCES USED SCOPE OF THE BOOK FRAMEWORK OF THE BOOK

3 3 5 13 13 15 16 18

G E N E R A L CHARACTERISTICS OF GLACIGENIC SEDIMENTATION

19

DEPOSITIONAL PROCESSES IN THE GLACIGENIC ENVIRONMENTS Sedimentation by ice Subaqueous sedimentation Aeolian sedimentation Deposition from mass movements CHARACTERISTICS OF GLACIGENIC SEDIMENTATION The source of glacigenic sediments Grain size of glacigenic sediments Transgressive and regressive tendencies Influence of ice dynamics and extent upon sedimentation The character of the sediment input THE INFLUENCE OF CLIMATE ON GLACIGENIC SEDIMENTATION Role of temperature Role of precipitation Role of winds THE INFLUENCE OF ENDOGENIC FACTORS Vertical movements of the Earth’s crust Earthquakes The geothermal heat flux THE SEDIMENTARY FACIES Facies analysis FACIES INTERPRETATION Lithological characteristics Textural characteristics Occurrence Depositional mechanisms Sedimentary cycles GLACIGENIC FACIES MODELS TERMINOLOGY AND USAGE OF SYMBOLS Classification system used in this book

20 20 21 24 25 27 27 29 29 31 32 33 33 36 40 41 42 43 45 47 48 50 50 57 58 60 63 64 65 66

X

T H E SYSTEM O F GLACIGENIC DEPOSITIONAL ENVIRONMENTS

83

THE GLACIGENIC ENVIRONMENTS AS AN ENERGETIC ENTITY DEPOSITIONAL SEDIMENTARY ENVIRONMENTS AND THEIR GENERAL MODEL The glacial environment The periglacial environment

83 87 87 91

GLACIGENIC FACIES AND THEIR CHARACTERISTICS

93

THE MELTING-ICE F A C E S THE GLACIOFLUVIAL FACIES THE MARINE BOTTOM-CURRENT F A C E S THE GLACIODELTAIC FACIES THE GLACIOLACUSTRINE FACIES THE AEOLIAN F A C E S THE MASS-TRANSPORT F A C E S

96 101 106 107 111 116 120

TABLES O F T H E CONTINENTAL AND MARINE GLACIGENIC ENVIRONMENTS, SUBENVIRONMENTS, FACIES AND DEPOSITS

127

T H E CONTINENTAL SUPRAGLACIAL SUBENVIRONMENT (I-A) AND ITS DEPOSITS

131

SUPRAGLACIAL CONDITIONS ON ACTIVE ICE SUPRAGLACIAL CONDITIONS ON PASSIVE ICE GEOLOGICAL PROCESSES IN THE SUPRAGLACIAL SUBENVIRONMENT Sedimentation processes and supraglacial facies patterns DEPOSITS OF THE SUPRAGLACIAL MELTING-ICE FACIES (I-A-1) Supraglacial melt-out complexes (I-A-1-a) Supraglacial ablation tills (I-A-1-b) Supraglacial ice-raft deposits (I-A-1-e) DEPOSITS OF THE SUPRAGLACIAL FLUVIAL FACIES (I-A-2) Supraglacial fluvial complexes (I-A-2-a) Supraglacial tunnel-mouth deposits (I-A-2-c) Supraglacial stream deposits (I-A-2-d) Supraglacial sheet- and streamflood deposits (I-A-2-e) DEPOSITS OF THE SUPRAGLACIAL DELTAIC FACIES (I-A-3) Supraglacial deltaic complexes (I-A-3-a) Supraglacial deltaic topsets (I-A-3-b) Supraglacial deltaic foresets (I-A-3-c) Supraglacial deltaic bottomsets (I-A-3-d) DEPOSITS OF THE SUPRAGLACIAL LACUSTRINE FACIES (I-A-4) Supraglacial lacustrine complexes (I-A-4-a) Supraglacial lake-margin deposits (I-A-4-b) Supraglacial lacustrine bottomsets (I-A-4-c)

132 134 136 137 138 138 143 148 150 152 154 157 161 164 165 167 170 174 177 179 181 184

XI

DEPOSITS OF THE SUPRAGLACIAL AEOLIAN F A C E S (I-A-5) Supraglacial drift sands (I-A-5-b) DEPOSITS OF THE SUPRAGLACIAL MASS-TRANSPORT F A C E S (I-A-6) Supraglacial subaerial mass-transport deposits (I-A-6-a) Supraglacial crevasse deposits (I-A-6-b) Supraglacial subaqueous mass-transport deposits (I-A-6-c)

187 187 189 190 194 198

T H E CONTINENTAL ENGLACIAL SUBENVIRONMENT (I-B) AND ITS DEPOSITS

203

ENGLACIAL CONDITIONS IN ACTIVE ICE ENGLACIAL CONDITIONS IN PASSIVE ICE SEDIMENTATION PROCESSES IN THE ENGLACIAL SUBENVIRONMENT DEPOSITS OF THE ENGLACIAL MELTING-ICE FACIES (I-B-1) Englacial melt-out complexes (I-B-1-a) DEPOSITS OF THE ENGLACIAL FLUVIAL FACIES (I-B-2) Englacial meltwater-tunnel deposits (I-B-2-b) DEPOSITS OF THE ENGLACIAL MASS-TRANSPORT F A C E S (I-B-6) Englacial crevasse deposits (I-B-6-b) T H E CONTINENTAL SUBGLACIAL SUBENVIRONMENT (I-C) AND ITS DEPOSITS

204 206 207 207 208 213 214 218 218 223

SUBGLACIAL CONDITIONS UNDER ACTIVE ICE SUBGLACIAL CONDITIONS UNDER PASSrVE ICE SEDIMENTATION PROCESSES IN THE SUBGLACIAL SUBENVIRONMENT DEPOSITS OF THE SUBGLACIAL MELTING-ICE FACIES (I-C-1) Subglacial till complexes (I-C-1-a) Lodgement tills (I-C-1-c) Basal tills (I-C-1-d) Subglacial ice-raft deposits (I-C-1-e) DEPOSITS OF THE SUBGLACIAL FLUVIAL FACIES (I-C-2) Subglacial meltwater-tunnel deposits (I-C-2-b) DEPOSITS OF THE SUBGLACIAL DELTAIC FACIES (I-(3-3) Subglacial deltaic complexes (I-C-3-a) DEPOSITS O F THE SUBGLACIAL LACUSTRINE FACIES (I-C-4) Subglacial lacustrine complexes (I-C-4-a) DEPOSITS OF THE SUBGLACIAL MASS-TRANSPORT F A C E S (I-C-6) Subglacial mass-transport deposits (I-C-6-c)

224 230

T H E MARINE GLACIAL ENVIRONMENI' (I-D) AND ITS DEPOSITS

267

SEDIMENTARY PROCESSES IN THE MARINE GLACIAL ENVIRONMENT DEPOSITS OF THE MARINE GLACIAL MELTING-ICE FACIES (I-D-1) Marine glacial melt-out complexes (I-D-1-a)

269 270 271

230 23 1 234 237 246 249 252 254 257 257 260 261 263 264

XI1

DEPOSITS OF THE MARINE GLACIAL BO'JTOM-CURRENT FACIES (I-D-2) Marine glacial tunnel-mouth deposits (I-D-2-c) DEPOSITS OF THE MARINE GLACIAL MASS-TRANSPORT F A C E S (I-D-6) Marine glacial mass-transport deposits (I-D-6-c)

274 274 278 278

T H E CONTINENTAL TERMINOGLACIAL SUBENVIRONMENT (11-A) AND I T S DEPOSITS

281

SEDIMENTATION PROCESSES IN THE TERMINOGLACIAL SUBENVIRONMENT DEPOSITS OF THE TERMINOGLACIAL MELTING-ICE F A C E S (II-A-1) Terminoglacial till complexes (II-A-1-a) Terminoglacial ice-raft deposits (11-A-1-e) DEPOSITS OF THE TERMINOGLACIAL FLUVIAL F A C E S (II-A-2) Terminoglacial fluvial complexes (II-A-2-a) Terminoglacial tunnel-mouth deposits (II-A-2-12) Terminoglacial stream deposits (II-A-2-d) Terminoglacial sheet- and streamflood deposits (II-A-2-e) DEPOSITS OF THE TERMINOGLACIAL DELTAIC FACIES (II-A-3) Terminoglacial deltaic complexes (II-A-3-a) Terminoglacial deltaic topsets (II-A-3-b) Terminoglacial deltaic foresets (II-A-3-c) Terminoglacial deltaic bottomsets (II-A-3-d) DEPOSITS OF THE TERMINOGLACIAL LACUSTRINE F A C E S (II-A-4) Terminoglacial lacustrine complexes (II-A-4-a) Terminoglacial lake-margin deposits (II-A-4-b) Terminoglacial lacustrine bottomsets (II-A-4-c) DEPOSITS O F THE TERMINOGLACIAL AEOLIAN FACIES (II-A-5) Terminoglacial drift sands (II-A-5-b) DEPOSITS O F THE TERMINOGLACIAL MASS-TRANSPORT FACIES (II-A-6) Terminoglacial subaerial mass-transport deposits (II-A-6-a) Terminoglacial crevasse deposits (II-A-6-b) Terminoglacial subaqueous mass-transport deposits (I-A-6-c)

282 284 285 288 293 293 296 305 308 312 313 316 318 321 324 327 330 333 341 342 344 346 349 352

T H E CONTINENTAL PROGLAC 1A L SUBENVIRONMENT (11-B) AND ITS DEPOSITS

361

SEDIMENTATION PROCESSES IN THE PROGLACIAL SUBENVIRONMENT DEPOSITS OF THE PROGLACIAL MELTING-ICE F A C E S Proglacial till complexes (II-B-1-a) DEPOSITS OF THE PROGLACIAL FLUVIAL F A C E S (XI-B-2) Proglacial fluvial complexes (II-B-2-a) Proglacial stream deposits (II-B-2-8 Proglacial sheet- and streamflood deposits (I-B-2-e)

363 365 365 367 373 376 385

XI11

DEPOSITS OF THE PROGLACIAL DELTAIC F A C E S (II-B-3) Proglacial deltaic complexes (II-B-3-a) Proglacial deltaic topsets (11-B-3-b) Proglacial deltaic foresets (II-B-3-c) Proglacial deltaic bottomsets (11-B-3-d) DEPOSITS OF THE PROGLACIAL LACUSTRINE F A C E S (11-B-4) Proglacial lacustrine complexes (II-B-4-a) Proglacial lake-margin deposits (11-B-4-b) Proglacial lacustrine bottomsets (II-B-4-c) DEPOSITS OF THE PROGLACIAL AEOLIAN FACIES (11-B-5) Proglacial aeolian complexes (11-B-5-a) Proglacial drift sands (11-B-5-b) Proglacial dunes (11-B-5-c) Proglacial coversands (11-B-5-d) Proglacial loesses (11-B-5-e) DEPOSITS OF THE PROGLACIAL MASS-TRANSPORT FACIES (11-B-6) Proglacial subaerial mass-transport deposits (11-B-6-a) Proglacial subaqueous mass-transport deposits (11-B-6-c)

391 392 395 400 403 408 409 413 417 423 424 427 43 1 437 443 447 448 453

T H E CONTINENTAL EXTRAGIAACIAL SUBENVIRONMENT (11-C) AND ITS DEPOSITS

459

SEDIMENTATION PROCESSES IN THE EXTRAGLACIAL SUBENVIRONMENT DEPOSITS OF THE EXTRAGLACIAL AEOLIAN FACIES (11-C-5) Extraglacial aeolian complexes (II-C-5-a) Extraglacial drift sands (II-C-5-b) Extraglacial dunes (11-C-5-c) Extraglacial coversands (II-C-5-d) Extraglacial loesses (II-C-5-e) DEPOSITS O F THE EXTRAGLACIAL MASS-TRANSPORT FACIES (II-(3-6) Extraglacial subaerial mass-transport deposits (II-C-6-a)

460 463 463 466 469 475 480 490 49 1

THE MA K.1N E T E HM IN OG LAC I A L SUBE N V I RONM E NT (1I -D) AND ITS D E P O S I T S

499

SEDIMENTATION PROCESSES IN THE MARINE TERMINOGLACIAL SUBENVIRONMENT DEPOSITS OF THE MARINE TERMINOGLACIAL MELTING-ICE FACIES (II-D-1) Marine terminoglacial ice-raft deposits (II-D-1-e) DEPOSITS OF THE MARINE TERMINOGLACIAL BOTTOM-CURRENT FACIES (II-D-2) Marine terminoglacial tunnel-mouth deposits (II-D-2-c) Marine terminoglacial bottom-current deposits (II-D-2-d)

501 502 504

508 510 513

XIV

DEPOSITS OF THE MARINE TERMINOGLACIAL DELTAIC FACIES (II-D-3) 517 517 Marine terminoglacial deltaic complexes (IT-D-3-a) 519 Marine terminoglacial deltaic topsets (II-D-3-b) 521 Marine terminoglacial deltaic foresets (II-D-3-c) 523 Marine terminoglacial deltaic bottomsets (II-D-3-d) DEPOSITS OF THE MARINE TERMINOGLACIAL MASS-TRANSPORT FACIES (II-D-6) 526 527 Marine terminoglacial mass-transport deposits (II-D-6-c)

T H E MARINE PROGLACIAL SUBENVIRONMENT (11-E)

533

THE: MARINE EXTRAGLACIAL SUBENVIRONMENT (11-F) AND ITS DEPOSITS

537

SEDIMENTATION PROCESSES IN THE MARINE EXTRAGLACIAL SUBENVIRONMENT DEPOSITS OF THE MARINE EXTRAGLACIAL BO'M'OM-CURRENT FACIES (11-F-2) Marine extraglacial bottom-current deposits (II-F-2-d) DEPOSITS OF THE MARINE EXTRAGLACIAL DELTAIC FACIES (11-F-3) Marine extraglacial deltaic complexes (II-F-3-a) Marine extraglacial deltaic topsets (II-F-3-b) Marine extraglacial deltaic foresets (II-F-3-c) Marine extraglacial deltaic bottomsets (II-F-3-d) DEPOSITS OF THE MARINE EXTRAGLACIAL MASS-TRANSPORT FACIES (II-F-6) Marine extraglacial mass-transport deposits (IT-F-6-c)

538 539 539 544 545 546 548 550 552 552

EPILOGUE

555

REFERENCES

561

S U B J E C T INDEX

669

Introduction

1

INTRODUCTION Glacial activity involves both erosion and sedimentation (Ehlers, 1981; Vinogradov, 1981),thus shaping the Earth (a.0. Embleton and King, 1968, 1975,1977; Sugden and John, 1976). Either erosion (Fig. 1)or sedimentation prevail in specific places, but the two may alternate in time as well (Lindner and Ruszczyfiska-Szenajch,1979).The entire area covered by an ice sheet or under its immediate influence (as indicated by the presence of a permafrost layer or some kind of sediment related to the ice cover) is therefore considered here to form part of the depositional environment. The ice-related sediments - and their depositional processes - from both areas are often called 'glacigenic' (sometimes 'glaciogenic': Visser, 1989).

Fig. 1. Typically glacial valley (SWNorway), owing its U-shape to erosion ofwalls and bottom by a Pleistocene glacier. The material eroded here now forms probably part of many glacigenic deposits, partly many hundreds of kilometres away from the source area.

2

Introduction

As will be explained in more detail later, the most useful approach in practice is to consider these glacigenic areas as an entity, but to distinguish between the area covered by ice (termed the 'glacial environment') and the area in front of it (termed the 'periglacial environment'; the latter environment is characterised - at least under continental conditions - by a permafrosted soil and it may also occur in regions where no ice cover is nearby). It should be emphasised here that all depositional processes in the continental part of the glacial environment and the final deposits that they are responsible for are called 'glacigenic'. The continental part of the periglacial environment, as well as both the marine glacial and periglacial environments, are characterised by a combination of glacigenic and non-glacigenic processes and sediments. The non-glacigenic processes and deposits are considered beyond the scope of this book. The glacial and periglacial environments are characterised by specific physical, chemical and biological conditions (Ruszczyrkka-Szenajch, 1981a,b). Since these conditions show important variations within a n environment, it is general practice (a.0. Reineck and Singh, 1980) t o distinguish between subenvironments; these can be subdivided into sedimentary facies that are characterised by prevailing processes, resulting in more or less mutually related types of deposits. The facies is commonly considered t o be the sum of all primary features of a depositional unit, on the basis of which the conditions and mechanisms of deposition can be interpreted (Gressly, 1938; Krumbein and Sloss, 1963; Reading, 1978a; Walker, 1978; Miall, 1983; Gradziiiski et al., 1986). This approach will also be followed here. This approach results in a classification of environments, subenvironments, facies and deposits (cf. Brodzikowski and Van Loon, 1987), sometimes with further subdivisions on the basis of lithological or sedimentological characteristics. Classification schemes of natural phenomena such as facies and deposits are, almost by definition, arbitrary and controversial. Each investigator is inclined to adapt a system according to the needs of the specific research object. The investigator may also be bound t o general rules developed for a specific project or for field work carried out by a specific institute. This implies that numerous classification schemes are in current use. An additional complication is that theoretical considerations and practical applicability pose different requirements. It must be emphasised, however, that unambiguous communication between investigators is only possible if the terminology used is unequivocal. It often is not: the present authors found much data in the literature that could not be placed in a general glacial framework without raising considerable

Early literature and problems in terminology

3

doubt as to their correct position. Interpretations of data in the literature, even of vertical sections, may therefore be incorrect. The present authors apologise to authors quoted here who find t h a t t h e i r m a t e r i a l is incorrectly interpreted. We would greatly appreciate being informed of such - hopefully rare - misinterpretations. Perhaps these will act as a stimulus for the investigator involved to reconsider the way the material is described or illustrated. EARLY LITERATURE AND PROBLEMS IN TERMINOLOGY It will only be possible t o unravel the relationships between the many different types of sediments within the intriguing glacial and periglacial depositional environments if researchers can rely fully on their interpretation of work carried out by colleagues. It thus seems useful to direct attention to the terminology used earlier, in order to improve the accessibility of the older literature. The scientific literature concerning glacial sediments is very rich and has its roots i n the beginning of the 18th century (see, among others, Scheuchzer, 1723). It became much more abundant, however, in the course of the nineteenth century (Lyell, 1830, 1840a,b; De Charpentier, 1835; Murchison, 1836; Schimper, 1837; Agassiz, 1838a,b, 1840, 1842, 1847; Buckland, 1840; Godeffroy, 1841; Rendu, 1841). The terminology touching glacigenic deposits that developed gradually was partly descriptive, with terms such a s 'boulder clay' (Croll, 1870; the term is still being used: Cailleux, 1965; Olszewski, 1974) or ground moraine (Zilliacus, 1990).

From Diluvium to diamicton There were, however, also genetic terms. A still common example of such a genetic term is 'Diluvium', which term was based on the presumed deposition by the Biblical Flood (Buckland, 1823; Trimmer, 1831).Such a catastrophic process was thought necessary to explain the occurrence of uncommonly large boulders in otherwise more fine-grained layers. These boulders were - correctly - thought t o be derived from remote places; this explains their name 'erratics' (De Charpentier, 1835; Agassiz, 1838a,b), which name is still in use (see, e.g., Von Huene et al., 1973; Dalland, 1977). No other mechanism than the Biblical Flood was considered capable of transporting the huge erratics. Some geologists approached the phenomenon of erratics more scientifically. Lyell assumed for some time that the erratics were brought along

4

Introduction

by gradually melting icebergs. This view led to the term 'drift' for deposits containing erratics (Fig. 2) (it is interesting t o note how terminology can change with new scientific ideas). The term 'drift' was mainly used in the nineteenth century for more or less structureless glacial deposits (A. Geikie, 1863); this usage has gradually become obsolete, although it has not been completely abandoned (see, e.g., Okko, 1955; W.H. Johnson, 1964; Harris and Wright, 1980; Stewart and Van Hees, 1983). The usage of the term 'stratified drift' (see the interesting review by Jopling, 1975), used in the nineteenth century particularly for what are known now as glaciofluvial, glaciolacustrine and glaciodeltaic deposits (Salisbury, 1896), seems more generally accepted nowadays (Gustavson and Boothroyd, 1982; Sharpe and Barnett, 1985). It is interesting that the term 'drift' is now also applied to glaciomarine deposits (Armstrong and Brown, 1954; Pevear and Thorson, 1978). When the genesis of 'non-stratified drift' had been unravelled, and when the deposition of erratics from melting landice covers had been recognised (which was already widely accepted in the 19th century; see, among others, Jamieson, 1860; Close, 1867; Dakyns, 1872; Goodchild,

Fig. 2. Huronian ( > 2 Ga) deposits near Cobalt (Ontario, Canada), interpreted as a glaciomarine till.


5

1875; C.H. Hitchcock, 1879; Penck, 1882; Penck and Bruckner, 1909), the term 'drift' became gradually replaced by the - also genetic - term 'till' ('tillite' if lithified: Reading and Walker, 1966; Howarth, 1971; Dreimanis, 1974; Hambrey and Harland, 1979, 1981; Max, 1981; Hambrey, 1983; Dreimanis and Schluchter, 1985). The term 'till' found wide acceptance and almost all researchers working in this field have applied the term. It is interesting in this respect to mention here some of the more or less recent prominent workers who used this term: Holmes (1941, 1960),West and Donner (1956),Dreimanis (1961, 1971, 197613, 1982b, 1983, 1988), Willman et al. (1963, 1966), Kauranne (1967), Sitler (1968), Dreimanis and Vagners (1969, 1971, 19721, Frye et al. (1969), Warnke and Richter (19701, Andrews (19711, Evenson (19711, Goldthwait (1971), Niewiarowski (1971), Nobles and Weertman (1971), Pettyjohn and Lemke (1971), Ramsden and Westgate (19711, Drake (1972, 1974, 1977), Mark (1973, 1974), Boulton and Dent (1974), Boulton et al. (1974), Boulton (1976a, 1978), Mills and Mark (1976), Scott (1976), Shilts (1976, 1978), Garnes and Bergersen (19771, Gillberg (1977), Lundqvist (1977), Shaw ( 1 9 7 7 ~ Drozdowski )~ (1979a), Ehlers and Stephan (1979, 1983), Garnes (1979), Mickelson et al. (1979), Van der Meer (19791, Gibbard (1980), Haldorsen (1981, 1983a,b), Hutter and Olunloyo (1981), Kemmis (19811, Wickham-Sprecht and Johnson (19811, Baermann et al. (1983), Ehlers (1983d), Eriksson (1983), Hall (19831, H.G. Johansson (1983a), Muller (1983b), Nielsen (19831, Nielsen and Houmark-Nielsen (1983), Punning and Raukas (1983), Riezebos (1983), Ringberg (1983), Bouchard and Martineau (1984),Bouchard et al. (1984), Dreimanis and Lundqvist (1984), Rappol (1985), Bouchard and Salonen (1988) and Sharpe (1988). In spite of the wide acceptance of the term 'till' it was generally felt that a descriptive term might be more appropriate than a genetic one. The term 'synmictite' was therefore introduced by Flint et al. (19601, but this term found no general adherence. The term 'diamict' ('diamictite' if lithified: Frakes, 1978; Hambrey, 1982; Fairchild, 1985; Visser, 1989) was then introduced in the seventies (Flint, 1975; Frakes, 1975). A relatively new development is the usage of the term 'diamicton' (Lawson, 1981; Easterbrook, 1983; Gravenor, 1985).

Landforms versus sediments Another general problem in the literature results from confusion between geomorphology and sedimentology. Several, possibly even most, of the thousands of publications in this field include terms such as 'esker

6

Introduction

deposits', 'ice-pushed ridge sediments', etc. It cannot be sufficiently emphasised that such a terminology does not contribute t o a better understanding of the genesis of the lithological units involved. On the contrary, one should even consider the possibility of a morphological misinterpretation. There are some landforms that are wholly or partly due t o glacial processes. The recognition of such landforms may help in interpreting the genesis of the deposits involved. Three types of landforms are of special interest in this respect: moraines, drumlins and fluted moraines (drumlins, fluted moraines and other subglacially formed features will be discussed in more detail in the chapter on the subglacial subenvironment). Several other types of landforms result from glaciofluvial processes in (previously) glaciated areas. These forms may also be helpful in interpreting the genesis of the sediments that they contain. The most relevant forms in this context are kames, eskers, sanders and - to a lesser degree pradolinas. Moraines

The term 'moraine' is a confusing one, because it is generally used in a very loose sense. Two different meanings are found frequently in the literature: a glaciological one and a morphological one. In a glaciological sense, the term is generally used for "all rock debris incorporated in, in transit on or in, or carried and eventually deposited by glaciers'' (Visser, 1980). In a morphological sense, it is used for the ridges of debris that have accumulated at sites where a stagnant glacier has deposited debris (Fig. 3); such ridges may be pushed somewhat during a next phase of glacier advance, but it is generally agreed upon that there is a difference between moraines and ice-pushed ridges. Many authors use the terms 'moraine' and 'till' or 'glacial deposit' as more or less synonymous (e.g., Chamberlin, 1883; Bishop, 1957; Bjsrlykke, 1967; Gaigalas, 1969; Boltunov, 1970; Lavrushin, 1970a,b; Drozdowski, 1974; N. Eyles and Rogerson, 1978a,b; Ahmad, 1979; Rabassa and Aliotta, 1979; Ellis and Chalkin, 1983; Espizua, 1983; Figge, 1983; Hall, 1983; Hofle and Lade, 1983; Bouchard, 1989). Other authors, however, use the term 'moraine' in a truly morphological sense, even though sedimentary characteristics of this landform may be described (e.g., Lampluch, 1911; Goldthwait, 1951; King, 1969; King et al., 1972; N. Eyles and Rogerson, 1977b; Chinn, 1979; Habbe, 1979; Haselton, 1979; Rabassa et al., 1979, 1981; Rothlisberger and Schneebeli, 1979; Serrat, 1979; Von Husen, 1979; Warren, 1979; Rains and Shaw, 1981; Wakahama


7

Fig. 3. Ridges of debris left in front of the Columbia Icefield (Rocky Mountains, Canada) during subsequent stages of retreat of the Athabasca glacier.

and Tusima, 1981; Rogerson and Batterson, 1982; Ruszczyiiska-Szenajch, 1982a; Butler et al., 1983; Meyer, 1983b; Persson, 1983; Wilke and Ehlers, 1983; Maizels and Petch, 1985; Sharpe, 198813;Bouchard et al., 1989). It seems justified to prevent any possible confusion by using the term 'moraine' exclusively in its geomorphological sense, and by using the term 'glacial deposit' (or till, diamict) for the sediments. Indeed a clear distinction between 'till' and 'moraine' is made in the literature by several authors (e.g., Grube, 1983a; Johansson, 1983b; Lundqvist, 1983), but Boulton and Eyles (1979) raise confusion with their term 'supraglacial morainic till complex'.

Drum1 ins Drumlins are, commonly elongated and flat, mounds of glacial material, usually subglacial deposits (note: a drumlin of glaciofluvial material has been described by Shaw and Kvill, 1984). The elongation of the drumlins, which occur often in so-called drumlin fields (Haavisto-Hyvarinen, 1987; Haavisto-Hyvarinen et al., 1989), indicates the direction of ice movement (Fig. 4).

8

Introduction

Fig. 4. Orientated drumlins in the main part of the Peterborough drumlin field (Canada). After Sharpe (1987).Courtesy: A.A. Balkema (Rotterdam).

The genesis of drumlins, being the result of interaction between debrisrich ice and substratum, has been studied by many researchers (e.g., Smalley and Unwin, 1968; Whittecar and Mickelson, 1977; Boulton et al., 1979; Menzies, 1979, 1981; Seret, 1979; Shaw, 1980, 1987b; Dardis, 1981; Boulton, 1982; Menzies and Rose, 1987; Rose, 1987; Smalley a n d Piotrowski, 1987; McCabe and Dardis, 1989; Piotrowski, 1989). Regional studies with interesting sedimentological data are numerous. They include studies by Von Schaefer (1969), Lundqvist (1970), Karczewski (1976),Whittecar and Mickelson (1979), Menzies (19821, Hillefors (19831, Piotrowski (1986), Hanvey (1987), Sharpe (1987), Piotrowski and Smalley (1987) and Hanvey (1989). Fluted moraines Fluted moraines (Hoppe and Schytt, 1953; Baranowski, 1970), also termed lflutes' (Paul and Evans, 1974; Boulton, 1976b; Morris and Morland, 1976) and 'flutings' (Gravenor and Meneley, 1958; Shaw and Freschauf, 1973; Shaw, 1980; N. Jones, 1982), are elongated and flat mounds formed subglacially; they consist usually of subglacial till material. Fluted moraines are relatively small - if compared with drumlins - but their


9

origin is probably closely connected. They show a clear orientation, just like drumlin fields, indicating the direction of ice movement. Their genesis is a result of combined subglacial deposition and deformation (Menzies and Rose, 1987,1989).

Kames Geomorphologists tend to describe all irregular fluvioglacial 'highs' in the glacial environment as 'kames' (Shaler, 1884; H.L. Fairchild, 1896; Holmes, 1947; Szupryczyiiski, 1965; Karczewski, 1974; Schwan and Van Loon, 1979; Lewandowski and Zielifiski, 1980). Detailed sedimentological analysis could still show eventually t h a t such highs may consist of englacial sediments and supraglacial glaciofluvial, glaciodeltaic and glaciolacustrine deposits (Fig. 5) which owe their shape to denudation and erosion (Bartkowski, 1967; Grzybowski, 1970). 'Kames' therefore cannot be considered as more or less equivalent t o 'supraglacial and/or englacial crevasse deposits'. From a sedimentological point of view the use of the term 'kame' should therefore be avoided. From the geomorphological point of view, however, the term 'kame' could be preserved. The term 'kame' was introduced by Jamieson (1874) and is derived from the Scottish word 'kaim', used to indicate steep-sided ridges. Charlesworth (1957) considered kames as a special kind of esker, and termed the forms developed as subaqueous marginal moraines 'true kames' (also see

Fig. 5. Phases of development (A-E) of a kame. The numbers indicate successively younger deposits. From: Krzemiiski (1974). Courtesy: Societas Scientiarum Lodziensis.

10

Introduction

Francis, 1975). Several crevasse deposits have been described in the literature, usually as kames or kame terraces (Schwan and Van Loon, 19791, and their palaeogeographic development has been much analysed (e.g., Bartkowski, 1967; Klatkowa, 1972; Baraniecka, 1975; Brodzikowski, 1982a). All these authors stress that (geomorphological) kames consist of various deposits; it is also clear that most analyses point to a deglaciation or a t least a n oscillational retreat during kame building.

Eskers A special phenomenon within the glaciofluvial facies are the englacial and subglacial meltwater tunnels. Englacial tunnels generally are the natural prolongation of crevasses and, if they extend far enough, end at the base of the ice then become subglacial tunnels. Meltwater tunnels are therefore present i n both t h e englacial and t h e subglacial subenvironment. Subglacial meltwater tunnels are not always the continuation of englacial tunnel and crevasse systems, but could also have formed due to thermosubrosion resulting from the geothermal heat flux. They may be located either in the ice or in the incised substratum, depending on the thermal conditions and on the hydrological regime (Rothlisberger-type or Nye-type channels; see Sugden and John, 1976; Denton and Hughes, 1981; N. Eyles and Menzies, 1983). The englacial (and subglacial) tunnels commonly contain meltwater carrying and depositing debris. There is no abrupt transition between the englacial and the subglacial subenvironment. There is a difference in the type of deposits, however: those formed in the englacial tunnels are most strongly related to the englacial crevasse deposits (and generally have been deposited on a strongly inclined substratum consisting of ice), whereas the sediments formed under subglacial conditions are commonly true channel deposits. Deposits accumulated in meltwater tunnels tend to be relatively narrow, and high with respect to their width. After melting, this results in remarkable, elongated ridges called 'osar', 'aasar' or - most commonly 'eskers' (e.g., Sollas, 1883; Michalska, 1969; Radlowska, 1969; Allen, 1971; Banerjee and McDonald, 1975; Saunderson, 1975, 1977a,b, 1982; Saunderson and Jopling, 1980; Ringrose, 1982; Terwindt and Augustinus, 1985; Visser et al., 1987). The term 'esker' only has a morphological meaning (J.Geikie, 1877,1894; Charlesworth, 1957; Flint, 1971; Banerjee and McDonald, 1975; Sugden and John, 1976) and was defined by Francis (19751, who produced a compromise between the definitions proposed by J. Geikie (1894) and by Chamberlin (1894): "Eskers are glacial features


11

made up of morainic material deposited in contact with glacier ice as ridges whose trends tend t o conform in general with a direction of ice movement, and whose composition is dominantly, but not necessarily exclusively gravel and sand". A more detailed, but less precise definition was given by Saunderson (1975) after a thorough sedimentological study of Canadian eskers. In most cases the terms 'esker formation' and 'paraesker formation' may be clear from their context. In contrast, the meaning of 'esker deposit' is usually rather vague. Since the term 'esker' is typically geomorphological, we consider it incorrect to speak of 'esker sedimentation' (Shulmeister, 1989) or 'esker deposits'; a more general term e.g., 'subglacial channel deposit(ion)'is preferred. Although it is difficult to distinguish objectively between subglacialchannel and tunnel-mouth deposits, one must realise that there are indeed certain differences. Boulton (1972b) has explained that deposition starts in zones where the influence of tunnel-mouth conditions is minimal. If the subglacial channels embouch in a lake, a subaqueous fan may result (Aario, 1972).

Sanders Fans are present wherever meltwater streams reach depositional areas where they are split into branches. Most of the resulting fans are comparatively small but they can have a considerable morphological impact in the proglacial subenvironment (facies II-B-2) and in the terminoglacial marine subenvironment (facies II-D-2). The meltwater streams originating in the glacial environment and passing the terminoglacial subenvironment form fans in the proglacial subenvironment. The sediments of this fan facies often form part of land forms that are known as sanders; these are also called 'sandrs' or, more frequently, 'sandurs' (Krigstrom, 1962; Church, 1972; Klimek, 1972; Bluck, 1974; Boothroyd, 1976; Ward et al., 1976; Ruegg, 1977; Maizels, 1983; Landvik and Mangerud, 1985) or 'outwash plains' (Mc Donald and Banerjee, 1970,1971; Augustinus and Riezebos, 1971; Eynon and Walker, 1974; N.D. Smith, 1974; Kozarski, 1975; Fraser, 1982; Cherven, 1984; Cheel and Rust, 1986). Sanders have been considered in detail by many authors (a.0. Fahnestock, 1963; P.F. Williams and Rust, 1969; Church, 1972; Rust, 1972, 1978; Bluck, 1974; Gustavson, 1974; Boothroyd and Ashley, 1975; Church and Gilbert, 1975; Clague, 1975; Rust and Romanelli, 1975; Miall, 1977, 1978,198313;Boothroyd and Nummedal, 1978; Zieliiiski, 1980,1987b).

12

Introduction

Proglacial sanders can be sedimentologically subdivided into proximal, middle and distal parts. Horizontal grain-size differentiation is common, the coarsest units indicating the proximal facies. The thickness of the separate sets decreases from the proximal towards the distal part,but channel infillings may locally disturb this simplified picture. The deposits constituting sanders are formed by processes that do not really differ from those building up fans without any influence of an ice sheet. Particularly in the upper part of the outwash plains, there are frequent mass flows that alternate with sheetfloods and streamfloods. These relatively short-lasting processes may make a large contribution to the final volume of the fan, although the more regular glaciofluvial sedimentary processes inside channels prevail for most of the time. There is commonly a slightly undulating relief on sanders in the 'interchannel areas'. This makes that locally overbank deposits can accumulate after flooding. Such overbank deposits are much rarer on sanders, however, than in river valleys such as ice-marginal streamways (pradolinas).

Pradolinas Pradolinas are river valleys that run parallel to the ice front, as rivers flowing from an upland area towards the ice must change their course. Such valleys formed frequently in the European lowlands during the Pleistocene glaciations (Lewandowski and Zielifiski, 1988). Most commonly, the rivers in such valleys had a low-sinuosity stream pattern (Charlesworth, 1957; Woldstedt, 1957; Kozarski, 1967, 1969; Galon, 1968); recent proglacial river valleys show characteristics that are sedimentologically very similar (P.F. Williams and Rust, 1969) (Fig. 6). It is commonly possible t o distinguish four main depositional levels in pradolinas, just as in most valleys of large-scale braided or low-sinuosity rivers. The lowermost level is that of the active channels, where bars are exposed only during low-water stages. The second level is characterised by the presence of a few channels that are active only during flood stages; there may be a sparse vegetation cover (Williams and Rust, 1969; Miall, 1977, 1983b). The third level has channels where low-energy water currents flow during flood stages; moderate vegetation covers are common in humid areas. The uppermost, fourth level consists mainly of islands and interfluves; there may be dense vegetation, but there may also be areas of aeolian deflation and dune migration (Fig. 7). Terraces may be found that are independent of the glaciofluvial depositional levels, as a result of changes in the local erosion base (Kozarski, 1962, 1965; Galon, 1968; Zieliiiski, 1980b).


positions of ice margins

13

pradolinas

Fig. 6. Pradolinas in the central European plain. Adapted from J a h n (19751, after Woldstedt (1950). Courtesy: PWN.

Ambiguous descriptions A third general problem, also frequently encountered, is the use of insufficiently detailed facies descriptions. The term 'till' has little meaning if it is not indicated whether the diamict results from debris on top of, in, underneath, or at the front of the ice mass, and if the process of sedimentation is not made clear. Similar insufficiently accurate descriptions are, for instance, 'glaciolacustrine' (where?) and 'glaciomarine' (also: where?). Nomenclature

Some new terms will need t o be introduced in this book. Wherever possible, however, use is made of existing and widely applied terms. One of the main reference works in this respect is the nomenclature published in five languages by the Royal Geological and Mining Society of The Netherlands (Visser, 1980). The more specific sedimentological and glaciological terminology, and the classification schemes foIlowed have been established taking into account the following important works in this field: Chamberlin (1894), Woodworth (1899), Kuenen (1950, 19531, Ksiazkiewicz (19541, Bouma (1962), Allen (1963, 1966, 1968, 1970a,c,

14

Introduction

Fig. 7. Schematic model of an ice-marginal streamway (pradolina) with four depositional levels in the main channel. Modified after Kozarski (1962, 1965), Williams and Rust (1969) and Brodzikowski and Van Loon (1987).

19821, Diulyiiski (1963a), Schumm (1963),Washburn et al. (1963), Bouma and Brouwer (1964), Diulyiiski and Walton (1965), Middleton (19651, Nagtegaal (19651, Jahn (1970, 19751, Flint (19711, Katasanov (19731, Popov (19731, Allen and Collison (1974),Carter (1975),Miall (1977,19781, Rust (1975,1977,19781, Boulton (1976c, 1980a1, Sugden and John (1976), Aario (19771, Embleton and King (19771, Boulton and Eyles (19791, Embleton and Thorns (1979), Laverdiere et al. (19791, Lowe (1979), Prior and Coleman (19791, Schluchter (1979), Dreimanis (1980, 1982b, 1988), Reineck and Singh (19801, Rukhina (19801, Ruszczyliska-Szenajch (1982b), Ehlers (1983), Evenson et al. (19831, Eyles (19831, Kingston et al. (1983), Gravenor et al. (19841, Gradziiiski et al. (1986) and Brodzikowski and Van Loon (1987).

Sources used

15

SOURCES USED The material just mentioned, mainly overviews of various aspects dealt with in the present volume, was one of the important sources of information. Papers from a vast number of journals were a n even more abundant source. In spite of confusing terminologies, the present authors have tried to interpret these sources as correctly as possible. Only some of the sources were selected for listing in the reference list. This list covers mainly depositional aspects; the disturbances (Fig. 8) caused by glaciotectonism (see, e.g., Gripp, 1979), a common phenomenon that may hamper investigations considerably but that also may provide indications about the direction of ice movement may be provided (Hicock and Dreimanis, 1984), will be discussed i n another volume. Coverage of the literature for the present book was extensive as one of the authors (K.B.) could study a wealth of material in languages from Eastern Europe. The Polish and Russian literatures have provided extremely interesting data. The data are of great interest, not only from the point of view of the regions covered, but particularly because the

Fig. 8. Glaciotectonically deformed glaciofluvial megaripple (Uelsen area, Federal Republic of Germany).

16

Introduction

relative lack of exchange of information between the Western and Eastern countries has gradually resulted in the development of largely diverging views on several problems. The authors aimed at unification of these data into models satisfactory with respect to both views. Although some discrepancies in interpretations and models could not be entirely reconciled, the authors feel that some progress has been made. The choice of how much the interesting Eastern European literature should be included was a difficult one. Most such data not only will remain inaccessible for the majority of researchers in the Western world but the less recent (and sometimes even the new) books, monographs, journals, etc. may be difficult t o purchase. The reference list would almost have doubled in length had all relevant material from Eastern Europe been included. The authors thus decided to include only the most important work that is of relatively easy accessibility.

SCOPE OF THE BOOK The sedimentology of the glacial and periglacial environments covers a wide field. It is impossible t o deal in detail with all aspects without making this book of an impractical length. We therefore have chosen t o present a general framework, with emphasis on practical aspects such as the recognition, correct description and logical interpretation of glacigenic sediments, all within the context of an understanding of the glacigenic conditions and depositional processes. It is not intended to present regional overviews of glaciated areas; the reader is referred to more appropriate literature for such material (e.g., De Ploey, 1961; J.D. de Jong, 1965; Vorren, 1973; Matwiejew, 1976; Raukas, 1978; Hantke, 1979; Schubert, 1979; Campy, 1983a; De Jong and Maarleveld, 1983; Ehlers, 1983b,c; Lundqvist, 1983a,c; Mangerud, 1983; Rasmussen, 1983; Sjorring, 1983; Sorensen, 1983; Ter Wee, 1983a,b; Andrews et al., 1984; Clague, 1986). Sediments of Pleistocene age will be emphasised because they are the most frequently occurring, and because Pleistocene glacigenic sediments generally differ from other Pleistocene deposits more than older glacigenic sediments differ from their non-glacigenic counterparts (due t o processes like consolidation, lithification and/or metamorphism); criteria for the distinction of old glacigenic deposits have been provided by Gamundi and Amos (1983). Not all types of glacigenic sediments will receive equal attention. The authors have tried to concentrate on those types of sediments that either give rise to most problems during field work (e.g., tills) o r are most

Scope of t h e book

17

common. Generally rather badly exposed sediments such as Pleistocene glaciomarine deposits or rare types such as subglacial lacustrine deposits therefore will be dealt in relatively less detail. Some aspects related to glacigenic geology are considered out of the scope for t h e present book. The most important topics of this kind a r e glaciotectonism (see, among others, Schwan and Van Loon, 1981; Gripp, 1983; Maarleveld, 1983; Lea, 1985; Kozarski and Kasprzak, 1987), glacigenic morphology (Dylikowa, 1952; Hoppe, 1959; Bik, 1960; Tricart and Cailleux, 1967; Dionne, 1968; Embleton and King, 1968, 1975, 1977; Reid, 1970a; Ryder, 1971; P.G. Johnson, 1972; Price, 1973; Clayton and Moran, 1974; Sugden and John, 1976; Butzer, 1977; Church, 1977; Moran e t al., 1980; Croot, 1981; Stow, 1981; N. Eyles and Paul, 1983; N. Eyles e t al., 1983a; Grube, 198313; Houmark-Nielsen, 198313; Kruger, 1983; M. Sharp, 1985a; Bouchard, 19891, ice extent (Fannin e t al., 1979; N. Eyles and Westgate, 1987), stratigraphy (Van der Hammen and Maarleveld, 1952; Von Steinmuller, 1973; Von Jerz, 1979; Kozarski, 1981; Nelson, 1981; Vandenberghe, 1981; Vandenberghe and Krook, 1981; Ehlers and Iwanoff, 1983; Graf, 1983; Lagerlund, 1983; Stephan et al., 1983; McCabe, 1987), palaeontology (J.D. Shaw, 1972; Anderson, 1975; Allison, 1978; Kellogg e t al., 1979; Osterman, 1982; Brandani, 1983) and palaeoecology (Frenzel, 1959; Wagner, 1959; Martin and Wilczewski, 1970; Lord, 1979; Balazarini, 1983; Drozdowski, 1986; Eissmann, 1990) of glacigenic deposits, the sea-level changes induced by ice ages (Easterbrook, 1963; Roeleveld and Van Loon, 1979; Beard e t al., 1982; Tikkanen, 1989), the causal factors behind the occurrence of ice ages (Milankovitch, 1930, 1938; Van Loon, 1980,1982),and the chronology of ice ages (Hamelin, 1969; Frazier, 1974; Eisbacher, 1981; Brugger et al., 1983; Josenhaus, 1983). These subjects (and a few other, less important ones) will only be mentioned occasionally in the present book, where considered appropriate. No attention is paid to applied aspects. Readers interested in these aspects are referred to, among others, Baker (1974), Collins (1981) and Lloyd (1983)for hydrological aspects and water supply; to Richards (1976), Cocksedge (1983), Depiante (1983), N. Eyles (1983b), Money (19831, Somerville (1983)and Strachan and Dearman (1983) for civil engineering; and to Evenson e t al. (1979), De la Grandville (19821, Stephens e t al. (1983) and N. Eyles and Kocsis (1989) for mining aspects. I t is hoped that, in spite of all these restrictions aimed at making the book a practical reference, the reader will find most of the information sought for. Should t h e reader find that specific aspects a r e inadequately dealt with, the authors would greatly appreciate receiving any critical comments.

18

Introduction

FRAMEWORK O F THE BOOK The book is subdivided into two main parts. The first deals with the various general aspects important for facies interpretation (particularly touching glacigenic facies) and gives a survey of the parameters that determine these facies. Terminology and use of symbols are one of the principal aspects dealt with in this first part. This first part also deals with the differentiation of the glacigenic environments into subenvironments, their facies and the specific types of sediments. This is done by means of models, systematic analyses and definitions. This part is structured so as t o provide a relatively simple key for establishing the various types of genetically related sediment types. The second - most extensive - and, in our opinion, the most important part deals with the various glacigenic facies and provides detailed descriptions of all types of glacigenic deposits and their characteristics, together with photographs; it is also illustrated with sections from a large number of sources. Lithological characteristics, textural characteristics, occurrence and depositional mechanisms are detailed as far as considered appropriate. This framework was chosen to aid those who are not yet familiar with glacigenic sediments, but also to facilitate field interpretations for more experienced investigators.

General characteristics of glacigenic sedimentation

19

GENERAL CHARACTERISTICS OF GLACIGENIC SEDIMENTATION What is so far the most complete monograph on the Quaternary was published in 1957 by Charlesworth, who analysed the development of knowledge regarding glacigenic sediments: Although the clays, sands and gravels belong t o the youngest and most accessible formation, their apparently chaotic state and seeming lack of interest made them the last to be investigated: they were for long a synonym for confusion, and except for their fossil shells and bones seemed unattractive and unimportant. The 'extraneous rubbish' was a troublesome hindrance in examining the 'solid' geometry. Long after Agassiz had revived the glacial theory, official state surveys ignored them. Thus the British drifts were passed over almost without scrutiny until most of Southern England had been examined. They were first mapped in Norfolk by J. Trimmer. Their mapping was only undertaken when, somewhat belatedly, their connection with agriculture, drainage, dwelling sites and engineering problems had been recognized". This view from the time of Charlesworth is now, while only a few decades old, a thing of the past. Glacial geology now receives much more attention and new research methods continue t o be developed. One of the characteristic differences between Charlesworth's and our time is the present emphasis on facies analysis. In spite of the rather recent tendency towards facies analysis and sedimentary models, some early reports on glacigenic deposits showed a fairly modern sedimentological approach. Such reports received, however, less attention from glacial geologists than they deserved. Some of these early reports concern glaciolacustrine sediments (A. Smith, 1832; Hitchcock, 1841; Jamieson, 1863); other works concerned sandur plains and glacigenic deltaic sediments (Gilbert, 1885, 1890; Davis, 1890; Salisbury, 1896) and the sedimentology of glacial diamicts (Agassiz, 1840; A. Geikie, 1863; Jamieson, 1865; Goodchild, 1874; J. Geikie, 1877, 1894; Torell, 1877; Chamberlin, 1894; Crosby, 1896). These reports might even be considered as the predecessors of the more recent publications that devote much attention t o facies associations and sedimentological patterns (e.g., Potter and Pettijohn, 1963; Broussard, 1975; Bull, 1977; Collinson, 1978; Friedman and Sanders, 1978; Reading, 1978a; Reineck and Singh, 1980; Leeder, 1982; Miall, 1984; Gradziiiski et al., 1986). 'I...

20


Recent requirements for studies on glacial sedimentology include the reconstruction of the palaeogeographic development of the ice-covered area (e.g., Bouchard and Martineau, 1985). DEPOSITIONAL PROCESSES IN THE GLACIGENIC ENVIRONMENTS The t w o glacigenic environments show distinct variations in the predominance of the depositional processes. Material may be transported by ice, water, wind or due to gravitation. Deposition may take place from active or passive ice, in running or stagnant water, by large-scale or local winds, and along steep or barely inclined, subaqueous or subaerial slopes. This results in a complex pattern that changes rapidly in both time and space. The frequent facies changes depend heavily on, for instance, the dynamics of the ice sheets and on their sediment supply. Sedimentation by ice The feature most characteristic of the glacial environment is, from a sedimentological point of view, the deposition of debris supplied by the ice mass (Fig. 9). The most common depositional process is the settling of material from melting ice. This process often leaves poorly sorted sediments (diamicts) in which the larger clasts may still show a preferred orientation that, although commonly vague, is in accordance t o their position within the ice mass. Deposits thus formed are commonly indicated by the (genetic) term 'till'; a special type are the ice-raft deposits and related types of sediments, which contain clasts derived from a melting ice cover on top of a water body. Melting of ice takes place during both active (forward moving) and passive (gradually retreating due t o ablation) stages of the ice. The resulting sediments show somewhat different characteristics, mainly due to differences in the original flow lines of the ice, the rate of melting, the character of ablation, etc. More common characteristics stem from the precise place of deposition, the local topography, climatic factors and the occurrence of endogenic processes. The combination of all these parameters gives rise to sediments (tills) that generally have a diamict character (tills were previously often called 'boulder clays', or something similar in several countries).

Depositional processes in the glacigenic environments

21

ice movement

climatic conlrol

Fig. 9. Relationships between the main agents that influence sedimentation by glacial ice. Dashed arrows indicate main relationships, black arrows indicate intermediate ones, and white arrows indicate minor relationships.

Subaqueous sedimentation Glacigenic areas are commonly characterised by poorly permeable soils. This is due, particularly directly in front of ice caps, t o the permafrost and to the occurrence of sediments such as diamicts or loesses with low permeability. Undulations in the topography therefore easily lead t o lakes. Another lake-forming process is the irregular movement of ice lobes, resulting in dammed-off meltwater streams. Whatever is the origin of a lake, one of the main characteristics is the (almost) stagnant water in which even the finest sediment particles may settle. The water in glacial lakes is due only for a minor part to local melting of ice. Most of the water is supplied by meltwater streams originating a t a more or less remote place. Such meltwater streams tend to have a braided character, indicative of changes in water supply and thus of stream

22


velocities and channel depth. These circumstances result in deposits much more irregular than those formed in lakes. Sedimentation from running water

Considerable quantities of meltwater may be formed in the ablation zone of a n ice sheet if the climatic conditions are favourable. The meltwater streams can be found on top of the ice, in tunnels within the ice and underneath the ice. They finally leave the ice mass and flow, with often large amounts of debris, into the foreland of the ice mass, where the material is deposited sooner or later. The dynamics and the transport capacity of the meltwater streams are fairly variable in time and space (Ostrem, 1975), being determined by the ablation rate, local topography, type of material transported, etc. The deposits formed from such streams are all designated by t h e general (genetic) term 'glaciofluvial deposits' (see, e.g., German et al., 1979; Williams and Wild, 1984); synonyms used less frequently are 'glaciofluvial deposits' (Paul and Evans, 1974), 'fluvioglacial deposits' (Augustinus and Riezebos, 1971), 'glacifluvial deposits', 'meltwater deposits' (Ehlers and Grube, 1983) and 'melt-water deposits' (Pessl and Frederick, 1981). Glaciofluvial deposits (Fig. 10) generally constitute the major part of all glacigenic sediments and they show most of the same characteristics as fluvial deposits of non-glacial origin. Most glaciofluvial deposits a r e relatively coarse-grained because the flow rate is temporarily too high for settling of the finest particles, but also because the fine-grained material is trapped in pools and lakes. The final characteristics are mainly determined by a limited number of parameters (Leopold et al., 1964; Allen, 1982; Gradzinski et al., 1986): bed geometry, amount of water, flow velocity, water depth and type of substratum. These parameters show interrelations and are largely influenced by the ablation conditions of the ice. Sedimentation in stagnant water

An irregular topography may, as well as ice lobes, damm off meltwater streams (cf. R. Gilbert, 1971) and thus form pools and lakes. Most of such glacial lakes are rather small (up t o a few kilometres in diameter, at most) and of short duration, but very large lakes may occur and survive for several thousands of years. The sediments formed in such lakes are most commonly called 'glaciolacustrine deposits', although the terms 'glacilacustrine' and 'glaci(o)


23

Fig. 10. Glaciofluvial channel fill (Balderhaar, Federal Republic of Germany; exposure known as 'wall of the angry farmer'). Note the channel lag with angular pieces of unconsolidated sand. These sand pebbles were transported in frozen form.

limnic' are also used. Sediments in glacial lakes may be derived from melting ice along the lake margin, from meltwater streams embouching in the lake or from dust-bearing winds. Most glaciolacustrine sediments are relatively fine-grained because the water is stagnant or has a low flow velocity, so that even the finest particles may settle. Factors responsible for the final depositional process are: settling out of suspension (wind-blown material, surficial currents), bottom currents, wave action (either or not in combination with tides) and mass movements along the slopes. The lithological characteristics depend on the prevailing process(es), but one commonly finds relatively coarse lake-margin deposits and fine-grained bottomsets; the latter frequently show varves (Fig. 11): graded layers that may originate from seasonal settling when a n ice cover melts in the spring, but may also be due to turbidity currents.

24


Fig. 11. Varves in a glaciolacustrine succession of Drenthian age (overburden of the BekhaMw browncoal mine, central Poland).

Aeolian sedimentation The presence of large ice caps has a considerable influence upon the atmospheric circulation. Thermal inversion occurs frequently and cold air masses from above the ice cap meet the warmer air from the foreland a t the ice margin. These conditions are favourable for the production of intensive winds. Wind action in the area in front of the ice (Kida, 19851,where braided streams flow between subaerially exposed fluvioglacial sediments, results in wind erosion which happens even more easily since no or almost no vegetation is present. Snow storms may even erode particles larger than sand size, thus giving rise to relatively coarse niveo-aeolian deposits (cf. Baranowski and Pekala, 1982).The eroded material may be blown away over extremely large distances (dust that has originated now from the African Sahara can be traced in Western Europe and in the United States), but commonly results in a zone of coversands (Fig. 12)followed by a zone of the finer-grained (silty) loesses. The final depositional extent and


25

Fig. 12. Coversands of Vistulian (Weichselian) age (14,000-10,000 years BP) exposed i n a browncoal mine in central Poland (Kleszczow graben, near Lodi). Note the frequent alternations of coarse and fine laminae, resulting from phases with higher and lower wind velocities respectively. Photograph: J. Gokdzik.

grain-size distribution of the coversands and loesses (Smalley and Leach, 1978) depend on the wind velocity, prevailing wind direction, nature of the eroded material, topography of the area (both coversands and loesses tend t o level off height differences in the depositional area), vegetation, etc. (Catt, 1977). Aeolian deposits become quite commonly reworked (Mucher and De Ploey, 1977), either by new wind activity or by surficial currents (rain water). Some more or less classical loess areas (e.g., in southern Poland) even turned out to have few original loesses but mainly glaciolacustrine sediments that were almost entirely derived from loess.

Deposition from mass movements Each slope, either subaerial or subaqueous, easily induces mass transport. Inclinations of less than one degree may be sufficient for processes like subaqueous slumping but other forms of mass transport may require steeper slopes.

26


Subaqueous mass movements Rivers may induce rock fall by undercutting the walls but the smaller or larger blocks thus formed in the river bed have almost no preservational potential if consisting of unlithified material. Well preserved mass-movement deposits are much more common in glaciolacustrine facies (Fig. 13) where the supply of sediment from meltwater streams may build up unstable slopes. Slumps, slides, mudflows and turbidity currents will then result. A special type of sediment is formed by material that enters the lake more or less directly from the ice, commonly by plastic flowage (flow till). Subaerial mass movements Subaerial mass movements are quite common all over the glacigenic environments, though they are not evenly distributed. All types of sediments (glacial, fluvioglacial, glaciolacustrine and aeolian) may undergo

Fig. 13.The irregular surface of sediments on top of dead-ice bodies, due to collapse after melting of buried ice, triggers subaerial mass-transport processes (Hornsund area, Svalbard).Photograph:J. Cegka.

Characteristics of glacigenic sedimentation

27

such reworking, particularly if sedimentation or erosion has created differences in height (a slope of a few degrees is enough) and when the soil is wet, e.g., after rainfall or when any other process has reduced the mechanical strength of the sediment. The intensity of the mass-movement process determines i n how far the original sedimentary characteristics of the reworked material will be preserved. The deposits that have undergone subaerial reworking have been named 'slope deposits', but the reworking has commonly been so slight t h a t there seems t o be no reason t o consider them as a separate group of sediments; one might even consider some slight subaerial reworking as part of the more general pedological processes. CHARACTERISTICS O F GLACIGENIC SEDIMENTATION The general characteristics of the glacigenic facies depend largely on the nature of the material supplied. Lack of specific grain sizes, for instance, will result in the absence of specific sedimentary structures. The source areas therefore influence the glacigenic facies, but other factors (transgressive or regressive tendencies, tectonic activity, isostatic movements, climate, intrabasinal processes such as reworking, compaction, etc.) also play a role. Knowledge of glacigenic sedimentation has greatly profited from the current interest in the environmental conservation of relatively undisturbed regions. This has resulted in more frequent earth-science research in areas such as Antarctica (see, e.g., Jacobs et al., 1970; Hughes, 1975, 1982; Moyan, 1976; Macharet, 1981; Lennon et al., 1982; Lindner et al., 1982; McKelvey, 1982; Kristensen, 1983; Rabassa, 1983; Domack, 1985) and Spitsbergen - often called 'Svalbard' in the literature - (Gripp, 1929; Klimaszewski, 1960; Kozarski, 1982; Szczypek, 1982; Kida, 1985).

The source of glacigenic sediments There are three main sources for the debris transported by glaciers and ice caps, i.e. material eroded from the substratum (and if present, valley walls: Larsen and Mangerud, 1981; Rastas and Seppala, 1981), detritus falling from nunataks (due to, e.g., frost weathering: Fig. 14; see also Brockie, 1973; Reheis, 1975; Latridou and Ozouf, 1982) on the ice surface, and particles that were supplied by the wind. The last type of debris is commonly of minor importance, while i t is t h e first type t h a t predominates (Fig. 15).

28


Fig. 14.Irregular rock shapes due to frost weathering at an altitude of some 3300 m in the Zillertaler Alps (S. Austria).

Fig. 15. Sources of mineral particles in the glacigenic system, and main interrelationships of the factors influencing the characteristics of the glacial debris.


29

Material from nunataks and wind-blown particles start their glacial transport on the ice surface. They may become incorporated in the ice when transported by supraglacial meltwater streams disappearing in englacial crevasses and tunnels, but also when fresh snow forms new covers. The material eroded from the substratum may be transported at the ice base, but may also become incorporated in the ice due to shearing that takes place in the ice mass. These processes imply that all debris transported by the ice become more or less mixed, which is one of the reasons for the diamict character of most tills (the breakage and pulverisation of clasts during transport are another reason: Hallet, 1981; Nahon and Trompette, 1982). The only glacial deposits that commonly show rather specific (non-mixed) characteristics are tills formed in the ablation zone by melting of ice with debris that has been eroded shortly before and that had no time to be mixed with other material or to be pulverised; such tills can show characteristics that resemble local pre-glacial surface deposits. Grain size of glacigenic sediments

The mixing of detritus during transport by ice results in poor sorting. This, however, does not imply that all glacigenic sediments have equal characteristics. Differences may occur due to, for instance, variations in time of source area, the prevailing transport mechanism and the position of clasts within the ice. Even though the glacigenic facies may thus vary, they commonly show debris of all grain sizes, particularly if the ice cap has eroded continental lowlands. A characteristic diamict is formed if ice containing debris of all these fractions should melt. It should be kept in mind, however, that meltwater streams may wash out such deposits; since the clay fraction and the boulders are most difficult to erode, it is quite common that a typical 'boulder clay' is left and that most sand and silt is washed away and deposited elsewhere in a fluvioglacial facies. If the meltwater streams are strong enough, no till will be formed or previously formed tills will be eroded and material comprising all grain sizes will be deposited in the fluvioglacial facies that commonly shows alternating layers of coarser and finer material, representing flows with more and less energy respectively. Transgressive and regressive tendencies

An ice sheet or glacier constitutes an energy system. The development and disappearance of such a system are lengthy processes. Growth of the ice

30


mass and, consequently, a n increase in energy are, on the long term, mainly determined by climatic developments. An increased accumulation of snow, gradually converted into ice, is commonly due to a lowering of the temperature and a n increased precipitation rate. Only if a certain threshold has been passed, does the mere existence of the ice body itself influence climatic development (less precipitation through dry atmospheric conditions, high albedo): the climate becomes colder and dryer. The resulting decrease in precipitation implies that ablation may start predominating over accumulation, and that the energy level is distinctly lowered, Thus, transgression changes into stabilisation or even regression. This development is complicated by t h e time l a g between t h e occurrence of specific processes and the final effects that they induce. In fact, a wetter and cooler climate existed for a long time before a n ice mass shows a real transgressive behaviour; on the other hand, the transgression can continue if the climatic conditions already favour a regression. This example - many more are available - indicates that all dynamics of glacigenic processes, evidently including glacigenic sedimentation, depend on complex mass-balance relations. An additional complication is that the processes that determine the sedimentary pattern are different during the transgressive, stable and regressive phases (Fig. 16). Transgressive phases are characterised by prevailing erosion, with incorporation i n the ice body of much rock detritus eroded from the substratum. The relatively low level of energy output (mainly in the form of meltwater) means that sedimentation plays

[ L

-3 7

- ~ _ _ _ glacial retreat

&JI

preservational potential preservat2::l

potential

P

Tie,: o climatic change

,

~

glacial advance

little deposition

much erosion

+2-,

preservational potential

Fig. 16. Main factors controlling the preservational potential of glacigenic sediments in relation to ice advance and retreat.

p


31

a comparatively minor role: fairly few sediments are formed and their preservational potential is limited. On the contrary, deposition prevails during regression of the ice when melting of debris-laden ice increases due to more intense ablation or t o lack of 'fresh' ice as a result of decreased snowfall in the accumulation areas. The deposits formed during retreats of the ice have a fairly good preservational potential, although they may soon afterwards become eroded during a recessional re-advance of the ice (Schliichter, 1983). The large net deposition during regression of the ice mass is only a small part of the total energy output of the glacial system under these conditions; much more energy is lost in the form of meltwater. Influence of ice dynamics and extent upon sedimentation Debris transported by ice can be found far beyond the outer limit of the farthest ice extent because meltwater streams and winds take over the transport activity. Truly glacial deposits, however, can only be found in the areas covered by the ice. The ice cap is not simply moving towards a final point then again retreating during one ice age: there are many oscillations with extending (transgressive) ice masses, separated from each other by recessions (regressions). Although much is known about the physics of glaciers (see, among others, Paterson, 1981), the fluctuations in ice extent are still a matter of speculation (Mickelson et al., 1981). Both regional uplift or subsidence of the Earth's crust and sea-level changes (which themselves are partly a result from glaciation and deglaciation; see, e.g., Walcott, 1970; Andersen, 1979; Vorren and Elvsborg, 1979; Sollid and Reite, 1983) may play a role (Edwards, 1978; Miall, 1984). It is likely, however, that autocyclic and allocyclic large-scale climatic changes are much more important. The mechanism behind these changes is still under discussion, although Milankovitch's (1920, 1930, 1936, 1938) views concerning astronomical factors now seem fully justified; the main problem is that other factors must also play a role, but these are not yet well enough known to be included in clear and detailed models. The climatic fluctuations contribute much to the characteristics of the glacigenic facies because they induce sea-level changes and isostatic movements, influencing both the erodibility of the source area and the characteristics of the depositional basins. A direct relationship between climate and, for instance, glaciofluvial deposits is nevertheless not really traceable. An important role is most probably also played by weather fluctuations (difference between day and night, and seasonal changes), but this role is even much more difficult to specify.

32


The character of the sediment input Accumulation of glacial deposits may take place gradually if the dynamics of the ice remain more or less stable and if there are no major changes in climate. A much more abrupt type of deposition may occur if debris concentrated in englacial crevasses is suddenly set free, for instance by rapid melting of a last remnant of ice underneath the crevasse, a process which can be triggered by complex factors such as ice characteristics, ablation conditions and local topography (Fig. 17). Such 'triggered' sedimentation is relatively common in ice lobes that extend considerably in front of the main ice mass. It will be obvious from the data presented above that transgressive conditions are characterised by a more or less uninterrupted sediment input, whereas an input of this type occurs in pulses during periods of regression.

Fig. 17. Character of sediment input in the glacigenic system, and major interrelationships between the factors that control the input.


33

THE INFLUENCE OF CLIMATE ON GLACIGENIC SEDIMENTATION Both the accumulation of snow in the firn basin and the melting of ice in the ablation zone are largely controlled by the climate. The dynamics of the ice mass depend on the energy balance that results from snow accumulation and ice melting (Fig. 18),which implies that the alternation of ice advances and retreats during a glacierisation also depends on this factor. Climate and weather thus influence the possibilities and character of glacial deposition and lead to differences between the various glacigenic facies. Temperature, precipitation and wind activity are considered the most important meteorological factors. Role of temperature Changes in the air temperature affect the ablation rate immediately, not only in the frontal area but over the entire supraglacial area that thus becomes covered with scattered detritus or even with a more or less continuous layer of debris (Sugden and John, 1976).

thermal regime of ice

net energy balance of

Fig. 18. Interrelationships between the main factors that determine the dynamics of an ice sheet or glacier.

34


A much more complex aspect is the influence of air-temperature changes, together with other meteorological and climatic elements, upon the thermal regime of the ice (Fig. 19). This regime must be considered as a complex function of the energy balance at the ice surface (Fig. 20). A cold or moderate regime influences the type of deposits formed (by influencing the depositional processes), whereas rapidly varying regimes (a common feature: Baranowski, 1977; Brodzikowski, 1987) dominate the dynamics and the changes in the depositional processes in the entire glacial environment. The ice dynamics are also strongly influenced by the thermal regime (Boulton, 1972a, 1979; Embleton and King, 1977). The four most characteristic ice-regime situations are presented in Figure 21, which is based on studies in recently glaciated areas and on studies carried out in the European Lowlands where Pleistocene glaciations left their imprints.

conditions of snow and ice accumulation

3

k

net energy balance of ice body

m

thermal regime of the ice body

depositional conditions

Fig. 19. Most important factors controlling the thermal regime of a n ice body.

The influence of climate on glacigenic sedimentation

of water v a p o u r

I/

i

35

friction in i c e

V heat i n p u t

L

1

heat loss t o atmosphere

net energy balance of ice surface

freezing of water

outflow Of ‘ w a r m ’ water

ablation

A

I1

vertical m o v e m e n t of i c e m a s s e s

equilibrium

zone

lht3 I

100 krn

cold thermal regime

I

Pielstocene glaclatton of N Asia and Canada

polar continental (high latitude)

B

cold

equilibrium

thermal

IhW

100 km

u

illtie preclpltatlon

polar continental (middle latitude)

Plei~toceneqiaciation of Europe

Pleistocene glaciation of Middel Europe I Southern Canada zone 01

D

50 km

surgtng

L_

/ / / / / / / / / / / / / / / / / / / / , , / , / I // / / / / ,

’,/

I / ,

,

Pleistocene glaciation of mountains and their forelands

Fig. 21. Hydrological and thermal regimes of large ice bodies (after Baranowski, 1977). The four possible (main) possibilities (A-D) are presented in simplified form.

36


Role of precipitation The glacigenic environments are characterised by subpolar, polar and Arctic climates. It is most important whether a cyclonal or an anticyclonal circulation prevails (Fig. 221, since this factor influences strongly most of the meteorological parameters, particularly the rate and type of precipitation, which parameters determine the type and the intensity of the ablation process.

Fig. 22. Reconstruction of two phases of ice extent in northern Europe, with emphasis on the pattern of cyclonal circulation. Black arrows indicate prevailing routes of the cyclones; dashed areas indicate the ice covers. Above figure: Karelo-Barentz ice sheet has grown together with the Scandinavian ice sheet; new centres of glaciation are developing in Ireland. A subarctic climate prevails in middle Europe; cyclonal activity becomes minimal and thermal continentalism increases. The position of the Karelo-Barentz anticyclone area is very stable. Snow accumulation decreases distinctly.


37

Cyclonal atmospheric circulation tends to result in a high precipitation rate. This means more snow in the accumulation area and more rain in the ablation zone. Increasing precipitation rates have a complex effect on the energy gradients of meltwater streams (Sugden and John, 1976; Baranowski, 1977; Embleton and King, 1977) and therefore also on the characteristics of the glaciofluvial deposits. Detailed palaeoclimatological reconstructions (regarding the palaeocirculation in particular) have shown that, during the Pleistocene, some

Fig. 22 (continued). Above: Phase of maximum ice extent in Europe. The Karelo-Barentz ice sheet has grown considerably, but has also split up locally. The maximum gradients in atmospheric pressure are situated between the centre and the margin of the ice sheet. The precipitation on the ice-covered area decreases again; the cyclonal circulation in middle Europe increases.

38


middle European areas were under the influence of prevailing cyclonal circulation for more than half of the year. This type of circulation produced meteorological conditions (Fig. 23) that caused very specific palaeoglaciological circumstances (Fig. 24), for instance wet and dynamic conditions over the entire extended ablation zone of the ice (which had a temperate thermal regime). These conditions resulted in a fairly constant and high accumulation rate.

Fig. 23. Palaeoclimatic reconstruction of the most common weather conditions i n middle Europe during the optimum of the Drenthian ( = maximum Pleistocene) ice extent, based on a palaeosynoptic model. A, B, C, D: precipitation zones. ACA = arctic cold air; PTA = polar temperate air.


A

L external zone 01

I

I

I

C

model o f ablation area

I

ice sheet

39

palaeocirculation

A

--

1 - zone 01 polar cyclons

--I

Fig. 24. Palaeoclimatological reconstruction, based on palaeosynoptic models, of the ice-marginal zone in the DDR and the Sudetic Mountains during Elsterian a n d Drenthian times (maximum Pleistocene ice extent: Dnieprovian). A: palaeoglaciological model of the ice-marginal zone. B: ablation and accumulation. C: palaeosynoptic model. PTA = polar temperate air; PCA = polar cold air; WF = warm front; CF = cold front; AF = arctic front; As = altostratus; Cb = cumulonimbus.

There is much less precipitation if anticyclonal circulation prevails. In combination with low temperatures, such circumstances induce a significant increase of ice sublimation in the ablation zone. This process affects the position of clasts in the upper layers of the ice and, if supraglacial deposits are finally formed, could lead t o particular lithofacies characteristics (Sugden and John, 1976; Shaw, 1977a). Such conditions prevailed in Eastern Europe during the Pleistocene periods of maximum ice extent. It is most probable that the zone where the Dniepr lobe was situated (the Dnieprovian is comparable to the Western European Drenthian) in particular witnessed a dominant anticyclonal circulation throughout the year (Fig. 22). The climate was therefore dry, cold and sunny (Fig. 25) and the ice sheet was characterised by a continuously cold thermal regime (Fig. 26). The intensity and the dynamics of the depositional processes were much lower than those in Middle Europe.

40


Fig. 25. Palaeoclimatological reconstruction of the most common weather conditions in Eastern Europe during the maximum Pleistocene ice extent (Dnieprovian), based on a palaeosynoptic model. See Figure 23 for explanations.

Role of winds Winds are primarily a result of atmospheric circulation. It should be emphasised, however, that the transitional zones between ice-covered and ice-free regions affect the wind pattern and the wind intensity. The ablation zones are often characterised by much wind activity, especially in the zone of cyclonal circulation (Fig. 21). Winds are not only responsible for the formation of regional or local aeolian deposits, but also change the surficial humidity in the sedimentary basins by vaporisation

41

The influence of endogenic forces

model of ablation area

L external zone of ice sheet A

L -

-

-150-250 km

0

~

C Lm

palaeocirculation

L L

low-pressure

co

45%N

oess

polar tropopause

co 50e N

Fig. 26. Palaeoclimatological model of the ice-marginal zone of the extremely continental Dniepr lobe (Soviet Union). See Figure 22 for explanations.

of surface waters (and of glacial ice as well if the temperature is low and sufficient insolation takes place). If surface waters are rare or absent, the dry winds may easily carry away the finest particles from the sediment cover in front of the ice. This may result in dust clouds that can be transported over hundreds of kilometres. Winds therefore greatly influence the depositional pattern in large parts of the periglacial environment (Jahn, 1950,1970; Cegla, 1972; Rozycki, 1979).

THE INFLUENCE OF ENDOGENIC FACTORS Extending ice sheets do not discriminate between tectonically active and more stable regions. Consequently, the ice caps may cover a rising or subsiding substratum, accompanied or not by earthquakes. The upheaval or subsidence of the substratum is of special importance in this context because it influences the depositional pattern under the ice, whereas it also determines to a large degree the preservational potential of the glacial deposits.

42


Another endogenic factor, of even more importance for the behaviour of the ice, is the Earth's heat flux, which greatly influences the energy balance of the ice. Since ice is a good heat insulator, much of the heat coming from the Earth's interior is absorbed by the ice, sometimes giving rise to melting of considerable masses and thus to subglacial streams and spaces where sedimentation can take place. It can be stated that, in general, endogenic processes affect not only the physiography of glacigenic sedimentary basins, but also the intensity of local thermal - subglacial - subrosion, the volume of the sediment output, the ratio between sedimentation and erosion, the character of redeposition processes and the frequency of facies changes in space and time. Vertical movements of the Earth's c r u s t Vertical movements have three important aspects in the framework of glacigenic sedimentation: they influence the energy input into the system (Fig. 27), they influence the preservational potential of the deposits underneath the ice cover, and they (may) influence the lateral extent of the ice masses. Crustal movements tend to influence the borders of sedimentary basins in general. The same holds for glacial sedimentary basins. The location of the movements (under the ice sheet or in front of it) is obviously of the greatest importance. If upheaval takes place underneath the ice, erosion of the substratum will increase, thus enriching the subglacial zone in detritus. The eroded particles may later become part of the englacial subenvironment (by transport along shear planes). In general, deposition in the periglacial environment will profit from these circumstances. In contrast, subsidence of the substratum underneath the ice will diminish erosion, finally possibly resulting in a reduced sediment supply to the periglacial environment. The same subsidence provides better depositional circumstances, however, within the subglacial environment, thus increasing the preservational potential of the subglacial deposits. If crustal movements take place in front of the ice (e.g., because of isostatic compensation), the depositional pattern in the periglacial environment may be affected. Subsidence in front of the ice results in basins where meltwater deposits may accumulate, but a t the same time such a subsidence may accelerate the ice advance, thus resulting in a n overriding and possibly in erosion of the sediments deposited earlier. Upheaval in front of the ice may result in stagnant ice because the barrier thus formed cannot be passed by the ice until the barrier can be overriden.


from the glacial system

43

output of meltwater from the glacial System

Fig. 27. Energy input in the glacigenic depositional system by vertical tectonic movements of the substratum.

Earthquakes Earthquakes affect the glacial environment in two ways. First, they represent sudden movements (faults) of the Earth's crust, forming o r reactivating zones of weakness where a n increased heat flow from the Earth's interior t o the ice mass may take place. This accelerates melting of the ice in the ablation zone or may induce melting where this process would not otherwise have taken place. Earthquakes may also disturb the equilibrium within the ice or the sedimentary cover; the latter (Fig. 28) may result in distinct structural changes (Brodzikowski et al., 1987b,d).The ice movement may thus undergo a sudden pulse a t the beginning of a changing thermal regime. The pulse may become visible because of a relatively fast advance of the ice front over several kilometres. Disturbance of the equilibrium within

44


IV

0

rnax

r l ~

P o

d

b

D

o

rnax

Fig. 28. Deformation horizons (D) within Elsterian and Saalian sediments due to earthquakes in the Kleszcz6w graben (central Poland). 1 = glacial till; 2 = fluvioglacial sediment; 3 = glaciolacustrine sediment; 4 = relative scale for intensity of the endogenic activity; 5 = endogenic activity; 6 = earthquake-induced deformation horizons; 7 = distinct changes in sedimentary conditions; 8 = relative scale for abruptness of facies transitions; 9 = facies transitions; 10 = sedimentary cycle; 11 = sedimentary subcycle; 12 = horizon with large-scale deformations.

the sediment may be expressed by mass movements from topographic heights or by destruction of barriers responsible for the existence of glacial lakes. A sudden outflow of lake waters may not only result in specific (‘catastrophic’)deposits but may also affect the depositional pattern in the terminoglacial and proglacial environments.


45

The geothermal heat flux The geothermal heat flux may be relatively high in zones with active faulting, but it may also be high in zones with other endogenic activity (6ermak and Rybach, 1979). Such zones tend to be of limited extent. Other areas, however, do not have a uniform (lower) heat flux, but show regional or even local differences. This has as a result that the values for the heat flux form a kind of mosaic (Fig. 29). Consequently, the heat flux has a diverse influence on the permafrost, the subglacial ablation, the quantity of waters in the subglacial zone and the dynamics of the ice masses. Extreme situations can result in an increased regional advance of the ice and in high flow velocities. The ice dynamics and the ablation rate are generally dependent on the heat flux, thus influencing the depositional conditions (Fig. 30).

2 Fig. 29. Geothermal heat pattern (in mW.m- ) in Europe (modified after Eermak and Rybach, 1982).

46


The heat flux may induce collapse of the basal ice masses if part of the ice has melted away (Rubulis, 1983). This process may be restricted to a few centimetres, but in extreme cases some tens of metres may be involved. If collapse structures are found, i t is difficult - were it possible to determine the mechanism responsible because other processes can result in similar structures (Eissmann, 1975, 1981). It is obvious, however, that collapsing in the subglacial subenvironment will greatly influence the depositional conditions.

Fig. 30. Influence of the geothermal heat flux upon glacigenic depositional conditions.

The sedimentary facies

47

THE SEDIMENTARY FACIES The history of the term 'facies' reaches back a long way. According to Walker (1984),it was first used in geology in 1669 by Nicolaus Steno but it received its modern meaning from Gressly (1838). Much more precise elaborations appeared later (by, e.g., Walther, 1894; Teichert, 1958; Weller, 1958; Krumbein and Sloss, 1963); Middleton (1978), Reading (1978b) and Walker (1984)have provided the most recent definitions, with comments and discussions. In spite of the clear definitions available, the term has been applied (and misapplied) in several ways. It should, however, always indicate a geological unit or a number of geological units with specific features in common. Lithofacies, biofacies and geochemical facies are examples definable by means of parameters that can be determined unambiguously. The type of facies that should be used depends on the purpose of the research involved. The sedimentologist Middleton (1978) states that it is understood that (the facies) will ultimately be given an environmental interpretation". The present authors use the term 'lithofacies' (cf. N. Eyles et al., 1984b, 198813; Shaw, 1987a) where the rock type is considered (mineralogical composition, grain-size distribution) and the term 'sedimentary facies' where an environmental analysis is involved. The sedimentary facies, the main type of facies dealt with in this book, is not entirely unambiguous, for it is determined by the prevailing depositional process(es) and thus requires interpretation. As will be seen, this use leads t o specific facies types, e.g., 'melting-ice facies' and 'proglacial Iacustrine facies', being distinguished. This approach thus requires more than mere description of specific characteristics that can be observed directly. Specific characteristics may help in determining the correct type of sedimentary facies, but more frequently the lateral and vertical transitions of facies must be studied t o ensure a reliable interpretation (Brodzikowskiand Van Loon, 1983,1987). The term 'sedimentary facies' is thus intermediate between 'environment' and 'lithofacies'. This bridge function is essential in sedimentological analyses, because a (sub)environment may include a wide variety of lithofacies (e.g., the subglacial subenvironment with poorly sorted diamicts, varved glaciolacustrine deposits, etc.), whereas specific lithofacies (e.g., poorly sorted diamicts) may occur in a wide variety of (sub)environments (supraglacial, englacial, subglacial and terminoglacial). It is thus necessary t o provide details about environmental conditions, depositional mechanisms and lithological characteristics; this is achieved by means of the sedimentary facies. 'I...

48


Facies analysis Lithofacies are represented by actual deposits. Specific lithofacies have names that are usually, although not always, sufficiently clear. It is obvious, for instance, that a deposit consisting mostly of quartz grains that have a grain-size distribution almost entirely in the 64-2000 micron range can be called a sandstone. It is less generally known that a matrixsupported, massive, sheared sediment with a wide grain-size range can be called a diamict. On the other hand, it is far from obvious from the lithofacies data presented above to which sedimentary facies the sandstone belongs, while the other sediment is obviously a lodgement till, thus belonging to what is termed here the 'subglacial melting-ice facies'. Sedimentary facies are thus described and analysed in order t o establish which parameters determine the regularities and variations within the sediments. The depositional basin is therefore investigated as regards its environmental characteristics, including the palaeogeographic reconstruction and interpretation of the depositional mechanisms (Potter and Pettijohn, 1963; Reading, 1978b; Miall, 1984). Although a sedimentary basin is in many respects an entity, a wide variety of sediment types may be deposited. This is due to physiographic differentiation and to changes in prevailing processes in time and/or space. The various facies also are not stable in time once they have been formed: complete lithological units may become eroded and the boundaries between adjacent facies types may shift (e.g., due to the gradual growth of a delta). Sedimentary facies differ from each other by their lithology, extent, structures, energy vectors, etc. (Allen, 1970b, 1982; Friedman and Sanders, 1978; Leeder, 1982; Miall, 1984). All these parameters are related to processes or combinations of processes that may have changed, either gradually or abruptly, simultaneously or one by one, either t o reach a new stable value or t o continue changing. Important parameters that may be changed are the amount and nature of sediment supplied, the climate, the height of the sea level and the stability or instability of the substratum. Facies analysis as understood by Selley (1970), Miall (1973) and Walker (1984) is based largely on statistical methods. Selley (1970) aims a t the presentation of facies associations and sequences in a clear, objective, graphic manner characterising both facies interrelationships and facies patterns. This can be done by tabulating the numbers of specific transitions observed, converting these numbers into relative frequencies, calculating a matrix with the assumption of the null hypothesis (i.e. that such transitions are random and that they depend only on the relative

The sedimentary facies

49

abundance of the facies that are being studied) and, finally, by establishing random probabilities t o produce a matrix emphasising any differences from random which are at large (Walker, 1984). A detailed analysis of local or regional facies changes could be hampered by a lack of outcrops. Lowland areas (where most Pleistocene glacigenic sediments have been studied) particularly tend to be poorly exposed. This makes i t all the more necessary to have plentiful information about facies associations and sequences. If there are too few exposures and if borings cannot provide the data required, it might be useful t o first study comparable facies i n a better exposed area or i n hard-rock equivalents that have already been investigated in detail (cf. Vanney and Dangeard, 1976). Descriptions of the characteristics of lithified glacigenic rocks concern glacigenic conditions from many ages. Facies data from Precambrian glacigenic sediments have been provided by, among others, Coleman (19071, Bjorlykke (1969), Lindsey (1969, 1971), Roscoe (19691, Young (1970, 1973, 1974, 1978, 1981), Aalto (1971, 1981), Spencer (1975a,b, 1981), Deynoux and Trompette (1976, 19811, Edwards (19761, Nystuen (1976), Sumartojo and Gostin (1976), Nystuen and Seather (1979), Anderton (1980, 1982), Gravenor (1980), Boulton and Deynoux (19811, Chumakow (19811, Edwards and Foin (1981), Legun (1981), Link and Gostin (19811, Donaldson and Munro (1982),Hambrey (1982),Stupavski et al. (1982),Anderson (1983),Christie-Blick (1983),C.H. Eyles and N. Eyles (1983b, 19851, Fairchild (1983,1985),Miall (1983a, 19851, Gravenor et al. (19841, Dowdeswell et al. (1985) and Fralick (1985). Similar data on Cambrian and/or Ordovician glacigenic rocks were presented by, among others, Spjeldnaess (1973), Tucker and Reid (19731, Davies and Walker (1974), Deynoux (1980), Hein and Walker (1982) and Fortuin (1984). Facies data on Carboniferous and/or Permian glacigenic deposits a r e numerous; to mention only a few: Rattigan (1967), Frakes and Crowell (1969), Le Blanc Smith and Eriksson (1979), Bull et al. (1980), Davis and Mallett (1981), W.K. Harris (1981), Jackson and Van de Graaff (19811, Rogerson and Kadybka (1981), Casshyap and Tewari (1982),Visser (1982, 1983b), Visser and Kingsley (1982), Coretelezzi and Solis, 1983; Cuerda (1983), Gonzalez (1983), Gravenor and Rocha-Campos (1983), Stauffer and Peng (1984), Visser et al. (1984, 1986, 1987), Visser and Hall (1985), Visser and Loock (1987), and S.Y. Johnson (1989). Glacigenic facies of Tertiary age have been described by Plafker and Addicott (19761, Dalland (1977), Plafker (1981), Barett and Powell (19821, McKelvey (19821, Minicucci and Clark (1983) and C.H. Eyles (1985). An overview of the chronology of glaciations has been provided by Harland (1981).

50


FACIES INTERPRETATION The analysis and interpretation of facies require both careful sampling of the data, and the development (or application) of a model which must fit the various data. This, of course, also holds for glacigenic sediments. The data that should be collected in the field and the laboratory comprise the lithology (including grain-size analysis, mineralogy and petrography), inventary of sedimentary structures and of early-diagenetic deformations, geometry and size of the various units, signs of erosional surfaces, type of contacts between the various units, and palaeocurrent or ice-movement directions (Hill and Prior, 1968). These data must first be interpreted in terms of depositional (and erosional) processes. Once the interpretation is completed, a logical framework must be found to explain the vertical and horizontal transitions. This means that much attention must be directed to the rclative abundance or scarcity of specific features, their associations and other interrelationships.

Lithological characteristics The lithology of sedimentary units must be determined because it can facilitate the correlation between various outcrops. Determination of the extent and of the lateral and vertical transitions is important for a n environmental reconstruction. Grain size, mineralogical composition, sedimentary structures, deformations and palaeocurrent indicators are also helpful tools if the depositional history is t o be reconstructed. Size and geometry of the units The size of lithological units (thickness and areal extent) depends on the size of the depositional basin and on the basin development, the depositional rate (net sedimentation rate), the duration of the depositional process(es) and the possible erosion afterwards. There is generally insufficient information about these parameters to estimate their relative contribution. Nevertheless, it seems worth paying more attention to these aspects as our insight into the depositional process of glacial sedimentation might thereby be much improved. Why, for instance, are most Pleistocene tills only a few metres to maximally some tens of metres thick (admittedly, there are Wisconsinan till sequences with a thickness of several hundred metres) whereas Precambrian tillites often seem to reach much greater thicknesses? Also how did some Miocene glaciomarine

Facies interpretation

51

tillites accumulate to several thousands of metres? Much research must still be done t o find answers t o questions such as these. The geometry of a deposit depends on the shape of the depositional basin, the depositional mechanism, the interrelationship with adjacent depositional areas, and erosion. The same problems as mentioned for the size of the deposits however still arise. In spite of these uncertainties, size and geometry together can give rather reliable information, particularly if trends in grain-size distribution are also taken into account. Contact characteristics The characteristics of the contacts between adjacent lithofacies are important for facies analysis. Aspects that should be investigated in particular are the geometry of the contact plane, the type of contact (erosive or non-erosive), deformation of the contact plane, etc. Erosive contacts - Erosion is part of almost all depositional processes. Sedimentary breaks are therefore very common. It is most important, however, t o recognise the erosional contacts that point to a process other than sole alternation of sedimentation and erosion as an ongoing process. The importance of 'real' erosional contacts had already been emphasised by Walther (1894). Nevertheless, the interpretation of this phenomenon still receives insufficient attention. Erosion in glacigenic sediments is a most important feature because glacial erosion may indicate various stages of ice (re)advance whereas, in various other types of deposits, erosion may indicate subaerial exposure and mass wasting along a slope. Non-erosive contacts - Non-erosive contacts can be either sharp or gradual (sometimes also called 'progressive'). The nature of the contact may be an indication of transgressive or regressive development: transgressive phases give commonly rise t o relatively many sharp contacts whereas regressive phases tend to lead t o more gradual contacts. The nature of the contact cannot be taken as a criterion, however, because the underlying processes (commonly changes in hydrodynamic properties) occur in a rather unpredictable way. Alternations of sharp and gradual contacts in fine-grained sediments may give more indications, as in the case of varves. Obviously, non-erosive contacts may show structures that point t o at least a small break in sedimentation. Such structures include outwash phenomena, sole marks (groove casts, prod marks, etc.), raindrop

52


imprints, etc. Other changes in the sediment, e.g., in colour, consistency or cementation, may also indicate sedimentary breaks. This implies that the analysis of the sedimentary history requires more than a rough impression of erosional or non-erosional contacts (cf. Reading, 1978a; Reineck and Singh, 1980; Allen, 1982; Gradziiiski et al., 1986). This problem has been dealt with in more detail by Twenhofel (19391, Shrock (19481, Kuenen and Menard (1952), Sanders (1960), Ksiqikiewicz (19611, Diulyiiski and Sanders (1962), Diutyliski (1963b, 1965), Diu€ydski and Walton (1965) and several others in more recent years.

Grain size A grain-size analysis may reveal possible sedimentation mechanisms (or, a t least, exclude some mechanisms) Wisher, 1969). The grain-size data, although rarely unambiguous, may thus be used for hydraulic interpretation (Glaiser et al., 1974). Experiments in this context seldom yield reproducible data (Harms and Fahnestack, 1965),partly because granulometry depends on various parameters such as bed form and local flow regime. The distance from the source also plays a role (Teisseyre, 1975),so that, if other data are lacking, granulometric data may also be used t o reconstruct a palaeocurrent direction (Middleton, 1965; Reineck and Singh, 1980; Gradzifiski et al., 1986), because the coarsest particles will generally remain closest to the source. Much less is known about the relationship between grain-size distribution and glacigenic melt-out or subaerial mass movements (although an increasing number of detailed studies into glacigenic diamicts have been published in the last few years); some sedimentary structures found in deposits formed under glacigenic conditions can therefore not be explained properly. Analysis of the grain size is also important because of the influence on the geotechnicaUengineering characteristics of the sediment (see, among others, Boulton, 1976a; Lee and Focht, 1976; Brand and Brenner, 1981; Browzin, 1981).

Mineralogy

A mineralogical analysis of the sediments (or a petrological analysis if coarse clasts are concerned) is useful for the determination of the source area (Zandstra, 1983). Recognition of the source area is important for palaeogeographical reconstructions because it allows transport routes to be found (Di Labio and Shilts, 1979). It should be kept in mind, however,


53

that the mineralogical composition of a glacigenic sediment is almost always the resultant of erosion in the source area, erosion during ice movement, and erosion in the neighbourhood of the final depositional site, thus giving a mixture of assemblages, each of which must be recognised as such. Mineralogical analyses are commonly restricted to heavy minerals (which give rather reliable and easily obtainable results). Much more time-consuming and specialised equipment, requiring trace-element or trace-mineral analysis, can however provide much more precise data. The petrological characteristics of clasts especially may change from bottom to top within one lithological unit, either because of mixing ice masses from different sources, or due t o different processes occurring within one ice mass, dependent on the location of the clast (sub-, en- or supraglacial, embedded in a relatively rigid ice mass or located in a shear zone, etc.) (Haldorsen, 1977; Hallet et al., 1978; Slatt and Eyles, 1981; Houmark-Nielsen, 1983a). Nevertheless, the petrography may give indications about the source area and thus about the ice movement (Meyer, 1983a; Schuddebeurs and Zandstra, 1983).

Sedimentary structures Sedimentary structures (see, e.g., Collinson and Thompson, 1982) should be inventoried because they give valuable information about both the depositional process(es) and the palaeocurrent directions. This is particularly true for aeolian sediments and deposits formed in current water. Much less is known about the significance of the various, often vague and rather irregular, structures that can be found in the most characteristic glacigenic deposits: the diamicts. It is not unlikely, however, that the lack of generally accepted interpretation of such structures has lessened the interest of researchers who are not primarily interested in this specific problem. The authors are of the opinion, based on their own field investigations, that a much more systematic inventory of structures in diamicts might contribute greatly to a better understanding of the genesis of these sediments.

Deformation structures Deformation structures are fairly common in glacigenic sediments. They range from simple undulations to complex multi-phase discontinuities and may be formed by a process or a number of processes that can be grouped (cf. Van Loon, 1990) into the following ten categories: bioturbations, cryoturbations, glaciturbations, thermoturbations, graviturbations,

54


hydroturbations, chemoturbations, atmoturbations, endoturbations and astroturbations. A detailed analysis of the penecontemporaneous and postdepositional early-diagenetic deformations (Fig. 3 l),including determination of their relative age and frequency, could give an insight into the dynamics of the environment during (or shortly after) deposition, or into the processes affecting the sediments afterwards. This body of data could form a n elegant though not always reliable basis for determining the genesis of the sediment in a particular (sub)environment or facies when there are insufficient other data. As an example, some deformational structures, while they are not diagnostic, can be characteristic of specific circumstances. This holds, e.g., for the joints formed in subglacial diamicts due t o loading and subsequent unloading by the overlying ice cover; the type of discontinuities and their spatial distribution may help englacial and subglacial diamicts to be distinguished. Some types of structures are signs of 'en masse' reworking before final deposition of the sediments.

Fig. 31. Deformations in a sand quarry near Ossendrecht (The Netherlands), possibly due to a combination of load casting and cryoturbation.


55

Palaeocurrents Palaeocurrent indicators are most important for reconstruction of the palaeogeography. One should keep in mind, however, that traces left by palaeocurrents can vary widely: meandering streams, for instance, have current directions that may be opposite, even a t relatively small distances from each other (Teisseyre, 1977, 1978a,b, 1980, 1984). It is therefore essential t o measure as many palaeocurrent indications as possible if a reliable regional picture with prevailing directions is to be obtained. Palaeocurrent directions may be reconstructed in various ways, mainly depending on the depositional and/or erosional processes that took place. This implies, for instance, that the approach in melt-out facies must be different from that in glaciofluvial, glaciolacustrine or aeolian facies. Consequently, a vertical section may require different analytical methods for the various units (cf. Gradzifiski et al., 1986). A proper analysis of the palaeocurrent data should not only yield information about prevailing transport directions (and thus about the direction of the source area) but also should show the relative frequency of changes in the hydrodynamic regime, in the morphology of the substratum and in the dynamics of the depositional process. There are large numbers of traces from which palaeocurrent directions may be deduced. These include: the orientation of the foresets (Fig. 32) in current ripples or wind ripples (Momin, 1968; Kumar and Bhandari, 1973), gradual horizontal changes in the average and/or maximum grain size (Agterberg et al., 1967; Miall, 1974), orientation of objects (imbrication of pebbles, orientation of shells: Van Loon, 19721, depositional 'shades', sole marks such as flute casts (Pelletier, 1965), etc. .4 large number of structures may, however, only show the axial direction of the transport, such as glacial striae (Fig. 33) (Von Brunn and Marshall, 1989; Visser, 1990), prod marks, sole marks such a s groove casts, parallel orientation of plant debris, etc., thus requiring additional data if a definite conclusion is to be drawn. It is most probable that some soft-sediment folds may be induced or a t least influenced by palaeocurrents (Johansson, 1965; Griffiths, 1967; Parkash and Middleton, 1970; Teisseyre, 1975; Potter and Pettijohn, 1963,1977). The morphology in the glacial and periglacial environments is, in general, rather complicated (Embleton and King, 1975; Sugden and John, 1976). This results in a complicated pattern of palaeocurrent directions, which can be interpreted correctly only if sufficient data are available. Use of a n unduly small amount of palaeocurrent data could hide rather than unravel the palaeogeography.

56


Fig. 32. Climbing ripples (ripple-drift cross-lamination in glaciofluvial sands (quarry Eggestedt Nord, 20 km north of Bremen, Federal Republic of Germany). The ripples are good palaeocurrent indicators (current from left to right).

Fig. 33. Glacial striae made by a Pleistocene mountain glacier in the wall of a valley near Tabescih (central Pyrenees, Spain).


57

Textural characteristics Textural characteristics of the sediments include the nature of the surfaces of the grains, their shape, their roundness and their orientation (fabric) within the sediment. These characteristics can be studied in the field as far as clasts are concerned, but grains of sand size or smaller need to be studied with a binocular, hormall microscope or even with a SEM (scanning electron microscope) (Bull, 1981). Such textural studies are not specific for glacigenic sediments and will therefore be dealt with briefly. In general, texture may give indications about the processes that the sedimentary particles have undergone. Surface analysis, for instance, may provide indications of aeolian transport whereas the roundness may provide information about the transport of the particle by currents or waves. An aspect that is quite typical of glacial sediments and therefore deserves special attention is the degree of weathering of large clasts. The occurrence of strongly weathered clasts (often granitic boulders) that crumble as soon as they are isolated from the deposit (Fig. 34),is a fairly

Fig. 34. Completely weathered erratic in a Weichselian moraine near Wartenberge (Federal Republic of Germany).

58


strong argument favouring transport of the clast while embedded in ice (Embleton and King, 1977; Embleton and Thorns, 1979). The fabric of diamicts, and of other glacigenic sediments also, is important because it allows the prevailing stress conditions during sedimentation to be reconstructed (Richter, 1930,1932; Seifert, 1954; Lawson, 1979; Prange, 1983; Dowdeswell and Sharp, 1986). One should, however, do this with care because postdepositional processes (ice pushes, compaction, etc.) may affect the original fabric.

Occurrence Palaeogeographical reconstructions of glacigenic areas require that the spatial ( = lateral and vertical) relationships of a specific unit with other deposits be established. The preservational potential (units may have disappeared completely by erosion) is most important in this context.

Preservational potential Deposits formed under different conditions tend to have varying preservational potentials (Reading, 1978a; Gradziiiski et al., 1986).Few sediments are preserved without being affected by erosion (or at least abrasion). Various glacigenic sediments tend to have a rather small chance of surviving erosion (N. Eyles, 1983~). Energy changes are a factor of prime importance as far as the preservational potential of a sediment is concerned. Moving ice masses represent a giant amount of energy, which implies that it is the sediments that are directly or indirectly influenced by active ice that tend t o undergo erosion. Deposition may prevail locally in the 'shadow' of a barrier, if the substratum is subsiding or if the erosional base is changed (e.g., by a eustatic sea-level rise). This is of more importance than momentary climatic or meteorological conditions for determining the preservational potential. A limited preservational potential will usually be expressed by a relatively large number of erosional phases. The ratio between erosive and non-erosive contacts might therefore be a measure of the preservational potential but insufficient data are available to rank the various types of glacigenic deposits according t o this parameter.

Horizontal and vertical facies associations The various facies and their deposits in the glacial and periglacial environments commonly show well recognisable relationships touching


59

their horizontal and vertical transitions into each other. This is, of course, due t o the gradually changing boundaries between the facies resulting from the logical succession of depositional and erosional processes. Walther's facies law already recognised this in the 19th century. Glacigenic facies face yet another changing parameter: climate. Even relatively small fluctuations in the average temperature or precipitation may induce significant changes in depositional patterns and should therefore be considered as an important factor influencing the distribution of facies in space and time (Boulton, 1972a; Sugden and John, 1976). The normal sedimentary processes and the climatic fluctuations are the main reasons for the common occurrence of closely interrelated facies, both vertically and horizontally. Such groups of facies that apparently have a number of elements in common, are called 'facies associations'. A well known example is the association of proglacial deltaic and lacustrine facies with scattered erratics supplied by melting ice masses in the lake. Such deposits from associated facies may, in turn, become included in the tills of an advancing glacier (N. Eyles, 1983a) and become part of diamicts. Sequences - The term 'sequence' is commonly used when facies associations form a vertical succession. A sequence consists of a succession of lithological units with gradual, sharp or erosive contacts, formed by an uninterrupted, more or less predictable series of depositional processes which occurred at a specific place due to a set of depositional conditions that changed according t o a logical depositional model. Characteristic examples of sequences are the coarsening upward sediments of deltas (Oomkens, 1967; Van Loon, 1972) and the fining upward fluvial sequence (Allen, 1965; Kessler and Cooper, 1970; Leeder, 1973; Harms et al., 1975, 1982; Cant and Walker, 1978; Bluck, 1980) (Fig. 35). Sequences need not be characterised by changing grain sizes: the sedimentary structures can also change as to frequency, nature or direction, or fossil assemblages may appear or disappear, etc. Whether such changes occur gradually or suddenly, and whether they take place frequently or rarely, is an indication of the underlying processes and therefore often provides a clue for the interpretations of environmental changes. In general, the wide variety of facies associations can only give clues for detailed interpretation if additional data are gathered. This is also true for glacial sequences (Crowell, 1978; Schwan et al., 1980; Beard et al., 1982), although the interpretation may raise severe discussions (Dreimanis, 1984b; N. Eyles et al., 1984a; Karrow, 1984a; Kennis and Hallberg,


60

I

f--

1

rooilet zone

coliche nodules

overbank floodplain deposits

E 0 Ni IC)

point- bar deposits

1

channel

cross-bedde d sondstonss

conq lome r a t ! c sandstone w i t h introclasts srosionol baae

Fig. 35. Idealised fining-upward fluvial sequence, as commonly found in the various glaciofluvial facies. Adapted from Pettijohn (1975).

1984). However, if the sequences iesult from a distinct and logical succession of depositional processes, they will usually be a key to the genetic interpretation.

Depositional mechanisms The glacigenic conditions are so diverging that a wide variety of depositional processes play a role. Each subenvironment and each facies is characterised by a specific combination of prevailing depositional processes, but the local conditions are so important that it is not possible to base a facies interpretation on the mere relative importance of the various processes that are presumed t o have formed the pertinent deposits. It should also be kept in mind that periods of 'normal' sedimentation may alternate with phases of 'catastrophic' processes. There is no general relationship between the relative frequency or duration of these different situations and the impact that they have on the final sediment. On the


61

other hand, 'rare' deposits may represent 'common' depositional conditions (with a low net sedimentation rate) and vice versa.

Normal and catastrophic sedimentation 'Normal' sediments are the net result of the depositional and erosional processes that prevail a t a given location under regular conditions. Such sediments increase in thickness at a rate that corresponds with the prevailing rate of net deposition, which factor depends on the general energy level. There may, however, occur short-term, incidental processes with greatly different energy levels, resulting in what are commonly called 'catastrophic' sediments (Reading, 1978a; Gradzifiski et al., 1986). The glacial melt-out process and the englacial and subglacial deposition of diamicts are examples of 'normal' processes, whereas subaerial slumps or subaqueous suspension currents are examples of the 'catastrophic' category. Drumlins may be associated with catastrophic subglacial floods (Menzies, 1989; Shaw et al., 1989). Vertical cross-sections through glacigenic deposits commonly show both types, suggesting that 'normal' and 'catastrophic' processes alternate more or less regularly. This is not the case, however, because long periods with 'normal' sedimentation and erosion can easily result in a much thinner succession than one momentary 'catastrophic' event. The relative abundance of 'catastrophic' sediments is therefore no indication of the frequency of such events but indicates only the energy and the transport capacity involved under these extreme conditions, and the preservational potential of both categories of deposits. This, however, does not exclude the possibility that 'catastrophic' events occur frequently; subaqueous slope sediments, for instance, may become reworked, redeposited, again reworked, etc. Flow tills may be composed of a number of lithological units that have undergone a n increasing number of reworking phases with increasing age; consequently, the oldest sediments in such a flow till may show a much more irregular character t h a n the youngest sediments involved, even if the most recent flowage process had not affected them in different ways.

Exceptional conditions - 'Catastrophic' sedimentation is commonly but not necessarily due to exceptional conditions; on the other hand, some exceptional situations may be difficult to reconstruct because they leave no traces or because what traces are left cannot easily be recognised as such. Nevertheless, recognition of exceptional conditions may be most important if the development of a glacigenic area is to be reconstructed.

62


Exceptional situations differ from catastrophic situations in that the latter may distinctly interrupt the normal depositional process but nevertheless be a part of the regular development. It is t o be expected and is therefore not exceptional that, e.g., areas with almost no vegetation may undergo fairly catastrophic sheet flooding at more or less regular intervals. Exceptional situations will arise, for example, if well developed vegetation arises locally in a sheltered area nearby the ice cover, resulting in organic-rich deposits. Such exceptional traces may have a strong influence on the reconstruction of a glacigenic development; one could even state that, in general, the more exceptional a find is, the more attention should be given t o the fitting of such data into the general model. It should always be kept in mind, however, that exceptional (or catastrophic) events may result in deposits that are not or that are hardly to be distinguished from 'normal' deposits. On the other hand, an exceptional combination of 'normal' factors may result in apparently exceptional deposits. There is no unambiguous method available to distinguish with 100%certainty 'catastrophic' deposits from 'normal' deposits. Common and rare types of deposits It is clear from the literature on glacigenic sediments that the various types of deposits involved occur with strongly varying relative frequency and extent. Local and regional differences are common but obviously there are also some general trends as regards the probability of finding a specific type of deposit. The two main reasons for this are well known: the frequency and extent of the various types of deposits may differ, and their preservational potential may be different. Original differences in frequency may represent differences in the dynamics of the prevailing processes, especially differences in energy gradients. Rapid alternations of high-energy and low-energy processes tend to lead to much erosion, resulting in a sediment of restricted thickness and extent - if any sediment is left (Zielinski, 1982b). The sediments formed along the margin of an ice cap are exposed to this set of conditions. Consequently, the remaining sedimentary pattern is often quite chaotic and difficult t o interpret. This implies that it may be helpful t o reconstruct the dynamics on the basis of the relative frequency of deposit types.

63


Sedimentary cycles Sedimentary cycles are due to repetitions of sedimentary conditions and therefore of sedimentary sequences. The sediments that form part of such cycles have been given various names, e.g., 'cyclites' and 'rhythmites' (e.g., Duff and Walton, 1962; Duff et al., 1967; Reineck and Singh, 1980; Gradzifiski et al., 1986).The sediments of one specific cycle are commonly called 'cyclothems'. Cyclic sedimentation has been described from several facies, among others from fluvial facies (Allen, 1964, 1970a; Beerbower, 19641, deltaic facies (Moore, 1959; Oomkens, 1967),lacustrine facies (Lambert and Hsu, 'nl

C

tm

0

5

lo I

10

E

9 3

10

6

12

Fig. 36. Characteristic examples of glaciodeltaic and glaciolacustrine cycles in the Jaroszow Zone (SW Poland). 1 = structureless coarse and medium sands; 2 = idem with cross-bedding; 3 = idem, with trough sets; 4 = structureless fine sands; 5 = idem with horizontal lamination; 6 = idem with climbing ripples; 7 = fine and medium sands with small-scale cross-bedding; 8 = silts with horizontal lamination; 9 = silts and clays with wave ripples; 10 = silts with wavy lamination; 11 = clay; 12 = varved clay; 13 = small-scale deformations. After: Brodzikowski and Van Loon (1983).

64


1979a,b), aeolian facies (Hunter and Rubin, 1983) and submarine fans (Maldonado and Stanley, 1976,1979). The characteristics found for these cycles under non-glacigenic conditions apply in principle also t o similar deposits formed in glacigenic areas (Fig. 36). Although sedimentary cycles occur relatively frequently - in glacigenic (Miller et al., 1977; Crowell, 1978; Beard et al., 1982) and glaciomarine (Mode et al., 1983) deposits also - precise interpretation often seems difficult and the controversies are evident from the literature. These controversies stem often from the differences in opinion regarding the position of the base in each cycle and thus the real cycle of processes. Some authors have thus suggested that the term 'cyclic sediments' be replaced by 'repeating sediments'. Attempts to approach the cyclicity problem on a more methodological basis (Zeller, 1964) have thus far not found much support. Instead of this, the main trend in sedimentology during the past twenty years has been the application of mathematical (statistical) procedures such as Markov chain analysis, factor analysis and probabilistic calculations. Cycles in glacigenic sediments can be found on a macro-, a meso- and a microscale, which implies that cycles may show their own subcycles. This is quite plausible since many characteristic glacigenic deposits are formed in the neighbourhood of the ice front, and the position of this front is subject to a large number of both smaller and larger fluctuations. Each fluctuation may result in a cycle (in fact a sequence) and each cycle may include deposits with their own cycles (e.g., varves in glaciolacustrine sediments).

GLACIGENIC FACIES MODELS Sedimentological field work, and basin analysis in particular, requires that models be established, verified in the field and finally rejected, or accepted as useful for further research. A generally accepted facies model constitutes, in its widest sense, a summary of a specific depositional environment (or subenvironment) or a closely related group of (sub)environments (Walker, 1984). Numerous facies models have been developed by sedimentologists throughout the world. Comparison of large numbers of such models shows that there exist models that seem t o be well applicable for most situations within a specific sedimentary environment. These descriptive models can be used as basis for more detailed and perhaps speculative models for a particular area.

Terminology and use of symbols

65

Models must combine all the information that can be derived from field data such as lateral and vertical facies transitions, the occurrence of sequences and/or cyclothems, energy gradients, erosional phases and sediment supply. This implies that glacigenic models must deal not only with the area (and the processes taking place) in front of a n ice cap, but also with the area on top of, within and underneath the ice. There do not yet exist good methods to study sedimentary processes within or underneath a n ice cap. Glacigenic models then, of necessity, include uncertainties, perhaps even more uncertainties than the models from any other sedimentary environment. Fortunately, our insight into glacigenic processes has increased considerably in the last few years and even though some details a r e impossible to verify, existing models appear to be sufficiently accurate to have a fairly good predictive value when regional studies are initiated. Laboratory experiments have been of great help for understanding the processes and the resulting sedimentary characteristics, although i t must be emphasised that such experiments are commonly carried out on a small scale (Rozycki, 1958); there are several indications that extrapolation of the experimental results t o full-scale conditions is not always feasible. The same holds for experiments and observations in 'natural laboratories' such as waste-dumping areas, alluvial fans in sand pits, tailings, etc., although observations made in such 'laboratories' under polar or subpolar conditions can indeed give reliable information about relatively smallscale processes. Experiments, field observations and theoretical analyses all have contributed t o the models of glacigenic facies. Such models obviously become less accurate as they become more detailed. The models t o be presented in this book will therefore be of two types: rather general models that can apply superficially to each situation dealt with, and much more detailed models that have as primary aim to show actual situations on a smaller scale.

TERMINOLOGY AND USAGE O F SYMBOLS Descriptions of glacigenic lithofacies by different authors are difficult to compare because each author tends to develop a terminology t h a t is most suitable for (1)his specific research interest and (2) the region of his work. Lithofacies are most commonly designated by letters and/or numbers: lithofacies A, B, C or 1, 2 , 3 or (with subdivisions) A-1 etc., when referred to in literature.

66


Several attemps have been made to improve communication between researchers by devising a generally applicable terminology. Miall (1977) and Rust (1978) designated lithofacies by a two-letter code characterising the lithology and the structure. These proposals made it possible to carry out relatively simple comparative studies and the concept was developed further by Miall (1978,1983a, 1985) and Eyles (1983,1985).

Classificationsystem used in this book A much more detailed classification system was elaborated in some steps by the present authors (Brodzikowski and Van Loon, 1980, 1983, 1987). This classification involved (1)environments and subenvironments (based on the spatial relation with the ice cap), (2) the depositional facies (based on the depositional conditions, in particular the depositional processes), and (3) the glacigenic deposits (based on the depositional mechanism). This classification proposal raised important discussions with fellow researchers, most of whom considered the approach very consistent and easily applicable in practice. There were, however, useful suggestions for adaptations. The authors therefore decided t o follow the same approach in the classification scheme in the present book, although with a number of adaptations. Four-level subdivision

The classification proposed by Brodzikowski and Van Loon (1987) comprises four levels, indicated by Roman numbers, capital letters, Arabic numbers and lower-case letters, respectively. The reader is referred t o following sections of this book for details. Only some schematic explanations will be provided in this subsection. The first level distinguishes between the glacial (I) and the periglacial (11) environments. The glacial environment is roughly the area with a continuous ice cover. The periglacial environment is not covered by ice (or is covered in a discontinuous way), but is still under the influence of the ice regime (meltwater streams, loess deposition or comparable features); the continental periglacial environment is the region characterised by a permafrosted soil. The second level (subenvironments) distinguishes parts of the two environments on the basis of their spatial relation t o the ice cap. The (continental) glacial environment, for instance, includes a supraglacial (I-A), an englacial (I-B) and a subglacial (I-C) subenvironment, situated on top of, within and underneath the ice cap, respectively.


67

The third level refers to the facies on the basis of the most characteristic depositional conditions (processes). Some adaptations of the earlier proposal (Brodzikowski and Van Loon, 1987) were made at this level: the suggestion that each specific facies type be indicated by the same Arabic number, irrespective of the subenvironment in which it occurs, was followed. Not all facies are present in all subenvironments, so that the consequence of this adaptation is the existence of 'empty' places in the scheme. The following facies were distinguished: melting-ice facies (Arabic number 1; the supraglacial (continental) melting-ice facies is therefore denoted as 1-A-l),fluvial facies (2), deltaic facies (3), lacustrine facies (41,aeolian facies (5) and mass-transport facies (6). The fourth level indicates with a lower-case letter the deposits formed by a specific mechanism within a particular facies. For example, three types of deposits can be distinguished in the (continental) terminoglacial fluvial facies, viz. terminoglacial tunnel-mouth deposits (II-A-2-c), terminoglacial stream deposits (II-A-2-d) and terminoglacial sheet- and streamflood deposits (II-A-2-e). Moreover, a terminoglacial fluvial complex (II-A-2-a) is introduced for those cases where a mixture of the just mentioned fluvial deposits exists, or where it is impossible t o determine for a specific fluvial deposit t o which type it belongs. Further subdivision

That it may be useful t o handle the sedimentary characteristics of a deposit in an equally systematic way has become apparent from the work of various authors, in particular of Miall (1977, 1978, 1983b, 1985), Rust (1978) and N. Eyles (1983b, 1985, 1987). This implies that additional codes must be used. It should be emphasised that such a n approach implies that one is leaving the sedimentary facies and entering the lithofacies. The additional codes t o be mentioned here at a 'lower-than-fourth' level therefore do not inform about the sedimentary facies as such, but may be helpful in the inventorising of lithofacies data for the various sedimentary facies. The lithofacies codes applied by the above mentioned authors are simple and easy t o work with but all show inconsistencies that make later comparisons with other lithofacies ambiguous. The present authors have therefore developed a lithofacies code scheme that is definitely based on previous work, in particular on that by N. Eyles (1985), but with adaptations that not only make the coding itself more consistent but also render it consistent with the approach followed in the coding of the sedimentary facies.

68


It seems most appropriate to place additional codes as superscripts and subscripts behind the code for the last level. According t o N. Eyles (1985) one could code: (1)the grain size, (2) the composition, (3) the sedimentary structures and (4) the bedding characteristics. Eyles also provides a code for the supposed genesis of the deposit, but such an additional code is superfluous in our classification because the genesis is already clear from the main (4-level)coding. It is much less feasible t o base a subdivision upon a systematic grouping at these sublevels than at the main four levels. The authors thus found it useful t o follow Eyles' suggestion for coding by means of 'recognisable' letters, in principle the first letter of the word that characterises the property involved. It is expected that there will usually be no need t o indicate all subcodes simultaneously and the following notation might therefore be applied: grain size with a capital superscript, composition with a lower-case superscript, sedimentary structures with a capital subscript and bedding characteristics with a lower-case superscript.

Codes for grain size - A rough distinction can be made between deposits consisting mainly of boulders, gravel, sand and 'fines' (silt and clay). There may, of course, also exist mixed deposits (in practice these are even the most common). The same distinction (and the same codes) should be used for lithified counterparts. The superscript B should be used for sediments that appear to consist mainly of boulders (Fig. 37-A). One problem is that truly coarse deposits are not suited for reliable grain-size analyses; it therefore seems acceptable from a practical point of view to apply this code in cases where material coarser than sand (over 2 mm) dominates and where boulders seem t o constitute the greater part of the coarse particles. There will commonly be a fine-grained matrix, so that most of the sediments of this category may be called 'diamicts'. The superscript G should be used for sediments that consist mainly of gravel, although scattered boulders may be present. A gravelly supraglacial ablation till would thus be indicated by the code I-A-1-bG. The superscript S is applied for sandy deposits (Fig. 37-B). Larger clasts, as well as finer particles, may be present but the sand fraction should account for at least 50% (if possible, as determined in the laboratory). It is important t o mention in this context that a deposit tends t o have a sandy appearance in the field only if the fraction of silt and clay is low (generally less than about 25%);this implies that laboratory analysis of grain-size should be used to check the field data if one is not experienced in estimating the grain size of a sediment.


69

I

I.

I, . :.

,L

Fig. 37. Various typical types of glacigenic deposits with different grain sizes. A: densely packed boulders and cobbles. B: glaciofluvial sands. C: horizontally laminated silts and clays (lacustrine bottomsets). D: typical diamict.

70


In practice, mainly silty material (a rare phenomenon, but loesses may belong to this group) is difficult t o distinguish from mainly clayey material, particularly when there is some admixture of sand. It was therefore decided, as suggested by other investigators, to group silt- and clay-sized deposits (Fig. 37-C) within one category, indicated by the superscript F (fines). Glacigenic sediments, and tills in particular, are commonly characterised by extremely bad sorting: particles ranging from clay to boulder size may be present. Such badly sorted material (Fig. 37-D) - if fines, sand and coarser particles are all present in significant quantities - should be designated by the superscript D (diamict). Diamicts may result from a direct depositional process, or from postdepositional processes. It is also possible that deposits are relatively well sorted, but with an average grain size more or less a t the boundary between two fractions, or they may be composed of material belonging to two grain-size classes. A combination of the code letters could be used in this case, e.g., superscript SG for a sandy-gravelly deposit. Codes for composition - Most sediments in the glacial and periglacial environments are siliciclastic. Other types of sediments may occur as well, however, and their presence can provide interesting information about the geological (climatological) development. It therefore seems useful to use a specific code for such sediments. As mentioned before, a lower-case superscript will be used for the purpose. Organic material may be designated by the superscript 0. Sediments with such a composition tend t o be of rather limited extent, both horizontally and vertically. They are most commonly peaty levels; such peat may be either in situ o r reworked (Petersen, 1983) in, for instance, the proglacial or extraglacial subenvironment (Fig. 38). Sediments of chemical origin are denoted by the superscript c. Such sediments are rather rare in the glacigenic area; if present, they have often been formed diagenetically, e.g., by transport in solution and subsequent precipitation of iron in the form of oxides and hydroxides. Such precipitates may form crusts, especially in the contact zone with a layer of low permeability. Diagenetically formed carbonate layers may also occur, especially if surrounding sediments contain limestone clasts or calcareous shells (Fig. 39). It does not seem justified t o attribute the code for chemical sediments t o veins that have been formed and filled inside glacigenic sediments because such veins do not form part of the sedimentary succession in a strict sense. Layers that consist mainly of concretions, however, might be denoted with a superscript c.


71

Fig. 38. Peat horizon within an aolian deposit (terrace of Kopanica river, Poland). Such organic deposits are indicated with superscript '0'.Photograph: J. Burdukiewicz.

Fig. 39. Limonite horizon (dark lower band) formed due to precipitation of iron-rich percolation water on top of an impermeable, fine-grained layer. Such chemical units are indicated with subscript 'c'. Photograph: J. Burdukiewicz.

72


Palaeosoils or comparable pedogenic levels, though not necessarily bedparallel, are important types of levels. They are most important for the reconstruction of the palaeogeographic development of an area and should therefore be indicated in stratigraphic sections. These levels are often made up of specific sedimentary layers that show characteristic colours due t o leaching and concentration of specific elements as a result of the pedogenesis. It is useful, in such a case, to give the superscript p to the layers that represent a soil horizon (Fig. 40).

3:

Fig. 40. Example of a soil horizon (to be indicated with subscript 'p') within fluvial deposits of Holocene age. Photograph: J. Burdukiewicz.

Codes for sedimentary structures - Sedimentary structures are one of the main keys for unravelling the sedimentary mechanism and the lateral and vertical changes in the depositional processes. In our opinion it is not practical to give codes for all types of sedimentary structures but the most meaningful structures do deserve such notation as a capital subscript. Current- or wind-induced cross-bedding (see, e.g., Jopling, 1965; J.R.L. Allen, 1968, 1973a,b, 1980a,b; Boersma et al., 1968; N.D. Smith, 1972; Banks, 1973b; Hunter, 1977) is a most important structure because it allows the direction of the palaeocurrent to be measured. Cross-bedding (Fig. 41),designated by the subscript C, can be found, for instance, in drift sands. Trough-shaped cross-bedding can be found in sandy dune stratifi-


73

Fig. 41. Regular cross-bedding (indicated by subscript 'C') in glaciofluvial sands. Photograph: A. Hahszczak.

cation deposited under a low flow regime. Planar cross-bedding may be found in fluvial outwash deposits of sand size and in gravelly or sandy deltaic material. Low-angle cross-bedding (less than 10") is often formed under upper flow-regime conditions. Cross-bedding resulting from scour-and-fill processes, thus indicative of alternating erosional and depositional phases, is designated by the subscript S. The same symbol can be used for the inclined lamination that can be found in channel infillings (Picard and High, 19731, as well known from supraglacial stream deposits (Fig. 42). Subscript R is attributed to ripple-drift cross-lamination, also called climbing ripples, because of the specific depositional circumstances. Such structures (Jopling and Walker, 1968; Allen, 1970c, 1971; Hunter, 1977) are commonly found in proglacial lake-margin deposits (Fig. 43) and wherever currents and settling from suspension occur simultaneously. Wave ripples (Davidson-Arnott and Greenwood, 1974; Piper et al., 1983) may be designated by subscript W. They are found in, e.g., terminoglacial lacustrine deposits (Fig. 44). Graded bedding is designated by the subscript G. This structure may occur as a result of turbidity currents (Kuenen and Migliorini, 1950), for instance from a proglacial deltaic slope to the bottomsets in front (Fig. 45).

74


I

i b.

I: Fig. 42. Inclined laminated (indicated with subscript 'S) in a channel within glaciofluvial deposits.

Fig. 43.Ripple-drift cross-lamination (subscript 'R).Photograph: A. Hahszczak.


75

Fig. 44. Irregular wave ripples (subscript 'W) in the marginal deposits of a glacial lake.

Fig. 45. Normal, i.e. upward, grading (indicated with subscript ' G ) in proglacial bottomsets.

76


Grading may also be reversed (Sallenger, 1979; Broster and Hicock, 1985). A varved succession (Kempe and Degens, 1979; Schluchter, 1979a,b; Schove, 1979; Sturm, 1979; Striimberg, 1983), commonly consisting of graded layers resulting from seasonal deposition (but aeolian varves are also known: Stokes, 1964) - alternating or not with turbidites - is denoted by the subscript V (Fig. 46). This code will be applied most commonly for varved bottomsets in glacigenic lakes.

Fig. 46. Typically varved (subscript 'V') glaciolacustrine deposits. Some dropstones are also visible.

Parallel lamination (Fig. 47) is to be designated by the subscript L. This quite common structure (McBride et al., 1975; Boyko-Diakonow, 1979; Mackiewicz, 1983; Mackiewicz et al., 1984) may have different origins, but distinguishing between them is considered beyond the scope of the present discussion. Laminated terminoglacial tunnel-mouth deposits formed under a high flow regime, subglacial channel deposits with a lamination due t o a low flow regime and proglacial lake-margin deposits that are laminated by swash and backwash thus only warrant their notation on the basis of a description of the structure and not of interpretation of their genesis. A special code is also considered useful t o indicate the presence of deformation structures within a layer. Such deformation structures


77

Fig. 47. Parallel lamination (subscript 'L'),formed during transport of sand grains under upper flow-regime conditions.

(Anketell et al., 1970; Van Loon and Wiggers, 1975, 1976; Prescott and Lisowski, 1977; Boulton and Jones, 1979; Parriaux, 1979; Doe and Dott, 1980; Funder and Petersen, 1980; Krtiger and Humlum, 1980; Schwan et al., 1980b; Boulton, 1981; Mills, 1983; Van Loon et al., 1984, 1985; Van Loon and Brodzikowski, 1987) are quite common in water-saturated sediments, especially if there is a high silt content or a relatively large amount of organic material. Various types may occur as a result of plastic deformation but liquefaction is also common. The code applied for all these structures is the subscript D (Fig. 48). Apparent absence of sedimentary structures is also worth mention. The subscript M could be applied for such massive units (Fig. 49). There may be, e.g., englacial melt-out tills that could be described by this code. A specific unit may of course be characterised by a number of different sedimentary structures. All pertinent codes might be used in such a case; the order of the codes should indicate the relative importance of the various structures. Codes for bedding characteristics - The nature of the contacts between successive layers may be useful for the interpretation of the depositional history. It is therefore considered appropriate to add a specific code (a

78


Fig. 48. Plastic deformation and liquefaction (sedimentary deformation structures are indicated with subscript 'D') in the foresets of a proglacial delta.

Fig. 49. Apparently structureless ( = massive; subscript ' M ) of unknown glaciofluvial origin.


79

lower-case subscript) in some cases in order to indicate the nature of the lower boundary of the layer. The lower boundary may be erosive, designated by subscript e, indicating that the layer involved was deposited by a process related t o an erosive force (there are two contacts a t the same place if the erosive process had nothing to do with the layer involved) (Fig. 50). The contact may also be influenced by tectonic activity (glaciotectonic push, regional endogenic forces). In this case it is useful t o indicate the non-sedimentary nature of the contact by the subscript t (Fig. 51). A rather sharp contact without any sign of a sedimentary break is denoted by subscript s. This may be the case, for instance, if a terminoglacial mass-flow deposit is laid down on top of other sediments (Fig. 52). Gradual contacts are more common, indicating that the sedimentary processes did not change abruptly. Such contacts, denoted by the subscript g, may be present in e.g. coversands where slight changes in wind intensity or direction influenced the sedimentary succession (Fig. 53). Deformed contacts due to early diagenetic processes like load casting are quite common, especially so in water-saturated sediments with

Fig. 50. Deformed sediments (centre), being a remnant of a layer that had deeply incised the clay underneath. The light-coloured layer was then eroded itself, being preserved only in the erosion depressions made before. Such erosive contacts (the contact here is partly erosive in a duplicate way) are designated with subscript’e’.

80


Fig. 51. Succession with several tectonic contacts (subscript 't') due to shearing as a result of glaciotectonism.

W

I m I :,-

. .

.

.

Fig. 52. Sharp contact (subscript 's') between a unit of silts and fine sands, and a sand layer of probably turbiditic origin.


81

i

Fig. 53. Gradual grain-size transition (subscript'g') in coversand.Photo: J. Cegia.

(temporary) high sedimentation rates and alternating grain sizes. Such deformed contacts will be denoted by subscript d (Fig. 54). Relevance of coding

An outcrop in glacigenic sediments may consist entirely of sandy material. It is superfluous t o code each layer with the superscript S in such a case. Codes should be used only where appropriate and relevant. This implies that codes should be used in cases where they are necessary (or a t least helpful) for the interpretation of the sediments or where they may serve t o distin uish between various lithological units. The code I-A-3-CgFcCefor a layer can easily be understood by readers as referring t o a layer in supraglacial deltaic foresets, consisting of sand with a relatively large amount of fines (a considerable part of the particles consisting of small concretions), with current ripples and an erosive base. It is questionable, however, whether such detailed information should always be provided, even though field work implies that the investigator does make all these observations. It does not seem practical to provide generally applicable guidelines touching the details of coding. A short description may improve readability and be equally useful. Each researcher must decide how and in how

82


Fig. 54. A diapir, representing an extreme form of deformed contacts (subscript '$1.

much detail coding should be used. The framework sketched above should therefore be considered only as a tool to facilitate communication among scientists. Use of incomplete coding

In practice, lack of data may make it impossible to establish the specific type of deposit within a particular facies. In such a case one might still use all codes that are considered correct and relevant. If one is not sure, for instance, whether a specific laminated deposit from a supraglacial deltaic facies should be interpreted as a supraglacial stream deposit or a supraglacial deltaic foreset, the pertinent deposit might be referred t o as I-AL, thus deleting the code for the specific facies and type of deposit. An erosive, massive diamict of unknown nature in the subglacial melting-ice facies might be referred t o as I-C-lMeD .

The system of glacigenic depositional environments

83

THE SYSTEM OF GLACIGENIC DEPOSITIONAL ENVIRONMENTS Continental glaciations are the final result of complex interactions between lithosphere, hydrosphere and atmosphere. Specific conditions are required for the formation of an ice cap in the contact zone between the three. Even then a continental ice cap remains a fairly unstable phenomenon, although the mere fact of the existence of such a n ice cap has its own - considerable - impact upon all three spheres. An ice cap has its own energy (mainly potential and kinetic energy) and there is a continuous energy exchange with the lithosphere and atmosphere; energy exchange of the ice with the water-phase part of the hydrosphere is largely concentrated in the frontal zone of the ice sheet. The energy exchange results in a rather complex and sensitive energy balance.

THE GLACIGENIC ENVIRONMENTS AS AN ENERGETIC ENTITY A geosystem comprises the various processes inside the Earth, at its surface and in the atmosphere that contribute to the local formation, deformation or removal of material belonging t o the lithosphere. The dynamics of these processes result in gradients for the various parameters involved and imbalances give rise to changes in intensity of all processes involved. This may result in, among others, physical changes (deformations) or chemical processes (dissolution/precipitation, etc.). An important aspect of any geosystem is its energy balance (see also Chernova, 1981).The principal inflow of energy in the glacigenic system is constituted by solar irradiation and the geothermal heat flux; the energy balance is also largely influenced by the potential energy in the system, resulting from the gravitational force. These parameters affecting the various components in a geosystem are time- and space-dependent, resulting i n a complex pattern of gradients (Fig. 55) t h a t largely determine the nature and the velocity of the changes that take place in the system. The geosystem of continental glacierisations is a typical example of a n open system. There is a n inflow of energy (e.g. solar radiation, geothermal

84


of solar energy

energy

climatic changes

changes of climate

1 sea level changes

Fig. 55. Nature and intensity of surficial (geological and geomorphological) processes as a result of various inflows of energy into the geosystem. Modified after Embleton and Thornes (1979).

heat flux) and material (debris, precipitation) into the system and an outflow of water (vapour, liquid, ice), debris and thermal energy (Fig. 56). Whether the ice advances or retreats depends on the predominance of either inflow or outflow. It should be kept in mind, however, that inflow and outflow alternate frequently during the phases of glaciation and deglaciation, although there may be considerable variations in intensity. The situation is even more complex because the glacigenic system always precipitation

rock detritus

gravity

solar irradiation

geothermal heat

Fig. 56. Input and output of energy in the glacial system. Modified after Sugden and John (1976).

The glacigenic environments as an energetic entity

85

comprises an area of mass and energy accumulation (the alimentary zone) and an area of mass and energy losses (the ablation zone). These areas are separated by a continuously shifting equilibrium line which is in fact a curved plane that corresponds more or less to the local snow line (Figs. 57, 58). Deposition of material by ice or by streams in the glacial area is one of the possible ways in which the glacigenic energy system can diminish an energy imbalance. Sedimentation thus represents a transfer of energy from the glacial system to the immediate surroundings. This implies that all areas where material from the ice-covered area is deposited belong t o the glacigenic energy system (Fig. 59); there are similar reasons for the area with permafrost t o be considered part of this system (cf. Jahn, 1950, 1970, 1975; Rbzycki, 1970, 1979; Flint, 1971; Washburn, 1973; Embleton and King, 1978). accumulation zone

@

ablation zone

T

7

I

ice thickness

I equilibrium line

upward tendence

/

ward tendence

mass and energy

\

a+ ,

f

+

energy t@

output 'wedge' of mass and energy accumulation

snow line - - ablation zone

,input

of mass and energy

-

Fig. 57. Model of the glacial mass balance. Ice flowage is required to maintain an equilibrium surface profile (A). The flow lines are indicated in cross-section (B). Modified after Sugden and John (1976).


86

accumulation zone Y n p u t of m :a

I I

marine ablation zone

and energy

I

continental ablation zone

DreciDitation lequilibrium line

eauilibrium line I

I

snow line / I

1 \

t

t

t

I

%

output of ice

output of detritus//

water; outflow

-

/ye{:

outflow of water

/

Fig. 58. Idealised cross-section through an ice cap, showing the various forms of mass and energy input and output.

glacigenic system

Y

A

marine periglacial environment

glacial environment A

continental periglacial environment k

-i

-

I

-4

Fig. 59. General distribution of glacigenic elements within the glacigenic system of environments.

Depositional sedimentary environments and their general model 87

DEPOSITIONAL SEDIMENTARY ENVIRONMENTS AND THEIR GENERAL MODEL The glacigenic depositional environments (Fig. 60) might be considered as that part of the glacigenic energy system in which sedimentation of debris is the predominant factor in the energy balance (other parts of the system as a whole are the systems where erosion, accumulation of ice, or deformation prevail) (Fig. 61). It was mentioned earlier in this book t h a t sedimentation may occur in a variety of places. A first - rough - distinction can be made between the glacial depositional environment and the periglacial depositional environment. The environments can be subdivided into subenvironments on the basis of various criteria. As described by Brodzikowski and Van Loon (19871, the spatial relationship t o the ice was chosen as the criterion for subdivision (Fig. 62).

The glacial environment The glacial environment (code I) is formed by the entire area that is covered by an uninterrupted sheet of glacial ice or a glacier. This uninterrupted ice body may be either active or passive. The external margin of the glacial environment is formed by a well defined front in both cases. The periglacial environment starts from this point on, where

GLACIAL ZONE PER IGLACI AL ZONE

Fig. 60. Overview of the transition zone between the glacial and periglacial deposi. tional environments in Greenland (photograph J. CegSa).

88


glacigenic energy system

I

glacigenic depositional system

I

W

T V

glacigenic deformational system

glacigenic erosional

sedimentary infilling of glacial basins

erosion and denudation of the substratum

glacitectonism

Fig. 61. General model of the glacigenic depositional, erosional and deformational systems as components of the glacigenic energy system.

I.

eng ac s ~ o e n ,ronment

s-oglac R s,oenv ronrnerit

ma, no g ac a s,oenvironmen[

millme term ,109 acia ice-raft (lepus IS

Fig. 62. Schematic model of the glacial and periglacial environments. Modified after Edwards (1978b) and Brodzikowski and Van Loon (1987).


the outflow of meltwater and glacial debris is well organised. An uninterrupted sheet rarely represents a passive stage although this may be the case under specific palaeogeographical conditions (Brodzikowski, 1982b, 1987; Brodzikowski and Van Loon, 1983,1985a). If the ice is in movement, or if it is passive but not separated from an active zone, a well defined front may be formed. In both cases the ice induces depositional, erosional and deformational processes that are completely different from those in other glacigenic areas. The processes within the glacial environment are greatly influenced by the distance from the ice front and by the distance to the top and bottom of the ice. The first aspect mainly accounts for differences in the deformational and erosional processes, hut also influences the physico-mechanical properties of the deposits. This means that it is the state, and not the nature of the deposits that is influenced. The second factor is responsible for the nature of deposits; lithofacies changes may occur because of different hydrologic regimes between the separate Yloors' within an ice sheet or glacier. These differences are most obvious in the various subenvironments of the glacial environment (Fig. 63). They are one of the supraglacial subenvironmen! (1-A)

I

terminoglacial subenvironment

(11-4

; I

V

subglacial subenvironment

(1-C)

Fig. 63. Schematic model of the glacial environment under continental conditions. The three subenvironments and their main facies are indicated with their hierarchic codes. Modified after Brodzikowski and Van Loon (1987).

90


most important criteria for distinguishing between the three continental glacial subenvironments (supraglacial = I-A, englacial = I-B and subglacial = I-C). The glacial environment is also found under marine conditions (I-D), viz. where a n uninterrupted 'eternal' sheet of glacial ice extends into the sea, as, for instance, around the Antarctic continent; the extent of this ice fluctuates with the seasons (Cooke and Hays, 1982). Since clastic material is supplied via the moving ice mass, glacial deposits of significant size can accumulate in the marine glacial environment, due t o settling of debris after thermosubrosion at the contact of the ice body with the sea water underneath. The area reached by glacial mass-transported material is considered t o belong the marine periglacial environment (Fig. 64).

marine periglacial environment

-

marine glacial environment

c

A

h

I

I restricted influence of marine currents

\

'1 1 1 1 1 1 1

\ '

I

I

floating ice shelf

I

A

r,

crevasses

I I 'mass1

/

I ! i

\

Y

/

tunnel mouth predominance \

I I

,

-

\

-

deposits formed due elting of the , glacial ice

I _ I1 0 UIIUelIII

uninterruph?d ice shelf

,

I

.

I

.

-; . I I continental glacial I environment

Fig. 64. Schematic model of the contact zone between the continental and the marine glacial environments.


The periglacial environment The situation in the periglacial environment (code 11) is different. This environment has been defined in many ways, but the most generally accepted definition is that the periglacial environment, as far as the continent is considered, covers all areas with a permafrosted soil (Washburn, 1951, 1979; Dylik, 1962, 1964; Pewe, 1969; French, 1976; Maarleveld, 1976; Hofle, 1983) and cryogenic processes (Troll, 1944; Dylik, 1952; Sekyra, 1960; Grigoryew, 1962,1966; Dostovalov and Kudryavtsev, 1967; Popov, 1967; etc.). This implies that the permafrosted area is often situated in front of an ice cap or glacier. This need not be the case (cf. S.A. Harris, 1982) (there are large permafrosted areas without any direct relationship with an ice cover, for instance in Canada: N. Eyles, 1977),but the term Iperiglacial'is used in the present book exclusively in the sense of 'belonging to a perrnafosted area i n front of a n ice cover'. Buried, discontinuous ice masses and isolated dead-ice blocks may be present in this environment, particularly during a recession of the ice. Buried deadice blocks may be found far away from the ice front if the retreat is sufficiently fast and if the climatological conditions are favourable. The processes in the periglacial environment are strongly dependent on the distance from the ice front, a factor which is responsible for the main differences in depositional, erosional and deformational phenomena which characterise the various subenvironments. The periglacial environment comprises both continental and marine subenvironments (a permafrost may be present under marine conditions: Lachenbruch, 1957; MacKay, 1972), all of which are characterised by specific associations of deposits. The continental periglacial environment can be subdivided into a terminoglacial (11-A),a proglacial (II-B) and an extraglacial (II-C) subenvironment. The marine equivalents of these continental subenvironments are coded II-D, II-E and II-F, respectively (Fig. 65 on page 92).

92


per ig lacial environ men t (I I)

-

h-

A

extraglacial aeolian facies

ice marginal streamway

(11-GI)

1

proglacial Ian facies (11-6 1)

I I

outwash

terminoglacial lacustrine facies (11-A 1)

proglacial fluvial deposits (11-B-1-c)

-

J

~

Y-

(11-A)

(11-B)

(11-C) -

i

terminoglacial subenvironment -,

proglacial subenvironment

extraglacial subenvironment

-

-

,r

A

proglacial lacustrine IaCleS

A

I i

terminoglacial terrestrial facies

proglacial deltaic facies

proglacial fan facies

terrninoglacial fluvial facies (11-A 2)

zone of ice-marqinal streamways

marine periglacial environment (11) ~

undercurrents

0

deposits

glacioma

ice raft deposits (11-D-1 -e)

marine terminoglacial subenvironment (11-D)

marine terrninoglacial mass transport deposits (I1 A 6 c)

marine glacial environment (I-D)

Fig. 65. Schematic model of the periglacial environments. A: under continental conditions with well developed terminoglacial lakes, outwash plains and pradolinas. B: under continental conditions with well developed terminoglacial fluvial and masstransport facies, and proglacial fluvial, deltaic and lacustrine facies. C: under marine conditions with grounding ice, a floating ice shelf and ice rafts. Modified after Brodzikowski and Van Loon (1987).

Glacigenic facies and their characteristics

93

GLACIGENIC FACIES AND THEIR CHARACTERISTICS The glacigenic subenvironments, distinguished on the basis of their spatial relationship t o the ice sheet or glacier, can be subdivided into facies that the proposed system of hierarchic coding indicates by a n Arabic numeral. The basis for this subdivision is the relationship between the natural conditions within a subenvironment; the same basis for subdivision is applied by most other authors but often with a different nomenclature. Each subenvironment is subdivided into facies on the basis of the most characteristic depositional conditions; in turn, types of deposits may be discerned within the various facies (see also the Tables on pages 128-129). The two environments, their subenvironments, facies and deposits have no random distribution. Natural changes in climatic conditions result in logical successions in time and space. This means that a phase of glacierisation will commonly be expressed in the geological record by a sedimentary sequence; the glacial sequence was indeed one of the first to be recognised and understood as such. A typical continental glacial sequence (Fig. 66) starts with extraglacial sediments (e.g., loess), followed successively by proglacial sediments (e.g., meltwater deposits), terminoglacial sediments (e.g., tunnel-mouth deposits), subglacial melt-out sediments (till) that have often an erosive base; because of the conditions of active ice, this lower part of the sequence will commonly not be followed by en- and supraglacial deposits. When the ice recedes, supra- and englacial deposits may be formed, followed by subglacial sediments. All these glacial sediments (s.s.) will again be covered by subsequently terminoglacial-, proglacial and extraglacial sediments (Fig. 67). A complication that arises during ice recession is the presence of dead-ice bodies that may greatly influence the depositional pattern. The sedimentary sequence just described constitutes a n expression of the changing depositional conditions in time, and thus forms the basis for the general model concerning glacigenic deposition. It must, however, be emphasised that 'ideal' sequences are extremely rare (if present at all) and that they cannot serve as the only basis for a model: the present-day distributions of facies, the prevailing conditions and the resulting types of deposits must also be studied in order to understand how the preserved


94

loesses hills

extraglacial proglacial

----

supraglacial

:Fr

_ _ _ _ _

- lakes etc.

- --

stream deposits dead - ice

flow tills

- ~ ~ _ _ _ englacial subglacial ~

terminoglacial

~

coversands varvites

tunnel deposits

active ice

diamicts incorporation tills

active ice _

_

_

.

flow tills tunnel-mouth deposits

fan

I

ice retreat

J7

sheetflood deDosits

braided-river deposits outwash plain channel-bar deposits proglacial/

lake

channel fills varvites

pradolina

channel-bar deposits

extraglacial

i_extraglacial

}adzce

delta

_____

coversands loesses substratum

Fig. 66-A. Idealised glacial sequence, formed as a result of gradual facies shifts during successive periods of ice advance and retreat.

Fig. 66-B. Relationship between spatial facies distribution, ice advance and resulting sedimentary succession.

95


+i

1

...........

retreat

1

7--

A

time

extraglacial

6

proglacial

C

terminoglacial

G

supraglacial

E

subglacial (passive ice)

C

terminoglacial

borderline of the glacigenic depositional system

relative extent of deposits

*

advance

/J

A

extraglacial

I

1 maximum extent

j of the ice front I

Fig. 67. Schematic representation of the most common succession and relative extent of glacigenic deposits. Note that the englacial and supraglacial deposits are not commonly formed during ice advance.

material fits within the framework of the originally much more extensive sedimentary succession. A problem is, of course, that while the periglacial environment is relatively well accessible - so that the geologically relevant processes can be studied - the glacial environment is rather hostile and furthermore processes like subglacial lodging, plucking due to water-pressure variations (Rothlisberger and Iken, 1981), erosion and sedimentation can in fact not be studied directly a t all, thus leaving some 'white areas' in our knowledge of glacial processes. Consequently, models concerning subglacial and englacial sedimentation must be based almost entirely on theoretical considerations and on the interpretation of field data. It can therefore be expected that when knowledge increases models of glacial sedimentation will need more adaptation than models for periglacial sedimentation. The problem is not a new one and it has generated much discussion and design of models (e.g., see Sugden and John, 1976; Embleton and King, 1977; Ehlers, 1983; N. Eyles, 1983). In spite of all the shortcomings of our present-day knowledge, it seems justified to distinguish between environments, subenvironments, facies and deposits, as mentioned before (see the Tables on pages 128-129).

96


The general environmental conditions determine the nature of the facies within each glacigenic subenvironment. It appears that six types of facies occur frequently in the glacial and the periglacial environment (cf. Flint, 1971; Jahn, 1975; Sugden and John, 1976; Embleton and King, 1977; Edwards, 1978; N. Eyles, 1983b; Gradzinski et al., 1986) and sometimes in a number of subenvironments. These types are: the melting-ice facies (called 'melt-out facies' by some authors, among others Boulton, 1970b; Ruszczyiiska-Szenajch and Lindner, 1976; Drozdowski, 1982,1983; Haldorsen and Shaw, 1982; Shaw, 1982, 1983), the fluvial facies (under these conditions commonly called 'glaciofluvial facies'; the bottom-current facies under marine conditions is considered here as a n equivalent with the same number (2) as code), the glaciodeltaic facies, the glaciolacustrine facies, the aeolian facies and the mass-transport facies. It is quite remarkable that these facies have been recognised as being of primary importance from the very beginning of the study of glacigenic sediments (cf. Woldstedt, 1954; Charlesworth, 1957; Flint, 1971; Klimaszewski, 1976). THE MELTING-ICE FACIES Melting-ice facies occur under continental conditions where debris, previously embedded in a n ice cap or glacier, is set free by ablation or thermosubrosion and accumulates usually more or less in situ (Fig. 68). Most of the deposits thus formed consists of mineral particles varying in size from clay fraction to boulders; organic material supplied by the ice may also be present but this is a very rare phenomenon because the source area tends to contain no organic clasts (some organic material may be eroded, however, from the substratum when the ice mass passes). The accumulated material is commonly called 'till' ('tillite' if lithified), although 'glacial diamict' ('diamictite' if lithified) would be more consistent with the generally applied rules of terminology (cf. Eyles, 1983; Eyles et al., 1983). The depositional process implies that these facies are most common in the supraglacial (facies I-A-1) and terminoglacial (facies II-A-1) subenvironments, but englacial (1-B-1) and subglacial (1-C-1) melting-ice facies occur as well. Ice melts during phases of both advance and retreat of a glacier or ice cap. This implies that sediments formed in a melting-ice facies are t o be found all over the area that has been glaciated. In fact, the sediments even reach somewhat further, because mass movements of material set free in the glacial environment by ablation, may reach the fluvioglacial fans and lakes of the terminoglacial subenvironment in front of the main ice body.

The melting-ice facies

97

Fig. 68. Melting-ice facies at de Glacier des Bossons near Chamonix (Switzerland). Note the concentration of debris set free by ablation.

Melting-ice facies also occur in the marine counterparts of the continental terminoglacial, proglacial and extraglacial subenvironments (facies 11-D-1,II-E-1and 11-F-1,respectively). These facies (Fig. 69) differ from the continental equivalents in that undermelting is generally more important in the process of melting than solar irradiation. All melting-ice facies are characterised by the fact that erosion - if present at all - is small, whereas sedimentation is a more or less continuous process. This does not imply, however, that the melt-out processes are uniform: differences occur due t o variations in prevailing conditions, mainly determined by the distance to the top and the base of the ice body and by the meteorological or climatological conditions. Subglacial melting, for instance, may be a result of the thermal heat flux or may represent pressure-melting. Englacial melting is mainly a result of thermosubrosion (primarily by meltwater streams that were warmed supraglacially and subsequently penetrated the ice mass via crevasses). Supraglacial melting is commonly the direct consequence of solar irradiation, but thermoerosion by supraglacial meltwater streams and vaporisation also contribute to the ablation. A special type of this

98


process is the gradual melting of ice blocks floating in a glacial or periglacial lake o r in a sea; this process is commonly called 'undermelting'. Although tills are the most characteristic glacial sediments, their genesis is, in general, less well understood than that of other glacigenic deposits. This dearth of knowledge is due mostly to the inaccessibility of the depositional sites. It is obvious, however, that the characteristics of tills from the various melting-ice facies are quite different, but it should be emphasised that tills from one specific facies may also show a large variation in characteristics. Among the reasons for the differences are the ablation rate, the drainage pattern, the size of the clasts, the prevailing pressure and the original iceklasts ratio.

Fig. 69. Melting-ice facies under marine conditions.

The melting-ice facies

99

The various types of melting processes result in more or less different conditions for the accumulation of the glacial debris set free. Field work, however, does not generally reveal very distinct features t o reconstruct the precise melting process from the sedimentary characteristics. This is probably due t o reworking of the material by meltwater streams after the ice has retreated or by collapse of sediment if some buried dead-ice melts away. The deposits formed by melting processes must therefore be interpreted on the basis of other characteristics (overconsolidation of subglacial tills, place in a sequence, lateral facies transitions, etc.). The deposits from melting-ice facies have now been studied in detail for more than a century (Hutton, 1795; Lyell, l830,1840a,b; A. Geikie, 1863; Jamieson, 1865; Goodchild, 1874; J. Geikie, 1877, 1894; Torell, 1877; Miller, 1884; Upham, 1891; Chamberlin, 1894; Crosby, 1896; Garwood and Gregory, 1898). Research was intensified in the 1930s and tills still receive much attention, with a more sedimentological approach having been followed since the 1960s. Recent work by Canadian researchers (e.g., Eyles, 1983; Eyles et al., 1983) has used this approach very sucessfully. Melt-out sediments may be classified according to their presumed place of genesis within a subenvironment (the approach followed in this book), or on the basis of sedimentary characteristics, presumed precise process of deposition, degree of reworking or morphological expression (see also Dreimanis, 1988). These various approaches have resulted i n a great many names; the most important ones are referred to in the following chapters. Five types of deposits are distinguished withing the melting-ice facies. They are: melt-out complexes (denoted by code 'a'), ablation tills (b), lodgement tills (c), basal tills (d) and ice-raft deposits (e). The melt-out complexes (see, a.0. Shaw, 1979, 1982; Boulton, 1980a,b; Dreimanis, 1988) are not due t o a specific depositional process, but they consist of a mixture of other types of melt-out deposits that cannot be classified otherwise, either because they lack sufficiently clear characteristics, or because post-depositional processes (e.g., glaciotectonism) have mixed a number of different melt-out deposits. Such mixtures are common in the supraglacial, englacial, subglacial and terminoglacial (continental) subenvironments (Fig. 70). Ablation tills form subaerially, more o r less i n situ. They occur therefore only in the supraglacial and terminoglacial (continental) subenvironments. The primary process that results in their formation is the melting (or sublimation) of ice due to a temperature rise resulting from contact with air masses or from penetration into the ice of solar heat (cf. Drozdowski, 1977; Shaw, 1977a,b; Klatkowa, 1982).

100


t supraglacial till complex

0

I m

2m

Fig. 70. Supraglacial till complex a t the margin of the Breidamerjokull (SE Iceland), with various types of deposits, including ablation till (l), flow till (2) and glaciofluvial sand (3).Modified after Boulton and Eyles (1979).

Lodgement tills are restricted to the subglacial subenvironment. They are deposited during the movement of a n active-ice body in the contact zone between the ice and the substratum. The deposition is a result of successive frictional retardatation and pressure-induced melt-out of individual bedrock particles and/or debris aggregates against the substratum (Boulton, 1972a, 1975a; Mickelson, 1973; N. Eyles and Menzies, 1983). Lodgement tills show a wide variety of characteristics; such different types are commonly indicated by different names, but nomenclature is very inconsistent. Basal tills are subglacial diamicts formed in situ. They represent the lowermost till in a complete sequence and they consist of debris that was set free by melting of the ice as a result of - predominantly - t h e geothermal heat flux. Basal tills are more commonly formed during a stage of passive ice than during active ice movement (Olszewski, 1974; Drozdowski, 1979). Ice-raft deposits are subaqueous deposits formed as a result of melting of a n ice mass covering - or floating in - a water body, thus releasing the embedded debris. Such deposits are well known from supraglacial and terminoglacial lakes (it cannot be excluded that they are also formed enand subglacially, but such deposits have no preservational potential or will not be recognised), and from the marine glacial environment and the

The glaciofluvial facies

101

marine terminoglacial subenvironment. The occurrence of ice-raft deposits outside the marine periglacial environment is a result of the long distance that icebergs may travel before they have been melted away completely (they have been observed in subtropical seas). THE GLACIOFLUVIAL FACIES These facies, characterised by meltwater streams, extends over truly large areas. They are found in both the glacial and the periglacial environments. The supraglacial fluvial facies (I-A-2)is most conspicuous in the glacial environment, but the englacial (I-B-2) and subglacial (14-2) meltwatertunnels may also be considered as a particular type of glaciofluvial facies. The terminoglacial (11-A-2)and proglacial (11-B-2)fluvial facies are most characteristic in the periglacial subenvironment (Fig. 7 1). No glaciofluvial facies exists in the extraglacial subenvironment because both the water and the debris carried along may be largely of non-glacigenic origin. Meltwater streams are found under conditions of both active and passive ice. Obviously, the fluvial influence decreases with diminishing ablation rate, but it is unlikely t h a t there should be periods totally without glaciofluvial sedimentation. Glaciofluvial conditions commonly resemble 'normal' fluvial conditions and the depositional pattern of meltwater streams is much the same as under non-glacigenic conditions. The currents erode, transport and deposit debris and form straight, braided or meandering streams, depending on the conditions of slope, amount of water and amount of debris (cf. Colby, 1963; Flint, 1971; Klimaszewski, 1976; Sugden and John, 1976; Edwards, 1978; Gradzinski et al., 1986). It seems justified to refer the interested reader to more specialist works on fluvial sedimentation, if detailed information on the sedimentological aspects is required; some useful handbooks are those by Reading (1978a), Reineck and Singh (1980), Collinson and Lewin (1983), Miall (19841, Walker (1984) and Gradzinski et al. (1986). More specific d a t a on sedimentation in braided streams is t o be found i n works by, among others, Chien (1961), Doeglas (1962), Ore (1963), Boothroyd (1970), N.D. Smith (1970, 1971), Costello and Walker (1972), Banks (1973a), Cant (1975, 1976, 1978a,b), Gilbertand Asquith (19761, Miall (19761, Gilbing and Rust (1977),Hein and Walker (19771,Osterkamp (1978), Blodgett and Stanley (1980), Vos and Tankard (1981), Bluck (1982), Turner (1983,

102


Fig. 71. The proglacial fluvial facies in Svalbard. Top: overview of the facies in front of the Werenskjold glacier. Photograph:J. Cegla. Bottom: idem, detail of some braided channels. Photograph:J. Bierohski.


103

1984) and Rust (1984). Details on sedimentation in meandering rivers are given by Leopold and Wolman (1957), Bernard and Major (1963), Langbein and Leopold (1966), Bluck (1971), Shelton and Noble (1974),Jackson (1975, 1978), Gustavson (1978), Ori (1979, 1982), Nanson (1980) and Stewart (1981). Englacial and subglacial meltwater streams are interesting exceptions because they may be truly different from subaerial streams: the tunnels may be completely filled with water and where the water may flow, sometimes even in a n upward direction, as a result of pressure exerted by water penetrating the ice mass from above. Relatively little is known about the sedimentary processes in englacial and subglacial tunnels; the most detailed sedimentological reconstructions are possibly those by Banerjee and McDonald (1975), Rust and Romanelli (1975) and Saunderson (l975,1977b), but interesting material has also been provided by Durand (1951, 1953), Guy e t al. (19661, Acaroglu and Graf (1968), Babcock (1970), Elliott and Gliddon (1970), K.C. Wilson (1970), Wilson and Brebner (1971) and McDonald and Vincent (1972). The deposits interpreted as having been formed under such conditions show relatively poor sorting, but are few in number and have undergone deformations due t o ice movement, collapse o r other processes. Eskers are interpreted by some authors as being deposits formed i n subglacial channels, but field data and theoretical considerations make i t more likely that the eskers were formed not in the tunnels themselves, but in the tunnel mouth; according t o this view the considerable length of eskers is due to the gradual retreat of the ice (and of the tunnel mouth), thus leaving behind a real trace of the previous tunnel-mouth position (and therefore probably also of the original subglacial channel). Subglacial tunnels are almost inaccessible in present-day ice caps, so that next to nothing is learned from actual field inventories. Experiments carried out to solve complex technical problems (e.g., flowage through closed pipelines) have not yet contributed much to what is known about sedimentation in glacial tunnels (cf. Saunderson, 1977), but experiments into this problem are still in progress. Much more is known about the subaerial glaciofluvial sedimentation. In general, the supply of both water and debris is rather irregular under glacigenic conditions and braided streams therefore seem most common, especially near the ice front. In more distal zones, however, many such streams may have combined to form straight channels or large meandering rivers. Overbank deposits are nevertheless rare as compared to the situation in more moderate climates.

104


The precise type of sedimentation depends on the local morphology, while the morphology, in turn, is commonly due t o the glaciofluvial processes (Fig. 72). Glaciofluvial outwash plains generally grow laterally (and slightly vertically) by more or less continuous deposition in small channels, whereas real fluvial sequences may be formed in deeply incised channels between the dead-ice blocks that occur in the terminoglacial subenvironment. The variation in conditions gives rise to a great diversity in glaciofluvial subfacies and types of deposits. The conditions in the proglacial subenvironment, particularly near the transition t o the extraglacial subenvironment, resemble those in non-glacigenic areas. The relief and the amount of meltwater and of debris largely determine the size and shape of the glaciofluvial channels, the thickness and extent of the deposits and the dynamics of the depositional process. The depositional pattern is also influenced, particularly with respect to the vertical accretion and the average grain size, by the energy gradient in the region (slope) and the distance from the ice front.

Fig. 72. The supraglacial fluvial facies on top of a glacier at Svalbard. The course of the streams is largely determined by depressions resulting from collapse of ice above cavities. These cavities are mainly due to thermosubrosion in subglacial drainage systems that are, in turn, partly determined by the supply (via crevasses) of surficial meltwater. Photograph: J. Cegta.


105

For reasons of practice, only the following general types of deposits are distinguished here: fluvial complexes (denoted by code 'a'), tunnel deposits (b),tunnel-mouth deposits (c), stream deposits (d),and sheet- and streamflood deposits (e). Fluvial complexes consist of obviously fluvial deposits that cannot be further determined, or of bodies that apparently consist of a number of different types of fluvial deposits. They are found in the (continental) supraglacial, terminoglacial and proglacial subenvironments. Tunnel deposits are formed only under englacial and subglacial conditions; the difference with subaerial (or subglacial) stream deposits is the increased pressure exerted by the weight of water in crevasses at higher locations, in direct contact with the tunnels. Tunnel-mouth deposits are formed where englacial or subglacial streams leave the ice body. This situation is characterised by high energy gradients resulting in rapid vertical accumulation. These deposits form the transition from stream or tunnel deposits t o deltaic, lacustrine or marine deposits. Tunnel-mouth deposits occur in the continental supraglacial and terminoglacial subenvironments and in the marine terminoglacial subenvironment. Stream deposits are the most 'classical' type of fluvial deposits. They are formed in straight, braided and meandering rivers, and are condidered here to include overbank deposits (Schumm and Lichty, 1963; Klimek, 1974; Steel, 1974; Stear, 1978; Nanson, 1980; R.M.H. Smith, 1980), fans, etc. They occur in the supraglacial, terminoglacial and proglacial subenvironments (and possibly also in the englacial and subglacial subenvironments, but without any significant preservational potential); marine equivalents (marine bottom-current deposits; these are often indicated in literature by different names, e.g., 'marine channel deposits': C.H. Eyles, 1986, 1987) are found in the marine terminoglacial and proglacial subenvironments. Sheet- and streamflood deposits, which have been studied particularly in non-glacigenic environments (McGee, 1897; Chawner, 1935; McKee et al., 1967; Rahn, 1967; G.E. Williams, 1971; Karcz, 1972; Bryhni, 1978; Tunbridge, 1981, 1983; Hogg, 1982), occur when the normal drainage channels have insufficient capacity to transport a sudden amount of water (e.g., after heavy rain, snow melting or lake break-through). Such deposits occur in the (continental) supraglacial, terminoglacial and proglacial subenvironments. Some investigators assume the occurrence of sheet and stream floods in the subglacial subenvironment as well (Shaw and Kvill, 1984; Shaw, 198713; Shaw et al., 1989), and also under lacustrine conditions, where they would form 'subaqueous high floods'.

106


THE MARINE BOTTOM-CURRENT FACIES Bottom currents under marine conditions have many aspects in common with fluvial streams, especially regarding the depositional processes. The marine bottom-current deposits are therefore denoted in the present book, for reasons of practice, with the same code as the fluvial facies (arab number 2); the codes of the various types of deposits are also consistent with their continental equivalents. The main sedimentological difference between marine and fresh-water currents is that the salt content of sea water affects the specific weight of the water and thus the effective specific weight of the particles that are transported. Many data have been presented in the last decades by Normark and Piper (19691, Ness and Kulm (19731, Walker (1975, 1978, 19841, Eriksson (1982), Andrews and Matsch (1983); Gilbert (1983), Miall (1983b, 19851, Vorren et al. (19831, Wright et al. (19831, Clifton (19841, Hein (19841, McCabe et al. (1984, 19871, C.H. Eyles (1985,1986,1987) and N. Eyles et al. (1985) and others. The marine bottom currents considered in this book are restricted to those with a glacigenic character. This implies that the majority of such currents originate from subaqueous englacial o r subglacial tunnel mouths. Where the meltwater streams embouch in the sea, the current velocity drops suddenly, so that all coarser particles are deposited in the direct vicinity of the tunnel mouth. The result is a subaqueous fan that is in some respects comparable to a subaerial outwash plain. The main difference is that subaqueous fans have commonly a much steeper surface. This inclined surface induces faster currents that, in turn, may incise the fan deeper than is common under subaerial conditions. The channels become gradually filled with deposits that were supplied partly grain-bygrain, partly in the form of mass transport (Eriksson, 1982; C.H. Eyles, 1987). Large supplies of debris-laden meltwater currents facilitate, just as does a complete filling of channels, more areal sedimentation so that thin, flat covers are formed. These blankets may later be incised again. Other interesting data about marine bottom-current activity have been presented by Dowdeswell (1987) who stated that variable current velocities may be responsible for bedforms such as sandwaves and megaripples in water with a depth of less than 110 m. Most effective seem bottomcurrent velocities of 40-70 cm s-'. Deposits from marine glacigenic bottom-current facies have been reported from the terminoglacial and the extraglacial subenvironments. The glacigenic character disappears with increasing distance from the source, so that such deposits need not necessarily be distinguished in the extraglacial subenvironment.

The glaciodeltaic facies

107

THE GLACIODELTAIC FACIES The continuous supply of meltwater, combined with impermeable subsoils (ice, permafrosted soil, loam) and an irregular topography, leads to the formation of lakes in most glacigenic subenvironments. Meltwater streams embouching in such lakes tend to form deltas (Fig. 73) because the current velocity decreases and the debris carried along by the streams comes to rest. As such, the glaciodeltaic facies forms a gradual transition between the glaciofluvial and the glaciolacustrine facies. The glacigenic deltas are most commonly found in the supraglacial subenvironment (facies I-A-3) and the proglacial subenvironment (II-B-3),but may also be found a t the margins of the predominantly lacustrine facies of the subglacial subenvironment (facies 1 4 - 3 ) and the terminoglacial subenvironment (facies 11-A-3);subglacial deltas may become embedded in the ice mass (Harris and Bothamley, 1984),but it seems not justified to distinguish them as a type of englacial deltas. Glacigenic deltas are also found in the marine terminoglacial (11-D-3) and the marine extraglacial (11-F-3)deltaic subenvironments.

Fig. 73. Small proglacial delta in Greenland, some 10 km from the ice front. Photograph: J. Ceg4a.

108


It is common for glaciodeltaic deltas under continental conditions to develop simultaneously at several places at the border of a lake, because meltwater streams may approach the lake from all sides. This implies that more debris is usually transported per unit time to a glacigenic lake than to non-glacigenic lowland lakes, and glacigenic deltas therefore may soon grow together t o form an irregular lake margin. The water is usually (almost) stagnant even when the lake has an outlet, which means that the deltas develop without being eroded by currents. Wave action may be present, but if the lakes are small, the waves are not strong enough t o affect the shape of the delta. Consequently, well developed deltas may be found with classical sequences and a classical grain-size distribution in a lateral direction. The precise shape of the delta depends on local conditions, but there are to be found classical subaqueous fan deltas, supraaqueous deltas, Gilbert-type deltas and Salisbury-type deltas, the latter being a transitional form between the Gilbert-type delta and the supraaqueous Hjulstrom-type delta (see Gilbert, 1885; Salisbury, 1892; Hjulstrom, 1952; Bates, 1953; Fisher et al., 1969; Aario, 1972; Broussard, 1975; Church and Gilbert, 1975; Galloway, 1976; Clemmensen and Houmark-Nielsen, 1981; Schwab and Lee, 1983, 1988; Schwab et al., 1987). The gradual transition from glaciofluvial deposits t o glaciodeltaic topsets, as well as the gradual transition from the glaciodeltaic bottomsets to glaciolacustrine bottomsets, emphasises the transitional character of the glaciodeltaic facies. Most of the above descriptions apply equally well to terminoglacial and extraglacial marine deltas if 'sea' is read instead of 'lake'. It should be kept in mind, however, that the conditions in polar seas are usually much rougher than in glacial and periglacial lakes. This implies that the deltas commonly develop in a less classical way, that the channels in the topsets may be influenced by tides, etc. Sedimentation in glacigenic deltas is essentially the same as under other conditions (Fig. 74). The currents slow down when they reach the lake, and increasingly finer material is deposited. Locally ongoing accumulation may force the current t o shift its course, thus finally creating the characteristic delta shape. Debris-rich currents are usually relatively heavy, and thus follow the slope of the depositional basin. If, however, a meltwater stream is relatively warm (e.g., because of insolation), it may be lighter than the water in the lake and a plume of 'river water' may extend along the surface of the lake where it gradually takes on the temperature of the surrounding water and becomes mixed with the lake water, due t o

The glaciodeltaic facies

109

turbulences in the water. Such plumes are formed most easily if the meltwater streams do not contain many heavy particles, for instance because they are already far from the source area or because the debris was derived from an eroded clayey sediment. The 'ideal' pattern of vertical and lateral grain-size zones can be disturbed by various processes. Mass flows over the foresets occur quite commonly; such flows often contain relatively coarse material from the topsets that is redeposited in the bottomsets. There may be so many comparatively coarse-grained mass-flow deposits side by side - transported via more or less parallel channels in the foresets - that a belt is created that might be compared with a piedmont at the foot of a mountain range. Other disturbances may result from wind-induced currents or waves, resulting in ridges or comparable structures, or in a relative enrichment of sand in zones affected by wave action. Moreover, fine-grained deposits may form in the coarse belt of the delta's topset within interdistributary bays. Another, different factor can affect the zonal buildup of the deltas. This is the (dis)continuity of parameters such as the supply of water and debris and the water level in the lake, which may fluctuate because of changes in the height of the barriers surrounding the lake (sometimes ice masses) or because of tectonic activity. Such tectonic movements in the substratum may distinctly influence the channel pattern in the deltaic topsets, thus affecting the lithofacies characteristics. If the disturbances dominate the sedimentation pattern in a delta, it can become almost impossible to distinguish between topsets, foresets and bottomsets. In such a case it might be advisable to distinguish only the deltaic complex as an entity.

7

m\

channel

,,:I

stream depbs8Is

/

Sheet andstreamflood deposits

lake margin deposits

dropstone

lake margin deposits

Fig. 74. Schematic section of glaciodeltaic and associated deposits, as interpreted from exposures in Lower Silesia (SWPoland).

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Deltaic sedimentation has received much attention in the past. Some general data have been provided by, among others, Van Straaten (1960), Moore (1966), Oomkens (1967), McGowen and Scott (1974), Wright et al. (1974),Coleman and Wright (1975), Sutton and Ramsayer (1975),McCabe and Jones (19771, Prior et al. (1981) and Coleman et al. (1983). Regional studies with interesting observations were published by, among others, Moore (1959), Coleman and Cagliano (19641, Gole and Chitale (19661, Agterberg et al. (1967), Bouma and Bryant (1969), McGowen (1970), Van de Graaff (19721, Van Loon (1972),Flores (1975), Roberts et al. (1976), Vos (1977,1981), Cherven (1978), Stanley and Surdam (1978), Ricci Lucchi et al. (1981), Kostaschuk and Smith (1983) and Porqbski (1984). Specific characteristics of glacigenic deltas were described by, among others, Gustavson et al. (1975),Galloway (1976), Cohen (19791, Jorgensen (1982), Leckie and McCann (1982) and Torreson and Schwab (1987). All these studies indicate that deltaic sedimentation is extremely complex; i t comprises typically fluvial and lacustrine sediments, but mass movements are also common. However, most authors agree t h a t it is sedimentologically most appropriate t o distinguish either the deltaic complex (a) as a n entity, or t o distinguish between deltaic topsets (b), foresets (c) and bottomsets (d). This subdivision is also followed in the present book. The term 'deltaic complex' is used here wherever a deltaic deposit is too small t o distinguish between the top-, fore- and bottomsets, or where the transition between clearly fluvial (including tunnel-mouth) deposits and clearly lacustrine or marine deposits is gradual or vague. Such complexes are found in the continental supraglacial, terminoglacial and proglacial subenvironments, and in the marine terminoglacial and proglacial subenvironments. If topsets, foresets and bottomsets are distinguished, they occur, of course, in the same subenvironments. Glaciodeltaic topsets represent the most elevated parts of the classical Gilbert-type delta, with distributary channels, levees, bars, etc. The deposits from these subfacies form usually a complex pattern. Glaciodeltaic foresets are constituted of the inclined units that prograde as sedimentation continues. The foresets are incised by channels, and mass movements take place frequently. Deltaic foresets are absent in Hjulstom-type supraaquatic deltas (which are sometimes called 'fans'). Glaciodeltaic bottomsets are formed by the fine-grained material that could be transported far enough. The bottomsets alternate with massmovement deposits; they pass gradually into the 'normal' lacustrine or marine deposits. The bottomsets are generally very complex as regards their internal structure.

111

The glaciolacustrine facies

THE GLACIOLACUSTRINE FACIES Lakes occur throughout the glacigenic environments; as already mentioned their frequent occurrence is due to the impermeable soil present in many places, the irregular topography, and the abundance of meltwater streams (Fig. 75). Most glacigenic lakes are t o be found in the periglacial environment, where ice masses and irregular topography have a strong influence on the flowage pattern (Fig. 76), namely in the terminoglacial subenvironment (facies 11-A-4)and the proglacial subenvironment (facies 11-B-4). By definition, glacigenic lakes do not occur in the extraglacial subenvironment. Lakes are somewhat less abundant i n t h e glacial environment but do occur in the supraglacial subenvironment (facies I-A4) and the subglacial subenvironment (facies 1 4 - 4 ) (see, a.0. E.H. Walker, 1967; R.G. Walker, 1967; May, 1977; Ashley, 1975; Church and Gilbert, 1975; Gustavson, 1975; Shaw, 1975a, 1977b; Merta, 1978; Shaw and Archer, 1978, 1979; Shaw et al., 1978; Waitt, 1980; Gilbert and Shaw, 1981; C.H. Eyles and N. Eyles, 1983a; Pickrill and Irwin, 1982, 1983; Quigley, 1983; N. Eyles and Miall, 1984; Weirich, 1982; N. Eyles, 1987; N. Eyles et al., 1987a; Brodzikowski, in press).

terminoglacial lake fed directly by the ice cap supraglacial lake moraine lakes fed by meltwater streams

I

I

crevasse with subqlacial lakes

i

Fig. 75. Schematic overview of various types of glacigenic lakes.

small enalacial

112


Fig. 76. Supraglacial lake on the Werenskiold glacier (SW Svalbard). The lake was formed because meltwater could not flow away from a depression in between ice-cored ablation tills. Photograph: J. Czerwifiski.

The glaciolacustrine deposits formed in the proglacial subenvironment have a reasonable chance of survival (although they may become deformed or partly eroded); the preservational potential becomes increasingly less for the deposits formed in the terminoglacial, supraglacial and subglacial subenvironments, respectively. Glaciolacustrine deposits with a large extent and formed within a deep (preferably deepening) lake have good chances of survival (C.H. Eyles and N. Eyles, 1983a; N. Eyles and Clark, 1986; N. Eyles et al., 1987a, 1988a); some glacial lakes with a length of over one hundred kilometres are known t o have survived thousands of years. The deposits formed in such lakes may be traceable in outcrops over large areas, even when the deposits have been deformed by glaciotectonics or other processes. On the other hand, the much more frequently formed sediments in small lakes or even in pools have only a small chance of survival; should they survive, their precise place in the lithostratigraphic record is often difficult t o determine because such


113

deposits cannot be traced from one outcrop to another. However, they can play an important role in detailed palaeogeographical reconstructions. The particles for these deposits are supplied by large meltwater streams (sometimes via tunnel mouths in either dead-ice blocks or the ablation zone of the ice cap) and by small streams that run off the surrounding sedimentary surface. This results in three types of 'real' lacustrine deposits: those where a tunnel mouth enters the lake (more or less comparable to a river entering a lake), marginal lacustrine deposits and lacustrine bottomsets. The far greatest part of the debris accumulating in glacial lakes is supplied by meltwater streams (Gradziiiski et al., 1986). Minor amounts are supplied by melting of debris-loaden, floating ice masses and by winds that are strong enough to transport dust (or even fine sand). An even less important role is played by erosion of the borders of the lake by wave or in large lakes - current action. Several sedimentary processes are responsible for almost all glaciolacustrine deposits that are preserved in geological records (Fig. 77). The main process is the deposition of bed-load material, mainly along the lake margin, from currents that slow down gradually. Such deposits are characterised by current ripples. The second process is the settling of fine particles from suspension, particularly in the centre of the lake. Such

Fig. 77. The sediment of glacigenic lakes is supplied by supraglacial, englacial and subglacial meltwater streams and by melting ice rafts. The depositional pattern depends on the homopycnal, hyperpycnal or hypopycnal character of the inflowing, debris-containing water.

114


particles may have been supplied by meltwater streams or by the wind. The deposits formed by settling of fine particles often show graded bedding (varves), which is considered to be - at least partly - a type of seasonal fluctuation in the sediment supply: wind and surficial streams continue t o supply particles if the lake is covered by ice during the winter. Melting of the ice in springtime can thus result in a sudden supply of clastics, and the finest particles then released can only settle during the next winter when the water becomes completely quiet as a result of a new ice cover. This explanation of varve formation is, however, not really satisfactory since the presumed depositional process is not known from actual glaciolacustrine conditions: the formation of true varves has at least not been observed directly. Various authors have therefore discussed the origin of the grading (e.g., Kuenen, 1951; Woldstedt, 1954; Merta, 1978) but no convincing alternative has yet been proposed for the often regular grading. Turbidity currents are certainly responsible for some of the graded layers (Schwan et al., 1980, even described a 'double-source turbidite') but it is doubtful whether this can be considered t o be the main mechanism (Shaw et al., 1978). Anyway, redeposition of material may take place along the lake margins. This may occur in the form of slumps or mudflows, but turbidity currents are considered t o play an important role. The resulting turbidites may constitute a considerable part of the graded beds in a glacial lake (Shaw, 1977b; Shaw and Archer, 1978,1979; Shaw et al., 1978). All other depositional processes are of minor importance, although they may contribute considerably in specific cases. An example of such a process is the melting of floating ice masses, resulting in debris falling down (the deposits thus formed are termed here 'ice-raft deposits' and are grouped in the melting-ice facies). The larger clasts thus forming part of much finer grained glaciolacustrine sediments are commonly called 'dropstones', a name that is also used in the literature t o indicate the deposit in which such clasts occur. The result of the combined sedimentary processes is that a relatively coarse-grained marginal belt, with a gradually inward-fining sediment more in the centre, is commonly found. A lateral shift in the facies may therefore lead to vertical changes in the average grain size in glaciolacustrine successions. The spatial distribution of the fine-grained and coarse-grained deposits may be disturbed by various processes, e.g., longshore currents in large lakes, catastrophic events like sudden outbursts of meltwater through a barrier surrounding the lake, tectonic subsidence of the substratum, etc. Moreover, the average grain size of the material supplied may change due t o retreat or advance of the ice front, increased


115

slope instability leading to mass movements, changes in the atmospheric conditions, etc. Subglacial lacustrine conditions may differ greatly from those sketched above, but little is known about them because of the inaccessibility of the facies. Moreover, it is likely that deposits from subglacial lakes are not commonly recognised because they resemble the surrounding diamicts, particularly after deformation by the active ice. Some authors suggest a subglacial lacustrine origin for laminated diamicts that are supposed to result from thermosubrosion of the roof of the empty space. Lacustrine sedimentation has been investigated i n detail, partly because of economic reasons. Some general information has been presented by Gilbert (1885), Yuretich (1982) and Hakanson and Jansson (1983). Case studies with interesting data have been described by, among others, Ludlam (1974,1979,1981),Clemmensen (19781, Clemmey (1978), Hesse and Reading (1978), Heward (1978), Link and Osborne (19781, Muller and Wagner (19781, Sturm and Matter (19781, Ashley (1979), Horie (1979), Gilbert and Church (1983) and Spallatti (1983). Typically glaciolacustrine sediments were investigated and discussed by Hardy and Legget (1960), Nichols (19601, Ferrians (1963), Howarth (1968), Harrison (1975), McDonalds and Shilts (1975), Eschman (19791, Schluchter and Knecht (19791, Schwan et al. (1980), Smith and Syvitski (19821, Smith et al. (1982), Campy (1983b), C.H. Eyles and N. Eyles (1983a, 1984b), N. Eyles et al. (1983b),Dreimanis (1984a),Gravenor (1984),Karrow (1984b1, Sharpe (19841, N. Eyles and Clague (1987) and Shaw (1988). It is generally agreed upon that differentiation of lake deposits is only useful if the lake has a diameter of a t least some kilometres (very large lakes like those of Wisconsinan age in North America may even show a a n equally wide variety of characteristics as found under marine conditions). The deposits of relatively small lakes are - which is consistent with most literature - considered in the present book a s a n entity, termed 'lacustrine complex' (denoted with code 'a'). However, there is no general agreement as to the types of deposits to be distinguished in large lakes. The most simple subdivision is followed here, distinguishing only between lake-margin deposits (b) and lacustrine bottomsets (c). The lacustrine complexes comprise the sediments of lakes that are too small to have distinctly differentiated central and marginal deposits, or commonly distorted - mixtures of both types. Such complexes are commonly found in the supraglacial, terminoglacial and proglacial subenvironments, and more rarely also in the subglacial subenvironment. There is no reason t o believe that englacial lacustrine complexes are not formed, but unambiguous descriptions or interpretations are lacking, because of the

116


inaccessibility of the englacial subenvironment and because the preservational potential of such deposits is practically zero. The lake-margin deposits comprise all lacustrine deposits t h a t a r e formed in the relatively shallow marginal zone, where other processes than settling in (almost) stagnant water prevail. The only exception is formed by the deposits that are formed by the direct influence of the agent that supplies the bulk of the clastic particles and of the water (deltaic deposits, tunnel-mouth deposits). Lake-margin deposits have been described frequently from the supraglacial, terminoglacial and proglacial subenvironments. Lacustrine deposits in the subglacial subenvironment can as a rule not be differentiated into lake-margin deposits and lacustrine bottomsets. The lacustrine bottomsets are those formed more or less in the centre of a lake, particularly by settling from suspension. Seasonal coverage of the water surface by ice may favour seasonal sedimentation cycles: t h e grading upward varves. Graded units within lacustrine bottomsets may, however, also be due to turbidity currents originating in the marginal zones and reaching the lake's centre. Lacustrine bottomsets are found i n the same subenvironments as lake-margin deposits. THE AEOLIAN FACIES Aeolian sedimentation takes place over the entire glacigenic area, as far a s i t is exposed subaerially. A relatively dry climate, t h e lack of vegetation and a windy climate favour erosion, transport and deposition of particles by wind (Fig. 78). All subaerial glacigenic subenvironments undergo these processes. This results in the existence of more or less well developed aeolian facies in the supraglacial (I-A-5),terminoglacial (II-A5), proglacial (11-B-5) and extraglacial (11-C-5) subenvironments (windblown particles may also settle in the sea but, there, they mix with other material and cannot be distinguished as separate deposits). Glaciofluvial sedimentation, however, generally predominates over deposition of wind-blown material. This implies that a truly aeolian facies can develop only where meltwater streams have little or no influence, or where inter-channel areas are dry enough t o be affected by wind action. The extraglacial subenvironment thus represents the best conditions for aeolisation, but there are many reports on distinct aeolian deposits from the supraglacial, terminoglacial and proglacial subenvironments as well (e.g., Pew6,1955; R. Gilbert, 1983; Pewe and Journaux, 1983). Wind can easily erode the soil in the glacigenic area because the climate is relatively dry (dry soils are more easily eroded than wet soils)

The aeolian facies

-3

L

- L

117

IIIIII

6

111111

Fig. 78. Erosion, transport and deposition of aeolian dust. A = cool and dry air of an advancing cold front. B = air of a dry zone, warmed by insolation. C = humid air of the upper atmosphere. D = humid, temperate zone. E = annual precipitation. a, a1 and az: ascending air currents that carry dust into the air. b zone of aeolian segregation of sand and dust. c: zone of rolling and saltation of sand. d: zone of falling dust loaded with atmospheric moisture. dl: zone of maximum loess accumulation. e: zone of uniform, gradually decreasing loess accumulation. 1 = direction of cold front. 2' = direction of inflow of upper humid air masses. 3 = wind of intermediate altitude. 4 = upper winds, carrying dust. 5 = dust carried up by ascending air massers. 6 = falling of dust loaded with atmospheric moisture. 7 = surficial winds. 8 = accumulated loess. 9 = low winds, reworking accumulated loesses. 10 = dunes. 11 = direction of sand rolling. 1 2 =upper humid air masses. From: R6iycki (1979). Courtesy: Polish Academy of Sciences, Branch Office in Poznari; Committee of Quaternary Research.

and because vegetation is scarce or even absent. The source areas of the large coversand and loess belts that were formed during the Pleistocene have been much discussed. Gradients in grain-size distribution, sedimentary structures and heavy-mineral analyses were studied to approach the problem. Sufficient data are now available in most cases t o determine the source area with certainty, but more detailed knowledge is still lacking about which phenomena in the Pleistocene aeolian deposits should be ascribed t o the original supply of the wind-blown material, and which phenomena are due to later processes that reshaped and redistributed the original aeolian sediments.

118


The main processes contributing to aeolian sedimentation are gravityinduced falling from the air of wind-blown particles, accumulation in the form of ripples as a result of traction processes, saltation of grains and avalanching (resulting in the leeward aggradation of dunes). These processes have been investigated and described in detail, both from glacigenic and other areas (e.g., Bagnold, 1941; Chepil, 1945; Butler, 1950; Mason and Folk, 1958; G.P. Williams, 1964; T.R. Walker, 1967; Glennie and Evamy, 1968; Smalley and Vita-Finzi, 1968; Stokes, 1968; Cegla, 1969; Glennie, 1970; McKee et al., 1971; I.G. Wilson, l971,1972a,b, 1973; Bigarella, 1972; Cooke and Warren, 1973; Brookfield, 1977; Hunter, 1977, 1980, 1981; R.G. Walker and Middleton, 1977, 1979; Ahlbrandt and Andrews, 1978; Ahlbrandt et al., 1978; Reading, 1978a; Sarnthein, 1978; Smalley and Krinsley, 1978; Ahlbrandt, 1979; Fryberger and Ahlbrandt, 1979; Fryberger and Dean, 1979; McKee, 1979a,b; Reineck and Singh, 1980; Ahlbrandt and Fryberger, 1981,1982; Fryberger and Schenk, 1981; Kocurek and Dott, 1981; Kocurek and Fielder, 1982; Kocurek, 1981a; Horowitz, 1982; Rubin and Hunter, 1982, 1983; Mucher, 1986; Schwan, 1987). Experiments contributing t o the knowledge on aeolian processes were carried out by Kuenen (1960), Bowen and Lindley (1977), De Ploey (1977) and Whalley et al. (1982) among others. Regional studies proving good insight in aeolian depositional processes, the resulting deposits and their sedimentary structures were carried out by, among others, McKee and Tibbits (1964), Stokes (1964), Kozarski et al. (1969), Folk (1971),Glennie (1972,1983),McKee and Bigarella (19721, Steidtmann (19741, Yaalon and Dan (1974), Ahlbrandt (1975), Coetzee (1975), McKee and Moiola (1975), Ahlbrandt and Andrews (19781, Brookfield (19791, Fryberger et al. (19791, McKee (1979c, 1982), Kocurek (1981b),Koster (1982),Mader (1982) and Blakey and Middleton (1983). Wind-blown material settling from the air may fall into the sea, but is most usually not voluminous enough to be traced as an aeolian contribution to *normal*marine sedimentation (even though it is supposed that the recent red deepsea oozes owe their colour t o wind-blown, ferruginous desert sand); the aeolian facies is thus not present as such in the marine periglacial environment. Wind-blown material may, however, also settle in lakes, where it forms part of the *normal'lacustrine sediment. Most aeolian fine-grained material is found as loess covers in terrestrial areas (Fig. 79), especially where high-pressure and catabatic atmospheric conditions prevail. Such conditions are quite common in the extraglacial subenvironment, which means that these aeolian sheets may cover most of the extraglacial area, interrupted only by fluvial sediments in valleys and by material reworked by slope processes.

The aeolian facies

119

Field data indicate that sedimentation of aeolian covers during the Pleistocene was not a more or less continuous process but that, instead, relatively short periods of intensive aeolian deposition alternated with periods of almost no deposition or even of net aeolian erosion. The cyclicity is obvious, for example, from the palaeosoils that could develop in these deposits, but there seem t o be discrepancies between the cycles within loess and those within coversands and dunes. There are also local differences in cyclicity. It can therefore be deduced that not only the largescale climatic conditions played a role, but that small-scale meteorological conditions had an impact. This topic needs much more research, although there has been considerable progress in the last few years. Apart from the aeolian complexes (denoted by code 'a'), four main types of aeolian deposits can be distinguished: drift sands (irregular masses of accumulated, wind-blown material; denoted by code lower-case 'b'), dunes (denoted with 'c'), coversands (denoted with Id') and loesses ('e'). Drift sands (Dewers, 1934) are thin sheets, small accumulations or other units of limited size and commonly restricted preservational potential. They are well known from present-day aeolian facies, e.g., from

Fig. 79. Characteristic, undulating loess area, south of Dunidin (NewZealand). Photograph: J.D. de Jong.

120


Svalbard, Greenland, Baffin Island (Aksu and Piper, 1979), Alaska and Iceland. Such drift sands may be found on top of both aeolian and nonaeolian deposits. Dunes are much larger sand masses (sometimes they consist of clay: Bowler, 1973; Dare-Edwards, 1982) and show distinct relief. They move under the influence of the wind over the substratum (if it is dry enough), mainly by avalanching of saltating grains at the leeward side. These dunes are commonly found in pradolinas and in the distal parts of outwash plains. Much about the genesis of glacigenic dunes results from experiments (e.g., Logie, 1981) or from investigations of non-glacigenic dunes (Cornish, 1879; Friedman, 1961; Shepard and Young, 1961; McKee, 1966; Hand, 1967; Roberts et al., 1973; Bigarella, 1975; Goldsmith, 1978; Howard et al., 1978; Hesp, 1981; Lancaster, 1982). Coversands are widespread sheets of aeolian sand that tend t o level the original relief (Maarleveld, 1960). The sand is mainly supplied in nearsuriace traction, but part of the - finer - particles may have been blown high in the air. The formation of coversands requires dry conditions, such as they occur in polar and subpolar areas. Relatively small coversands are known from Greenland and from sanders on Baffin Island; the most extensive deposits are known, however, from the last glaciation; they form more or less continuous belts in the - then - proglacial subenvironment (Ducker and Maarleveld, 1957). Loesses occur also in the proglacial subenvironment but they are more characteristic for the extraglacial subenvironment (Dewers, 1932); they consist for a large part of silt-sized particles, which allows transport over large distances (Smalley, 1966, 1975, 1980; Goudie, 1978). Loesses form commonly homogeneous covers over wide areas, in front of the coversands (which are deposited in a much narrower zone). Recent loesses have been described from Svalbard and Alaska, but the large, fertile loess belts in the United States, Europe and Asia date from the Pleistocene glaciations. THE MASS-TRANSPORT FACIES The term 'mass-transport facies' as used here applies to the facies characterised by reworking of previously deposited material on subaerial or subaqueous slopes. Such conditions may be present in the proglacial and extraglacial (continental) subenvironments but the reworked material then usually forms a minor part of other facies. Several parameters facilitate the mass transport of sediments. The most important of these parameters are the existence of slopes, water-

The mass-transport facies

121

saturated material, and deposits with a high content of silt (or clay). These conditions frequently prevail in glacigenic areas and, consequently, masstransported material is quite commonly found. It is therefore considered appropriate t o distinguish a mass-transport facies in the context of glacigenic sedimentation. Mass transport may take place in numerous ways. As it is not a typically glacigenic process it seems justified t o deal here with mass transport in a schematic way only. Three main types of mass-transported glacigenic deposits are therefore considered: those formed subaerially (denoted by lower-case 'a'), those formed by a mixture of subaerial and subaqueous processes (mainly in crevasses; denoted by 'b'), and those formed under subaqueous conditions (denoted by 'c'). The subaerial mass-transport processes represented in glacigenic nonconsolidated rocks (Fig. 80) include rock fall, creep (De Ploey and Moeyersons, 1975), solifluction (under glacigenic conditions often called

Fig. 80. Subaerially reworked supraglacial debris. The ongoing melting of the ice cores induces instable slopes along which the debris rolls, slumps, slides and creeps. Werenskiold glacier (SW Svalbard).Photograph:J. Czerwifiski.

122


'gelifluction'), slumping and sliding and flowage (e.g., Reid, 1969; Addison, 1981). The resulting, reworked sediments may still be found on the slopes where they originated, but most accumulation takes place in depressions at the foot of such slopes, e.g., in lakes or valleys. The intensity of the reworking is determined by the grain-size distribution of the material, the degree of consolidation, the amount of water in the pore spaces, the inclination of the slope and the meteorological conditions (Jahn, 1970, 1975; Washburn, 1973; Embleton and King, 1977). These factors also affect the final extent of the individual deposits and that of the facies as a whole. By definition, crevasses occur only in the supraglacial (facies I-A-61, englacial (facies I-B-6) and terminoglacial (facies II-A-6) subenvironments. There is a clear difference between the supraglacial crevasses (exposed directly t o the atmosphere) and englacial crevasses (representing the more or less vertically positioned spaces within the ice; the more horizontal ones are called 'tunnels'). The crevasses, under favourable conditions, can be filled with clastic material supplied by water flowing over the glacier's surface or through the often complex system of englacial tunnels and crevasses; the crevasses can also be filled by wind-blown material or by material that reaches the crevasses during subaerial mass transport. Crevasses a r e thus commonly filled, a t least partly, by sediments that form a mixture of subaerially and subaqueously masstransported material. When the ice has melted away, the deposits thus originated may form - often complicated and disturbed - sedimentary bodies (Klatkowa, 1972; Sugden and John, 1976; Sharp, 1985a,b). A s mentioned, crevasses occur not only a t the ice surface (Fig. 81) but also within the ice. One might argue that the term 'crevasses' does not apply if there is a prevailing horizontal component; in our opinion the term 'crevasse' should apply to all open spaces embedded in the ice with a n overall inclination of over 45". After melting of the ice it is generally impossible to distinguish between crevasse and tunnel infillings; both result in irregular lenses or layers of sands and gravels, embedded within englacial melt-out tills. Englacial crevasses may occur throughout the ice, but their deposits are often found on top of subglacial tills, suggesting that crevasses occur, mainly during advanced deglaciation, in the lower part of the ice also. Analysis of the microstructures occurring in englacial crevasse deposits and their relations to englacial tills allows an environmental reconstruction of the crevasses. I t is found t h a t the crevasses can be formed in the upper part of an ice sheet, where they are formed under conditions of static pressure below the plastic limit of the ice. The active plastic strain of the flowing ice increases in relation to depth. This results


I

/

I

-

123

m

till ridges

traces of crevasses

-~ _- foliation Fig. 81. Interrelationships between the orientations of foliation, crevasses and till ridges at the western margin of Eybakkajokull (Iceland). Modified from Sharp (198513).

in small, only partially opened crevasses in the lower part of the active-ice body; the shape and position of these crevasses may change during their existence, since intracrystal dislocations (sliding mechanism of flowing ice) give rise t o specific confining pressure and dynamics. The changeability is greater than that in more surficial crevasses, since the ice in the top part is almost passive. The hydrodynamic regime of crevasses has been discussed by various authors (e.g., Nye, 1965; Dewart, 1966; Stenborg, 1968; Shreve, 1972). An idealised, completely static model of an ice sheet or glacier section (see Fig. 14.11 of Sugden and John, 1976) can be compared with the hydrogeological situation in a karst area (Shreve, 1972). The mechanism involved in each, however, has not yet been analysed. This may be due to the complicated aspects of the permeability factors in an ice body at the melting point; there is a 3-dimensional network of fissures around the single ice grains, as shown by Nye and Frank (1973). The analysis by Sugden and John (1976) unfortunately does not deal with the confining pressure which increases with depth, or with the changes in plasticity.

124


A special type of crevasses are the so-called frontal crevasses (Fig. 821, reported from the terminoglacial part of ice masses. They are narrow, commonly deep, vertical spaces between the front of an active ice mass and the foreland. These spaces become filled by sliding or otherwise masstransported debris. Similar crevasses exist around nunataks. Supraglacial crevasses may easily be filled with sediments, while englacial crevasses may be closed again before sedimentation ends or even starts. An increasing effect is caused by the melting of the ice: this process will take place more quickly in supraglacial than in englacial conditions; the amount of debris that becomes available is more or less proportional t o the volume of melted ice. Supraglacial streams, carrying along larger or smaller amounts of detritus, may encounter a crevasse on their way. The water will then fall down and form a pool at the bottom of the crevasse where the sediments can settle until the crevasse is completely filled. The sediments in such crevasses will consist of particles of various sizes; the largest particles may have tumbled down, clasts of intermediate size may have been transported as bed load in a current, and the finest particles may have settled out of suspension in the water-filled pool. Moreover, variations in

Fig. 82. Frontal crevasses in the Gefrorne Wand glacier (SW Austria).


125

the steepness of the crevasse walls facilitate subaqueous reworking of previously deposited material. The resulting sediment is therefore a mixture with a wide variety of sedimentary characteristics. If the crevasse does not form a closed system from the ice surface downwards but is connected with the englacial system of tunnels and crevasses the meltwater stream will keep flowing, which can imply an increase in the relative amount of sand-sized particles that are deposited (finer particles will be transported along). The sedimentary processes in the englacial subenvironment are not essentially different, but it should be emphasised that most englacial meltwater streams have a supraglacial origin. This means that most of the coarse clasts will not reach the englacial crevasses, so that settling from a slackening current is a less important process than settling from suspension where more or less stagnant water occurs. Mass transport occurs even more easily on subaqueous than on subaerial slopes. An inclination of less than one degree may be sufficient t o initiate mass movement and even less inclined surfaces will favour the continuation of a mass-transport process. Subaqueous mass transport occurs under glacigenic conditions in many ways. Plastic deformation (e.g., in the form of slumps) is quite common and may occur in lakes and under marine conditions. Slumps may, if the conditions are favourable and if the length of the slope is large enough, pass into mudflows and subsequently in turbidity currents. Such turbidity currents (Banerjee, 1966) are not only responsible for the deposition of graded beds in glacigenic lakes (where they may resemble season-induced varves), but also for the occurrence of glacigenic material in the marine extraglacial subenvironment. If the liquid limit of the material has not been passed during the subaqueous mass transport, the resulting deposits are sometimes called 'flow till' (Hartshorn, 1958; Marcussen, 1973, 1975; Paul, 19731, also under glaciomarine conditions (Powell, 198313).The term is also applied by some authors t o similar deposits that have been transported subaerially. The processes responsible for subaqueous mass transport have been studied extensively during the last few decades, and they are relatively well understood now, although some aspects, particularly touching mass transport of diamicts, need more study (Dreimanis, 1988). Papers dealing with glacigenic subaqueous mass transport and their deposits were presented by, among others, Wright and Anderson (1982), Broster and Hicock (1985), N. Eyles (1987), N. Eyles et al. (1987a, 1988a) and Schwab et al. (1987). Some general data on the formation of such deposits have been provided by Heezen and Ewing (1952), Kuenen (19521, Shepard

126


(1954), Middleton (1966a,b,c, 19671, Johnson (1970), Van Loon (1970), Skipper (1971), Ricci-Lucchi (1975), Carlson and Molnia (1977), Normark (1978), Skipper and Bhattacharjee (1978), Nardin et al. (1979), R.G. Walker (1979), Lee et al. (1981), Prior et al. (1982), Shanmugan and Moiola (1982), Bugge (19831, Weaver and Kuijpers (1983), Mutti et al. (1984) and many others.

Tables of glacigenic subenvironments, facies and deposits

127

TABLE OF THE CONTINENTAL GLACIAL ENVIRONMENT, WITH SUBENVIRONMENTS, FACIES AND DEPOSITS supraglacial subenvironment (1-A)

melting-ice facies (I-A-1) fluvial facies (I-A-2)

deltaic facies (I-A-3)

lacustrine facies (I-A-4)

melt-out complexes (I-A-1-a) ablation tills (I-A-1-b) ice-raft deposits (I-A-1-e) fluvial complexes (I-A-2-a) tunnel-mouth deposits (I-A-2-c) stream deposits (I-A-2-d) sheet- and streamflood deposits (I-A-2-e) deltaic complexes (I-A-3-a) deltaic topsets (I-A-3-b) deltaic foresets (I-A-3-c) deltaic bottomsets (I-A-3-d) lacustrine complexes (I-A-4-4 lake-margin deposits (I-A-4-b) lacustrine bottomsets (I-A-4-c)

aeolian f. (I-A-5) drift sands (I-A-5-b) mass-transport facies (1-A-6)

subaerial mass-transport deposits (I-A-6-a) crevasse deposits (I-A-6-b) subaqueous mass-transport deposits (I-A-6-c)

englacial

melting-ice facies melt-out complexes (I-B-1-a)

su benvironment

fluvial f. (I-B-2)

(1-B)

mass-transport f. crevasse deposits (I-B-6-b)

subglacial subenvironment (I-C)

melting-ice facies (I-C-1)

till complexes (I-C-1-a) lodgement tills (I-C-1-c) basal tills (I-C-1-d) ice-raft deposits (I-C-1-e)

fluvial f. (I-C-2)

meltwater-tunnel deposits (I-C-2-b)

deltaic f. (I-C-3)

deltaic complexes (I-C-3-a)

meltwater-tunnel deposits (I-B-2-b)

lacustrine f (I-C-4)lacustrine complexes (I-C-4-a) mass-transport f. mass-transport deposits (I-C-6-c)

128


TABLE OF THE CONTINENTAL PERIGLACIAL ENVIRONMENT, WITH SUBENVIRONMENTS, FACIES AND DEPOSITS terminoglacial subenvironment

(II-A)

meltin -ice facies 81-A-1)

till complexes (II-A-1-a) ice-raft deposits (II-A-1-e) fluvial complexes (II-A-2-a) fluvial tunnel-mouth deposits (II-A-2-c) facies (11-A-2) stream deposits (II-A-2-d) sheet- and streamflood deposits (II-A-2-e) deltaic complexes (II-A-3-a) deltaic deltaic topsets (II-A-3-b) facies (11-A-3) deltaic foresets (II-A-3-c) deltaic bottomsets (II-A-3-d) lacustrine complexes (II-A-4-a) lacustrine lake-margin deposits (II-A-4-b) facies (II-A-4) lacustrine bottomsets (II-A-4-c) aeolian f. (11-A-5) drift sands (II-A-5-b) subaerial mass-transport deposits mass-trans ort (II-A-6-a) facies(I1-A-%) crevasse deposits (II-A-6-b) subaqueous mass-transport deposits (II-A-6-4

proglacial

melting-ice facies till complexes (II-B-1-a)

subenvironment

fluvial facies (11-B-2)

(II-B)

deltaic facies (II-B-3) lacustrine facies (II-B-4) aeolian facies (11-B-5)

mass-transport facies (II-B-6)

extraglacial subenvironment (II-C)

aeolian facies (11-C-5) mass-transport facies (11-C-6)

fluvial complexes (II-B-2-a) stream deposits (II-B-2-d) sheet- and streamflood deposits (II-B-2-e) deltaic complexes (II-B-3-a) deltaic topsets (II-B-3-b) deltaic foresets (II-B-3-c) deltaic bottomsets (II-B-3-d) lacustrine complexes (II-B-4-a) lake-margin deposits (II-B-4-b) lacustrine bottomsets (II-B-4-c) aeolian complexes (II-B-5-a) drift sands (II-B-5-b) dunes (II-B-5-c) coversands (II-B-5-d) loesses (II-B-5-e) subaerial mass-transport deposits (II-B-6-a) subaqueous mass-transport deposits (II-B-6-c) drift sands (II-C-5-b) dunes (II-C-5-c) coversands (II-C-5-d) loesses (II-C-5-e) subaerial mass-transport deposits (II-C-6-a)


TABLE OF THE MARINE GLACIAL ENVIRONMENT, WITH FACIES AND DEPOSITS marine glacial environment (1-D)

melting-ice facies (I-D-1)

melt-out complexes (I-D-1-a)

bottom-current facies (I-D-2)

fluvial complexes (II-A-2-a) marine glaciai tunnel-mouth deposits (I-D-2-c)

mass-transport facies (I-D-6)

marine glacial mass-transport deposits (I-D-6-c)

129

130


TABLE OF THE MARINE PERIGLACIAL ENVIRONMENT, WITH SUBENVIRONMENTS, FACIES AND DEPOSITS

marine terminoglacial subenvironment (11-D)

marine proglacial subenvironment

(11-E) marine extraglacial subenvironment (11-F)

melting-ice facies (11-D-1)

marine terminoglacial ice-raft deposits (II-A-1-e)

bottom-current facies (11-D-2)

marine terminoglacial tunnel-mouth deposits (II-D-2-4 marine terminoglacial bottomcurrent deposits (II-D-2-D)

deltaic facies (11-D-3)

marine terminoglacial deltaic complexes (II-D-3-a) marine t. deltaic topsets (II-D-3-b) marine t. deltaic foresets (II-D-3-4 marine terminoglacial deltaic bottomsets (II-D-3-8

mass-transport facies (11-D-6)

marine terminoglacial mass-transport deposits (II-A-6-a)

no glacigenic facies

no glacigenic deposits of sedimentological significance

bottom-current facies (TI-F-2)

marine terminoglacial bottomcurrent deposits (11-D-2-D)

deltaic facies (11-F-3)

mass-transport facies (11-F-6)

marine terminoglacial deltaic complexes (II-D-3-a) marine t. deltaic topsets (II-D-3-b) marine t. deltaic foresets (II-D-3-4 marine terminoglacial deltaic bottomsets (II-D-3-d) marine terminoglacial mass-transport deposits (II-A-6-a)

The supraglacial subenvironment

131

THE CONTINENTAL SUPRAGLACIAL SUBENVIRONMENT (I-A)AND ITS DEPOSITS The supraglacial subenvironment (see the Table on p. 127) - Sharp (1949) and Dreimanis (1989) use the term 'superglacial' - comprises the ablational area from the ice surface downwards, as far as it is still subaerially exposed (crevasses), influenced by atmospheric processes or influenced by processes taking place within the sediment cover. Sedimentation is fully determined by local ablation, relief and - if present - vertical tectonic movements of the substratum. The former is responsible for the rate of melt-out processes, their character and the amount of meltwater; the latter determines the exact position of deposition, the flow pattern of the meltwaters and their erosive and transport capacity (Fig. 83). The

position of accumulation

* relief

pattern of meltwater flows

a

3

erosional and transport capacity

'

4

deformation 01 substratum

tectonic act~vityof Substratum

Fig. 83. Main agents that influence the rate and nature ofsupraglacial sedimentation.

132


important characteristic of this subenvironment (melting of ice) is reflected in the type of deposits that are predominantly deposited in water. Various facies can be discerned within this subenvironment (see also Boulton, 197213;Shaw, 1972a; N. Eyles, 1979; Boulton and Deynoux, 1981; Rains and Shaw, 1981; Brodzikowski and Van Loon, 1983, 1987; Paul, 1983): the supraglacial melting-ice facies (1-A-l),the supraglacial fluvial facies (I-A-2), the supraglacial deltaic facies (I-A-3), the supraglacial lacustrine facies (I-A-4), the supraglacial aeolian facies (I-A-5) and the supraglacial mass-transport facies (I-A-6). It should be emphasised that this subenvironment can be studied under present-day conditions, with the result that the information is much more accurate than that available for the englacial and subglacial counterparts. This relatively good accessibility holds for studies on both active and passive ice (Boulton, 1967, 1968, 197213; N. Eyles and Slatt, 1977; Shaw, 1977c;Boulton and Eyles, 1979; N. Eyles, 1979,1983a). SUPRAGLACIAL CONDITIONS ON ACTIVE ICE Studies of the sedimentation on top of active ice (e.g., Boulton, 1972a,b) indicate that the subenvironment is characterised by momentary deposition. Debris may be present, particularly in the accumulation area, in the form of supraglacial moraines consisting mainly of particles fallen down from frost-weathered rocks (Fig. 84). It can also be set free by surficial ablation or be supplied via meltwater streams through crevasses towards the glacier's surface. Most of this debris is commonly transported by supraglacial streams t o the ice front, where it can come t o rest in a terminoglacial or proglacial deposit. Debris may, however, also be deposited in local depressions on top of the ice, where real landscapes may develop if the conditions are favourable (Kozarski and Szupryczynski, 1973; Drozdowski, 1977; N. Eyles, 1979; Brodzikowski, 1984). The preservational potential of such supraglacial deposits is rather small because ice movement carries the sediments on top of the ice away with varying speeds, and because the sediments may again become fully or partly incorporated into the ice mass. Supraglacial sediments formed on top of a rather flat ice sheet that melts slowly on a substratum with a favourable topography are the most likely to survive. An example of such favourable conditions is a very slight inclination towards the ice front (Fig. 851, because this hampers the formation of rapidly flowing meltwater streams carrying the debris away from the glacier. Pleistocene supraglacial deposits formed under such conditions are frequently found in lowland areas (Klatkowa, 1972, 1982; Drozdowski, 1974; Karczewski,

Supraglacial conditions on active ice

133

Fig. 84. Glacier in the Bernina Massif (Italian Alps). Note the presence of supraglacial debris (right foreground).Photograph:J.F.Th. Schoute. supraglacial

supraglac~aldrln

deltaic complex

Fig. 85. The supraglacial subenvironment under active-ice conditions, with a substratum gently inclined towards the centre of the ice mass.

134


1974; Kozarski, 1981; Brodzikowski and Van Loon, 19831, but are also present in mountainous regions (Boulton and Eyles, 1979; Van der Meer, 1982; N. Eyles, 1983d; N. Eyles et al., 1987a). SUPRAGLACIAL CONDITIONS ON PASSIVE ICE When precipitation in the alimentary zone of a n ice sheet diminishes, the frontal ice zone may face a situation without a supply of fresh ice. The ice stops moving and remains in a passive condition. Ablation will continue in the frontal zone, resulting in gradual retreat of the ice mass. This deglaciation process generally takes place somewhat irregularly, leaving blocks of passive ice isolated from the main ice sheet. Such isolated blocks are termed 'dead-ice bodies' or 'dead-ice blocks'. The depositional conditions on top of passive ice are quite different from those on top of active ice. Ablation results in a continuous setting free of debris, and meltwater streams may run in directions quite other from supraglacla subaerlal mass-tran~poit

supraglacial deltaic complex

supraglacial

Fig. 86. The supraglacial subenvironment under passive-ice conditions during a stage of advanced deglaciation (modified after Brodzikowski and Van Loon, 1987).

Supraglacial conditions on passive ice

135

those on top of the active ice (Fig. 86).These streams may carry away most of the debris, but passive ice commonly soon becomes entirely covered by a layer of debris; this layer protects the underlying ice from direct solar irradiation (Fig. 87), thus reducing the ablation rate. On the other hand, some relatively deep supraglacial channels may develop in which surficial waters may flow until they penetrate the ice through crevasses or small joints (Fig. 88). The englacial meltwater streams that arise in this way cause the interior of bodies of passive ice to melt a t an increased rate as a result of thermosubrosion (Baranowski, 1977; Brodzikowski, 1987). Debris-covered passive ice may be reached by meltwater streams coming from the active zone. This can result in gradually thickening supraglacial sediments. If ablation continues and the ice retreats, debriscovered ice masses are left in front of the ice sheets. Such buried dead-ice blocks then form part of the terminoglacial subenvironment, where they may survive for a long time. A subsequent readvance of the ice may either override such dead ice masses, push them away or incorporate them into the main ice body (Brodzikowski and Van Loon, 1987).

Fig. 87. Glacier table near Hornsund (Svalbard), formed because a large block protected the underlying ice from irradiation-induced ablation. Photograph: J.D. de Jong.

136


Fig. 88. Large crevasse, acting as a pathway for supraglacial meltwater to the englacial subenvironment.

GEOLOGICAL PROCESSES IN THE SUPRAGLACIAL SUBENVIRONMENT The supraglacial subenvironment is the most complex of all glacial subenvironments because almost all possible geological processes may be involved. Deposition from meltwater is by far the most important, as shown by the supraglacial facies patterns. Deposition takes place partly on top of the ice (for another part in crevasses), thus making these supraglacial sediments easily accessible to erosional processes. Their preservational potential is therefore rather small. If these sediments survive, they will be found as a more or less continuous cover or as isolated patches (due to erosion) on top of the 'transgressive' part of the glacial sequence. The position in the sedimentary sequence is not the only characteristic: the deposits are - by definition - formed on top of ice that melts away during deglaciation. Consequently, the ice 'substratum' of the supraglacial sediments is unstable and many deformational structures (collapse structures, normal faults, slump-like structures) may be found. Irregular

Geological processes in the supraglacial subenvironment

137

melting of the ice also accounts for considerable differences in time ands space of the hydrological conditions. Thermosubrosion may induce collapse over englacial interstices (Brodzikowski and Van Loon, 1979, 1980,1983; Brodzikowski, 1982a, 1984).

Sedimentation processes and supraglacial facies patterns Supraglacial sedimentation can take place not only directly from melting ice, but also from running or stagnant water, from waning winds and as a result of mass transport along weathered and eroded nunataks (Sugden and John, 1976; Embleton and King, 1977). All types of glacigenic sedimentary facies are therefore present in the supraglacial subenvironment (Fig. 89). The schematic model of Figure 89 describes the manner in which

deposition from running water

deposition from stagnant water

deposition by gravity-induced mass movements

deposition from melting ice

deposition by aeolian activity

Fig. 89. Main interrelations between the supraglacial depositional processes and the supraglacial facies and deposits.

138


the instability of the ice substratum affects the final pattern of the deposits formed under these conditions. It is obvious that sediments from all facies may become reworked again as long as ice is present in the substratum. Indeed, most supraglacial deposits remaining after retreat of the ice show signs of more or less intense reworking (Flint, 1971; Boulton, 1972b; N. Eyles, 1979; Paul, 1983). In addition, the deposits may be affected by meltwater currents. Thus, undisturbed supraglacial deposits are not commonly found. In spite of this, the original melt-out, glaciofluvial, glaciodeltaic, glaciolacustrine o r aeolian character of supraglacial deposits can generally be recognised as such (although these characteristics are generally better preserved in the periglacial environment). DEPOSITS OF THE SUPRAGLACIAL MELTING-ICE FACIES (1-A-1) Three types of deposits can be discerned in this facies (Fig. 90) on the basis of their genesis (also see the table on pages 128-129). The first type is the supraglacial melt-out complex (I-A-1-a),which is a mixture of till types or which cannot be identified more accurately. The second type consists of supraglacial ablation tills (I-A-1-b),formed fundamentally as a result of the melting of debris-containing ice under the influence of solar irradiation. The third type, supraglacial ice-raft deposits (I-A-1-el, consists of material that settled on the bottom of a supraglacial lake as a result of gradual undermelting or sudden overturning of ice rafts with debris accumulations. Ice-raft deposits are usually not well developed in the supraglacial subenvironment and it is difficult to recognise them as such in Pleistocene and older deposits. They will therefore be dealt with only briefly in this section; more details will be provided in the corresponding sections on iceraft deposits in the chapters on the continental and marine terminoglacial subenvironments.

Supraglacial melt-out complexes (I-A-1-a) These complexes represent material on top of a glacier or ice sheet t h a t has been affected by ablation; nevertheless, these complexes do not show the characteristics of typical ablation tills (see a later subsection). The complexes generally consist of ablation tills, fluvioglacial material and sometimes ice-raft deposits that are intensely interfingering and that are commonly also mixed by deformation processes such as collapse after

Deposits of the supraglacial melting-ice facies

139

Fig. 90. The supraglacial melting-ice facies (Hornsund area, Svalbard). A cover of debris rests on top of the ice, which is incised by meltwater streams and which contains abundant small pools and lakes. Most of the debris should be considered as supraglacial ablation till (1-A-1-b),but there are also some irregular masses of supraglacial melt-out complexes (I-A-1-a) where the ablation material is mixed with streamflood deposits. Supraglacial ice-raft deposits (I-A-1-e) may be formed in the lakes where melting ice rafts drop debris. Photograph: J. Cegfa.

melting of ice underneath (Dumanowski, 1961; Klatkowa, 1972, 1982; Szczepankiewicz, 1972; Brodzikowski, 1984). The characteristics mentioned above imply that most supraglacial melt-out complexes in the sense used here are comparable with a combination of the lowered till, melt out till and sublimation till (Shaw, 1977c) described by Hambrey and Harland (1981), and with a combination of the sublimation till and melt out till described by Boulton (1980a). The complexes are fairly common in the recent supraglacial subenvironment and have frequently been described in a morphological sense as 'complex supraglacial moraines' (a.o., Klimaszewski, 1960; Szupryczyfiski, 1965; Boulton, 1967; Karczewski and Wiiniewski, 1975), as 'supraglacial ablation drift' (Flint, 1971), and as 'supraglacial morainic till' (Boulton and Deynoux, 1981).

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Lithofacies characteristics

The supraglacial melt-out complex is commonly a typical diamict with irregular concentrations of relatively fine or coarse particles. There may be lenses or - more or less vague - layers with diverging grain size, but the complex may also have a massive appearance (Fig. 91). As a rule, there are frequent and rapid changes of grain size in lateral directions. The average grain size obviously depends on the size of the particles supplied by the ice. Pleistocene supraglacial melt-out complexes in the European lowlands tend to have a relatively small average grain size, but there are a great many exceptions. Recent complexes on Svalbard are usually relatively coarse, with minor admixtures of fine-grained particles (Szupryczydski, 1963; Klatkowa, 1982). Matrix-supported complexes show abundant sedimentary deformations such as load casts and water-escape structures (Drozdowski,1983).

Fig. 91. Supraglacial melt-out complex (I-A-1-a) in the Jordan6w region (SWPoland). Colour differences indicate variations in grain size. The central part is relatively homogeneous but there are dark bands of clay that may represent fluvial activity. The largest clasts have a diameter of 5 cm and the average content of clay- and silt-sized material is less than 10%. Height of photograph: approx. 1.5 m.


141

Textural characteristics The shape of the clasts in the complexes depends on their history. If an ice cap has eroded a soft-sediment substratum with fluvial gravels, most of the clasts will be rounded; if hard-rock erosion fragments dominate, angular clasts will prevail. Large numbers of clasts may show glacial striae in both cases. The clasts commonly have a random orientation. A preferred orientation is an indication of post-depositional reworking; the longest axis (the a-axis) of the clasts then tends t o become more or less horizontal. Continued post-depositional reworking may finally lead to a horizontal position of the clasts' ah-planes. If the complexes show a relative shortage of large clasts, a preferred horizontal position of the a-axis or the ah-plane may be primary as no obstacles prevented these clasts from taking this position during gradual melting of the ice.

Occurrence Supraglacial melt-out complexes form the highest part of glacial sequences as far as they developed before local retreat of the ice (if these complexes were formed and if they were not eroded afterwards). The thickness of these deposits ranges from less than a decimetre to a few metres. Thicker supraglacial till units are commonly built up by an alternation of ablation and flow tills, sometimes with glaciofluvial deposits in between. The complexes show a more or less random lateral distribution, as they accumulate mainly in depressions on top of the ice. The individual lenses usually have a diameter of less than one kilometre, and often much less. A few larger exceptions are known, among others from the Sudetes foreland, but in such cases the thickness of these complexes varies considerable in a lateral direction, mainly because of the irregular topography on which they rest. It has been emphasised by several of those working in the field (Ostrem, 1959,1962; Szupryczyiiski, 1963; Klajnert, 1966; Boulton, 1967; Szupryczydski and Kozarski, 1970; Kozarski and Szupryczyiiski, 1978; Klatkowa, 1972; Szczepankiewicz, 1972; Brodzikowski, 1982a, 1984)that supraglacial melt-out complexes are rarely found in association with deposits other than those from the melting-ice facies. There have been reports of transitions into the fluvial and mass-transport facies but the processes characterising these two facies usually are so destructive that there remain no traces of melt-out complexes.

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Depositional mechanism Supraglacial melt-out complexes are formed as a result of temperatureinduced melting and sublimation of the ice surface (Shaw, 1977~).The gradual disappearance of surficial ice implies that the embedded clasts are set free, forming a layer of material that goes downward together with the ice surface. The clasts are therefore replaced, even if there is no horizontal component; this replacement means that the original position of the clast within the ice is not generally preserved. The higher the ablation rate, the more meltwater is formed and the greater the chance of meltwater streams affecting the deposits. This is the main reason why melt-out complexes tend to be a combination of ablation tills and fluvioglacial material: the latter may be ablation material that has been transported over very short distances only and that originally belonged to the same deposit as the ablation material with which it again becomes mixed. Concentrations of fine-grained material in layers may result in levels with a low permeability, thus increasing the water content in the layer above. If there is also a high content of silt-sized material, the depositional process is commonly accompanied by synsedimentary and postsedimentary deformation processes; th;t: i s one more reason for the intermixing of fluvial and melt-out material that may take place.

A

debris band

slumped debris

crops out as ,,m d y

L

\

protects Ice a m ablation l!l!1

,

c

~.

debris band

1

\

Fig. 92. Development of small supraglacial melt-out complexes from a high-angle debris band in the ice. Ridges such as in A and B are common on glacier surfaces. Figure D is a hypothetical reconstruction of C, with the ice core removed. Modified, from Boulton (1967).


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. proximal

distal 4

abundant debris on tread

thrust planes

t

moratnc ridge

It11 hummocks

passive State

~ c ecore undergoing progreu,ve ablallo"

Fig. 93. Origin of large supraglacial melt-out complexes, as interpreted from field studies i n t h e Antarctic. From: Rains and S h a w (1981). Courtesy: J o u r n a l of Glaciology.

Boulton has proposed a depositional mechanism that seems particularly applicable t o supraglacial melt-out complexes of relatively small size (Fig. 92). Rains and Shaw (1981) provided another explanation, based on studies in the Antarctic and applicable to complexes of a larger size (Fig. 93). Both explanations stress the simultaneous occurrence of ice melting and sublimation, transport by meltwater streams, sliding over gentle slopes, and mass movements due to depressions within the ice substratum. Finally, frost action and mass transport result in mixing of the various types of deposits into these complexes (N. Eyles, 1979).

Supraglacial ablation tills (I-A-1-b) Supraglacial ablation tills (called 'ablation end moraines' by Kozarski, 1981) often form the uppermost truly glacial deposit in a glacial sequence (Flint, 1957; Lavrushin, 1976; Sugden and John, 1976; Embleton and King, 1977; Minell, 1979; Klatkowa, 1982; M.G.C. de Jong, 1983; Rappol, 1983; and others). The tills commonly are of limited thickness but have a large surface. They are fairly common in glaciated lowland areas and have been decribed frequently (e.g. Jahn, 1952; Flint, 1957,1971; Dumanowski, 1961; Drake, 1971; Pressl, 1971; Steward and McClintock, 1971; Boulton, 1972a, 1976c, 1980a; Lavrushin, 1980; Rappol, 1983; Brodzikowski and Van Loon, 1987; Dreimanis, 1989).

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The ablation conditions (a relative abundance of water; particles of different sizes; water-saturated deposits) facilitate mass-movement processes of these tills, particularly by flowage. While Boulton (1971) considers the reworked tills as a normal type of ablation till, such reworked material is commonly called 'supraglacial flow till'. It is considered here as another type of deposit (with subtypes: cf. Klajnert, 1966,1978; Klatkowa, 1972,1982; Olszewski, 1974; Nalewajko, 1982) and will be dealt separately, as one of the deposits belonging t o the category of mass-transport facies. It must be noted that the term 'ablation till' is frequently used in the literature to indicate a combination of supraglacial tills such as melt-out tills, sublimation tills, lowered tills and flow tills. This implies that the meaning in such cases is almost identical to that of our supraglacial meltout complexes (I-A-1-a).The terminology in the literature is, however, so inconsistent, that it is often impossible to unravel what precise type of supraglacial till the researchers mention. The supraglacial ablation till as defined in the present book has been distinguished as a separate type of deposit because of its specific lithofacies chartacteristics, occurrence and diagnostic value. The meaning of this type of deposit for facies interpretation has been recognised by many workers. Shaw ( 1 9 7 7 ~described ) it as 'lowered till and sublimation till'; Boulton (1970b, 1980b), Drozdowski (1974), Dreimanis (1980), Lavrushin (1980) and Shaw (1985) have described this type of till as 'supraglacial melt out till'. Lithofacies characteristics

Most supraglacial ablation tills have a massive appearance and consist mainly of sand-sized clasts. Both boulders and fine-grained material (dust and colloidal material) tend to be of minor importance, though there are exceptions in the form of gravelly ablation tills from which the sand and the fines have been washed out (Fig. 94); fine- or coarse-grained levels may also be present in predominantly sandy deposits. Laminations, lineations and current-induced structures a r e usually absent. Sedimentary deformations may be present, often resulting from water-escape processes, loading or similar processes. Textural characteristics Most material in supraglacial ablation tills has undergone a long transport process, partly embedded in ice, and partly not. The frequently


145

Fig. 94. Supraglacial ablation till from the Jordan6w area (SWPoland). This relatively coarse example is a massive, clast-supported diamict. The apparent layering a t the base and top of the succession is not of primary origin but is due to a small-distance reworking after collapse of a cavity in the underlying ice substratum. The band at the left in the middle of the succession is probably a flow till.

rounded shapes of boulders and other clasts suggests that transport by currents (in the supraglacial subenvironment or not) is important. The rounded shape may be destroyed by frost action, leaving two or more parts of the original particle with a partly rounded, partly irregular surface with sharp edges. Frost splitting is a common phenomenon in all supraglacial facies. Particles of sand size may show signs of aeolian transport, as clearly visible on SEM photographs. Pebbles and larger clasts commonly seem t o have a more or less random orientation, though tills are known that have a rougly horizontal orientation of the ah-planes of flattened clasts. Orientation of the a-axes, in particular if the preferred orientation is found with respect t o both the icemovement direction and the dip, is explained as being a result of postdepositional processes or of current activity that took place simultaneously with the melt-out process.


146

The other textural properties are not very helpful for the reconstruction of the genesis of the sediment: comparable characteristics are found in englacial melt-out tills. Occurrence

Supraglacial ablation tills can be found in the upper part of glacial sequences, or in the middle part if the 'regressive' sediments are also considered. The thickness of these tills ranges from a few tens of centimetres (thinner deposits are difficult t o recognise as such) t o several metres; the thickness may vary rapidly in a lateral direction: such variations in thickness are the rule, rather than an exception, even in one outcrop. Variations of this kind result from the fact that clasts concentrate m

0-

.-

x

.

a

.

.

*

.

.

1-

2-

. 3-

4-

. . x

-

Fig. 95. Supraglacial ablation till in the neighB bourhood of Chorzesz6w, - near L6di, central Poland. The till rests upon a glaciofluvial unit (I), consisting of medium and coarse, bedded sands. The ablation till (II)starts with fine-grained material (clay to fine sand) with a few gravels (2), followed by badly sorted sand (3). This is followed by silt without visible bedding (41, strongly deformed medium grained sands with internal lamination parallel to the deformed boundaries i (5), and mainly mediumgrained sands (partly finegrained sands) with some gravel (6).The ablation till is overlaid by a top series (III) consisting of sands of various grain sizes, with some gravel (7) and some I humus-rich sand (8).From: ; Klatkowa (1982). Courtesy: Societas Scientiarum Lodziensis.


147

in local depressions in the ice. Consequently, it is commonly impossible to make a sharp distinction between this type of sediment and supraglacial crevasse deposits. The thickest ablation tills tend to occur at the location of the maximum ice extent (this extension may be maintained for some time and the deposit is not destroyed by moving ice masses in such a case). Thinner deposits may be found over large parts of the glaciated area (Fig. 95), though small patches are more common, due t o partial erosion of the tills, especially by meltwater streams during deglaciation and by reworking. Depositional mechanism

Supraglacial ablation tills are formed by surficial melting of an ice mass, resulting in a concentration of previously embedded material accumulating without substantial horizontal dislocations. Because there are depressions at the ice surface which may be filled with ablation till, and because sediments in such depressions may be covered by other supraglacial sediments (e.g. flow tills, glaciolacustrine and glaciodeltaic deposits), the preservational potential is fairly large. Whether the till will be affected by water streams on the ice surface depends on the local conditions and characteristics of the deposits. Percolation of water may wash out the finest particles, which is the reason for the generally low content of fines. Consequently, the ablation tills tend to be the most permeable type of till, though much of the pore space may again be filled after burial. The post-depositional influence of water upon the properties of the ablation till has prompted theories that many deposits interpreted as ablation tills are in fact only a remnant of any supraglacial sediments that were transported and washed out by meltwater streams. The absence of current ripples etc. is, however, an indication that this interpretation is incorrect (cf. Klatkowa, 1982; Nalewajko, 1982; Morawski, 1984). One should keep in mind that ablation conditions imply - by definition the presence of meltwater streams, so that ablation tills may easily be affected by such streams and there is thus no sharp transition between ablation tills and supraglacial stream deposits. As mentioned before, this also holds for the transition between supraglacial ablation tills and crevasse deposits. Consequently, there are several intermediate types; Klatkowa (1982) mentions: (1) supraglacial bedded diamicts (well sorted, sometimes with gravel), (2) sandy sediments with a subtle, commonly deformed, lamination, with scattered and clearly oriented pebbles, and (3) massive lenses of diamicts formed in supraglacial pools. These intei

148


mediate forms indicate that the characteristics of supposed ablation tills should be analysed carefully because it is very likely that more than one depositional mechanism was involved. The basic depositional mechanism of supraglacial ablation tills consists of the gradual melting at the ice surface and the consequent accumulation of glacial debris that is set free (simultaneously, meltwater from the ice surface percolates through the thickening layer of debris). Much of the melting is a direct result, not of air temperature, but of the short-wave solar irradiation that heats the mineral particles; these particles gradually transfer the absorbed heat t o the ice. Sublimation may be important in extremely cold and dry areas but, in this case, the ice disappears much more slowly than during melting (Shaw, 1977c, 1981). Sublimation results in perfectly autochthonic sublimation tills that show no signs whatsoever of clast replacement. Sublimation tills also tend to contain more fine-grained material than do other types of ablation till, because of the absence of meltwater t o remove the fines. The good - theoretical - preservational potential of ablation tills has been confirmed by field observations, both in the case of Pleistocene sediments (a.0. Brodzikowski and Van Loon, 1983,1987; Morawski, 1984) and for recent conditions, e.g. on Svalbard (Szupryczyiiski, 1963; Boulton, 1971, 197233; Karczewski and Wihiewski, 1976) and in Iceland (Szupryczyriski and Kozarski, 1970; Kozarski and Szupryczyriski, 1978).

Supraglacial ice-raft deposits (I-A-1-e) Deposits formed in supraglacial lakes may contain lenses or layers of sediments formed as a result of the melting of ice rafts. These ice-raft deposits should be considered to belong to the melting-ice facies, although they are genetically strongly related to the glaciolacustrine facies. Supraglacial ice-raft deposits are fairly rare and are moreover difficult to recognise as such. Relatively good conditions for their formation are found in large, deep lakes. These lakes can be formed supraglacially especially when a continuous meltwater current is dammed off by, for instance, a moving ice mass or a huge subaerial mass flow that comes t o rest in the exit of the lake. Deposits thus formed during the maximum extent of the Scandinavian ice cap have been found in Middle Europe, where the ice front was stopped by mountain ranges (Sudetes, Carpathians, Erzgebirge) and where large lakes formed by ablation. The characteristics of supraglacial ice-raft deposits are only schematically described in the following subsections. More attention will be paid t o


149

similar deposits in the chapter on the continental and marine terminoglacial subenvironments because this is where these deposits are found most frequently. Lithofacies characteristics

The category of ice-raft deposits as described here includes several types of sediments, among which three main types can be distinguished. The first consists of a layer with an average grain size remarkably coarser than is normal in a specific lacustrine sediment. The second type consists of isolated (floating) clasts within a n otherwise fine-grained, 'normal' glaciolacustrine deposit. The third type consists of irregular lenses of coarse material within a 'normal' glaciolacustrine deposit. The genesis of these three types is explained under the heading 'Depositional mechanisms'. This third type of ice-raft deposit may be particularly difficult to distinguish from some types of mass-transport deposits. Textural characteristics

One of the characteristics of the ice-raft deposits in general is that the large clasts do not show a prefered orientation. There are several reports that the a-axes of the large clasts show a tendency to incline over 45". Wherever lenses of gravel-sized ice-raft deposits occur, the gravelly material is commonly still matrix supported. There is generally no sign of layering or sorting, thus leading to a homogeneous appearance. It is quite common t h a t the lenses overlie finer-grained glaciolacustrine deposits that are strongly deformed as a reaction to either the deposition of the gravel lenses or to the resulting unstable density distribution within the sedimentary succession. Such deposits have been described from the Fraser Canyon area (British Columbia, Canada) by N. Eyles (1987)and N Eyles and Clague (1987). Occurrence

Supraglacial ice-raft deposits can be found wherever supraglacia lacustrine deposits occur. They can extent over the entire area of glaciolacustrine deposition but are more commonly of restricted lateral extent. The ice-raft deposits can be found a t each level of a glaciolacustrine sequence; there may be several levels of ice-raft deposits within such a sequence, and two or more units may be found without a n intercalation of 'normal' glaciolacustrine sediments.

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Depositional mechanisms The three types of supraglacial ice-raft deposits just mentioned are formed by the following mechanism. The first type (relatively coarse layers) is a result of a more or less frequent passage of debris-laden ice rafts. These rafts undergo undermelting so that debris is continuously set free at the base of the raft. This debris commonly has an 'average' grain size, whereas the sediments settled in the centre of a lake are relatively fine. Consequently the 'rain out' of debris from the ice rafts results in a relatively coarse layer. A period with a low 'normal' glaciolacustrine sedimentation rate and a frequent occurrence of ice rafts can thus result in a real layer of relatively coarse material. The second type (a unit with isolated pebbles) consists mainly of 'normal' lacustrine sediments. The floating pebbles (dropstones) are supplied from occasional ice rafts that gradually melt away. The third type (lenses) usually results from sudden tumbling over of ice rafts after their position has become unstable due t o undermelting. Tumbling over results in a sudden release of the debris that had concentrated on the ice surface as a result of melting due to solar irradiation. The extent of these deposits is thus restricted to the location of the ice raft involved; this explains why such deposits (also called dump deposits) are commonly restricted to relatively small lenses.

DEPOSITS OF THE SUPRAGLACIAL FLUVIAL FACIES (I-A-2) Supraglacial topography is commonly irregular, often hummocky, and supraglacial streams tend to be discontinuous. They may disappear in crevasses or embouch in ponds or lakes without outlet, but they may also be formed suddenly by an outburst from an overpressurised englacial meltwater stream (Fig. 96). Elevated parts of a supraglacial sediment are easily eroded by supraglacial streams, but deposition - particularly in depressions - is much more common. This results in the frequent occurrence of supraglacial fluvial deposits, at least in modern environments such as on Svalbard and Baffin Island and in Iceland and Greenland. However, there are relatively few detailed descriptions of such deposits (see, among others, Flint, 1971; Boulton, 1972b; Sugden and John, 1976; Embleton and King, 1977; Paul, 1983; Goldthwait and Matsch, 1989). Much more

Deposits of the supraglacial fluvial facies

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Fig. 96. The supraglacial fluvial facies on the Werenskiold glacier (Hornsund area, Svalbard). Note the irregular topography determining the course of the streams, and the thick layer of glaciofluvial debris. A small spring, fed by pressurised englacial meltwater streams, is visible on the foreground. Photograph: J. Ceg+a.

information is available about Pleistocene sediments of this kind (e.g., Kozarski, 1962; Klajnert, 1966,1978; Bartkowski, 1967; Klatkowa, 1972; Drozdowski, 1974; Brodzikowski and Van Loon, 1983; Ehlers, 1983d; Brodzikowski, 1984; Kozarski and Kasprzak, 1987). Facies I-A-2 contains four types of sediments: supraglacial fluvial complexes (I-A-2-a), supraglacial tunnel-mouth deposits (I-A-2-c), supraglacial stream deposits (I-A-2-d) and supraglacial stream- and sheetflood deposits (I-A-2-e).These types are all fairly common and can be found frequently in modern supraglacial environments (Hartshorn, 1952; Klimaszewski, 1960; Szupryczyiiski and Kozarski, 1970; Loomis, 1970; Boulton, 1972b; Klimek, 1972; Kozarski, 1975; Kozarski and Szupryczyfiski, 1978; Boulton and Eyles, 1979; N. Eyles, 1979; Klysz a n d Lindner, 1982; Lindner et al., 1982). The most characteristic are the fluvial complexes and the stream deposits.

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Supraglacial fluvial complexes (I-A-2-a) The character of supraglacial currents changes so easily, both in space and time, that the resulting sediments can often not be attributed to one specific type of fluvial deposits. This occurs commonly under modern supraglacial conditions (Boulton, 1967, 197213; Kozarski and Szupryczyfiski, 1978; N. Eyles, 19791, and is clearly reflected by 'fossil' (Pleistocene) supraglacial fluvial deposits (Bartkowski, 1967; Drozdowski, 1974). Lithofacies characteristics

Supraglacial fluvial complexes are usually mixtures of sandy and gravelrich material, sometimes with finer-grained lenses, although these are more commonly absent (Fig. 97). Irregular layering is widespread, but massive units may also be present. The occurrence of structures like

Fig. 97. Supraglacial fluvial complex of Drenthian age in the neighbourhood of Boledawiec (SWSilesia, Poland). Note the alternation and interfingering of sediments of different fluvial origin, viz. stream deposits (predominantly in the upper left part), streamflood deposits (lower right) and - possibly - some overbank deposits (finergrained, deformed layers). Photograph: J. Czerwifiski.


153

current ripples and horizontal or wavy lamination depends on the depositional mechanisms involved but such structures are nevertheless commonly found. Rapid lateral changeability are main characteristics of the complexes. Sequences and cycles are absent.

Textural characteristics In general, the complexes consist mainly of debris set free by ablation and transported over some distance by supraglacial streams and floods. There may thus be clasts (mainly the larger ones) with predominantly signs of glacial transport (striae) and other clasts with predominantly fluvial characteristics (rounding). It is not unusual that a large percentage of the clasts shows relatively fresh surfaces along which they had broken as a result of frost scattering after final deposition. Both pebbles and sand-sized grains may be orientated in the same way as in non-glacigenic fluvial deposits. This implies that clast-supported units may show a well developed imbrication, whereas elongated sand grains tend t o have a horizontal position of the ah-plane, with the a-axis perpendicular to the current direction.

Occurrence Supraglacial fluvial complexes are formed most easily if there is no well developed channelised outflow of meltwater. Their formation is also favoured by a rapidly changing supply of water and continuous changing of the drainage pattern, so that no mature stage of deposition is reached. The consequence is t h a t the complexes have a restricted extent, a n irregular thickness, an irregular (though usually elongated) shape and vague boundaries; there is often a gradual transition into stream deposits, sheet- and streamflood deposits or crevasse deposits. The complexes can be found in the glacial sequence wherever supraglacial fluvial deposits are present; they generally form only a limited part of the supraglacial fluvial facies and individual complexes are a few metres thick at most.

Depositional mechanisms There is, obviously, no single specific depositional process responsible for the formation of the complexes. Their genesis is a result of more or less important contributions by the various fluvial processes.

154


Massive complexes (or parts of complexes) are commonly due to sudden deposition of material around springs, although i t cannot be excluded t h a t some of the massive deposits are due to debris flows (and might thus be assigned to the mass-flow facies); the generally elongated, narrow shape of the latter might be a criterion for distinction. The not uncommon reversed grading in some layers of the complexes indicates a grain-flow origin due to currents in the upper flow regime. Normal grading, due to waning flows, and the frequent occurrence of small-scale ripples indicate, however, that a lower flow-regime is the more common situation. The complexes seem to be formed most easily during early stages of deglaciation. However, observations on Svalbard (Klimaszewski, 1960; Boulton and Deynoux, 1981; Lindner et al., 1982) evidence that temporary formation on top of active ice is also possible. There is a tendency for the drainage system t o become better organised in the course of time during ongoing deglaciation. This explains why complexes are then less common. The main reason for the better organised drainage pattern is that ongoing deglaciation diminishes the hummocky nature of the ice because it tends to level the surface, both by better exposure of elevated parts to solar irradiation and by mass wasting, so that the conditions gradually become more favourable for increasingly 'structured' processes, which may result in stream deposits (I-A-2-d) and sheet- and streamflood deposits (I-A-2-e).

Supraglacial tunnel-mouth dep3sits deposits (I-A-24) The surface of ice caps and glaciers is so irregular in the ablation area that meltwater streams in englacial tunnels can easily reach the surface; in addtion, overpressurised water may be forced in the englacial drainage system to flow upwards, thus occasionally forming springs. Tunnel mouths may aldso result, either subaerial or subaqueous. The debris-transporting currents that embouch in a lake or on the ice surface spread out over a wider surface and thus undergo a drop in current velocity and consequently deposit some of the material carried along. If the currents were large enough and transported sufficient material, this series of processes results in tunnel-mouth deposits which differ mainly from deltaic deposits because the 'pipeline' character of the tunnels results in flow characteristics that are truly different from those in a surficial stream. If the englacial currents were relatively small and if little debris was transported, the resulting tunnel-mouth deposits could be insufficiently large and characteristic to be recognised as such.


155


Supraglacial tunnel-mouth deposits tend t o consist of irregular gravels with admixtures of sand (and sometimes boulders also). Fine-grained material is usually absent because it is removed by the current involved. The fairly chaotic sedimentary situation influences the characteristics of the final deposits, which may show no regularity whatsoever. There may also be, however, some kind of layering and the layers may even show grading. These supraglacial tunnel-mouth deposits commonly show a gradually better developed stratification in the downcurrent direction, thus passing into supraglacial stream deposits (I-A-2-d) without distinct transition. Cross-bedding is rare in deposits that should still be attributed to the tunnel-mouth area. Supraglacial tunnel mouths may be numerous, but they also tend t o be relatively small. This implies that their sideward extent, as a rule, is restricted and that no lateral changes can be found within the deposit. Still one may find, in the distal part of deposits formed in front of relatively large tunnel mouths, tabular cross-stratified sets with reactivation surfaces and aggradation ripples. Textural characteristics

These deposits have no particular textural characteristics (they are primarily recognised on the basis of their location and their relatively coarse appearance). Large clasts in the deposits may show imbrication. Occurrence

The main diagnostic feature of these deposits is their occurrence in the form of commonly small and irregular cones. These cones are often of more or less the same length in the longitudinal direction and in cross-section. Most of the units are less than one metre thick (Fig. 98). The deposits are surrounded by other supraglacial fluvial and masstransport deposits, and the same types of deposits are commonly found underneath and above. One finds often a series of tunnel-mouth deposits on top of each other, or separated by only thin units of different origin. Depositional mechanism

The main reason for the formation of these deposits is the sudden drop in current velocity of englacial streams that transport debris in both traction

156


Fig. 98. Gravel-rich supraglacial tunnel-mouth deposit from the L6di region (central Poland). The coarsest fragments were left behind without visible sorting when the currents left the tunnel mouth under conditions of the upper flow regime. There is a gradual transition towards the right into 'normal'fluvial deposits.

and suspension when they reach the ice surface. The sudden decrease of current depth which commonly accompanies the embouchure also favours deposition of the debris. It is quite a normal situation that current conditions in the tunnel mouth change rapidly from the upper flow regime to the lower part of the lower flow regime, but it is uncommon that the current velocity drops so much that fine sand and smaller particles can come to rest. The remaining current velocity is, at least in the proximal part of the area, generally still sufficient t o erode the substratum. That proximal tunnel-mouth deposits nevertheless exist indicates that variations in the water supply occur. If there has been erosion and if no subsequent filling of the depression has taken place, flow tills and other mass-transport deposits may contribute to the local sedimentation. There are only few descriptions of the depositional mechanism and the resulting sediments. Most data in the literature refer t o similar conditions under terminoglacial conditions (e.g., Banerjee and McDonald, 1975; Rust and Romanelli, 1975).


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Supraglacial stream deposits (I-A-2-d) It is well known that large parts of the supraglacial sediments consist mainly of bedded sands and gravels. Very coarse or fine parts are less common but do occur as lenses, mainly in depressions. Their layering may be irregular due t o settling along an inclined substratum, but deposition takes place mainly in a regular way, through gradual slowing of the supraglacial water currents. Supraglacial streams can exist on active ice but are more common on passive ice, where ablation prevails and a thick cover of debris may hide the ice (Fig. 99). The channels are generally much deeper in passive ice than in active ice (Fig. 100); a result is that supraglacial channels on active ice usually have commonly banks of debris, whereas channels on top of active ice are usually incised through the debris cover into the ice itself. After melting of the ice, (which may disturb the original sediment by collapse) the supraglacial stream deposits form topographic heights that are affected by erosion. Several deposits of this type will therefore not be

Fig. 99. Supraglacial channels on the Werenskiold glacier (SW Svalbard). The ice is active and the channels on top of it are relatively shallow (less than 2 m); they are incised mainly in the supraglacial debris. Photograph:J. Cegta.

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Fig. 100. Contact zone between active ice (foregound) and passive ice (background) at the Werenskiold glacier. Note the sudden increase in depth (from 1.2 m to 6 m) where the stream passes from the active ice to the zone of passive ice. The channel in the zone of passive ice is incised not only in debris but also in the ice itself. Photograph: J. Ceda.

preserved very long and if they are preserved, it may be extremely difficult to recognise them as such. Supraglacial stream deposits have been described frequently from both modern and ancient (mainly Pleistocene) environments, particularly from Iceland (Kozarski and Szupryczyiiski, 1978; N. Eyles, 19791, Svalbard (Klimaszewski, 1960; Szupryczyiiski, 1963; Boulton, 1972b; Klimek, 1972; Lindner et al., 1982) and several mountainous areas (Sugden and John, 1976; Embleton and King, 1977). Pleistocene sequences from central Europe that contain such deposits were analysed in detail by Bartkowski (1967), Grzybowski (1970), Klatkowa (1972) and Brodzikowski and Van Loon (1980,1983) and several others. Deposition of supraglacial stream deposits takes place under conditions of both active and passive ice, but the preservational potential of the former is rather low because of the more active erosion.


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Supraglacial stream deposits are well stratified, but the stratification may be strongly disturbed postdepositionally as a result of collapse of the underlying ice. Most current-induced sedimentary structures indicate a lower flow regime, with tabular cross-stratified sets, horizontal lamination and some trough-shaped beds being the most common. Gravel and sand are the dominant grain sizes, but local conditions (supply and current conditions) determine the precise granulometry. Most of the material generally consists of medium to coarse sand, with stringers of gravel in between (Fig. 101). N. Eyles (1979) described such deposits (mainly gravel-rich sand) from Iceland, where they form a supraglacial outwash plain, passing into sandy channel fills in the downcurrent direction. Studies of Pleistocene deposits carried out by Grzybowski (1970) revealed a complex depositional framework on an outwash plain (Fig. 102).

Fig. 101. Supraglacial stream deposits from the Jarosz6w zone (SW Poland). The sediment consists typically of sand layers with different grain sizes, with some gravel strings in between. The layers are usually erosive and some show current-induced cross-bedding.


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N

Fig. 102. Section through supraglacial stream deposits on an outwash plain near Kalisz Kaszubski (Poland). From: Grzybowski (1970). Courtesy: Acta Geol. Polonica.

Textural characteristics Field descriptions are consistent, in that there are no truly diagnostic characteristics of these deposits, but it appears that soft-sediment deformations are more common than could have been expected on the basis of grain size only.

Occurrence Supraglacial stream deposits are probably the most common type of deposit in the supraglacial subenvironment. They are made up of (the remnants of) sediments that formed in supraglacial channels. Since the drainage pattern changes continuously, mainly because of changes in the supraglacial topography, the entire supraglacial subenvironment - a t least as far as affected by ablation - is covered or was once covered by a channel. This implies that stream deposits will have been present everywhere but that much of these deposits may have been eroded afterwards. In spite of erosion, most supraglacial stream deposits consist of units that reflect their origin by their irregular, elongated shapes. These channel sediments may interfinger with, or be bordered laterally by deposits


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from the melting-ice and mass-transport facies. Flow tills, crevasse deposits and sheet- and streamflood deposits are particularly often close company. The same relationships are found in a vertical profile. This implies that these deposits may be encountered in the glacial sequence over the entire section that represents the supraglacial subenvironment. Depositional mechanism The deposits are formed in the same way as their counterparts in nonglacigenic channels. The main differences are that the geometry of the channels on the ice surface is partly determined by processes within the ice substratum, and that the currents in the channels receive a continuous supply of local meltwater during winter, whereas water supply may stop completely during the winter; these factors, however, do not really affect the depositional processes. The supraglacial conditions most usually result in low-sinuosity channels that are well comparable with braided systems, but several side bars, transverse bars and overbank deposits can be distinguished. The currents in the interbar channel zones usually belong t o the lower flow regime and produce current ripples. These ripples tend to be shortlived because later phases with higher-energy currents tend to destroy them.

Supraglacial sheet- and streamflood deposits (I-A-2-e) Supraglacial sheet- and streamfloods are well known from recent environments, particularly in regions with dead-ice (Fig. 103). The main reason for the occurrence of sheet- and streamfloods is the sudden supply of water, for instance after breakthrough of an ice or sediment barrier. The process has been described in the literature, particularly as it is seen in arid climates and in the terminoglacial and proglacial subenvironments. These processes and deposits will therefore be considered in the relevant chapters of the present book. Lithofacies and textural characteristics These deposits consist of gravels, gravelly sands and sands. There is a tendency for streamflood deposits (Fig. 104) to be somewhat coarser than sheetflood deposits (Fig. 105) but the locally available material is more important as determining factor for the grain size of these deposits.

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Fig. 103.Final stage of a streamflood on the Werenskiold glacier (SW Svalbard), with conditions of a waning flow. Coarse material was left behind in depressions; finer debris was carried along during the flood stage but some fines are deposited i n the slackening water, especially in pools that remain between the gravel concentrations. Photograph: J. Ceg+a.

The deposits are partly formed under conditions of the upper flow regime and antidune stratification has been reported from modern environments; there are no such reports touching 'fossil' supraglacial equivalents. Deposition of gravel under conditions of upper flow regime may also result in apparently structureless, massive units. Imbrication of pebbles and cross-stratification in sands is more common. Both are due to transport under conditions of the lower flow regime. Such conditions prevail during the final stages of the stream- and sheetfloods; the high-energy structures formed earlier are usually destroyed during the periods of waning flow. Occurrence Supraglacial sheet- and streamfloods are frequently occurring phenomena and their deposits are consequently numerous. They form sheets of irregular thickness, but gradually thin out to pass into stream deposits.


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Fig. 104. Streamflood deposits of Drenthian age in the Strzelin upland (SW Poland). The base is eroded into an ablation till. The lower part of the streamflood deposits consists of horizontally stratified sands; this part is followed by massive gravels with pebbles that have mostly horizontal &-planes.

Sheetflood deposits may be formed especially as a result of pulses of water from tunnel mouths, so that there is frequently a relationship between tunnel-mouth deposits and sheetflood deposits, both laterally and in a vertical section. Glacial sequences often contain sheet- and streamflood deposits embedded in stream deposits and mass-transport deposits. Szupryczyliski (1963), Boulton (1968, 1972b) and Klimek (1972) have described some examples found in such a context. Depositional mechanism Formation of these deposits is largely determined by the supraglacial relief and the pulsatile character of the water flows. A sudden outburst of water results in high-energy conditions and the situation of upper planebed conditions arises. Small-scale obstacles are, however, sufficient t o disturb the uniform conditions and antidune structures and scours and similar structures may result.

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Fig. 105. Relatively thick supraglacial sheetflood deposit of Drenthian age near Mokrzesz6w (SW Poland). The horizontally stratified sands pass into crevasse deposits.

By definition, the upper flow regime is followed by a lower flow regime characterised by more 'normal' currents, destroying the earlier bottom topography but commonly leaving a 'decapitated' unit of horizontally laminated material, which is covered by a unit with current ripples. The deposits formed during the upper flow regime may show reversed grading within some laminae, but normal grading prevails. DEPOSITS OF THE SUPRAGLACIAL DELTAIC FACIES (I-A-3) The supraglacial drainage system is largely determined by the relief, which is commonly hummocky. Consequently, there are many lakes and ponds and supraglacial streams may embouch in these (Shaw and Archer, 1979; Brodzikowski and Van Loon, 1987). The typically transitional character of the supraglacial deltaic facies (Fig. 106) was pointed out earlier by several authors, among which Ashley (19751, Banerjee and McDonald (19751, Church and Gilbert (19751, Gustavson (1975), Rust and Romanelli (1975),Saunderson (1975) and Shaw (1975).

Deposits of the supraglacial deltaic facies

165

Four types of deposits can be distinguished within t h i s facies: supraglacial deltaic complexes (I-A-3-a), sometimes called 'supraglacial paradeltaic deposits', supraglacial deltaic topsets (I-A-3-b), supraglacial deltaic foresets (I-A-3-c)and supraglacial deltaic bottomsets (I-A-3-d).

Supraglacial deltaic complexes (I-A-3-a) Small lakes and ponds reached by supraglacial streams result in more or less sudden slackening of the current, thus inducing sedimentation. As a principle, the sediments form deltas but their size may be insufficiently small t o allow a distinction between top-, fore- and bottomsets. Such small deltaic complexes or fans are most commonly of the Hjulstrom type. There are almost no descriptions of these complexes in the sedimentological literature. Studies that deal mainly with palaeogeographical and morphological aspects (e.g., Bartkowski, 1967; Aario, 1972, 1977; Klatkowa, 1972) provided a certain amount of data. supraglacial subenvironment (I-A)

I I,

uninterrupted buried dead-Ice (glacial environment)

!

supraglacial deltaic bottomsets

~

supraglacial deltaic

subglacial

5lream deposits

6

Fig. 106. Schematic model of the supraglacial deltaic facies (I-A-3). A: reconstruction based on field studies carried out near Zary (western-Poland). B: detail of A based on observations made near Konin Zagaliski (near Zary).

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Lithofacies and textural characteristics The grain size of the fans depends mainly on the material supplied, but there is commonly an intricate interfingering of sand and finer-grained layers (Fig. 107) with, possibly, some gravels in the very proximal part, occurring particularly as channel lags. The coarse-grained proximal part tends to be relatively massive (except for the possible presence of some channels) but more sedimentary structures appear in a distal direction: first, sands with lamination parallel t o the substratum (commonly accompanied by units with cross-stratification), followed by a zone with alternations of laminated silts and sands with ripple-drift cross-lamination and, finally, a fine-grained zone with mainly parallel lamination. Disturbances resulting from gravity-induced mass movements may be found throughout the complex. Occurrence The complexes form irregular, fan-shaped bodies with a longitudinal and a transverse extent of no more than a few tens of metres. Their thickness usually does not exceed two metres. In a vertical section, the complexes overlie supraglacial lacustrine deposits and are overlain by supraglacial fluvial deposits.

Fig. 107. Part of a supraglacial lacustrine complex of Drenthian age in the Jarosz6w zone (SWPoland). Note the discontinuity of the laminae and the differences in grain size.


167

Complexes of this type are found frequently in Pleistocene sediments as part of kames and kame plateaus.

Depositional mechanism The fundamental reason for the genesis of the complexes is always the slowing down of current velocity when a stream embouches in a body of more or less stagnant water, thus diminishing transport capacity. The result is that the coarsest particles are deposited a t once, whereas finergrained material may be transported somewhat further. The frequent occurrence of ripple-drift cross-lamination nevertheless indicates that not only are the bed-load particles soon deposited, but that sedimentation by settling from suspension also takes place. The gradual accumulation of water-saturated sediments will generally result in unstable conditions, so t h a t mass movements (debris flows, slumps, turbidity currents, etc.) occur. As a result, relatively coarse particles may be found in the distal part of these complexes. Cross-stratification occurs mainly due to currents with a more or less constant transport of sand as bed load. Fine-grained sands seem t o form current ripples more easily than do coarser sands, but it is likely that the former are easier to recognise, particularly because finer foresets in the current ripples may differ in colour from the rest.

Supraglacial deltaic topsets (I-A-3-b) In practice i t is almost impossible to distinguish between supraglacial stream deposits and deltaic topsets (unless lateral facies transitions are clear). This implies that some of the characteristics mentioned under I-A2-d could also apply to the deposits described here. Supraglacial deposits with this kind of uncertain origin have been investigated in detail in the Jaroszow area (Lower Silesia), where they were formed during the last stage of the Elsterian glaciation (Brodzikowski and Van Loon, 1983; Brodzikowski, 1984). A good description of supraglacial deltaic topsets was provided by Shaw and Archer (1979) in their clear elaboration of supraglacial deltaic sedimentation in the Okanagan valley (British Columbia, Canada).

Lithofacies and textural characteristics The lithofacies and the textural characteristics of these deposits depend mainly on the character, type and size of the delta. The topsets are most

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commonly sandy with abundant cross-stratification (Fig. 108). Thin drapings of finer material are frequently present. In large-size deltas, there may be distinct lateral transitions between coarse-grained channel deposits and fine-grained interdistributary bays. The size of supraglacial deltas is, however, rarely large enough t o sustain such 'bays' for a long time. Shaw and Archer (1979) have described very coarse topsets (Fig. 109). These include gravels with inclined (locally horizontal) stratification, much resembling proximal glaciofluvial outwash plains. Some finegrained intercalations are also present. These topsets represent Salisburytype deltas (Salisbury, 1896; Church and Gilbert, 1975). Occurrence

The topsets form the beginning of the transition between supraglacial streams and lakes. They can therefore be found intercalated, both lateral-

Fig. 108. Sandy current ripples in supraglacial deltaic topsets of Pleistocene age near Konin Zagaliski (Zary upland, western Poland). Note the irregular character, which is due to continuous changes of the flow regime. Intervals of stagnant water resulted in settling of clay and silt particles, forming drapes.


169

0 m

S

N 1 I

6m

2 I

3 6 m

I

20m

4 I

5 50m

diamictite: competent, sand-silt matrix, light buff. clasts

I

7

6 3Um

- well-rounded

I

100m

I

rocks, rip-up silts, disturbed sands

diamictite: friable, silty sand matrix, white-grey, c l a s t s - predominantly well-rounded rocks, internal stratification common diamictite: compacted, sandy matrix, blue-grey, closts

- predominantly

well-rounded rocks, massive ( ? )

gravels: poorly stratified, matrix-rich composition common, cobbler lreqvently maior component, well-rounded gravels: well rlrotified, open-work structure common, woll-rounded sands: frequently well stralified sand a n d silt: well stratified. frequently rhythmically laminated silt ( a n d clay): (rhythmically) laminated diomictic body: very ltiobte,randy s i l t matrix, clasts - rip-up silts predominate, some well.rovnded grovels

a

diamictic boulder talus, slump material

Fig. 109. Several profiles through supraglacial deltaic topsets near Harrop, British Columbia. From: Shaw and Archer (1979).Courtesy: A.A. Balkema (Rotterdam).

ly and vertically, between supraglacial stream deposits and supraglacial lake-margin deposits. In a lateral direction, the topsets gradudly wedge out and interfinger, not only with the stream and the lake-margin deposits, but also with supraglacial deltaic foresets.

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The thickness of the topsets is commonly less than one metre; Shaw and Archer (1979) described units that locally exceeded two metres. Topsets could possibly reach even far greater thicknesses in large lakes that deepen during deltaic sedimentation as a result of thermokarst. Depositional mechanism

The depositional mechanism does not differ greatly from that involved in supraglacial stream deposits. Most of the material is supplied, transported and deposited by distributary channels. The deltaic surface is only slightly inclined and vertical accretion of the topsets is commonly slow (cf. Clemmensen and Houmark-Nielsen, 1981);particularly if the flow regime is low. Gilbert-type deltas will thus be the type that will most commonly be formed. Salisbury-type deposits may be formed if the material is supplied mainly via sheet- and streamfloods. There is a high-energy flow regime in this case, and vertical accretion of the topsets tends t o be fast. This situation occurs particularly if the delta develops in the vicinity of a supraglacial tunnel mouth.

Supraglacial deltaic foresets (I-A-3-c) As in non-glacial lakes, the geometry of the deltaic foresets depends on water temperature, the nature of sedimentary particles and the amount of sediment supplied. The deltaic surface is usually rather flat, with a slope that is constant over almost the entire part of the foresets; the inclination may almost disappear in the distal parts of the foresets (cf. Kuenen, 1951; Gustavson, 1975; Gradziiiski et al., 1986). Cross-bedding may be visible throughout the succession; both tabular and trough-shaped forms are usually observed. Small current and wave ripples may be found in the finer-grained layers. Sequences may be found, as in the bottomsets, and are presumably due t o annual cycles. The sequences begin with relatively coarse material and end with silty or clayey laminae. The grading is, however, much less distinct than in e.g., the varves the deltaic or truly lacustrine bottomsets. A well-developed example of supraglacial deltaic foresets was found (and was used for palaeogeographic reconstruction) at Konin Zagaiiski (near Zary in Western Poland). The transitions between the various fluvial, deltaic and lacustrine deposits are well visible in this outcrop (see also Fig. 106). An excellent example of coarse-grained supraglacial deltaic foresets was presented by Shaw and Archer (1979) (Fig. 110).

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supraglacial stream deposits

ake level

I

shear planes

basal debris-rich ice

meltbout from below

Fig. 110. Rapid accretion of supraglacial deltaic foresets a t the junction between active and passive ice. Modified from: Shaw and Archer (1979).

Lithofacies and textural characteristics Supraglacial deltaic foresets most usually consist of gravel and sand, but their exact granulometry depends on parameters such as supply, flow regime and distance from the apex of the delta. Small-scale foresets deposited under high-energy flow conditions, for instance, tend t o be rich in gravel, with abundant coarse sand, less fine sand and a minor amount of silt and clay. Large-size foresets commonly show a decreasing grain size in distal direction. The structure of these foresets is, like that of other glacigenic and nonglacigenic equivalents, very characteristic: the stratification that is distinctly inclined, up t o 30°,is the most obvious feature (Fig. 111).This large-scale foreset stratification might be considered as being a kind of giant tabular cross-bedding (Jopling, 1962, 1965a,b; Church and Gilbert, 1975; Harms et al., 1975; Elliott, 1978). The inclined units are often laminated parallel to the overall (inclined) stratification. The distal parts of the foresets may contain ripple-drift crosslamination as a result of a combination of bottom currents and settling from suspension. This indicates relatively quiet conditions. However, conditions of more powerful current may still exist in the distal part, as shown by the occurrence of backflow ripples (Jopling, 1961; Clemmensen and Houmark-Nielsen, 1981).It is not uncommon that coarse units within the foreset are separated from one another by finely laminated layers of

172


fine-grained material. The latter might represent yearly periods of less accretion, due to the lack of debris being supplied, e.g. as a result of freezing of supraglacial meltwater (Shaw, 1977b; Shaw and Archer, 1979). The foresets most often show irregular layers formed as a result of mass flow over the inclined sedimentary surface (Fig. 112). The nature of the mass flow may range from short-distance sliding to turbidity currents, only a small part of which comes t o rest within the deltaic facies; the turbidity currents may reach the opposite side of the supraglacial lake, and currents coming from different points may become mixed. This implies that mass-transported layers within supraglacial deltaic foresets may show diverging palaeocurrent directions and not necessarily consist of material of the delta itself. Occurrence

These foresets are commonly found in the marginal parts of Pleistocene kames. They are usually at least twice as thick as the overlying topsets

Fig. 111. Typically inclined supraglacial deltaic foresets (thickness approx. 7 m) in Pleistocene deposits near Konin Zagafiski (western Poland), showing small-scale channels.


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but may be much thicker: the authors found several 6-8 m thick examples in exposures near Konin Zagaiiski (western Poland). The foresets generally have an irregular shape and the lower and upper contacts with the other deltaic units are gradual. The inclination of the foresets decreases with increasing distance from the source. Lateral contacts are seen mostly with stream deposits, other deltaic deposits and lacustrine deposits. Vertical contacts generally occur in the same way (Gustavson et al., 1975; Elliott, 1978), but there may also be direct contact with subaerial mass-transport deposits.

Depositional mechanism The depositional mechanism of these deposits does not differ from that operating for similar foresets formed under other conditions. The inclined overall stratification is due to grain avalanche and grain flow over an inclined surface.

Fig. 112. Distal part of supraglacial deltaic foresets near Konin Zaganski. Note the irregular layer in the centre, which represents mass transport over the inclined sedimentary surface.

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Changes in water turbulence determine whether fine-grained particles will settle whereas the bottom-current velocity determines the maximum grain size of the particles carried along. Rapid accretion leads to slope instability, resulting in frequent mass transport of various kinds.

Supraglacial deltaic bottomsets (I-A-3-d) The supraglacial bottomsets are deposited in the transitional zone between the deltaic foresets (I-A-3-c) and the supraglacial lake-margin deposits (I-A-4-c).Again, comparison with the lateral and vertical deposits may be necessary to allow the decision as to whether a particular deposit is of this specific type. These deposits often show a rhythmic (cyclic) sequence that several authors have studied in detail (e.g. Morgan et al., 1968; Aario, 1972; Ashley, 1972,1975; R. Gilbert, 1975; Gustavson et al., 1975; Shaw, 1975a, 1977b; Shaw and Archer, 1978,1979; Shaw et al., 1978; Clemmensen and Houmark-Nielsen, 1981). However, not all the studies were concerned

Fig. 113. Supraglacial deltaic bottom sets from the Jaroszbw zone (SW Poland). Note the interlayering of fine sand and clayeylsiltyunits. The varves are not well developed.


175

with deltaic bottomsets of a doubtlessly supraglacial nature. Typical supraglacial deltaic bottomsets have been described for the Jaroszow zone in SW Poland (Brodzikowski and Van Loon, 1983) and the Zary area in western Poland (Brodzikowski and Van Loon, 1980).

Lithofacies and textural characteristics The small grain size (often with a relatively high silt percentage) and distinct horizontal lamination are characteristic of almost all bottomset deposits (Fig. 113). Current ripples and ripple-drift cross-lamination may be observed in the coarser parts with fine sand. The deposits normally show the graded bedding of typical varves (see, e.g., Ashley, 1975); these varves may be a few cm or even a metre thick. The varves represent a one-year cycle, ending with a thin (1-2 cm) band of clay. Supraglacial bottomsets may also show 'subcycles' due to seasonal variations in sediment supply, but seasonal differentiation is commonly absent from small deltas.

Occurrence Supraglacial deltaic bottomsets are found in the glacial sequence between supraglacial lacustrine deposits and supraglacial deltaic foresets. The same relationships is seen in the horizontal direction (Fig. 114). It is also possible that the bottomsets are directly underlain by till. It may be difficult t o recognise the deltaic character of these deposits (lacustrine bottomsets are comparable in many respects), particularly if there is no unambiguous context. The supraglacial character may also be questionable, as similar deposits show identical characteristics in nonglacigenic and other glacigenic (sub)environments. Thus, sufficient data about lateral facies transitions and vertical contacts are required if the interpretation is to be reliable.

Depositional mechanism The grain size of the material that reaches the most distal part of the deltaic facies may vary. If sand is supplied, the bottom currents are relatively strong - for a t least part of the time. Current velocities in the lower flow regime are sufficiently great t o produce current ripples. Tractional grain flow, representing the upper flow regime, may also be present and, if present, results in high-energy planar-bed stratification. While slower currents will generally not supply much sand, the amount

176


A. stable terminal ice margin

B. retreating ice margin

. \ -

I

,

\

delta foresets

slide backslope

b. /

,

, \

-

,

I

I-,’

’

I

I

f- intermediate+-

- I/:

\,I

/

-

\

- \ ‘/.

rotated and subsided delta

sand-slide/turbidity-

D.

\

distal varves

high

d--

low

qualitative ice-retreaVsedimentation ratio

Fig. 114. Lateral and vertical relationships between supraglacial deltaic bottomsets and related sediments. From: Shaw (1977b). Courtesy: Geografiska Annaler.

may be just sufficient to produce climbing ripples. If no sand is supplied, settling of silt- and clay-sized particles from suspension is the dominant process. Massive or contorted layers indicate that mass-transport mechanisms are also present in the distal parts of the supraglacial facies.

Deposits of the supraglacial lacustrine facies

177

DEPOSITS OF THE SUPRAGLACIAL LACUSTRINE FACIES (I-A-4) As mentioned earlier, meltwater in the supraglacial subenvironment is commonly concentrated in ponds and lakes (Fig. 115). There are several types of bodies with stagnant water in which glaciolacustrine sedimentation may occur (cf. Brodzikowski, in press). The types include melt-out depressions, thermokarst depressions (also see N. Eyles and Rogerson, 1977a), widened parts of supraglacial channels and wide, shallow crevasses. Lakes may also be abundant if a rock barrier is met by the active ice, thus with drainage towards the foreland prevented (Fig. 116). This situation occurred frequently in Middle Europe during the Pleistocene glaciations and the resulting lakes were long-lived and developed in a characteristic way: they started as small ponds and lakes in the earliest stage of ice blockage, gradually grew and joined to form huge lakes during the phase of maximum glaciation, and finally broke up into hundreds of

Fig. 115. Small supraglacial lake (Werenskjoldglacier, Svalbard). The lake existed for several years then suddenly disappeared via a crevasse. The sediments in the lake were supplied partly by winds and inflowing rainwater, but mainly by sliding and slumping from the surrounding debris layer. Photograph:J. Ceda.

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Fig. 116. Supraglacial lakes formed due to blockage of the meltwater by a rock barrier in front of the ice. This situation occurred several times in Middle Euiope during the Pleistocene, when ice masses from t h e North encountered mountain ridges ( e x . Sudetes, Carpathians) along the way.

lakes of greatly varying sizes when the ice retreated (cf. Bartkowski, 1967; Szczepankiewicz, 1972; Szponar, 1974,1986; Eissmann, 1975,1981; Liedtke, 1975). It is quite usual that kames and kame terraces contain supraglacial lacustrine sediments (Klatkowa, 1972; Shaw, 1972b; Petelski, 1978; Schwan and Van Loon, 1979; Schwan et al., 1980a). The sediments of this facies can be studied adequately under present day conditions but, rather surprisingly, most descriptions and analyses of depositional processes and the resulting sediments concern Pleistocene material (see, among others, Jewtuchowicz, 1962, 1965; Szupryczydski, 1963, 1968; Czeppe, 1966; Boulton, 1968, 1972b; Szupryczyfiski and Kozarski, 1970; Karczewski and Wiiniewski, 1975,1977; Kozarski, 1975; Kozarski and Szupryczyfiski, 1978; Wis’niewski and Karczewski, 1978; Klrysz and Lindner, 1981,1982; Marks, 1981; Schwan and Ritzema, 1982). Three types of deposits are distinguished within this facies: the supraglacial lacustrine complexes (I-A-4-a), the supraglacial lake-margin deposits (1-A-4-b)and the supraglacial lacustrine bottomsets (I-A-4-c).


179

Supraglacial lacustrine complexes (I-A-4-a) Many of the supraglacial bodies of water are small. Ablation-induced meltwater streams, carrying along debris set free by ablation, are commonly seen. The combined occurrence of these two characteristics implies that several of the water bodies are relatively soon filled with detritus. Sedimentation usually takes place so quickly that there is insufficient time for the depositional processes to develop a differentiation of the sediments into lake-marginal deposits and bottomsets. Consequently, there occur many bodies of supraglacial lacustrine sediments that are best described by the term 'supraglacial lacustrine complexes' (the term 'supraglacial paralacustrine deposits' has also been used). Such deposits are known from both modern lakes and Pleistocene deposits (Bartkowski, 1967; Klatkowa, 1972, 1982; Brodzikowski, 1982a; Brodzikowski and Van Loon, 1983). Most of the reports that deal in detail with the palaeogeography and the morphology of glaciated areas also mention such deposits (e.g., Schluchter, 1979a; Ehlers, 1983; Evenson et al., 1983; Menzies and Rose, 1987; Goldthwait and Matsch, 1989). Lithofacies and textural characteristics The complexes generally consist of predominantly sandy material (most of the gravel comes t o rest in the deltas or in the streams, and there is usually less finer material) but coarse intercalations and layers o r lenses of fine material are not uncommon. The complexes generally convey a n impression, if they are not deformed later, of a simple alternation of deposits with grain sizes varying from silt to sand (Fig. 117). The lack of differentiation into lake-margin deposits and lacustrine bottomsets implies that sediments with the characteristics of bottomsets (particularly varvites) are absent or rare, since the water remained to turbulent to allow settling of the finest particles from suspension. On the other hands, current or wave ripples may be draped locally with fines. The bottom currents are often strong enough to produce ripples, but structures pointing to a high flow regime are almost always absent. Lowenergy parallel lamination is generally abundant and wave ripples are also commonly found. Occurrence The complexes tend to form irregular, lense-shaped bodies of stratified sands and sandy muds, sometimes interfingering with mass-flow deposits.

180


Fig. 117. Deposits of a lense-shaped supraglacial lacustrine complex in the Kleszcdw graben (central Poland). The lake was probably formed due to local subsidence of the hard-rock substratum (graben activity). Postdepositional deformation has disturbed the originally 'quiet' character of the sediments. Photograph:A. Hahszczak.

They are found in a vertical section commonly between or just on top of deposits belonging to the supraglacial melting-ice facies. The deposits are usually overlain by supraglacial stream deposits or by sediments of the melting-ice facies. The horizontal relationships are the same; lateral transitions may be either abrupt or gradual. It is important from a palaeogeographical point of view t h a t the complexes are most frequently found where the ice melted away rapidly. Such conditions often result in sudden and fast modifications of the supraglacial relief, thus providing the depressions required for glaciolacustrine sedimentation. Moreover, fast ablation also implies a considerable amount of meltwater that may supply proportionally large amounts of debris. The irregular relief caused by fast ablation also leads t o numerous small lakes and ponds rather than t o a limited number of large bodies of stagnant water, so that the conditions in this subenvironment are not very favourable for differentiation between lake-margin deposits and lacustrine bottomsets.


181

Depositional mechanism

Supraglacial lacustrine complexes represent fast deposition, which, however, occurred under low-energy rather than high-energy conditions. Material is supplied partly by supraglacial streams and is deposited by lacustrine bottomcurrents as a result of slackening currents. Fine particles may settle from suspension if the water is not very turbulent. It is likely that mass flows from the surrounding, irregular relief contribute significantly t o the deposition. Material supplied by winds may be important in specific cases but it usually plays only a minor role.

Supraglacial lake-margin deposits (I-A-4-b) For the present purposes, lake-margin deposits are considered t o consist of all sediments directly influenced by processes related to sedimentation within the glaciolacustrine facies, with the restriction that they surround a central part of the lake where accumulation takes place mainly by settling from suspension and/or from supply by mass movements. Sediments grouped in the deltaic and fluvial facies are also excluded, although they may otherwise fulfill the requirements. There are two main types of supraglacial lake-margin deposits: the first type consists of sediments that may be exposed alternately t o subaerial and subaqueous conditions (e.g. as a result of wind-induced rise of the water on one side of the lake); the second type consists of sediments that are always in a subaqueous position but are still distinctly influenced by waves and/or relatively strong bottom currents. It is obvious that the first type requires rather large lakes and is thus rarely seen in the suprnglacial subenvironment. The second group occurs much more commonly (Reineck and Singh, 1973; Reading, 1978; Walker, 1984; Gradziiiski et al., 1986). While supraglacial lakes are frequently found, lake-margin deposits are less common, at least in 'fossil' sediments. The first reason is that the lakes tend to be short-lived, thus providing little opportunity for differentiation of bottomsets and lake-margin deposits; the second reason is that lake-margin deposits are easily destroyed by mass-transport mechanisms and other processes affecting supraglacial sediments of restricted thickness. The above reasons explain why relatively few lake-margin deposits are recognised as such in 'fossil' sediments and described in detail in the literature. Such descriptions do exist however (Brodzikowski and Van Loon, l983,1985a, 1987; Brodzikowski, 1984).Much more is known about such sediments in the terminoglacial and proglacial subenvironments

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(C.H. Eyles and N. Eyles, 1983a; Walker, 1984; C.H. Eyles, 1986; N. Eyles and Clark, 1986; N. Eyles and Clague, 1987). Deposits of this type will therefore be dealt with here in more detail in the relevant sections of the chapters on terminoglacial and proglacial sediments. Lithofacies and textural characteristics Well-developed supraglacial lake-margin deposits were described by Brodzikowski and Van Loon (1983, 1985a) from the Jaroszdw zone (SW Poland). They consist of laminated material, ranging from silty clay t o gravelly sand. The lamination may follow the slope of the lake. The 'subaerial' lake-margin deposits (Fig. 118) are much less regular than the purely subaqueous ones. They may resemble the supraglacial lacustrine complexes, and small frost fissures may be present. The truly subaqueous sediments (Fig. 119) tend t o be somewhat more regular (with a longer continuation of laminae), to have a smaller average grain size, and t o be more susceptible to penecontemporaneous deformations.

Fig. 118. Supraglacial lake-margin deposits from the Jarosz6w zone (SW Poland). The gravel-rich layer (upper left) was possibly exposed to subaerial conditions during a period of low water, and was subsequently slightly reworked subaqueously.


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Fig. 119. Supraglacial subaqueous lake-margin deposits of Wartanian age from the Twardog6ra Hills (Poland). The originally regular laminae show penecontemporaneous deformation. The sandy layer (lower part) may represent a collision flow. Thickness shown approx. 50 cm.

Occurrence It is difficult t o recognise these deposits if there is no clear context. The situation is optimal if there is a continuous lateral exposure from supraglacial fluvial and deltaic deposits to supraglacial 1acustrine.bottomsets. The same is true for recognition in a vertical section, where these deposits may show thicknesses of the order of one metre. However, there may also be lateral and vertical contacts with deposits that belong to the supraglacial melting-ice and mass-transport facies.

Depositional mechanism Bottom currents and wave action are probably the factors most strongly determining for sedimentation. Material may be supplied by mass movements, supraglacial streams or deltas, or by rain- or meltwater-induced wash-off. The presence of material sorted by current action and waves will result exactly as along non-glacigenic shores; a net transport of coarse

184


material towards the coast and net transport of fine material to the centre of the lake will occur. Sedimentation most probably takes place continuously under lowenergy conditions, with the exception of sudden mass movements t h a t may represent temporary high-energy levels.

Supraglacial lacustrine bottomsets (I-A-4-c) Of the deposits from the supraglacial subenvironment, this type is one of those best described (see, e.g., Shaw, 1972b, 1975a, 1977b; Shaw and Archer, 1978, 1979; Brodzikowski and Van Loon, 1980, 1983, 1987; Brodzikowski, 1984; N. Eyles, 1987; N. Eyles et al., 1987a). The sediments are formed in the central, deep parts of lakes. Particles settle out from suspension during periods of quietness (for instance during winter when the lake is covered by an ice layer). Another important source is sediments that have accumulated at the lake margin, became unstable and were transported en masse to the deeper parts. Aeolian influence may be recognisable if supply from other sources was limited but wind-supplied particles are not often traceable in the sediment. The commonly silty character of these deposits makes them most susceptible to deformations. Melting of the underlying ice may be the trigger for penecontemporaneous deformation processes. Lithofacies and textural characteristics Varves (sensu De Geer) are characteristic of these sediments. According t o classical theories, the grading in the varves (Fig. 120) is due to the deposition of coarse material during summer and of fine material (silt and clay) in the winter; winter deposition takes place in the quiet water below a n ice cover. Depending on the supply, the thickness of one varve may vary from a fraction of a millimetre to several centimetres (Shaw, 1977b; Shaw and Archer, 1978; Brodzikowski and Van Loon, 1983). Complex horizontal lamination may be seen instead of varves. The parts formed during summer are relatively thick and may contain current ripples. More usually, however, fine-grained silty laminae predominate, which were deformed by penecontemporaneous and postdepositional processes in such a way that little of the original lamination is left. There are also supraglacial lacustrine bottomsets t h a t do not show typically varved units, but mainly simple horizontal lamination (Fig. 121). Such deposits often consist of fine sand and coarse silt and there may be small current ripples and wide, shallow channels.


185

Fig. 120. Typically varved supraglacial lacustrine bottomsets from the Jarosdw zone (SW Poland). Note the regularity of the undisturbed (dark) part and the intense metadepositional deformation of the (light-coloured) unit with alternations of fine sand and silt.

Occurrence

These deposits are easily recognisable in the supraglacial part of a glacial sequence because they contain varves. They may be associated with supraglacial lake-margin deposits, but the latter may not have been preserved properly. The bottomsets tend, instead, t o be associated with deposits grouped in the melting-ice, the fluvial, the deltaic and the masstransport facies. Horizontal relationships were investigated in detail by Brodzikowski and Van Loon (1983) who described sediments from the Jarosz6w zone (SW Poland). Deltaic bottomsets in these deposits gradually pass into lacustrine bottomsets within a horizontal distance of 100-150 m. The lacustrine bottomsets, in turn, pass into lake-margin deposits. The thickness of the deposits depends mainly on the depth of the supraglacial lake and the duration of sedimentation. A thickness of a few metres may be reached, but one of a few decimetres is much more common

186


Fig. 121. Handpiece of the bottomsets that fell dry after disappearance of the lake shown in Fig. 115. Note the regular lamination and the barely incised channel fill in the top part. Photograph: J. Cegla.

(the bottomsets of the lake shown in Fig. 115 had a total thickness of sixty centimetres).

Depositional mechanism Much work has be done in attempts to explain the varved nature of most of these sediments (e.g. by Kuenen, 1951; Woldstedt, 1957; Ashley, 1975; Gustavson et al., 1975; Shaw, 1977b; Merta, 1978). It is now generally agreed that turbidity currents are responsible for the formation of part of the graded layers (Schwan et al., 1980a even described a 'double-source turbidite') but it is doubtful whether they can be considered to be the main mechanism (Shaw et al., 1978) under supraglacial conditions. Many of the laminae with normal grading seem to result from parapelagic accumulation (Brodzikowski, in press) caused by a pulsed input of suspended particles into the lake. Some of the suspended material may be autochthonous, i.e. whirled up in the basin during shocks due t o thermosubrosion.

Deposits of the supraglacial aeolian facies

187

Aeolian supply, while it may also play a role, seems t o be of minor importance for most supraglacial lakes. However, particles accumulated through wind action on the ice-covered lake surface during winter may certainly constitute a source for deposition after melting of the ice cover. It can be concluded that a number of processes are involved. Neither the classical theory of season-determined settling, nor the later theory of predominantly turbidity currents seems satisfactory. Both play a role, together with wind action. More details will be considered in the relevant sections on terminoglacial and proglacial lacustrine bottomsets. DEPOSITS OF THE SUPRAGLACIAL AEOLIAN FACIES (I-A-5) The climatic conditions in glaciated areas favour wind directions away from the ice, rather than from the periglacial environment toward the ice. There is consequently no regular supply of windblown material. Moreover, the processes that affect the supraglacial subenvironment tend t o prevent the preservation of sedimentary covers, so that coversands and loess covers of any significant extent or thickness cannot develop. A relative scarcity of sand implies that dunes are generally not formed and if they are present, their preservational potential is practically nil. This, however, does not mean that sand and silt accumulated in the supraglacial subenvironment cannot be redistributed by wind action. Wind may affect the surficial sand and silt, particularly during late summer and autumn, before snowfall, when the sediments are relatively dry and when storms occur frequently. However, aeolisation is only temporary and its final effect is small, almost negligible from a geological point of view. Local accumulations of windblown material may exist in spite of this, particularly on the leeside of obstacles such as large boulders or glacier tables. Such supraglacial drift sands (I-A-5-b)are the only type of deposit in this facies. Observations on the Werenskjold glacier (Svalbard) indicate that such drift sands contain predominantly sand, with a minor amount of silt.

Supraglacial drift sands (I-A-5-b) Supraglacial drift sands have been described from several locations, from SW Svalbard among others (Jahn, 1961a,b; Kida, 1981,1985; Baranowski and Pqkala 1982). Taber (1953), Rozycki (1957), Ohlson (19641, H.T.U. Smith (1964), Peve (1968), Akerman (1980) and Bryant (1982) have described them from other places.

188


No descriptions of 'fossil' supraglacial drift sands are known to the present authors, but there are reports of supraglacial deposits containing grains that have a texture indicating transport by wind ( J a h n , 1975; French, 1976; Washburn, 1979).

Lithofacies characteristics Supraglacial drift sands are usually massive or horizontally stratified bodies of sand with a n admixture of silt. The silt content in individual bodies can vary from about 10% to about 30%, which sometimes makes them resemble a sandy loess (Kida, 1981,1985). Sandy bodies may show cross-bedding, but the small size of the bodies and the irregular internal structure make the precise buildup and geometry of the cross-bedded units difficult to unravel. The general lithofacies characteristics are strongly influenced by local conditions (Ceda, 1972; Pekala, 1980).

Textural characteristics The relatively short periods of aeolisation commonly do not result i n classical aeolian forms; the textural characteristics of supraglacial drift sands are therefore not very aeolian-like. Some of the grains, particularly the larger sand grains, may show an aeolian surface but this is commonly not well developed.

Occurrence Supraglacial drift sands generally occur randomly. The deposits form small concentrations, sheets or ridges only a few centimetres thick and no more than a few metres in diameter. Large obstacles may allow sand heaps to form with a height up to about 15 cm (Kida, 1985). Drift sands are unstable, change positions continuously and usually are finally mixed with other material (e.g. supraglacial stream deposits). Drift sands have the best chance of surviving for some time in depressions with a wetted bottom. This implies that drift sands are aften found in abandoned channels. The preferred accumulation on wet surfaces results from the larger force required to remove the material from such surfaces (Cegla, 1972). I t was mentioned earlier t h a t no reports are known t h a t contain descriptions of 'fossil' supraglacial drift sands. Accordingly, the position of such deposits in the glacial sequence is only hypothetical. One might,

Deposits of the supraglacial aeolian facies

189

however, expect them to occur - assuming t h a t they a r e present - in association with supraglacial stream deposits and supraglacial lakemargin deposits. Depositional mechanism

The most characteristic process is short aeolian transport, after erosion of sand and silt from moraines, talus cones, sanders and abandoned channels. The strong winds may suddenly blow up the material, transport it over some tens or hundreds of metres, then concentrate it along the flow lines of the wind over the ice surface or on the leeside of obstacles. Much of such wind-blown material will ondergo repeated aeolisation of this type. The most common end of an aeolian history is deposition of the grains in a lake or a stream where they become incorporated into the sediments of another facies.

DEPOSITS OF THE SUPRAGLACIAL MASS-TRANSPORT FACIES (I-A-6) The supraglacial subenvironment offers excellent opportunities for mass transport. Several types of mass movements do occur and are not truly different from those occurring under non-glacigenic conditions. Some of the processes take place subaerially. While the resulting sediments are grouped together as supraglacial subaerial mass-transport deposits (I-A-6-a), a wide variety of processes and deposits are included. The supraglacial crevasse deposits (I-A-6-b) that result from commonly complex combinations of subaerial and subaqueous mass movements and are often also influenced by meltwater currents (Fig. 122) are more characteristic of the supraglacial subenvironment. The third type consists of sediments t h a t were formed by subaqueous mass movements; these supraglacial mass-transport deposits (I-A-6-c) include the sediments that are commonly - but incorrectly - termed 'flow tills'. Most of the sediments of this facies consist of material t h a t had previously been deposited in the supraglacial deltaic and lacustrine facies, b u t reworking of material from the supragflacial melting-ice and lacustrine facies is not uncommon. No data are available that indicate reworking by mass transport of supraglacial drift sands. Sediments of this facies have been described in detail by several authors, among others Boulton (1972b, 1976c, 1980a), Dreimanis (1976a,b, l978,1980,1982b, 1987,1989) and Evenson et al. (1977).

190


Fig. 122. Supraglacial crevasse 5-10 m deep and 50 m long in the Werenskiold glacier (Svalbard). The bottom of the crevasse is covered with debris supplied via several types of subaerial mass wasting mixed with material that was washed down from a small englacial t u n n e l (see centre left), and is slightly affected by temporary meltwater streams flowing over the floor of the crevasse. Photograph: J. Cegga.

Supraglacial subaerial mass-transport deposits (I-A-6-a) Most debris in the supraglacial subenvironment was originally embedded in the ice. The particles were set free by ablation and possibly later were transported by one or more agents. Ablation may be irregular and the relief of the ice mass may also be affected by collapse. The topography is thus commonly irregular, particularly in the ablation zone. The combination of slopes with water-saturated debris is ideal for the occurrence of subaerial mass movements. The processes involved include rock fall, sliding, slumping and so on (Fig. 123). The resulting deposits were termed 'flow tills' by Boulton (1968), but the only relationship with till (in its genetic meaning) is the final derivation from - in many cases - ablation till. The term 'flow till' is therefore best avoided. Use of the term 'diamict' for such deposits (Flint et al., 1960; Harland et al., 1966; N. Eyles et al., 1983c; Dreimanis, 1989) is also unfortunate, because the granulometry is not necessarily diamict-

Deposits of the supraglacial mass-transport facies

191

Fig. 123. Ablation zone in the Hornsund area (Svalbard), with an irregular ice mass piercing through a n equally irregular debris cover with a thickness of mostly 0.5-1 m. Single particles have rolled off the exposed ice mass; some slumping occurred a t the foot of the ice and some of the material was fluidised (foreground, left). Photograph: J. Cegta.

like. It is thus preferable to term these deposits 'supraglacial subaerial mass-transport deposits', which would seem much more consistent with the argumentation put forth by Lawson (1979, 19811, Gravenor e t al. (1984) and Shaw (1985). Lithofacies characteristics The deposits of this facies resemble their equivalents from non-glacigenic facies. The granulometry depends primarily on the type of material available, but a diamictic grain-size distribution is that most common. Cohesive debris flows may result in either clast-supported or matrixsupported deposits that form lobes with internal structures that reveal a flowage mechanism. Shear structures may be present, indicating turbulent flowage. The depositional mechanism may result i n local sorting of the material (Fig. 124). Normal grading is also common; it is even not uncommon that relatively large clasts are concentrated at the

192


Fig. 124. Supraglacial debris-flow deposit of Drenthian age, from the Jarosz6w zone (SW Poland). The high-density flow developed during deglaciation, reworking fluvial (coarse), lacustrine (fine) and melt-out (unsorted) deposits. Photo approx. 2 m high.

base of deposits that are fine-grained if considered as an entity (Dreimanis, 1989). Very dense debris flows generally result in deposits that do not show any internal lamination. Textural characteristics

Gravity-induced mass movements always result in orientation of the particles along the flow lines or the shear zones: the a-axis is commonly directed along the flow line but, in some cases, its orientation is perpendicular to the flow line. The general geometry of the unit involved should therefore be considered if clast orientation is used for palaeocurrent interepretation. The greater the flow velocity the better the orientation of the individual particles. N. Eyles et al. (1987a) also mentioned the occurrence of a slight imbrication, with a fabric parallel to the flow direction. The larger clasts may show glacial striae but no other specific textural characteristics are seen (Middleton and Hampton, 1973,1976).

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Occurrence

By definition, mass-transported deposits are to be found in depressions. The hummocky relief of the supraglacial environment provides abundant depressions, so t h a t a complex pattern of supraglacial subaerial deposits may be formed. However, the supraglacial relief is largely determined by irregularities in the ice thickness. After the ice has melted, the position of the supraglacial subaerial mass-flow deposits is therefore apparently random. Most of the deposits are nevertheless well recognisable because of their limited extent, lobe shape and internal structure. They are usually surrounded by supraglacial melt-out complexes and ablation tills. The thickness of the individual deposits varies widely, from less than a centimetre to several metres. Thicknesses of more than 2 m are, however, seldom seen. The total thickness may nevertheless be much greater because a number of mass-flow deposits may form on top of each other until the local depression has been filled (see also Fig. 125). The supraglacial subaerial mass-flow deposits in the glacial sequence are positioned in the topmost part of the supraglacial sediments, just under the (possible) deposits of terminoglacial origin formed during retreat of the ice. Depositional mechanism

The precise mechanisms of deposition, precisely as under non-glacigenic conditions, depend mainly on the grain-size distribution of the material involved, the water content of the material, the inclination of t h e sedimentary surface and the 'lubrification' of the contact plane between the moving mass and its substratum. The volume of the moving material also plays a role, but the volume may change (increase o r decrease) during the mass movement itself (due to erosion or deposition). The most common glacier slumping

;;Zldascial I

=

debris 50m

I

/,< 7\

' inferred upper limit 01

supraglacial subaerial mass-transport deposits

;;;dsand

buried glacier Ice

Fig. 125. Horizontal section through the terminal moraine of Dunerbreen (western Svalbard), showing a cover ofmass-flow deposits. Modified after: Boulton (1968).

194


mechanisms are rock fall, sliding and slumping. It is not uncommon that the transport mechanism changes in the course of time: clasts may fall, induce a slide and be transported further as a subaerial slump. Transport velocity may be either high (e.g., in the case of low-concentration debris flows) or low (very cohesive debris flows). Low velocity and small volume are the best conditions for rapid filling of depressions. Processes that require a relatively large amount of water (e.g. slumping) do not take place where the temperature is below the melting point of ice. This implies that supraglacial subaerial mass-transport deposits are rare outside the ablation zone (Boulton, 1972a,b; Sugden and John, 1976; Baranowski, 1977; Shaw, 1 9 7 7 ~ ) . Subaerial mass-transport deposits a r e far more a b u n d a n t i n t h e proglacial subenvironment. More detailed information on these deposits will therefore be given in the chapter on the proglacial subenvironment.

Supraglacial crevasse deposits (I-A-6-b) Supraglacial crevasses may result from tensional forces in the ice mass or from collapse above englacial cavities. The crevasses form elongated, relatively narrow depressions from the ice surface downwards, commonly in a vertical direction (crevasses may be inclined, but openings with slopes of less than 45" should be considered as tunnels rather than as crevasses). The depth of supraglacial crevasses may reach up to about 20 m on active ice, and even more on passive ice. Such depressions are gradually filled by mass-flow deposits and by particles left behind by meltwater currents in the crevasse. This implies that crevasses constitute locations of major depositional activity in the supraglacial subenvironment (cf. Flint, 1971; Sugden and John, 1976; Embleton and King, 1977). The sediments commonly show exceptionally rapid vertical and lateral lithofacies changes. These sometimes more or less chaotic sediments remain after deglaciation as topographic elevations (Fig. 126) called 'kames' (see pp. 9-10]. Much is known about the sediments found in kames (Holmes, 1947; MacKay, 1960; Bartkowski, 1967; Embleton and King, 1968; Flint, 1971; Worsley, 1974; Francis, 1975; Johnson, 1975; Klimaszewski, 1976; Schwan and Van Loon, 1979; Kurimo, 1980; Sharp, 1985). Boulton (1971, 1972a) compared Pleistocene crevasse deposits with kames in recently glaciated areas. There is generally rather irregular bedding in the lower p a r t of a crevasse infilling, with structures that show no preferred orientation. Whith gradual filling of the crevasse, the character changes into that of a 'normal' basin; current ripples and other structures then occur frequently.


A

isolated kame hill

kame terrace

195

kame plateau

C

lacustrine lacies

lodgement

lodgement tills

ICIIS

stream

mass-transport deposits

lodgement 1111

Stream deposits

stream deposits

llow ltl,

lodgement 1,Il

lodgement 1,115

stream deposm

subglacial channel deposits

subglacial channel deposits

lsc"sIrln.3

1011

svbglac,al channel deposils

cavities

Fig. 126. Schematic model for the formation of (A) kame terraces, (B) kame hills and (C) kame plateaus. Stages I-III indicate subsequent steps in the development. Unless indicated t o the contrary, all facies and deposits are from the supraglacial subenvironment. Adapted from: Brodzikowski and Van Loon (1987).

Since crevasses move along with the ice, most infillings become disturbed. If the ice no longer moves actively, deposition may take place in a gradually deepening crevasse, which also implies a certain degree of deformation.

Lithofacies characteristics The lithofacies characteristics of these deposits depend to a great extent on the depth and width of the crevasse, on the inclination of the crevasse walls, and on the relief in the direct vicinity. Deep, narrow crevasses are filled mainly by mass-transport processes. These processes may for some part show a subaqueous character, viz. if the crevasse is filled with water

196


because there is no connection with an englacial or subglacial drainage system. The result is a combination of deposits with widely varying characteristics, ranging from massive gravels and sands t o clays and silts with strongly deformed lamination or with vague remants of the internal structure that existed before fluidisation took place (Fig. 127). An even more complex lithofacies pattern can be found in crevasses that are much wider. Such crevasses, typical of a n advanced stage of deglaciation, may contain not only all kinds of mass-transported sediments but also deposits formed in stagnant or running water. The continuous supply of sudden sediment pulses then results in a partial or complete mixture of lacustrine and fluvial complexes with mass-transported material. It is not uncommon that such mixtures are reworked and reshaped again by glaciofluvial processes at the crevasse bottom. Another process that may affect the sediments in crevasses is the change in water level, due to damming off of water currents by ice or sediment barriers, and t o breakthroughs of the currents through such barriers. It can thus be concluded that supraglacial crevasse deposits have no other specific characteristics than being commonly chaotic, built up of

Fig. 127. Deposits formed in a supraglacial crevasse formed during an early stage of deglaciation. The crevasse deposits (near Sobbtka, Lower Silesia) are 15 m high, 20 m wide and some 100 m long. Note the occurrence of several lithofacies types.


197

units with largely varying lithologies that rapidly interfinger and pass into each other, both laterally and vertically. Deformations are the rule rather than the exception (Bartkowski, 1976; Klatkowa, 1972).

Textural characteristics Crevasse deposits have no specific textural characteristics; there are no preferred clast orientations, nor are there distinctly orientated fault systems (apart from normal faults in the margins of kames). Imbrication may be present locally where meltwater currents have affected t h e sediments a t the base of the crevasse.

Occurrence Supraglacial crevasse deposits cannot be recognised easily on the basis of their lithofacies or textural characteristics; lateral facies transitions and the position within kames may, however, give sufficiently reliable indications. The position of these deposits, in turn, may provide palaeogeographical information. Baraniecka (1975) pointed out that supraglacial crevasse deposits ('kames' in her terminology) in central Poland occur in a linear system that was largely determined by the geothermal heat flux. The local differences in the amount of geothermal heat, due to faults in the hard-rock substratum, influenced the pattern of ablation: the crevasse deposits therefore follow fault systems in the substratum. Many of the supraglacial crevasse deposits follow the patterns of englacial or subglacial drainage systems (collapse above such systems may result in supraglacial crevasses). The crevasse deposits are indeed often found on top of englacial or subglacial meltwater-tunnel deposits. A similar relationship was established by Szupryczyfiski (1963) for modern conditions on Svalbard. The deposits are commonly found as elongated bodies embedded i n other supraglacial or terminoglacial sediments. In the glacial sequence, they are found on top of englacial sediments (or subglacial sediments) and are often covered by other supraglacial sediments. If the cover of supraglacial sediments has been eroded, as frequently occurs, the crevasse deposits are overlain by terminoglacial sediments.

Depositional mechanism The depositional mechanisms that contribute to supraglacial crevasse deposits comprise all sedimentary processes t h a t a r e active i n t h e

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supraglacial subenvironment. The most important of these a r e masstransport mechanisms, both subaerial and subaqueous ones. Since the water level in crevasses may change frequently and unpredictably, subaerial and subaqueous processes alternate in a n apparently random rhythm. Processes that result in grain-by-grain transport (currents, settling in water or from the air) are, as a rule, of much less importance under these conditions.

Supraglacial subaqueous mass-transport deposits (I-A-6-c) The supraglacial subenvironment in the ablation zone is characterised by numerous ponds, lakes and crevasses. Accumulation of sediment on the sedimentary surfaces just around or under the water bodies may easily lead to unstable situations, particularly as melting of the ice substratum may give rise to still steeper slopes. Subaqueous mass movements will result in such a case. A term commonly used for such deposits is 'flow till'. It should be emphasised, however, that this term was proposed by Hartshorn (1958) and elaborated by Boulton (1968) to describe supraglacial subaerial masstransport deposits which occasionally came to rest in small supraglacial water bodies. The term is now generally applied to both subaerial and subaqueous mass-transport deposits in the supraglacial subenvironment. The term should nevertheless be considered an unfortunate one, because the depositional mechanism is not related directly to melt-out, while the term 'till' suggests the existence of such a relationship. The sediments are also described in literature under the equally inappropriate term, 'waterlain till' or 'waterlaid till' (Dreimanis, 1969; Francis, 1975; Morawski, 1984). The same term is applied t o describe similar deposits formed in the proglacial and the terminoglacial subenvironments (Dreimanis, 1969; Evenson et al., 1977; May, 1977; Kurtz and Anderson, 1979; Broster and Hicock, 1985). It is remarkable that, while these deposits have been mentioned frequently by, among others, Boulton (1968), Drozdowski (1974, 1985), Olszewski (1974),N. Eyles et al. (1982b),Haldorsen (1982),Haldorsen and Shaw (1982), Kjysz and Lindner (1982), Lindner et al. (1982), Brodzikowski and Van Loon (1983, 1987), Brodzikowski (1984) and Morawski (1985, 1989), their sedimentological characteristics and analysis have been somewhat neglected. Relatively detailed descriptions and analyses of gravity flows in large supraglacial lakes in western Canada were given by N. Eyles (1987) and N. Eyles et al. (1987a) and discussed by Shaw (1988) (Fig. 128).


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Gilbert-type tributary delta with failure shoreline malor drainage stream

W w t u r b l d i t y underflow from major drainage stream subaqueous debris flo rbidity or grain-flow buried ice’

Fig. 128. Occurrence of supraglacial subaqueous mass movements in a supraglacial lake. From: J. Shaw (1988).Courtesy: Sedimentology.

Lit hofacies characteristics The lithofacies characteristics obviously depend on the material that was transported. Many of these deposits are diamicts (e.g. some types of slumped supraglacial stream deposits) but fine-grained sediments occur as well (e.g. turbidites in supraglacial lakes). Diamict-type deposits may be either clast- or matrix-supported. The sediments may be massive (N. Eyles et al., 1987), show p a r a l l e l lamination, be intricately folded or show fluidisation characteristics. The characteristics may change from place t o place within one layer, for example because a slump partly turns into a mudflow that itself partly changes into a coarse turbidite fine-grained in its distal part. Such lateral changes a r e commonly much better developed t h a n the changes i n subaerial counterparts. The deposits may consist of layers that have a fairly constant thickness over relatively large distances (e.g., in the case of turbidites), but a lobelike geometry is also common, particularly in the case of slumps and mudflows (Fig. 129). In such a case the internal lamination is most usually parallel t o the outer boundary of the lobe (Fig. 130). Grading and parallel lamination may be present in turbidites; ripple-drift crosslamination and convolutions may also be present in these deposits. The thickness of such deposits ranges from a few millimetres to approx. 5 m, but several units may occur on top of each other.

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Fig. 129. Supraglacial subaqueous slump of Wartanian age (WlosMw, Zary upland, western Poland). The slump is lobe-shaped and consists of fine-grained supraglacial fluvial and lacustrine deposits. The lobe shape has now become somewhat exaggerated because of load casting.

Textural characteristics One of the most common textural characteristics of these deposits is that the orientation of clasts and grains is preferably parallel to the outer boundaries of the deposit (Evenson et al., 1977). The depositional process leaves the larger particles unaffected, so that they do not change. Striated clasts will retain their striae; rounded particles will remain rounded and broken particles will not be rounded.

Occurrence Most of these deposits are to be found within the supraglacial deltaic and lacustrine facies, but individual deposits may also form part of crevasse deposits, They can be found - and, indeed, are commonly present wherever there were supraglacial water bodies. They thus commonly interfinger with supraglacial fluvial, deltaic and lacustrine sediments, which also form the under- and overlying deposits in the glacial sequence.


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Fig. 130. Characteristic appearance of what is commonly called a 'flow till', even though only few gravel-sized fragments are present. Slumping took place in a plastic state and the originally horizontal stratification was preserved, but was reshaped following the boundaries of the slump head. As a result of fluidisation, the sandy parts do not show internal lamination.

Depositional mechanism All subaqueous mass-transport mechanisms are involved in the process, which is not different from the processes that occur under non-glacigenic conditions. Massive unsorted layers with coarse clasts are commonly formed by cohesive debris flows (Middleton and Hampton, 1976). Massive sandy layers are, as a rule, due t o grain flows where the support mechanism was not only dispersive pressure created by grain collision (Lowe, 1976) but also some fluid turbulence (Hiscott and Middleton, 1979) and fluidisation. The finding of buoyant lift provided by a fine-grained matrix is a reason to interpret such massive sands as being the result of a modified grain flow (Lowe, 1982; Nemec et al., 1984). Graded beds commonly result from turbidity currents (cf. Bouma, 1962; Bouma and Brouwer, 1964). Lowe (1982) mentioned 'surging mass flows'. Inversely graded beds suggest deposition from a traction carpet (R.G.

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Walker, 1975; Lowe, 1982), where the support mechanism consisted of both dispersive pressure - with upward dispersion of the larger clasts from the zones with the highest shear stress (the base of the traction carpet) and kinetic sieving (sensu Middleton and Hampton, 1976) involving upward displacement of smaller particles into free spaces between larger clasts. Layers with internal deformations are most usually due t o slumping. The internal structure diappears when the slumping character changes into a mudflow character. Some lamination may remain but, classically, fluidisation structures will be formed or all original structures will be 'erased'. The rapid deposition of water-saturated layers with different density on top of each other may easily result in instability. A seemingly insignificant event (e.g. a partial melting of underlying ice) may then be sufficient t o trigger deformation processes such as loading (Fig. 131).

Fig. 131. Succession of three 'flow tills' of Drenthian age (Jaroszow zone, SW Poland). The middle deposit was less dense than the lower and the upper flow tills. Some trigger mechanism (possibly melting of ice underneath) induced the loading of material from the upper unit into the middle one, resulting in tear-shaped lenses t h a t could be termed pseudonodules.

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The englacial subenvironment

THE CONTINENTAL ENGLACIAL SUBENVIRONMENT (I-B) AND ITS DEPOSITS The ice body itself, including deposits and isolated clasts between the lower and upper boundary of the ice (Small and Gomez, 1981) but excluding the zones affected by supra- and subglacial processes, constitutes the englacial subenvironment of the glacial environment (see the Table on p. 127). Deposition may take place during both the active and the passive stage but occurs mainly if no ice movement is present; the depositional processes tend to be discontinuous and temporary. The conditions and resulting processes are fully dependent on the energetics of the ice sheet and consequently also on the hydrology in places where melting occurs. The hydrological conditions determine the rate of englacial melting, the drainage pattern, the erosive and transport capacity of the englacial water and the places where englacial sedimentation takes place (Fig. 132).

1 melt-out processes

I 1

I

meltwater-tunnel processes

nglacial transport

Fig. 132. Main factors influencing sedimentation in the englacial subenvironment.

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The englacial subenvironment is the most poorly known of all glacigenic subenvironments. Most of our insight into it is based on models that have not been truly checked under natural conditions because of the inaccessibility of this subenvironment (video and photo cameras in boreholes are of only limited help: Koerner et al., 1981). The models are based on the deposits found in glacial sequences, as far as they can be ascribed t o sedimentation i n the particular subenvironment, but most of these deposits lack characteristics sufficiently unambiguous that they can be interpreted as being of this type without any doubt. The lack of knowledge and insight in the depositional processes is probably the reason why not all sedimentologists consider the englacial subenvironment as a separate unit. Some workers include i t partly in the supraglacial and partly i n the subglacial subenvironments (Boulton, l968,1972b, 1975a; Embleton and King, 1975; Edwards, 1978; Dreimanis, 1980,1989; N. Eyles, 1983b; N. Eyles and Miall, 1986). One should realise that englacial sediments are indeed known from modern areas (Goodchild, 1875; Crosby, 1896; Okko, 1955; Jewtuchowicz, 1962; Szupryczydski, 1965, 1968; Boulton, 197213; Healy, 1975; Baranowski, 1977; N. Eyles et al., 1982b; Haldorsen and Shaw, 1982) but also from earlier glaciations as well (Drozdowski, 1974; Olszewski, 1974; Shaw, 1982; Brodzikowski and Van Loon, 1983,1987; Brodzikowski, 1984).However, englacial sediments are relatively scarce and their geological importance is very restricted. Three depositional facies can nevertheless be distinguished within this subenvironment: the englacial melting-ice facies (1-B-l), the englacial fluvial facies (I-B-2)and the englacial mass-transport facies (1-B-6).

ENGLACIAL CONDITIONS IN ACTIVE ICE The conditions prevailing in active ice are known - though far from completely so - from studies of recent ice caps and glaciers. Deposition is usually momentary and restricted to so-called 'temperate' conditions. 'Temperate' glaciers (a better term t h a n 'warm glaciers') have local temperatures not far below the melting point; the ductility of temperate ice favours the rapid formation, deformation and disappearance of fissures, tunnels and interstices within the ice body (Crosby, 1896; Gripp, 1929; Sugden and John, 1976; Baranowski, 1977; Shaw, 19821, so that the drainage and depositional patterns also change rapidly. Well developed englacial tunnels, crevasses and cavities (Fig. 133)may, however, survive for long periods (including the final stage of deglaciation), particularly if they become filled with mineral debris and if the internal movement

Englacial conditions in active ice

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Fig. 133. Schematic model of the drainage system on top of, inside, and under active ice. Abundant englacial crevasses and tunnels form direct connections with the supraglacial and subglacial subenvironments. From: Brodzikowski and Van Loon, 1987.

(deformation) of the ice takes place slowly. Such englacial debris-filled cavities are known from modern ice masses, among others on Svalbard and in Greenland, but the englacial character of apparently similar sediments in 'fossil' sequences was never proven beyond doubt, although there are several descriptions (Okko, 1955; Shilts, 1978; Brodzikowski and Van Loon, 1980,1983,1987; Haldorsen and Shaw, 1982; Shaw, 1982). Sedimentation within active ice depends mainly on a number of glaciological and climatological factors, particularly since these factors affect the character and geometry of the englacial drainage system and the depositional processes that take place in these systems. These relationships seem to imply that sedimentation from active ice may be possible but that the resulting deposits have almost no preservational potential. Large accumulations of englacial debris may also affect the ice flow (Rusell, 1895; Shaw, 1971,1977a).

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ENGLACIAL CONDITIONS IN PASSIVE ICE The englacial drainage system existing under passive conditions is mainly due to thermosubrosion. The system is commonly well developed with crevasses, fissures and tunnels which are not affected by ice movements. This implies that these englacial interstices will grow gradually, until collapse of the overlying ice destroys part of the system. The englacial drainage pattern has a n irregular shape, forcing the meltwater flows t o change their current velocity. This results in the frequent occurrence of places where transported debris, including the finest particles, can be deposited. Continuous sedimentation, especially in widenings of t h e englacial tunnels (Fig. 134), may thus result in the formation of rather large sedimentary bodies (Haldorsen and Shaw, 1982; Shaw, 1982).

Fig. 134. Schematic model of the spatial relationship between sediments of the englacial subenvironment. Based on N. Eyles et al. (1982, 1983) and Van Loon and Brodzikowski (1987).

Deposits of the englacial melting-ice facies

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Ongoing thermosubrosion is responsible for the collapse of overlying ice, but also for the disappearance of ice masses from between the tunnels. Sudden connection of different spaces of the drainage system t o each other tends to induce deformation of the material deposited earlier. This is seen best in distinctly layered deposits, but one should realise that the original deposits are also commonly characterised by sedimentary deformations (e.g., load casts). Floating pebbles i n a finer-grained matrix are also common and they can easily become the starting point for sedimentary deformations.

SEDIMENTATION PROCESSES IN THE ENGLACIAL SUBENVIRONMENT Melting of ice, flowage of meltwater and mass transport of material in crevasses are the main processes that contribute t o deposition i n the englacial subenvironment (Fig. 134). Englacial stagnant waters do not seem to remain intact long enough to allow formation of typical lacustrine deposits; it cannot be excluded, however, that such deposits are formed, but they have almost no preservational potential and their general characteristics are almost identical to those of the surrounding diamicts formed by rain-out from the roof of such cavities. The dynamics of the depositional process are largely determined by the amount of water available and by the energy of the water. Steep gradients i n water energy may occur, resulting in considerable variations in time and space with respect t o deposition. However, the average net sedimentation rate is generally much lower than that under supraglacial and periglacial conditions. DEPOSITS OF THE ENGLACIAL MELTING-ICE FACIES (I-B-1) The englacial subenvironment is not affected directly by either solar irradiation or the Earth's thermal heat flux. Melting may nevertheless occur, mainly as a result of thermosubrosion by englacial meltwater streams with temperatures above the freezing point. Some water is also formed along the flow lines of the ice as a result of friction between the glacial ice and mineral particles. The consequence of englacial melting is that englacial debris is set free and accumulates (under favourable conditions). This results in englacial melt-out complexes (I-B-1-a),the only deposit found in this facies.

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Englacial melt-out complexes (I-B-1-a) Melt-out tills had already been mentioned by several wo1,ers in the last century (among them Goodchild, 1875; Crosby, 1896; Garwood and Gregory, 1898) but the first author to detail the occurrence of englacial such tills was P.W. Harrison (1957). While his concept referred to buried dead-ice bodies, it is most likely that comparable deposits can be formed in active ice through thermosubrosion (Eissmann, 1981). Several authors have mentioned the finding of such deposits in presentday continental glaciers, thus confirming Harrison's hypothesis, although not all workers agreed with Harrison's (and our) view regarding the geological significance of englacial deposition by melt-out processes. The study of englacial melt-out tills has nevertheless been pursued actively as evidenced by the many reports of e.g., Olszewski (19741, RuszczyfiskaSzenajch and Lindner (1976), Shaw (1979,1982, 1983), Haldorsen (1982) and Stephan and Ehlers (1983). Two subtypes of englacial melt-out complexes can apparently be distinguished, viz. a n upper and a lower one (Olszewski, 1974; Drozdowski, 1979a,b). These two subtypes - if present - can be distinguished by their ratio criteria (frequency of joints, sandklay ratio, sometimes the boulder fraction); the different ratios are explained by the existence of different sources for the material and, to a lesser degree, by a slightly different depositional mechanism. The upper unit is generally finer-grained than its lower counterpart, it shows much more frequent vertical joints and is commonly more porous; the lower units, on the other hand, a r e more compacted; (sub)horizontal joints dominate the vertical ones. There is rarely a distinct boundary between the two subtypes and most workers state that the complexes have a uniform appearance, due to the specific depositional mechanism. However, non-uniform examples have been described by, among others, Lundqvist (1969a,b), Haldorsen and Shaw (1982) and Shaw (1982). Lithofacies characteristics These till complexes are predominantly sandy/silty/clayey b u t may contain scattered pebbles and boulders of up t o several metres i n diameter. Strong cohesion of the material is a general characteristic. An - often vague - stratification or lamination may be present due to differences in main grain size; the lamination is irregular. If layers with a relatively high concentration of pebbles are present, the immediately adjacent layers tend to be poorer than average in coarse material. The


209

Fig. 135. Massive englacial till of Drenthian age from the Kleszcz6w graben area (central Poland). Gravel is scarce. Note the irregular subvertical joints. Photograph:R. Gotowa€a.

thickness of the layers commonly varies from a few centimetres t o almost a metre but may occasionally be much greater. Lenses and pockets with material of a rather uniform grain size are common. Sandy pockets often show current ripples; finer material may show similar structures, but these are often difficult t o observe in the field. Parallel lamination is also quite common. The primary sedimentary structures are commonly deformed, which is t o be expected because of the englacial site of genesis. Englacial melt-out tills, as decribed from the Pleistocene glaciations in Europe, tend t o be homogeneous, particularly on casual inspection (Fig. 135) but small irregularities (sticking, boulder-induced scours, stratification, etc.) may be identified on detailed analysis. Joints are found quite frequently (Fig. 136), both in vertical and (sub)horizontal directions (tunnels and crevasses often follow such planes). The vertical joints may occur in various directions, thus forming prisms, and are usually considered to be the result of early diagenetic processes (Wysokiiiski, 1967; Olszewski, 1974; Drozdowski, 1979). The horizontal joints may show a frequency gradient in the upward direction;

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Fig. 136. Several joint systems in a n englacial melt-out complex of Wartanian age from the Kleszczow graben area (central Poland).

they are considered t o result from both depositional processes a n d compaction. It appears that vertical joints occur particularly in the upper englacial melt-out complexes, whereas horizontal joints are more common in the lower subtype (Olszewski, 1974). So-called draping structures (Fig. 137) may be found around boulders that were gradually set free from the ice mass by melting (Shaw, 1979). The height differences of the draped layer are a n indication of the thickness of the ice that had melted (Shaw, 1983); it was calculated from such data that some ice masses originally contained clasts making up as much as 46%of their volume.

Textural characteristics The textural characteristics of the particles depend on their origin and on the main transport mechanism(s). It appears that bullet-shaped clasts are relatively common and that some elongated clasts have a 'keel'. Glacial striae are common (Dreimanis, 1989). A conspicuous feature is the preferred orientation of flattened clasts. The ah-planes are horizontal or show imbrication (dip 'upstream'). The a-


--

-.._._ ;.-.Ice layor ._._^

.:

__

0-0

’ llll

-

-

Sorlod sedimonl

211

Fig. 137. Draping around large clasts a s a result of melt-out processes. From: Shaw (1977).

>?&%Debris-rich ice Large clast axes of elongated pebbles have a preferred orientation parallel t o the current direction. The scattering that is found with respect t o this orientation can usually be ascribed to irregularities in the substratum or t o post-depositional rearrangement, possibly through ice movement. Fluidisation of the host material, another process that may destroy the original orientation of pebbles, is especially common if the material has a relatively high silt content. Occurrence Englacial tills form layers with thicknesses ranging from less than a metre to over ten metres. Variations in thickness are limited, at least at exposure scale (quarries),but the appearance may change rapidly in both

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a vertical and a lateral direction. Such variations are due t o local conditions during deposition. These tills are not a characteristic part of the glacial sequence. It is likely that englacial tills do not exist in all ice masses, and if they do form, are easily destroyed. This implies that there may be more or less complete sections in the field that lack this type of deposit. If present, it should be found between subglacial and supraglacial deposits. The contact between englacial and supraglacial tills is not always clearly visible, although it may be marked by subhorizontal shearing.

Depositional mechanism The genesis of these tills cannot be studied under natural conditions. Experiments are extremely difficult to perform under conditions that resemble the actual situation in the field. The description of the presumed depositional mechanism is therefore based on hypotheses themselves based on detailed field observations and general glaciological knowledge. It is now generally accepted that englacial tills are formed because of slow melting of an ice mass. The most common situation is melting of buried dead-ice, where the difference between the ice temperature and that of the surrounding sediment induces a heat flux from the latter towards the former. The other situation occurs in an ice body where penetrating waters caused the formation of englacial crevasses and tunnels, due t o thermosubrosion. The continuous supply of supraglacial water, warmed by solar irradiation, will result in heat transfer t o the ice in the englacial subenvironment, causing slow melting of the ice walls around the crevasses and tunnels. The embedded clasts are set free, fall and accumulate thus forming tills, or are transported by the englacial streams. The frequent orientation of the longest axes of elongated clasts may result from this transport, but the orientation may also be caused by 'postdepositional' rearrangement under the influence of shearing, as expressed by the occurrence of 'filation' (Fig. 138). The irregular depositional substratum tends to result in concentration of the clasts in depressions, although continuous layers are fairly common. Water percolation may play an important role. Detailed analyses of the depositional mechanism were discussed by Shaw (1971,1977a7c,1982) and Dreimanis (1989).

Deposits of the englacial fluvial facies

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Fig. 138. Foliation in the lower unit of an englacial till from the Jaroszow zone (SW Poland).

DEPOSITS OF THE ENGLACIAL FLUVIAL FACIES (I-B-2) Englacial meltwater-tunnel deposits (I-B-2-b)are the only type of deposit discernable in this facies. The englacial subenvironment contains spaces that are commonly elongated. These spaces are called 'englacial tunnels' provided their general dip is less than 45"; those with steeper dips are termed 'englacial crevasses'. The tunnels generally are the natural prolongation of crevasses and, if they extend far enough, end at the base of the ice where they become subglacial tunnels. These tunnels commonly contain meltwater carrying and depositing debris. The processes in these tunnels may be quite different from those occurring under subaerial conditions because the water may be transported under high pressure (Fig. 139). Such 'pipe flows' are usually characterised by a high energy level, so that much larger clasts may be transported in this case than in subaerial currents (Durand, 1951,1953; Loadwick, 1970; Saunderson, 1977b).

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Fig. 139. Small supraglacial spring (diameter 2 m; height approx. 0.5 m) on the Werenskiold glacier (Svalbard), formed as a result of breakthrough by an overpressurized englacial stream. The outflow lasted half an hour. Photograph: J. Cegla.

Englacial meltwater-tunnel deposits (I-B-2-b) Englacial meltwater-tunnel deposits are known from both modern ice sheets with a temperate thermal regime and from Pleistocene successions (Shaw, 1971, 1977a,c, 1982; Boulton, 1972b; Dreimanis, 1980,1989). The primary sedimentary structures have always been deformed, but it is commonly quite feasible t o reconstruct the original geometry.

Lithofacies characteristics The deposits described from modern glaciers (Haldorsen and Shaw, 1982) are predominantly coarse (gravelly) and form lenses i n the ice that are usually horizontally stratified or cross-bedded (Fig. 140). 'Fossil' equivalents are often somewhat less coarse, with sand-sized particles being the most common. Isolated clasts may occur within the stratified sands.


215

Fig. 140. Intercalations of stagnant glacier ice (Omnsbreen, Norway) and stratified englacial meltwater-tunnel deposits. From: Haldorsen and Shaw (1982). Courtesy: Boreas.

Successions with meltwater-tunnel deposits are usually fairly complex. There may occur alternations of homogeneous or horizontally stratified gravels or gravelly sands (eith clast- or matrix-supported), cross-bedded sands, and silts and clays with parallel lamination. Textural characteristics

The larger clasts in these deposits frequently show glacial striae. The clasts tend t o be rounded, but freshly broken clasts may also be present; such broken clasts are possibly partly 'inherited' from frost-weathered supraglacial material that was embedded in the ice after deposition in a supraglacial crevasse. Most clasts are flattened; t h e a/b-planes a r e commonly orientated parallel to the sedimentary surface (which need not be horizontal!) but imbrication occasionally occurs. There are no other specific textural characteristics (Shaw, 1977a,c, 1982; Minell, 1979; N. Eyles et al., 1982b; Dreimanis, 1989). Occurrence

Englacial meltwater-tunnel deposits are usually found surrounded by englacial till complexes. In longitudinal section, the latter form elongated bodies a few centimetres t o approximately one metre thick (Fig. 141). The thickness is rarely greater, because thermosubrosion can usually not cause sufficient ice to melt to form higher cavities. Many of the deposits

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Fig. 141. Longitudinal section through an englacial meltwater-tunnel deposit (light colour) of Elsterian age from the Jaroszo'w zone (SW Poland). The deposits are intercalated between englacial melt-out complexes.

have a tunnel-like shape, which implies that their cross-section is more or less rounded (Fig. 142), with a diameter of about the same size as the thickness of the unit. Wide and relatively thin deposits may also occur. The lower contact with englacial till is generally sharp as a result of the erosive capacity of the tunnel flows. The upper boundary is less sharp, due to processes such as rain-out from the roof of the tunnel. Most boundaries have been deformed. Depositional mechanism

The precise depositional mechanism depends on several factors, the most important being the position of the tunnel with respect to the relevant water level, the inclination of the tunnel, the continuity in the supply of water and debris, and the quantity of water and debris. A 'high' position of the tunnel, a gentle slope, and regular supply of limited quantities of water and debris result in deposits that are fairly well comparable with most deposits formed in subaerial channels. The


217

Fig. 142. Transverse section through a small meltwater-tunnel deposit (width approx. 10 cm) of Elsterian age, consisting of deformed sands (Jaroszow zone, SW Poland). The surrounding englacial melt-out complex consists of much finer material.

resulting sedimentary structures commonly indicate t h a t low-energy conditions prevailed. A much more characteristic situation occurs if the tunnel is situated below the water level. The tunnel is then completely filled with water and the current depends on overpressure exerted by the mass of water bodies in a higher position. This type of current is commonly called 'full-pipe flow' or 'tunnel flow' (Newitt et al., 1955; Acaroglu and Graf, 1968; McDonald and Vincent, 1972; Saundersson, 1975, 1977c) and so-called 'sliding bed' conditions prevail under such conditions. High-energy conditions may be present, particularly if the tunnel is steep, and the sediments often show a parallel lamination formed under the conditions of the upper flow regime. More details on tunnel flow will be presented i n t h e section on subglacial meltwater-tunnel deposits (chapter o n subglacial deposits), since tunnel flow is more common under subglacial conditions.

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DEPOSITS OF THE ENGLACIAL MASS-TRANSPORT FACIES (I-B-6) It is not unlikely that subaqueous mass transport takes place under englacial conditions. However, the process cannot be studied in the field and the resulting deposits (if formed) may not survive later processes; it is also very well possible that such deposits were not recognised by the researchers. Much more is known about the deposits that are formed englacially by a combination of transport by water and 'dry' falling in crevasses. Such englacial crevasse deposits (I-B-6-b) gradually pass into englacial meltwater-tunnel deposits; their most obvious difference is the average slope.

Englacial crevasse deposits (I-B-6-b) Crevasses do not occur only a t the ice surface but are also found within the ice (cf. Fig. 133). Whether the term 'crevasses' applies if there is a prevailing horizontal component is debatable, but in our opinion the term should apply to all open spaces embedded in the ice with a n overall inclination of over 45"(if the inclination were less one should call the open spaces 'englacial tunnels; see facies I-B-2). Once the ice has melted, it is impossible t o distinguish between more horizontal and more vertical crevasse infillings; both result in irregular bodies of sands and gravels, embedded within englacial melt-out tills. Englacial crevasses may occur throughout the ice, but their deposits are often found on top of subglacial tills; this suggests that the crevasses also occur, mainly during advanced deglaciation, in the lower part of the ice. Analysis of the microstructures observed in englacial crevasse deposits and their relations t o englacial tills allows a n environmental reconstruction of the crevasses. It is found that the crevasses can be formed in the upper part of a n ice sheet under conditions of static pressure below the plastic limit of the ice. The active plastic strain of the flowing ice increases in relation to depth. This results in small, only partially opened crevasses in the lower part of the active-ice body; the shape and position of these crevasses may change during their existence, since intracrystal dislocations (sliding mechanism of flowing ice) give rise t o specific confining pressure and dynamics (Fig. 143). Changeability is greater here than in more surficial crevasses, since the ice in the top part is almost passive. One specific result is that supraglacial crevasses may easily be filled with sediments, while englacial crevasses may be closed again before sedimentation ends or even starts. An increasing effect is caused by melting of the

Deposits of the englacial mass-transport facies

ice lhickness i 40 m

11 the sudicIaI crevasse is filled by supraglecial sediment the mechanical property of layer (A) may change dislinclly

total vel~clty

\

219

of the ice body

elasllc limit plaslic limit

substratum

I

total displacement

I

Fig. 143. Schematic representation of ice displacement in a temperate glacier. The position of the plastic limit within the ice body largely determines the partially active, partially passive and partially sliding behaviour of the ice.

ice: this process will take place more quickly in supraglacial t h a n in englacial conditions; the amount of debris that becomes available is more or less proportional to the volume of melted ice. The hydrodynamic regime of crevasses has been discussed by various authors (e.g., Nye, 1965; Dewart, 1966; Stenborg, 1968; Shreve, 1972). An idealized, completely static model of an ice sheet or glacier section (see Fig. 14.11 of Sugden and John, 1976) can be compared with the hydrogeological situation in a karst area (Shreve, 1972). The mechanism involved in each has, however, not yet been unraveled in detail, possibly because of complications concerning the permeability factors in a n ice body at the melting point; there is a 3-dimensional network of fissures around the single ice grains, as shown by Nye and Frank (1973). The analysis by Sugden and John (1976) unfortunately does not deal with the confining pressure which increases with depth, or with the changes in plasticity. Lithofacies characteristics These deposits usually present no specific lithological or structural characteristics. Their grain size may vary from silty clay t o sandy gravel,

220


depending on the type of material supplied. Internal structures are rarely present. Deformation of the deposits is the rule rather than the exception, but is usually visible only because of the deformed contacts with adjacent sediments (Fig. 144)or by the presence of vague flowage structures.

Textural characteristics Englacial crevasse deposits formed under active-ice conditions usually show a preferred orientation of flattened clasts with their ah-planes more or less horizontal, as a result of compression and shearing under the influence of the overlying, moving ice mass. Deposits formed during a deglaciation stage lack the above orientation of clasts. Most deposits formed under these conditions shows no preferred orientation of grains or clasts whatsoever, but there are reports of such deposits with clasts showing a slight preference for a-axes orientated in a more or less vertical position.

Fig. 144. Englacial crevasse deposit (light colour in the central upper p a r t of t h e photograph) within an englacial melt-out complex t h a t overlies stratified englacial meltwater-tunnel deposits. Jarosz6w zone, SW Poland.

Deposits of the englacial mass-transport facies

22 1

There are no other specific textural characteristics, but the glacigenic origin is well reflected by the common occurrence of striated clasts.

Occurrence Englacial crevasse deposits form bodies of irregular shape, often with a greater height than width. These bodies are commonly elongated, so that the overal geometry is lense-like, intrusion-like or pipe-like. The size varies from less than half a metre in each direction t o several metres in width and height and tens of metres in length. It may be difficult t o determine the outer boundaries of the deposit because of lateral changes in lithology, but the contact with englacial melt-out complexes is generally clear. There are several works on englacial crevasse deposits in active ice (Bartkowski, 1967; Flint, 1971; Embleton and King, 1975; Klimaszewski, 1976; Sugden and John, 19761, but no detailed descriptions of such deposits (existing reports deal with deposits the englacial crevasse origin of which is not without doubt). This seems to indicate that this type of deposit has so far received insufficient sedimentological attention.

Depositional mechanism The character of the sediments suggests that the infilling of englacial crevasses takes place in several stages, mostly through subaqueous mass movements but also through falling of debris without water; on the other hand, temporary changes in the configuration of the crevasse, resulting from differential ice movement, often cause the mass-transported deposits in the crevasse t o be mixed with sediments that might be considered as englacial meltwater-tunnnel deposits (thus deposited by water currents). The deposits are commonly surrounded by englacial tills, but may rest on subglacial tills and be covered by supraglacial tills. There may be lateral transitions into englacial meltwater-tunnel deposits.

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The subglacial subenvironment

223

THE CONTINENTAL SUBGLACIAL SUBENVIRONMENT (I-C)AND ITS DEPOSITS The contact zone between the glacial body and its substratum is known as the subglacial subenvironment (see the Table on p. 127). An important difference of this zone from the 'passive' supraglacial and englacial subenvironments is that many more new particles are produced within the subglacial subenvironment and that grains are much more frequently modified by crushing and grinding as a result of inter-particle and particlebed contacts (Boulton, 1975b, 1979; Sharp and Gomez, 1986); breccias are also formed (Menzies, 1990). This implies that stone counts made for source analysis (Visser et al., 1986) may be less reliable than similar counts made in material from the other glacial environments. The general conditions are determined mainly by the thermal regime of the ice, the nature of the substratum and the amount of debris present. The depositional conditions depend on the subglacial ablation and the velocity of the ice. While there have been a few very rare chances of making direct observations (Kamb and La Chapelle, 1964), relatively little is known about present-day subglacial conditions and even the range of subglacial sedimentation is the object of controversy; it would seem that no subglacial areas are fundamentally excluded from depositional processes, though according to some authors sedimentation takes place i n the ablation area only (downwards of the equilibrium line). It is known that sedimentation can occur during both glaciation and deglaciation phases. Subglacial conditions have been and are being much studied, not only by sedimentologists (Boulton, 1975a, 1979; Schluchter, 1979a; Evenson et al., 1983; Menzies, 1986; Dreimanis, 1989; Goldthwait and Matsch, 19891, but also by geomorphologists; the formation of drumlins, for instance, is of interest t o both disciplines (N. Eyles and Menzies, 1983; Menzies and Rose, 1987,1989; Habbe, 1988,1989; Menzies 198913). Five facies can be distinguished within this subenvironment: the subglacial melting-ice facies (1-C-l), the subglacial fluvial facies (I-C-2), the subglacial deltaic facies (I-C-3),the subglacial lacustrine facies (I-(2-4) and the subglacial mass-transport facies (1-C-6). The subglacial deltaic and the subglacial lacustrine facies are of minor importance and relatively badly known from Pleistocene and older deposts.

224


SUBGLACIAL CONDITIONS UNDER ACTIVE ICE The base of a n active ice mass may move over the substratum or may be frozen to it. In the first case, the ice may override the substratum smoothly but, more generally, will destroy it, breaking fragments off and transporting them away; a soft-sediment substratum can also be deformed easily. Fragments that have been taken away may be redeposited i n the 'shadow' of some obstacle but may also be transported over a thousand kilometres or more, especially if the fragments become incorporated into the central part of the ice mass via shear zones. Deposition, erosion and deformation may be equally active i n the subglacial zone. The question of which process will prevail is largely determined by the thermal regime a t the ice/substratum interface. The thermal conditions may be rather complex as they depend on the geology of the substratum and on yearly changes in the net energy balance. Several models of subglacial sedimentation under active ice have been developed. Subglacial erosion (Fig. 145)is always a n important aspect of such models. Both theoretical considerations and field observations amount and behaviour 01

I

subglacial

-..,. .."

. Fig. 145. Diagram showing the relationships between the main parameters influencing subglacial processes such as erosion. Based on Sugden and John (1976), Embleton and King (1977)and Embleton and Thornes (1979).

225

Subglacial conditions under active ice

indicate (Haldorsen, 1981) that abrasion and crushing of grains strongly influence the lithofacies characteristics of subglacial deposits, although other processes also play a role (Fig. 146). Mineral particles may be transported in the subglacial subenvironment in traction or in suspension; the transport mechanisms can be compared with those for bed load and suspended load in subaerial streams, but the resulting deposits differ somewhat (Boulton, 1975a; N. Eyles and Menzies, 19831, particularly because of the influence of the type of substratum and the thermal regime at the icehedrock interface (Fig. 147). It is quite possible that subglacial diamictic material will be concentrated locally under these conditions if one takes into account the debris that is available in the subglacial subenvironment, the behaviour of the composition of original glacial drift

I

and t r a n s p o r t A

I

+?--

removal of fines

crushinq

of fines

removal ............. crushing

!

.- - - - -

I

.r - - - - - -

deficiency of silt and clay

selective enrichment of

deficiency of feldspar and sheet silicates

quartz enrichment in the sand

1

1-

, subglacia! glaciofluvial sediments

relative quartz enrichment in the remaining material

I

quartz enrichment

;

crushing ......... .....~~. ..:

'

;,_.

abrasion

abrasion

.~~. .. ...

quartz enrichment

I

removal of fines

Fig. 146. Genetic development of subglacial sediments (left) and two alternative ways quartz is enriched in such sediments (right). Slightly modified after Haldorsen (1983b).


226

water flow

-4 0

0

p suspended load

0

bed load -.

0

C

o

O

P 0

2 cm; white areas: net deflation of > 2 cm. Slightly modified from: Good and Bryant (1985).

Deposits of the proglacial aeolian facies

443

AeoIian activity goes hand in hand with local redeposition, viz. due t o changing winds. Good and Bryant (1985) found that the deposits remained active all the time, although areas could be distinguished that have net accumulation and others that show net deflation (Fig. 271).

Proglacial loesses (11-B-5-e) Loesses are the most typical type of deposit of the extraglacial subenvironment (Jahn, 1950,1956; Dylik, 1954; Smalley, 1966; Cegta, 1972; Jersak, 1973; Smalley and Leach, 1978; Rdiycki, 1979) but they are also found in the proglacial subenvironment (Pbwe, 1951,1955,1975;Ugolini, 1966; French, 1976; Embleton and King, 1977; Washburn, 1979; Akerman, 1980; Baranowski and Ppkala, 1982; Szczypek, 1982; Kida, 1986). The most detailed descriptions of proglacial loesses are those of Pewe (1975) for Alaska and of Bryant (1982)for Svalbard (Fig. 272). More general data regarding silty deposits on Svalbard have been reported by Jahn (1961), Wojtanowicz (1972, 1976) and Kida (1976). Icelandic proglacial loesses have been described by Ashwell (1966, 1972), Bogacki (1970) and Pekala and Wojtanowicz (1987),those in Greenland by Nichols (1969) and those in the area of the Sondre Stromfjord by Bocher (1949). Descriptions of proglacial loesses and related silts provide no or little sedimentological data. Sedimentation rates of 0.2-2.0 mm per year have however been measured in Alaska by Pewe (1975) and rates of 1.0 mm per

Fig. 272. Distribution of proglacial loesses in the Lower Adventdalen (Svalbard). Slightly modified from: Bryant (1982).

444

The proglacial subenvironment

year were found by Fernald (1965). These values are consistent with measurements carried out at Svalbard by Czeppe (19661, Baranowski and Pekala (1982) and Szczypek (1982),who found quantities of niveo-aeolian deposits (with a high silt concentration) of 29-400 g m-2 per year. The wide spread of these values must be attributed to periodical changes in wind activity. The Alaskan loesses have been described by numerous authors, among them Spurr (1898),Eardley (1938),Tuck (1938),Taber (1943,1953,19581, Pew6 (1951,1955,1965a,b),Trainer (1961) and Hopkins (1963). Alaskan loesses that were reworked by fluvial, lacustrine, estuarine or marine processes have been described by several other authors.

Lithofacies characteristics All loesses have comparable lithofacies characteristics: they have a n exceptionally high silt content, resulting in steep cliffs when erosion takes place (man-made 'hollow roads'). Many of the loesses seem massive at first sight, but detailed analysis (or a favourable type of weathering and/or erosion in an outcrop) reveals fine laminations, usually of an undulatory character, parallel t o the sedimentary surface (Fig. 273). Individual laminae may be up t o several

Fig. 273. Probably proglacial loesses of Wartanian age, showing indications of temporary fluvial influence. The lamination is well visible as a result of differential erosion resulting from minor differences in grain size. Outcrop near Trzebnica (SW Poland).

Deposits of the proglacial aeolian facies

445

millimetres thick, but are commonly less than 1 mm thick. Sand admixtures may be seen at the base of some laminae or as separate laminae. Pew6 mentioned the occurrence of proglacial loesses with a fine, locally discontinuous but always distinct horizontal lamination on sedimentary surfaces that are usually slightly inclined. He found abundant smallscale metadepositional deformations, partly a result of cryoturbation. The extraglacial subenvironment provides a wealth of further detail on the lithofacies characteristics of loesses.

Textural characteristics Characteristic loesses such as those described by Pewe (1975) contain 8090% of particles in the 0.005-0.5 mm range. It appears that the average particle size decreases with increasing distance from the source. This helps with reconstruction of the source of Pleistocene loesses. The relationship between source area and grain-size gradients is confirmed by conditions in Alaska, where the alluvial plains form the main source areas (Davidson et al., 1959; Trainer, 1961; Pewe and Holmes, 1964). However, Bryan 11982) found much larger variations in grain size, viz. from 0.25 to 9.0 (phi-scale)in Svalbard, variations not very different from the variability mentioned by Portmann (1969) for comparable deposits. The mineral and chemical composition reflects differences in the composition of the substratum in the source area. As proglacial loesses are generally deposited not very far from the source area, there is usually a much more distinct relationship between proglacial loesses and the substratum on which they rest then between extraglacial loesses and their substratum. Textural studies by Dement (1962)have indicated that the clay content is relatively stable. Clayey strikes may be present that cannot be ascribed to soil processes (which often affect loesses). The carbonate content, while it varies, may be considerable regionally. It is found that abrasion of grains increases with distance from the source. Sorting shows the same tendency, but proglacial loesses may contain sandy intercalations even in the most distal parts (Fig. 274). Sanning-electron microscope studies have shown that many of the grains are covered by upturned plates and that a number of fine particles adhere to a 'core' (Cegla et al., 1971). Analysing the fabric in thin sections, Bryant (1982) found that the grains show a distinctly preferred orientation. He also found that iron is present as hematite coatings on quartz grains rather than as granular fragments. This author also found abundant organic material in the thin sections, mainly as plant roots.

446


Fig. 274. Possibly proglacial loesses in a sandpit near Opole (South Poland). Note the irregular lamination and the patches offine sand. The structures suggest t h a t deposition of the wind-blown material took place in very shallow (1-5 cm) pools, possibly of a temporary nature (e.g., after a rainfall). Photograph: J. Cegh.

Occurrence Pew6 (1975) found that loess covers in Alaska are thickest near streams that drain glaciated areas, and that loess is currently deposited at the highest sedimentation rate near streams in the outwash plains. He found loess as high as 760 m above sea level, but occurring most commonly below 450 m. Not only is deposition at these higher levels quantitatively limited, but loess settled on hill tops and slopes is also easily removed again by surficial drainage. It is thus transported and deposited in the valleys where massive silt units accumulate, sometimes with large amounts of organic material. The thickness of proglacial loesses can reach 95 m in valleys (where reworked loesses are included), but PBwe encountered no successions of primary loess thicker than 61 m to the north of the Tanana River. He found that most deposits along the many rivers valleys are 3-12m thick, but thinner layers were found covering dunes.

Deposits of the proglacial mass-transport facies

447

The basal contacts of loess deposits are usually sharp, unless they form the natural, vertical, continuation of extraglacial loesses (phase of ice advance). The same is true for the upper boundary. There are also lateral contacts both with extraglacial loesses and with other proglacial aeolian deposits, and also with glaciofluvial and glaciolacustrine deposits. The position in the glacigenic sequence is determined by the relationships just mentioned.

Depositional mechanisms The non-reworked loesses are formed by settling of (mainly silt-sized) particles from the air during periods of decreasing wind velocity. Most reports state t h a t massive loesses are formed in this way and t h a t laminated deposits are a result of redeposition, e.g., by surficial run-off. The present authors consider it very likely, however, that lamination can also be formed (and preserved) if silt-sized particles that have already settled are blown in pulses over the surface. Bryant (1982) described laminated loess covers that he interpreted as the result of interaction between settling from the air and fluvial activity. The abundant presence of micro-scale current ripples in any case proves that water currents have affected many (if not almost all) proglacial loess deposits. There are, however, few sedimentological analyses available touching these deposits. Much more is known about the deposition of loesses in the extraglacial subenvironment.

DEPOSITS OF THE PROGLACIAL MASS-TRANSPORT FACIES (11-B-6) Locally, the proglacial subenvironment has an irregular relief under both subaerial and subaqueous conditions, and this facilitates the occurrence of mass-transport processes (Fig. 275). These processes are controlled by gravity, which implies that their frequency and quantitative importance increase with inclination and with the length of slopes. Two groups of sediments result from these processes: proglacial subaerial mass-transport deposits (II-B-6-a)and proglacial subaqueous masstransport deposits (II-B-6-c). The former group is represented by a wide variety of deposits, including debris and mud flows on proglacial fans, slumping of permafrosted valley banks, landslides, slumps, etc. The m-ouo of subanueous deoosits is connected orimarilv with the lacustrine

448


Fig. 275. Irregular solifluction lobe in the proglacial subenvironment of Svalbard (Hornsund area). Photograph: J. Cegta.

and deltaic facies, where slope instabilities or other situations trigger the reworking of previously deposited material in the form of high- and lowviscosity flows. Both the subaerial and the subaqueous mass-transport processes have long been known in the proglacial subenvironment and descriptions abound (e.g., Charlesworth, 1957; Beatty, 1974; Boothroyd and Ashley, 1975; Carter, 1975; N. Eyles, 1983; Miall, 1983).

Proglacial subaerial mass-transport deposits (II-B-6-a) Geomorphologists rather than sedimentologists have studied these deposits and the processes involved. The field experiments designed to unravel the relief-forming capacities of these processes were also carried out primarily by geomorphologists. Several classifications have been proposed for these deposits. This diversity is not surprising as there occur many gradual transitions between series of deposits with distinctly different end-members. In addition, more or less comparable topographic forms may result from


449

different processes and vice versa. The most commonly applied classification in this field is followed here. A first important group is represented by debris flows and mudflows. These frequently occur on proglacial fans, particularly in the upper and middle zones where there are water currents with a high concentration of glacial debris. Water-saturated layers of debris on these slopes may easily be triggered t o move 'en masse' along the slope. The resulting deposits have been described by Blissenbach (1954), Augustinus and Riezebos (1971), Boothroyd (1972,1976),Bull (1972,1977),Church (1972), Klimek (1972), Beatty (1974), Boothroyd and Ashley (19751, Rust (1975, 1977, 1978), Miall (1977, 1978), Ruegg (1977), Boothroyd and Nummedal (1978), Erikson and Vos (1979), Nemec and Steel (1984, 19881, Nemec et al. (1984), Drewry (1986) and Brodzikowski and Van Loon (1987) among others. The second important group comprises large-scale masses of material transported in the form of slumps or landslides. The deposits in this group are better known from extraglacial conditions, but have also been described several times from the proglacial subenvironment (Embleton and King, 1975; Jahn, 1975; French, 1976; Klimaszewski, 1976; Sugden and John, 1976; N. Eyles, 1983; Drewry, 1986; M.J. Clark, 1988). Quantitatively, this group appears to be the most important. The third group of deposits is a result of fast surficial mass-transport processes connected with the presence of a permafrost.The most common processes involved are skinflow (slope failure of the active layer), mudflow, local debris flow (e.g., due t o heavy rainfall), blockfall, rockfall, multiple retrogressive flow or slide, slumping and gliding of the active layer (Lewkowicz, 1988). The last group is represented by slow surficial mass-transport mechanisms, also connected with the presence of permafrost in the substratum. Creep of the permafrost, frost creep, needle-ice creep and gelifluction are the most important processes and often result in very specific deposits and forms (Jahn, 1975; Washburn, 1979; Lewkowicz, 1988). The last three categories are best known from the extraglacial subenvironment and will therefore be discussed here only briefly. More attention will be devoted t o the first category, which is of more or less equal importance in the proglacial and the extraglacial subenvironments. Lithofacies characteristics

Many of these deposits have a diamict-like grain-size distribution, but - in contrast to most tills - scarce or absent boulder-sized material (Fig. 276).

450


Fig. 276. Deposit resulting from a liquefied debris flow (Hornsund area, SW Svalbard). Photograph: J. Bierofiski.

The deposits are clast- or matrix-supported, depending on the relative proportion of fines available. The deposits in the distal part of this subenvironment are therefore more frequently matrix-supported than those in the proximal part. The grain-size distribution within one deposit may vary from place t o place. Differential movements during transport can result in elongated zones with a relatively large amount of fine or coarse material. In addition, some processes result in vertical grading, either normal or reversed. The situation with regard to large-scale slumps and landslides (Fig. 277) is much more complex as large masses of already varied material are usually involved. The original lithological characteristics seen in such a case not be greatly altered. Parts of the moving mass will, however, break loose and the bottom parts will be sheared. This affects the general characteristics and the 'head' of a landslide may thus be coarser than the tail, depending on the influence of the water content of the moving mass. Most of the fast and of the slow surficial mass-transport processes tend to result in stratified deposits, with thin laminae or crude layering several centimetres thick. Flowage structures tend to be well developed (Fig. 278). Surface slopewash produces distinctly laminated deposits.


45 1

Fig. 277. Slump scarp in the Hornsund area (Svalbard). The top of the permafrost acted as a (non-affected) gliding plane. Photograph: J. Cegfa.

Fig. 278. Distinct flowage lobes in a very fine-grained solifluction lobe in the proglacial zone of Svalbard (Hornsund area). Photograph: J. Cegfa.

452


Textural characteristics Because of the wide variety of processes involved and the different textural characteristics of the parent material, these mass-transported deposits show no diagnostic textural characteristics. Sorting is usually bad, but this is certainly not an exceptional feature of sediments formed under glacigenic conditions. Depending on the type of transport, the larger clasts show random orientation, or have their ah-planes preferentially parallel to the flow lines. Slow transport processes such as creep and solifluction may produce an orientation of flat particles with their ah-planes subparallel t o the bounding surfaces. The grain surfaces show no specific textural characteristics as they are not altered during the mass-transport processes. Occurrence These deposits can be found throughout the proglacial subenvironment, except in the presence of surficial waters. They may occur on truly horizontal surfaces, but the source areas must, of course, be inclined (a 2" slope suffices for gelifluctionj; it is therefore not surprising that these deposits most frequently appear at the transitions from inclined surfaces (ice-pushed ridges, ice-cored moraines, etc.) to more horizontal areas (e.g., alluvial plains). The deposits fill depressions in the sedimentary surface, particularly if the material transported had a high water content, and thus contribute to levelling of the relief. Large slumps and landslides may produce isolated 'heaps' of masstransported material, but small-scale deposits may be so frequent that these deposits can cover a large area. The deposits interfinger with all other types of proglacial deposits, but do so especially with aeolian deposits and with fluvial deposits a t river terraces (cf. Jahn, 1975; Washburn, 1979). The areal distribution of these deposits as just described implies that they can be found throughout the proglacial interval of the glacigenic sequence, often occurring as local 'interruptions' of deposits from other facies. Depositional mechanisms Subaerial mass flows have been elaborately discussed by Lawson (1979b, 1981b), primarily in the case of the supraglacial subenvironment (the


453

subenvironment is however of no importance for the transport mechanism). Lawson also distinguished the four categories of subaerial masstransport deposits that were mentioned above. Sediment resting on an inclined surface will start mass movement when the forces applied exceed the internal strength of the material. This may occur, for instance, if rain infiltrates the debris layer and reduces cohesion or the number of intergranular contacts. The internal porewater pressure during water infiltration may then reach values that result in the start of fluidisation, especially if the material contains fine particles that cause local retention. If the threshold value has been passed and mass movement has started, the velocity will increase as long as the inclination of the slope does not diminish. Mass flows with a very high density and a relatively low pore-water coctent will usually move rather slowly, while material saturated with water will move much faster. The amount of pore water thus plays an important role but is not decisive; several other parameters, e.g., mean grain size, relative proportion of fines, presence of a n impermeable subsoil, thickness of the moving layer, shear strength, porosity, bulk density, dynamic viscosity also play a role. The slow types of surficial mass transport are controlled by seasonal melting of the top part (active layer) of the permafrost and by the water content, the latter parameter fluctuating especially during warm periods of the year. Temperature changes as a result of the dayhight cycle also seem t o affect the susceptibility of the surface layers to mass transport. Creep of the permafrost and gelifluction (the term is used to indicate solifluction under glacigenic conditions) are the best known of these processes.

Proglacial subaqueous mass-transportdeposits (11-B-6-c) Some subaqueous mass-transport processes in the proglacial subenvironment do take place in rivers but the greater majority occur in lakes. Deltaic foresets are preferred, but the slopes of the lake itself are also frequently affected. Mass-transport processes occurring under these conditions are exceedingly well comparable t o the equivalent processes operating in the terminoglacial subenvironment, with the exception that no direct influxes from the glacial environment take place. The terminoglacial subenvironment being defined on the basis of mass movements starting in the glacial environment, this implies that the frequency of occurrence of these deposits is much greater in the terminoglacial than in the proglacial subenvironment.

454


Two main groups of subaqueous mass-transport deposits can be distinguished in this subenvironment: those formed as a result of slumping (which may, but need not, change into a high-density turbidity current and, occasionally, later into a low-density turbidity current), and those formed directly as a result of - usually small-scale - low-density turbidity currents triggered by local circumstances. These two groups of deposits include a wide variety of processes and resulting deposits. The authors found that many such different types, among others deposits from subaqueous debris flows, large-scale grainflow, fluidised flow, turbidity currents and fluxoturbidites, occur all together within deposits formed in proglacial lakes in the Kleszczow graben (central Poland).

Fig. 279.Proglacial subaqueous mudflow deposits between deltaic bottomsets of Elsterian age in the Kleszczow graben (Poland).


455

Lithofacies characteristics Most of these deposits consist of predominantly fine-grained material, with some larger floating particles (Fig. 279). These larger particles may be small stones (e.g., in fluxoturbidites and slumps) but more commonly are intrabasinal fragments of unlithified deposits. The reworked deposits then have a diamict-like appearance. There are also deposits that consist exclusively of particles no larger than sand. Several turbidites are good examples of this and the most distal parts even contain only fine sand, silt and clay. There is no need to discuss the structures of turbidites here, but the normal grading is again a n important feature. Proglacial turbidites are often incomplete (Shaw , 197713; Shaw and Archer, 1978, 1979; Shaw et al., 1978) and tracing turbidites from one outcrop t o another has revealed that deposits may eventually be left with grading as the only characteristic. Such graded, fine-grained layers cannot be distinguished from the true, seasoninduced, varvites. Other deposits, on the contrary, are relatively coarse and contain a number of clasts. They are formed as a result of transport in high-density suspension clouds (these may be triggered by, for instance, earthquakes) and often have a massive appearance. Massive sandy intercalations may also be present in other sediments and usually represent fluidised flows or grainflows; they occur in the proximal part of lakes in front of a delta rather than in the distal parts. Details touching the lithofacies characteristics of proglacial subaqueous mass-transport deposits have been provided by numerous authors (Banerjee, 1966,1973; R. Gilbert, 1971,1975; Ashley, 1972,1975; Gustavson et al., 1975; Shaw, 197713, 1988a; Shaw and Archer, 1978, 1979; Shaw et al., 1978; Sturm and Matter, 1978; Sturm, 1979; Brodzikowski and Van Loon, 1980, 1983, 1985c, 1987; Hahszczak, 1980, 1982; Gilbert and Shaw, 1981; Brodzikowski, 1982, 1984; N . Eyles, 1983b; Quigley, 1983; N. Eyles et al., 1987a, 1988a). Textural characteristics The characteristics are very similar to those of the corresponding deposits in the terminoglacial subenvironment, although the average grain size is somewhat smaller. Boulders are usually absent (they can be found if frost-weathered blocks supplied by, e.g., subaerial mass movements reach a lake: Fig. 280); even pebble-sized material is rare. This implies that sorting is usually somewhat better here than in the terminoglacial sub-

456


Fig. 280. Subaerial debris flow on Svalbard (Hornsund area), which has deposited part of its material on the shore around a proglacial lake. Part of the diamictic flow slid down into the lake, forming a pebble-rich subaqueous mass-flow deposit. Photograph: J. Cegta.

environment, but this generalisation does not necessarily hold for individual deposits. Relatively light fragments, e.g., reworked parts of more or less consolidated unlithified sediments may, however, be present. The flattened soft-sediment clasts are orientated either subparallel t o the sedimentary surface or parallel to the flow lines (in the case of a highdensity flow). There may be imbrication of the pebbles at the base of these deposits if sufficient clasts are present (cf. Van Loon, 1970), and orientation may be more or less horizontal orientation in the case of concentrations of clasts at the top of a unit (reversed density gradient as a result of high-energy turbulent flow). As hard-rock glacial material of pebble size is absent, glacial striae proving transport by ice are also usually absent. There will be no glacial fragments whatsoever if the deposits were formed in front of a n advancing ice cap.


457

Occurrence The deposits are found almost exclusively in proglacial lakes, principally in front of deltas. They form distinct intercalations in the lake sediments, usually characterised by an appearance coarser than that of the deposits where they are embedded. The thickness of individual beds railges from less than a centimetre to half a metre (thicker layers are exceptions), but several layers may occur on top of each other (or with very thin units of autochthonic sediments in between) so that successions of several metres thickness can be found. Thick units are usually a result of plastic flowage, resulting i n a relatively limited areal extent (e.g., slumps), whereas turbidites may cover almost the entire surface area of the lake but be correspondingly thinner. The deposits are found in the glacigenic sequence intercalated between proglacial lacustrine and deltaic deposits, and much more rarely between fluvial deposits. We observed that the deposits can constitute up to 20% of proglacial deltaic and lake-margin deposits, and up to 50% of proglacial lacustrine bottomsets; this difference is the result of a much smaller supply of 'regular' lacustrine material rather than a larger supply of mass-transported material in the distal zones.

Depositional mechanisms The main mechanisms involved were described above in the corresponding section of the chapter on terminoglacial deposits. The main difference between the deposits from the two subenvironments results from the fact t h a t the proglacial subenvironment generally has a somewhat less pronounced relief, so that the resulting flows have, on the average, a somewhat lower energy. Low-density turbidity currents are therefore somewhat more common in the proglacial than in the terminoglacial subenvironment. According t o Postma (19861, both cohesive and non-cohesive turbulent low-density flows can produce a sequence with Bouma intervals, reflecting the general decrease in current velocity: the low strength of the cohesive flows (higher strength values will decrease turbulence, thus changing the flow from turbulent into laminar: Enos, 1977) allows sand grains t o settle before clay particles. The resulting deposit will thus resemble the Bouma sequence until the tangential shear stress has reached the yield strength of the plastic flow; a massive mixture of fine sand, silt and clay then ends the distribution grading. Flocculation of clay particles in muddy suspensions may cause pronounced segregation of clay

458


and silt particles, and results in well developed laminae (Stow and Bowen, 1980) in the upper part of the Bouma sequence. The currents stop should the supply of the waterlsediment mixture become insufficient, e.g., because a decrease in inclination of the slope stops autosuspension. Numerous micro-scale turbidites therefore have only a very restricted extent.

The extraglacial subenvironment

459

THE CONTINENTAL EXTRAGLACIAL SUBENVIRONMENT (11-C) AND ITS DEPOSITS This subenvironment (p. 128) is the most external part of the periglacial environment. It starts in front of the terminoglacial or proglacial subenvironment and ends where a permafrost is no longer present (cf. Washurn, 1973; Jahn, 1975; French, 1976; Embleton and King, 1977; M.J. Clark, 1988) (Fig. 281). The problem is, of course, that it is not always possible t o find out whether there has been a permafrost in some specific place at a specific time. The occurrence of 'periglacial' structures (e.g., cryoturbation) is therefore of great help if the former transition line between the extraglacial subenvironment and the non-glacigenic, adjoining environment is to be reconstructed. The continental extraglacial subenvironment, which is absent in the Southern hemisphere because the potentially extraglacial zone is covered by ocean waters, forms a wide belt, regionally more than 2500 km wide, in

-~ v

~

pertglacial environment 0)

Fig. 281. Extent of the extraglacial subenvironment and generalised depositional sites.

460


Hochpolaiei F i O S l S C h ~ l t ~ e i i g i s i i d r

Fig. 282. The periglacial subenvironment on the Northern hemisphere. Slightly modified from: Karte (1979).

the Northern hemisphere. This vast region can be subdivided into several climatic and vegetational subzones (Fig. 2821, which also differ a s t o intensity and nature of the prevailing cryogenic processes; these processes and the structures and landforms (Fig. 283) resulting from them have been studied in detail since the important practical applications of geocryology have become obvious. The most important geological and geomorphological distinction that can be made is between zones with continuous and discontinuous permafrost. Permafrost is also present in areas that are not under the influence of a n ice cap, viz. at high elevations. While typically cryogenic processes and extraglacial deposits may be formed there (Karte, 1979),the extraglacial features described in the present book are restricted t o those of the belt surrounding - at some distance - land-ice caps and glaciers. SEDIMENTATION PROCESSES IN THE EXTRAGLACIAL SUBENVIRONMENT In view of the distance of this subenvironment from the related glacial ice mass, i t follows that only part of the material supplied is related with the glacierisation involved: most of the supply will be from non-glacigenic sources. This implies t h a t the same types of depositional agents a r e

Sedimentation processes in the extraglacial subenvironment

46 1

Fig. 283. Schematic block diagram showing the main structures and landforms under modern periglacial conditions. From: Karte (1979). Courtesy: Bochumer Geografische Arbeite.

present as in the proglacial subenvironment (Fig. 284), but that the glacigenic nature of rivers, deltas and lakes is usually so small that the resulting deposits cannot be considered predominantly glacigenic. There are therefore only two significant typical glacigenic facies in this subenvironment: the extraglacial aeolian facies (11-C-5) and the extraglacial mass-transport facies (11-C-6); the latter contains predominantly reworked aeolian material. Of course, there may occur fluvial, deltaic and lacustrine deposits with glacigenic characteristics, but they are rare and therefore not described here . The characteristics of such sediments are not essentially different from their proglacial counterparts. Although the characteristics of the extraglacial deposits have been studied for a long time, the underlying processes are still not yet fully understood. New insights, especially with respect t o coversands, have developed in the past decade (Ruegg, 1983b; Schwan, 1986,1987,1988a1, but the genesis of extraglacial dunes and loesses still needs detailed study. It is obvious, and completely understandable, that a general decrease of grain size is found in aeolian deposits in an off-ice direction (transition

462


Fig. 284. The extraglacial subenvironment in Greenland. Note the occurrence of streams and small lakes that are not, however, considered to contain extraglacial deposits. Deflation zones are the source of extraglacial aeolian deposits that, in turn, can be reworked to form extraglacial mass-transport deposits. Photograph: J. Ceda.

from coversands to loesses), but areas under marine influence also tend to have a relatively large share of dunes and coversands, whereas loesses seem to occur more frequently under continental conditions. Both the coversands and the loesses may cover extensive areas: the most famous of these is the fertile Saalian and Weichselian (and equivalent) loess belt from Western Europe (where the loesses have a thickness of some metres) via Central and Eastern Europe (where the thickness commonly reaches several tens of metres) to the eastern coast of China (where the loess may be several hundred metres thick locally: Tungsheng, 1988). The precise sedimentation processes of both coversands and loesses remain highly controversial. Recent field data are accumulating that seem to indicate that much of the material was originally laid down by winds but was more or less intensively reworked by surficial currents and/or mass-flow movements. In contrast, the undulating topography with a permafrosted soil must have given rise to a large number of shallow

Sedimentation processes in the extraglacial subenvironment

463

lakes where wave action may also have affected the sediments and where wind-supplied particles settled gradually according t o their size and shape. In spite of these non-aeolian influences, coversands and loesses tend to level the original relief. They may reach considerable thicknesses where the subsoil shows depressions like those in the original valleys. As fluvial, deltaic and lacustrine deposits are not considered here to form part of the extraglacial sediments, the mass-transport deposits in the extraglacial subenvironment are considered as glacigenic only in so far as they consist of subaerially reworked aeolian material. Reworking may be intensive in some places, so that considerable accumulations of reworked deposits can form. The active layer of the permafrost plays an important role, emphasising the glacigenic character of these mass-transport deposits. DEPOSITS OF THE EXTRAGLACIAL AEOLIAN FACIES (11-C-5) It follows from the descriptions in preceding sections that, while the glacigenic aeolian facies is not restricted to the extraglacial subenvironment (aeolian deposition is well known from other subenvironments), in a glacigenic context, it is most characteristic of extraglacial conditions. Much research has therefore been devoted to this facies and descriptions of the processes involved and of the resulting deposits thus abound. Five types of deposits can be distinguished in this facies. The most characteristic, and most widespread and studied type is that consisting of extraglacial loesses (II-C-5-e),which usually forms a belt surrounding the narrow zone of coversands (11-C-5-d). Extraglacial dunes (11-C-5-c) are found locally, where sufficient sand is available (e.g., along sea shores, but also in inland areas), and extraglacial drift sands (11-C-5-b) are formed where so little sand is supplied by the wind that no vast bodies can be formed. Extraglacial aeolian complexes (11-C-5-a) consist of a mixture of two or more of the types of deposits just mentioned.

Extraglacial aeolian complexes (11-C-5-a) These complexes are mainly a result of changing conditions, resulting in alternation of depositional processes and resulting landforms. Several complexes are composed of - partly eroded - dunes with intercalated units of cover sand and/or loess, with lateral transitions into material that can be best described as drift sands. The complexes are therefore not very specific, from either a sedimentological or a geomorphological point of

464


view and have so far received little attention. There are no studies devoted exclusively to these complexes and descriptions are only found as part of more general studies. No such complexes have been described from lithified rocks.

Lithofacies characteristics The mixed nature of these deposits accounts for the wide variation in lithology found in even a single complex. Parts that might be considered as remnants of extraglacial dunes can show large-scale cross-bedding, whereas coversands and loesses usually show irregular lamination, sometimes with horizons characterised by current-induced cross-bedding. Parts that might be interpreted as having been deposited in sand drifts show predominantly horizontal lamination, but may also show small-scale cross-bedding. Apparently massive zones (commonly a result of grain-size differences insufficient t o expose the internal structures) are found in most complexes (Fig. 285).

Fig. 285. Part of an extraglacial aeolian complex showing massive parts, sets with irregular horizontal lamination, and paleosoils. Odra river valley (west-central Poland). Photograph:B. Nowaczyk.

Deposits of the extraglacial aeolian facies

465

The nature of the complexes implies that there frequently are sharp boundaries between the various units, but there may also be gradual transitions (e.g., from fine coversand to coarse loess). This complex history, with erosional phases being common, is expressed in the presence of numerous erosional surfaces (including steeply inclined truncations) and in the development of paleosoils. The latter, however, are not usually very well developed because the exposed layers again became covered by freshly supplied wind-blown material.

Textural characteristics Extraglacial aeolian deposits usually underwent much more long-lasting aeolian transport than the corresponding proglacial deposits. The extraglacial aeolian complexes reflect this general rule, as some (specifically the coversand units) are made up of grains with a much better rounding; the parts that can be interpreted as remnants of dunes or drift sands, however, do not show such a pronounced difference. Another difference from proglacial aeolian complexes is that the loesskoversand ratio tends to be much higher in the extraglacial examples than in the terminoglacial complexes. Sorting of the complexes as a whole is usually bad, but improves greatly if only individual units are considered. The effect of surficial water currents may lessen the extent of sorting because sand and finer particles of different origins become mixed and settle simultaneously if depositional conditions are favourable. The surfaces of the individual grains show more aeolian characteristics than do the proglacial counterparts; this applies particularly to the grains forming part of the coversand and loess units.

Occurrence The complexes are found throughout the extraglacial subenvironment. It must be emphasised in this context that the low degree of vegetation, the abundance of wind-exposed sediments and the climatological conditions favour erosion, transport and deposition through wind activity. According to palaeogeographic analyses, however, the typical coversand and loess belts are very uniform, so that relatively few aeolian complexes will be found embedded there. The most favourable conditions are therefore found in the 'proximal' part of the extraglacial subenvironment, where zones with such complexes form the more or less natural continuation of similar zones in the terminoglacial or proglacial sub-

466


environments. This implies that the extraglacial aeolian complexes are found in the glacigenic sequence most frequently located i n the zone adjacent t o the aeolian proglacial deposits. The thickness of the complexes is restricted and depends on the precise definition applied. It does seem reasonable, however, to distinguish aeolian complexes only if the thickness of the individual units is much smaller than under well developed conditions. The complexes therefore should not comprise units of dune height. Thus, with this approach the thickness of the complexes is limited to a few (lo?) metres at most. The horizontal extent may be much greater, as changing conditions may affect large areas. However, field observations indicate that the complexes are usually recognised as such only if they contain dune remnants, so t h a t the complexes rarely exceed the width of 'normal' extraglacial dunes or, maximally, the joint width of a limited number of dunes (order of 100 m). The complexes interfinger with other aeolian deposits and, if the relief was favourable, also with extraglacial subaerial mass-transport deposits. Interfingering with non-glacigenic sediments (for instance, of fluvial nature) is found in addition. Depositional mechanisms

The genesis of these complexes does not differ from t h a t of proglacial aeolian complexes, which are in fact more characteristic. The reader seeking more detailed information on the depositional mechanisms will therefore be referred to the relevant section in the chapter on proglacial deposits.

Extraglacial drift sands (11-C-5-b) Extraglacial drift sands constitute such a small fraction of the extraglacial aeolian deposits and exhibit so few characteristic properties that they have received very little attention from sedimentologists. They are nevertheless fairly common, as only a small quantity of sand needs to be affected by wind t o form such a deposit. These sands abound in modern extraglacial zones, but are much more rarely recognised in Pleistocene deposits (there are no descriptions from older glaciations). The reason for this may be that drift sands are usually so thin that they are not recognised, or that they are not considered of enough importance to be described as typical intercalations in other deposits. In addition, pedogenesis may affect both these deposits and underlying deposits of different origin. The fact that all these deposits are

Deposits of the extraglacial aeoIian facies

467

affected may it make impossible to trace 'fossil' extraglacial drift sands in practice. Finally, the thin drift sands may easily be eroded by wind action before they are covered by younger sediments. Some Pleistocene drift sands have been identified from the textural characteristics of sand grains in buried frost edges (Gozdzik, 1970, 1973, 1976,1981,1983).

Lithofacies characteristics These deposits usually have a massive appearance, probably as the result of either secondary processes (such as pedogenesis) or insufficient differences in grain size. Irregular streaks (the term 'laminae' would not be justfied here) of coarser and finer material (Fig. 286), representing wind pulses of different velocity are found if outcrops are favourable, or on careful study of lacquer peels. Current observations, under modern extraglacial conditions, indicate that true horizontal lamination may occasionally be present at the base of the drift sands. This finding is explained by subcritical climbing of wind ripples.

Fig. 286. Extraglacial sand drifts with a vague, irregular stratification due to streaks of somewhat coarser grains. Photograph: J. Cepta.

468


Textural characteristics Sorting of these deposits is moderate to good, with grains rarely outside the 0-3 phi range, but with distinct preference for grains of the order of 2 phi. It appears that relatively coarse deposits show the best sorting; this can be explained by removal of the fine particles by wind activity. Many of the sand grains in glacigenic deposits have undergone several phases of aeolian transport. The mainly fluvial deposits that function as a source for the extraglacial drift sands therefore already contain many well rounded grains with an aeolian surface texture (Goidzik, 1980; GoLdzik and Mycielska-Dowgialto, 1988), so that short extraglacial aeolian transport does not necessarily imply a lack of aeolian texture. On the other hand, the extraglacial sand drifts also contain abundant grains that show almost no sign of aeolian transport. In addition to this, individual grains may have been broken after deposition, as a result of frost activity. Most of the grains analysed appear t o consist of quartz, with minor amounts of feldspar (Goidzik, 1973,1981). Occurrence The thermal, wind and moisture conditions of the extraglacial subenvironment render almost all deposits susceptible to wind erosion and thus t o become suitable sources of drift sands. Consequently, sand drifts are found throughout this subenvironment: on moraine-covered uplands, in extraglacial valleys, around dune fields, etc. The bodies are thin, usually with a thickness of less than one metre, often of less than a decimetre and occasionally of less than one centimetre. The thickness depends on the local topography and tends to level microdifferences in relief. This implies that these deposits can often be found as isolated patches - usually as concentrations of patches - in topographic depressions. It is well possible that the limited thickness of these deposits is due to the fact that somewhat thicker accumulations grow rapidly in a vertical sense, thus forming dunes. There are indeed several indications that extraglacial dunes have somewhat diverging deposits (drift sands?) at their base (Dylikowa, 1964, 1969; Galon, 1969; Kozarski et al., 1969; Gozdzik, 1973; Nowaczyk, 1986). The lack of fossilised drift sands suggests that the preservational potential of these deposits is very low, or that their recognition in fossil sediments is very difficult. Under modern conditions these deposits are often found in the direct vicinity of extraglacial dunes. This type of loca-


469

tion might serve as clue for the recognition of 'fossil' specimens, both their areal distribution and their position in the glacigenic sequence being factors.

Depositional mechanisms The depositional mechanisms involved are practically the same as those responsible for the accumulation of extraglacial coversands (the mechanism involved in this type of deposits will be described in more detail). Sand grains are taken up by the wind, mainly from areas with no vegetation and with a dry surface, and transported as a 'carpet' of grains rolling over the substratum and in saltation. A decrease in wind speed or the passage of a wet area will cause some of the grains to be left behind. Both adhesion and settling of saltating particles are thus important mechanisms. The combination of these two processes is responsible for the formation of climbing ripples, which may be either subcritical or supercritical. The deposits thus formed may become homogenised through post-depositional processes such as soil formation, growth of plant roots, percolation of rainwater, etc.

Extraglacial dunes (II-C-5-c) There are numerous extraglacial dunes (Fig. 287) in North America, Europe and Asia from the end of the Pleistocene, and in somewhat less abundant numbers in South America and New Zealand. Rozycki (1979)

Fig. 287. Extraglacial dunes, dating from the Pleniglacial, in the Wielkopolska Lowland (west-central Poland). Photograph: B. Nowaczyk.

470


stated that most dunes of this origin are formed if the ice front is about 500 km away, if there is little precipitation and if the climate changes from very cold t o somewhat more moderate (also see Dylikowa, 1958, 1969; Galon, 1958, 1969; Henderson, 1959a; Stankowski, 1963; H.T.U. Smith, 1965; Urbaniak, 1967; Gawlik, 1969; Kozarski et al., 1969; Wojtanowicz, 1969; Rotnicki, 1970; Rozycki, 1972; Perrin et al., 1974; Maarleveld, 1976; Krajewski, 1977; Washburn, 1979; Vandenberghe, 1981, 1983a, 1985; Nowaczyk, 1986). These conditions are, however, not essential ones as there are Pleistocene dunes in the extraglacial subenvironment of Alaska where they formed a t distances less than 200 km from the ice front (Carter and Robinson, 1978). Furthermore, modern extraglacial subenvironments in the Antarctic, Greenland, Svalbard, Iceland, Alaska and Arctic Canada also have dunes in areas less than 100 km away from the ice. Nevertheless, it remains striking that by far most of the Pleistocene dunes are found in locations distant from the ice front. It is also remarkable that many Pleistocene extraglacial dunes date from the period of 13,000-9,000 years BP. This suggests that the Pleniglacial represents an interval most favourable for dune formation, but similar conditions must also have occurred during, for instance, Weichselian (Wisconsinan, Valdaian, Devensian) times. While no buried extraglacial dunes are described in the literature, there are numerous reports of 'fossil' aeolian deposits of glacigenic, and probably extraglacial, nature with characteristics that make it likely that they originated as extraglacial dunes. Descriptions such as these concern Precambrian deposits of Canada (Ross, 1983a,b), South Africa (Meinster

Fig. 288. Sand quarry in a dune close to that shown in Figure 287. The dune sands are being exploited because they consist of relatively well sorted fine to medium sands. Photograph:B. Nowaczyk.


471

and Tickell, 1976) the North Atlantic region (Nystuen, 1985), Mali (Deynoux et al., 1989), Mauretania (Deynoux, 1982) and Scotland (N. Eyles and Clark, 1985), and Ordovician deposits in the Sahara (Deynoux, 1982). These data are, however, much less detailed than those available for Pleistocene extraglacial dunes of North America and Europe. Field observations from these continents indicate t h a t most extraglacial dunes had parabolic forms, with numerous transitions into (or from) longitudinal dunes. Isolated longitudinal dunes are less frequent, and barchans are still rarer. There are very few mentions of erg-like forms (Galon, 1958, 1969; Nowaczyk, 1986) or of simple straight-crested transverse dunes (Stankowski, 1963). Most studies dealing with these dunes were carried out by geomorphologists, whereas sedimentological data are rarely found.

Lithofacies characteristics Most extraglacial dunes consist of fine t o medium sand (Fig. 288). They show several types of sedimentary structures and of stratifications. Five main types of strata can be distinguished (some aspects of these were dealt with earlier in the corresponding section on proglacial dunes). The first important type of stratification is produced by avalanching of sand over the crest, followed by slumping of sand masses downslope over the dune foreset. The resulting strata are up to 5 cm thick. The consequence of the foreset-parallel deposition is that the dunes show high-angle aeolian cross-stratification on the lee sides, whether o r not barchans, parabolic dunes or transverse dunes are involved. This type of stratification was described in detail by Borowka (1979, 1980) who provided fine examples of the slumps. Hunter and Richmond (1988) found that daily cycles may be found in this stratification if the wind activity is subject to d a y h i g h t cyclicity. The large-scale sandflow cross-stratification has a tabular character and is well comparable to typical foreset bedding, as mentioned by several researchers (Stankowski, 1963; Dylikowa, 1969; Nowaczyk, 1976; Hunter, 1977b; Fryberger et al., 1979,1983,1984, 1988; McKee, 1979; Kocurek, 1981; Kocurek and Dott, 1981; Rubin and Hunter, 1983; Kerr and Dott, 1988; Kocurek, 1988). The second type of strata results from grains settling from the air on the depositional surface. Settling results from either flow separation behind the top of the dune or from a diminished transport capacity in shadow zones behind an obstacle (a height of a few centimetres is suffices for this) such as a wind ripple. Wind gusting produces separate layers (Fryberger et al., 1983). The grainfall process can also produce cross-

472


stratification in the lower foresets of very small dunes; the laminae are then less continuous and wedge out rapidly in all directions (Fryberger et al., 1979,1983,1988;Hunter, 1981). The third type of strata is represented by layers built up from material left behind by passing wind ripples. Such layers may extend a few metres downwind. Ripple-produced strata may also be due t o wind gusting that resulted in an alternation of coarser and finer layers on the lee side of ripples. Ripple foresets rarely exceed a length of 2.5-5 cm (Fryberger et al., 1983). Fine examples of such rippled foresets were described by Hunter (1981). The fourth type, adhesion-produced strata, results from the cohesive effects of a water-saturated sand surface on a carpet of sand moving across the surface (Hunter, 1973,1980b; Kocurek, 1981; Kocurek and Dott, 1981; Kocurek and Fielder, 1982; Fryberger et al., 1983, 1988; Hummel and Kocurek, 1984; Lindquist, 1988; Olsen et al., 1989).The stratification thus formed usually shows vague horizontal lamination with several microscale irregularities. The fifth type, representing flat plane-bed stratification, is produced by very strong winds. This type of horizontal lamination is more common in sand sheets and in coversands than in dunes.

Textural characteristics As can be seen, there are numerous studies touching the textural characteristics of dune sands in general (Bagnold, 1941; Mason and Folk, 1958; Pernarowski, 1959; Kuenen, 1960; Friedman, 1961; Kozarski, 1962b; Stankowski, 1963; Hand, 1967; Urbaniak, 1967; Galon, 1969; Glennie, 1970; Ahlbrandt, 1975, 1979; Nowaczyk, 1976; Borowka, 1980; Logie, 1981; Ahlbrandt and Fryberger, 1982). Reports concerning extraglacial dunes more specifically indicate that these are well sorted, with most grains in the 0.1-0.5 mm range. Studies of Polish dunes have indicated that the average grain size diminishes slightly in a dune area in the downwind direction. Rounding of the grains is much better in these dunes than in dunes from other glacigenic subenvironments. The rounding is a result of aeolian abrasion; the rounding process is first rapid, but slows down during continued aeolian transport. This is partly due to the fact that other grains also become better rounded, so that the abrasive capacity diminishes. Rounding of the grains is important from a sedimentological point of view, as higher wind velocities are needed for spheres than for grains with irregular surfaces.


473

Occurrence

Extraglacial dunes are found mostly where a sandy surface (e.g., in a fluvial plain) is present without protection from vegetation. Favourable locations such as these, preferably far away from the ice front (see a previous section) may occupy vast areas: extraglacial dunes of Vistulian age in Poland have been traced over a distance of more than 500 km (measured in the direction from the ice away), extending from North to South Poland. Palaeogeographic reconstructions indicate that the dunes started t o develop more or less simultaneously with the uppermost coversands and loesses of southeastern and eastern Poland (Cegja, 1972; Rbzycki, 1972,1979; Nowaczyk, 1986). The zones with extraglacial dunes have an extent comparable with that of coversands. The last glaciation, for example, left large dune fields in North America, Northern Europe and Northern Russia. Depositional mechanisms

Wind velocity, wetness of the substratum and vegetation are the main parameters that determine dune formation. However, variations i n wind direction and wind velocity are the final factors that control the detailed depositional process (Fig. 289). Most researchers agree that dune formation starts as a result of changing flow lines over a n obstacle at the sedimentary surface. Vegetation may be such a n obstacle, behind which there is a shadow zone where sand accumulates. The bodies thus formed have been analysed in detail (Hunter, 1977b; Hesp, 1981; Clemmensen, 1986; Guanatilaka and Mwango, 1989). The dunes start to migrate when the original obstacle has been fully covered by sand. During migrations, the dunes grow by accretion. Most extraglacial dunes are 8-15 m high, but may grow t o over 20 m if enough sand is supplied. Substratum conditions determine whether parabolic dunes or barchans are to be formed. Parabolic dunes are the more frequent and the largest type found in the extraglacial subenvironment. Variable wind conditions, as usually present far from the ice (i.e. at the outer edge of the extraglacial subenvironment), however, result in a higher barchadparabolic-dune ratio, but the barchans seen under these conditions are usually small and tend to be reshaped into longitudinal or irregular transverse dunes. Datings from the European Weichselian indicate t h a t extraglacial dunes are formed within relatively short intervals. Two periods of dune formation in the middle European lowlands can be distinguished during

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Fig. 289. Wind streamlines and resulting stratification of dunes. From: Hunter (1981). Courtesy: Journal of Sedimentary Petrology. R = climbing ripples; US = upper slipface deposits (grainfall deposits and/or slump masses); LS = lower slipface deposits (sandflow deposits, grainflow deposits and/or slump masses); S = undifferentiated slipface deposits; G = grainfall deposits. a: dune on which grainfall deposition extends beyond the slipface (occurs frequently on small dunes). b: dune affected by lee eddy winds (most common on lee slopes oblique to the wind direction). c: dune intermittently affected by wind reversals (reversing dune). d: dune without a zone of flow separation on its lee side (occurs only where the lee slope is relatively gentle).


475

this glaciation, and both periods probably lasted less than a thousand years. This implies that the net vertical accretion was of the order of 20 mm per year, which is high when compared t o most other rates of sedimentation, but much less than has been observed in modern areas with desert dunes. The migration rate of dunes with average size was approx. 0.3-3 km per year during the Weichselian periods of increased aeolisation.

Extraglacial coversands (II-C-5-d) The climatic conditions are such that vegetation is scarce in a widespread area of the periglacial zone. As a consequence, wind can easily affect the sediments, blowing them ovei' considerable distances and gradually letting them settle when wind activity decreases. The result is that it is mainly sands that are deposited in the 'proximal' part. The sands form a cover that tends t o reduce topographic differences but are usually also present, though thinner, in elevated areas. These coversands have always been much studied, especially by geomorphologists although purely sedimentological work has been published lately (e.g., Ruegg, 1983b; Schwan, l986,1987,1988a, 1990). Extraglacial coversands have been described mostly from the European lowlands, where they were deposited in large belts during the Saalian but especially the Weichselian glaciations (Koster, 1982a). Descriptions of sych coversands abound (Rutten, 1954; Diicker and Maarleveld, 1957; Maarleveld, 1960, 1968, 1976; Crommelin, 1964, 1965; Roep, 1968; Ten Cate, 1969; Veenstra and Winkelmolen, 1971; Borsy, 1972; Cailleux, 1972, 1973, 1974; Ruegg, 1975, 1981, 1983b; Nowaczyk, 1976, 1986; Gullentops et al., 1981; Koster, 1982b; Kolstrup, 1983; Schwan, 1988b). There are also several descriptions of coversands formed under modern extraglacial conditions (French, 1976; Pissart et al., 1977; Washburn, 1979; Good and Bryant, 1985). Extraglacial sands having several features shared with coversands are found in extraglacial interdune areas. Such sands are mostly described from the mid-European lowlands (Galon, 1958, 1969; Kozarski, 196213; Nowaczyk, 1967,1976,1986; Dylikowa, 1969; Gawlik, 1969; Borsy, 1972, 1974,1978;Karte, 1979; Borsy et al., 1981a,b,c). Lithofacies characteristics

These sands are usually horizontally stratified deposits, showing (sub)horizontal lamination and good sorting (Straw, 1963; Puritz, 1972; Nowaczyk, 1976; Buckland, 1982; Vandenberghe, 1983a; Schwan, 1986).

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This type of bedding is also termed 'horizontal alternating stratification' (Haest, 1985; Schwan, 1986). Schwan (1986) distinguished eight types of stratification in Dutch coversands. The first type is represented by poorly sorted sand with a coarse, indistinct lamination, or with a massive appearance; particles of up to granule size are present locally. The second type consists of well sorted sands with coarse, indistinct lamination or with a massive appearance. The third type is formed by well sorted fine sands with a uniform lamination strictly parallel to the sedimentary surface or with low-angle cross-lamination; individual laminae are clearly developed and are less than 1 mm thick. The fourth type comprises silts to fine sands with a coarse, indistinct and irregular lamination or with a massive appearance. Type five consists of well developed silt t o fine sand beds that contain abundant wavy or parallel lamination; the individual laminae are thin, often not distinctly developed, and crinkly in detail; cross-lamination due to the climbing of adhesion ripples is found between units with other structures. The sixth type is represented by silt t o fine sand with discontinuous wavy or even roughly parallel lamination, without distinct bedding planes or laminar surfaces. Type seven is formed by sand to silt with crosslamination produced by the climbing of adhesion ripples or by surficial currents; non-climbing adhesion ripples are rare. The last type consists of sand t o silt with scour-and-fill structures; the beds may appear massive but abundant horizontal lamination is found. There are comparable data from other researchers (e.g., Ruegg, 1983b) but much less detailed. The frequent occurrence of small-scale cryogenic deformations has also been mentioned repeatedly. However, there may also occur deformations that resemble those formed by cryogenic processes but have another origin (Schwan, 1990). The original lithofacies characteristics may have been eliminated by diagenetic processes such as soil formation. Intensive soil formation in extraglacial coversands is known from Hungary, among other places. The coversands in this case often show thick beds (10-20 cm), termed 'iron-pan layers' (Borsy, 1972), the characteristics of which are fully determined by intensive pedogenesis (Fig. 290).

Textural characteristics The textural parameters show wide variability. It is therefore remarkable that most coversands have in common a yellow-grey t o whitish-grey colour. Grain size may range from coarse sand t o fine silt and sorting is equally variable. Schwan's (1986) stratification types 2 and 3 are very


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Fig. 290. Exposure near Aranyosapati (Hungary) with typical coversand succession. 1 = brown forest soil with iron-pan layers; 2 = coversand from the Older Dryas; 3 = palaeosoil of Bolling age; 4 = loess; 5 = coversand from the beginning of the Upper Pleniglacial. From: Borsy e t al. f19Sl). Courtesy: Acta Geographica Debrecina.

well sorted, types 4, 5 , 6 and 7 are moderately sorted, and types 1,2 and 8 are badly sorted. Van der Hammen ( 1951) considered bimodal grain-size distributions, with peaks in the 70-75 and the 105-150 micron ranges, characteristic of coversands, but unimodal distributions have also been reported, usually from distal places, where only fines dominate (Ducker and Maarleveld, 1957, and Ruegg, 1981, have reported unimodal distributions with peaks in the 16-63 micron range). It appears that the sorting of the coversands tends to be much poorer than that of extraglacial dunes. The grains are moderately to well rounded, with a tendency to better rounding towards the top of a coversand section. The surface of the grains shows typical aeolian abrasion marks. Occurrence The coversands mostly form a belt covering the proximal part of the extraglacial subenvironment. The deposits form widespread blankets, typically 0.5-3 m thick, that slightly level the underlying deposits. It is fairly rare, however, t o find height differences of over 1.5 m and slopes of over 1". Coversand ridges, elongated in the wind direction and with a height of 4-5 m above the surrounding are exceptions. The sands usually have a sharp base, although gradual transitions from underlying fluvial deposits have been reported; vertical alternations of fluvial deposits and coversands, resulting from changing environmental conditions, are common.

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Most coversands in the glacigenic sequence are found between loesses and fluvial (occasionally lacustrine) deposits; there may also be contact with extraglacial dunes and mass-transport deposits. Depositional mechanisms

The involvement of a number of depositional processes has been suggested as regards the formation of coversands. The most probable interpretation of this complex genesis has been proposed by Schwan (1986, 1988), who analysed the eight lithofacies types summarised above, It seems justifiable t o use his interpretations here. The sands of Schwan's first type of lithofacies represent niveo-aeolian deposition. The snowstorms must have been severe, as shown by the granule size of the particles. Micro-scale deformation of the substratum (due t o liquefaction and loading) can be the result of water saturation following melting of the snow. The second lithofacies type is also due to niveo-aeolian activity, probably responsible for deflation of nearby sands. A combination of three conditions is responsible for the characteristics of these deposits, including their good sorting: (1)absence of excessive wind speed, (2) a local source which was itself deposited by snow-free winds, and (3) the incorporation of sand in drifting snow moving near the surface, so that interception of suspended grains is excluded or considerably reduced. Lithofacies type 3 is a result of vertical accretion as a result of aeolian supply. The plane-bed lamination and subcritically climbing tranlatent stratification (cf. Hunter, 1977b) are due to tractional deposition at wind velocities too high to form ripples (cf. Bagnold, 1941; Fryberger et al., 1983). The result of this depositional mechanism is the formation of thin sand sheets or very flat lenses. Lithofacies type 4 is formed by settling from suspension during periods of low wind speed. The silty material may settle on a dry or a damp surface, or may fall into shallow pools with stagnant water (cf. Cegla, 1969,1972). The fifth type is interpreted as a result of settling from the air onto a moist surface. The moist surface is due to capillary rise from a seasonally thawing part (active layer) in the permafrosted substratum. The lamination is identical t o plane-bed adhesion lamination (cf. Kocurek and Fielder, 1982),which was termed 'quasi-planar adhesion stratification' by Hunter (1980b). The relatively poor structural development in the sixth lithofacies type is interpreted by Schwan (1986) as the result of greater textural uni-


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formity and/or less variation in - a generally low - wind speed during deposition. Lithofacies type 7 (Fig. 291) is mainly a result of short-lived periods during which thin beds are formed, intercalated between laminated strata, as a result of adhesion. Melting snow results in superficial runnoff, which is expressed by units with small-scale current ripples. The last, or eighth, lithofacies type is primarily a result of superficial water currents. These are generally weak (and shallow), but short periods of increased current velocity allow them t o scour the substratum; a successive period of waning flow results in the filling of the scoured depressions. The alternation of all these processes usually results in a complex succession of lithofacies types, which also pass into each other in lateral directions. The periglacial climate, including the presence of a permafrost, is an extremely important factor (Rutten, 1954; Ducker and Maarleveld,

Fig. 291. Coversand lithofacies type 7 of Schwan (1986). The lower part shows predominantly climbing adhesion ripples, whereas the upper part is dominated by irregular current ripples. Photograph:J. Schwan.

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1957; Cailleux, 1973; Maarleveld, 1976; Pissart et al., 1977; Kolstrup, 1983; Ruegg, 1983b; Schwan, l986,1987,1988a,b).

Extraglacial loesses (II-C-5-e) In the idealised model, extraglacial loesses are found in a belt that surrounds the area with extraglacial dunes and coversands (Cegfa, 1972; French, 1976; Embleton and King, 1977; Rozycki, 1979; Washburn, 1979; Schwan, 1986). A large and uninterrupted loess belt extends from China, passes via Russia and Central Europe to reach the Atlantic coast. Similar deposits are also commonly found in North and South America. Loess deposits are highly characteristic of the periglacial zone (Jahn, 1950, 1956,1975; Dylik, 1954; Pew6,1969; Willmann and Frye, 1970; Rozycki, 1972, 1979; Smalley, 1975; Mucher, 1986; Pye, 1987). Most descriptions concern deposits from the two latest phases of the Pleistocene glaciation, but there are also reports of Precambrian loessites from Norway and Svalbard (Edwards, 1979), Palaeozoic loessites from the USA (Murphy, 1987; Johansen, 19881, Jurassic loessites from the USA (Rautman, 1975), Neogene deposits from the central plains of the USA (Hunt, 1985; Winkler, 1987) and Neogene deposits from NW Colorado (Johnson, 1989). All these 'fossil' loesses and loessites are interpreted as having been formed under extraglacial conditions. The genesis and structures of these deposits, especially of those in Poland, have been studied in detail by Jahn (1950,1956), Sawicki (1952), Mojski (1965) and Jersak (1973) among others. Facies characteristics were described by Dylik (1954), Malicki (1961b), Mojski (1965) and Cegia (1972), but without a distinctly sedimentology-oriented interpretation. More recent studies (Mucher, 1986; Pye, 1987) have provided further sedimentological data, but several of the phenomena seen in loess profiles have not yet been explained satisfactorily. The term 'loess' was introduced by an amateur geologist, Karl Caesar von Leonard (probably 1823-1824). Lye11 (1834) introduced the term into English. A first hypothesis regarding an aeolian origin was proposed by Virlet d'Aoust (1857). This interpretation was eventually accepted through the efforts of Richthofen (1882) who had travelled t o China, where the existence of loesses (huang tu) had been described already more than 2000 years ago (see Liu Tung-sheng et al., 1985). A type of deposit so long known and studied (it yields one of the most fertile soils) should not be the object of controversy. However, few deposits have received so much attention and, indeed, raised so much controversy. Even the definition of the term 'loess' is not generally agreed upon. Pye


481

(1987) defines loess as a terrestrial wind-blown deposit of silt consisting chiefly of quartz, feldspar, micas, clay minerals and carbonate grains, in varying proportions. Heavy minerals, phytoliths, salts and volcanic-ash particles may be locally important constituents. Fresh (i.e. non-weathered) loess is, according t o the definition of Pye (19871, homogeneous, not or only vaguely stratified and highly porous. It is usually buff coloured, but may also be grey, red, yellow or brown. Under dry conditions, it forms standing scarps (these are very characteristic of loess areas), partly because of a tendency to fracture along systems of vertical joints (Fig. 292). Saturation with water considerably reduces shear strength so that the material is then subject t o flowage and sliding. Another topic that has now been discussed for a century and a half, and still remains controversial, is the question of what should be considered as a 'typical loess'. Pye (1987) stated that typical loesses have a grain-size distribution with a distinct peak in the 20-40 micron range, and that there is positive skewness (towards the finer sizes). However, fine sand often makes up about 10%; with a sand content of over 20% the deposit no longer a 'typical loess' but becomes a 'sandy loess'. Up t o 20% clay may

Fig. 292. Loess scarp from Szczebrzeszyn (SE Poland) with characteristic columnar jointing. Photograph: J. Cegla.

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also be present in a typical loess. This approach, as used by Pye (1987), is far from being completely satisfactory. Weathered loesses, for instance, may contain up t o 60% clay, often in buried paleosoils. The term 'loessoid' is sometimes applied t o mixtures of soil components, alluvial material and aeolian dust and sand. The term 'loess-like' is sometimes used t o indicate deposits with the grain-size distribution of a 'typical loess' but of nonaeolian origin; such deposits may be overbank silts, lacustrine bottomsets and some colluvial deposits (Pye, 1987). They may, however, also be formed in a warm climate (Kriger, 1965; Yaalon, 1965, 1969; Yaalon and Ganor, 1966; Yaalon and Ginsbourg, 1966; Smalley and Vita Finzi, 1968; Kukal and Saadallah, 1970; Fedorovich, 1972; Yaalon and Dan, 1974; Smalley and Krinsley, 1978; Sneh, 1983). This short review of the terms used indicates once more how much difference there is between the sedimentological and the geomorphological or soil-science terminology. Loesses have remarkable characteristics that make their study a highly interesting one. Frequent experiments thus have been carried out for the purpose. There is consequently an abundant literature on topics such as chemical composition, grain-size distribution, texture, susceptibility t o erosion, porosity and permeability, depositional mechanism and source area (Cegla, 1969,1972; Jahn, 1970,1975; Washburn, 1973,1979; Smalley, 1975; Catt, 1977; De Ploey, 1977; Embleton and King, 1977; Mucher and De Ploey, 1977; Embleton and Thornes, 1979; Whalley et al., 1982; Mucher, 1986). Many questions remain t o be answered, in spite of the numerous studies carried out so far. One of the most interesting questions is that about the cause of the extreme thickness of loess covers in some areas that underwent periglacial conditions during the Pleistocene. The loesses in The Netherlands are generally rather thin covers, exactly like those in Great Britain, Belgium, France and Western Germany; loess in Poland may (exceptionally) reach thicknesses of some 50 m, those in Czechoslovakia some 100 m and those in Hungary more than 100-150 m, while the loesses can reach up to 200 m in the Soviet Union (Ukraine area), and those in China reach up t o 400 m! Most researchers now agree that a combination of several factors determines the regional thickness of a loess cover (Jahn, 1956,1975; Goudie, 1978; Smalley and Leach, 1978; Rozycki, 1979,1986; Washburn, 1979; Smalley, 1980; Pye, 1987). There are also loesses from modern areas in the extraglacial subenvironment, for example Svalbard (Bryant, 1982; Kida, 19861, Alaska (Pewe, 1975),Greenland and Arctic Canada. There are several reports on modern Siberian loesses (Melnikov, 1966; Popov, 1967, 1973; Sudakova, 1969; Volkov et al., 1969; Fedorovich, 1972; Konishchev, 1972, 1973;


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Tomirdiaro, 1972, 1975a,b; Danilova, 1973; Pewe, 1973a,b,c; Tomirdiaro et al., 1974; Pbwe and Journaux, 1983). These studies are particularly important in t h a t they provide a chance t o analyse the depositional mechanisms in the field.

Lithofacies characteristics Particles finally deposited as loesses may travel over very large distances. This implies that there may be difference in lithofacies between 'proximal' and 'distal' loesses. The former will, for example, usually have a higher sand content than the latter. Most extraglacial loesses, however, share a massive appearance, a yellow, brown or yellowish-grey (olive grey) colour and a high silt content. Sand grains in the loesses are usually randomly distributed, although thin laminae may sometimes be found that are clearly enriched in sand (sand concentrations may also be found, particularly in 'proximal' loesses i n shadow zones behind obstacles, e.g., vegetation remnants). The lamination is usually not very distinctly developed, either in the 'proximal' loesses - where sandy laminae are found embedded in silt - or in the 'distal' loesses where relatively coarse silt may be found concentrated in the overall fine-silty to silty-clayey sediments. Lamination is usually discontinuous and irregular, and the thickness of the laminae varies from less than a millimetre t o about one centimetre. There are also extraglacial loesses that show much clearer lamination (Fig. 293). Detailed studies of laminated loesses have been reported by J a h n (1950,1956), Cegla (1961a,b, 1964,1965, 1971,1972), Jersak (1965, 1973))Lindner (1967))Borowiec and Nakonieczny (1968) and Kida (1981, 1984, 1986) for Silesia and the surroundings of Lublin (Poland). The horizontal lamination in these loesses may have a composite character and is then usually continuous; the sediments are composed of sets that each contain about 5-15 laminae that are 1-5 mm thick. There also occurs a non-composite (simple) laminated loees type, most frequently in the 'distal' parts. The lamination (up to 5 mm thick) in these clayey loesses is discontinuous and individual laminae give the impression of being slightly concave upwards (fills of shallow depressions). A third lithofacies type shows variable cross-bedding. This type is most common i n sandy loesses; the grain-size differences make small-scale cross-lamination easily visible. These probably current-induced ripples are often accompanied by adhesion structures. Normal grading is sometimes found i n sandy loesses, but the differences in grain size between the base and the top parts of the graded units are usually so slight t h a t they can only be

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Fig. 293. Loess with distinct horizontal lamination. The lower parts has been deformed by cryogenic processes. Photograph:J. Cegh.

detected microscopically or, under favourable conditions, as a result of differential erosion. One also finds secondary (post-depositional) structures in loesses in addition t o the primary structures mentioned. One of the most important is the lamination-like structure resulting from subhorizontal fracturing due t o the melting of ground-ice lenses (Popov, 1967). Furthermore, one frequently finds load structures that result from unstable density gradients (Butrym et al., 1964), flowage and fluidisation structures, collapse structures, and water-escape structures (Liszkowski, 1971; Cegta, 1972). Cryogenic deformations such as cryoturbations (Fig. 293) and frost fissures (Jahn, 1950, 1956, 1975; Mojski, 1965; Goz’dzik, 1973) are a less common finding. The massive appearance of loesses may also be a secondary feature, e.g., as a result of pedogenesis (Jersak, 1965,1973).

Textural characteristics A typical (unweathered) loess is mainly composed of particles in the 10-50 micron range (Browzin, 1985; Tsoar and Pye, 1987), usually with a peak in


485

the 20-30 micron range. There often are gradual transitions in aeolian deposits between silty sands and sandy silts (cf. Ravikovitch, 1953; Lugn, 1962, 1968; Kes, 1984), so that sand grains dispersed between the silt particles are a normal feature. These characteristics, contrary to what is commonly assumed, result in loess being a relatively poorly sorted sediment with sorting values (Folk and Ward, 1957) of the order of 1 t o 3 and an almost exclusively positive skewness (indicating a tail of fine particles) in the range between 0.3 and 0.7. The proportion of aggregated fine particles is significant (Gillette and Walker, 1977; Whalley and Smith, 1981). The shape of the individual loess particles depends on the mineral composition, the crystallographic structure, the processes responsible for their formation and the influence of diagenetic processes, including weathering (Pye, 1987). Quartz grains derived from sediments actively affected by chemical weathering (e.g., soils) often show edge rounding due to partial solution and reprecipitation, whereas fresh grains formed by, e.g., subglacial crushing or frost action have sharp edges and surfaces characterised by conchoidal fractures. Feldspar grains have a shape largely determined by mechanical breakage following cleavage, such grains often forming equidimensional particles. Phyllosilicates occur predominantly as platy grains. Fragments of volcanic glass are often angular and have a smooth surface (Cegla, 1969; Pye, 1987). The surfaces of loess grains are often covered by adhering clay-sized particles and/or amorphous aluminum silicates. While Smalley and Cabrera (1970) interpreted these coatings as a result of glacial comminution, similar coatings have been produced experimentally by non-glacial processes (Pye and Sperling, 1983). Minervin (1984) reported that some Russian loesses from central Asia contain micro-aggregates or globules (10-100 micron in diameter) with a quartz (occasionally feldspar) core surrounded by concentric layers of amorphous silica gel and carbonate, with an outer layer of clay minerals, iron hydroxides, amorphous silica, finely dispersed quartz and carbonates. This author interpreted such features as being the result of cryogenesis during transport and deposition under mainly glaciofluvial conditions. The textural characteristics of loesses have been well reviewed by Pye (1987).

+

+

Occurrence Extraglacial loesses are widely distributed. They are found, for instance, in a more or less continuous zone 200-800 km wide, forming a belt parallel to the ice-marginal zone that existed during the phase of maximum ice

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extent in the Pleistocene. Regionally, the zones may have a width even greater than 800 km (North America, Asia). As the extraglacial subenvironment comprises - in the framework applied in this book - only two facies, and taking into account the fact that loesses form the most distal aeolian deposits, it can be concluded that the most distal glacigenic deposits are made up of loesses with intercalations of subaerially formed mass-transport extraglacial deposits. Loesses are much more frequent than mass-transported sediments, which implies that they usually are the outermost sediments that can be considered as glacigenic. It also implies that extraglacial loesses tend t o be found as both the lowermost and the uppermost deposits in the glacigenic sequence. The areal distribution of loess is still under discussion. It appears that several relatively thick loess accumulations are found in relatively high areas (these areas also having been elevated during deposition of the loess). Interpretating thick loess accumulations as being the result of slope processes thus cannot be valid, at least not for all accumulations (Woldstedt, 1954; Jahn, 1956; Charlesworth, 1957; Jersak, 1965, 1973; Cegla, 1972). Accumulations in elevated areas should instead be considered as resulting from regional wind patterns.

Depositional mechanisms The most important parameters influencing ( o r determining) loess deposition are the nature of the source material, the position of the source area, wind direction, fluctuations in wind velocity (including the occurrence or lack of heavy storms), the type of transport, the conditions prevailing at the depositional site (including the presence or absence of vegetation, the relief, the groundwater table), and the intensity of precipitation. Even a single one of these parameters, the provenance of the fine silt particles, has been the object of numerous - often interdisciplinary studies. It was shown that probably more than ten processes make a considerable contribution to the presence of silt-sized particles as a source of loess. Tutkovski (1899, 1910) and Geikie (1898) had suggested that glacial processes could produce silt-sized particles, but it was discovered much later that only glacial grinding is capable of transforming large quantities of sand (and coarser particles) into silt (Smalley, 1966, 1980b; Smalley and Vita Finzi, 1968; Boulton, 1978; Smalley and Krinsley, 1978). This, latter, possible origin of silt has been confirmed experimentally by Vivian (1975), Whalley (1979), Haldorsen (1981, 1983) and Sharp and Gomez (1985) among others.


487

The importance of frost weathering in the case of silt production was a concern of Zeuner (19491, St. Arnaud and Whiteside (19631, Smalley et al. (1978), Konischev (19821, Smalley and Smalley (1983) and Pye and Paine (1984). Experiments by Martini (1967), Brockie (1973), Moss et al. (1981) and Lautridou and Ozouf (1982) demonstrated the effectiveness of the frost-weathering process. Phyllosilicates in fine-grained parent rocks may be, according to Kuenen (1949), a source of the silts that form loesses. Blatt (1967, 1970) found that granitic rocks may be sources of both sand and silt. Moss et al. (1973) have claimed that fluvial comminution plays a n important role. However, it is generally considered that aeolian abrasion is a more effective process (Smalley and Vita Finzi, 1968), especially because of many experiments with the latter process were consistent with this latter interpretation (Knight, 1924; G.E. Anderson, 1926; Kuenen, 1960; Krinsley et al., 1981; Whalley et al., 1982; Krinsley and Greeley, 1986). Salt weathering has also been mentioned as a process of local importance (Goudie, 1977,1985,1986;Goudie et al., 1979; Goudie and Day, 1980; Cooke, 1981; Pye and Sperling, 1983; Fahey, 1985). According to several authors salt weathering in combination with frost action may produce large amounts of silt (Goudie, 1974; Hudec and Rigbey, 1976; McGreevy, 1982) This field observation was experimentally confirmed by Williams and Robinson (1981). The role of chemical weathering in the reduction of particle size during pedogenesis was recognised by Van der Waals (1969), Crook (1968), Eswar a n and Stoops (1979), Nahon and Trompette (1982) and Pye (1983a, 1985). Diurnal o r seasonal temperature changes may induce such large thermal gradients that mechanical weathering occurs as a result. Insolation can have the same result (Roth, 1964; Peel, 1974). The possible production of silt by this process has been discussed in detail by Yaalon (19741, Rice (1976), Smith (1977) and Winkler (1977). Alternation of drying and wetting has also be proposed as a process that would result in the formation of silt-size particles (Blackwelder, 1925; Griggs, 1936). Silt produced as a result of biogenic activity has been mentioned by several researchers. Silt-sized biogenic opal was found by Yeck and Gray (1972), Norgren (1973) and Wilding et al. (19771, whereas the finding of silt-sized fragments of radiolarians and echinoderms was referred to by Jones et al. (1964). Pollen and spores may also be present in loess a s siltsized particles (Horowitz et al., 1975; Melia, 1984).

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Clay may form silt-sized aggregates (pellets) through a number of mechanisms (Lancester, 1978a; Dare-Edwards, 1982,1984). Most researchers now agree that local factors determine which of the many processes mentioned will predominate in a specific area. Climate, relief and life forms are thus important factors. For example, glacial and fluvioglacial abrasion may, together with frost action, be the determining mechanisms under cold conditions. As mentioned earlier in this section, loess formation also depends on conditions at the depositional site. Pye (1987) argued that wind-blown silt particles will settle from the air if wind velocity and turbulence decrease o r if the particles are 'captured' by collision with rough, moist or electrically charged surfaces, if the particles become charged and form aggregates or, finally, if the particles are washed out together with precipitation. Small particles may accumulate on a smoooth surface if they become immersed in the thin laminar layer just above the sedimentary surface (Owen, 1960). Another important process is the rapid deposition of dust t h a t occurs when a silt-laden cloud meets a barrier, e.g., when passing from bare soil to a vegetated area (Oke, 1978; Pye, 1987b). Forests thus force dust to settle much more effectively than do steppes and tundras. Dust is also easily deposited both on the lee side of topographic obstacles (Jackson and Hunt, 1975), and where it becomes fixed by a wet sedimentary surface. It is interesting t o note that this 'wet deposition' reduces the evaporation of groundwater, so that the groundwater table will gradually rise, wetting the freshly deposited sediments, thus facilitating further accumulation (Cegla, 1969,1972). Goossens (1985b) pointed out that grains settle collectively in the form of a more or less coherent mass rather than individually from dust clouds with high concentrations of particles. He demonstrated experimentally that a n 'explosion point' is reached during dust fall; lateral diffusion then reduces particle concentration to a level that allows particles to settle individually. It should, however, be kept in mind that dust concentrations in nature reach Goossen's values only during storms of exceptional intensity (Pye, 1987a). It is well known that dust may remain suspended in the atmosphere for long periods. A long-lasted stay in the air can result in collision and aggregation of the fine particles, simply as a result of Brownian motion, laminar shear or turbulent motion (Friedlander, 1977; Suck et al., 1986), but a build-up of bipolar electrostatic charges may also play a role in aggregation (Greeley and Leach, 1979; Marshall et al., 1981). Aggregates thus formed have a higher mass/surface ratio, which makes them less


489

susceptible t o air turbulences, with the result that they will settle earlier than non-aggregated particles. Several authors have stressed the importance of washout by rain or snow as regards deposition of air-suspended particles (Itagi and Koeunuma, 1962; Ganor, 1975; Graedel and Franey, 1975; Knutson et al., 1977). Two mechanisms of washout were distinguished by Pasquill and Smith (1983):immediate collection of dust by rain, hail or snow falling through the cloud, and capture of the finest mineral particles by cloud droplets, with subsequent deposition of the particles still adhering to the rain droplets. The structures encountered in loesses facilitate the reconstruction of the specific depositional process(es) involved. A lack of structures, i.e. a massive appearance, may be due to postdepositional processes but may also result from deposition on a substratum that was not entirely dry so that the particles immediately became stabilised without a chance of moving laterally. This type of situation cannot exist when winds a r e blowing; particles thus fixed will, as a rule, not start t o move when the wind velocity increases, as this requires a relatively strong increase in energy level. In this context, it must be remembered that the average sedimentation rate of loesses is some 2 mm per year so that there will be long-lasting periods without sedimentation even if there are years with a much more rapid accumulation. Rain, snow and wind may affect the upper sediment layer during these periods and rain drops falling on the surface may destroy any lamination that may have existed, thus homogenising the sediment. A truly wet sedimentary surface, with films of water of the order of 1 mm will almost immediately catch particles touching t h e ground, although the grains may float for some time at the surface of water bodies of sufficient depth. Winds will blow these grains to areas where they reach the mineral bottom, so that they become stabilised, forming micro-ridges lense-shaped in transverse section. This depositional mechanism is assumed t o be responsible for the discontinuous lamination found i n loesses deposited on uplands without vegetation. The primary lamination i n loesses shows great variation and can generally be attributed to three mechanisms. The first type of lamination is formed when dust settles in small ponds. Grain-size differences of the particle involved will then result i n grading. Successive phases of deposition thus result in laminae that are each characterised by grading on the micro-scale. The second type of lamination is formed when particles settling from the air reach flowing water in overbank areas. The particles will accumulate there through fluvial processes. Such 'loessial' overbank

490


deposits can reach thicknesses of over 2 m, as the authors observed in the valley of the Bobr river (western Poland). The deposits usually show a n admixture of fluvial elements that confers 'non-typical loess' characteristics. The third type of lamination is due to splash erosion, which can be considered as a form of rain erosion. Only a slightly inclined surface is required ( < 1")to force rain to run off over the impermeable subsoil. Many of the loess deposits show this type of fine lamination. Niveo-aeolian deposition will not produce distinct lamination because melting of the snow during summer or other warm periods will destroy the original structures. A number of loess deposits show some type of coarse lamination or bedding as a result of slope processes (Jersak, 1963, 1973; Cegfa, 1972; Pewe, 19751, b u t t h e present authors consider deposits formed by slumping, mudflows, etc. as belonging the extraglacial mass-transport facies (11-C-6). Lamination of a secondary (diagenetic) nature may also be seen. It appears from the literature that the distinction between primary and secondary lamination is usually a difficult one t o make. Pseudolamination or pseudostratification are often the result of upward (capillary), downward (infiltration) or horizontal (percolation) movements of water. Several types of smudges, shadows, laminae and lenses may thus be produced (Cegla, 1972). So-called pseudoparalamination may also result from the melting of lenses of ground-ice (Popov, 1967; Jersak, 1965,1973; Cegla, 1972; Pye, 1987). Most of the studies devoted t o this topic originated from the USSR and the USA.

DEPOSITS O F THE EXTRAGLACIAL MASS-TRANSPORT FACIES (11-C-6) The mass-transport deposits found in the extraglacial subenvironment are considered as glacigenic only if they consist, at least partly, of reworked material of glacigenic origin, and if they result from processes controlled by cryogenic conditions. Consequently, this facies comprises only one type of deposit: the extraglacial subaerial mass-transport deposits (II-C-6-a). This facies has long attracted interest from geomorphologists (see reviews in J a h n , 1975; French, 1976; Embleton and King, 1977; Washburn, 1979). In recent years, sedimentology has concerned itself more Eyles, 198313; N. Eyles and Paul, 1983; frequently with these studies (N. N. Eyles and Clague, 1983; N. Eyles and Kocsis, 1988; N. Eyles et al., 1 9 8 8 ~ )Some . of these studies, however, have yielded controversial results

Deposits of the extraglacial mass-transport facies

49 1

(N. Eyles et al., 1990; Mandryk and Rutter, 1990). It is therefore of the utmost importance that the facies be studied in greater detail, particularly because its presence may be taken as a clue for establishing the boundary line between glacigenic and non-glacigenic environments. Extraglacial subaerial mass-transport deposits (II-C-6-a) The same three categories of deposits can be distinguished here as in the proglacial subenvironment. The first category comprises deposits formed by large-scale, fast mass movements such as mudflows and debris flows. These processes can involve huge quantities of material during even a single phase of transport (N. Eyles and Paul, 1983; N. Eyles and Kocsis, 1988; N. Eyles et al., 1 9 8 8 ~ ) The . second category includes deposits formed by small-scale, usually surficial, fast mass movements such as skinflow (McRoberts and Morgenstern, 1973, 1974; French, 1976; Brown et al., 1981; Carter and Galloway, 1981; Strangl et al., 1982), mudflows (Shilts, 1978; Egginton, 1979; French and Egginton, 19841, multiple regressive flows (McRobertsand Morgenstern, 1973; Chatwin and Rutter, 1978; Brown et al., 19811, multiple retrogressive sliding (Brown et al., 19811, retrogressive-thaw flow slide (Hughes et al., 1973), active-layer gliding (Hughes, 1972; Hughes et al., 1973; MacKay and Mathews, 19731, slumping (Hughes, 1972; MacKay and Mathews, 1973; McRoberts and Morgenstern, 1973), rockfall (Luckman, 1976), blockfall (Brown et al., 1981), debris flow (Rapp, 1960,1975;Larsson, 1982),ground-ice slumping (MacKay, 1966; French, 1976) also known as 'thaw slumping' (Washburn, 1980), and bimodal flow (McRoberts and Morgenstern, 1973, 1974). The third category is represented by deposits formed through slow surfical mass-transport processes such as permafrost creep (Williams, 1979, 19821, frost creep (Washburn, 1967; Benedict, 1970; Price, 1972), needleice creep (Washburn, 1969, 1980; Beaty, 1974; French, 1976; Nicholson, 1978) and gelifluction (Andersson, 1906; Washburn, 1980). Many of the researchers have also emphasised the important depositional role of slopewash, a process that (according t o Lewkowicz, 1988) can be considered to be a mechanism including both surface wash (Jahn, 1960,1961; Czeppe, 1965; Pissart, 1967; Journaux, 1976) and subsurface wash (Lewkowicz, 1981; Lewkowicz and French, 1982). All the processes mentioned above are very active in the extraglacial subenvironment. Deposits formed by these processes however often have a limited extent (both as to thickness and areal extent) because the deposits thus formed - commonly on slopes - are easily affected (and again removed) by successive slope processes. Most deposits of this type are

492


therefore found where the slope has decreased (Fig. 294), or where there were depressions in the subsoil. Thick accumulations of mass-transported sediments may be found in such locations, where they are also easily protected against erosion by the presence of a cover of aeolian sediments. These deposits have attracted much attention from geomorphologists. The 'grezes litees' (Dylik, 1956,1960; Journaux, 1976; De Wolf, 1988) are typical examples. However, the terminology applied t o these deposits is highly inconsistent and absolutely does not conform with geological (and particularly sedimentological) nomenclature.

Lithofacies characteristics The parent material and the precise mass-transport mechanism largely determine the lithofacies characteristics of these deposits. Large-scale debris flows (N. Eyles and Kocsis, 1988), for example, contain mainly massive, matrix-supported diamicts that may show vague grading, either normal or inversed. The thickest layers (up t o about 10 m) contain deformed slabs and may show a weakly developed stratification. Crude stratification may also result from superimposed deposits of massive beds. The slow, surficial mass-transport processes usually generate strongly deformed diamicts, sometimes with vague stratification. Downslope transport is commonly so slow that other deformation processes (cryo-

Fig. 294. Cross-section through the margin of the Mroga valley, which is part of the Warsaw-Berlin pradolina system. From: Jahn (1975; after Dylik, 1967). 1 = fine sand; 2 = gelifluction deposit; 3 = silt (reworked loess?); 4, 5 = slope deposits; 6 = extraglacial fluvial sand; 7 = block of permafrost; 8 = young slope deposit.


493

turbation as a result of gradients in the cryostatic pressure, water escape, frost heave, etc.) may affect the same material simultaneously and so strongly influence the lithofacies. The so-called 'grezes litees' form exceptionally well stratified sediments, which is exceptional for subaerial mass-transport deposits. Most fast surficial mass-transport processes produce relatively thin, m a s s i v e layers, but these usually occur as a repetition, on top of each other,giving the impression of a stratified succession. Only slopewash deposits are sometimes themselves distinctly laminated. Processes such as rockfall and blockfall generate breccias or diamicts with floating clasts of the parent sediment embedded in the reworked material. The parent material is by far the most important parameter of the lithofacies characteristics of these deposits.

Textural characteristics The mass-transport process involved determines the textural characteristics of these deposits. Debris-flow deposits as described by N. Eyles and Kocsis (1988) show a wide range of textures, partly because of strongly varying clast/matrix ratios, partly because of the grain-size distribution in the matrix and in the clast range. The clasts, which may have all types of lithology and shape, show a strongly preferred orientation in a subvertical direction and give the impression of being positioned according to the flow lines of eddies or swirls. However, some debris-flow deposits show a basal unit with imbricated clasts or subhorizontal clasts with a preferred orientation of the a-axes. The combination of clasts and matrix results in a distinct bimodal granulometry, an important characteristic of debris-flow deposits. The mud content varies widely, from less than 10% to over 50%. It is not uncommon that the percentage of mud matrix is overestimated, as part of the fine material may in fact belong t o mud clasts t h a t have the same lithological characteristics as the matrix. Detailed studies of solifluction lobes (Worseley and Harms, 1974; Benedict, 1976; Matthews et al., 1986) have evidenced that the textural characteristics of these deposits do not differ much from those of debrisflow deposits; the main difference is that the clasts in solifluction lobes usually have a much more pronounced preferred orientation (Fig. 295). The deposits formed by micro-scale mudflows, liquified flows or slopewash show a wide range of textural characteristics. The 'grezes litbes' in particular show an extremely well developed preferred orientation of the clasts (sub)parallel to the slope surface, with occasional imbrication. The


494

'0

2

1

3

metres Fig. 295. Fabric diagrams for turf-banked (top) and stone-banked (below) solifluction lobes from the Niwot Range, Front Range, Colorado. Slightly modified from: Benedict (1976). 1 = cobbles and boulders (to scale); 2 = rose diagram of a-axis orientation of fifty stones; 3 = rose diagram of two hundred elongated sand grains in thin section. Circles in roses represent 10% frequency.

rock fragments in the coarser units are angular and reach sizes up to 2.5 cm (Washburn, 1969). The matrix of these coarse 'grkze litees' is laminated, due t o the rapid vertical alternation of average grain size. The matrix may be predominantly sandy, but may also consist of mud. 0 cc u rrence

The existence of a slope is a prerequisite for the formation of all these deposits. Some processes need escarpments (rockfall, blockfall), some need steep slopes (slumps) and some need only slight inclinations (gelifluction). However, the intensity of almost all processes increases as the inclination increases, whereas the process slows down or stops if a certain lower threshold value has been exceeded. This explains why most


495

deposits are found a t the foot of hills or inside topographic depressions. Successive deposits may form thick successions (Dylik, 1960; De Wolf, 1988). The precise transport mechanism determines whether thick units of limited areal extent or whether widespread, thin deposits will be formed. The total volume of material obviously is another parameter that influences thickness and extent. Individual beds seldom reach thicknesses of over a metre; thicknesses over five metres are definitely exceptional. The deposit formed by mass-transport processes usually has a sharp lower boundary. However, there may occur a gradually more pronounced reworking from the base upward, particularly if processes are involved in which friction along the detachment plane plays a role. In such cases, there may be a zone - occasionally characterised by shearing - which is intermediate between the original deposit and the reworked material. On the other hand, the general mass-transport process may stop at a certain moment, whereas the topmost part may still undergo some movement. The upper boundary will be not very distict in this case. The deposits may occur in such high concentrations that, as a n entity, they cover large areas. This is the main reason why i t is considered appropriate to group these deposits as a separate facies. Single deposits, or groups of very few deposits, may also be found in isolated positions; interfingering with other types of deposits is much more common in the latter case. Most lateral (and vertical) contacts are with extraglacial loesses. It can even often be difficult to discover whether a specific unit with loess characteristics has been affected by mass-transport processes, particularly because the reworked material (so-called 'slope loesses') tends to show laminations t h a t need not necessarily differ from the primary lamination in loess. There may also be lateral and vertical contacts with non-glacigenic sediments. This emphasises the fact that these deposits are found in the transitional zone between the glacigenic and non-glacigenic environments. The same relationships are found in vertical directions in the glacigenic sequence.

Depositional processes The processes involved in subaerial mass transport in general have been much studied (e.g., Hampton, 1975; Larsen and Steel, 1978; Lowe, 1979; Nemec and Muszyiiski, 1982; Waters, 1983; A.M. Johnson, 1984; Nemec and Steel, 1984, 1988b; Nemec et al., 1984; Shultz, 1984; Rappol, 1985; Van Dine, 1985; Postma, 1986). The findings from these studies can also

496


be applied t o the reconstruction of the depositional mechanisms that play a role under extraglacial conditions. Deposition from debris flows occurs when the driving stress of gravity, responsible for mass emplacement, falls to below the strength of the debris, so that 'freezing' occurs. A so-called 'rigid plug', characteristic of debris flows, plays a n important role in both transport and deposition. The rigid plug is carried on top of the subaerial debris that is in the state of flow. The thickness of the plug is proportional to the strength, and inversely proportional to the density and to the angle of the slope. For instance, a decrease in slope angle results in increased thickness of the plug or decreased thickness of the flow; the flow stops moving if the thickness of the plug equals that of the flow (Johnson, 1970; Middleton and Hampton, 1976). There may b several subtypes of subaerial debris flows under extraglacial conditions. Some represent typical mudflows, with a high proportion of fine particles. The energy of such flows is very high and flowage will continue over long distances, even when the inclination is only slight. It is almost impossible for water to filter through such muddy flows and the shear strength is extremely low, so that the flow only comes t o rest if the terrain is very flat. If completely saturated with water, oarser material may be transported as a typical grainflow; occasionally such flows even show thixotropic behaviour, i.e. if there is a n admixture of fine-grained particles. The more typically extraglacial processes are related to the presence of permafrost in the substratum. Several processes may occur and Lewkowicz (1988) subdivided them into fast and slow processes. Both are usually small-scale surficial features that produce finely laminated deposits. Permafrost shows very slow (up to 3 mm per year) creep, more a continuous downslope deformation that takes place under stress conditions well below the stress value required for flowage, sliding or rupture (Williams, 1979). It develops due to the presence of liquid (water) in the otherwise frozen subsoil and is most effective in ice-rich soils on steep slopes. Morgenstern (1981) argued that permafrost creep is associated with localised shearing in ground-ice lenses, although in this case the process is slower than ice creep. The process is highly characteristic of extraglacial conditions but has not been studied in any detail. This lack of interest seems in part due to the relatively small effect: the extremely slow movement results in transport rates that are much less than one cubic metre per year, even for large masses involved. Frost creep is often taken to be a special form of gelifluction, but some authors (e.g., Lewkowicz, 1988) consider it a n entirely different process. It


497

results from the tendency of soil particles to be heaved during freezing in the direction normal to the (inclined) sedimentary surface, whereas the particles sink more or less vertically during periods of thaw. The net movement is thus downslope and the process is only active where there is a rapid alternation of soil freezing and thawing (Washburn, 1967). A comparable mechanism is involved in gelifluction (as defined by Washburn, 1980). Gelifluction is the glacigenic form of solifluction (as defined by Andersson, 19061, i.e. the slow downslope flowage of masses saturated with water. This process results in characteristic lobes, which are frequently found in Pleistocene deposits (Fig. 296). Loesses, with their typical grain-size distribution, seem exceptionally susceptible t o this process. Mechanically, needle-ice creep is a similar process but the graias rise to above the sedimentary surface during frost, as they are taken along by gradually growing needle-ice crystals. If the crystals melt during a thaw, the grains fall and come to rest at a location slightly lower than where they were before. Repetition of this process eventually results in a net downslope movement. While the process is well known from modern extraglacial areas, very small quantities of material are involved that are not likely to have any preservational potential whatsoever. In addition, it

Fig. 296. Characteristic lobe-shaped extraglacial solifluction deposit, separating loess units of different Vistulian ages, near Babor6w (Gtubczyce upland, S. Poland). Photograph: J. Cegta.

498


seems highly unlikely that 'fossil' sediments that have undergone this type of mass-transport will be identified. Skin flows are shallow mass flows over a plane, inclined surface. They can be considered as a kind of layer gliding and develop when ice in the active layer melts rapidly. The failure plane is usually the so-called frost table, i.e, the boundary between the frozen and the unfrozen soil (McRoberts and Morgenstern, 1973, 1974; French, 1976; Strangl et al., 1982). The moving mass is saturated with water formed by melting of ice lenses in the soil, but melting of a snow cover can also contribute t o water-saturated conditions (Shilts, 1978; Egginton, 1979). The process involved has been described by Middleton and Hampton (1976). Multiple retrogressive flow, rather than a single process, is a repetition of flowage phases at a specific place, usually as the result of seasonal melting of the active layer. Multiple retrogressive sliding (Brown et al., 1981) is comparable, but has a somewhat lower water content in the sediment. Active-layer sliding is a sliding process in which a large 'block' of the active layer is involved, with a minimum of internal disturbance (MacKay and Mathews, 1973). Blockfall is the downward movement of detached blocks. This process occurs especially in permafrosted areas where the walls of river valleys are undercut. It also takes place in ice-rich coastal escarpments where thawing and erosion result in the development of thermoerosional niches (Czudek and Demek, 1970; Brown et al., 1981). Rockfall is a process that results from mechanical weathering (usually frost weathering under the conditions described): joints are formed and the isolated blocks thus formed no longer have a fixed position, so that they tumble down as soon as their position becomes unstable. The descriptions of processes and the terminology just presented are based mainly on geomorphological studies. Taking the sedimentological point of view, it could be useful to group the various processes otherwise, i.e. in a category with extremely slow movement (permafrost creep), a category with slow movement ('gelifluction', comprising solifluction and needle-ice creep), and a category with fast movement (with debris flows, mudflows, sliding, slumping, blockfall and rockfall). Slopewash resulting from surficial currents of rain water, could possibly also be considered as a separate category.

The marine terminoglacial subenvironment

499

THE MARINE TERMINOGLACIAL SUBENVIRONMENT (II-D)AND ITS DEPOSITS During its flow downward, active ice may meet a sea or ocean. In this case, the glacial ice may reach into the marine environment, forming floating ice shelves, tidewater glacier tongues or ice-walled shores. The marine area covered by the uninterrupted sheet of glacial ice is termed the marine glacial environment (1-D),and the area t o its front is termed the marine terminoglacial subenvironment. This subenvironment is thus only present if glacial ice reaches the sea (see the discussion in the chapter on the marine proglacial subenvironment), so that mass flows originating in the glacial environment enter the sea immediately (without passing a continental terminoglacial zone). Icebergs can consequently be found in this subenvironment. The marine terminoglacial subenvironment is the most 'proximal' of the marine periglacial subenvironments (also see the table on p. 130). The boundary between the marine glacial environment and the marine terminoglacial subenvironment should be drawn along the imaginary line on the ice shelf whence mass movements may start that reach the area in front of the ice. As will be explained in the chapters following ,this there is no boundary between the marine terminoglacial and the marine proglacial or extraglacial subenvironment: the terminoglacial subenvironment passes directly into the non-glacigenic marine environment. The outer zone reached by mass movements of glacial origin. The marine terminoglacial subenvironment can be hundreds of kilometres wide and can nowadays be found around the Antarctic, Greenland, Baffin Island, Svalbard and off the coast of numerous glaciated regions such as Alaska. Several authors classify parts of continental slopes and abyssal plains with this subenvironment (Piper et al., 1973; Barrett, 1975; Molnia and Carlson, 1978; Powell, 1981; Elverhoi et al., 1983; C.H. Eyles et al., 1985). In principle, the boundary could be located between the marine terminoglacial subenvironment and the nonglacigenic marine subenvironment where the last icebergs disappear. However, the influence of sedimentation from these icebergs upon the marine sediment is so slight as to be untraceable after a travel distance of maximally a few hundred kilometres. It therefore seems justifiable t o locate the boundary with the non-glacigenic marine environment a t the

500


imaginary line that may be reached by turbidity currents originating from the marine glacial environment. This location is consistent with the definition of the continental terminoglacial subenvironment and, moreover, the area involved can then reasonably be assumed to be identical to the area in which dropstones from floating icebergs are found. The great width of this subenvironment results in sedimentation that differs considerably from place to place (Fig. 297), partly as a result of variations in direct particle supply, partly as a result of the decrease in size and frequency of debris-laden icebergs that occurs if the distance

Fig. 297. Model for marine terminoglacial sedimentation (based partly on N. Eyles, 1983; C.H. Eyles e t al., 1985; Brodzikowski and Van Loon, 1987). FIS = floating ice shelf; SF = submarine fan; S = slumping; UM = undermelting; R = redeposition; TC = turbidity current; CH = (submarine) channel; IRD = ice-raft deposition; DD = dump deposit; t r = traction. 1 = active ice; 2 = supraglacial channel; 3 = flow lines within the ice; 4 = crevasse; 5 = medial moraine; 6 = lateral moraine; 7 = nunatak; 8 = englacial debris; 9 = sealevel; 10 = marine glacial sedimentary surface; 11 = submarine moraines representing a previous grounding line of the ice; 12 = glacial marine mass-transport deposits (I-D-6-c); 13 = marine terminoglacial sedimentary surface; 14 = marine terminoglacial ice-raft deposits (11-D-1-e);15 = rain-out process; 16 = dump deposits; 17 = subglacial deposits; 18 = glaciotectonically deformed seafloor; 19 = hard-roek substratum.


501

from the ice increases. Sedimentological analyses of this subenvironment therefore require that it be subdivided in much more detaile. Several authors have made a rough distinction between proximal and distal zones (Molnia and Carlson, 1978; Elverhoi and Roaldset, 1983; Miall, 1983; Molnia, 1983b; C.H. Eyles et al., 1985; Dowdeswell, 1987), but these two zones differ only essentially in the average grain size of the sediments. At least two subzones can be distinguished in the most proximal zone. The first subzone is located at the very end of the floating ice, where mass transport of glacial material plays a dominant role and where wave action may affect bothglacial ice and sea bottom (Boulton and Deynoux, 1981; Anderson et al., 1982, 1983a,b; Anderson, 1983; Andrews and Matsch, 1983; Molnia, 1983; C.H. Eyles et al., 1985; N. Eyles and McCabe, 1989a). The second subzone is characterised by grounding ice masses in shallow seas (C.H. Eyles et al., 1985; McCabe and Eyles, 1988; N. Eyles and McCabe, 1989; N. Eyles et al., 1989). The somewhat less proximal zone can be subdivided into the subzone of the inner shelf and the subzone formed by fjords. The inner-shelf subzone (Barrett, 1975) occurs under open-marine conditions and in wide gulfs; the fjord subzone (Richards, 1976; Glasby, 1978; Powell, 1981; Elverhoi et al., 1983; Gilbert, 1983; Drewry, 1986) is very specific, being strongly influenced also by non-glacigenic or paraglacigenic agents, mostly through largescale mass-transport processes. The distal part of the marine terminoglacial subenvironment can be subdivided into the mid-shelf zone, the outer shelf zone, and the heads of fjords. Numerous models have been proposed to describe all, or some of these conditions. Most models emphasise the complex sedimentation pattern in this subenvironment. SEDIMENTATION PROCESSES IN THE MARINE TERMINOGLACIAL SUBENVIRONMENT The depositional processes seen under marine conditions resemble those occurring in large lakes (N. Eyles, 1983; C.H. Eyles et al., 1985). The main differences concern the scale and the intensity of the processes involved. Another difference from the processes in large lakes is that the thickness of glacigenic deposits formed under marine conditions clearly diminishes in an off-coast direction (Elverhoi et al., 1980,1983; Barrett et al., 1983), partly as a result of decreasing energy (Drewry and Cooper, 1981) whereas, in lakes, such deposits often show no pronounced thickness gradients in lakes. While the salinity of sea water is another aspect

502


(Lake and Walker, 19761, water circulation is often a less variable parameter under marine conditions (Gilbert, 1983; Drewry, 1986) than in lakes. Sea-ice seems a more important factor for the sedimentation pattern (Kovacs and Miller, 1974; Reimnitz et al., 1978) than does nonglacial ice in lakes. This is due t o the interaction between seasonal packice, fast ice and shearing. Acoustic methods have often been used to study the character of deposition in this subenvironment. This approach resulted in the distinction of several 'acoustic facies', which are interpreted as being characterised by specific depositional processes. Several sedimentation models have been proposed on this basis (Piper et al., 1983). Other methods have, however, also been applied in the course of various projects and numerous types of - = deposits from modern marine terminoglacial areas have been described in detail. Studies such as these have made major contributions t o the recognition of Pleistocene deposits and of older deposits formed under similar conditions (Elverhoi et al., 1980, 1983; Orheim and Elverhoi, 1981; Powell, 1981a, 1984; J.N.J. Visser, 1982, 1983a,b,c; Elverhoi and Roaldset, 1983; Gilbert, 1983; Molnia, 1983; McCabe et al., 1984, 1986, 1987; C.H. Eyles et al., 1985; Drewry, 1986; Dowdeswell, 1987; Visser and Loock, 1987; McCabe and Eyles, 1988; N. Eyles and McCabe, 1989). Most studies distinguish facies that, according t o the classification system used in this book, could be grouped into the marine terminoglacial melting-ice facies (11-D-11, the marine terminoglacial bottom-current facies (11-D-Z),the marine terminoglacial deltaic facies (11-D-3) and the marine terminoglacial mass-transport facies (II-D-6).

DEPOSITS OF THE MARINE TERMINOGLACIAL MELTING-ICE FACIES (II-D-1) The marine terminoglacial subenvironment is characterised by a shifting ice front and the frequent occurrence of ice blocks that float away from the main ice body as a result of calving. Such icebergs or ice-rafts give rise t o ice-raft deposits in the same way as do their counterparts in glacial lakes. The marine terminoglacial melting-ice facies has been studied in rocks of Quaternary, Tertiary, Palaeozoic and Precambrian age, distributed over all continents, but the marine melt-out processes are being studied, in addition, in Arctic and Antarctic regions. Interesting descriptions and analyses of these deposits are to be found in reports by Gadd (1971), Edwards (1978), Anderson et al. (1980a,b, 1984), Domack (1982, 1984),

Deposits of the marine terminoglacial melting-ice facies

503

Gilbert (1983), Miall (1983a), Visser (1983a), Eyles and Eyles (1984), Fortuin (1984) McCabe et al. (1984, 1987), Eyles et al. (1985a,b), Benn and Dawson (1987),Dowdeswell(l987)and Visser and Loock (1987). The intensive melt-out processes in the marine terminoglacial subenvironment induce a high sedimentation rate (Molnia, 198313) that gives rise t o deposits that are rather unstable, particularly if they are on an inclined part of the sea bottom, so that mass-transport processes occur frequently. McCabe et al. (1984) and C.H. Eyles et al. (1985b) have presented the most detailed models of this facies and the interfingering with other facies. Several authors have stressed that sedimentation that results from melt-out from icebergs is different from sedimentation resulting from the undermelting of an uninterrupted ice shelf (Wright and Anderson, 1982; Andrews and Matsch, 1983; Vorren et al., 1983; Anderson et al., 1984; C.H. Eyles et al., 1985; Visser and Hall, 1985). The differences are due partly to a gradual decrease in iceberg intensity in a direction away from the ice (Fig. 298) and partly to differences in supply and sedimentation

Fig. 298. Schematic distribution of marine terminoglacial ice rafts in front of a floating ice shelf. UMZ = undermelting zone (marine glacial); IRZ = ice-raft zone. 1 = active ice; 2 = flow lines within the ice; 3 = direction of ice movement; 4 = crevasse; 5 = wave erosion of ice; 6 = iceberg; 7 = dropstone; 8 = submarine channel; 9 = marine terminoglacial bottom-current deposit; 10 = sediment-covered marine terminoglacial tunnel-mouth deposit; 11 = marine terminoglacial tunnelmouth deposit under formation; 12 = undermelting i n a subglacial lake; 13 = lodgement till; 14 = marine terminoglacial ice-raft deposit; 15 = substratum.

504


rate (Fig. 299): undermelting of a n ice shelf results in deposition of all material that is set free at random from the ice, whereas icebergs may tumble around, producing dump deposits (Thomas and Summers, 1982, 1985) that are commonly coarser, because finer particles are partly washed away from the ice beforehand. Melting of icebergs and ice-rafts being the only process that contributes to sedimentation of particles set free from glacial ice, the only type of deposits distinguished in this facies is marine terminoglacial ice-raft deposits (II-D-1-e).

Marine terminoglacial ice-raft deposits (II-D-l-e) These deposits a r e characteristic of the marine terminoglacial subenvironment, even though they cover only about 20% of the seabottom (mass-transport deposits account for most of the other space). Nevertheless, ice-raft deposits are encountered almost throughout this subenvironment but are intercalated with the mass-transport deposits. Annual cycles sometimes appear in these deposits as lowering of the temperature during winter more or less stops the undermelting, whereas the higher water temperature during summer results in a relatively large amount of debris being set free. An average sedimentation rate of up to one millimetre per year is thus reached; this value is identical to that for undermelting of ice shelves (Elverhoi et al., 1980,1983). Icebergs are unstable bodies. They may tumble over if they have become unstable as a result of thcrmosubrosion, thus suddenly releasing masses of debris concentrated at their surface; such masses fall, forming dumpstones, often accompanied with deformation of the muddy seabottom (Thomas and Connell, 1985). Icebergs are also driven along by wind and sea currents (Fig. 300); entering a shallow part of a sea, they may drag along the bottom and bring back into suspension particles that had already settled; this process is accompanied by the formation of grounding structures and other types of glaciotectonic deformations (Dreimanis, 1979; Elverhoi et al., 1983; Vorren et al., 1983; Eyles and McCabe, 1989). I t was found that icebergs account not only for most of the deposition in this subenvironment, but also for much of the erosion (Vorren et al., 1983; Weeks et al., 1983; Barnes et al., 1984; Thomas and Connell, 1985; Drewry, 1986). There is also considerable interaction with pack ice and fast ice, particularly during winter. Accumulation of snow i n shear zones also plays a role in shallow seas (Barnes and Reimnitz, 1974; Reimnitz et al., 1978).

Deposits of the marine terminoglacial melting-ice facies

THERMAL REGIME Of GLACIAL SOURCE

LITHOLOGY

RATEIKa

T T

modern

mudidiamictirand

3750m

0 0 1 Ma

mudidmmlct

35m

Molnia and Carlron 119781

T T

0 0 1 Ma ^u 2 Ma

mud/diamict

2-3rn lm

Molnta and Carlwn 119781

mudldiamicthandr

Alaskan, proximal to ridewater ~ c margin e

T

modern

mudidiamict

MlOm

Powell 119811

Spirrbergen. proximal to tidewarer ice margin

T

modern

mudidiamcct

Alaskan. proximal 10 fiord head delta New Zealand. mid fiord

T

modem

mud

T T

modern

silty rand

0 84-4 3 n

Glaiby 11978)

Norway. m i d fiord

modern

mud

Spilrbeigen, ouier fmrd

T

0 6 - 2 55m 01-20m

Elverhoi el al., 119831

Coastal embaymenti

SHELF

inner shelf proximal 10 tidewater ice margins outer shelf outer shelf

ANTARCTICA CONTINENTAL SHELF AND UPPER SLOPE ABYSSAL PLAIN

REFERENCES

A~~

SETTING

GULF OF ALASKA: CONTINENTAL

505

Outer rhelf and u ~ o e rlom r outer shelf and upper dopf

< m e ( shelf

rediment-sIarved inner shelf adlacent to I C shelf ~ sediment starved inner shelf adlacent to ~ c eshelf

F

F T

~008Ma

mud

1lXOm

Elverhoi e1 a1 , 119831 Hostin and Burrell 119721

mudldiamict rnud/dlamict mudidiamlct

2-7cm

modern

Platter and Addicon 119761

lOOOm

25-5 Ma

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