Glacigenic Sediments

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.

Early literature and problems in terminology

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

Early literature and problems in terminology

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

Early literature and problems in terminology

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

Early literature and problems in terminology

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).

Early literature and problems in terminology

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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)

Depositional processes in the glacigenic environments

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

General characteristics of glacigenic sedimentation

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

Depositional processes in the glacigenic environments

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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.

Characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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

Characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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.

Characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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.

The influence of climate on glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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.

The influence of climate on glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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.

The influence of endogenic forces

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

General characteristics of glacigenic sedimentation

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.

The influence of endogenic forces

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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,

Facies interpretation

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

General characteristics of glacigenic sedimentation

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.

Facies interpretation

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

General characteristics of glacigenic sedimentation

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).

Facies interpretation

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

General characteristics of glacigenic sedimentation

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

Facies interpretation

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,

General characteristics of glacigenic sedimentation

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

Facies interpretation

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

General characteristics of glacigenic sedimentation

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

Facies interpretation

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

General characteristics of glacigenic sedimentation

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

General characteristics of glacigenic sedimentation

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.

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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.

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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.

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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-

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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.

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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.

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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.

Terminology and use of symbols

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

General characteristics of glacigenic sedimentation

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

The system of glacigenic depositional environments

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).

The system of glacigenic depositional environments

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

The system of glacigenic depositional environments

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).

Depositional sedimentary environments and their general model 89

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

The system of glacigenic depositional environments

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.

Depositional sedimentary environments and their general model 91

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

The system of glacigenic depositional environments

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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

+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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

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.

The glaciofluvial facies

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

Glacigenic facies and their characteristics

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.

The glaciofluvial facies

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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

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).

110

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

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

The glaciolacustrine facies

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

Glacigenic facies and their characteristics

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

The glaciolacustrine facies

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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

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

Glacigenic facies and their characteristics

'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

The mass-transport facies

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

Glacigenic facies and their characteristics

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).

The mass-transport facies

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

Glacigenic facies and their characteristics

(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

Tables of glacigenic subenvironments, facies and deposits

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)

Tables of glacigenic subenvironments, facies and deposits

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

Tables of glacigenic subenvironments, facies and deposits

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

The supraglacial subenvironment

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

The supraglacial subenvironment

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

The supraglacial subenvironment

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 supraglacial subenvironment

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).

140

The supraglacial subenvironment

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.

Deposits of the supraglacial melting-ice facies

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.

142

The supraglacial subenvironment

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).

Deposits of the supraglacial melting-ice facies

143

. 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).

144

The supraglacial subenvironment

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

Deposits of the supraglacial melting-ice facies

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.

The supraglacial subenvironment

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.

Deposits of the supraglacial melting-ice facies

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

The supraglacial subenvironment

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

Deposits of the supraglacial melting-ice facies

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.

150

The supraglacial subenvironment

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

151

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.

152

The supraglacial subenvironment

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.

Deposits of the supraglacial fluvial facies

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

The supraglacial subenvironment

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.

Deposits of the supraglacial fluvial facies

155

Lithofacies characteristics

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

The supraglacial subenvironment

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).

Deposits of the supraglacial fluvial facies

157

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.

158

The supraglacial subenvironment

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.

Deposits of the supraglacial fluvial facies

159

Lithofacies characteristics

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.

The supraglacial subenvironment

160

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

Deposits of the supraglacial fluvial facies

161

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.

162

The supraglacial subenvironment

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.

Deposits of the supraglacial fluvial facies

163

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.

164

The supraglacial subenvironment

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).

166

The supraglacial subenvironment

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.

Deposits of the supraglacial deltaic facies

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

168

The supraglacial subenvironment

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.

Deposits of the supraglacial deltaic facies

169

0 m

S

N 1 I

6m

2 I

3 6 m

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5 50m

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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.

170

The supraglacial subenvironment

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).

171

Deposits of the supraglacial deltaic facies

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

The supraglacial subenvironment

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.

Deposits of the supraglacial deltaic facies

173

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.

174

The supraglacial subenvironment

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.

Deposits of the supraglacial deltaic facies

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

The supraglacial subenvironment

A. stable terminal ice margin

B. retreating ice margin

. \ -

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delta foresets

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,

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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.

178

The supraglacial subenvironment

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).

Deposits of the supraglacial lacustrine facies

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

The supraglacial subenvironment

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.

Deposits of the supraglacial lacustrine facies

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

182

The supraglacial subenvironment

(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.

Deposits of the supraglacial lacustrine facies

183

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

The supraglacial subenvironment

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.

Deposits of the supraglacial lacustrine facies

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

The supraglacial subenvironment

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

The supraglacial subenvironment

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

The supraglacial subenvironment

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

The supraglacial subenvironment

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).

193

Deposits of the supraglacial mass-transport facies

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

The supraglacial subenvironment

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.

Deposits of the supraglacial mass-transport facies

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

The supraglacial subenvironment

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.

Deposits of the supraglacial mass-transport facies

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

198

The supraglacial subenvironment

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).

Deposits of the supraglacial mass-transport facies

199

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.

200

The supraglacial subenvironment

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.

Deposits of the supraglacial mass-transport facies

201

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.

202

The supraglacial subenvironment

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.

203

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.

204

The englacial subenvironment

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

205

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).

206

The englacial subenvironment

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

207

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.

208

The englacial subenvironment

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

Deposits of the englacial melting-ice facies

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;

210

The englacial subenvironment

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-

Deposits of the englacial melting-ice facies

--

-.._._ ;.-.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

212

The englacial subenvironment

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

213

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).

214

The englacial subenvironment

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.

Deposits of the englacial fluvial facies

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

216

The englacial subenvironment

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

Deposits of the englacial fluvial facies

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.

218

The englacial subenvironment

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

The englacial subenvironment

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

The subglacial subenvironment

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

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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

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, subglacia! glaciofluvial sediments

relative quartz enrichment in the remaining material

I

quartz enrichment

;

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

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abrasion

.~~. .. ...

quartz enrichment

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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).

The subglacial subenvironment

226

water flow

-4 0

0

p suspended load

0

bed load -.

0

C

o

O

P 0

<

stable b e d

,~

interstitial water pressure

traction zone

suspension

zone

Fig. 147. Particle transport by water and by sliding ice. Top left particle transport a t the base of running water. Bottom left: suspension and traction zones at the base of ice sliding on rock. Top right: suspension and traction zones with ice sliding over sediments (e.g., till) with low pore-water pressure, where effective normal pressures exceed the shear stress so that a stiff bed results. Bottom right development of thick traction zone of deforming bed materials where the effective normal pressure is exceeded by the glacier shear stress. Modified after Boulton (1975a) and N. Eyles and Menzies (1983).

pore water and the free running water a t the icehedrock interface (Menzies, 1982).Fluted moraines (Fig. 148-A),drumlins (Fig. 148-B),rock channels (Fig. 148-C) and rogen moraines (Fig. 148-D) may result from the geological processes that take place under active ice (Shilts et al., 1987; Aylsworth and Shilts, 1989; Zilliacus, 1990). An important aspect, particularly from a palaeogeographical point of view, is t h a t such subglacial features are distributed in space according to rules that are connected with ice thickness, velocity of ice movement, relief of the substratum and the lithological character of the icehedrock interface (Fig. 149). Models of the processes involved will be discussed in a subsequent section. A schematic model of the entire subglacial subenvironment, applicable t o the ablation zone of temperate glaciers and ice sheets with a soft-rock substratum, is presented in Figure 150. The complicated subglacial situation may be expressed in the geological record by a complex pattern of interfingering subglacial tills and meltwater deposits (the latter may be found in channels eroded in subglacial till deposited earlier). The subglacial streams may find their way dammed off by ice or sediment, with the result that stagnant pools are formed where fine-grained subglacial lacustrine sediments may settle.

227

Subglacial conditions under active ice

I

I

L

shear plans with debris-loaded 1 ~ above e

1-

I&;*&

I

Stre31 buildup

*hear plane 'take-on'and atrear release

shear plane Uhe~an'and IlreJI relaare

Fig. 148. Schematic representation of the formation of (A) fluted moraines, (B) drumlins, ( C ) rock channels and (D) rogen moraines as a result of subglacial processes taking place under active ice. Slightly modified after: Sugden and John (1976).

Other subglacial lakes may be found in the frontal zone of the ice if the substratum has a general inclination towards the ice, so that drainage is hampered. It is likely that subglacial sediments are reworked frequently, since the exact locations of net deposition and net erosion may change quickly in time and space. The erosional potential of temperate ice masses is shown both by the landscapes with U-shaped valleys after retreat of the ice, and, on a smaller scale, by the frequent occurrence of glacial striae in the hardrock substratum (Fig. 151). The combined deformational and erosional processes affect the freshly deposited sediments continuously, so that they are eventually partly or - more frequenly - totally destroyed. The preservational potential of subglacial deposits is thus generally low, but a

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outer edae 01 aiaclal advance

ice thickness

bedforw elongation ratio (a/b).

Y

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Subglacial conditions under active ice

229

Fig. 151. Glacial striae left by a modern Norwegian glacier. Photograph:J.D. de Jong.

favourable relief of the substratum and a favourable ice flow can result in preservation. The situation is much less complex in an ice mass with a cold thermal regime. The temperature at the icehbstraum interface usually varies between -3 and -18 "C, a temperature which is too low to allow pressure melting. The base of the cold ice is therefore continuously frozen t o the substratum (although there may be local spots with temperate conditions at the icehedrock interface, e.g., above fault zones) and deposition of debris is impossible under such conditions. The erosional capacity of polar ( = cold) glaciers and ice sheets is also low (Weertman, 1961), and the icehedrock contact is preserved during erosion: fragments of the bedrock are sheared and incorporated in the ice as glacial (glacitectonic) rafts. The shearing and incorporation processes are considered by several authors t o be a form of deformational glacial activity (e.g., Ruszczyiiska-Szenajch, 1976; Sugden and John, 1976). Temperate (or wet-based) glaciers have, according to N. Eyles (1983b), a temperature that is near the melting point of the ice throughout the ice body (-1t o -3 "C at the base). Pressure melting is the dominant mechanism operating under such conditions. This often results in intensive lodgement

230

The subglacial subenvironment

and in the formation of subglacial drainage networks. Drainage may take place through channels or in the form of thin films. Some authors (a.o., Sugden and John, 1976; Denton and Hughes, 1981; N. Eyles, 1983b) assume the existence of such drainage patterns under cold ice sheets also, although in a much less well developed form. The temperature at the icelsubstratum interface may reach values of about 0 "C. Melt-out processes are intensive under such conditions, and glaciofluvial material accumulates together with melt-out tills. Lodgement obviously does not occur in such a situation. It is obvious in view of all these considerations that cold-based glaciers have relatively little impact on subglacial deposition in a strict sense, particularly if the substratum consists of unlithified material. Still, the thermal conditions at the base of ice sheets vary considerably, both areally and temporary, so that deposition underneath a n ice mass presents a truly complex picture (cf. Boulton, 1972a; Sugden, 1977; Goodman et al., 1979; Denton and Hughes, 1981; N. Eyles and Menzies, 1983). SUBGLACIAL CONDITIONS UNDER PASSIVE ICE The subglacial conditions i n passive ice a r e commonly those of a temperate thermal regime. This implies that glaciofluvial and glaciolacustrine processes play a major role in sedimentation. Thermosubrosion also takes place within the ice and englacial and subglacial sediments thus may easily become mixed. This implies t h a t it is usually very difficult (if even possible) to distinguish between sub- and englacial sediments from passive ice masses. Analysis of such sediments is also hampered by the frequent occurrence of deformation structures, t h e formation of which is favoured by the water-saturated conditions t h a t frequently occur a t the bottom of dead-ice masses with a warm thermal regime. A cold thermal regime sometimes seems to occur i n dead-ice bodies also, and will not facilitate sedimentation other than through melt-out processes. The sediments characteristic of the subglacial subenvironment (tills) are therefore absent but there may be a gradual transition to flow and ablation tills that can develop along the margins of a dead-ice body. SEDIMENTATION PROCESSES IN THE SUBGLACIAL SUBENVIRONMENT The character and intensity of the subglacial depositional processes depend on the dynamics of the ice, on the amount of debris in the sub-

Deposits of the subglacial melting-ice facies

231

glacial zone and on thermosubrosion. This implies that relief (Crabtree, 1981) and thermal regime are the main parameters determining melt-out, glaciofluvial and glaciolacustrine processes. Still other processes with an important impact on the subglacial subenvironment are temperatureinduced melt-out, and lodgement, which is conditioned by the pressuremelting mechanism and by successive frictional retardation (against the substratum) of individual bedrock particles and/or aggregates. These processes will be dealt with in more detail in a later section. Sedimentation in subglacial channels (Galon, 1965) differs from 'normal' fluvial sedimentation in that there are closed tunnels, often entirely filled with water under pressure. Subglacial stagnant pools are quite common, so that streams embouching in these pools may form small deltaic complexes and most of the coarse material will be trapped in such pools. If the flow regime becomes sufficiently low (which may be the case if the stream is dammed off and finds another course), even the finest particles may settle out of suspension. One should keep in mind, however, that such glaciofluvial and such glaciolacustrine sediments are both very easily affected by the moving ice, so that, if they are preserved at all, they will usually be severely deformed. DEPOSITS OF THE SUBGLACIAL MELTING-ICE FACIES (1-C-1) This facies is characteristic of the subglacial subenvironment. The deposits grouped in this facies result from processes that include not only deposition but erosion and deformation of the substratum as well. The result of this complex combination of processes is a rapidly varying sedimentary cover and a characteristic subglacial relief (see, e.g., Menzies and Rose, 1987,1989). The main processes related to the melting of ice that contribute t o deposition of debris in the subglacial zone are lodgement, in situ melt-out, undermelting and pressure melting; lodgement, particularly as defined by Boulton (1972b, 1975a) and Kruger (1979, 1984), is to be considered as a combination of pressure melting, deformation and in situ melt-out. Each of the above processes results in a specific type of diamict. Several models have been developed in attempts to explain precisely the relationship between subglacial melt-out deposits, the subglacial environmental conditions and the processes involved in the genesis of these diamicts (Fig. 152). Among the various types of melt-out deposits, those formed in the subglacial subenvironment have been discussed most intensively and various proposals have been made touching their classification (e.g.,

232

The subglacial subenvironment

Elson, 1961; Dreimanis, 1969, 1976, 1980, 1988; Lavrushin, 1970a,b, 1976, 1980; Flint, 1971; Ruchina, 1973; Olszewski, 1974; Boulton, 1976, 1980; Drozdowski, 1979). Also, countless names have been applied to specific genetic types of subglacial tills or tills with specific characteristics, e.g.: comminution till (Elson, 1961), deformation till (Elson, 1961), lodgement till (Boulton, 1972b), subglacial flow till (Boulton, 1972a), basal till (Drozdowski, 1974; Olszewski, 1974), stoss side till (Hillefors, 1973), lee side till (Hillefors, 1973), shear till (Grube, 1979; Stephan, 1989), melt out subglacial till (Gibbard, 1980; Shaw, 1987b), shear zone till (Rappol, 1983), undermelt diamicton (Gravenor et al., 19841, till developed as the result of morainization (Lundqvist, 1985), dislodged till (Seret, 1985), undermelt till (Drozdowski, 1985), diagenetic till (Warren, 1987) and squeeze till (Dreimanis, 1989, after Goldthwait, 1980). The rather chaotic nomenclature for subglacial tills is not a n accident. The simultaneous activity of depositional, deformational and erosional processes, taking place both under active and under passive ice, result in a 1

STIFF BED

Y'

DEFORMING BED

EXTENT O F WET-BASE0 SLIDING

I M I N I M A L EROSION AREPL ABRASION ABRASIONIDEPOSITION ENTRAINMENTIOEPOSITION

J

J

J

J

J

4

J

J

J

/FLOW

LINE

a

SHIELD ROCKS

=THICK

SEOUENCE O F DIRTY ICE

M MELTED

M G MELTING

SEDIMENTARY ROCKS E EFFECTIVE N O R M A L PRESSURE

F FROZEN

FG FREEZING

C. GLACIER SHEAR STRESS

Fig. 152. Basal thermal regimes under a large ice sheet and corresponding zones of erosion and deposition. The inset shows deposition below a melted ice margin with abrasion or lodgement depending on the values for effective stress. Deformation of bed material occurs in areas of low effective stress. Note the destruction of ice in melting and melted bed areas and formation in freezing zones. Adapted from: N. Eyles and Menzies (1983).

I

Deposits of the subglacial melting-ice facies

233

wide variety of deposits that are difficult t o group according to a logical and practically applicable scheme. Deposition under active ice, for instance often results in complex deposits of mixed appearance which pass laterally over short distances into typical lodgement, squeeze, comminution, shear, deformation and incorporation tills. The question of which type is formed a t a specific place seems to be determined by local microenvironmental conditions. The type of material plays also a role; the mineral debris available may be derived from lodging, grain abrasion, shearing of the (hard-rock o r soft-rock) substratum or of previously deposited till, incorporation of single grains or sediment rafts, pressure melting, regelation, etc. Parameters such as relief, lithology of the substratum, availability of meltwater, ice dynamics, etc., will determine which aspects will predominate a t a specific place. The genetic classification of tills proposed by Dreimanis (1989) is based on the assumption that some of the subglacial tills are of primary origin, whereas others are secondary. The primary tills would result directly from subglacial ice melting, and the secondary tills would be reworked older tills. This seemingly logical approach is, however, not applicable in practice because there are inconsistencies due to gradual transitions. For instance, Dreimanis considers lodgement as a purely primary process, but comminution, incorporation, shear, deformation and squeeze till a r e grouped in the category of secondary tills. Moreover, Dreimanis (1989) apparently accepted the definition of lodgement given by Boulton (1972b, 1975a) and Kruger (1979, 1984), which implies that processes such as shearing, deformation and comminution must be regarded as included in lodgement. All these processes take place under more o r less similar conditions of confining pressure, which causes matrix of the diamicts to be deposited mainly due t o pressure melting. This process occurs during pressure shearing and is thus also related to incorporation tills. In addition, pressure-induced mass transport takes place during the formation of lodgement till. A practically applicable subdivision of subglacial melt-out deposits would therefore require either a very detailed and complex analytical approach or relatively rough grouping. The present authors prefer the latter approach. Only three types of subglacial melting-ice deposits are therefore distinguished here. These are: subglacial till complexes (I-C-1a), lodgement tills (1-C-1-c), basal tills (I-C-1-d) and subglacial ice-raft deposits (I-C-l-e).

234

The subglacial subenvironment

Subglacial till complexes (I-C-1-a) Most subglacial tills show a mixed character and should therefore be considered as till complexes. The characteristics of lodgement tills or basal tills are absent o r dominated by other characteristics. It is not always easy to interpretat the subglacial origin on the basis of lithofacies characteristics, but the relationship with other glacial sediments may reveal the nature of the till (Elson, 1961; Flint, 1971; Olszewski, 1974; Drozdowski, 1979; Minell, 1979; Rappol, 1983). Lithofacies characteristics

The complexes usually have a massive appearance. Matrix-supported tills are more common than clast-supported ones (Fig. 153); material finer than sand usually dominates, as a result of grain abrasion. The relatively heavy overburden of the ice results in most cases in strong consolidation. This may explain in part the common lack of

Fig. 153. Subglacial till complex from the Kleszcz6w g r a b e n ( c e n t r a l Poland). The till is predominantly sandy and clasts are scarce. The apparent lamination is not a result of stratification, but is due to a more or less horizontal s h e a r system.

Deposits of the subglacial melting-ice facies

235

primary structures, but foliation, secondary lamination and (sub)horizontal shear planes may be found. Some deformational structures (predominantly small-scale drag folds) may also be present. Small sediment rafts and sandy smudges are fairly common features; there are no reliable diagnostic characteristics of the subglacial origin.

Textural characteristics The strong consolidation due to the weight of a thick overlying ice mass commonly results in a schist-like appearance of the 'basal' part of the terminoglacial till complex, similar to the 'comminution till' of Elson (1961). There may also be a system of fissures in a dense network (if compared to overlying deposits). This system can be very complicated: four main sets can sometimes be distinguished. There have been many reports of a 3-dimensional fissure network (usually horizontal, vertical and inclined; see, e.g., Karczewski, 1966, 1967) and most authors agree that the horizontal system is the oldest one. The development of these systems is facilitated by the existence of a more or less pronounced lamination in the tills. The inclined fissures are considered of intermediate age. They are explained by Karczewski as being a result of pressure due t o the weight of the stagnant ice. The youngest (vertical) set is related t o epigenetic changes, a conclusion also reached by Wysokifiski (1967). The grains in these till complexes usually show a preferred orientation; this is due to either the direction of ice movement or compaction. The stoss sides of clasts are distinctly directed towards the source area of the ice, whereas the lee sides are situated in the opposite direction (Kruger, 1979; Dreimanis, 1989; O'Brien, 1989). The surfaces of pebbles and boulders, if present, show frequent striations and, sometimes, a polished appearance. On the other hand, crushing and abrasion of particles are very important subglacial processes, so that angular fragments are common. There is a convincing explanation by Haldorsen (1981) why larger particles in these tills tend t o be more rounded than the smaller particles. A bimodal grain-size distribution is common.

Occurrence The complexes usually form subcontinuous beds 0.5-1.5 m thick, and are found in the glacigenic sequence as the lowest unit in the part that represents glacial conditions. If lodgement and/or basal tills are also present, the complexes are usually found under the other subglacial tills.

236

The subglacial subenvironment

In such a case, the complexes most frequently form elongated lenses, within depressions in the substratum (Fig. 154). The lower contact is sharp, frequently even erosive, and the upper boundary is often marked by the disappearance of dense subhorizontal foliation and by the beginning of a semi-vertical joint system due to relaxation processes (Olszewski, 1974; Drozdowski, 197913).

Depositional mechanism All geological processes involved in subglacial melt-out probably play a part in the formation of the subglacial till complexes. Pressure melting, undermelting and in situ melt-out are possibly no less important than 'true' depositional mechanisms. Squeezing, lodging, shearing and folding are usually simultaneous; if they are not directly involved in the genesis of the complexes, they often affect the complexes shortly after their formation.

Fig. 154. Small-scale depression in the substratum (possibly formed by subglacial meltwater streams with strong separation eddies: cf. Sharpe and Shaw, 1989), filled with a subglacial till complex. The channel is about one metre deep and was incised in lacustrine sandy silts. The length of the till complex, as found during excavation of this quarry near Zary (western Poland) was a t least 70 m.

Deposits of the subglacial melting-ice facies

237

Lodgement tills (I-C-1-c) The term 'lodgement till' ('lodgment' is also found in the literature) was introduced by Upham (1892); two years later the term also appeared in a paper by Chamberlin (1894) who interpreted the genesis of this type of till as a result of subglacial push-up and dragging of sediment, and subsequent lodging in bedrock within the submarginal zone. This interpretation has been confirmed regularly ever since, though some minor reinterpretations appear t o be necessary, as much more detailed information becomes available. Lodgement till is, in any event, commonly considered as the most characteristic type of sediment found in glacigenic environments. It should be kept in mind, however, that characteristics may vary, mainly as a result of local conditions, although several authors state that the term 'lodgement till' should be reserved for a very specific, well defined type of till (see the relevant references in Dreimanis, 1989). On the other hand, Boulton (1972b, 1975a, 1982), Kriiger and Marcussen (1976), Kriiger (1979, 1984) and N. Eyles et al. (1982a) describe a wide range of characteristics and pathways of genesis. It thus appears that one might distinguish between lodgement types in the strict sense and in a wider sense. The former is a relatively rare type of deposit with very well specified lithological characteristics and formed under well defined conditions. The lodgements tills in a wider sense, which have also been termed 'paralodgement tills', show a much wider range of lithological characteristics and include more or less all types of melting-ice deposits formed subglacially under active-ice conditions, with the restriction that processes other than melt-out must have influenced the depositional process. In this book, the term 'lodgement till' will be used in this wider sense, which implies that several subtypes can be distinguished. Various authors have described such 'diverging' lodgement tills, often designating them by different names. According t o the present authors, not all types described in literature are truly diverging. Lodgement tills that show a truly different structure and those which are characterised by their occurrence under restricted conditions only will be described separately in the next subsection.

Lithofacies characteristics Lodgement tills are cohesive, matrix-supported diamicts (clasts are usually floating) with distinct shears on a micro- to meso-scale (Fig. 155). The mineralogical composition tends t o be fairly constant, particularly in the matrix; the larger clasts usually show an increase of local material

The subglacial subenvironment

238

direction of ice movement

4 major shear planes with slickensides / / major foliation planes ,@

contained clasts

Fig. 155. Diagram showing the shear structures in lodgement till. The structures are commonly considered to result from stress under moving ice. From: Boulton (1970b). Courtesy: International Glaciological Society.

towards the base of the till. The granolumetry depends largely on the transport distance and on the granulometry of the source material; a multimodal particle-size distribution is characteristic (Dreimanis, 1989). The structure appears massive at first glance, but closer examination often reveals the presence of structures with preferred orientation (Fig. 156), due to the stress applied by the moving ice mass. Four 'subtypes', representing the most important 'diverging' lodgement tills, are now described. Stoss-side tills (Hillefors, 1973) are formed against the stoss side of elevations in a hard-rock substratum (Fig. 157).This variety consequently shows extreme hardness, a very dense network of fissures and extremely low porosity. Sediments of this type form layers or lenses of only a few to maximally a few dozen centimetres in direct contact with the substratum. No stoss-side tills have so far been found in lowland areas, which is not unexpected as the unconsolidated substratum is insufficiently resistant

Deposits of the subglacial melting-ice facies

239

Fig. 156. Lodgement till of Wisconsinan age from the vicinity of the Farwell Canyon (Fraser Plateau, British Columbia). The strongly compressed diamict shows two subhorizontal shear directions. The vertical joints (lower part of the photograph) have a n early-diagenetic origin unrelated to the genesis of the deposit.

Fig. 157. The formation of stoss-side (right, black) and lee-side (left, inset) tills. From: Hillefors (1973). Courtesy: Bulletin of the Geological Institute, Uppsala University.

240

The subglacial subenvironment

and becomes deformed if the pressure is too high. Stoss-side tills can best be differentiated from the 'classical' lodgement till by their much higher frequency and the much greater length of their internal shear planes. The difference from the so-called shear till (Grube, 1979; Stephan, 1989) is the local occurrence and the lense-shaped body with concave shears of the stoss-side tills; shear tills form continuous sheets with well developed continuous, horizontal or even slightly convex shear planes. Local tills are made up only of local material, without any indication of transport over a greater distance. Two subtypes can be distinguished. The first, described by Elson (19611, Broster et al. (1979) and Broster and Dreimanis (1981), is formed on a hard-rock substratum and equivalent types are found in most contemporarily glaciated areas. Broken and crumbled fragments of the substratum material are found in a matrix of pulverised particles with the same petrographic composition. This sometimes gives the appearance of a strongly compacted breccia, which led to the introduction of the term 'exaration till' by R6iycki (1976). The second subtype of local tills, described by Schluchter (1977) and Stephan and Ehlers (1983), is found in areas with a n unconsolidated substratum. The presence of material from the local substratum is so dominant that the till exhibits all the characteristics (e.g., colour) of the substratum. The most important difference between local tills and 'classical' lodgement tills is that the former do not show distinct internal shearing, while subhorizontal joints due to overconsolidation are common. Incorporation tills have often been described (e.g., Virkkala, 1952; Weertman, 1961; Lundqvist, 1967; Moran, 1971; Boulton, 1972a; Drozdowski, 1974; Ruszczyiiska-Szenajch, 1976, 1981b) from areas that have been glaciated during the Pleistocene. The tills are characterised by the presence of material (unconsolidated substratum, a mixture of substratum and buried ice) that was embedded in the base of the ice mass. The resulting sediment is a mixture of clasts with sharp edges and finegrained melt-out material (Fig. 1581, a characteristic which makes them quite different from the 'classical' lodgement tills. These tills are usually massive and show a high degree of compaction and many joints. Deformation tills (Elson, 1961) are represented by small-scale (0.5-2 m) accumulations of fine-grained subglacial material in the form of folds, diapirs, breccias, and other deformations, as formed during subglacial deposition due t o melt-out. Elson considered the genesis of these tills to be a reaction of the substratum t o the differential stresses exerted by the tangential forces of the active ice. The extreme deformation (Fig. 159) is the most diagnostic feature of this subtype. Deformation tills have been described, though not always under this name, by Lundqvist (19671,

Deposits of the subglacial melting-ice facies

241

Fig. 158. Incorporation till (Strzelin area, Lower Silesia), with fragments of glaciolacustrine deposits embedded in the main diamict.

Dreimanis (1969), Ruchina (1973), Olszewski (1974), Lavrushin (1976, 1980), Van der Meer (1980), Morawski (1981) and Rappol (1983) among others. Some shear tills (in the sense used by Grube, 1979, and Stephan, 1989) might be considered to belong to this subtype of lodgement till. Textural characteristics

Lodgement tills, including the various subtypes, usually show specific textural characteristics, e.g., predominance of subangular to subrounded clasts, but this depends partly on the nature of the substratum. The exact rounding of the larger clasts depends mostly on the distance of transport within the basal zone of traction. This transport also results in the relatively frequent occurrence of bullet-shaped ('flat-iron', 'elongate pentagonal') clasts. Some of the clasts show a keel at their base. Glacial striae are often visible on the clasts, particularly if the clasts consists of hard, lithified fine-grained material. Striae usually occur parallel t o the a-axes of elongated clasts unless these clasts were lodged or transported by rolling; lodged clasts are striated parallel t o the direction

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The subglacial subenvironment

Fig. 159. Deformation till from t h e Kleszczow graben (central Poland). The deformation can be reconstructed to have been formed under conditions of active ice. Photograph: R. Gotowata.

of lodgement and can show impact marks on their upper and lower surfaces. Crushed, sheared and streaked-out faceted clasts are also common. A macrofabric (i.e., orientation of the larger clasts) with the a-axes parallel to the local direction of ice movement is characteristic, but a transverse orientation may be present occasionally as a result of folding and shearing. The development of the fabric depends on, among others, the grain-size distribution of the till material, the abundance of clasts and postdepositional modifications. The origin of the fabric in lodgement tills is generally complex. It may be produced by the lodging process (in the strict sense) or by deformation. A well developed fabric will result if both stress directions coincide. It is typical that the ah-planes dip slightly in a n upglacier direction if lodgement only was involved. The microfabric (i.e., the orientation of the particles in the matrix) is commonly not less developed than the macrofabric (Dreimanis, 1989). Most lodgement tills, and the poorly sorted matrix-supported varieties in particular, are overconsolidated. Bulk density, cohesion and seismic

Deposits of the subglacial melting-ice facies

243

velocity are usually high, whereas permeability is relatively low (cf. Minell, 1979; N. Eyles and Sladen, 1981; Haldorsen, 1981; N. Eyles et al., 1982b, 198713;Sladen and Wrigley, 1983). Occurrence

Lodgement tills often form the lowest truly glacial part of a glacial sequence, though subglacial till complexes, basal tills and - in rare cases subglacial flow tills may be found underneath. It is generally agreed that lodgement tills form widespread sheets. There are, however, contradictory analyses of the thickness of such tills. Several authors (e.g., Boulton, 1976c; Kriiger, 1979; Boulton and Deynoux, 1981; N. Eyles and Sladen, 1981; N. Eyles et al., 1982b), while they report considerable thicknesses, present detailed characteristics only for the lowermost 1-2 m and state that the higher parts are similar. Some authors (e.g., Sugden and John, 1976) doubt whether really thick lodgement tills can be formed during one depositional phase when the fact is taken into account that these tills are formed under active ice,

1

Zonation

of

basal ice

@

Sources o f mobility b e t w e e n sub-zones of traction and susoension

Lodgement by gram-by-grain

Fig. 160. General model of the lodgement process. a: distinction between debris in traction over the bed and debris in suspension above i t in the basal transport zone (this zone is characterised by the ease of transition between traction and suspension mechanisms of pressure-meltinghegelation and flow folding). b: grain-by-grain lodgement on rigid and soft substrata. c: lodgement of debris-rich masses. From: Boulton (1982).

244

The subglacial subenvironment

conditions under which erosion also plays a role. Several other authors emphasise that the lodging process finally results in a deposit that is the net result of a balance between deformation, erosion and deposition (Olszewski, 1974; Marcussen, 1975; Boulton and Paul, 1976; Dreimanis, 1976a; Kruger and Marcussen, 1976; N. Eyles and Slatt, 1977; Schluchter, 1977; Bjornbom, 1979; Drozdowski, 1979a; Miiller, 1983a). An elegant overview of the occurrence of lodgement tills was provided by Dreimanis (1989). Under conditions of advancing glaciers, the tills occur lodged over older (pre-advance) sediments and over 'glacitectonites', unless these have been eroded; under conditions of retreating glaciers, the tills form the lowermost depositional unit if the deposits related to the ice advance have been eroded, although they may be underlain locally by meltwater-channel deposits. The base of the lodgement till is important, as it represents the sliding base of the glacier. There is generally a plane contact with an unconsolidated substratum, although grooves may occur. A hard-rock substratum is usually abraded, particularly on stoss-sides of elevations. Since the sliding base of the glacier represents a large shear plane, sheared and strongly attenuated material from the substratum may be found as a thin basal layer, incorporated locally (by shearing) in the till material (see also the references in Dreimanis, 1989). Depositional mechanism All lodgement-type tills are formed under active ice by the simultaneous activity of several erosional, deformational and depositional processes. Boulton (1982) developed a model of the lodgement process (Fig. 160). The pure lodgement process, i.e. the process that produces lodgement tills in the strict sense, develops due to the tangential forces excerted by moving ice. Basal debris may be lodged grain by grain, possibly resulting in concentration of rock particles (Fig. lSl), but sorting need not take place (Hallet, 1981).Sheets of debris-rich ice may also become lodged, viz. if the force imparted by the ice flow is insufficient t o maintain forward motion; continuous lodgement on top of these sheets buries them (Boulton, 1970b, 1971; Lindner and Ruszczyiiska-Szenajch,1979).While such lenses of till, resulting from negative regelation (Lavrushin, 1976, 1980), are considered by various authors as lodgement tills in a wider sense, they should be termed 'basal tills' (I-C-1-d) according t o the classification applied in the present book. Deformation of already deposited lodgement till by glacial drag under the glacier sole often accompanies the accretional lodgement process

Deposits of the subglacial melting-ice facies

c

245

glacier sole

0

Fig. 161. Grain-by-grain lodgement. a-c: lodgement of a clast against the till matrix a s a result of ploughing through this matrix and of matrix consolidation. d-e: lodgement of a clast against an already lodged clast (note the pattern of deformation of the matrix). Slightly modified after Boulton (1975b).

(Boulton, 1979, 1982, 1987; Dreimanis, 1987, 19891, especially if dilatation has occurred. Additional microstructural changes take place by drainage through the dilatant system (Miiller, 1983a). The depositional mechanism and its influence on the final sediment cannot be observed directly. The process was interpreted by Boulton (1972a, 1975a) on the basis of theoretical considerations and of observations made in areas that are now glaciated and have a hard-rock substratum. A substratum of this type implies that the large amount of fine-grained material in the till presumably consists of grains that were formed by abrasion of larger clasts. Abrasion continues until grain sizes have been reached that are no longer liable to further destruction (Vivian, 1970; Boulton, 1974; Sugden and John, 1976). An unconsolidated substratum, as found in the European lowland areas that were glaciated during the Pleistocene, greatly complicates the interpretation of the depositional mechanism. It is difficult t o explain how a lodgement process can take place under such conditions. Indeed, it

246

The subglacial subenvironment

appears that subglacial tills in lowland areas are sometimes difficult t o distinguish from englacial melt-out tills, though the degree of compaction may present a solution t o this problem. Nevertheless, much more work is needed if the precise lodgement process is t o be understood. Further problems are encountered if the genesis of the various subtypes of lodgement tills is t o be explained. There are several models, but none that can explain all characteristics in detail. However, convincing explanations, at least for their genesis in general, have been provided for several subtypes. The development of stoss-side tills, for instance, has been described in detail (also see Fig. 157) by Hillefors (1973) and Boulton (1975a). These tills are interpreted as developing as a result of vertical accretion which has to take place when the moving ice meets a n obstacle (lee-side tills are forced for similar, but opposite, reasons in the 'shadow' of obstacles). The depositional process of stoss-side tills is thus mainly determined by pressure melting (if, however, the obstacle forms a steep, concave ridge, some of the basal ice may be frozen to the rock within the cavity, leaving a basal till on melt-out). Models for the incorporation of rafts or grains in a cold-base glacier have been presented by Weertman (1961) and Boulton (1972a). The formation of local tills (and exaration tills in particular) is explained by a model developed by Moran (1971). The general process of subglacial till deposition as a result of pressure melting or melt-out (due to heat supplied by either traction or the geothermal heat flux) occurs simultaneously with deformation of previously deposited material. Deformation is most commonly reflected by squeezing and shearing. The genesis of the resulting deformation tills has been described by a number of authors, among them Hoppe (1952), Okko (1955), Elson (1961), Kamb (1978), Boulton (1979), Schluchter (1983) and Seret (1985). Boulton (1979) also performed experiments in which the deformation was made visible.

Basal tills (I-C-1-d) The term 'basal till' as commonly used in the literature in its loose sense, frequently applies only to 'the lowest till in a glacial sequence'. In our opinion, the term should be restricted to those subglacial tills that are formed by a melt-out process only. Basal tills conforming to this description have been reported by, among others, Shaw (1971), Olszewski (1974) and Drozdowski (1979). Two subtypes of basal tills may be distinguished. The first is formed under conditions of active ice, when fragments of debris-rich basal ice are

Deposits of the subglacial melting-ice facies

247

sheared and frozen to the substratum in protected cavities. Subsequent melt-out may result in the deposition of basal till, simultaneous with the deposition of lodgement till on top of it (cf. Shulmeister, 1989b). The second type of subglacial till is formed under conditions of passive ice, where debris-rich ice melts and the debris set free accumulates without being affected by ice movement. It is evident that subglacial tills formed under passive ice can also be found in the terminoglacial subenvironment, where surficial or buried dead-ice masses are present. Basal tills from these two subenvironments seem almost impossible to distinguish but, in principle, tills formed in the terminoglacial subenvironment are not overlaid by a lodgement till (there is no active ice). One should keep in mind, however, that a lodgement till found on top of a basal till may have been formed during a phase of readvance of the ice mass.

Lithofacies characteristics These tills most commonly have a massive appearance, although vague lamination may be visible. This impression may be due partly to streaks of relatively sandy material, and partly also to the frequent occurrence of (sub)horizontal joints in the cohesive material. The density of the joints may vary over short distances but the network of joints is generally less distinct than in lodgement tills. The subglacial melt-out process is not really different from that taking place under englacial conditions and, as a result, basal tills have much in common with englacial melt-out complexes. They form relatively homogeneous mixtures of sand, silt and clay, with a varying number of floating pebbles (Fig. 162). Roiycki (1970) and Morawski (1981) described small deformation structures in basal tills, which they considered to be the result of glacidynamic disturbances. However, the disturbances are generally unimportant and debris banding present in the ice may be preserved after melting, although the sharp contacts tend to become more gradual as a result of fluidisation and other comparable processes. Dreimanis (1989) stated that debris banding may be destroyed by both the melt-out process itself and subsequent deformation. The till will in such a case have a more massive appearance. Basal tills may contain some clasts of the unconsolidated substratum. The grain-size distribution in these tills shows more variation than that in lodgement tills. Haldorsen (1981) mentions that subglacial tills are, as a rule, coarser than the related lodgement tills.

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The subglacial subenvironment

Fig. 162. Basal till from the Kleszcz6w graben (central Poland), formed under active ice. Note the subhorizontal joints, the floating gravel and the presence of more sandy streaks.

Textural characteristics The orientation of flattened boulders and pebbles is the most conspicuous characteristic of basal tills. The orientation is with the ah-planes either (sub)horizontal, or slightly imbricated (if a favourable concentration of flattened clasts is present). The orientation of the clasts is inherited from the active-ice stage where they were aligned according t o the glacial stress field (usually parallel t o the ice flow but perpendicular t o the flow direction if compressive stress conditions are present), but the more distinct orientation during that phase becomes somewhat weakened during the melt-out process under passive-ice conditions. Short transport or local origin of the clasts is testified to by the occurrence of material that could not have survived long-distance transport. A number of clasts show abrasion signs, but distinctly faceted or freshly broken clasts are relatively rare. Basal tills are usually overconsolidated, particularly if they are overlain by lodgement tills. The high degree of compaction is an important diagnostic feature. The strength of the material is generally much greater

Deposits of the subglacial melting-ice facies

249

for basal tills formed under conditions of active ice than for those formed under passive ice.

Occurrence Basal tills are often preserved in small, shallow depressions. They form then lenses or layers with a thickness of commonly less than half a metre if formed - below a lodgement till - underneath active ice. They may be thicker if formed under passive ice; the thickness then depends on the concentration of clasts in the zone that underwent melt-out. Basal tills could be considered to be the most common sediment formed during periods of deglaciation. They indeed form continuous layers where formed under passive ice but their overall volume is generally less than that of lodgement tills from the same glacierisation. Basal tills occur at the basal part of the truly glacial part of the glacigenic sequence, but alternation of phases with ice retreat and ice readvance often results in basal tills intercalated between (and interfingering with) lodgement tills. Depositional mechanism Basal tills are deposited as a result of debris being set free by subglacial melt-out at the base of stagnant ice or in a stagnant zone under active ice, as a result of slow melting at the base of a temperate ice mass, or through so-called negative regelation (Lavrushin, 1976, 1980). The main energy flow responsible for the melting is that provided by geothermal heat but friction due t o basal sliding also contributes (Drewry, 1976). The characteristics (horizontal and lateral grain-size distribution, position of the boulders) are affected by running meltwater during the melt-out process. As meltwater is an integral part of the melt-out process, the re-orientation of pebbles, the partial removal of the finest particles and comparable processes should be considered t o be part of the depositional mechanism.

Subglacial ice-raft deposits (I-C-1-e) Subglacial ice-raft deposits may be formed in water-filled depressions under an ice mass. Whether the deposits formed under such conditions should be termed 'subglacial lacustrine deposits' is a matter of debate as a truly lacustrine development seems exceptional. The relatively rare cases of subglacial deposition within stagnant water bodies consist mainly of

250

The subglacial subenvironment

flow tills (here considered to belong with the mass-transport facies) and melt-out deposits (in this case, termed 'subglacial ice-raft deposits'), which belong with the melting-ice facies. The subglacial ice-raft deposits consist of debris set free by undermelting of the ice. The particles fall into the water formed underneath and are embedded in the other mineral material that accumulates a t the bottom of the 'lake' (Fig. 163). The term 'subglacial ice-raft deposits' should thus be used exclusively for deposits that probably consist for more than 50%of undermelt material (cf. Gibbard, 1980; Gravenor et al., 1984; Dreimanis, 1989),so that material supplied by subglacial streams or mass movements plays only a minor role (cf. Boulton, 1982; Drewry, 1986). The melt-out origin of the deposits led several authors t o use terms such as 'subglacial melt-out diamicts', 'waterlain melt-out diamicts' and 'subglacial undermelt diamicts' (Lavrushin, 1976; Gibbard, 1980; Dreimanis, 1982b, 1989; Gravenor et al., 1984; Morawski, 1984) t o describe these deposits.

Fig. 163.Subglacial ice-raft deposits (upper left) from the Kleszczow graben (central Poland). Note the dropstones, also seen i n the underlying subglacial lacustrine complex.

Deposits of the subglacial melting-ice facies

251

Lithofacies characteristics These deposits usually appear massive, with floating clasts in a poorly sorted matrix. The lower boundary tends to be relatively sharp, whereas the upper part often passes gradually into other subglacial deposits. Load structures are a frequent finding. The deposits often show deformations. The deformations result partly from sudden pressure exerted by falling clasts and partly from instability of the material (reversed density gradient, inclined sedimentary surface, etc.). Reversed density gradients may be present if subglacial channels embouch in the stagnant water body and supply pulses of material with a granulometry different from that set free by undermelting, or if specific types of debris flows reached the subglacial lake. Textural characteristics A polymodal grain-size distribution is characteristic of these deposits. All grain sizes found in related subglacial deposits are present in comparable proportions (Gibbard, 1980; Dreimanis, 198210). The coarser clasts show the characteristics of subglacial and englacial clasts: they may vary from angular t o rounded, may include wedge- and bullet-shaped specimens, and glacial striae may be seen. The orientation of the larger clasts is variable. There may be no preferred orientation whatsoever, but the ah-planes may also show a preference for a (sub)horizontal positlon or follow the deformational patterns of the deposit itself. Occurrence The deposits occur in the glacigenic sequence a t the same level as lodgement tills and basal tills. Gibbard (1980) emphasised the transition into lodgement till (Fig. 164). They are mainly preserved as small sediment bodies in shallow depressions. Depositional mechanism By definition, the deposit is formed under temperate ice masses as a result of undermelting. It is unimportant whether active or passive ice is affected by the melt-out process, but the ratio between true melt-out material and other debris (e.g., supplied via meltwater streams) is usually higher under conditions of active ice.

252

The subglacial subenvironment

stratified till formed by debr's release

by basal melting and meltwarer activity

PLG

Fig. 164. Undermelting of ice and relationship between deposition of rain-out material in depressions and the occurrence of lodgement till. From: Gibbard (1980). Courtesy: Boreas.

DEPOSITS OF THE SUBGLACIAL FLUVIAL FACIES (I-C-2) The heat flow from the interior of the earth and various other factors may cause the lower parts of glaciers or ice sheets t o melt, This implies t h a t subglacial streams may arise in which fluvial deposits can accumulate. Additional and commonly even larger quantities of water are supplied by supraglacial and englacial water that percolates via crevasses and tunnels until i t arrives in the subglacial subenvironment. Such meltwater tunnels may end in the terminoglacial subenvironment where tunnel-mouth deposits may be formed either subaqueously or subaerially. This facies may thus form a gradual transition in the direction of terminoglacial facies (cf. Szupryczyfiski, 1965; Flint, 1971). The water-pressure regime in the subglacial fluvial facies remains rather puzzling. J.R.L. Allen (1971a) had suggested t h a t the water pressure is normal, just as in the englacial crevasses and meltwatertunnel facies. Other authors have argued for high hydrostatic pressures (Rothlisberger, 1972; Lister, 1973). Sharpe (1987b), Shaw (1988b) and Sharpe and Shaw (1989) have described the production of erosional bedforms in the hard-rock substratum (Fig. 165) by rapid, sediment-laden,

Deposits of the subglacial fluvial facies

253

turbulent subglacial meltwater flows with separation eddies. Hydraulic aspects of this so-called 'full-pipe flow' or 'tunnel flow' have been both investigated experimentally and much discussed (Durand, 1951, 1953; Guy et al., 1966; Acaroglu and Graf, 1968; Elliott and Gliddon, 1970; Loadwick, 1970; K.C. Wilson, 1970; Wilson and Brebner, 1971; McDonald and Vincent, 1972). Shaw (1971) considered the high hydrostatic pressures t o present a complicated problem as regards closed englacial and subglacial channels in active ice. He mentioned various delicate interrelations between the plastic strain in the ice mass and the thermal properties of the ice, between flowing water and its frictional property, etc. Englacial and subglacial crevasses and tunnels will be (partly) filled by sediments as a result of these interrelations and as a result of melting of the ice. Reconsidering the work of Shaw (1971),Sugden and John (1976) stated that the ''... interactions have not been adequately investigated, but they are certainly of fundamental importance for an understanding of the process of meltwater deposition..." (cf. Shaw, 1983; Shaw and Kvill, 1984). Subglacial meltwater-tunnel deposits (I-C-2-b) are the only type of deposits found in the subglacial fluvial facies.

Fig. 165. Hard-rock substratum sculped by subglacial meltwater flows. Cantley (Ontario).

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The subglacial subenvironment

Subglacial meltwater-tunnel deposits (I-C-2-b) Field studies indicate that these deposits form during stages of passive rather than active ice although, theoretically, they may still form during a n active stage (Szupryczyiiski, 1965) underneath temperate glaciers and ice sheets (N. Eyles, 1983d). Deposition under active ice is only temporary and the chance of the deposits of being preserved is small (cf. Boulton, 197213,1979).The situation is different if the subglacial channels are very large (Galon, 1965; Kozarski, 1966/1967;Liedtke, 1981; Ehlers, 1983b), as in this case deposits may be preserved in the deeply incised valleys. Data obtained from boreholes in the European lowlands indicate that even in this case most subglacial channel deposits are relatively small and occur between subglacial tills. Large and deep subglacial channels are usually erosional features, indicative of the hydrostatic pressure conditions during their genesis. These processes can only occur if the crevasse or channel bottom is inclined in the direction of the ice flow. However, these surfaces may have a n opposite direction, especially in the ablation zone, with two possible outcomes: (1)the complete space may become filled with water that will be subjected t o hydrostatic pressure due to the plastic strain in the ice; (2) open spaces may be preserved, especially a t the beginning of subglacial melting, but infiltration from higher locations and plastically moving ice will make these spaces disappear. Investigations carried out in Lower Silesia (e.g., Brodzikowski, 1982a, 1984,1987; Brodzikowski and Van Loon, 1983) showed that it is extremely difficult t o reconstruct inclined surfaces in the proximal zone. These surfaces are rare in more distal zones, but remain well recognisable. This is probably due to the hydrostatic pressure in subglacial channels, which prevents waters from flowing very fast. Therefore, most of the channels with water in a phreatic state are not filled with channel deposits only but more commonly contain subglacial tills and other deposits. This explains why these subglacial channel deposits are relatively scarce (cf. Brodzikowski and Van Loon, 1983). Nevertheless, stratified examples can be found, especially in erosion gullies within lodgement and basal tills (I-C-1c and I-C-1-d). Occurrence of the deposits in this position means that, after melting of the ice, they usually give rise to a much less marked geomorphological relief than do typical eskers that originate i n tunnel mouths (N. Eyles et al., 1982b). An abundance of new data on subglacial drainage has been presented during the last decade. Is has become clear, for instance, that some of the drumlin fields consist primarily of subglacial meltwater-tunnel deposits

Deposits of the subglacial fluvial facies

255

(Muller, 1974; Dardis and McCabe, 1983; Shaw, 1983; Dardis et al., 1984; Shaw and Kvill, 1984; Sharpe, 1987). Dardis and McCabe (1987) considered some of the subglacial fluvial deposits to be a result of sheetwash. Subglacial 'full-pipe' flow was discussed by W.H. Graf (1971), McDonald and Vincent (1972),Saunderson (1975,1977b) and Drewry (1986).

Lithofacies characteristics Subglacial meltwater-tunnel deposits show widely varying lithological and structural characteristics but usually differ from fluvial deposits formed under subaerial conditions (Saundersson, 1975, 197713; J.R.L. Allen, 1982). Experiments aimed at better understanding of the flowage mechanism in subglacial pipes (McDonald and Vincent, 1972) showed that, if the concentration of particles in the flow increases, pressureinduced upward flowage at a given discharge results in a succession of current ripples, dunes, dune planes, plane beds and once more dune planes. Antidunes are not formed because the standing waves required are suppressed by the ice roof. Scour-and-fill structures, parallel lamination and irregular crossbedding are commonly found. Stratification is usually distinct but is irregular, due to the intercalation of layers with pebbles and even cobbles between much finer layers. There also occur massive, graded, pebbly and laminated sandy successions that might be interpreted as a result of perhaps local sheetwash (Dardis and McCabe, 1987; Sharpe, 1987).

Textural characteristics The textural characteristics of these deposits are not truly different from those of the subaerial counterparts. A difference may be that subglacial fluvial transport takes place over what are usually relatively short distances so that the average rounding of the clasts is less well developed. The rounding of the original material, however, still is very important. Striations are common on the larger clasts. Orientation of the pebbles may be with the a-axes either parallel or perpendicular t o the current direction; imbrication is rare. Sorting of the material is relatively bad, particularly because of the rapidly changing energy conditions (Fig. 166).

Occurrence Subglacial meltwater-tunnel deposits are found in the lower part of the glacial section of glacigenic sequences, where they occur as elongated

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The subglacial subenvironment

Fig. 166. Subglacial meltwater-tunnel deposit from the Kleszczow graben (central Poland), showing bad sorting, rapid lateral and vertical grain-size changes and badly developed stratification.

lenses embedded in lodgement till, intercalated betweem basal and lodgement tills, or situated in depressions within the substratum (N. Eyles et al., 1982a; Shaw, 1983). As shown by Galon (1965), the deposits form complex patterns. Most deposits formed in what appears to be an ongoing channel are discontinuous, particularly if formed under active ice. Those formed under passive ice are generally more continuous, but need not necessarily show a distinct relationship with the previous direction of ice movement.

Depositional mechanism The main parameters that control the depositional mechanisms in subglacial meltwater tunnels are the diameter of the 'pipe', current velocity and sediment concentration. The importance of the relationship between grain size and current velocity for the transport mechanism in pipes was investigated by Newitt et al. (1955). These authors presented a diagram giving a hydrodynamic explanation of the actual deposits (cf. Saunderson, 1977).

Deposits of the subglacial deltaic facies

257

Several models have been proposed for the subglacial fluvial facies (by, e.g., Klatkowa, 1972; Banerjee and McDonald, 1975; Shaw and Kvill, 1984; Hanvey, 1989), i n the case of both active a n d passive ice. Saunderson (197713) stated t h a t sliding beds may be a diagnostic sedimentological characteristic of subglacial tunnel deposition. He assumed that, as proposed by Newitt et al. (19551, there are four principal transport regimes during sedimentation in pipes: (1)Stationary beds occur at low velocities. There is bedload under such conditions and some particles move by saltation. The major bedforms are ripples. (2)As velocity increases, the bedload may begin to move (slide), particularly if i t consists of coarse material. Under these conditions the bed moves as a n entity along the channel bottom. The movement may be continuous or not, depending on the exact flow regime and the sediment concentration. (3)Heterogeneous suspension of all particles takes place a t still higher velocities, but size sorting occurs so that the average particle size decreases from the bottom upward. (4)Homogeneous suspension without vertical grain-size differentiation characterises the highest velocities. Subglacial fluvial activity was discussed also by Shaw and Kvill(1984). DEPOSITS O F THE SUBGLACIAL DELTAIC FACIES (I-C-3) Subglacial deltaic sediments (Fig. 167) are rare because exceptional conditions are required: an ice sheet or glacier with a temperate thermal regime, a relatively large subglacial depression filled with more or less stagnant water, a subglacial sediment-laden meltwater stream embouching in the stagnant water, and negligible ice dynamics. This implies that such deposits are most likely to be found under passive ice near the ice front (cf. Gjessing, 1960; Brodzikowski and Van Loon, 1987). There are no reports in the literature suggesting that it might be useful to distinguish between various types of deposits within subglacial deltas. Only one type of deposits is therefore included here: the subglacial deltaic complexes (I-C-3-a),sometimes termed 'subglacial paradeltaic deposits'. Subglacial deltaic complexes (I-C-3-a) These complexes constitute the natural transition between subglacial meltwater-tunnel deposits and subglacial lacustrine complexes. They

258

The subglacial subenvironment

I

ablalion zone of active ice sheet

I

I(--

desinlegration of stagnant lobe (passive frontal zone)

I I i I

surlicial crevasses supraglacial

1 I

I

waterlaid subglaci subglacial lacustrine deposits

deltaic complex

Fig. 167. Relative position of subglacial deltaic sediments. Modified after Sugden and John (1976) on the basis of field evidence from the Kleszczow graben (central Poland).

have been encountered in the Kleszczow graben (central Poland), where tectonic lowering of the substratum provided the depressions required for stagnant water bodies under the Elsterian ice cover. This deposit had a longitudinal section of about 200 m and a thickness of up to 15 m.

Lithofacies a n d textural characteristics Sand predominates in most complexes but finer and coarser material is present as admixture. The stratification is as a rule complex, although the overall appearance may resemble that of lnormal' deltaic foresets with large-scale cross bedding. This overal picture is complicated by t h e occurrence of smaller structures such as scour-and-fill structures, parallel lamination, small-scale current ripples, ripple-drift cross-lamination and commonly abundant - deformation structures (Fig. 168), and by the frequent admixtures of undermelt material. Some parts of the complexes follow channels incised in the underlying glacigenic sediments by subglacial meltwater streams during periods of high energy. The lower part of such channels is commonly filled by gravels and/or massive sands. Mass-transported units that followed the channel pattern are also very common. The material in the complexes rarely shows good sorting. Large clasts, which are commonly found, may result from undermelting of the ice roof.

Deposits of the subglacial deltaic facies

259

Fig. 168. Subglacial deltaic complex of Elsterian age, developed in a subsiding part of the Kleszczow graben (central Poland). Shear planes give the - incorrect - impression of horizontal boundaries. Photograph: R. Gotowata.

Occurrence These complexes are found in the glacigenic sequence a t the same level of subglacial tills and fluvial deposits. They are usually embedded in tills, or are intercalated between basal or lodgement tills under, and englacial or supraglacial sediments on top of them. The complexes often lack the classical deltaic form because subglacial deltaic sedimentation depends strongly on the local relief of the substratum. The thickness is rarely greater than a few metres, but a complex with a n exceptional thickness of 15 m was found in the active Kleszczow graben in central Poland. The lower and upper boundaries of the complex are commonly sharp; lateral transitions into meltwater-tunnel deposits in a n upstream direction and into lacustrine deposits in a downstream direction are much more gradual and thus difficult t o define precisely.

260

The subglacial subenvironment

Depositional mechanism The complexes are formed because of a drop in the current velocity of sediment-laden subglacial streams, at the point where they embouch in bodies of stagnant water. The result is a gradually prograding delta, which generally moves relatively rapidly forward because the depth of the depression t o be filled tends to be limited under subglacial conditions. The inflowing, particle-rich currents have a distinctly hyperpycnal character, so that bottom currents dominate. This aspect also contributes to the rapidly prograding character of the delta. The resulting decrease in grain size in a down-current direction is often masked by a continuous supply of particles of all sizes falling from the ice above the water body as a result of undermelting. Gravity-induced mass movements following the slope of the depression also mask the grain-size differentiation in a longitudinal direction. DEPOSITS OF THE SUBGLACIAL LACUSTRINE FACIES (I-C-4) Widenings of subglacial channels may form in lakes with almost stagnant water. The plasticity of the ice (whether in a passive or in an active stage) makes it impossible for large subglacial cavities to persist very long. Most subglacial lacustrine deposits are thus formed under a relatively thin ice cover, which implies that they belong to the ablation zone. Melting of the ice results in both isolated stones falling from the roof (resulting in dropstones) and in masses that flow down (termed 'waterlain tills' by some authors). This implies that the subglacial lacustrine deposits commonly have intercalations of other deposits. It is generally impossible to establish the subglacial character of a glaciolacustrine sediment with any certainty, especially since the original lateral and vertical contacts of these deposits will usually be destroyed or deformed (e.g., during a recessional readvance of the glacier). According t o several studies (May, 1977; Dreimanis, 1979,1980,1982a; Huddart, 1983; Morawski, 1984) this type of deposit is more common in Pleistocene sediments than was previously supposed. The deposits can occasionally be recognised because they show fissures in patterns similar t o those in subglacial tills, as a result of overburdening by the ice. The specific fissuring seems difficult to explain otherwise than by overconsolidation; the fissures form much more regular patterns than those due to, e.g., shrinkage fissures in ablation tills (I-A-1-b). It seems acceptable t o interpret these lacustrine deposits as also being subglacial, especially

Deposits of the subglacial lacustrine facies

261

when such fissures are found in lacustrine sediments in the direct vicinity of other subglacial sediments. Subglacial lakes seldom seem to be larger than a few hundred metres in diameter. It therefore seems preferable from a practical point of view not to distinguish between various types of subglacial lacustrine deposits, but to group them all together as subglacial lacustrine complexes (I-C-4-a). Deposits of this type have often been described, e.g., by Gjessing (1960), Gibbard (19801, Morawski (1984) and Dreimanis (1989). It appears from these descriptions that subglacial lacustrine complexes are much more common than had so far been supposed.

Subglacial lacustrine complexes (I-C-4-a) The impact of the seasons is much less at the bottom of subglacial lakes than at the bottom of supra- or proglacial lakes. This implies that the graded layering (varves) is usually less well developed (a possible reason for not recognising these deposits as such); some cyclicity may be present but seems mostly due t o variations in the input of sediment-laden meltwater currents. The typical lacustrine characteristics of the complexes may be masked by the relatively frequent occurrence of dropstones (supplied by an undermelted ice roof) and reworked subglacial tills.

Lithofacies characteristics Most complexes show a high silt content; clay and fine sand usually play a minor role but may be important if present locally in the substratum. Small current ripples and current-induced parallel lamination may be visible in sandy parts (Fig. 169) indicating that the water need not be completely stagnant, even in subglacial lakes. Gibbard (1980) and Dreimanis (1982a) have detailed the lithological and structural characteristics. Graded bedding, current ripples and, particularly, parallel lamination are the most common structures, but they have usually undergone deformation due to ice dynamics and unstable density gradients within the sediments. These deformations make the deposits difficult to recognise. However, the abundance of dropstones and reworked diamicts is also characteristic of this assemblage of deposits.

Textural characteristics The textural characteristics are not or almost not different from those of other subglacial sediments (sorting, for instance, is usually not very good;

262

The subglacial subenvironment

Fig. 169. Subglacial lacustrine complex from the Kleszczow graben (central Poland) with a sandy lower part (with cross-bedding) and a finer upper part (with isolated sandy ripples). Photograph: R. Gotowala.

striae occur on many clasts). However, the orientation of the larger clasts seems t o be random, being largely determined by the mechanical characteristics of the sediment underneath. 0 cc u rrence

The complexes can be formed under temperate ice masses only. Their precise location depends primarily on the relief of the substratum and on the ice dynamics. The deposits are therefore found in the glacigenic sequence in the part representing ice retreat rather than in the part formed during ice advance. The complexes form lenses that are usually elongated, with an irregular boundary, and are often overlain by englacial or supraglacial deposits. Lodgement till may also overlie the subglacial lacustrine deposits but, in this case, it represents a phase of ice re-advance. Ice re-advances are the rule rather than the exception. Most subglacial deltaic complexes have therefore been deformed severely. The preservational potential for complexes with almost no deformation is relatively

Deposits of the subglacial mass-transport facies

263

high in areas where tectonic subsidence takes place during the glaciation. The complexes, which are rarely more than 3 m thick, may become much thicker under such conditions: a complex found in the (active) Kleszczdw graben had a thickness of ten metres.

Depositional mechanism Several depositional mechanisms play a role, the most important being related to sediment supply by subglacial streams. The sediment is usually supplied in pulses as a result of fluctuations in the undermelting of the debris-laden ice base and because of changes in the subglacial drainage pat tern. Streams embouching in a subglacial lake leave most of the debris they carry in deltaic complexes or comparable fans, but the remaining streams are heavy enough to form undercurrents that spread particles througout the stagnant water body. Even the finest particles can settle when these undercurrents have slowed down sufficiently. The 'normal' grain-size distribution (relatively coarse near the source, increasingly finer in a down-current direction) is disturbed by sevarl processes, among them material falling from the ice roof and by influxes of mass-transported sediment.

DEPOSITS FROM THE SUBGLACIAL MASS-TRANSPORT FACIES (I-C-6) Subglacial conditions, particularly a n irregular topography, watersaturated sediments and instable density gradients within the deposits because of poor sorting or intercalations of layers with different grain sizes, favour the development of mass-transport processes. Reworking in this facies takes place exclusively under subaqueous conditions. Detailed studies have been devoted t o these deposits, although reworked diamicts have attracted most interest. Such reworked (and deformed) subglacial diamicts were termed 'secondary tills' by Dreimanis (1989). He included i n this group the subglacial squeeze tills of Goldthwait (1980), the shear tills of Grube (1979) and Stephan (1989), the dislodged till of Seret (1985), the squeeze flow till of Dreimanis (1989) and the subglacial flow till of Boulton (1980a) and many others. If this group - which we term 'subglacial mass-transport deposits' - is defined as consisting of material that has been reworked subglacially mainly as a result of gravity-induced mass-transport processes, i t should then include the lee-side tills of

264

The subglacial subenvironment

Hillefors (1973), the waterlain (subglacial) tills of Dreimanis (1989) and the watermorainic (subglacial) deposits of Morawski (1984). Subglacial mass-transport deposits (I-C-6-c) include all the varieties mentioned above and are thus the exclusive representative of the subglacial mass-transport facies.

Subglacial mass-transport deposits (I-C-6-c) Not all deposits formed subglacially as a result of mass-transport processes should be called 'flow tills': many show no till-like grain-size distribution. In addition, the term 'till' implies direct deposition by ice. These requirements need not be fulfilled (subglacial lakes may contain finegrained, varve-like turbidites), but diamicts a r e not only t h e most frequent but also the most obvious and the most spectacular representatives.

Lithofacies characteristics Most of these deposits have a diamictic granulometry, with matrix-supported types being more common than clast-supported ones . The lower contacts of mass-transported sediments a r e sharp, sometimes even erosive. Grooves and other sole marks may be found a t the base, especially if the mass flow contained relatively large clasts. The upper contact may be either sharp or gradual. Sharp upper contacts are often a result of sudden (vertical) changes in grain size; such sharp litliological boundaries are often reflected in the frequent occurrence of deformations (e.g., load casts) and undulating contact planes. More gradual upper contacts are commonly represented by planar, undeformed surfaces. The grain-size distribution of the deposits depends primarily on the material supplied and much less on the proximal or distal position (transport is rarely over a sufficient distance to result in lateral grain-size differentiation). The sedimentary structures present in the deposit depend on the depositional mechanism. The most common structures are parallel lamination, flowage and fluidisation structures, but the mass-transport process may also result in apparently structureless deposits.

Textural characteristics Sorting is commonly bad unless turbidity currents are involved or if wellsorted sediments (e.g., lacustrine deposits) have been reworked. Grain

Deposits of the subglacial mass-transport facies

265

sizes range from clay t o cobbles or even boulders, giving many deposits a diamict character (this explains the commonly used term 'flow till'). The larger clasts are often striated and usually show a widely varied lithology, thus resembling the clast composition seen in other subglacial sediments. The orientation of the larger clasts is, if they are sufficiently flattened, usually with the a/b-planes parallel t o the bedding o r - if deformation has taken place - parallel to the deformation structure, but random orientation can also occur. Occurrence The deposits can be found in the field in combination with subglacial fluvial, deltaic and lacustrine deposits. They have the same relationships in the glacigenic sequence. If formed during ice advance, they can also be found between the substratum and the lodgement till and, if formed during ice retreat, they may cover a lodgement till. Most of the deposits are found as layers of distinctly different character within subglacial lacustrine complexes. In situations such as these they may have the same granulometry as in lacustrine deposits, viz. if lacustrine material has been reworked (Fig. 1701, which is especially prone to occur along the slopes of the basin or if the substratum is unstable. Depositional mechanism Gravenor et al. (1984) proposed a systematic classification for diamict-like glacigenic deposits resulting from gravity-induced mass-transport deposits (also see Drewry, 1986). These authors distinguished three main types: subglacial debris flows, subglacial slurry flows and subglacial chaotic debris flows. The approach followed by Gravenor C.S. was not based on common sedimentological concepts. It is probably justified t o state t h a t all the various subaquatic mass-transport processes described by several authors can play a role. In this approach, the precise mechanism involved depends mainly on the ratio between mineral debris and water, and includes subaqueous rockfall (not a true mixture of particles and water), slumping (water-saturated sediment), mudflow (liquid limit has been passed) and turbidity current (relatively low concentration of sediment in the water). The latter process may result in deposits that resemble varves (Lajtaj, 1967; Banerjee, 1973; Nystuen, 1976) but, as season-dependent varves are not likely to occur under subglacial conditions, an incorrect interpretation is unlikely.

266

The subglacial subenvironment

Several workers have emphasised that cavities between a n ice mass are easily filled by mass-transport processes. Boulton (1982) presented a scheme in which the most important processes were visualised.

Fig. 170. Subglacial lacustrine deposits deformed during mass transport processes that were induced by subsidence of the substratum i n the Kleszczdw graben (central Poland). The height of the wall shown is approximately 6.5 m. Photograph: R. Gotowata.

The marine glacial environment

267

THE MARINE GLACIAL ENVIRONMENT (I-D) AND ITS DEPOSITS This environment (see the Table on p. 129) includes the entire area occupied by a n uninterrupted shelf of 'eternal' glacial ice (also see p. 90). These shelves of glacial ice represent the tongues of continental glaciers or land-ice caps (cf. Charlesworth, 1957; Flint, 1971; Embleton and King, 1975; Sugden and John, 1976; N. Eyles, 1983b; Drewry, 1986). Continental (or island) glaciers t h a t extend into the sea lower the water temperature, thus favouring freezing of sea water. It is therefore not uncommon to find sea-ice around glaciated islands (e.g., parts of Svalbard) or along the shores of glaciated land (Antarctica). Season-induced freezing of seawater results in a clearly shifting boundary of the sea-ice (Reimnitz and Kempema, 1987), but this is of no significance for the boundary of the marine glacial environment, which is based on the much less fluctuating extent of the glacial ice. The boundary of the marine glacial environment could be drawn at the average position of the margin of the glacial ice during summer. One should keep in mind, however, that it is not always possible to distinguish between sea-ice and glacial ice, particularly if the surface is covered by snow. Moreover, the ice extent during the summer depends on various factors including atmospheric and water temperature, the occurrence of supercooling during autumn storms (Reimnitz et al., 1987) and erosion by seawater (Reimnitz and Kempema, 1987). Such floating ice shelves, surrounded by sea-ice, were fairly common during the Pleistocene (also see Thomas, 1979). The equilibrium line between the accumulation and the ablation zones of glacial ice was situated far from the shore in several parts of Europe and North America and several islands that were covered by glaciers were surrounded by a wide shelf of glacial ice and an even wider zone of sea-ice (Charlesworth, 1957). Sediments from both the glacial and periglacial marine environments are usually termed 'glaciomarine sediments' (R.D. Miller, 1973; Drewry and Cooper, 1981; Elverhoi and Roaldset, 1981; R. Gilbert, 1982b; Nelson, 1982; Kaye, 1984; Benn and Dawson, 1987; N. Eyles and McCabe, 1989a,b). Various other names are in use, however, e.g., 'glacimarine' (Mackiewicz, 1983; Powell, 1984 among others), 'glacial marine' (Carey and Ahmad,l961; Frakes and Crowell, 1975; Domack e t al., 1980;

268

The marine glacial environment

Domack, 1982; Kravitz, 1982; Stauffer and Peng, 1984) and 'glacialmarine' (Kravitz, 1983; Molnia, 1983a; Powell, 1983a). Most authors, when they deal with glaciomarine sediments, do not distinguish clearly between the marine glacial and the marine periglacial environment - if they do make this distinction at all. A number of authors group ice shelves, tidewater ice tongues and grounding ice together as masses that can be termed 'glaciomarine' (Anderson et al., l979,1980a,b, 1982, 1983a,b; Clark et al., 1980; Orheim and Elverhoi, 1981; Thomas, 1981; McKelvey, 1982; Andrews and Matsch, 1983; Domack, 1983; Elverhoi et al., 1983; Molnia, 1983b,c; Vorren et al., 1983; C.H. Eyles et al., 1985b; Sharpe, 1989). However, the non-glacial ice formations have as common feature only that they are affected by undermelting, which does not seem important enough to be considered as a criterion for classification as a separate environment. Other authors do not distinguish a marine glacial environment at all, the main argument in this case being that, under marine conditions, all ice would represent ice-contact conditions, so that there would only be a terminoglacial (periglacial) situation (Domack, 1982; McCabe, 1986; McCabe et al., 1986; McCabe and Eyles, 1988; N. Eyles et al., 1989). However, the uninterrupted nature of ice shelves, while it is different from the nature of the terminoglacial subenvironment, is characteristic of the glacial environment. The marine glacial environment, as the present authors apply the term, has such a restricted variability that there is no reason t o distinguish between (supra-,en- and subglacial) subenvironments. Whatever processes may occur on top of, within or at the base of the permanent cover of glacial ice, only a few of the types of sediment accumulations that are formed have a chance of surviving. These sediments consist largely of particles that settle after undermelting (by thermosubrosion) then become embedded in 'regular' marine deposits where they sometimes constitute such a small fraction that they are hardly recognisable (French, 1976; Pye, 1977; Washburn, 1979); there are also tunnel-mouth deposits if subglacial channels embouch in the sea at the grounding line of the ice. Furthermore, inclined surfaces may favour mass-transport processes that result in mass-transport deposits that show several similarities with those formed under marine terminoglacial conditions.

Sedimentation processes in the marine glacial environment

269

SEDIMENTATION PROCESSES IN THE MARINE GLACIAL ENVIRONMENT The only relevant glacigenic sedimentary processes t h a t occur i n t h i s environment are, a s stated above, melting of ice and surging of meltwater from subglacial tunnel mouths (Fig. 171). The depositional conditions t h u s resemble those which occur on a lake bottom under glacigenic conditions (cf. Gibbard, 1980; Boulton and Deynoux, 1981; C.H. Eyles e t al., 1985b; Dreimanis, 1989). The deposits resulting from marine undermelting, however, resemble in some respects those of the (continental) terminoglacial subenvironment

Fig. 171. Sedimentary facies a t the inner grounding zone of a floating ice shelf, according to Drewry (1986). Undermelting and deposition in front of tunnel mouths are the most important depositional processes. Courtesy: Edward Arnold (London).

270

The marine glacial environment

more than they do those of the (continental) subglacial subenvironment (May, 1977; Gibbard, 1980; N. Eyles et al., 1982b; Dreimanis, 1982a, 1983; McCabe, 1986; McCabe et al., 1986, 1987; Brodzikowski and Van Loon, 1987; McCabe and Eyles, 1988; N. Eyles et al., 1989). This is due t o the fact that subglacial conditions are exceptional (often overpressurised water flows, a relatively thin layer of meltwater, movement of the ice mass over the substratum, etc.). Furthermore, marine glacial ice, while it contains debris that was embedded after glacial erosion of the substratum, also contains minor amounts of other particles such as wind-transported material (Reimnitz and Kempema, 19871, although individual 'foreign' grains commonly show predominant signs of other transport mechanisms (Hodel et al., 1987). An aeolian supply such as this is uncommon in the continental glacial environment (although there may be considerable concentrations of wind-blown material in specific areas: see, e.g., Kempema et al., 1989). I t is somewhat less exceptional to find predominantly siltand clay-sized particles in marine glacial ice (Reimnitz and Barnes, 1974) exactly as is seen in continental terminoglacial situations. There are also reports of considerable transport of fine-grained sediments by ice rafting (Barnes et al., 1982), and slush ice also accounts for some sediment transport as indicated by experiments carried out by several investigators, among them Clayton et al. (1990). Marine glacial ice may leave distinct traces. Local grounding of ice masses may result in the deformation and erosion of the substratum (Reimnitz and Kempema, 1982, 1984). Furthermore, Reimnitz and Kempema (1983) described 'strudel scours', craters in the sea floor as large as 25 m wide and 6 m deep, that were excavated by vertical drainage flow during the yearly spring flooding of vast reaches of shorefast ice t h a t surrounds arctic deltas. Surficially deformed marine sediments with floating 'erratics' may also be a clear indication of a marine glacial genesis (also see Reimnitz et al., 1978,1988). If the features just described are taken into account it will not be surprising that three facies can be distinguished in this environment, viz. the marine melting-ice facies (1-D-l), the marine glacial bottom-current facies (I-D-2) and the marine glacial mass-transport facies (I-D-6).

DEPOSITS OF THE MARINE GLACIAL MELTING-ICE FACIES (1-D-1) Undermelting occurs wherever glacial ice is found under marine conditions, implying that debris embedded in the ice is gradually set free. In the case of a n uninterrupted ice shelf such as exists in the marine glacial

Deposits of the marine glacial melting-ice facies

27 1

environment, there is no such thing as overturning icebergs. The exclusively glacial sedimentation thus occurs in the form of rain-out (Fralick and Miall, 1989). No other relevant mechanism being active, it is justified to group all deposits in this facies as marine glacial melt-out complexes (I-D-1-a).

Marine glacial melt-out complexes (I-D-1-a) The complexes are well comparable with lacustrine ice-raft deposits but with a n incomparably larger extent and thickness. Relatively few Pleistocene or younger deposits of this type are in exposures where they can be studied, so that most of our present knowledge of these features is based on observations made in outcrops of older (lithified, sometimes metamorphosed) deposits. Modern marine glacial environments, however, supplied the data reviewed byMolnia and Carlson (1978), Elverhoi et al. (1980,1983), Powell (1981) and C.H. Eyles et al. (1985). One problem is that these deposits often show no - or only vague - stratification, further complicating the analysis of lithified examples. In addition, irregularities t h a t arose in the sedimentary surface during deposition (for instance as a result of uneven debris concentrations within the melting ice shelf) may easily result in reworking over short or greater distances. The presence of reworking is often difficult to prove because of random orientation of the larger clasts or, with lithified material, reorientation as a result of compaction or pressure. This sometimes makes questionable the reliability of analyses of ancient marine glacial melt-out complexes. There are reports, for instance, of such deposits being several kilometres thick. Seen in the light of Pleistocene or present-day conditions, such interpretations are, to say the least, debatable: there are no such equivalents known from Quaternary times.

Lithofacies characteristics It is typical that the deposits form diamicts, the ratio between larger and smaller particles being dependent on the source material. There may be very large boulders but a matrix-supported diamict with irregularly distributed pebbles and coarser clasts is the more common find. The lithology resembles that of subglacial tills in many respects. The grain-size distribution shows no distinct lateral tendencies: only the material that happens to be present in the melting ice determines the local granulometry. This implies that there may be irregular patches or zones with relatively small or large average grain sizes.

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The marine glacial environment

Primary sedimentary structures are usually absent, but some deformation structures may be found. The most common of these structures is an 'impact crater' where a large clast fell on the soft substratum (Fig. 172); the water-saturated nature of the sediments usually results in the prompt disappearance of such a crater. Reversed density gradients, resulting e.g., from an interval with deposition of a relatively large number of clasts after an interval with predominantly silt and clay sedimentation, may result in load casts, flame structures and associated deformations. Slumping of these deposits definitely does occur, but there are few descriptions available of unambiguous mass transport of material that forms part of these deposits.

Textural characteristics The fabric of the larger clasts appears t o vary widely. Lithified examples cannot be assumed to give a reliable impression because the clasts most probably underwent reorientation in the course of time. Relatively little is

Fig. 172. Dropstone in glacial marine sediments formed in the former Champlain Sea (Ontario, Canada).

Deposits of the marine glacial melting-ice facies

273

known about the fabric in modern deposits, partly because the samples studied are usually not large enough to present a proper picture of the fabric, partly because the sampling procedures almost inevitably disturb the original fabric. Some data were obtained with geophysical methods, but the most reliable data are therefore based on observations of nonlithified, Pleistocene examples that have now become exposed due, e.g., t o upheaval resulting from isostatic compensation. The Pleistocene deposits studied by the authors show no distinct fabric, although several flattened clasts seem to prefer a n orientation with the a/b-plane more or less parallel t o the substratum but truly random fabrics were also occasionally found. A relatively large number of clasts usually show glacial striae, but this is not necessarily the case. If glacial ice is gradually pushed further into the sea, the basal part will be that first affected by undermelting. The originally englacial part, which contains relatively few striated clasts, may undergo undermelting further from the coast, so that - in principle - a zone with few striated clasts might be found around a zone with more striated clasts. The authors are, however, not aware of studies aimed at this kind of analysis. Occurrence

The complexes form continuous sheets that may extend over very large areas. Many of these deposits, however, are of only limited extent, e.g., where a glacier extends into the sea. The thickness of the deposits depends on the concentration of debris in the ice, the melting rate of the ice and the time span involved. A number of the deposits are less than 10 m thick but thicknesses of up to 100 m are not uncommon. As mentioned above, thicknesses of several kilometres are purported to exist but there is still much debate whether these interpretations are entirely correct (there may be tectonic doubling that is hard to recognise). In the glacigenic sequence, these complexes are rarely found on top of, or underlying continental glacigenic sediments. This is completely understandable as a rise of the sea level accompanies ice retreat, so that no glacial deposits can be formed on top of the marine sediments. On the other hand, the fact that lowering of the sea level takes place during periods of glaciation implies that the depositional basin retreats away from the ice, so that no marine deposits will be found on top of glacial deposits. It must therefore be deduced t h a t the vertical succession of marine and continental glacial deposits represents deposition during subsequent cycles of glacierisation.

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The marine glacial environment

Depositional mechanisms Undermelting of a floating shelf of glacial ice is the most important (if not the exclusive) depositional mechanism. Material is gradually set free and sinks t o the bottom. As the water under an ice shelf is usually relatively quiet, the finer particles do not need an very long period of time to settle, so that both fine and coarse particles are continuously accumulating. Reworking, may have a n effect on the material, but plays a minor role.

DEPOSITS OF THE MARINE GLACIAL BOTTOM-CURRENT FACIES (I-D-2) Several models have been proposed for deposition by current activity underneath a floating ice shelf (C.H. Eyles et al., 1975; J.N.J. Visser, 1982; Domack, 1983; Miall, 1983; also see Drewry, 1986). These models were developed on the basis of theoretical considerations rather t h a n based on field data because direct observations are impossible to make in modern marine glacial environments. In addition, the few data obtained from modern environments are highly controversial (Barrett, 1975; Kellog et al., 1979), so that field data used for sedimentation models are derived mostly from ancient (lithified) rocks (Young, 1978; Miall, 1983). This raises a n additional problem, as the glacial marine nature of the ancient rocks is not always unambiguous. Several of the models seem fairly reliable in spite of the above considerations, especially because there are no sound reasons t o assume that t o assumethis facies (and its deposits) will differ greatly from the corresponding facies (and its deposits) in the marine terminoglacial subenvironment (11-D);only rain-out material will much more frequently be found intercalated in the former facies. It is also possible that subaqueous fans are more common under marine glacial conditions (cf. Rust, 1988; Sharpe, 1988) than under marine terminoglacial conditions, where such fans are more easily affected by other processes. On the basis of all the considerations just mentioned, it seems only logical to distinguish this facies as a separate one, with one type of deposit: the marine glacial tunnel-mouth deposits (11-D-2-c).

Marine glacial tunnel-mouth deposits (11-D-2-c) The ancient sediments now interpreted as marine glacial tunnel-mouth deposits are formed as a result of subaqueous meltwater surges from the

Deposits of the marine glacial bottom-current facies

275

ice, more or less at the grounding line, i.e. at the border between continental and marine glacial environments. A phase of deglaciation will result in the most frequent occurrence of meltwater surges, so that the marine tunnel-mouth deposits will usually be formed during a phase of ice retreat. The gradually retreating tunnel mouth, often accompanied by a more or less equal backward shift of the grounding line, results in the formation of submarine ridges that are very well comparable to continental eskers. Subaqueous sanders will be formed in front of the tunnel mouth, usually in the form of fans. The resemblance with the esker beads formed under continental conditions has been stressed by Sharpe (1988), who also provided a n interesting model of the depositional mechanism and the internal structure of these bodies (Fig. 173). Lithofacies characteristics The deposits are coarse-grained, particularly so a t the base. Material of boulder and pebble size is the most common type found. There is gradual fining towards the top and the massive, gravelly lower part may thus pass into stratified gravelly sands or pure sands in the upper part.

0

c o a r s e ( i r d ) and fine (flocculated) sediment

ice front,

overflow

Fig. 173. Tentative depositional model for deposition in subglacial tunnel mouths during ice retreat. Slightly modified from Sharpe (1988).

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The marine glacial environment

The sandy units frequently show (sub)horizontal stratification, but massive units and units with tabular cross-bedding are also present. The thickness of the cross-bedded units tends to decrease in a n upward direction, which suggests gradually diminishing energy levels. This interpretation is supported by the occurrence of small-scale ripples, and sometimes also of mud drapes, in the topmost parts of some of these deposits. The lithofacies characteristics do not seem very different from those of similar deposits formed under marine and continental terminoglacial conditions. More details are provided in the relevant sections of the chapters on these two subenvironments.

Textural characteristics Sorting of these deposits, like that of most 'proximal' fan-like deposits, is fairly bad, particularly a t the base (Fig. 174). Gradually better sorting is found in the upward direction, but, except for some drapes, no well sorted units are found.

Fig. 174. Marine glacial tunnel-mouth deposit of Pleistocene age from Ontario (Canada).

Deposits of the marine glacial bottom-current facies

277

The clasts in the basal part show no preferred orientation, because deposition during the surges of meltwater was usually too fast. Imbrication or a preferred (sub)horizontal position of the flattened clasts can however be found in the units where 'regular' current activity played a role (as evidenced by, for instance, tabular cross-bedding). Neither the larger clasts nor the grains show specific textural surface characteristics, but striations are commonly found. This can be ascribed to the predominantly subglacial transport (at least in the last stage of transport) of the material involved. Occurrence The deposits are found at the grounding line of the glacial ice. The length of the deposits thus depends largely on the rate of retreat of the grounding line, whereas the width depends on the curvature of the subglacial channel that is gradually 'exposed'. The tunnel-mouth deposits interfinger with marine glacial melt-out complexes (1-D-1-a), or what is left of them after removal of the finest particles (by bottom currents or tidal processes). The deposits also interfinger with marine glacial mass-transport deposits (I-D-6-c),which consist for a large part of tunnel-mouth deposits that became unstable, for instance as a result of reversed density gradients or collapse (Sharpe, 1988). Depositional mechanism The depositional mechanism has been described in detail in the section devoted to the depositional mechanism of tunnel-mouth deposits in terminoglacial lakes. Much information has also been provided by Saundersson (1977b, 19821, Shaw and Kvill (1984),Sharpe, 1988) and Shaw et al. (1989). The most important factor regarding deposition is the occurrence of meltwater surges, which results in a sudden, strong increase in energy level, followed by a somewhat slower, though still fairly fast, drop of the energy level. The period of decreasing energy accounts for the upward fining of the material. Successive surges may result in the repeated erosion of the top part of the last deposit. This results in a final predominance of coarse-grained material that represents only part of the original deposition.

278

The marine glacial environment

DEPOSITS OF THE MARINE GLACIAL MASS-TRANSPORT FACIES (I-D-6) This facies is characterised by subaqueous slumps, slides, debris flows, turbidity currents and associated processes that may occur at the bottom underneath a floating ice shelf. While the processes are largely a consequence of the irregular topography of the sea bottom, also of glacial sedimentation, they tend to even out the height differences by seeking for the lowest places. As all process occur - by definition - under subaqueous conditions, only one group of deposits can be distinguished: the marine glacial masstransport deposits (I-C-6-c). Marine glacial mass-transport deposits (I-D-6-c) The submarine mass-transport processes tend t o be rather complex and t o pass into one another, as explained in models proposed by Stanley (1981), Stanley and Maldonado (1981) and Wright and Anderson (1982). It is predominantly the substratum characteristics, including inclination, that decide whether there will occur cohesive debris flow, mass flow, grain flow, fluidised flow or whatever else. According t o most authors, the majority of processes should be considered as a mixture of the various possible flow types. These circumstances do not really differ from those prevailing under glaciolacustrine conditions, especially those in the terminoglacial and proglacial subenvironments.

Lithofacies characteristics Most of these deposits have a diamictic character (Fig. 175), but can nevertheless be easily distinguished from the rain-out deposits that usually surround them. The shape and internal structures (flow lobes, folds, shadows in matrix-supported units) reveal the direction of transport. There are also frequent deformation slabs, particularly in sandy or siltykandy material that has undergone slumping. Flow lobes may also deform the substratum during transport and shortly after deposition.

Textural characteristics Sorting is usually bad. Deposits other than turbidites usually lack a preferential orientation of clasts, except for the basal and the top parts.

Deposits of the marine glacial mass-transport facies

279

Fig. 175. Precambrian diamicts (Poudingue de Donville), possibly representing glacial diamicts slumped down under marine glacial conditions (Donville beach, France).

There are no diagnostic textural characteristics, as these deposits are very much like those formed in glacigenic lakes or under marine terminoglacial or proglacial conditions. Rounding of the particles depends on the parent material. Glacial striae being common in subglacial debris, such striae are also found on many clasts in the mass-transport deposits. However, striae need not be present if the material involved consists predominantly of englacial or supraglacial material set free by undermelting of the floating ice shelf,. Occurrence These deposits are found throughout the marine glacial environment, but in very variable concentrations: some successions consist mainly of masstransported material, whereas there may also be areas (particularly if somewhat elevated) that lack such deposits. The thickest successions are commonly found at the base of fans formed in front of tunnel mouths near the grounding line. The thickness of individual layers may range from less than one centimetre to several metres, although a thickness of more than three metres

280

The marine glacial environment

seems t o be exceptional. A succession of such deposits on top of each other may easily reach tens of metres; some Precambrian successions of well over one hundred metres thickness have also been interpreted as resulting from submarine glacial mass-transport processes.

Depositional mechanisms Slumping, passing gradually into complex mass flowage, has been reported t o be the most important of these mechanisms. It is common that such flows change into high-concentration and then into low-concentration turbidity currents, even though successions consisting of well developed turbidites with the characteristic A-E intervals (Bouma, 1962) have not been described in the literature. More details are presented in the relevant sections of the chapters on the continental and marine terminoglacial subenvironments.

The terminoglacial subenvironment

281

THE CONTINENTAL TERMINOGLACIAL SUBENVIRONMENT (II-A)AND ITS DEPOSITS This subenvironment (see p. 128) has been given many names, among others 'transitional zone' and 'marginal zone' (Embleton and King, 1977; Edwards, 1978; N. Eyles, 1983b), but also 'ice-marginal zone' (Shaw, 1972a, 1975a; N. Eyles and Slatt, 1977; Cohen, 1983; Kozarski and Kasprzak, 1987), 'glaciomarginal zone' (Kurimo, 1980), 'ice-contact zone' (N. Eyles et al., 1989) and 'glacier-contact zone' (Boulton, 1986). It is a subenvironment that forms the frontal and marginal parts of the ice cover (Fig. 176), from the (imaginary) line connecting the points situated most inwardly from the ice cover, from which sediments of glacial origin may reach, via a mass-transport process, the area in front of the ice, to the (imaginary) line connecting the farthest points that may be reached by mass-transported sediments that developed in the glacial environment. Several authors group this subenvironment with the glacial environment (particularly because it is partly covered by ice), but processes occurring

Fig. 176. Terminoglacial subenvironment in front of the Werenskiold glacier (Svalbard). Photograph: J. Cegta.

282

The terminoglacial subenvironment

just alongside the ice masses dominate the depositional pattern and it therefore seems more justified t o include this subenvironment in the periglacial environment. The terminoglacial depositional zone is narrow as compared to the adjoining glacial and periglacial subenvironments. It nevertheless shows a wide variety of morphological forms and of sedimentological processes. These two parameters influence each other, but the determining factor for most processes is the amount and changeability of meltwater and debris carried along by the meltwater streams. The extremely high energy gradients that are present in most parts of this subenvironment are the most conspicuous characteristics. The prevailing conditions result in six facies being frequently present (Fig. 177): the terminoglacial melting-ice facies (11-A-l), the terminoglacial fluvial facies (II-A-2), the terminoglacial deltaic facies (II-A-3), the terminoglacial lacustrine facies (II-A-4), the terminoglacial aeolian facies (11-A-5) and the terminoglacial masstransport facies (II-A-6). The amount of meltwater formed by ablation does not generally depend to any great extent on seasonal changes or dayhight alternation. The supply t o the foreland may nevertheless be quite irregular, mainly because much water can be trapped in pools, tunnels, etc. in the various glacial subenvironments. The presence of unduly large amounts of water stored in these areas may make barriers collapse suddenly, giving rise t o heavy floods in the foreland, even in the absence of specific meteorological reasons. Such floodings may of course leave traces much more distinct than their relatively rare occurrence warrants; this is mainly due t o the sudden strong increase in the energetic level, affecting the substratum and morphological forms that are considered as obstacles.

SEDIMENTATION PROCESSES IN THE TERMINOGLACIAL SUBENVIRONMENT Sedimentation in meltwater streams is the most common form of deposition in this subenvironment. There is also accumulation of debris as a result of melt-out processes and sedimentation in lakes between or just in front of the ice masses. The lacustrine sediments may comprise some floating pebbles derived from ice rafts in the lake that are undergoing insolation and thermosubrosion. The lacustrine sediments can be intercalated with subaqueous mass-flow deposits formed as a result of sliding of supraglacially accumulated material. Such mass movements

Sedimentation processes in the terminoglacial subenvironment

terminoglacial

,

283

surficial meltwater outflow

fluvial facies

,

terminoglacial subaerial mass-transport deposits (11-A-6-a)

tunnel-mouth deposits

subglacial melting-ice facies

(II-A- 2 - C )

terminoglacial s ubenvi ronment (II-A) A

(I-c-1)

I I

/

I

/

\

I‘

,

’ \

,, / \ ,

‘

-

\

\ / I ‘

c

4

\

I

terminaglacial subaquatic lacustrine bottomsets fan (II- A - ~ - c )

terrninoglacial tunnel-mout)l deposits (II-A- 2 -c)

f

subglacial melting-ice facies

(I-c-1)

Fig. 177. Schematic model of the terminoglacial subenvironment under continental conditions. Top: subaerial facies and deposits. Below: subaqueous ( = lacustrine) facies a n d deposits.

284

The terminoglacial subenvironment

parts may also form depositional sites for wind-blown material. The fluvial sedimentation processes are characterised by a n alternation of relatively short periods of high energy levels, resulting in erosion and transport of material by meltwater streams - followed by downstream deposition - and longer periods with a lower energy level. It appears t h a t the high-energy periods are responsible for most of the net sedimentation, possibly because the deposits formed during low-energy periods are relatively easily eroded during a subsequent phase of high-energy currents that at supply much fresh material. The morphological consequences are comparable: most forms result from high-energy processes that reshape the forms that were built up during the longer intervals of quietness. The high-energy processes (strong meltwater currents, etc.) have a strong impact, not only because of the erosive power, but also because of the large amount of debris that they bring along. Consequently, the often small depositional basins i n the terminoglacial subenvironment are short-lived and show frequent lateral and vertical facies changes. Isolated ice bodies, resulting from an irregular ablation pattern a t the ice front, are also affected by erosion but sedimentation also takes place on their top surface, thus changing the originally subaerial ice masses into so-called buried dead-ice bodies. The sediments formed on top of such bodies might be considered as supraglacial, but the present authors feel it is more appropriate t o consider them as terminoglacial, on the basis of their position relative to the main ice mass. The buried ice masses gradually melt away if the climate ameliorates; the resulting cavities are commonly filled by collapse of the overlying sediment. Deterioration of the climate, followed by re-advance of the ice, may result in glaciotectonic push of the buried dead-ice, but the dead-ice may also become embedded in the 'fresh' ice masses.

DEPOSITS

O F THE TERMINOGLACIAL MELTING-ICE FACIES (II-A-1) The tills in the terminoglacial subenvironment are formed mainly along the frontline of the main, continuous ice sheet and tills formed by melting of dead-ice masses. Ablation tills and flow tills are found just in front of the ice mass and are formed as a result of processes taking place in the glacial environment. It nevertheless seems justified to assume that, a s a rule, these types belong to the general category of terminoglacial till complexes (I-A-1-a).

Deposits of the terminoglacial melting-ice facies

285

The category of 'true' terminoglacial tills also includes terminoglacial ice-raft deposits (II-A-1-e) formed by melting of icebergs o r ice-rafts in terminoglacial lakes (cf. Boulton, 1972b; N. Eyles, 1979; C.H. Eyles et al., 1985a,b; Drewry, 1987).

Terminoglacial till complexes (II-A-1-a) The size of dead-ice blocks can vary from barely visible t o many hundreds of metres (or even kilometres) in diameter. This has important implications for the deposits that are formed in connection with these ice bodies: supraglacial, englacial and subglacial melting, combined or not with the occurrence of flow tills, may form originally well distinguishable deposits if large and thick ice masses are melting but become fully mixed if the ice masses involved were originally small, or had mostly melted away. All these tills are treated here as a single type of deposit - terminoglacial till complexes (11-A-1-a)- because of the essential feature that these tills - if originally present - eventually become mixed due to the natural process of continuous melting. These complexes may be formed under either subaerial or subaqueous conditions (Dreimanis, 1989). By definition, the subglacially formed parts of terminoglacial till complexes are comparable to the basal tills (I-C-1-d) discussed earlier. If dead-ice blocks are melting, the detritus is concentrated to form a till that may show foliation and orientated clasts. The englacial contributions to the terminoglacial tills are formed within the ice body by gradual melting of debris-laden ice masses. They resemble the englacial melt-out till complexes (I-B-1-a) of the englacial subenvironment in nearly all aspects. Surficial melting results in 'supraglacial' material: ablation till and flow till. Both types may be changed into a transitional type due t o secondary processes such as frost action. Drozdowski (1979a) named this type 'ablation till in situ', but the term coined by Boulton (19711, 'paraautochthonic flow till', seems more justified in this case (cf. Klajnert, 1966, 1978; Klatkowa, 1972,1982; Nalewajko, 1982). Lithofacies characteristics The above will have made it clear that it is very difficult to define the 'normal' lithological characteristics of terminoglacial till complexes. Much depends on the extent and thickness of the melting ice mass, on the main agent that induces melting (there is a large supraglacial contribution from solar irradiation; a dense network of englacial tunnels and crevasses

286

The terminoglacial subenvironment

facilitates englacial melting; a high geothermal heat flux favours subglacial deposition of debris), and on the position (buried or surficial ice) but various other factors play also a role. The combination of all these factors may lead to a lithology t h a t is fairly similar t o that of a clearly different type of till (except lodgement till), but a mixture of characteristics is more common (Fig. 178). This complicates recognition in the field. An additional complication may be that the specific conditions, which commonly include a water-saturated subsoil, easily result in disturbances such as flowage structures, waterescape structures and load casts. The most common lithology is a variable matrix (clay to coarse sand) with floating pebbles and even boulders which may be several metres in diameter. There may be some more or less horizontal bands of a relatively large extent that seem enriched in coarse clasts (the 'sveg till' of Lundqvist, 1969b).The finer the matrix, the better observable (in some cases) is the orientation of the larger clasts (Glen et al., 1957). Other tills form a relatively homogeneous mixture of sands, silts and clays with floating pebbles that show a clear orientation. The erratics in some of these tills

Fig. 178. Terminoglacial melt-out complex from the Jaroszow zone (SW Poland). Note the extremely variable matrix and apparent stratification. Height of photo: 6 m.

Deposits of the terminoglacial melting-ice facies

287

occur most frequently a t the contacts between the 'basal' and the 'englacial' contributions. The presence of irregular lenses and layers of sandy and gravelly material is characteristic of the 'englacial' components.

Textural characteristics The textural characteristics of the complexes are determined mainly by the ablation process, thus by the thermal regime in the ablation zone. Typically temperate conditions result in the predominance of gravityinduced mass movements that form flow tills; percolation of meltwater through the deposits usually results in the total absence of preferred orientations of single grains. Cold and dry conditions, on the other hand, result in sublimation and evaporation, processes which leave intact the existing fabric in the mineral debris. Cold conditions also tend to result in the formation of tills t h a t have 'englacial' or 'supraglacial' textural characteristics. The texture of the 'englacial till' is characterised by the occurrence of foliation, the occurrence of sandy or even gravelly lenses and the orientation of the larger clasts. Moreover, the ratio between coarse clasts and matrix is higher than that in the upper tills. The 'supraglacial' parts most likely have textural characteristics originally identical to those in the supraglacial subenvironment. The sudden changes that occur in the terminoglacial subenvironment are the reason for the destruction of most of the original characteristics of these sediments by mass transport or by meltwater streams.

Occurrence By definition, the till complexes formed under the isolated ice blocks in this subenvironment can only be discontinuous. It is impossible to correlate the various single deposits because the local conditions result in quite different characteristics of the tills. The position of these complexes in a vertical section is much clearer: dead-ice blocks do not originate in front of an ice mass but are left behind when a n ice mass retreats. This implies that, in a glacigenic sequence, terminoglacial till complexes are found only on top of the glacial part. These complexes most commonly interfinger with a large variety of glaciofluvial and glaciolacustrine deposits.

288

The terminoglacial subenvironment

Depositional mechanisms While the 'englacial tills' are a result of slow thermic melting, geothermal energy as well as thermosubrosion by englacial waters may play a part. Most authors suggest however that the required temperature increase is primarily due to climate-induced surface conditions (Boulton, 1972b; Shaw, 1976; Brodzikowski, 1987). The most important process is melting of the debris-laden ice. The debris consists mainly of fine-grained material which gives rise to a predominantly fine deposit with coarser (sandy) laminae or lenses and floating pebbles and boulders. Glaciofluvial activity, which is especially present under surficial conditions, may lead t o subsequent removal of the fine-grained debris by meltwater, so t h a t boulders and sand become relatively enriched. It is important to notice that terminoglacial ice blocks may be bordered wholly or partly by lakes. Flow tills can easily develop under such conditions, but subaerially mass-transported till units may also be part of the terminoglacial till complexes.

Terminoglacial ice-raft deposits (II-A-1-e) These deposits are formed within terminoglacial lakes and have been abundantly described (N. Eyles, 1983b, 1987; N. Eyles et al., 1983c, 1987a; N. Eyles and Miall, 1984; N. Eyles and Clark, 1985); most ice-raft deposits dealt with by the various authors have, however, a glaciomarine origin (see, among others, Anderson et al., 1980b, 1983b; Drewry and Cooper, 1981; Domack, 1982, 1983; Dreimanis, 1983,1989; N. Eyles, 198313; C.H. Eyles et al., 1985a,b). The sedimentation processes, and the structural characteristics of ice-raft deposits in general, have been discussed extensively by Vorren et al. (1983) and Thomas and Connell(1985). Icebergs and ice-rafts in terminoglacial lakes will gradually melt away a s a result of both solar irradiation and subaqueous undermelting (gradual subaqueous melting by heat transfer from the lake water to the ice). This results in the gradual loss of embedded material (Fig. 179). Fine particles will settle slowly, but material of sand size and larger will drop down rapidly, forming 'xenoliths' in the regular glaciolacustrine sediment o r even more or less continuous sheets of melt-out material that has fallen t o the lake bottom. 'Normal' lacustrine deposits with a distinct number of 'xenoliths' are commonly named 'dropstones' (cf. Thomas and Summers, 1982), although the term 'rain-out diamicts' also is used frequently in the literature.

Deposits of the terminoglacial melting-ice facies

289

RAIN OUT

..

sp

RESE DIMEN TAT ION

CURRENT REWORKING

Gm/Sm

1

4

Fig. 179. Depositional model for diamicts formed as a result of undermelting of lake ice. Slightly modified from: N. Eyles and Menzies (1983).

Lithofacies characteristics These deposits are usually massive, matrix-supported diamicts with a number of remarkably large, floating clasts (pebbles, cobbles or even boulders). The concentration of the clasts depends on their relative contribution t o the debris in the mother-material (ice-raft). The deposits may form a lithological entity (if the rain-out process was the predominant sedimentary process) but it is more common t o find rainout material embedded in material with 'normal' lacustrine lithology (Fig. 180). If the 'normal' lacustrine processes had prevailed, the deposits can show current ripples, wave ripples, etc. Flowage structures may be present if the bottom of the lake had been sufficiently inclined. Large clasts falling from the ice-raft may deform the commonly very sensitive lake-bottom material (Fig. 181).

290

The terminoglacial subenvironment

Fig. 180. Abundant Tertiary xenoliths in a Drenthian sediment deposited i n a terminoglacial lake near Gtogow (SW Poland). Mass-transport processes have resulted in a number of deformations.

Fig. 181. Distinct deformation underneath a xenolith, formed by the impact of the falling clast that was set free after undermelting of a n ice-raft in a terminoglacial lake. Late Drenthian succession, Kleszczow graben, central Poland).

Deposits of the terminoglacial melting-ice facies

291

Textural characteristics Ice-raft deposits usually have a fine-grained matrix, with abundant sand and silt. The floating larger clasts (which may be of boulder size) do not show any preferred orientation. Flattened sand grains may show a preferential orientation of the a/b-plane parallel to the sedimentary surface, particularly if they had been deposited a t the base of a season-induced cycle (after winter). According to the description by N. Eyles et al. (1987a) the aaxes of the grains show a random orientation (within the preferred a h plane), but there also occur several grains with an inclined ah-plane, in which case the a-axes tend t o follow the direction of maximum dip (longitudinal imbrication). Occurrence Terminoglacial ice-raft deposits form usually widespread, planar bodies that show a gradual thinning in the distal direction. The deposits are commonly found downstream of subaqueous tunnel-mouth deposits and have been described in a morphological sense as 'subaqueous sanders' or 'subaqueous glacilacustrine fan deposits' (C.H. Eyles and N. Eyles, 1983a, among others). Depositional mechanisms The essential depositional mechanism is undermelting of ice rafts (or ice covers formed during winter) in terminoglacial lakes. Embedded material is thus set free and the larger particles drop down to become embedded in the bottom sediment. The finer particles settle under quiet conditions then become part of the 'normal' lake sediment from which they cannot be distinguished. Relatively large, local concentrations of unsorted debris within terminoglacial ice-raft deposits represent the sudden overturning of icebergs as a result of thermosubrosion-induced instability; this results i n sudden dumping of the surficially concentrated debris that had been set free by melting of the ice due to insolation. The final deposit usually is also affected by other processes, of which resedimentation and traction-current activity are the most important. Local and temporary conditions determine which process will prevail. C.H. Eyles et al. (1985b) have described the manner in which the various processes affect the final characteristics of the deposit (Fig. 182).

292

The terminoglacial subenvironment

R A I N OUT

+

(ICE-RAFTING A N D PELAGIC FALL.OUT)

DOMl N ANT PROCESSES

1

RAIN OUT

Predominantly massive ( D m m l Planar geometries

iz 2.

RAIN OUT+ RESEOIMENTATION

izd 3.

RESEDIMENTATION

A/

-

1. RESEDIMENTATION + TRACTION

r/

D l A M l C r LIlHOFACIES

ACTIVITY

-

Massive and stratified diainicts with flow structures, abundant silt and clay clasts. rafts (Dmm, Dms)

Stratified and massive diamicts with flow structures. abundant s i l t and clay clasts. rafts. variable grading characteristics (Drnm, Dms, Dmg). Fills and flattens irregular topograpliy. Predominantly stratified diamicts with evidence of resedimentation and traction current activity (wiii iiowcd units, silt a i i t l sand stringers. rippled sands, flow structures, variable grading) IDms. Dcsl

ASSOCIATED LITHOFACIES

Massive muds ( F n i l

Massive muds ( F m ) , laminated S i l t s and clays (turbidites) (FI. Fld)

Laminated silts and clays (turbidites) (FI. Fldl, graded and massive sands (Sg, Sm).

Traction bedded sands ISr, St, Sp) graded sands (Sg) and deformed units of silty sand (Sd)

TRACTION CURRENT ACTIVITY

Winnowed dianiicts. predominantly stratified (Dcs. Dins) Cliaiinelired geometries

Traction bedded qravels and s i i i i d s ((;in, Sr, S t . S i l l

6. TRACTION CURRENT ACTIVITY

Massive and stratified diamicts with s i l t and sand stringers, rippled sands, some winnowed units (Dmm, Dms. Dcs)

Traction bedded sands and gravels (Sr. St, Sp, Gm). Some inud drapes ( F m )

5

.HICCI)

+RAIN OUT

Fig. 182. The three main processes involved in the genesis of terminoglacial ice-raft deposits, and their consequences for the final lithology. From: C.H. Eyles e t al. (1985b). Courtesy: Elsevier Science Publishers.

Deposits of the terminoglacial fluvial facies

293

DEPOSITS OF THE TERMINOGLACIAL FLUVIAL FACIES (11-A-2) The transitional zone between the glacial and proglacial subenvironments is characterised by the melting of ice, resulting in streams, especially where subglacial streams appear or where ablation plays an important role. The longest course of such streams will be in the proglacial environment, but they may flow for a few kilometres through the terminoglacial subenvironment with its high energy gradients. This course may occur partly in the proximal parts of outwash plains. The most obvious characteristic of the terminoglacial subenvironment, viz. the high energy gradient, is expressed in the fluvial facies by extreme variability, in both time and space, of the balance between erosion and sedimentation. The terminoglacial relief therefore also changes continuously: channels several metres deep may be incised within one hour during periods of sudden floods and may also disappear within a few days or even a few hours due to filling. Periods with such intense changes are usually short-lived and separated from each other by longer intervals of more quiet erosion and sedimentation. The deposits in this facies show a wide variety of features and characteristics. They can be grouped into terminoglacial fluvial complexes (11-A2-a), terminoglacial tunnel-mouth deposits (II-A-2-c), terminoglacial stream deposits (11-A-2-d) and terminoglacial sheet- and streamflood deposits (11-A-2-e).The deposits usually show gradual transitions from one group t o the other, and together form large sediment bodies (outwash plains, termed 'outwash fields' by Habbe and Rogner, 1989) that often have a flattened conical shape (sanders).

Terminoglacial fluvial complexes (11-A-2-a) The continuous changes in the position of the ice front, the rapidly changing hydrodynamic conditions, the unstable relief and the intermittent water supply in the terminoglacial subenvironment result in fluvial deposits that are affected by many processes and that do not develop according t o simple fluvial models. These conditions result in the formation of terminoglacial fluvial complexes that show no particular characteristics or show a variety of characteristics that change rapidly, in both the lateral and the vertical direction. For instance, rapidly flowing currents may be partly dammed off, which will result in a decrease in transport capacity and a diminishing average grain size; pools may be formed upstream, catching part of the sediment; shallow fluvial beds may be affected by streamfloods, etc.

294

The terminoglacial subenvironment

The overall result is a deposit which cannot be well characterised, with lithologies ranging from fine sand to badly sorted sand with pebbles and cobbles. Such complexes of Pleistocene age have often been described (commonly under names such as 'ice contact deposits') by, among others, Bartkowski (1967), Klatkowa (1972a), N. Eyles and Slatt (1977), McCabe et al. (1984) and N. Eyles (1987). Other authors (Szupryczydski, 1965; Boulton, 1972a; Boulton and Eyles, 1979; N. Eyles, 1979) have described these complexes from modern environments.

Lit hofacies characteristics The complexes are usually coarse-grained, with gravels or diamict-type mixtures being common at locations near the ice front (Fig. 183).Distinct stratification tends to be rare (Boulton, 1972; Kozarski and Szupryczyfiski, 1978; N. Eyles, 1979; Kozarski, 1982). If horizontal stratification is present, it may be due to either subsequent phases of deposition or to more or less continuous sedimentation

Fig. 183. Terminoglacial fluvial complex of Drenthian age from the Strzelin Hills (SW Poland). Note the bad sorting and the irregular distribution of the clasts. Slight imbrication indicates a relatively strong current.

Deposits of the terminoglacial fluvial facies

295

in the upper flow regime. Strong currents are also evidenced by the occasional occurrence of imbricated clasts, particularly a t the base of the complexes (Fig. 183), a s reported by Shaw (1972a), Banerjee a n d McDonald (19751, Rust (1975, 1984), Saunderson (19751, Brodzikowski and Van Loon (1983), Ruegg (1983), and Kozarski and Kasprzak (1987).

Textural characteristics Bad sorting, particularly a t the base of the complexes, is a major characteristic. Rounding of the clasts is variable, but the commonly short fluvial transport distances result in the frequent occurrence of subangular t o subrounded clasts. The basal parts of the complexes may show imbrication, with the a-axes of flattened pebbles dipping upstream. If no imbrication is present, the aaxes of the larger clasts seem t o have no preferred orientation, though the ah-planes in such a case tend to be parallel to the sedimentary surface. There is usually great spatial variation in petrological and mineralogical composition, especially if compared to terminoglacial stream deposits.

Occurrence The complexes usually form irregular bodies; they are found mostly in a relatively narrow belt parallel to the ice margin, but this belt is discontinuous. The discontinuity is due not only to local deposition, but also t o erosion (the preservational potential of the complexes is relatively small). Chances for preservation are best during ice retreat, when there is a gradual increas in distance of the complexes away from the ice front where sudden floods may exert destructive effects. The complexes often form the distal prolongation of tunnel-mouth deposits and a sharp boundary cannot easily be drawn. Other complexes are found between dead-ice blocks where the topography is very irregular. Because of the differences in preservational potential between phases of ice advance and ice retreat, most terminoglacial fluvial complexes are found in the 'regressive' part of the glacigenic sequence, commonly on top of ablation tills or associated sediments. Lateral and vertical contacts with mass-transported deposits are frequent.

Depositional mechanisms The complexes usually result from rapidly changing flow conditions, with high-energy conditions frequently playing a role. Waning flows result in a

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fining-upward succession, but the finer-grained upper p a r t is often destroyed by a subsequent phase of high-energy conditions (Church and Gilbert, 1975; Rust, 1975), so that coarse - often massive - deposits are usually dominant. Deposition of the complexes takes usually is rapid: outbursts (e.g., by a breakthrough of a n ice dam) result in sudden floods (from zero t o several cubic metres per second within one bed) that have large erosive capacity, thus increasing their sedimentary load and depositing large amounts of generally unsorted material when the current velocity drops. Saunderson (1975, 1977) suggested that most sediments that we term 'terminoglacial fluvial complexes' might be deposited under full pipe flow conditions. The horizontal stratification due to current activity in the upper flow regime is the most common sedimentary structure a n d Saunderson (1977) therefore terms this the 'sliding bed facies'. It should be noted that horizontal lamination found in the complexes may also be due to currents in the lower flow regime, but it is striking that current ripples are relatively scarce.

Terminoglacial tunnel-mouth deposits (II-A-2-c) Tunnel mouths a r e found both subaerially and under subaqueous conditions (viz. where a subglacial or englacial meltwater stream embouches in a lake under the water level: Fig. 184). These different conditions result i n somewhat different deposits, but the similarities seem more important than the differences, so that all these deposits can be grouped together. The most obvious characteristic is the high sedimentation rate of - usually unsorted - material as a result of a sudden decrease in the energy level of the meltwater current. The rapid deposition is also responsible for the usually very limited lateral extent of tunnel-mouth deposits. Together, all tunnel-mouth deposits contribute t o the formation of eskers. The morphological term 'esker' (see also p. 10) was defined by Francis (1975), who arrived a t a compromise between the definitions proposed by Geikie (1894) and by Chamberlin (1894): "Eskers are glacial features made up of morainic material deposited in contact with glacier ice as ridges whose trends tend to conform in general with a direction of ice movement, and whose composition is dominantly, but not necessarily exclusively gravel and sand". Saunderson (1975) gave a more detailed, but less precise definition after a sedimentological study of Canadian eskers. Sets of lithologically different deposits can be found inside an esker, deposited and sometimes subsequently partly reworked in various subenvironments and facies. It is clear, for instance, that eskers may consist

Deposits of the terminoglacial fluvial facies

297

lounderrn9 01 channels

Deformed Substrate Sandy Rain-out Facies from Elllux Jet

Muddy Rain out Facies from Plume

0

Debris Flow Facies Subaqueous Oulwash Facies

Fig. 184.Model of a terminoglacial, subaqueous tunnel mouth and related phenomena. Slightly modified from: N. Eyles and McCabe (1982).

of subglacial meltwater-tunnel deposits (I-C-2-b) and of terminoglacial tunnel-mouth deposits (II-A-2-c). The morphology may be identical but the internal structure will show sedimentological differences. There appear to exist lateral and longitudinal successions in the sedimentary structures within eskers (Terwindt and Augustinus, 1985). Tunnel-mouth deposits formed in terminoglacial lakes show relatively large variability in both the horizontal and the vertical direction, though most of the sediment is rather coarse. There may be dammed-off parts in the streams, a feature that emphasises the relationship with lakes. The sediments may show decreasing grain size and increasingly regular lamination towards the centre of the temporarily dammed-off locations. Five situations have been described in which terminoglacial tunnelmouth deposits largely determine the relief in the area near the ice front. This concerns two subaerial and three subaqueous situations. In the case of a subaerial tunnel mouth, small fans are usually formed, which tend t o follow a limited number of channels. If there is a n ongoing retreat of the ice front, accumulation of these fans results in eskers such as those described by Szupryczyiiski (1965). If the ice front is more or less stable, a complex zone of thick channel fills is formed, within channels

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The terminoglacial subenvironment

that were deeply incised in the previously deposited glacigenic material. In this case separate fans and the position of the tunnel mouth can rarely be correlated unambiguously because each successive outburst of water from the tunnel mouth results in erosion of older tunnel-mouth deposits, formation of new gullies and sedimentation within these new gullies. The final result is a complex, coarse-grained body that resembles i n some respects the terminoglacial fluvial complex (II-A-2-a) described earlier (cf. Flint, 1971; Embleton and King, 1975; Saunderson, 1975; Klimaszewski,

Fig. 185. Subaqueous esker-bead formation during recession of the grounding-ice front. Modified after: Banerjee and McDonald (1975). TM = subaqueous tunnel mouth; SFT = subaqueous flow till; SEB = subaqueous esker-bead (single form);IRD = ice-rafted debris; DD = dump deposits; R = redeposition after reworking; tr = traction; TC = turbidity currents; CH = subaqueous channels. 1 = active ice; 2 = flow lines within the ice; 3 = supraglacial crevasses; 4 = hanging englacial tunnel mouth; 5 = subglacial subaqueous tunnel mouth a t the grounding line; 6 = floating icebergs; 7 = subglacial channel; 8 = rain-out diamicts; 9 = traction-current deposits; 10 = erosive base of subaqueous fan; 11-12 = channel deposits; 13 = laminites; 14 = sands deposited by waning flows; 15 = melting-ice facies; 16 = sands and gravels; 17 = parapelagic cover; 18 = transgressive deposits; 19 = lodgement tills; 20 = diamicts within the flow lobes; 21 = substratum.

Deposits of the terminoglacial fluvial facies

299

1976; Sugden and John, 1976).A situation that involves an advancing ice front does not result in tunnel-mouth deposits with a good chance of being preserved. The three situations with subaqueous tunnel-mouth deposits result in esker beads (Banerjee and McDonald, 19751, subaqueous sanders (Rust and Romanelli, 1975) or in paraesker bars. These bodies were studied by the present authors in the foreland of the Sudetic Mountains (SW Poland), where they occur along the previous front of the Pleistocene ice cover during its maximum extent. Esker beads, as described by Banerjee and McDonald (1975) may form subaqueously during the step-by-step retreat of the ice front. Each short period of stagnation results in the formation of a type of fan. The final

Fig. 186. Formation of subaqueous sanders during the continuous retreat of a grounding-ice front in a terminoglacial lake. After: Rust and Romanelli (1975). TM = tunnel mouth; SFT = subaqueous flow till; IRD = ice rafting of glacial debris; DD = dump deposits; SS = subaqueous sander deposits. 1 = active ice; 2 = flow lines within the ice; 3 = marginal crevasses; 4 = subglacial tunnel mouth; 5 = ice raft; 6 = supraglacial mass flows; 7 = terminoglacial flow till; 8 = traction-current and rain-out deposits; 9 = terminoglacial subaqueous ablation tills; 10 = non-specific diamict; 11 = deposits formed during one phase of large influx; 12 = sands and gravels of the subaqueous sander; 13 = infilling of large subaqueous channel; 14 = idem, in longitudinal section; 15 = diamicts of a previous phase of glacierisation; 16 = terminoglacial lacustrine bottomsets; 17 = subglacial till; 18 = substratum.

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The terminoglacial subenvironment

Fig. 187. Formation of subaqueous paraesker bars during retreat of the grounding-ice front. TM = tunnel mouth; PP = parapelagic deposition; tr = traction; CH = subaqueous channel. 1 = active ice; 2 = englacial tunnel; 3 = flow lines within the ice; 4 = active ice/water contact; 5 = subglacial tunnel mouth; 6 = sandy fan covering the bar; 7 = parapelagic cover; 8 = reworked parapelagic deposits; 9 = erosive channel; 10 = undermelting in subglacial channel; 11 = coarse tunnel-mouth deposits; 12 = coarse core of the bar; 13 = proximal fan deposits formed during periods of high water; 14 = distal fan deposits; 15 = traction-current deposits formed onder waning-flow conditions; 16 = secondary channels with cross-bedded sands; 17 = current ripples; 18 = parapelagic mud; 19 = erosive base of the bar; 20 = substratum.

result is a series of isolated hills, located along the pathway followed by the tunnel mouth (Fig. 185). Subaqueous sanders (Rust and Romanelli, 1975) are formed when a number of tunnel mouths are present one next to another, and when the ice front retreats slowly but without distinct interruptions. The result of sedimentation under these conditions is a large number of small fans superimposed proximally upon each other, both parallel and perpendicular to the previous ice front (Fig. 186). Paraesker bars are much smaller but very continuous. The authors have found them in areas that had been dominated by terminoglacial ponds that were dammed off on one side by grounding ice and on the other

Deposits of the terminoglacial fluvial facies

301

side by a mountain range. Such elongated sediment bodies with a 'fandeltaic' character form during a rapid, continuous retreat of the ice front. The internal structure of these bars is much more complex than that of subaqueous sanders but simpler than that of esker beads (Fig. 187). The general characteristics just described indicate that terminoglacial tunnel-mouth deposits form a relatively large and complex group of glacigenic deposits. They have great palaeogeographic significance, particularly because they indicate both the position of a previous ice front and the presence of a stage when ice retreat was dominant.

Lithofacies characteristics The most complex lithofacies is to be found in tunnel-mouth deposits that form subaqueous esker beads. The proximal parts often have a gravelly or massive diamictic appearance (Fig. 188) - sometimes with either massive or stratified sands and gravels in the top part - while the distal parts usually consist of stratified gravelly sand. There is a clear decrease in average and maximum particle size in the distal direction, and there may

Fig. 188. Massive, diamictic proximal part of terminoglacial tunnel-mouth deposits from the Sudetic foreland (SW Poland), formed during the Drenthian.

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The terminoglacial subenvironment

be current ripples and horizontal lamination where sand predominates (Fig. 189). Although some antidunes have been reported few structures are usually left by currents in the upper flow regime; back ripples due t o flow separation are commonly seen (Kennedy, 1963; Banerjee and McDonald, 1975). Subaqueous paraesker bars are also complex, though somewhat less. Massive gravelly and sandy lithofacies dominate in the proximal zones, gradually passing into more stratified deposits with a more or less similar grain-size distribution in the distal direction. Current ripples are fairly common in the distal part and often show remarkably steep foresets. Subaqueous sanders are much less complex and consist mainly of gravels and sands. An upper flow regime can often be reconstructed on the basis of the sedimentary structures, but there are equally abundant ones representing a lower flow regime. Horizontal lamination is the most common, but tabular and trough-crossbedded sets of ripples are usually found as well (Fig. 190). Boulder beds may also be present (Rust and Romanelli, 1975; Visser et al., 1987). Tunnel-mouth deposits formed under subaerial conditions are usually massive bodies of gravelhoulder mixtures or clast-supported diamicts. Stratification is rare, and - if present - often vague and discontinuous.

Fig. 189. Four vertically stacked successions in the Peterborough esker. Current from right to left. Length of longitudinal section 600 m. Slightly modified from: Banerjee and McDonald (1975).

Deposits of the terminoglacial fluvial facies

303

Fig. 190. Section, through a subaqueous sander (as defined by Rust and Romanelli, 1975),from the Scinawa ice-dammed lake (Trzebnica Hills, SW Poland) that existed during the Wartanian glaciation. Most beds are vaguely laminated or show badly developed trough-like cross stratification.

Textural characteristics

These deposits are usually badly sorted, especially if formed under subaerial conditions. Better sorting is found in the distal parts of subaqueous deposits. The usually bad sorting (occasionally with particles over one metre in diameter) is a result of relatively short transport, which is also expressed by the frequent occurrence of angular clasts although subangular to subrounded clasts tend to dominate. The surfaces of the larger particles often show glacial striae. Sand grains and larger particles tend to be randomly orientated in the most proximal parts of tunnel-mouth deposits, but there is often imbrication or a (sub)horizontal orientation of the ah-planes in the distal parts, where organised bottom currents play a clearcut role. The mineralogical and petrological composition of the deposits shows no regularity; in this respect, the composition of the deposits is as variable as that of terminoglacial tills.

3 04

The terminoglacial subenvironment

Occurrence The position of terminoglacial tunnel-mouth deposits depends heavily on overall environmental conditions. The occurrence of these deposits in comparison t o that of other types of terminoglacial sediments can vary from relatively rare t o frequent. If formed subaerially, the deposits have an irregular lens shape that shows overall thinning in a distal direction. Most of these deposits are covered by diamicts and show frequent interfingering with terminoglacial sheet- and streamflood deposits. Some have been found in the lower parts of kames. Tey are much more common, however, as elongated ridges, termed 'eskers'. Subaqueously formed tunnel-mouth deposits are usually much less well expressed geomorphologically, except for esker beads, which may form distinct ridges a t the bottoms of large lakes. Such ridges are commonly embedded in glaciolacustrine deposits, but may be in direct contact with terminoglacial tills. Paraesker bars are usually discoidal bodies about 2-5 m thick, with a most irregular spatial distribution. Their longest axis is orientated perpendicular t o the ice front. Subaqueous sanders are much more continuous, but this type is rare because of the special depositional conditions required for their formation. These rare subaqueous sanders, which are thin in comparison to their extent and very continuous, usually are also covered by glaciolacustrine deposits. Depositional mechanisms A pulsational character, often even resulting in the development of sheetor streamfloods, is a very important aspect of the meltwater currents flowing out from tunnel mouths. A very high energy level is reached very soon after the outflow has started to increase; slowing down of the current usually takes much longer, but the intervals with high current velocities are relatively short. The energy gradients thus change very rapidly. The gradual slowing of the debris-laden current results in the deposition of relatively thick sedimentary bodies. These will generally be destroyed during a subsequent phase of ice advance. However, if the ice should recede, the preservational potential is very high (Banerjee and McDonald, 1975; Rust and Romanelli, 1975; Visser et al., 1987), particularly because erosion by a new outburst of meltwater from a tunnel mouth is unlikely: the curved shapes of meltwater tunnels result in positions of the tunnel mouth that change laterally during ice retreat and, in addition, the currents loose much of their power within short distances.

Deposits of the terminoglacial fluvial facies

305

The sedimentary particles are transported in the subglacial channels as both bed load and suspended load. If the current velocity is high, the bed loads consists only of coarse material; the suspensions may be either homogeneous or heterogeneous (Newitt et al., 1955; Acaroglu and Graf, 1968; Elliott and Gliddon, 1970; Saunderson, 1977). The current velocity drops suddenly if the subglacial, channelised current spreads out in front of the ice. The bed load that was carried along will be deposited almost immediately, forming diamicts or even pure gravels. The finer material will be transported further away, forming a gradually thinning, fining body in the distal direction. The tunnel-mouth deposits that form the middle part between the proximal and the distal part often reflect the current energy best: the base of a unit may be erosive and the unit itself may be made up of - from bottom to top - coarse, horizontally stratified sands that represent the upper plane bed phase, tabular cross sets and isolated current ripples that represent the lower flow regime, and - lastly - horizontally laminated fine sands o r muds that represent the lower plane bed phase.

Terminoglacial stream deposits (II-A-2-d) Typically channelised flows may occur both as rivers or brooks, and as subaqueous flows in deltas or lakes. The latter group of 'subaqueous' currents is very similar t o 'subaerial' currents - particularly from a hydrodynamic point of view - and several sedimentologists therefore include such currents in the fluvial facies. The present authors do not follow this approach as we consider, like most other sedimentologists, bottom currents in lakes and deltas to be phenomena that are a normal part of the lacustrine and deltaic facies. The stream deposits described here are therefore restricted to the fluvial facies in its strict sens, i.e. under 'subaerial' conditions. The meltwater flows coming from the ice may join streams of nonglacial origin that flow - due t o the general relief - towards the ice front, where the water cannot continue its course. The streams are then diverted and run more or less parallel to the ice front in so-called ice-marginal streamways or pradolinas. Such situations have been reconstructed in detail for various phases of the European Pleistocene glaciations. Another possibility is that the meltwater currents follow a general topographic inclination away from the ice; such streams first follow irregular paths between dead-ice blocks, moraines, etc. (Fig. 191) then become gradually more 'organised', finally to change into regular braided or even meandering streams in the proglacial subenvironment.

306

The terminoglacial subenvironment

Fig. 191. Terminoglacial area in front of the Werenskjold glacier (SW Svalbard). The ridge at the foreground is an ice-cored moraine. An irregular drainage pattern, where irregular stream deposits accumulate in rapidly changing channels, is visible in the centre. Photograph: J. Cegia.

Lithofacies characteristics Most streams in the terminoglacial subenvironment have a braided character. Debris of all sizes is supplied by the meltwater currents; the currents have rapidly changing transport capacities, so that especially the coarser particles (sand and gravel) settle whereas most of the fines are carried along. The terminoglacial stream deposits formed in braided channels do not differ from their non-glacigenic counterparts. This implies t h a t the general appearance is from massive to irregularly stratified, with finegrained intercalations and some levels in which current ripples have been preserved (Fig. 192). Large-scale cross-bedding occurs occasionally (Rust, 1975, 1978; Miall, 1977) and is encountered particularly in interbar channel fills. The channel bars may also show tabular cross-stratification. It is usually impossible to interpret braided-stream deposits as terminoglacial without taking into account the lateral and vertical facies

Deposits of the terminoglacial fluvial facies

307

Fig. 192. Terminoglacial stream deposits from the surroundings of Swidnica (Lower Silesia, SW Poland). Note the superimposed ripples on the bar surface.

relationships: there are no truly diagnostic features. The vicinity of an ice mass can sometimes be reconstructed from glaciotectonic deformations caused by a temporary re-advance of the ice.

Textural characteristics In these deposits, sorting may range from good to very bad; this depends mainly on the number of transport and depositional cycles (each subsequent cycle tends to result in somewhat better sorting). This is also expressed as the degree of rounding of the larger clasts: there is a wide variety but, i n general, the most proximal deposits comprise mostly angular and subangular clasts, whereas the most distal parts also show numerous rounded and subrounded clasts. Most of the flattened clasts are orientated with their ah-planes parallel t o the sedimentary surface, but imbrication may be present. This structure is more common, however, in deposits that were formed under the relatively rare condition of a meandering stream.

308

The terminoglacial subenvironment

Occurrence

The terminoglacial stream deposits are t o ,e found in the relatively narrow belt in front of the ice mass, where the terrain is often irregular due to the presence of dead-ice masses, moraines, etc. The configuration of the stream deposits in this subenvironment is therefore irregular, but the irregularity decreases in a distal direction. The horizontal extent away from the ice may range from only a kilometre (e.g., in front of the glaciers a t Svalbard: Baranowski, 1970; Boulton, 1970, 1972; Drewry, 1986) t o several tens of kilometres (e.g., in the European lowlands in front of the Pleistocene ice cap: Bartkowski, 1967). In the field, the terminoglacial character of the deposits may become clear from the frequent interfingering with typically glacial deposits (e.g., supraglacial, englacial or subglacial tills). Another indication is the almost unavoidable interfingering with glaciolacustrine deposits, due to the typically irregular terminoglacial topography. Depositional mechanisms

The stream deposits are formed as a result of waning meltwater currents in front of the main ice mass. The depositional mechanisms that play a role have been described extensively for non-glacigenic conditions by numerous authors. Massive o r even diamictic units may be formed during periods of highenergy currents, when coarse material can be supplied. Such currents may be erosive and, indeed, the massive and coarse units are most commonly found as the lower parts of channel fills. A gradual decrease in energy of the currents results in less coarse deposits, often with some kind of stratification, even with current ripples and low-energy parallel lamination. Fine-grained lenses represent pools that were formed when sedimentation led t o local damming-off of the braided channel pattern.

Terminoglacial sheet- and streamflood deposits (II-A-2-e) These deposits are formed when sudden outbursts of meltwater result in a rapid change from 'normal' current conditions to high-energy conditions. Streamfloods (Reading, 1978) represent a situation in which the water level is high enough to fill the channel pattern completely but not high enough t o make the water overflow the banks that surround the channelincised area. Sheetfloods occur when the water level is even higher, so that all the area is covered with a layer of rapidly flowing water.

Deposits of the terminoglacial fluvial facies

309

Streamfloods occur frequently in the terminoglacial subenvironment, usually as a result of upstream breakthrough of a sedimentary body or a n ice mass that dammed off a meltwater current or that formed the wall of a lake. Channels in the terminoglacial subenvironment undergo sudden increases in meltwater supply so frequently that they are usually deeply eroded. This implies that these channels gradually become better adapted to transport sudden water masses. Streamflood deposits are therefore most commonly found in the downstream part of the terminoglacial fluvial area. Sheetfloods are rarer, simply because of the usually high capacity of the channels. However, if sheetfloods occur, sedimentation takes place everywhere as soon as the current velocity drops. The result is a widespread, relatively thin cover of stratified but badly sorted material.

Lithofacies characteristics There is a clear sedimentological difference between the lithological characteristics of sheet- and of streamfloods in general, including those in this subenvironment.

Fig. 193. Wisconsinan terminoglacial streamflood deposits covered by a sheetflood unit (canyon of the Fraser River, British Columbia,Canada).

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The terminoglacial subenvironment

Streamflood deposits consist mostly of relatively coarse particles, and successions may show pure gravel at the base that gradually changes into gravelly sand at the top (Fig. 193). The sands often show a cross-bedded lower part and a horizontally laminated upper part, reflecting the gradual decrease of the current's Froude number. The sheetflood deposits are thinner and more widespread. They usually show a well developed horizontal stratification, interrupted by frequent scour-and-fill structures. The make-up of these deposits may range from gravel with occasional cobbles to muddy material. The gravelly parts usually include strings of pebbles that give the appearance of good stratification (Fig. 194);the sandy parts show distinct layers, often having levels with abundant current ripples; the muddy deposits often show a fine lamination, indicating that current velocity had dropped t o almost zero (possibly by damming off of a channel) so that even the finest particles could settle in the almost or completely stagnant water.

Fig. 194. Terminoglacial sheetflood deposits of Drenthian age from the Sudetic foreland (SW Poland).

Deposits of the terminoglacial fluvial facies

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Textural characteristics The textural characteristics depend largely on regional and local conditions. This is expressed by extremely variable degrees of sorting, with bad sorting being the most common. Rounding of the clasts is equally variable, but subangular t o subrounded particles seem t o dominate in most deposits. The orientation of the larger clasts depends mainly on the shape of the clasts. Flattened clasts are mostly orientated parallel to the sedimentary surface (although imbrication may be present locally), but non-flattened particles seem not t o have any preferred orientation. One should keep in mind, however, that the common alternations in grain size facilitate the genesis of sedimentary deformations which may hide an original clast fabric or, on the contrary, introduce a new preferential orientation.

Occurrence Streamflood deposits are found as sedimentary bodies in which the configuration of the drainage pattern is well visible (if there are sufficient outcrops or if a present-day situation is observed). The deposits show a highly irregular thickness and frequent interfingering with numerous other types of terminoglacial deposits, particularly stream deposits, glaciolacustrine deposits and mass-transported sediments. If no erosion has destroyed the previously deposited material, sheetflood deposits form an almost continuous cover; each cycle may contribute a few centimetres up to about two metres t o the entire thickness. The blankets formed by the sheetflood deposits thin very gradually in the distal direction, where they interfinger with the same types of deposits as do the streamflood deposits. Terminoglacial sheet- and streamfloods are usually found in the last few metres of the glacigenic sequence, under the first tills and above the last tills.

Depositional mechanisms Both sheet- and streamflood deposits represent high-energy conditions. Deposition of coarse material may take place rapidly during such periods of high energy; the 'finer' (predominantly sandy) parts were usually also formed in a short time, viz. when the high-energy meltwater current gradually lost its power. A series of sheet- or streamflood deposits on top of each other thus represents a period with short intervals of rapid sedimen-

312

The terminoglacial subenvironment

tation and much longer intervals without significant sedimentation or even with erosion. The fundamental mechanism behind the genesis of these deposits is thus the irregular, pulse-like supply of debris-laden meltwater.

DEPOSITS OF THE TERMINOGLACIAL DELTAIC FACIES (11-A-3) The terminoglacial subenvironment has the largest concentration of meltwater streams and of bodies with (more or less) stagnant water. It is therefore not surprising t h a t deltas develop frequently. Deltas a r e complex sedimentary bodies. The classical concepts regarding deltas, e.g., those distinguishing between the Gilbert-type, the Hjulstrom-type and the Salisbury-type delta (see, among others, Aario, 1972; Le Blanc Smith and Eriksson, 1979; Kostaschuk and McCann, 1983; Syvitski et al., 1984) have recently been adapted t o make them consistent with new insights. Numerous authors now distinguish primarily, as far as coarse-grained deltas a r e concerned, between fan deltas and braid deltas (see, e.g.,

Braided river-

EraidplainJSandur

Fig. 195. Difference between a fan delta (upper left) and braid delta (others). Coalescing braid deltas of sander origin are seen on the right. Slightly modified from: McPherson e t al. (1988).

Deposits of the terminoglacial deltaic facies

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McPherson et al., 1987,1988; Marzo and Anadon, 1988; Nemec and Steel, 1988a,b; Postma and Cruickshank, 1988; Wood and Ethridge, 1988), the first group generally representing conditions of higher energy (McCabe et al., 1987; McCabe and Eyles, 1988; N. Eyles et al., 1989). Fan deltas (as defined originally by Holmes, 1965) are formed by the progradation of a n alluvial (i.c. glaciofluvial) fan directly in a standing body of water, whereas braid deltas are formed by the progradation of a purely fluvial (i.c. glaciofluvial) braid plain on to a standing water body (Fig. 195). According to this criterion, braid deltas seem more typical of the proglacial subenvironment, whereas fan deltas will usually be present in the terminoglacial subenvironment (cf. Schumm, 1968; Cotter, 1978; Eriksson, 1978; Minter, 1978; Vos, 1981; Fuller, 1985; Bergh and Torske, 1986; McPherson et al., 1988; Nemec and Steel, 198813). If the deltas are insufficiently large t o show internal differentiation, they form terminoglacial deltaic complexes (11-A-3-a).If the deltas are larger, terminoglacial deltaic topsets (II-A-3-b), terminoglacial deltaic foresets (11-A-3-c)and terminoglacial deltaic bottomsets (11-A-3-d)can be distinguished.

Terminoglacial deltaic complexes (11-A-3-a) Few of these complexes have been described, although several workers have referred to their existence under terminoglacial conditions (Gilbert, 1972a; Mitchell et al., 1973; Shaw, 1977; Shaw and Archer, 1979; Haluszczak, 1982; Gustavson and Boothroyd, 1987; Brodzikowski, in press). The small complexes are generally short-lived, or form the early stage of much larger deltas; this explains why they have not been mentioned as frequently as their larger counterparts.

Lithofacies characteristics Complexes of fan deltas consist mainly of gravels and gravelly sands (Fig. 196). Their internal structure depends strongly on the topographic situation; a strongly inclined lake slope, for instance, will favour the formation of well bedded (inclined) layers that may show frequent signs of local mass transport. On the other hand, the upper part of the fan delta may temporarily be partly subaerial, and structures related t o streamfloods and sheetfloods may then be found (Hogg, 1982; Gustavson and Boothroyd, 1987). More sandy intercalations between the gravel-rich units become more frequent in a downslope direction and even truly muddy layers may be found at the foot (either or not due to mass movements).

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The terminoglacial subenvironment

Fig. 196. Terminoglacial deltaic complex from the surroundings of Strzelin (Sudetic foreland, SW Poland).

A relatively restricted supply of clastic material, in combination with a gentle lake slope, commonly results in a lack of distinct structures. A more 'normal' lake slope provides the opportunity for currents with sufficient energy to produce current ripples and parallel lamination. Textural characteristics The general abundance of gravel in fan deltas may result in a diamict-like appearance. Such a lithology is often matrix-supported, sometimes with a weakly developed grading (Postma, 1984a; McCabe and Eyles, 1988). The clasts are usually orientated with their ah-planes parallel to the sedimentary surface, but elongated clasts may also be directed in such a way that there is a n imbrication with the a-axis dipping either upslope or downslope (Volk, 1966; Postma and Roep, 1985; Major and Voigt, 1986; McCabe and Eyles, 1988). Fine-grained complexes may contain floating pebbles or pebble-rich lenses, often representing the head of a slumped mass. The glacial source of the clasts may be reflected in striated surfaces, but there are several reports of such deposits that show no striated clasts.

Deposits of the terminoglacial deltaic facies

315

Reversed grading may also be encountered, as a result of grain flow, in deposits formed in relatively deep water. Normal grading, usually weakly developed? may be found in complexes with a fine average grain size (Lowe, 1979,1982).

0cc u r rence The complexes form irregular bodies which interfinger, in their proximal part, with deposits of the terminoglacial or subglacial fluvial facies, and, in their distal part, with terminoglacial (or sometimes proglacial) glaciolacustrine deposits. The most common situation is an intermediate position between subglacial meltwater-tunnel deposits (I-C-2-b) or terminoglacial tunnel-mouth deposits (11-A-2-c) and terminoglacial lacustrine deposits (11-A-4-aor 11-A-4-b). I t is usually impossible t o draw a clear boundary between t h e glaciofluvial and the glaciodeltaic deposits on the one hand, and between the glaciodeltaic and the glaciolacustrine deposits on the other hand. This is due not only to the gradual transitions that are also known from nonglacigenic facies, but also t o the oscillations of the ice front and the frequent changes in the water level of the lake. In the glacigenic sequence, these complexes a r e most commonly covered by deposits from either a melting-ice or a glaciolacustrine facies from the terminoglacial subenvironment. In the first case there was obviously a n ice advance. The deltaic complexes are then usually strongly deformed by glaciotectonic pushing. Depositional mechanisms The main processes that influence the formation of these complexes are the dynamics of the ice front, the character of the meltwater stream, the supply of debris by the meltwater stream, and the nature of the lake in which the stream debouches. Subaerial parts of the deltas may undergo streamfloods and sheetfloods, but most sedimentation takes place through settling of particles in suspension and through interrupted transport of bed load in subaqueous channels. The channels especially also undergo mass-transport activities. A supply of predominantly fine material often results in continuous deposition that can be compared with an avalanche process. The limited are where all these processes take place, sometimes simultaneously? is the reason why no well developed delta was formed but a complex of relative small size.

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Terminoglacial deltaic topsets (II-A-3-b) These deposits have been described by Le Blanc Smith and Eriksson (1979), McCabe and Eyles (1988) and Fralick and Miall (19891, though only as a topic of secondary interest. Sedimentological data on deltaic topsets formed under conditions that correspond with those in the terminoglacial subenvironment were provided by Drewry (1986), Brodzikowski and Van Loon (1987), Nemec and Steel (1988b),Postma and Cruickshank (1988) and Postma et al. (1988a).

Lithofacies characteristics Coarse-grained (fan) deltaic topsets consist mainly of sandy gravels and gravelly sands, with frequent intercalations of debris-flow deposits that may have a more diamict-like lithology. The most proximal part of large deltas in particular tends to contain massive, gravelly strata with some laterally continuous sandy units (Fig. 197). Some of the gravel-rich beds might be interpreted as fluvial rather than as deltaic, and some of these

Fig. 197. Terminoglacial deltaic topsets of Elsterian age, formed in the lake within the Kleszczow graben (central Poland).

Deposits of the terminoglacial deltaic facies

317

beds are poorly stratified t o massive units, sometimes with scour structures and - rarely - solitary sets with planar cross-stratification (Marzo and Anadon, 1988). In their proximal part less coarse-grained topsets usually consist of horizontally stratified sands. These sometimes alternate with layers of pebbly sands that may be partly deposited by stream- and sheetfloods rather than by purely deltaic agents. There may be intercalations of matrix-supported debris-flow deposits under these conditions also. The massive and coarse character of the topsets gradually disappears in a distal direction, where fining-upward successions also become more common. Horizontally and cross-stratified units are also common here (cf. Miall, 1978; Eriksson and Vis, 1979; Le Blanc Smith and Eriksson, 1979; Marzo and Anadon, 1988).

Textural characteristics The textural characteristics are strongly influenced by the occurrence of fluvial and mass-transport intercalations. The purely deltaic units generally consist of badly sorted material, although both normal and reversed grading may be present. The material is usually so coarse that there is a clast-supported diamictic appearance. The larger clasts frequently show striations.

Occurrence These deposits form sheet-like bodies, usually with a much greater length than width (as a result of the progradation). The thickness is limited, often less than one metre (but sometimes over 5 m), and decreases in a downstream direction. The deposits form the top of the deltaic sequence within the terminoglacial interval of the glacigenic sequence, and has lateral contacts similar t o those described for the terminoglacial deltaic complexes (II-A-3-a). Contacts with the terminoglacial fluvial and mass-transport facies have been described by Le Blanc Smith and Eriksson (1979) and by McCabe and Eyles (1988).

Depositional mechanisms The terminoglacial deltaic topsets are formed primarily by deposition from running water, although mass-transport mechanisms may disturb the material deposited originally or supply material from a fluvial facies.

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Massive, gravelly units may result partly from the rapid deposition that takes place during streamfloods in the subaerial part of the delta; they commonly pass into more finely stratified beds that represent stages of waning flood. Fine-grained, widespread lenses represent deposition in interdistributary bays, and are incised by distributary channels with more coarse material and a typical channel sequence. The overall grain size increases from bottom to top, which results from a lateral facies shift, due t o the continuous progradation of the delta.

Terminoglacial deltaic foresets (II-A-3-c) There are great differences between coarse-grained (fan delta) and finergrained (usually Gilbert-type, braid delta) foresets (cf. Fisher et al., 1969; Galloway, 1976; McCabe and Eyles, 1988; Nemec and Steel, 1988a,b). There is a lack of specific information in the literature, but some aspects of these deposits have been described by Boothroyd (1976), Le Blanc Smith and Eriksson (1979) and McCabe and Eyles (1988). More adequate descriptions are available about comparable deposits formed under nonglacigenic conditions (e.g., Casey and Scott, 1979; Ethridge and Wescott, 1984; Kingsley, 1984; Fraser and Suttner, 1986; McPherson et al., 1987, 1988; Marzo and Anadon, 1988; Nemec and Steel, 1988a,b); the details provided in these reports are most likely to be also applicable t o terminoglacial deltaic foresets. Subaqueous fans (Fig. 198) have much in common with deltaic foresets. Relatively much is known about fans of this type that were formed under terminoglacial conditions (Rust and Romanelli, 1975; McCabe et al., 1984; McCabe, 1986; C.H. Eyles, 1987; Visser et al., 1987; N. Eyles et al., 1989).

Fig. 198. Longitudinal section through a subglacial channel with a terminoglacial subaqueous fan in front (a). A plan (b) shows ridges and troughs. From: Helm and Roberts (1984). Courtesy: Geological Journal.

Deposits of the terminoglacial deltaic facies

319

Lithofacies characteristics The internal, inclined structure in these deposits is striking. Relatively fine-grained foresets usually show large-scale cosets of high-angle to lowangle planar foresets. The inclination of the foresets may vary between almost nil t o about 30". Individual beds a r e often graded, and t h e uppermost parts of the foresets frequently show trough cross-bedding. The overall simple picture may be disturbed locally by the occurrence of debris set free from ice rafts; large boulders that fall on the foresets may induce mass movements. However, slumps of other origin are also common. The foresets in relatively coarse terminoglacial deltas are much more complicated (Fig. 199).This is due mainly to the predominance of gravityinduced mass movements, which alternate with subaqueous traction currents. An additional complication is, a s in the finer-grained equivalents, the occurrence of dropstones derived from melting ice rafts. The result is a n alternation of units with different lithological properties, varying from fine-grained layers via diamicts to massive, coarse beds. The upper part of the foresets tends to be coarser than the lower part; there is

Fig. 199. Terminoglacial deltaic foresets of Elsterian age, formed in a lake in the Kleszczow graben (central Poland). Height of the foresets: approx. 3 m.

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The terminoglacial subenvironment

also a slight tendency t o a change from coarse t o fine in a downstream direction. The same direction shows a gradual decrease in the inclination of the separate foreset units. The 'foreset appearance' is commonly due t o elongated bars (Nemec and Steel, 1984; Postma, 1984b). Subaqueous fans are usually coarse-grained. While massive diamicts seem t o be the most common type, gravelly and sandy intercalations are also seen. Gravels in proximal subaqueous fans usually have a massive appearance, although normal grading may be present. The inclination of the scparate units is determined by local conditions.

Textural characteristics Coarse-grained deposits of this type tend to be badly t o moderately sorted, whereas finer equivalents generally show good to moderate sorting. There are so many exceptions, however, that one should consider sorting as a variable criterion. Coarse deposits are usually clast-supported, with the clasts showing no preferred orientation or an orientation subparallel to the sedimentary surface. The matrix of the coarse units is also badly sorted and consists predominantly of sand, with minor amounts of coarser and finer particles. The somewhat finer-grained units often have slightly better sorting, and vague grading. Only units with no or almost no pebbles and coarser clasts may show relatively good sorting.

Occurrence The foresets form the middle part of the deltaic deposits. They are found in this position in the glacigenic sequence, but one should keep in mind that the underlying bottomsets are not always well developed (they may be interpreted as terminoglacial lacustrine bottomsets) and that the overlying topsets have reasonably often been eroded, particularly if a phase of ice advance followed the deposition. In this case, the terminoglacial deltaic foresets, which have a relatively high preservational potential, are immediately overlain by subglacial deposits. The deposits form relatively large, widespread bodies that are bordered at their proximal part by fluvial deposits and a t their distal part by lacustrine deposits; the sideward contacts may be with numerous other types of terminoglacial deposits. Terminoglacial subaqueous fans occur in a similar way. They form flattened, widespread bodies incised by channels that become wider and shallower in a downstream direction.

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321

Depositional mechanisms Several mechanisms contribute t o the deposition of terminoglacial deltaic foresets. Some of the material is supplied as a traction carpet, some settles from suspension due to a decrease in current velocity, and some is supplied by mass movements. Several types of mass-transport mechanisms take place under these conditions (Nemec et al., 1980,1984; Gloppen and Steel, 1981; Porqbski, 1984; C.H. Eyles, 1987; McCabe and Eyles, 1988). The main body of sediment is formed by avalanching and mass transport along the inclined sedimentary surface (cf. Postma, 198413; McCabe and Eyles, 1988; Postma et al., 1988a). The contribution from deposition from traction currents is considerable only in the middle and lower parts of the foresets. The lower parts also comprise abundant mass-transport deposits, formed by a number of processes that take place most frequently in channels incised in the delta front. Such channels may be filled later by new-generation mass-transport deposits. Deposition by rain-out processes is fairly continuous but, quantitatively, its contribution is usually small (McCabe, 1986; N. Eyles et al., 1989).

Terminoglacial deltaic bottomsets (II-A-3-d) Numerous authors have described terminoglacial fan and braid deltas from a broad palaeogeographical aspect (e.g., R. Gilbert, 1972a; Church and Gilbert, 1975; Gustavson et al., 1975; P.W.V. Harris, 1977; Shaw, 197713; Edwards, 1978; Shaw et al., 1978; Le Blanc Smith and Eriksson, 1979; Boulton and Deynoux, 1981; Jorgensen, 1982; N. Eyles, 1983, 1987; Ruegg, 1983; Miall, 1985; Drewry, 1986; Fralick and Miall, 1989; McCabe and Eyles, 1989). However, there is little information about the bottomsets of deltas in the terminoglacial subenvironment. The few d a t a available indicate that there is no true difference between these bottomsets and their non-glacigenic counterparts (cf. Bates, 1953; Nemec et al., 1980, 1984; Vos, 1981; Koster and Steel, 1984; Nemec and Steel, 1984, 1988a; Postma, 1986; Postma et al., 1988a; Kano and Takeuchi, 1989). I t is also clear from comparative studies that the terminoglacial deltaic bottomsets formed under lacustrine conditions do not differ significantly from those formed under marine conditions (C.H. Eyles and N. Eyles, 1983a; C.H. Eyles et al., 1985; C.H. Eyles, 1987; Visser et al., 1987). It is obvious from the studies just mentioned that terminoglacial deltaic bottomsets form complex deposits with many subtypes. The sedimentary body consists of prodeltaic bars, slump masses, laminites formed by settling from suspension, turbidites, channel deposits, etc. (Shaw, 1975a;

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The terminoglacial subenvironment

Shaw and Archer, 1978; N. Eyles, 1987a; N. Eyles et al., 198?a, 1988a; Marzo and Anadon, 1988; Postma et al., 1988a; Wood and Ethridge, 1988; Shaw et al., 1989). This complex composition invites misinterpretation, particularly if lateral facies transitions are not exposed or if erosion has destroyed the originally adjacent deposits.

Lithofacies characteristics The bottomsets are an intermediate (complex)deposit between the deltaic foresets - where coarse material and mass-transported units dominate and glaciolacustrine sediments where fine-grained material, settled from suspension, is predominant. Deltaic bottomsets (Fig. 200) are - as a whole - finer grained than the corresponding foresets, but channel deposits and mass-transported material often form coarse intercalations (cf. Shaw, 1975; Shaw et al., 1989).Melting of ice rafts results in the occurrence of pebbles and boulders that are obviously too large for deposition by the same mechanism as the surrounding sediments.

Fig. 200. Terminoglacial fan-delta bottomsets from the Kleszcz6w Zone (central Poland). Note the simultaneous occurrence of finely laminated muds, fine sands with current ripples and dropstones.

Deposits of the terminoglacial deltaic facies

323

Prodeltaic bars may be present and usually consist of sands and muds. The sandy parts show frequent ripple-drift cross-lamination. The sandy parts gradually pass into the more muddy parts, which are characterised by parallel lamination. The bars sometimes contain thin fining-upward cycles, overlain by parapelagic clays. The differences in grain size make the bar deposits sensitive to forces that result in soft-sediment deformations (e.g., load casting). The laminites that form part of the bottomsets and that can be considered as the most typical representatives of the bottomsets, usually consist of sandy to clayey mud; the precise grain-size distribution depends on the material supplied and the current velocities still present at the delta front. The often fine lamnination may be disturbed locally by dropstones. Slumped masses, of deltaic bottomset origin or not, are usually very thin but well recognisable because of the disturbed internal structure and the often diverging grain size (with pebbles commonly included). Other mass-transported deposits, such as turbidites, have lithological properties (and other characteristics) that are exactly the same as seen under non-glacigenic conditions (cf. Postma et al., 198813).They may alternate with other forms of mass-transported sediments, and there may be distinct changes from slump via mudflow to turbidite in a distal direction. The channel deposits consist mostly of gravelly to sandy material (C.H. Eyles, 1987), usually with either a massive or a horizontally stratified appearance. The individual layers and laminae may also be massive, but may show grading. Textural characteristics Most terminoglacial deltaic bottomsets have relatively good sorting, but there may be distinct granulometric differences between the various subtypes of deposits. Only the mass-transported units are commonly characterised by bad sorting. Pebbles generally have a preferred orientation, with their ah-planes subhorizontal (freezing disturbs the originally horizontal position). Elongated clasts usually have their a-axis parallel to the current direction if the clast is relatively large, and perpendicular t o the current if the clast is relatively small or light. Occurrence Terminoglacial deltaic bottomsets form the lowermost part of deltaic successions in the terminoglacial subenvironment. They usually overlie

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terminoglacial lacustrine sediments, with which they also interfinger in a distal direction. The bottomsets form relatively thin and widespread deposits surrounding (for some part) other deltaic deposits in a configuration that is ideally half a circle. There is gradual thinning of the deposits in a downstream direction, so that the passage into lacustrine bottomsets is gradual. Depositional mechanisms

Several processes contribute to the deposition of terminoglacial deltaic bottomsets and which one will predominate depends on local and temporary conditions. Sand-flow processes, resulting in laminated muddy sands, dominate in the most proximal part. A large sediment supply by sufficiently strong currents results in simple or compound tabular crossbedded units (Buck, 1985; R. Hunter, 1985; R. Hunter and Kocurek, 1986). Proximal channels are usually filled by massive, sometime gravelly, sands that represent mass transport. The upper parts of such channels are usually filled with finer material, commonly sand, which shows crossbedding and parallel lamination. Grading may also be present, with muddy clays as the top part, formed by settling from suspension under quiet conditions, as during winter time when an ice layer covers the lake. Ripples present in the prodeltaic bars are typically produced by rapidly decreasing traction. This may be compared with conditions of waning flow. As current velocity drops with distance from the place where a meltwater stream embouches in the terminoglacial lake, relatively quiet conditions may prevail downstream of the bars, so that even the finest particles can settle. Fine-grained turbidites may, however, constitute a significant part of the 'varved' bottomsets. Only dropstones, derived from melting ice rafts, may disturb the quiet appearance. DEPOSITS OF THE TERMINOGLACIAL LACUSTRINE FACIES (11-A-4) The terminoglacial subenvironment is richly supplied with lakes (Fig. 201). This is due to the continuous supply of meltwater on the one hand, and to the presence of topographic barriers (dead-ice, diamictic hummocks, etc.) on the other hand. The characteristics (e.g., size and depth) of these lakes are fully determined by the local relief dynamics (Edwards, 1978; Schluchter, 1979; Reineck and Singh, 1980; N. Eyles, 1983b; R.G. Walker, 1984; Drewry, 1986; Gradzinski et al., 1986). Deposition in the

Deposits of the terminoglacial lacustrine facies

325

Fig. 201. Terminoglacial lake in the Hornsund area on Svalbard. Note the irregular topography which is responsible for continuous changes in the number, size and geometry of lakes. Photograph: J. Cegfa.

basins is extremely complex, particularly because of the dynamics (high energy gradients) within the terminoglacial subenvironment. There is a strong influence of the melting ice masses, and mass-transport phenomena are common. The characteristics of the lake margins are also influenced by fluvial supply and by typical littoral and sublittoral processes (C.H. Eyles and N. Eyles, 1984b; N. Eyles and Miall, 1984; Drewry, 1986; C.H. Eyles, 1986; Brodzikowski and Van Loon, 1987). Most terminoglacial lakes have two remarkable features in common: the occurrence of marks left by ice rafts, and varves. Ice rafts, floating in the lakes after calving of the ice front, leave traces in the lakes in the form of dropstones (including dump sediments, etc.). The ice masses may also deform the lacustrine sediments, viz. if they are large enough to touch the lake bottom. Such 'ice gouging' is best known from marine terminoglacial conditions (Darwin, 1855; I.McK. Harris and Jollymore, 1974; L.H. King, 1976; Barnes et al., 1984; Woodworth-Lynas and Barrie, 1985; Woodworth-Lynas et al., 1985a,b), but the process has also been described from terminoglacial lakes (Thomas and Summers, 1982,1984; Thomas and Connell, 1985).

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Varves are graded layers that often extend over the entire area of the lake. They were long interpreted as being a result of season-induced sedimentation, but this simplistic view has now been abandoned (cf. Merta, 1978; Shaw and Archer, 1978; Hahszczak, 1980; Brodzikowski and Van Loon, 1983,1987; N. Eyles and Miall, 1984; R.G. Walker, 1984). Processes like turbidity currents and settling from a more or less continuous inflow of particle-laden meltwater are now known also to play important roles. It is interesting in this respect that meltwater inflows in terminoglacial lakes are almost exclusively hyperpycnal or homopycnal. This is due t o the fact that meltwater from the ice front contains a high concentration of fines (proglacial lakes more often have hypopycnal inflows although homopycnal and hyperpycnal inflows also occur). The character of the meltwater influx and the associated conditions of resuspension are major factors in the depositional processes in terminoglacial lakes (Drewry, 1986). In general, the water and sediment supply in the lakes depends mostly on the dynamics of the ice and on ablation conditions. This implies that there is a wide variety of meltwater influxes. N.D. Smith (1978) distinguished five inflow regimes on this basis: diurnal, subseasonal (duration of a few hours to several days, dependent on local weather conditions, and on glacial drainage), seasonal (related to snow melt in spring and to ice melt in the summer), annual, and exceptional events (such as jokulhlaups). This differentiation might be a key t o the cyclicity in many glaciolacustrine (particularly terminoglacial) lakes. It can occur that an advancing ice mass overrides a terminoglacial lake. If the lake is deep enough, there may be a timespan during which part of the lake is covered by a floating shelf of glacial ice. There will then be intensive undermelting, which strongly alters the water supply; the dynamics of the lake water will also be strongly influenced. There is, however, little information available about such conditions and their implications for sedimentation. Much more is known about the the influence of temperature differences between inflowing streams and lake water on the mixing conditions. It is well known that the mixing conditions affect the settling of suspended particles (Hutchinson, 1957; Gage, 1975; Jolly and Irwin, 1975; Pickrill et al., 1981; Irwin and Pickrill, 1982; Hakanson and Jansson, 1983). In spite of all variability in the depositional conditions within terminoglacial lakes, it is only useful to distinguish between two types of lacustrine sediments, plus an additional type for cases in which differentiation is not possible (Fig. 202). Thus, we distinguish terminoglacial lacustrine complexes (II-A-4-a),terminoglacial lake-margin deposits (II-A-4-b) and terminoglacial lacustrine bottomsets (II-A-4-c).

Deposits of the terminoglacial lacustrine facies

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Fig. 202. Schematic model of the terminoglacial lacustrine facies. Partly based on C.H. Eyles and N. Eyles (1984a), C.H. Eyles et al. (1985b) and Brodzikowski and Van Loon (1987).

Terminoglacial lacustrine complexes (II-A-4-a) The complexes are usually deposits formed in small lakes, so that a true sedimentological distinction between bottomsets and lake-margin deposits cannot be made. In other cases, large lacustrine sedimentary bodies have been glaciotectonically disturbed, so that the two subtypes have been mixed in such a way that it is no use separating the deformed units. The lacustrine origin of these complexes, which may be deformed together with fluvial and/or deltaic deposits, is commonly easy to reconstruct on the basis of the joint occurrence of varves and dropstones, without indications of much current action. There are numerous reports touching complexes such as these (e.g., Klatkowa, 1965, 1972a,b, 1982, 1987; Kozarski, 1967, 1981; Rbiycki, 1972; Drozdowski, 1974, 1979, 1986; Krygowski, 1974; Olszewski, 1974; Kozarski and Szupryczyfiski, 1978; Brodzikowski and Van Loon, 1980, 1983, 1987; Kfysz, 1981; Brodzikowski, 1982, 1984; Van der Meer, 1983; Dreimanis, 1989),but no detailed sedimentological analyses were carried

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out. It is, however, generally agreed upon that most of the complexes consist of a mixture of dropstone-containing varvites and mudflow- or slump-like deposits (flow tills, lacustrotills). Lithofacies characteristics Terminoglacial lakes usually form the first real trap for sediments set free from the ice mass by ablation. Consequently, the lakes are filled with large amounts of particles of all sizes, so that terminoglacial lakes are generally speaking - relatively coarse. This implies that terminoglacial lacustrine complexes do not show a clayey or muddy lithology, but generally have much sand included, even in the centre. The sandy parts often show current ripples and parallel lamination, sometimes wave ripples also. A high rate of sedimenting out from suspension may be reflected by climbing ripples. Dispersed dropstones may float within the sands. The more common alternations of sands and muds (Fig. 203) are generally somewhat less regular, but lamination - sometimes in the form

Fig. 203. Terminoglacial lacustrine complex from Potwor6w (Sudetic foreland, SW Poland).

Deposits of the terminoglacial lacustrine facies

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of varves - is commonly present. There may also be relatively irregular alternations of more sandy and more muddy units, sometimes in the form of lenses. It is not uncommon t o find that these complexes have sandy channel fills embedded in finer material. Coarse clasts may be present, both as channel lag deposits and as dropstones. Diamictic intercalations are a common phenomenon in terminoglacial lakes. They may consist of dump deposits, but are more often masstransport deposits. Silty lacustrine deposits may be broken up during mass transport along the inclined lake bottom; angular fragments thus formed may constitute an 'intrabasin breccia', that still shows - for instance - varves. When all these different lithologies are considered together, it can only be concluded that there is no such thing as a characteristic lithology of terminoglacial lacustrine complexes. Textural characteristics

The texture of these sediments is not very specific and depends on the local conditions and on the material supplied. There may be parts with bad sorting and parts with good sorting, dependent on the prevailing depositional mechanisms. Clasts are usually orientated with their a/b-planes parallel t o the sedimentary surface, but random orientations may occur in the case of dumpstones. Occurrence

Terminoglacial lacustrine complexes may be the most numerous glacigenic deposits, even though their joint volume may be comparatively small. Many small lakes and pools are found in the immediate front of the ice, and the irregular relief favours the genesis and termination of lakes. This is expressed in the geological record, as these deposits are frequently found a s relatively small sedimentary bodies embedded in other terminoglacial deposits. The complexes may be overlain directly by glacial deposits, because an advancing ice mass may override the lake. In such a case, the complexes are often deformed and possibly mixed with frontal crevasse deposits. The thickness of the bodies is highly variable, but it rarely exceeds 2-3 m; the authors are not aware of reports of complexes more than 10 m thick. This distinguishes these bodies, with their often chaotic lithology, from crevasse deposits (cf. Klatkowa, 1972a).

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The terminoglacial subenvironment

Depositional mechanisms The most important depositional mechanism is the action of traction currents, which is how most sandy units are formed. Settling from suspension usually plays a less prominent role, particularly because most lakes in which the complexes are formed are too small to have very quiet water masses. A result is that most deposits formed by settling from suspension have a somewhat irregular appearance, due t o the simultaneous activity of other depositional processes. Rapid sedimentation on inclined sedimentary surfaces promotes the formation of mass flows. Slumps and mudflows are more common than turbidity currents, mainly because the area involved is too small for turbidity currents to develop from mudflows. Winter periods, when the lake is covered with ice, may result in quiet intervals during which sediment may settle from suspension, thus forming varves. Generally, however, varves form only a minor constituent of terminoglacial lacustrine complexes, because bottom currents rework previously deposited material.

Terminoglacial lake-margin deposits (11-A-4-b) The water level of terminoglacial lakes frequently fluctuates. This implies that the extent of the truly lacustrine facies changes with time, and that there is a vague transition between the lake-margin deposits formed within the lake and the sediments that surround the lake. So as t o keep the discussion consistent, only the truly lacustrine deposits, i.e. those formed under subaqueous conditions, are dealt with here. Few sedimentological studies have been devoted t o terminoglacial lake-margin deposits, but reports from several more general studies mention such sediments (e.g., C.A.M. King, 1956; Hine et al., 1974; Nummedal et al., 1974; Nummedal, 1975; Hine and Boothroyd, 1978; N. Eyles et al., 1985; C.H. Eyles, 1986; N. Eyles and Clark, 1986). Pyokari (1981) mentioned that the borders of the lake can be damaged by pushes from lake ice; this process may result in overthrusts, simple horizontal faults, intrusions along zones of weakness, ploughing of the sediments, and the formation of ice-pushed ridges several metres high, consisting predominantly of lake-margin deposits. There are two reasons for such ice pushes: thermal expansion of the ice (Alestalo and Haikio, 1979) and wind action. If lake ice is pushed towards the lake margin, the effects can be increased by wave action and by cryostatic pressure (Dionne, 1976). Single ice rafts, left on the shore, may produce heaps of sediment after melting.

Deposits of the terminoglacial lacustrine facies

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The processes that determine deposition in the lake-marginal zone do not differ from those taking place under non-glacigenic conditions. However, wave action is usually of minor importance, because an ice cover may prevent wave formation during long periods of the year. It was nevertheless reported by N. Eyles and Clague (1987) that wave action can induce liquefaction of deposits at the water level (the same phenomenon is better known from anthropogenic basins: Jones et al., 1961; Pariseau and Voight, 1979; Schuster, 1979). It is also known that the dynamics of an active-ice front may cause large waves. Katabathic winds may induce seiches and more modest wave activity that suppresses longshore sediment transport (Drewry, 1986).

Fig. 204. Terminoglacial lake-margin deposits from the Kleszczow Zone (central Poland). The lenticular stratification was produced by currents in the upper part of the inclined sedimentary surface of the lake at its farthest point from the influx area. Tidal activity may have influenced the current pattern and, consequently, the extent and internal structures of the sediments.

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The terminoglacial subenvironment

Lithofacies characteristics The lithofacies characteristics seen under terminoglacial conditions are not very specific. The occurrence of hummocky cross-stratification is a criterion of sublittoral deposition (Harms et al., 1975; Murty and Polarvaparu, 1975; Van Dijk et al., 1978; Hamblin and Walker, 1979; R.G. Walker, 1984; C.H. Eyles, 1986). Barrier spit complexes are usually sandy, possibly with some gravel; these complexes may show swash-bar, berm-ridge, berm-top and beach-face types of stratification, depending on the size of the lake and of the spit (Hine and Boothroyd, 1978). Relatively small lakes are generally characterised by finely laminated sandy or sandy/muddy deposits; the bottom-parallel lamination is usually discontinuous. Periods of stronger currents may be reflected by sandy layers with cross-stratification (Fig. 204). The appearance may be disturbed by the presence of dropstones. Diamicts may be present, either as dumpstones or as mass-transported sedimentary bodies.

Textural characteristics Granulometry of these deposits depends strongly on the material supplied, but continuous reworking of the material is responsible for an overall good sorting per layer (except for mass-transported units). Orientation of larger clasts is parallel t o the lake bottom, with the exception of clasts in slumps and in dump deposits, which do not show a preferred orientation.

Occurrence These deposits have a limited areal extent and are generally no thicker than 1-3m. By definition, they surround lacustrine bottomsets but, due t o a gradual shallowing of the lake, may also be present on top of lacustrine bottomsets. There may also be lateral contacts with deposits of the fluvial, aeolian, deltaic and melting-ice facies. As in the case of terminoglacial deltaic complexes, these deposits may frequently be found in the terminoglacial interval of the glacigenic sequence. In an exposure, they may seem to be discontinuous because the irregular shape is responsible for only partial exposure.

Depositional mechanisms Wave action, current action and mass transport are the main mechanisms that determine lake-margin sedimentation. Sand flow plays an important

Deposits of the terminoglacial lacustrine facies

333

role in the formation of slightly inclined sandy layers. This process takes place under low-energy conditions. Settling from suspension requires even quieter conditions - and indeed occurs, mainly during the winter - but the laminated deposits thus formed are usually reworked in the course of time so that relatively few traces are left. Wave action may produce a hummocky stratification (Hunter and Clifton, 1982; Leckie and Walker, 1982; Clifton and Dringler, 1984; Leckie and Duke, 1984; Diem, 1985; C.H. Eyles, 1986; N. Eyles and Clark, 1986), but may also induce liquefaction through pressure exerted on water-saturated silty sediments. A result of this is that small turbidity currents may be formed, sometimes starting as simple grain flows that gradually develop into cohesive flows. Tidal action may play a role in lakes of extreme size (Reineck and Singh, 1980), but the dynamics of a floating or grounding ice front may induce processes that give similar results in smaller lakes. Melting of ice rafts in the lake results in dropstones and dumpstones. Ice rafts landed on the beach by wave or wind action will also produce badly sorted heaps of sediment.

Terminoglacial lacustrine bottomsets (11-A-4-c) These deposits, which are sometimes termed 'distal lacustrine deposits' in contrast to 'proximal' lacustrine deposits (Ashley, 1975; Merta, 1978; Shaw and Archer, 1978; Brodzikowski, 1984), are frequently described in general textbooks (among others by Jopling and McDonald, 1975; Sugden and John, 1976; Reading, 1978; Schluchter, 1979; N. Eyles, 1983b; R.G. Walker, 1984; Drewry, 1986; Gradzifiski et al., 1986) and even more frequently in more specialised sedimentological reports. The fact that so much attention is devoted t o these deposits is certainly due - at least partly - t o the fact that the majority of these deposits consist of varvites (Fig. 205), which can be considered as the most prominent terminoglacial deposit and, in addition, as one of the most characteristic glacigenic sediments. The question of whether these varvites are predominantly formed by settling from suspension or by turbidity currents is still much debated. This problem will be dealt with in more detail in the section on the depositional mechanisms of these sediments. Many of the terminoglacial lakes are very shallow, even i n their central parts. Sedimentation is thus easily affected by processes like wave action, temporary subaerial exposure, ice gouging, etc. The activities of so many processes have led researchers to distinguish between subtypes of deposits such as lake-bottom channel deposits, levees and lake-bottom

334

The terminoglacial subenvironment

Fig. 205. Varvites from the central part of a terminoglacial lake. Note the continuous character of the individual laminae and their grading. Irregularities result from metadepositional (penecontemporaneous) deformations. Exposure near Lekilisko (Kleszcz6w graben, central Poland).

plains (Shaw, 1975; Merta, 1978; N. Eyles, 1983b; N. Eyles and Miall, 1984; C.H. Eyles et al., 1985; Shaw et al., 1989). Even though the occurrence of several subtypes of deposits might suggest that terminoglacial lacustrine bottomsets have a very variable appearance, this is generally not the case: rhythmites (in this context sometimes termed 'varvites') usually predominate. It must be admitted, however, that the rhythmites in this context are often developed in a way somewhat different from those in the proglacial equivalents (II-B-4-d), possibly because the annual cycles of freezing and thawing are more pronounced at points farther away from the ice than only at the ice front (Ashley, 1972; Gustavson, 1972; Banerjee, 1973; R. Gilbert, 1973, 1975; Shaw, 1977b; Shaw et al., 1978; N.D. Smith, 1978; R. Gilbert and Shaw, 1981; N.D. Smith and Syvitski, 1982; N.D. Smith et al., 1982; Brodzikowski and Van Loon, 1983,1987).

Deposits of the terminoglacial lacustrine facies

335

Lithofacies characteristics Most terminoglacial lacustrine bottomsets consist of rhythmites built up of laminae (or layers) that go gradually from sandy mud a t the base to pure clay a t the top, Some sandy units may be present as a result of supply by bottom currents; this is reflected by the presence of current ripples, sometimes in the form of ripple trains (Fig. 206). There may also occur more massive sandy strata, but these are rare. Massive muddy units are much more common, and diamictic lenses are also frequently found. They result from processes not characteristic of lacustrine sedimentation. Another common phenomenon is the occurrence of levels with softsediment deformations. These may be the result of an unstable density gradient (e.g., load casts), or they may represent slumps. It is not uncommon that a specific level (e.g., a horizon with a high silt concentration) is highly sensitive to disturbances, so that even the smallest stress may trigger various processes simultaneously, such as flowage of loading material (Fig. 207).

Fig. 206. Terminoglacial rhythmites with a few sandy horizons showing distinct signs of current activity.

336

The terminoglacial subenvironment

Fig. 207. Deformed horizon in typical lacustrine bottomsets (Kleszcz6w graben, central Poland).

Textural characteristics

Good sorting of most units is characteristic, although dropstones - floating in the rhythmites like raisins in a cake - are commonly present (Fig. 208). The grading in the separate varves is a most conspicuous feature; the contact between the relatively coarse base and the relatively fine top of the underlying layer is usually sharp, especially in distal settings (Fig. 209), sometimes with load casts or comparable structures. Diamictic units show no preferred orientation of the clasts (the longest axis may even be directed vertically: Eyles et al., 1987b), but single dropstones, if flattened, are orientated with their ah-planes parallel t o the sedimentary surface. There may also be a preferred orientation of the clasts in some mudflows, parallel t o the internal flow lines. Occurrence

These deposits are, by definition, surrounded by terminoglacial lakemargin deposits but they also frequently form the downstream continua-

Deposits of the terminoglacial lacustrine facies

337

Fig. 208. Varvites with a high frequency of floating dropstones.

Fig. 209. Distinctly laminated (varve-like) terminoglacial lacustrine bottomsets in a typically distal setting (Kleszczow graben, central Poland).

338

The terminoglacial subenvironment

tion of terminoglacial deltaic bottomsets (11-A-3-d).Sharp boundaries with these types of deposits are rarely seen. Sharp boundaries do exist with a number of other types of deposits, depending on the local situation. If, for example, the terminoglacial lake is situated in front of a n advancing ice mass the lake will be overridden by the ice with the result that - if the ice advance was sufficiently fast - badly sorted glacial till are found directly on top of the lacustrine bottomsets. In the glacigenic sequence, these bottomsets - i n combination with lake-margin deposits - can be found all over the terminoglacial interval.

Depositional mechanisms

A wide variety of processes contributes to the sedimentation in the central part of a lake. The most important of these is settling from suspension clouds (Fig. 210). This process is a result of sediment supply by meltwater streams, particularly during the summer. Far fewer meltwater streams supply material during the winter, when an ice layer covers the lake,

Fig. 210. Schematic model of settling from suspension clouds in a glaciolacustrine lake. GFI = glaciofluvial input; SC = suspension current of overflow; W = wave activity; IF = interflow; BC = bottom current; Tllc = high-concentration turbidity current; T1.c = low-concentration turbidity current; PP = parapelagic sedimentation. 1 = overflow suspension sheets drifting across the lake; 2 = waves along the lake margin, triggering sandflows along the slope; 3 = sediment input from the lake shore; 4 = turbidity current or traction current; 5 = lateral extension of the suspension cloud in the thermocline zone; 6 = top of the suspension cloud; 7 = mixing zone; 8 = settling from interflow and overflow; 9 = secondary bottom currents; 10 = reworking by turbidity current; 11 = maximum suspension concentration; 12 settling of suspended particles from the mixing zone; 13 = substratum.

Deposits of the terminoglacial lacustrine facies

339

however subglacial streams may embouch in the lake during winter, below the ice level. Sedimentation by settling from suspension (see also Sturm, 1979) is increased when the winter ice layer melts away as the temperature rises. The ice layers contain much mineral material, particularly at their surface, as a result of wind activity and mass wasting. A relatively large amount of material may thus be set free within a short period. The coarser material sinks down immediately and the finer material follows gradually; i t is only during the following winter, when the lake is again covered by a n ice layer, that the water becomes quiet enough t o allow the finest clay particles to settle. The result of this sequence of processes is the formation of a season-bound graded two-unit layer, which is commonly termed 'varve'. The above interpretation of the depositional mechanism has long been considered correct, a t least until the 1950s. It was found in this period, when much sedimentological attention was given t o t h e concept of turbidity currents (Kuenen, 1950; 1953; Kuenen and Migliorini, 1950; Kuenen and Menard, 19521, t h a t many glaciolacustrine varves show

Fig. 211. Non-sorted dumpstone (upper part) on top of varved lacustrine bottomsets with two layers of dropstones. Kleszczow graben, central Poland.

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The terminoglacial subenvironment

characteristics that are well comparable with those of turbidites (Kuenen, 1951),for instance the presence of the so-called Bouma intervals (Bouma, 1962), with grading as one of the most conspicuous features. The opinion was held for some time that sedimentation, as seen in turbiditic fans (R.G. Walker, 1966, 1975, 1976, 1984; Stanley, 1969; Normark, 1970, 1974, 1978; Ricci-Lucchi, 1970, 1975; Mutti and Ricci-Lucchi, 1972, 1975; Walker and Mutti, 1973; Mutti, 1974, 1977, 1979; Ghibaudo, 1979; etc.), was the predominant mechanism. In sedimentology, new concepts lead to new interpretations. It was realised by several glacial geologists that many varves do indeed show some characteristics of turbidites, but that only very few have such an origin beyond any doubt. In addition, De Geer based an impressive late Weichselian chronology for Scandinavia on varves (e.g., De Geer, 1912); this chronology later appeared to be so consistent with absolute datings (in spite of some discrepancies) that a relationship between varves and yearly cycles cannot be denied. Most workers in this field are therefore of the opinion that most varves consist of season-bound cycles, and that only a minor part (in general) is due to turbidity currents. New insights into

Fig. 212. Varved terminoglacial lacustrine bottomsets with several levels of dropstones, formed by undermelting of ice rafts.

Deposits of the terminoglacial aeolian facies

341

the settling mechanisms have resulted from the many studies devoted t o this subject. The mechanisms of settling from overflows, interflows and underflows have been discussed in detail by N.D. Smith and Syvitski (1982) and by Syvitski et al. (1989) in a n attempt to explain why the fine particles settle faster than predicted by Stoke's law. There occur other types of sediments in addition t o the varves. For instance, there are massive muds, sands or diamicts that show no grading. If sandy, the massive units are interpreted as being the result of a mass flow t h a t froze suddenly when the shear stress in the mass decreased; sandy flows may be formed due to a hyperconcentrated flow, described as 'subaqueous floods' by Shaw (1975) and Shaw et al. (1989). If muddy, the massive sediment bodies may have formed by rapid settling from a large suspension cloud, such as may be triggered by a n earthquake or by the collision of two mass flows from opposite sides of the basin. Massive diamicts (Fig. 211) may be a result of mudflows (Middleton and Hampton, 1976) or, more appropriately, of sudden turning around of debris-laden ice rafts (dumpstones) (C.H. Eyles et al., 1985). Gradual undermelting of ice rafts results in pebble-rich horizons intercalated between the varved units (Fig. 212). DEPOSITS OF THE TERMINOGLACIAL AEOLIAN FACIES (11-A-5) The terminoglacial subenvironment is a source of windblown material rather than a depositional site. It is therefore not surprising that there are no detailed sedimentological reports on 'fossil' terminoglacial aeolian deposits, although aeolian accumulations are well known to exist under such conditions in modern environments. Most of those working in this field make detailed mention of the occurrence of aeolian deposits from the proglacial and extraglacial subenvironments, with occasional reference to such deposits in the terminoglacial subenvironment. Examples of such reports are, e.g., those for Svalbard (Andree, 1883; De Geer, 1901; Hogbom, 1911,1912; Nansen, 1920; Samuelson, 1921; J a h n , 1961b, 1975; Szupryczyiiski, 1963; Czeppe, 1968; Akerman, 1980; Baranowski and Pqkala, 1982; Szczypek, 1982; Migata and Sobik, 1984; Kida, 1986), Greenland (Hobbs, 1931; Teichert, 1935; Belknap, 1941; Flint, 1948; Troelsen, 1949, 1952; Nichols, 1969; Fristrup, 1953; Cegla, pers. comm., 1982), t h e Canadian Arctic (Bird, 1951; Pissart, 1966; French, 1972; Pissart et al., 1977), Alaska (Black, 1951; Pewe, 1955;Jahn, 1961a, 1966;Trainer, 1961; Fernald, 1965), the Antarctic (Webb and McKelevey, 1959; McCraw, 1967; Nichols, 1971; Lindsay, 1973; Selby et al., 1974) and Iceland (Bout, 1953;

342

The terminoglacial subenvironment

Bout et al., 1955; C.A.M. King, 1956; Ashwell, 1966,1972; Bogacki, 1970; Shunke, 1975). The terminoglacial subenvironment in general is wet, but frost conditions and wind activity may create dry areas, particularly in somewhat elevated places. Such dry places form deflation areas, where much material may be blown away. Transport is usually over short distances only, and most sand and silt grains accumulate in areas (sand drifts) where they may be aeolisated several times before they finally come to rest in drift-sand bodies. It appears that episodes of intensive aeolisation are generally short (Jahn, 1961a,b; Akerman, 1980; Kida, 1986; PFkala and Wojtanowicz, 1987): of the order of minutes or hours at most. Widespread accumulations of the order of some millimetres thick may form during such 'storm periods' (Ceg*a, pers. comm., 19821, whereas local heaps several centimetres or even decimetres thick may be formed (Akerman, 1980; Kida, 1986). There are no reports of true dunes being formed in this manner. Because of their very restricted size, the aeolian accumulations formed in the terminoglacial subenvironment should all be grouped together as drift sands (II-A-5-b).

Terminoglacial drift sands (11-A-5-b) The literature provides only sparse data about these deposits. Akerman ( 1980) has described them in his dissertation, and other data can be found in works mainly devoted to geomorphological aspects (Bogacki, 1970; Baranowski and Pykala, 1982; Szczypek, 1982; Kida, 1986).

Lithofacies characteristics Drift sands (Fig. 213) are usually massive or horizontally stratified (laminated). The first type is the more common one and occurs particularly where sand has accumulated in the shadow of a n obstacle (the same situation is found on modern tundras). It is not unusual t o find drift sands, particularly if they contain a relatively high proportion of silt, draped over irregularities (erratics, etc.) and filling the pores between gravelly deposits (e.g., subaerially formed mass-flow deposits). Deposits that consist mostly of sand may be 'reworked' by wind action, so that a thin surface layer of aeolian material (drift sand) originates with wind ripples. The preservational potential must be extremely low and no reports on 'fossilised' wind ripples from this subenvironment are known.

Deposits of the terminoglacial aeolian facies

343

Fig. 213. Small patches of drift sands in and around terminoglacial crevasses (Hornsund area, Svalbard).

Textural characteristics Individual sand grains in these deposits show relatively few signs of aeolian transport (or no such signs a t all), which is due to the usually very short distances travelled. The grains, which had previously undergone glacial grinding, frost splitting, glaciofluvial abrasion, etc.) are therefore rarely rounded; subangular-subrounded shapes are the most common. Due t o the limited transport, sorting is also bad in relation to that of most other aeolian deposits (Fig. 2141, so that a mixture ranging from silt (sometimes transported in pellets: Pye, 1987a) t o coarse sand is common.

Occurrence Nowadays, these deposits are found throughout the terminoglacial subenvironment, particularly in association with bars and hummocks. They may also be found on nunataks (Pekala, 1980; Baranowski and Pekala, 1982) and between dead-ice blocks (Migala and Sobik, 1984; Kida, 1986). The drift sands usually form irregular blankets only a few centimetres thick (Pekala and Wojtanowicz, 1987).

The terminoglacial subenvironment

344

O/O

-

o/o

A

LO-

LO

3020.

!

10-

o-

10 ,

r, ,

.

,

L1

,

-

-

Fig. 214. Granulometric composition of deposits in (A) a terminoglacial lake a n d ( B ) aeolian material in snow barchans. From: Migata a n d Sobik, 1984. Courtesy: Zeitschrift fur Gletscherkun-

Their small extent, particularly in a vertical direction, makes the drift sands exceptionally sensitive to erosion or reworking. This implies t h a t they will probably not be found in a n actual glacigenic sequence, although their occurrence can, of course, not be fully excluded. However, should they be present, it might be extremely difficult to recognise these deposits as being of aeolian origin.

Depositional mechanisms The deposits a r e formed because wind-driven particles, mostly i n saltation, come to rest when wind activity decreases or when a place on the lee-side of a n obstacle is reached. Smaller particles may be suspended in the air, but they generally do not make a large contribution t o terminoglacial drift sands because such dust may easily be blown further away (Pye, 1987a). Particles driven over the surface by the wind may be caught more easily by the substratum if it is wet (water-saturated soil), or if a water body is present; wind-blown particles settling in a pool or a lake are mixed with particles brought into suspension by other agents (e.g., inflow of meltwater), and cannot easily be distinguished - if at all - from nonaeolian particles.

DEPOSITS OF THE TERMINOGLACIAL MASS-TRANSPORT FACIES (11-A-6) The large quantity of meltwater in the terminoglacial subenvironment results in only a relatively small area not being covered (temporarily or

Deposits of the terminoglacial mass-transport facies

345

continuously) by flowing or stagnant water. The 'dry' areas in between do not receive much sediment but, due to their irregular topography (partly

Fig. 215. Sediments on top of dead-ice bodies in the terminoglacial subenvironment of Svalbard are affected by melting of the ice. This results in collapse structures (top) which are accompanied hy a large number of mass-transport phenomena, ranging from 'block fall' to fluidised flow (bottom). Photographs: J. Cegta.

346

The terminoglacial subenvironment

caused by phases of ice advance), undergo mass movements that can considerably alter the original position of large sedimentary masses (Fig. 215). These mass-transport processes are well comparable with those that occur in the supraglacial mass-transport facies (1-A-6). Three types of deposits can thus be distinguished: terminoglacial subaerial masstransport deposits (II-A-6-a), terminoglacial crevasse deposits (11-A-6-b) and terminoglacial subaqueous mass-transport deposits (II-A-6-c). All these types of mass-transport deposits have been described frequently, although with different emphasis and especially under a wide variety of names. Some of the more interesting general works on this topic are those by Klatkowa (1972a,b),Embleton and King (1975),Jahn (19751, Elson (1976), French (1976), Klimaszewski (1976), Sugden and John (1976), Lowe (1979, 1982), Washburn (1979), Cohen (1983), N. Eyles (1983b), Postma et al. (1983), R.G. Walker (1984), Broster and Hicock (1985), Drewry (1986), Postma (1986), N. Eyles et al. (1987a) and M.J. Clark (1988).

Terminoglacial subaerial mass-transport deposits (II-A-6-a) These deposits have mainly been studied by periglacial geomorphologists (Jahn, 1975; French, 1976; Washburn, 1979; M.J. Clark, 1988). The reason for this is that, of the agents supporting this mass-transport, an important one is ground ice, the presence of which results in the creep of permafrosted layers, frost creep, gelifluction, etc. (Lewkowicz, 1988). A disadvantage of having interest shown by geomorphologists rather than by sedimentologists is that characteristics of special sedimentological interest have not always been described sufficiently accurately, whereas much emphasis was placed on the consequences that subaerial mass transport could have for the relief (Jahn, 1960; P.J. Williams, 1966; Gardner, 1969; Benedict, 1970,1972; Chandler, 1972; Czudek and Demek, 1972; Harris, 1973, 1981; French, 1974; Mackay and Matthews, 1974; McRoberts and Morgenstern, 1974; L.W. Price, 1974; McRoberts, 1978; Varnes, 1978; Caine, 1980; Dyke, 1981; Lewkowicz, 1981; Mackay, 1981; Woo and Steer, 1982; Heginbottom, 1984). It is typical of this state of affairs that there are far fewer sedimentological (e.g., McCann et al., 1972; Church et al., 1979) than geomorphological studies. Lithofacies characteristics The lithofacies characteristics are obviously fully determined by the source material that undergoes mass transport. Under terminoglacial

Deposits of the terminoglacial mass-transport facies

347

conditions, subaerial mass transports usually involve badly sorted material with a wide range of grain sizes. Depending on the precise transport mechanism, this material can be deposited as structureless, matrixsupported diamicts, as badly sorted bodies with a vague stratification, which is, in most cases, of secondary origin (i.e., formed during the mass transport), or as silty slump balls. Material that has slid down along an inclined surface is, under these conditions, commonly termed lflow till', whether the transport process was slumping, sliding, flowage or another process, and whether the sediment is fine-grained, coarse-grained o r badly sorted. The main reason for applying this term is the usually diamictic character and the glacial origin of the clasts (Boulton, 1972; Marcussen, 1973; Van der Meer, 1983); subaqueous mudflows formed under glacigenic conditions are, however, also commonly termed Iflow till'. It can only be deduced that this term gives little information about the depositional mechanism and about the litholcgical characteristics. A wide range of grain sizes in the material being reworked may result in secondary lamination, i.e. when the grains are sorted and orientated as a result of internal shearing. There is, however, rarely regular and continuous lamination, but rather a series of smudges and shadows of sand or silt. Secondary lamination is formed most easily if mass transport takes place under water-saturated (but still subaerial) conditions, although it is also known from deposits formed due to a slow, very dense, creep-like transport mechanism (Jahn, 1975; Washburn, 1979; M.J. Clark, 1988). Experience has shown that most deposits of this type are coarse, but glaciolacustrine sediments that become reworked may be fine-grained (often with a high silt content, which favours the formation of softsediment deformations). Units with a high content of cohesive material in particular may form well developed disharmonic folds during the masstransport process. Textural characteristics

The terminoglacial subenvironment is characterised by an abundance of much badly sorted material. Subaerial mass-transport processes do not, in general, contribute to better sorting. It is therefore not surprising that most of these deposits are very badly sorted (Fig. 216). The orientation of the individual particles is often random, particularly if so many large clasts are present that the insufficiency of space does not allow them to find an optimum position. Mass-transport mechanisms

348

The terminoglacial subenvironment

Fig. 216. Badly sorted deposits-formed by subaerial mass wasting of till in the terminoglacial subenvironment (Zary area, western Poland). Note the vague layering, which may be an indication of deposition in several phases.

under water-saturated conditions may, however, result in a preferred orientation of the larger clasts with the a/b-plane subparallel t o the sedimentary surface. Imbrication has occasionally been reported. Graded bedding is not present, as the only mass-transport mechanisms leading to this type of sorting require subaqueous conditions, but bad mixing within a mudflow may result in vague, always very local, grading. The petrographic properties are determined by the source material. Occurrence

Most of the terminoglacial subaerial mass-transport deposits are found between deposits of the melting-ice facies and those of the fluvial facies. They also occur on top of (and as intercalations within) the terminoglacial deltaic and lacustrine facies. Most deposits form irregular bodies that gradually thicken in a downslope direction (there may also be a relatively thin 'tail' and a suddenly much thicker 'head', e.g., in the case of a slump). These irregular masses are often found associated with sediments that show collapse structures

Deposits of the terminoglacial mass-transport facies

349

formed by melting of dead-ice blocks. This is understandable, a s ice melting had changed the relief, thus causing, in most cases, a slope along which instabilised material might start to flow.

Depositional mechanisms Several subaerial mass-transport mechanisms affect the sediments in the terminoglacial subenvironment. The abundant inclined surfaces, the sudden changes in relief, the many places with non-permeable soil, the continuous supply of particles of many sizes, and several other conditions create a situation that is most favourable for mass transport. Sedimentologists commonly use a relatively rough classification (e.g., rock fall, creep, slippage, slumping, sliding, solifluction - often termed 'gelifluction' under these conditions -,mudflow, etc.). Some sedimentologists group all these processes together as 'subaerial debris flow'. Geomorphologists often go into more detail. Lewkowicz (1988) distinguishes, for instance, between two main groups of processes: slow a n d fast processes. The first category comprises the creep of permafrost, frost creep, needle-ice creep and gelifluction, whereas the second category includes skinflow, mudflow, multiple retrogressive flow, active-layer gliding, slumping, multiple retrogressive sliding, rock and block fall, debris flow and ground-ice slump. In addition, Lewkowicz considers slope wash (including both surface wash and subsurface wash) as a n important geological agent, particularly because i t contributes directly to both erosion and deposition. All the mechanisms and processes mentioned above are active in the entire periglacial environment. Tney commonly play a more important role in the proglacial subenvironment and will therefore be dealt with in more detail in the section on subaerial mass-transport deposits of the latter subenvironment.

Terminoglacial crevasse deposits (II-A-6-b) Deposits in terminoglacial crevasses are formed partly subaerially and partly under subaqueous conditions. The material accumulates i n the relatively narrow, sometimes deep, vertical or subvertical spaces between terminoglacial ice masses or between the ice front and a n obstacle in the foreland. This situation explains why the crevasses are sometimes filled with water, whereas they are dry at other times. As a consequence, there will be in most situations alternations of subaerial and subaqueous mass flows.

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The terminoglacial subenvironment

Bartkowski (1967) has described frontal crevasses t h a t must have existed during the Pleistocene between a front of active ice, during a stage of readvance, and dead-ice blocks in front of the ice (left after a phase of recession). It appears that terminoglacial crevasses of this type are rare in modern environments, and that the most important depositional site is between the front and a hard-rock barrier (cf. Walczak, 1968, 1970; Klatkowa, 197213;Szponar, 1974,1986).

Lit hofacies characteristics The absence of a characteristic lithofacies is the most characteristic feature: material present 'by chance' in the vicinity flows down into the crevasse by one or another mass-transport mechanism. The consequence is that most deposits have a diamictic character which may be either clasto r matrix-supported, but finely laminated and graded silts may also be found. It is the rule rather than the exception t h a t the deposits a r e

Fig. 217. Terminoglacial crevasse deposits (exposure near U d i , central Poland). There is an irregular mixture of deposits with diverging lithological and textural characteristics. Postdepositional deformation a s a result of collapse of the supporting ice wall has given the deposits a n even more chaotic appearance.

Deposits of the terminoglacial mass-transport facies

351

reworked to some extent after deposition, due to meltwater currents that may follow their way through the crevasse during periods of abundant water supply. The deposits are almost always deformed, as a result of both the depositional process and ice movements which may either press the deposits together or melt away thus removing supporting walls. It is therefore not uncommon that the deposits, deformed already during - for instance - a process of slumping, are deformed once more shortly after deposition. This may result in a chaotic appearance (Fig. 217) that is difficult to unravel.

Textural characteristics As mentioned just above, the deposits are usually extremely badly sorted. The grains show no specific traces of their mass transport, so that they exhibit the same characteristics (including geometry, rounding, etc.) that they had before the mass-transport process took place. Modern crevasse deposits sometimes show a preferred orientation of the larger clasts, but 'fossil' deposits rarely do so. This must be due to postdepositional rearrangement. Imbrication is found in a few cases, but is restricted to the base of the deposit involved. Occurrence Terminoglacial crevasse deposits are typically located in the transitional zone between the glacial and the periglacial environment. They have lateral contact with melting-ice facies, and usually also have such deposits (supra- and subglacial tills) underneath and above. They also interfinger with terminoglacial tills and terminoglacial fluvial deposits. The deposits often have a length several times their width, with a relatively high extension (not rarely over 10 m). Many such deposits may be found in elongated rows, deposited parallel to the (former) ice front, but the isolated position of such deposits may also be due t o partial erosion of larger bodies. The ice side of ice-pushed ridges often contains areas with crevasse deposits, but as locations such as these are open t o erosion after the ice has disappeared, 'fossil' ice-pushed ridges only show traces of such deposits as rare exceptions.

Depositional mechanisms All types of mass transport can contribute to the accumulation of these sediments: debris flow, mudflow, liquified flow, sliding, slumping, gliding,

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The terminoglacial subenvironment

creep and block fall are the most common forms, but other mechanisms (including small-scale turbidity currents) may also play a role. Crevasses may contain temporarily so much water that truly fluvial intercalations are found within the large body of mass-transported sediments. Some settling may also take place in pools. These 'fluvial' and 'lacustrine' deposits must be considered to form p a r t of the normal crevasse infilling.

Terminoglacial subaqueous mass-transport deposits (II-A-6-c) Terminoglacial lacustrine deposits not only contain material t h a t h a s been supplied grain by grain but, as already mentioned, quite often show intercalations of a mass-transport nature. Large slump-like and mudflowlike deposits, rather than turbidites, are the most prominent mass-flow deposits in this context. Various names have been proposed for these typically terminoglacial sediments (also see the review i n Dreimanis, 1989); 'flow till' is the most often used term (among others, by Boulton, 1968). These deposits were also often given names that refer to specific, local conditions: 'lacustrotill' (Odynsky et al., 1952; May, 1977), 'palatill' (Harland et al., 1966), 'aquatill' (Schermerhorn, 1966), 'waterlaid till' (Dreimanis, 1969), 'waterlain till' (Dreimanis, 1976, 1982), 'ablation submarine moraine' (Lavrushin, 1970), 'baseinova moraine' (Ruchina, 1973), 'subaquatic flow till' (Evenson et al., 1977) and 'watermorainic deposit' (Morawski, 1984) are examples of this way of naming. The descriptions of all the above types are more or less identical. Evenson et al. (1977) established a simple environmental model which fits all types. Classical sites with such deposits, e.g., the Scarborough Bluffs (Ontario, Canada), the Fraser River canyon (British Columbia, Canada), Ermelo and Elspeet (Veluwe, The Netherlands) and the Kleszczow graben (central Poland) show a wide variety of such deposits though still with the same general characteristics. These general characteristics are often intermediate between subglacial tills and 'normal' lacustrine deposits (May, 1977). The subaqueous mass-transport deposits in this subenvironment are most frequently seen in lakes, but also occur on subaqueous fans and on deltaic foresets. These deposits have been amply described i n t h e geological literature (Carin and Ahmad, 1961; Boulton, 1968, 1972; Lavrushin, 1970; Evenson et al., 1977; May, 1977, Postma et al., 1983; McCabe et al., 1984, 1987; Broster and Hicock, 1985; Dreimanis, 1989; etc.). It appears that a number of transport mechanisms are involved, and that most can be grouped as slumps, slides, mudflows or debris flows

Deposits of the terminoglacial mass-transport facies

353

(Dzulyrlski and Walton, 1965; Middleton and Hampton, 1973, 1976; Nardin et al., 1979; Postma, 19861, but other processes also play a role. Most of the descriptions and analyses of the mass-transport mechanisms involved are found in publications dealing with other than glacigenic deposits (e.g., Kuenen and Migliorini, 1950; Bagnold, 1954; Dott, 1961,1963; Bouma, 1962; Dzulynski and Walton, 1965; Curry, 1966; R.G. Walker, 1967,1973,1975a71978,1980;Johnson and Hampton, 1969; Van Loon, 1970, 1983; Middleton and Bouma, 1973; Middleton and Hampton, 1976; Nemec et al., 1980; Hiscott, 1981; Harms et al., 1982; Hein, 1982; Gravenor et al., 1984). These reports, while not specialised, a r e most useful for the understanding the processes involved under glacigenic conditions.

Lithofacies characteristics Most authors describing these deposits stress the variability of the lithofacies (Boulton, 1972; Evenson et al., 1977; Hein, 1982; Van der Meer, 1982; Middleton and Hampton, 1983; Rappol, 1983; Morawski, 1984; Brodzikowski and Van Loon, 1987). Mass flows entering a terminoglacial lake directly from the ice mass usually result in matrix-supported, massive or laminated diamicts with abundant fluidisation structures, water-escape structures, load casts and comparable sedimentary deformations. The average grain size of such diamicts may range from fine sand to coarse gravel. Deposits formed as a result of gravity-induced mass transport of till over a subaqueous slope, triggered by a seismic shock, overloading or any other mechanism, are usually slump or slide layers with a number of deformed slabs in a more fine-grained matrix. The slabs may range from a few millimetres to several metres. Both fluxoturbidites and turbidites show grading; they are generally coarse-grained but the former is usually somewhat more coarsegrained than the latter. Flows that do not originate on the ice (or on a till substratum) but on the subaqueous slope of a terminoglacial basin, usually also form diamicts, with massive, matrix-supported types dominating (Fig. 218). Deposits like these may contain a number of larger, floating clasts, which are often concentrated in the basal zone of the unit. The clasts may be hard-rock pebbles, but may also consist of eroded intrabasin material; silt and clay pebbles might be considered diagnostic according to May (1977). There are reports of such clasts occurring in the top part of a mass-transported layer, possibly as a result of movement of the clasts during transport along some kind of shear zones. These masses contain frequent

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The terminoglacial subenvironment

Fig. 218. Terminoglacial subaqueous mass-flow deposits from the Fraser Plateau (British Columbia). There are sandy laminated diamicts with pebble concentrations at the base, and masssive matrix-supported diamicts.

syndepositional deformation structures due to liquefaction, but there are also commonly seen syndepositional or metadepositional flowage structures, load casts with accompanying flame structures, and dish structures (Stauffer, 1967; Zubrzycki, 1972; Middleton and Hampton, 1973, 1976; Dreimanis, 1976; Evenson et al., 1977; Gradzifiski et al., 1986).

Textural characteristics The textural characteristics show as much variability as do the lithofacies characteristics. Descriptions of 'flow tills', for instance refer to material of all grain-size combinations, to all types of surface textures and to several types of fabric. However, most of these deposits are badly sorted, most contain a large proportion of angular fragments, and most show no preferred orientation of the larger clasts. Some orientation (either a slight imbrication or - more commonly - a subhorizontal position) is occasionally found in the basal part of mudflows and debris flows (N. Eyles et al., 1987a; Dreimanis, 1989). This type of orientation is similar to that found in non-glacigenic subaqueous cohesive

Deposits of the terminoglacial mass-transport facies

355

flows (cf. Middleton, 1970; Van Loon, 1970, 1983; Middleton and Hampton, 1973,1976; Carter, 1975; Hampton, 1975,1979; Lowe, 1976a,b, 1979; R.G. Walker, 1978,1984; Postma, l979,1983,1984a,b; Nemec et al., 1980; Porqbski, 1981; Postma et al., 1983). McCabe et al. (1984) found t h a t pebble-sized clasts in vaguely stratified matrix- to clast-supported diamicts can have a pronounced orientation of the ah-axis, subparallel t o the bedding plane (Fig. 219). There are also reports of clasts being orientated along the flow lines during deposition (Lindsay, 1968; Ramsden, 1970; May, 1977). Deposits formed by turbidity currents or any form of grain flow usually show graded bedding. This is normal grading in most cases, but reversed grading can also be found (Postma et al., 1988b). Pebbles present in such deposits are directed according t o the flow lines o r have settled to a (sub)horizontalposition. The lack of uniformity in the textural characteristics of these deposits is due t o the variations in the source-rock material, the divergence of mass-transport processes, the differences in transport distance, etc.

Fig. 219. Sandy laminated diamict with a distinct subhorizontal orientation of the pebbles (British Columbia).Note the.dense foliation, developedduring flowage.

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The terminoglacial subenvironment

Occurrence The deposits usually form irregular but continuous beds with a lensshaped cross-section and a number of lobes (slump-like deposits). These lobes (Fig. 220) may be up t o several metres thick (McCabe et al., 1984; Eyles et al., 1987a; Dreimanis, 1989) and the lobed deposits may be hundreds of metres long if the basin is large enough (Eyles and Miall, 1984; Eyles et al., 1987a). There are also beds that have a sheet-like appearance and very uniform thickness (turbidites, see the subsection on depositional mechanisms). The character of one specific layer changes often from its proximal part towards the distal part (e.g., from slump via mudflow to turbidite). A good model for such transitions, though not established for glacigenic sediments, was presented by Stanley (1981). A number of these deposits consist of material that had previously been deposited as till (the term 'flow till' is understandable i n this context), which started t o flow down a lake slope. Lateral contacts of the terminoglacial subaqueous mass-transport deposits with till are therefore com-

Fig. 220. Detail of the lobe of a slump-like deposit formed in a lake just in front of the Wartanian ice sheet (Glogow, SW Poland). The original lamination of the slumped material has not been destroyed.

Deposits of the terminoglacial mass-transport facies

357

mon. On the other hand, the flows spread out in the lake, so t h a t contacts with glaciolacustrine sediments are also common. An origin on deltaic foresets implies that there may exist direct relationships with terminoglacial deltaic deposits. The fine-grained p a r t s of these deposits i n particular may be again reworked as slurry flows or turbidity currents (Morgenstern, 1967; Hampton, 1972; Carter, 1975). In the glacigenic sequence these deposits may be found throughout the terminoglacial interval, wherever subaqueously formed deposits a r e present.

Depositional mechanisms These deposits represent all types of subaqueous mass-transport mechanisms. The most important ones being continuations of subaerially started mass flows, slumps and slides, cohesive debris flows, cohesionless debris flows, cohesive turbidity currents and 'normal' turbidity currents. The first category (subaerially started flows) is a specific type, because masses may slide or flow into the water with no time t o adapt their transport mechanism to the new situation before they come to rest. In such a case, sediment masses may be found embedded in subaqueous deposits and showing no signs of a subaqueous transport mechanism. It is also possible that the kinetic energy of such masses is so great t h a t transport continues long enough underwater to allow the material to adapt i t s transport mechanism. The final deposit will t h e n show characteristics no different from those of any other deposit formed through a similar subaqueous form of mass transport. 'Normal' turbidites are not very common in terminoglacial lakes, mainly because the size of such lakes is most usually insufficient for mass flows to develop into typical turbidity currents. The mechanism will therefore only be described below in the sections equivalent to these, on proglacial lacustrine bottomsets (11-B-4-d) and proglacial subaqueous mass-transport deposits (11-B-6-4. The distinction between 'normal' turbidites, cohesive turbidites, cohesionless debris flows and cohesive debris flow is based on the rheological characteristics (Postma, 1986) with five main parameters determining the laminar or the turbulent behaviour: velocity, density, cohesive strength, flow thickness and (apparent) fluid viscosity. Cohesive turbidity currents produce, as do 'normal' turbidity currents, a sequence of Bouma intervals that reflect waning flow conditions due to a gradual decrease in turbulence. The flow is turbulent except in the final stage and has a low concentration of particles. The low strength of the

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The terminoglacial subenvironment

current (a high strength value would damp the turbulence and transform the turbulent into a laminar flow: Enos, 1977) allows sand grains t o settle before clay particles. This results in a Bouma succession up to the point where the tangential shear stress has reached the yield strength of a plastic flow and where a massive mixture of very fine sand could be expected to terminate distribution grading (cf. Postma, 1986). Alternatively, flocculation of clay in muddy suspensions may cause a pronounced segregation of clay and silt particles, resulting in distinct laminae (Stow and Bowen, 1980) within the upper part of the Bouma succession (Postma, 1986; cf. Middleton and Hampton, 1973,1976). The term 'cohesionless debris flow' is used to group a number of flows that are known as fluidised flow, liquified flow, grain flow and modified grain flow (which comprises the flow of a traction carpet a t the base of a high-density turbulent flow). The final phase of these types of flows, i.e. deposition, is similar (cf. Bagnold, 1954; Wallis, 1969; Middleton and Hampton, 1973, 1976; Lowe, 1976a,b, 1979, 1982; Rodine and Johnson, 1976; Nardin et al., 1979; Hiscott and Middleton, 1979; Nemec et al., 1980; Postma et al., 1983, 1984; Postma and Roep, 1985). Cohesive debris flows have a laminar character usually following on fully sheared laminar or turbulent flow, although they may succeed flows that were not sheared (slide, slump or 'relict plug' flows). The flows have a high concentration of particles. They have been described from several glacigenic sites (Postma et al., 1983; C.H. Eyles, 1986). According to Lowe (1982) there are three stages of deposition from these high-concentration turbulent flows: a traction stage, a traction-carpet stage and a suspension-deposition stage. The traction stage occurs when the turbulent flow allows some of its load t o settle so as t o form a bed; interactions between flow and bed result in scour-and-fill structures and in several bedforms (Lowe, 197613; Gilbert, 1983; Gravenor et al., 1984; McCabe et al., 1984; C.H. Eyles, 1986; N. Eyles et al., 1987a; Visser et al., 1987). The traction-carpet stage occurs when the flow becomes unsteady; deposition takes place by vertical accretion of relatively thin graded layers that form a t the base of the still turbulent flow. The suspension-deposition stage occurs if the flow is separated, as a result of gravity, into a lower part with a high concentration of suspended particles and a more dilute and finer-grained upper part. Middleton (1967) demonstrated this process experimentally. The feature w a s also discussed by Postma et al. (1988) for non-glacigenic flows and by Broster et al. (1980), Postma et al. (1983) and C.H. Eyles (1986) for glacigenic flows. Deposition is more rapid if the cohesionless flow is gravelly than in the case of a sandy flow; this is due to the greater permeability and weight of the clasts.

Deposits of the terminoglacial mass-transport facies

359

Cohesive debris flows may originate from other flows that are either completely or not fully sheared. The flow character is laminar and the concentration of suspended particles is high. The maximum clast size in the deposits of such flows is a measure of the flow's maximum competence during flow (Hampton, 1975). Deposition takes place after shear has dropped below the yield strength. However, the velocity gradient is much higher a t the top than a t the base of the flow, so that deposition will start at the top and a rigid plug will thus be formed. Continuous shearing a t the base is promoted by a n increase in pore pressure due to the formation of the rigid plug and by squeezing of pore water from deposits underneath (Postma, 1986). The final deposit is thus characterised by a thin shear zone (which may locally be inversely graded) in which the lamination may be difficult to observe if there is little differentiation in the grain size of the particles involved; the grains in this zone are usually smaller than those embedded in the matrix-rich plug. Deposits t h a t have such characteristics have frequently been described from glacigenic settings (Broster et al., 1980a,b,c; C.H. Eyles et al., 1983a,b; C.H. Eyles and N. Eyles, 1984a, 1985; McCabe et al., 1984, 1987; Broster and Hicock, 1985; N. Eyles, 1987; N. Eyles et al., 1987a). If the cohesive flow has not been fully sheared, the deposits show usually ill-sorted debris in a muddy matrix. The sorting is extremely bad if large boulders are present. McCabe et al. (1974, 19771, Rodine and Johnson (1976) and Eyles et al. (1987a) have explained that such boulders may be transported by the flow if they are embedded in the rigid plug. The weight of the boulder is then supported by the contents of the plug and not by the flow, whereas the plug itself is supported by the mobile zone around it (see also the explanation regarding graded pebbly mudstones in Van Loon, 1970). If the state of the plug changes from rigid t o ductile, the boulder may become too heavy for the yield strength of the plug, so that the boulder sinks until it reaches the bottom. It can then be dragged along the sedimentary surface for some time as a result of the inertia stresses imposed by the fluid and the suspended grains. Such cohesive debris-flow deposits are usually characterised by diffuse lamination and an inversely graded basal shear zone. Several examples have been described by, among others, Evenson et al. (1977), Broster et al. (1980a,b,c),Broster and Hicock (1985), N. Eyles (1987), Eyles et al. (1987) and Dreimanis (1989).

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The proglacial subenvironment

36 1

THE CONTINENTAL PROGLACIAL SUBENVIRONMENT (11-B)AND ITS DEPOSITS There may exist a large zone in front of the terminoglacial subenvironment where dead-ice is present in the subsoil, left behind when the ice retreated (see also the Epilogue). The width of this zone may vary from nil to several dozens or even hundreds of kilometres. There is no cover of glacial ice whatsoever in this proglacial subenvironment (Fig. 22l), but other types of ice, e.g.,lake or river ice, may occur (Embleton and King, 1975; J a h n , 1975; French, 1976; Drewry, 1986). The subenvironment (see also the Table on p. 128) is characterised by relatively small (in comparison to the various glacial subenvironments and particularly to the terminoglacial subenvironment) and gradual changes i n the physical, chemical and hydrological conditions. The dynamics of the depositional processes (see also Fraser and Cobb, 1982) are largely determined by the amount of meltwater supplied, by fluctuations in this supply and the maximum energy levels reached, by the amount of debris supplied and the average and maximum sizes of the

subaerial slides

proglacial stream deposits (11-B-2-d)

Fig. 221. The proglacial subenvironment with some of the most common features.

362

The proglacial subenvironment

clasts, and by the topography and pattern of the depositional basins. This implies that the proglacial conditions have much in common with those in non-glacigenic areas, a fact which is reflected in the types of depositional processes and the resulting sediments. Several authors (e.g., Flint, 1971; Boulton, 197213;Klimaszewski, 1976; Sugden and John, 1976; Reineck and Singh, 1980; Edwards, 1978; Drewry, 1986) have stressed that the field conditions in this subenvironment allow much more detailed investigations than do the conditions in the terminoglacial subenvironment and the glacial environment. It appears that the general inclination of the proglacial area (towards the ice mass or in the opposite direction) also plays a major role, particularly because of the effects on the drainage system. A general inclination of the area towards the ice front restricts the length of the streams carrying glacial debris away from the ice; moreover, the continuous water supply from meltwater streams and the limited possibilities for water to find a way out of the area result in the frequent occurrence of lakes, swamps, etc. This type of situation may occur if, for instance, a mountain range is situated more or less parallel to the ice front (as was proglacial subenvironment (11-B) DrDlllaClai ,an ,ac,es

A

proglacial subenvironment (11-B)

Fig. 222. Schematic models of the proglacial subenvironment with the various proglacial facies and some deposits.

Sedimentation processes in the proglacial subenvironment

363

the case during the Pleistocene glaciations in central Europe: Carpathians, Sudetes, etc.; cf. Woldstedt, 1954; Bartkowski, 1967). In such a case the superfluous water may flow away via large valleys with marginal (extramarginal) rivers (Fig. 221), often called 'pradolinas' (Kozarski, 1962; Galon, 1972; Krygowski, 1972,1974; Stankowski, 198lb). If the area in front of the ice dips in a 'down-ice' direction, the proglacial area tends to be rather wide: the general inclination facilitates far transport of glacial debris. Gradual slowing down of the meltwater courses and mixing with non-glacial streams result in a generally vague transition into the extraglacial subenvironment. This occurred in eastern Europe (particularly in the European part of the USSR) and in northern America during the Pleistocene. Similar conditions have been reconstructed for the Permo-Carboniferous Patagonian glaciation (Frakes et al., 1969; Adie, 1975), for the Permo-Carboniferous Australian glaciation (Bowen, 1969; Stratten, 1971) and for the Ordovician glaciation in the Sahara (Beuf et al., 1971; P. Allen, 1975). Six depositional facies (Fig. 222) can be discerned: a proglacial meltingice facies (11-B-11,a proglacial fluvial facies (II-B-2), a proglacial deltaic facies (II-B-3), a proglacial lacustrine facies (II-B-4), a proglacial aeolian facies (11-B-5)and a proglacial mass-transport facies (II-B-6).

SEDIMENTATION PROCESSES IN THE PROGLACIAL SUBENVIRONMENT The proglacial subenvironment is very diverse. Several types of sedimentation processes can take place, but most deposition results from glaciofluvial, glaciodeltaic, glaciolacustrine, aeolian and mass-flow processes. The dynamics of sedimentation are largely determined by the supply of meltwater, the amount of debris carried, and the local relief. The proglacial subenvironment extends as far as buried dead-ice blocks, being remnants of a period of greater ice extent, are present. Meltwater streams are common in this zone. The combination of dead-ice blocks and meltwater streams largely determines the depositional pattern. A barrier for the meltwater streams, such as a mountain range, may also function as a distinct boundary of the proglacial subenvironment, i.e. if the ice cap was also unable to pass this barrier. If there is no clear feature of this kind, a boundary between the proglacial and extraglacial subenvironments is much more difficult to discern, especially in the case of 'fossil' deposits, but the presence of buried dead-ice should be considered as the main criterion for distinguishing between these subenvironments.

364

The proglacial subenvironment

This subenvironment may extend so far that meteorological conditions other than ice-dependent ones also influence the sedimentary pattern. Precipitation and the dynamics of the atmosphere play a major role. All these characteristics indicate t h a t the proglacial subenvironment resembles the non-glacial (sub)environments that exist elsewhere much more than do the subenvironments mentioned above. The same is true of the energy gradient, which does not differ significantly from non-glacial depositional environments at middle latitudes. The floods that are responsible for considerable glaciofluvial sedimentation in the terminoglacial subenvironment also occur in the proglacial subenvironment because vast amounts of water can be trapped temporarily in the terminoglacial lakes that are released suddenly if some barrier fails. The proglacial floods are generally much weaker, however, than their terminoglacial counterparts. I t is therefore understandable that truly coarse deposits are much less common in the proglacial subenvironment than further towards the ice front.

Fig. 223. Thick, irregular sediment cover on buried dead-ice blocks in the proglacial subenvironment in front of the Werenskiold glacier (Svalbard). Photograph: J. Czerwifiski.

Deposits of the proglacial melting-ice facies

365

Several workers in the field have described characteristics of the sedimentation processes and the resulting deposits. Most of these authors stress the importance of sedimentation processes on top of buried dead-ice and thus the presence of melting-ice deposits (Fig. 223) (among others, Czeppe, 1965; Klimek, 1972; Churski, 1973; Czerwiiiski, 1973; Kozarski and Szupryczyiiski, 1978; Shaw and Archer, 1979; K b s z and Lindner, 1981, 1982; Lindner et al., 1982), which implies that deposits deformed through collapse processes are not uncommon in this subenvironment. DEPOSITS OF THE PROGLACIAL MELTING-ICE FACIES (II-B-1) As the proglacial subenvironment is defined on the basis of the presence of buried dead-ice blocks, the deposits of the melting-ice facies are diagnostic, even though such deposits constitute only a minor part of all deposits in this subenvironment. The melting-ice deposits are found in the form of lenses, usually embedded in glaciofluvial material. Only one type of deposit can be distinguished here: the proglacial till complexes (II-B-1-a). Proglacial till complexes (II-B-1-a) These complexes are formed by the gradual melting of buried dead-ice blocks, left after the retreat of the ice cap. The deposits are formed at the base, alongside and on top of the ice mass, which gives the resulting sediment a mixed appearance. In addition, the melting of the ice creates open spaces, where part of the debris may be reposited in the form of flow tills. These considerations make i t justified to group all deposits together as till complexes. Part of the fines may be washed out by percolating waters, so t h a t the resulting deposits tend to be somewhat coarser than might be expected on the basis of the size of the debris originally present in the dead-ice blocks.

Lithofacies characteristics Many complexes are massive, matrix-supported diamicts with characteristics that are identical to those of till complexes formed by melting of passive ice in the terminoglacial subenvironment. Lamination may be present but is usually not very well developed; fluidisation structures, water-escape structures and shear planes are features that may hide the original structure.

366

The proglacial subenvironment

The complexes show a tendency t o become more massive towards the top, but the topmost part itself may have a different appearance because of thge frequent occurrence of reworked material (flow tills).

Textural characteristics The proglacial till complexes have no specific textural characteristics. Sand and finer material usually dominate, but pebbles and boulders may be present, usually floating in the finer-grained matrix. The clasts show no diagnostic features. The rounding is usually moderate, but very angular clasts may occur, sometimes together with well rounded clasts. Striations on the clasts are found frequently.

Occurrence Dead-ice blocks are only present in areas that were covered by the ice cap in a previous stage of ice advance. This implies that the complexes may be found in the zone from the line of maximum ice extent up to the transition into the terminoglacial subenvironment (where similar deposits can be found). Many dead-ice masses and resulting till complexes are incised by proglacial streams. There are thus frequent lateral contacts with fluvial deposits, but also with mass-transport deposits. The complexes are often also covered by proglacial fluvial and proglacial subaerial mass-transport deposits, but they rest usually on top of tills. The complexes are never found as widespread bodies, but always as isolated lenses.

Depositional mechanism The depositional mechanism is not different from that of terminoglacial counterparts. Much material is set free by melting of the ice, and comes to rest immediately a t the final depositional site. Other particles, however, may be replaced by mass-transport processes that take place in the spaces left by the melted ice. The open spaces also give the opportunity for meltwater streams t o flow through and around these deposits, partly eroding them and taking away relatively much fine-grained material.

Deposits of the proglacial fluvial facies

367

DEPOSITS OF THE PROGLACIAL FLUVIAL FACIES (II-B-2) The meltwater streams originating i n the glacial environment and passing through the terminoglacial subenvironment form several types of river deposits in the proglacial subenvironment. Net deposition takes place when channels gradually become filled, when overbank flow takes place, and by coalescence of fan-shaped bodies (Drewry, 1986). Fluvial conditions in the proglacial zone are well comparable t o those in nonglacigenic areas where pronounced changes in flow regime may occur. The changes in flow7 regime in the proglacial subenvironment are largely determined by seasonal melting of the nearby ice mass. Unstable conditions with rapidly evolving channels (Fig. 224) are therefore characteristic. Three subfacies can be distinguished, which gradually pass into one another. The most proximal subfacies is formed by glaciofluvial fans (which may, in turn, be subdivided into upper, middle and lower fan). The lower parts of these fans, if the conditions are favourable, pass into glaciofluvial braid plains. These braid plains may be very large, as is known from modern examples such as the Skeidararsandur (Krigstrom, 1962; Klimek, 1972; Churski, 1973) and from ancient - mainly Pleistocene examples (Kozarski, 1962, 1965; Bartkowski, 1967; Augustinus and

Fig. 224. Fluvial plain with shallow, rapidly shifting channels. Issortog area, western Greenland. Photograph: J. Cegfa.

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The proglacial subenvironment

Riezebos, 1971; Galon, 1972; Krygowski, 1974; Ruegg, 1977; J.R.L. Allen, 1982b; Landvik and Mangerud, 1985). The sediments of the fans and the braid plains in the proglacial subenvironment frequently build up landforms known as sanders (Fig. 225). Horizontal grain-size differentiation is common in these bodies, the coarsest units indicating the proximal facies. The thickness of t h e separate sets also decreases from the proximal towards the distal part. Such sanders are also described as outwash plains (e.g., Hjulstrom, 1954; Krigstrom, 1962; Augustinus and Riezebos, 1971; Church, 1972; Klimek, 1972; Churski, 1973; Bluck, 1974; Eynon and Walker, 1974; Gustavson, 1974; Boothroyd and Ashley, 1975; Clague, 1975; Ward et al., 1976; Hine and Boothroyd, 1978; Wisniewski and Karczewski, 1978; Haraldsson, 1981; J.R.L. Allen, 1982b; Casshyap and Tewari, 1982; Landvik and Mangerud, 1985). The braid plains, both those forming part of sanders and others, gradually pass into ice-marginal streamways (often of considerable length: see Fig. 6) which are commonly termed 'pradolinas' (Woldstedt, 1954; Charlesworth, 1957; Kozarski, 1962,1965; Galon, 1972; Krygowski, 1972,1974; Jahn, 1975; Klimaszewski, 1976; Edwards, 1978; Stankowski, 1981b), or into streams that are directed away from the ice. There is not always a braid plain between the fan a r e a and t h e pradolina area. The latter may also form the natural continuation of the fan area. In addition, all fluvial subfacies may change gradually into lacustrine facies, for instance by local blockage of a stream, a local depression or coalescence of a number of streams. On the other hand, lakes in the proglacial subenvironment may store large amounts of water, so that the outflow is generally more regular than the inflow (Kozarski, 1962, 1965). It is therefore not uncommon that streams enter the proglacial subenvironments as braided streams but leave it as meandering rivers. Low gradients in the river profile are characteristic of t h i s subenvironment (Drewry, 1986). In the case of proglacial streams in Alaska, Boothroyd (1986) mentioned gradients of 0.006-0.017 i n the proximal zone, decreasing t o 0.002-0.003 in the distal part. Individual channels are usually small, but may here and there reach widths of several hundred meters, with corresponding sizes of bars and islands around which the streams bifurcate (Drewry, 1986). The depth of the stream rarely exceeds a few metres. The low gradients imply that coarse material is generally left behind in the proximal parts of the fans (Fig. 226), although the downstream decrease in average grain size of course depends on specific characteristics

Deposits of the proglacial fluvial facies

0 =Coarse Gravel 0Fine Gravel DSand

369

1Okm

~ T I I I =Glacier

Ice

EEl Bedrock

Fig. 225. Some sanders, with their average grain-size distribution. The Skeidararsandur is in Iceland, the Malaspina and Scott systems are in Alaska. Slightly modified from: B2othroyd and Nummedal(1978).

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The proglacial subenvironment

Fig. 226. Upper fan in front of the Werenskiold glacier (Svalbard). Note the large clasts and the still larger blocks of ice, which could be transported further than stones of the same size because of their lower specific weight. Photograph: J. Cegia.

of the stream involved, such as current velocity, deptwwidth ratio, overall transport distance, turbulence, etc. (Fahnestock, 1963; Rice, 1982). The sanders and the pradolinas contain material that is usually much finer grained (Lied, 1953; Bartkowski, 1957,1967; Liedtke, 1957, 1961; Galon, 1961, 1968a,b; Kozarski, 1962, 1965, 1966; Lembke and Marcinek, 1965; Jewtuchowicz, 1967; Karczewski, 1968; Roszko, 1968; Marcinek, 1969): the gravelhand ratio, even in large pradolinas, decreases distinctly in the downstream direction to values much lower than 1/20. The horizontal differentiation into fan, plain and pradolina subfacies is consistent with Miall's (1978) vertical profile models for braided-stream deposits (Fig. 227). Coarse-grained systems form horizontal associations comparable t o the transition from the Scott type, via the Donjek type, t o the South Saskatchewan type (Bluck, 1974; Boothroyd and Ashley, 1975). The fine-grained systems form associations of the Donjek type (or even the Saskatchewan type) in the the proglacial fan, passing into the South Saskatchewan type or the Platte type in the area of the outwash plain. Associations of the Platte type or (for overbank deposits) even the Bijou Creek type are found downstream, in the pradolina area (Bartkowski,

Deposits of the proglacial fluvial facies

371

1957; Liedtke, 1961; Kozarski, 1962,1965; Jewtuchowicz, 1967).The same characterististics have been found for 'fossil' and modern situations (Hjulstrom, 1954; Krigstrom, 1962; Klimek, 1972; Hine and Boothroyd, 1978; Kozarski and Szupryczyfiski, 1978). Five processes are considered responsible for most of the sedimentary structures and for the grain-size distribution in deposits formed by braided rivers. These processes are: formation of longitudinal bars, bedform generation and migration, channel scour and fill, low-water accretion, and overbank sedimentation (Miall, 1977, 1978, 1984; Drewry, 1986). Collinson (1978) mentioned that most gravel deposits show imbrication, that most successions show a n upward fining tendency, and that most units show horizontal bedding, although there may also be crossstratification. Boothroyd and Ashley (1975) and Rust (1978) emphasised the common occurrence of massive gravelly units and clast-supported gravelly diamicts. Sandy units comprise tabular cross-bedded sets with occasional ripple-drift cross-lamination (Boothroyd and Ashley, 1975; Boothroyd, 1976; Miall, 1977, 1984; Walker, 1984). The infillings of channels often show eroded clays and silts, overlain by gravels (Picard and High, 1973). Draping of irregular bedforms is common, as a result of temporarily waning flows. The sedimentation pattern of the - mainly braided - streams in the proglacial subenvironment is rather complex, with numerous intercalations of bar and channel-fill deposits. Reconstruction of the palaeogeography

I

'

Fig. 227. Vertical profile models for braided-stream deposits. Arrows show small-scale cyclic sequences. Conglomerate clasts not to scale. From: Miall (1978).Courtesy: Canadian Society of Petroleum Geologists.

372

The proglacial subenvironment

therefore requires a detailed analysis of the nature of the bar and channelfill complexes (Collinson, 1970; Castello and Walker, 1972; Bluck, 1974, 1975/1976,1980; Eynon and Walker, 1974; Miall, 1977; Walker, 1984). Bars form a main constituent of proglacial fluvial deposits. Several types (e.g., longitudinal, side, lateral, diagonal, point and chute bars) have been described, but Miall (1977, 1978) has stressed that only those types should be distinguished by sedimentologists t h a t result in distinctly different lithostructural units. An additional problem is that bars are a geomorphological form, and that the terminology regarding these bodies is a very chaotic. Drewry (1986) stated with respect to this problem "... To add confusion the term bar is not necessarily ideal! For features which have spacings of 5-100 m, small heighthpacing ratios, and have characteristic flow velocities in the order of > 0.3-0.4 and < 0.7-0.8 m s-l (Harms et al., 1975) have proposed the term sandwave. Such bedforms lie between the dynamic ranges (defined by channel depth, water velocity a n d sediment size) of ripples and dunes Sedimentologiats, however, define sandwaves in a different way (Allen, 1968; Coleman, 1969; Collinson, 1970; Reading, 1978).Nevertheless, there is reasonable agreement t h a t it is useful from a sedimentological point of view to distinguish between three types of bars in the proglacial subenvironment: longitudinal, transverse and lateral bars. These types will be dealt with in more detail in the section on proglacial stream deposits (II-B-2-d). ...'I.

Fig. 228. Three-dimensional model of the principal features in the Donjek area. Numbers 1-4 refer to the respective stream channels and (or) levels. Slightly modified from: Williams and Rust (1969).

Deposits of the proglacial fluvial facies

373

Braided rivers in the proglacial subenvironment are often found i n areas with several distinctly different topographic levels that represent successive stages of downcutting (Fig. 228). Relief-rich areas where erosion is still active provide the best examples (Williams and Rust, 1969; Kessler and Cooper, 1970). Modern examples a r e the Scott and Yana glacial outwash rivers in southern Alaska (Boothroyd and Ashley, 19751, the South Saskatchewan river (Cant, 1975) and some sander rivers described by Church and Gilbert (1975). Pleistocene pradolinas with such levels have also be identified (Liedtke, 1957, 1961, 1981; Galon, 1961, 1968a,b; Kozarski, 1962,1965). There are many further important aspects of proglacial fluvial sedimentation, but their discussion is beyond the scope of this book. The most relevant aspects will be discussed in the relevant sections on specific types of deposits. Three general types of deposits are distinguished in this facies: proglacial fluvial complexes (II-B-2-a),proglacial stream deposits (II-B-2d) and proglacial sheet- and streamflood deposits (II-B-2-e).

Proglacial fluvial complexes (II-B-%a) Fluvial complexes are much rarer in this subenvironment than in the supraglacial and terminoglacial subenvironments. This is due primarily to the longer transport distances and results in much better differentiation of the material. However, as local conditions may result in narrow proglacial belts, fluvial complexes may appear, as exceptions (Jahn, 1956; Bartkowski, 1967; Klatkowa, 1972a,b; Szponar, 1974, 1976; Eissmann, 1975; Drozdowski, 1979b; Ehlers, 1981, 1983c; Klysz, 1981, 1986; Kozarski, 1981; Kozarski and Kasprzak, 1987). All palaeogeographic reconstructions show that specific conditions are required for proglacial fluvial complexes to be formed. It appears that such complexes are always associated with early stages of ice advance, when the proglacial subenvironment had not yet been levelled by periglacial (aeolian in particular) deposits. The fluvial complexes that result locally are thus always of limited extent. Lithofacies characteristics The complexes show a wide variety of lithologies. While sandy and gravelly strata predominate (Fig. 229), there may be numerous diamictic intercalations. Most of the gravelly units are massive and show no distinct sedimentary structures. The more sandy complexes (or strata within a complex) usually show horizontal lamination or cross-stratification and

374

The proglacial subenvironment

Fig. 229. Irregular fluvial deposits, to be considered as part of a proglacial fluvial complex. Surroundings of Suwaiki (NE Poland). Photograph: T. Zielifiski.

small-scale current ripples. Tabular sets are common but trough crossbedding and horizontal lamination are usually rare. The lithology changes rapidly in both a vertical section and lateral directions. The erosional phases inherent t o the formation of these bodies produce a generally complicated alternation of units with erosional surfaces with a slight upward tendency t o more regular sedimentation. Textural characteristics

Most of the larger clasts show a distinct, preferred orientation. This is usually with the a-axis perpendicular to the flow direction, although an orientation parallel t o the flow direction has also been observed (Rust, 1972,1975,1978; Boothroyd and Ashley, 1975). Imbrication (with a dip in upstream direction) is a very common phenomenon in both cases. The larger clasts usually show signs of significant fluvial transport in the form of subrounded t o rounded shapes with several angular specimens still present.

Deposits of the proglacial fluvial facies

375

The orientation of current ripples and other sedimentary structures (such as gravel stripes) is higly variable. It is likely that the orientation of these structures is determined to some degree by the discharge (Bluck, 1974). Occurrence Proglacial fluvial complexes are usually found in the lower part of glacigenic sequences, where they fill depressions in the substratum. The complexes are then overlain by terminoglacial deposits or by till - if the ice has eroded the terminoglacial sediments. The complexes may also be formed if repeated retreat and re-advance of the ice has resulted in a relief-rich foreland; this is reflected in the occasional presence of complexes in the middle of the glacigenic sequence. The complexes are much rarer in the upper part of the sequence, as ice retreat is seldom accompanied by depositional conditions that favour the formation of these complexes. The basal contacts of the complexes are sharp, almost without exception, and also erosional in most cases. In the lateral direction, the complexes interfinger with proglacial subaerial mass-transport deposits, proglacial deltaic, lacustrine and stream deposits and reasonably often with proglacial aeolian deposits. Lateral contact with deposits from other subenvironments are rare, except in the case of tills formed after an erosional phase of ice advance. The top part of the complexes generally passes gradually into other deposits, which most usually are also part of the proglacial fluvial facies. Depositional mechanisms The depositional mechanisms are no different from those operating under non-glacigenic fluvial conditions, where irregular and surge-like supply of water and debris predominates. Most of the coarse-grained deposits are formed during periods of high floods, when massive gravels with sandy admixtures are formed. These units may show rough, slightly inclined stratification. Units with trough-shaped stratification can be formed during all stages of a flood, but it seems that only units formed during the final stages have a good preservational potential. It is typical that rippled layers are formed during periods of relatively low discharge. The frequent alternations of units with the various characteristics described above indicate that the flow conditions change rapidly.

376

The proglacial subenvironment

Proglacial stream deposits (II-B-2-d) These deposits may be not those most characteristic of glacigenic conditions, but they occur very frequently and show a wide spectrum of properties. They are therefore interesting from a sedimentological point of view. The streams in the proglacial subenvironment are mainly braided but meandering streams occur as well. It appears that the deposits from braided streams have a better preservational potential in this subenvironment, probably because meandering rivers erode large parts of the previously deposited material, whereas braided streams more commonly result in relatively rapid vertical accretion. Consequently, one frequently finds channel-bar deposits and interbar channel deposits (J.R.L. Allen, 1972, 1982a; Reading, 1978; Reineck and Singh, 1980; Hey et al., 1982; Miall, 1984). Three main types of bars are generally distinguished: longitudinal, transverse and lateral bars; the last type comprises pointbars, side bars and alternate bars (Sundborg, 1956; Collinson, 1970; N.D. Smith, 1971a,b; Rust, 1972; Bluck, 1974, 1980, 1982; Shelton and Noble, 1974; Asquith and Cramer, 1975; Boothroyd and Ashley, 1975; McGowen and Garner, 1975; Cant, 1975, 1976; Cant and Walker, 1976; Miall, 1977; Vos and Tankard, 1981; Southard et al., 1984; Drewry, 1986; R.M.H. Smith, 1987).

Fig. 230. Gravelly and sandy units of a longitudinal bar formed in a Pleistocene proglacial stream. British Columbia.

Deposits of the proglacial fluvial facies

377

The longitudinal bars (also termed 'diamond bars') found in this subenvironment may be several hundred metres long, with a vertical relief of somewhat more than one metre (smaller specimens are much more common). These bars, which are mostly built of sand and gravel (Fig. 230), often show tangential to planar cross-lamination with frequent reactivation surfaces. They are usually formed in the middle of a channel, have a current-parallel shape, and migrate downstream by the development of slip faces along the bar edges. Migration is more frequent and more rapid in sandy than in gravelly bars. The top of the bar is often covered by ripples and transverse ribs, usually consisting of relatively coarse material (Boothroyd and Ashley, 1975; Miall, 1977; Reading, 1978; Walker, 1984). Rust (1978a) came to the conclusion that such forms are stable even during floods. So-called 'diagonal bars' must be considered as a variety of longitudinal bars; they developed in an oblique direction during the last part of a flood stage. The diagonal bars could also be considered as intermediate forms between longitudional and transverse bars. Transverse bars (Fig. 231), also known as 'cross-channel bars', have

Fig. 231. Part of a well developed transverse bar formed in the mid-fan part of a Pleistocene proglacial stream near Suwatki (NE Poland).Several reactivation surfaces are visible. Photograph: T. Zieliriski.

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The proglacial subenvironment

cusp- or spoon-like (linguoid) shapes. They can be several hundred metres long, with a height of up to a metre. They occur in the pebble-rich parts of low-sinuosity rivers (Bluck, 1974, 1980, 1982) and - more commonly - in the networks of sandy alluvial channels (Cant, 1975, 1976). Walker and Cant (1979) suggested that these bodies have a core that emerges from the water during low-water stages, and that they gradually grow by downstream extension of lateral 'horns'; there is also upstream expansion, although slower, by accretion of migrating sand dunes. Bars with a significant propcrtion of sand show laterally progressing slip-face migration. This process leads t o the development of tangential to planar crossbedding and pronounced ripple-drift cross-lamination (Doeglas, 1962; Fahnestock, 1969; N.D. Smith, 1971a; Picard and High, 1973; Bluck, 1974, 1980; Gustavson, 1974; Walker and Cant, 1984). The transverse bars occur in groups with an out-of-phase relationship between individual bars (Collinson, 1970; Miall, 1977; Vos and Tankard, 1981) so that the bars tend t o fill spaces between two preceding bars (Drewry, 1986). Lateral bars are formed in low-energy zones, typically in the distal braid plains and the pradolinas. They may be very small, but Collinson (1970) has described an example six kilometres long. The internal structure of lateral bars is usually complex, due t o the inherited structures of coalesced smaller units. Lateral bars may grow locally in a downstream direction, but the lateral accretion (side bars, pointbars) is the most important and is due t o a diagonal flow that results in accumulation along the margins of the bar. A specific type of lateral bar, although not a bar in the strictest sense, is the pointbar (also termed 'meander bar'), a characteristic feature of high-sinuosity meandering streams (McGowen and Garner, 1970,1975; N.D. Smith, 1974; Bluck, 1975/1976; Miall, 1977; Rust, 1978a; D.G. Smith, 1987; R.M.H. Smith, 1987). Well developed pointbars (Fig. 232) are relatively rare in the proglacial subenvironment. In addition to the bars, deposition takes place in interbar channels (Cant, 1976; Cant and Walker, 1976; Miall, 1977; Reading, 1978). The channels are continuously filled, or at least partly so, with meltwater, even during low-water phases. According t o Miall (19771, the morphology of the bottoms and slopes of braided channels show scour structures, dunes, etc. Small bar-like bodies may be found in these channels (Culbertson and Scott, 1970). Channel deposits usually have an erosional base and are characterised by a basal lag deposit of gravel (Miall, 1970; N.D. Smith, 1970; Cant and Walker, 1976). Cant (1976) suggested that the preservational potential of interbar channel deposits is relatively high because the large lower parts are protected against erosion by the smaller, overlying bedforms related to them.

Deposits of the proglacial fluvial facies

379

Fig. 232. Typical pointbar sequence with chute bars cutting lateral accretion deposits (lower unit with inclined stratification), chute channels and infills (lower dark sediment), channel lag (coarse, middle part) and channel infill (upper part). Proglacial deposits from the Kleszcz6w graben (central Poland). Photograph:A. Hahszczak.

Lithofacies characteristics The processes involved in fluvial sedimentation are many and varied, as already mentioned. The resulting deposits show an equally wide variety of lithologies. The precise distribution of the various lithofacies types depends on several factors, e.g., flow characteristics, energy level, distance t o the source (granulometry), position with respect to average water level, etc. This implies that there are generally differences between the predominant lithologies in the fans, those in the outwash plains and in the pradolinas. Boothroyd and Ashley (1975) have described an excellent model for Alaskan sanders (Fig. 233). Under modern conditions, the cyclicity in meltwater and sediment supply results in repeated sequences (Rust, 1978b) that are, however, found only occasionally in 'fossil' successions. The cycles are different for fine-grained and coarse-grained systems (Walker and Cant, 1984).

380

The proglacial subenvironment

Fig. 233. Downstream variation in facies and sedimentary structures in a proglacial fan. Slightly modified from Boothroyd and Ashley (1975).

The proximal stream deposits usually have a high percentage of gravel. The almost pure gravels rarely show internal structures, but may have an inclined tabular stratification if they contain sand admixtures. Units made of sand and not containing gravel are very rare in the proximal part of this subenvironment. A more granulometric variation starts in the midfan zone. Upward-fining successions in bars are already common in this zone and the upper, sandy parts commonly contain abundant current ripples. The interbar channel deposits in the mid-fan zone are generally somewhat coarser, with frequent trough cross-bedded gravelly or sandygravelly sediments. The lower fan, the outwash plain and the icemarginal streamways are finer-grained still, with abundant signs of the action of water currents, such as current ripples and horizontal lamination. Truly muddy facies are scarce, being occasionally found in depressions on top of bars, and as pointbar deposits in sequences of meandering streams.

Deposits of the proglacial fluvial facies

381

Textural characteristics A number of field studies have yielded short descriptions of the textural characteristics of these deposits (Williams and Rust, 1969; Church, 1972; Bluck, 1974,1982; Church and Gilbert, 1975; Rust, 1975). A wide range of sorting degrees and a general decrease of average grain size in a downstream direction are most frequently mentioned as the prominent features. The degree of sorting appears to be inversely proportional to the average grain size. Grains and clasts usually show distinct signs of attrition. The ah-planes of the larger clasts are commonly inclined. According to Laming (1966), the dip of this plane is often as much as 40" in ancient clast-supported deposits and less than 25" in matrix-supported sediments. Cailleux (1945) found that the inclination increased with decreasing flatness and decreasing grain size. However, relatively strongly inclined clasts occur (Johansson, 19651, particularly in coarse deposits and a strong inclination is also an indication of relatively deep water. In addition, a

Fig. 234. Vertical section through a proglacial longitudinal bar about 1 m thick, developed in the upper-fan zone. Note the orientation of the &-planes of the clasts parallel to the inclined stratification (British Columbia, Canada).

382

The proglacial subenvironment

strong inclination would be favoured by rounded (though less prolate) shapes of the pebbles (cf. White, 1952). Rust (1975) concluded from a detailed study that there is no clear relationship between pebble size and inclination. He also found that the dip is independent of the height of the pebbles within the deposit. Folk and Ward (1957)suggested that the fabric often is not a result of depositional processes but is an early-diagenetic phenomenon, due mainly to the removal of sand from the spaces between the individual pebbles. Rust (1975) considers this explanation as unlikely. Wadell (1936) and Johansson (1963,1965) showed that pebbles deposited on the inclined foresets of large ripples tend t o dip in a downstream direction but with an inclination that is, on an average, 15" less than the dip of the foreset itself. However, this difference in inclination need not exist (Fig. 234). Much more common is upstream imbrication of the clasts, particularly in clast-supported sediments. Williams and Rust (1969) had summarised the textural parameters in channel fills on the basis of detailed textural studies. They mentioned the following general trends: a consistent upward decrease in mean grain size and coarsest percentile, a consistent upward decrease in positive skewness, minor variations in sorting (which is usually poor to moderate), and insignificant variations in kurtosis.

Occurrence Proglacial stream deposits are found mainly in the lower part of the glacigenic sequence, but may also be present in the upper part, where they can overlie supraglacial and terminoglacial sediments. The stream deposits may reach considerable thicknesses (up t o tens of metres) and show frequent intercalations with proglacial sheet- and streamflood deposits. Lateral contacts with proglacial lacustrine, deltaic, mass-transport and aeolian deposits are also frequent. The fact that several subtypes of stream deposits can be found was mentioned earlier. They occur in more or less fixed zones rather than being randomly distributed. The streams in the upper-fan area have a braided character, due to the formation of longitudinal bars in the channels. These bars usually have a core of massive gravel and a gravellyhandy margin that migrates in a downstream direction, thus forming inclined, weakly stratified units (Boothroyd and Ashley, 1975; Bluck, 1980). These bars are best developed, and have their classical shapes, in the mid-fan area. Downcurrent, in the lower-fan area, the longitudinal bars tend t o be modified into side or lateral bars, which seem t o be the most frequent types in the proglacial subenvironment.

Deposits of the proglacial fluvial facies

383

Transverse bars are usually formed in the mid-fan area, but migrate downstream and reach their best development in the lower-fan area and in the braid plains. The core of these bars may, but does not necessarily, consist of a remnant of a longitudinal bar. These bars migrate downstream mostly by accretion of linguoid fronts. Formation of the lateral bars also starts in the mid-fan area, but they are best developed in the braid plains and pradolinas. Pointbars are formed, though rarely, along relatively small and sinuous channels within the lower-fan areas and the braid plains. Small-scale side bars are most typical of the mid- and lower-fan areas and of the braid plains; they are commonly associated with channels of fairly low sinuosity. Large-scale lateral bars, such as those described by Collinson (1970) are typical of the ice-marginal streamways and the distal parts of the proglacial rivers. The occurrence of specific bedforms in these deposits is important especially because the associated lithofacies are the most diagnostic characteristics seen in the proglacial fluvial facies (Rust, 1969; McDonald and Banerjee, 1971; Boothroyd and Ashley, 1975; Church and Gilbert, 1975; Jopling, 1975; Miall, l977,1978b, 1988; Teisseyre, 1978, 1984; Vos and Tankard, 1981; Zieliiiski, 1982). Depositional mechanisms

The depositional mechanisms of these sediments are relatively easy to reconstruct from an analysis of the sedimentary structures. J.R.L. Allen (1982a) has done this for alluvial sediments in general, on the basis of both field observations and experiments. It nevertheless seems worthwhile t o discuss the deposition of typically proglacial fluvial sediments. A first, important aspect is the fact that changes in discharge do not result in similar changes in water level throughout the proglacial fluvial facies. In general, a rise in water level will result in upper-fan and mid-fan zones that are fully covered with water (outside the channels also), whereas the lower braid plains and the pradolinas will show a rise of the water only inside the channels. A discharge that is only little greater may thus affect markedly the deposition conditions in the proximal part of the subenvironment, whereas hardly any effect will be visible further away from the ice (Church and Gilbert, 1975). Differences such as these are reflected, e.g., in the lithologies and sedimentary structures of the various types of bars that are formed within the three subzones (fan, outwash plain, pradolina). Longitudinal bars are formed mainly by vertical accretion, resulting in stacked layers of imbricated gravels. The individual layers are formed

384

The proglacial subenvironment

during the successive stages of maximal flood. The top part of the separate units may be sandy and may show horizontal stratification. N.D. Smith (1970), McDonald and Banerjee (1971), Gustavson (1974), Boothroyd and Ashley (1975) and Vos and Tankard (1981) have interpreted this as a result of deposition when the maximum flood had passed. However, Boothroyd and Nummedal (1978) stated that such horizontally bedded sands could also be formed under conditions of relatively deep water. It is remarkable that the horizontally bedded units are discontinuous and usually restricted t o very limited zones (Williams, 1979). The main body of large-scale longitudinal bars is built of tabular and trough cross-bedded, coarse-grained units (N.D. Smith, 1970; Vos and Tankard, 1981). This cross-stratification is formed during the decrease in current velocity after a flood stage. The sandy deposits on the lee sides of the bars, especially where the bars undergo downstream accretion, represent still later stages of flow decrease (Leopold and Wolman, 1957; Bluck, 1974; Boothroyd and Nummedal, 1978). These units consist of lowenergy tabular cross-stratified sets. The last depositional phase of a flood stage is represented by horizontally laminated units formed during a waning flow (Rust, 1972; Bluck, 1974). This phase is usually followed by a period in which small-scale channels are eroded i n the bar surface (Williams, 1971), a process that can be considered as a first step of bar modification. The depositional mechanisms involved in the formation of longitudinal bars have been both studied extensively and described in detail (by, among others, Doeglas, 1962; Fahnestock and Hamshild, 1962; Williams and Rust, 1969; Boothroyd, 1970; N.D. Smith, 1970; Williams, 1971; Rust, l972,1975,1978a,b;Picard and High, 1973; Gustavson, 1974; Boothroyd and Ashley, 1975; Miall, 1977, 1984; Boothroyd and Nummedal, 1978; Bluck, 1982; Fraser and Cobb, 1982; Zieliiiski, 1982, 1986; Walker and Cant, 1984). Transverse bars occur in a number of settings (Krigstrom, 1962; Collinson, 1970; N.D. Smith, 1970; McGowen and Groat, 1971; Vos and Tankard, 1981; Fraser and Cobb, 1982) and are usually formed in the low flow regime. These bars are generally considered typical of braided channels (Sundborg, 1956; J.R.L. Allen, 1968; Miall, 1977; Gustavson, 1978), particularly of those with a high concentration of traction material (Ore, 1963; N.D. Smith, 1971a; Picard and High, 1973). Most of the material is left on the bars when a flood stage has just passed (McGowen and Groat, 1971) and during low-water stages (Williams and Rust, 1969). The main body of transverse bars consists of a long lee-side depositional front that migrates rapidly in a downstream direction by an avalanche-like grain

Deposits of the proglacial fluvial facies

385

flow downslope from the bar margin (McDonald and Banerjee, 1971; Miall, 1977; Boothroyd and Nummedal, 1978; Rust, 1978b; Casshyap and Tewari, 1982; Fraser and Cobb, 1982).The thickness of the tabular crossstratified units thus formed equals the height of the bar. The top of transverse bars can consist of rippled or horizontally laminated sands (N.D. Smith, 1970; Williams, 1971; Picard and High, 1973) formed during the last stage of a flood, when the water level had already dropped so much that the top part of the bar is only just below the water level. There are reports of large transverse bars that had a complex history, which is expressed by a much more complex internal structure than was sketched above (Miall, 1977;J.R.L. Allen, 1982a). Lateral bars are present in the low-energy areas, where even the braided streams are fairly sinuous (J.R.L. Allen, 1965,1970; Bluck, 1971, 1975/1976; Klimek, 1972; N.D. Smith, 1974; Boothroyd and Ashley, 1975; Boothroyd and Nummedal, 1978). While side bars and composite lateral bars are very common in proglacial river valleys, pointbars are rare. Most of the lateral bars in this subenvironment are characterised by irregular vertical changes in granulometry, particularly if the bars are small (as is usually the case with alternate bars). These bars, including the pointbars, generally grow by both lateral and downstream accretion (Bernard and Major, 1963; Harms et al., 1963; McGowen and Garner, 1975; Jackson, 1976b).

Proglacial sheet- and streamflood deposits (II-B-2-e) Both sheetfloods and streamfloods are common phenomena in areas with an irregular supply of water and have therefore aroused the interest of sedimentologists working in fluvial facies. Most workers use comparable definitions of these types of water currents (e.g., Davis, 1938; Blissenbach, 1954; Potter and Pettijohn, 1963; Bull, 1972; Rust, 1975, 1978; Reading, 1978; Reineck and Singh, 1980; Zielidski, 1980,1982;Van der Meer, 1982; Gradzifiski et al., 1986). Sheetfloods are relatively shallow (some decimetres t o about three metres), high-energy flows that transport debris during periods of floods over widespread areas (Heward, 1978; Hogg, 1982) after having left the shallow channels used during periods of low water (Fig. 235). Sheetfloods may start suddenly, for instance after a heavy rain, and disappear equally fast (McGee, 1897). Gary et al. (1974) mentioned that sheetfloods of considerable duration may result in the formation of deeper channels in which the water currents eventually concentrate. Some researchers also apply the term 'sheetflood' to sheetlike, shallow flows (only a few centi-

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The proglacial subenvironment

Fig. 235. Sheetflood deposit on the midfan in front of the Werenskiold glacier (Svalbard). Note the limited thickness (lower right). Photograph: J. C e d a .

metres deep) which have a much lower current velocity (( 0.5 m s-') t h a n the 'classical' sheetfloods, which reach velocities of up to 10 m s-l (Hogg, 1982). Streamfloods (sometimes termed 'streamflows') have a lower energy than sheetfloods. They also result from a rise in water level but this rise is less pronounced, so that only the channels and the depressions between them are filled with water, while large areas remain unflooded. Streamfloods may sometimes form (during a short time interval) before a sheetflood arises, and they commonly form after a sheetflood has stopped (Krigstrom, 1962; Williams and Rust, 1969). The fact t h a t both sheet- and streamfloods are a result of water currents leaving their channels implies that such floods are rarely (or not at all) present in the upper parts of proglacial fans. They are, however, common in the lower areas where they leave their traces in the form of successive layers that each represents a cycle of rising and falling water. As a result, these floods, although characterised by a n increased flow regime, also leave overbank deposits that may consist of fine-grained material (cf. Czeppe, 1965; McKee et al., 1967; Fahnestock, 1969; Scott e t

Deposits of the proglacial fluvial facies

387

al., 1969; Williams and Rust, 1969; N.D. Smith, 1970; Augustinus and Riezebos, 1971; G.E. Williams, 1971; Klimek, 1972,1974,1981; Boothroyd and Ashley, 1975; Church and Ryder, 1975; McGowen and Garner, 1975; Butzer, 1976; Ruegg, 1977; Heward, 1978; Kumar and Singh, 1978; Miall, 1978b, 1981; Stear, 1978; R.M.H. Smith, 1980; Graham, 1981; Jackson, 1981; Tunbridge, 1981,1983; Vos and Tankard, 1981; J.R.L. Allen, 1982a; Fraser and Cobb, 1982; Hogg, 1982; Zieliiiski, 1982, 1986; Hiller and Stavrakis, 1984; Barvis and Hayes, 1985). Lithofacies characteristics

Sheetflood deposits, i n the classical sense, consist predominantly of horizontally stratified 'blankets' (Fig. 236) but there may also be lowangle units (Williams and Rust, 1969; Boothroyd and Ashley, 1975; Reading, 1978; Tunbridge, 1981; Hogg, 1982). Sands and gravelly sands are the most common feature (McKee et al., 1967; Tunbridge, 1981), but muddy intercalations may occur. Levels with current ripples are inter-

Fig. 236. Section through a series of Drenthian sheetflood deposits formed in the proglacial mid-fan zone. High-energy horizontal lamination predominates. Surroundings of Bolesfawiec (SW Poland).

388

The proglacial subenvironment

preted by Scott et al. (1969) as having been formed at the margins of the sheetflood. The base of single units of sheetflood deposits can show reversed grading, but the top parts often show normal grading (Bryhni, 1978). The high-energy conditions not only result in rapid vertical accretion (up to 1.5 m per flood) but also in the occurrence of antidunes, but these, however, are rarely preserved (G.E. Williams, 1971). Current ripples and horizontal laminations are common, particularly from the periods after the highest flood stage. Streamflood deposits are less widespread and more usually show lowenergy structures such as current ripples and low-energy horizontal lamination. Ripples tend to be of the trough type. The material involved ranges from gravelly in the channels to muddy in the depressions in between (Fig. 237). Sheet- and streamflood deposits have much in common, e.g., the abundance of soft-sediment deformations (Heward, 1978), usually a result of reversed density gradients due t o rapidly changing depositional conditions. Tunbridge (1983) mentioned that the deposits often form cyclo-

Fig. 237. Streamflood deposits of Drenthian age in Upper Silesia (Poland) formed on a proglacial mid-fan. The streamflood deposits (upper half of section) rest on bar deposits. Photograph:T. Zieliliski.

Deposits of the proglacial fluvial facies

389

thems, with a lower part represented by sands with horizontal lamination and current ripples (with climbing ripples being common), and an upper part with more massive muddy sands, sandy muds and clays, formed in stagnant pools (cf. Klimek, 1972). The preservational potential of this upper unit is small, however, since successive floods may again remove the top (Shumm and Lichty, 1963; Shumm, 1968). Textural characteristics Most proglacial sheetflood deposits are relatively well sorted and sorting improves in the downstream direction. Streamflood deposits are usually much less well sorted. Large clasts in sheetflood deposits have an ah-plane parallel to the sedimentary surface, and an a-axis either parallel or perpendicular t o the current direction. Streamflood deposits usually contain more clasts, so that imbrication is a common feature. Occurrence These deposits can be found in the glacigenic sequence in association with proglacial stream deposits. They form relatively widespread bodies that can usually be traced in the field from outcrop to outcrop. It appears that sheetflood deposits in the proglacial subenvironment are most numerous in the fan area, and less so in the outwash plains. Streamflood deposits, however, seem t o prefer the lowest part of this subenvironment (pradolinas),although they are also found in the outwash plains and, even more rarely, in the fan area. Horizontal transitions into other deposits show no preference: the depositional process results in a distribution that allows these deposits to interfinger with all other types of deposits to be found in the same area. Depositional mechanisms Proglacial sheetfloods transport most debris in the form of a traction carpet; minor amounts are transported in suspension. Although a series of sheetflood deposits usually results in a thick sandy succession, the sudden flow of water cannot infiltrate the bottom as the permafrost forms a barrier. Consequently, all water must be discharged surficially, which explains the widespread occurrence of the deposits. The high-energy laminar flow makes small-scale elongated scours in the depositional surface, which are filled by sheet-like bodies with horizontal lamination.

390

The proglacial subenvironment

Fig. 238. Well developed sheetflood succession of Vistulian age, deposited on a sander near Suwatki (NE Poland). The gravels represent deposition in gullies, horizontal lamination is predominantly due to transport in a traction carpet (partly in shallow scours),and ripples indicate low-energy flow conditions. Photograph: T. Zielifiski.

Fig. 239. Typical streamflood deposit showing signs of both high-energy and lowenergy flow conditions. Mid-fan of a Vistulian sander near Suwatki (NE Poland). Photograph: T. Zielifiski.

Deposits of the proglacial deltaic facies

391

The high-energy conditions are reached so quickly that the base of these deposits shows no signs of the lower flow regime; any low-energy structures that may have been formed will, moreover, generally be destroyed by the flood. However, a gradual decrease of the current velocity after the main stage can result in successions that show signs of the waning flow conditions (Fig. 238). The streamflood deposits are formed partly under high-energy conditions; coarse gravels and dunes reflect these conditions (Fig. 239). Most sedimentation, however, takes place once current velocity has decreased (Ostrem, 1975). In contrast t o the rapid deposition during sheetfloods, there are long intervals with slow deposition when a streamflood has passed its maximum. DEPOSITS OF THE PROGLACIAL DELTAIC FACIES (11-B-3) The proglacial subenvironment comprises fewer lakes than does the terminoglacial but some lakes do occur. Proglacial streams embouching in these lakes form deltas (Fig. 240). Descriptions of these deltas abound

active ice

Fig. 240. Schematic model of deltaic sedimentation in a proglacial lake.

392

The proglacial subenvironment

(G.K. Gilbert, 1885; Aario, 1972; Ashley, 1975; Church and Gilbert, 1975; Gustavson, 1975; Gustavson et al., 1975; Jopling, 1975; Shaw, 1975a, 1977b, 1988; Vos, 1977, 1981; Edwards, 1978; Shaw and Archer, 1978, 1979; Cohen, 1979; Eschman, 1979; Le Blanc-Smith and Eriksson, 1979; Parriaux, 1979; Hahszczak, 1980, 1982; Schwan et al., 1980a; Clemmensen and Houmark-Nielsen, 1981; Gilbert and Shaw, 1981; Brodzikowski, 198213, 1984; Leckie and McCann, 1982; Smith and Syvitski, 1982; Smith et al., 1982; Brodzikowski and Van Loon, 1983, 1985c, 1987; N. Eyles, 1983d; Kostaschuk and Smith, 1983; Ruegg, 1983; Syvitski et al., 1984; C.H. Eyles, 1986; Kelly and Martini, 1986; Brodzikowski et al., 1987a,c,d). Proglacial deltaic sediments differ from those in the terminoglacial subenvironment, especially by their finer grain size. This is primarily the result of most of the coarsest material being left behind before the lakes in the proglacial subenvironment are reached. Typical fan deltas, characterised by a coarse average grain size and bad sorting, are rare, although the finding of so-called Gilbert-type fan deltas has occasionally been reported (among others by Clemmensen and Houmark-Nielsen, 1981). Such relatively coarse-grained deltas are found mainly in the proximal zone of the proglacial subenvironment, where sorting processes during fluvial transport had not yet had the chance to be very effective. Proglacial deltas can range from very small t o fairly large. The smallest do not always show well developed sedimentological differences and are then best classified as proglacial deltaic complexes (II-B-3-a). Larger examples can be differentiated into proglacial deltaic topsets (II-B-3-b), proglacial deltaic foresets (II-B-3-c)and proglacial deltaic bottomsets (IIB-3-d). The characteristics of the large proglacial deltas do not differ from those of deltas formed under non-glacigenic conditions, except in their scarcity of organic material (Bates, 1953; Moore, 1959, 1966; Van Straaten, 1960; Coleman and Cagliano, 1964; Gole and Chitale, 1966; Agterberg et al., 1967; Morgan et al., 1968; Bouma and Bryant, 1969; Fisher et al., 1969; McGowen, 1970; Van de Graaff, 1972; Van Loon, 1972; Broussard, 1975; Coleman and Wright, 1975; Flores, 1975; Harms et al., 1975; Sutton and Ramsayer, 1975; Galloway, 1976; Roberts et al., 1976; Cherven, 1978; Heward, 1978; Massari, 1978; Stanley and Surdam, 1978; J.R.L. Allen, 1982a; Hakanson and Jansson, 1983; Miall, 198413).

Proglacial deltaic complexes (II-B-3-a) These deposits are relatively rare, because proglacial deltas, even smaller ones, usually have well developed top-, fore- and bottomsets. However, these complexes may occur, particularly where there are small and

Deposits of the proglacial deltaic facies

393

irregular depressions, and where the streams transport enough debris to fill the depressions rapidly. A high rate of sedimentation commonly results in a vague transition between the subaerial and the subaqueous parts of such a delta and, moreover, the original structure may easily be disturbed by synsedimentary and metadepositional reworking.

Lithofacies characteristics Two types of general lithofacies can be distinguished: that of relatively coarse complexes and that of finer-grained specimens. The group of relatively coarse proglacial deltaic complexes consist mainly of gravelly sand with admixtures of finer material. The individual units consist of slightly inclined parallel beds (Fig. 241) which can show a small decrease in average grain size in a downslope direction, usually from gravelly sand in the top part to sand in the distal part; this decrease may take place over distances as short as 10-15 m (intercalated masstransport deposits may occur that disturb this general picture). The sandy

Fig. 241. Deltaic complex of Elsterian age from the Kleszcz6w graben (central Poland). There is no distinct differentiation, but most of the material has foreset-like characteristics. The material was mainly deposited from traction currents.

394

The proglacial subenvironment

part shows distinct lamination in most cases, either parallel to the sedimentary surface or in the form of ripple foresets. These sands gradually pass into the finer material that constitutes the lacustrine deposits; no clear line can be drawn between the delta and the lake deposits. The finer-grained proglacial deltas consist mainly of sand with admixtures of silt and clay. The individual units tend to be somewhat more inclined (about 15") than those of the coarse-grained proglacial equivalents. The sands usually show distinct signs of transport by water currents, predominantly in the form of cross-bedded sets and horizontal lamination. Scour structures are a very common finding. Textural characteristics Most of these complexes show good t o moderate sorting as a result of current-induced transport. The moderate rounding of individual grains confirms this picture. The clasts found at the bottom of individual sets usually show a distinctly prefered orientation: those deposited in shallow currents (especially in the 'topsets') frequently show imbrication, whereas those deposited under a thicker layer of water tend to be orientated with the a b p l a n e parallel to the sedimentary surface. Occurrence By definition, the deposits are found between fluvial and lacustrine deposits, most of them surrounded by the latter. In a vertical section, the deposits are usually found on top of the same lacustrine deposits with which they interfinger laterally and are overlain by the same fluvial deposits the distal continuation of which they form in a lateral direction. These relationships also determine their position in the glacigenic sequence. The complexes form small bodies no more than five metres thick, and their longitudinal extent rarely exceeds thirty metres; the cross-section is usually even less (the greatest cross-section observed by the authors was fifteen metres). Depositional mechanisms The most important process that leads t o proglacial deltaic sedimentation is the sudden slowing of debris-laden proglacial streams when they embouch in a lake with more or less stagnant water. Most of the material

Deposits of the proglacial deltaic facies

395

settles very soon, thus contributing to the progradation of the delta front. Deposition on this front occurs through stopping of particles that were transported in a traction carpet, by avalanche, and by gravity-induced mass transport. If particles can no longer be transported in the traction carpet, they build up units that can be massive, horizontally laminated or graded, depending on the local conditions and the amount of material involved. Some horizontally laminated beds deposited in the 'topset' part of the complex are most probably formed under high-energy conditions (moderate current velocity but very shallow water, so with a high Froude number); the combination of a decreasing current velocity and an increasing water depth in the downstream direction implies that current ripples and low-energy horizontal lamination can be found in those parts of the complex t h a t might be compared with foresets and bottomsets. Settling of particles from suspension seems to play a very limited role; mass transport, while it is much more important, does not result in a net contribution to the complex formation, as material becomes unstable on the slopes of the delta front and rarely moves down beyond the deltaic area.

Proglacial deltaic topsets These deposits are formed in the area where parts of the delta develop under subaerial conditions, although most of the sedimentation takes place underwater. Several channels are separated from each other by bars and other parts of the alluvial deltaic plain. The channels may be embedded between levees, and interdistributary bays may be present in the plain. The result is a complex geomorphological pattern and a n equally complex mix of deposits (G.K. Gilbert, 1885; Salisbury, 1892; Bates, 1953; Brodzikowski and Van Loon, 1987). The distributary channels are mostly braided and with a usually restricted or very slow lateral migration. Mouth bars are often formed in the mouths of the distributary channels, thus affecting the current pattern, but vertical accretion is the main process influencing the courses of the channels (cf. Putnam and Oliver, 1980; D.G. Smith and Putnam, 1980; D.G. Smith and N.D. Smith, 1980; Putnam, 1982; D.G. Smith, 1983; Walker and Cant, 1984). Subsidence is a favourable condition for vertical accretion; this may have a tectonic origin but deltaic bodies are often so heavy that the underlying sediments undergo heavy compaction, so that a balance between compaction-induced subsidence and depositioninduced overloading is reached.

396

The proglacial subenvironment

The distributary channels are usually surrounded by low levees t h a t pass gradually into the deltaic plain. This plain comprises numerous depressions (interdistributary lakes or bays), where even the finest material may settle after flooding. The final nature of the topset deposits depends largely on the relative effects of current and wave activity. Most proglacial topsets are riverdominated, but wave action may become important where the topsets extend far into a large lake (Aario, 1972; Church and Gilbert, 1975; Elliot, 1978; Clemmensen and Houmark-Nielsen, 1981; Dixon, 1981; Tankard and Barwis, 1982; Miall, 1986).

Lithofacies characteristics Most proglacial deltaic topsets consist predominantly of sand. The distributary channels (Fig. 242) become filled with sets t h a t show medium-scale trough cross-stratification; these sets pass laterally in horizontally stratified sets and then, near the channel margins, in rippled sands. If the topsets are gravelly, the thickness of the separate sets

Fig. 242. Channel deposits formed in the topsets of a proglacial delta of Elsterian age (Kleszcz6w graben, central Poland).

Deposits of the proglacial deltaic facies

397

increases; medium-scale sets with tabular cross-bedding replace then the rippled sands. These sets will build up bars, thus decreasing the channel depth. On the channel side, the levees consist of rippled sands and, on the outer side, of laminated mixtures of sand, silt and clay. Levees may be lacking in coarse-grained deltas (Clemmensen and Houmark-Nielsen, 1981). The delta plain is formed by fine sand and mud with current ripples and horizontal lamination. Depressions may be filled with fine-grained deposits that resemble lacustrine sediments (Fig. 243), but there may also be concentrations of gravel that was supplied through crevasses. The authors have studied mouth bars in detail in the Kleszczbw graben, where they tend to consist of medium and fine sand (Fig. 244), with current ripples as the most common structure. The proximal parts of these bars contain horizontally stratified or slightly inclined units that pass into rippled sands both vertically and laterally. Trough crossbedding is found occasionally in these bars, especially at their base (cf.

Fig. 243. Finely laminated muds and fine sands formed in a depression of a proglacial deltaic plain. The laminated sediments are intercalated between coarser deposits formed during a flood stage. Kleszczdw graben, central Poland.

398

The proglacial subenvironment

Fig. 244. Detail of a proglacial distributary mouth bar of Drenthian age (Kleszczow graben, central Poland).

Moore, 1966). Subhorizontal lamination is also common. Mud drapes indicate that quiet periods have alternated with the intervals during which the bars underwent vertical accretion. Most topsets, particularly those in fine-grained proglacial deltas, show distinct cycles, due t o successive phases of flooding. The cycle starts with trough cross-bedded sands that represent channel-bottom deposits; these sands are followed by tabular cross-stratified or horizontally stratified sands, covered by a unit with small-scale cxurrent ripples and occasionally climbing ripples at the top. The top of such a cycle consists of either laminated or massive mud. Textural characteristics The textural characteristics of these topsets do not differ much from those in the distal proglacial alluvial plain. Sorting may, but need not be somewhat better. Gravel at the base of the channels may show imbrication. There are no textural characteristics that could be considered as diagnostic for these deposits.

Deposits of the proglacial deltaic facies

399

Occurrence

The topsets are usually a few metres thick (Clemmensen and HoumarkNielsen, 1981) and rarely exceed the thickness of five metres that was observed in the Kleszcz6w graben (central Poland). The lateral extent is more impressive: this may be hundreds of metres in the downcurrent direction, and a transverse section may exceed 100 m. The rather chaotic lateral contacts of the various subtypes of deposits (channel lag, levee, bar, interdistributary deposit, etc.) are a characteristic feature, but make recognition of 'fossil' topsets in badly exposed areas extremely difficult. The lateral contacts in a longitudinal section are equally characteristic and much more regular: the topsets form the continuation of proglacial stream deposits and themselves pass into proglacial deltaic foresets (Fig. 245). In the glacigenic sequence, these deposits form part of proglacial deltaic sequences that are found between proglacial fluvial and lacustrine deposits.

Fig. 245. Proglacial deltaic topsets, overlying the foresets of the same delta body. Lower Silesia, SW Poland.

400

The proglacial subenvironment

Depositional mechanisms

The depositional mechanisms do not differ from those in non-glacigenic environments. The most typical elements are the distributary mouth bars. These are formed only if the water level remains below certain values: low water is usually accompanied by an unduly small supply of mineral particles, whereas high-water conditions are characterised by erosion instead of deposition in the flow-separation zone (Ashley, 1975; Church and Gilbert, 1975; Clemmensen and Houmark-Nielsen, 1981). The intermediate water levels required for the formation of mouth bars also favour the development of trough cross-stratification, which is indeed the most common structure in these bodies (Moore, 1966). There may also be ripple-drift cross-lamination and wave ripples, which are much less important phenomena and are usually formed as a result of low-water periods. In most cases, these bars show a fining upward grading, and a sandy horizontal lamination in the topmost part (Tanner, 1960>,formed under low-energy flow conditions.

Proglacial deltaic foresets (II-B-3-c) The wide variety of the nature of proglacial deltaic foresets is evident from the literature on these deposits. They may be either coarse- or finegrained (Aario, 1972; Clemmensen and Houmark-Nielsen, 1981), gently inclined or steeply dipping (Fulton and Pullen, 1969; Aario, 1972), showing distinct cycles or lacking them (Hahszczak, 19821, resembling Gilbert-type deltas or showing gradual lithofacies changes in a downslope direction (J.R.L. Allen, 1968,1970d; Pharo and Carmack, 1979). As a rule, the foresets constitute the main delta body, being the direct result of delta progradation. The shape of the foresets depends largely on the nature of the water influx: a hyperpycnal flow will favour the formation of gently inclined, elongated foresets whereas a hypopycnal flow will generally result in much steeper foresets. In addition, hypopycnal flows may easily result in a cyclic succession. Gravity-induced avalanching is the most important process, but bottom currents, back-flow currents, settling from suspension and mass flows can distinctly influence the nature of the foresets. Lithofacies characteristics

Most of the proglacial deltaic foresets consist of sands or finer material. Gravel is a rare constituent. The most characteristic feature is the

Deposits of the proglacial deltaic facies

40 1

Fig. 246. Proglacial deltaic foresets from the surroundings of Swidnica (Sudetes foreland, SW Poland). Note the reactivation surfaces on the inclined sets.

distinct inclination of the individual sets (Fig. 246), each 5-20 cm thick. The sets may show a number of structures related to the depositional processes, such as scour-and-fill structures, horizontal ( = parallel t o the inclined sedimentary surface) lamination, and cross-bedding. Graded bedding may be found, resulting from both turbidity currents and gradually decreasing current velocities; mud drapes indicate periods of (almost) stagnant water. Current ripples indicating a flow direction opposite t o the hormal' one can be found, particularly in the lower parts of the foresets, although the ripples may also be present at higher levels (Clemmensen and HoumarkNielsen, 1981).They are interpreted as backflow ripples.

Textural characteristics Good sorting is a characteristic of most proglacial foresets. There is a n increase in average grain size in the inclined sets from the top part downwards, as the avalanching process of individual grains gives the largest grains the highest momentum. Foresets that dip only little may, however, show a slight decrease in grain size in the downslope direction, because the grains cannot roll over the surface but have t o be transported by water currents instead.

402

The proglacial subenvironment

Pebbles, if present and if flattened, are orientated with their ah-planes parallel t o the inclined sedimentary substratum, except for those which form a layer at the base of the foresets, where they are either orientated with their ah-planes in horizontal direction or are imbricated. There are no specific textural characteristics that can be considered diagnostic for these deposits. Occurrence The occurrence of these foresets is best explained by the model of a Gilbert-type delta. The contacts with the bottamsets underneath and with the overlying topsets is usually relatively sharp; the contact is very sharp if the foresets are overlain by lacustrine deposits. The lateral transition between the deltaic foresets and the related bottomsets is much less distinct: there are only gradual changes. The position of the proglacial deltaic foresets in the glacigenic sequence follows from the contacts mentioned above: they form the middle part of the deltaic succession that rests on top of proglacial lacustrine deposits and that is, in most cases, overlain by proglacial fluvial deposits. The foresets form relatively large sedimentary bodies. A delta that was described from the Lillooet Lake has an extent of about 1.5 km, in both longitudinal and cross-section, and is mainly built up of foresets. Proglacial deltaic foresets that were well exposed in the Kleszczow graben (central Poland) had extents of the order of 0.5 km and a thickness of up t o 40 m, which makes them by far the largest deposits there. It appears that the individual sets in these foresets show frequent reactivation surfaces, a feature which can be explained by an alternation of periods with and without substantial subsidence. Depositional mechanisms Grain avalanching is the process that makes the greatest contribution to sedimentation. This process results in the classical large-scale, inclined foreset stratification. The vast amounts of material thus deposited in apparently short periods indicate that the particles were usually supplied by high-energy currents (flood periods), although avalanching may continue under low-energy conditions. A hypopycnal character of the inflowing water also favours grain avalanching. Hypopycnal flow results in bottom traction and the foresets formed under conditions such as these will be less inclined. Gravity-induced

Deposits of the proglacial deltaic facies

403

avalanching only plays a minor role. The bottom currents that are then responsible for the deposition of particles can be compared with subaqueous sheet flows (Shaw and Kvill, 1984; Shaw et al., 1989). If the 'subaqueous sheetflood' is very strong, a traction carpet may be produced that will result in parallel lamination; reversed grading may also be present. Linear currents are rarer and occur during low-energy conditions. These currents produce slightly inclined foresets t h a t a r e massive or laminated parallel t o the bedding. Grading can be produced during gradual decreases i n current velocity. Even the low-energy bottom currents may result in scours. The elongated depressions thus formed are usually filled with rippled sands. Cyclic sedimentation is common i n proglacial foresets. There are several types of circumstances that produce cycles, and they may be superimposed on each other. The largest cycles are season-induced and result from the alternation of quiet periods during the winter (when the lake is covered with ice) and more turbulent conditions during summer. The winter period is often responsible for the gradual settling of suspended particles, partly in the form of mud drapes, whereas muddy intervals are usually not formed during the summer. Other cycles are the result of a suddenly increased supply of water and debris, for instance during spring when ice starts to melt. The most common cycles a r e related to smaller fluctuations in water supply due, for instance, to floods. Such cycles are usually characterised by fining-upward layers of some millimetres to some decimetres thick. Slumping is a process that makes a major contribution to sedimentation on the foresets. The slump deposits will be described in more detail in the section on proglacial mass-transported sediments.

Proglacial deltaic bottomsets (II-B-3-d) The distal parts of the proglacial fresh-water deltas gradually pass into lacustrine bottomsets. Deltaic and lacustrine bottomsets also occur i n other glacial subenvironments, where they tend t o show several comparable characteristics. The only way t o distinguish them is, i n most cases, by analysing t h e lateral and vertical facies transitions (cf. Brodzikowski and Van Loon, 1983). However, there are occasionally characteristics that simplify their recognition (Aario, 1972; Ashley, 1975; Clemmensen and Houmark-Nielsen, 1981; Brodzikowski and Van Loon, 1987). Proglacial deltaic bottomsets tend to be fine-grained and to extend in all directions, well away from the foresets (Houbolt and Jonker, 1968;

404

The proglacial subenvironment

Lara and Sanders, 1970; Aario, 1972; Ashley, 1975; Shaw and Archer, 1978,1979; Haluszczak, 1980,1982; Clemmensen and Houmark-Nielsen, 1981).I t is typical that they form rhythmites, but these do not all have the same origin (Antevs, 1925,1951; Kuenen, 1951; Granar, 1956; Lajtai, 1967; Agterberg and Banerjee, 1969; Gilbert, 1972b; Banerjee, 1973; Gustavson et al., 1975; Sturm and Matter, 1978; Brodzikowski and Van Loon, 1987).

Lit hofacies characteristics Deltaic bottomsets consist mainly of rhythmites. These are developed differently from the lacustrine equivalents. This is only a general trend but one that can sometimes be used as a criterion. The levels of rippledrift cross-lamination, the relatively high silt content, the parallel lamination and the many synsedimentary and penecontemporaneous deformations are characteristic of almost all bottomsets. The deltaic bottomsets tend t o show more flowage structures than load casts and related structures, whereas the reverse is usually observed in lacustrine bottomsets; it should be stressed again that this difference cannot be used as a reliable criterion for distinguishing between deltaic and lacustrine bottomsets (Brodzikowski and Van Loon, 1985b; Eyles and Clark, 1985). Well developed bottomsets were studied in the Kleszczow Graben by the authors. According to Hahszczak (1980) and Brodzikowski et al. (1980) the main constituents are silt and sand, while clay occurs principally as thin (1-2 cm) covers of the individual cycles. Similar, so-called 'deltaic varves' have been described by Gustavson (1975) and Ashley (1975). The rhythmic character is ascribed to seasonal influences, with subcyclicity of various origin. The subcycles start with a bottom part of fine-grained sand with ripple marks, followed by a n interval of alternating sandy and silty laminae and a top part of mainly silty laminae with the cover of clay mentioned above. There may be 4-10 of these subcycles visible within one main cycle, often showing well developed load casts and flame structures a t the base (Hahszczak and Brodzikowski, 1982; Brodzikowski et al., 1985; N. Eyles and Clark, 1985). The most common characteristics of the bottomsets are sandy horizontal lamination, massive sandy and silty layers, rippled sands, thick units with climbing ripples, finely laminated muds, and clay drapes (Fig. 247). Scours are frequently present; they have a gently concave base and are usually filled by massive or horizontally stratified sands sometimes by diamictic masses with deformation slabs (Gustavson et al., 1975; Shaw and Archer, 1979).

Deposits of the proglacial deltaic facies

405

Fig. 247. Proglacial deltaic bottomsets from the Kleszcdw graben (central Poland). Top (photograph: A. Hahszczak): characteristic proximal example. The overall picture is that of horizontally laminated muds, but several layers with current ripples and ripple-drift cross lamination are present. Clay drapes are found over some of the rippled surfaces. Bottom: distal bottomsets passing into lacustrine deposits. Note the dark clay layers deposited during winter and the thick units of summer deposits with intraseasonal cycles.

406

The proglacial subenvironment

It appears that typical turbidites, with their Bouma intervals, are much rarer than had been thought in the 1970s. However, turbidite successions can be found under exceptional circumstances, for instance with current ripples at the base, then a zone with ripple-drift crosslamination with an increasing angle of climbing in the upward direction, followed by wave ripples with occasional mud drapes, and, finally, horizontally laminated muds or graded sandy muds with a clay drape on top. Textural characteristics These deposits are fine-grained. Gravels are rare if not absent at all. The individual quartz grains are moderately t o well rounded and show bright surfaces. There is good sorting in each layer, but layers of different grain sizes can frequently alternate. Sorting within one layer is rarely present in a lateral direction, but grading is common, if not the rule. There may occur a few shallow channel-like depressions filled with sediments that show reversed grading, but this is an exceptional feature. There are no diagnostic textural characteristics, but by far most finegrained graded layers are found in deltaic and lacustrine bottomsets. Occurrence These deposits form the lowest part of proglacial deltaic successions. They usually rest on proglacial lacustrine deposits but the latter are not necessarily distinctly developed, so that the delta deposits may seem t o rest on other proglacial (or even non-glacigenic) sediments. The deposits are widespread, forming a wide belt around the associated deltaic foresets, and gradually pass in a downcurrent direction in proglacial lacustrine bottomsets. The thickness of these deposits varies widely, as local factors have a tremendous influence: depth of the lake, supply of particles, density of the inflowing water, frequency of floods all determine largely where, when and how much material will be deposited. In spite of this, there is a rough rule that the thickness is less than 50% of the thickness of the associated foresets if there are no 'external' factors (e.g., a subsiding basin) to disturb the normal sedimentation pattern. The general picture of these deposits is one of quietly deposited fines on a widespread subaqueous plain. This picture is, however, not correct in its details, while prodeltaic bars can be found in the proximal part of the bottomsets (if relief conditions and supply of material are favourable).

Deposits of the proglacial deltaic facies

407

The bars themselves (up to some tens of metres long, occasionally about 100 m) consist mainly of rippled sand, and the channels between the bars are also filled with relatively coarse material that shows current ripples and frequent ripple-drift cross-lamination. The more distal parts conform much more closely the general picture: some thin units can be traced over several hundred metres (R. Gilbert, 1972b).

Depositional mechanisms The main depositional mechanism is settling: from bottom currents, from suspension and from gravity-induced mass movements. Bottom currents may occur in combination with quiet water above the current, so that particles settling from suspension and from bottom transport in the form of current ripples result in climbing ripples. This process is responsible for fast vertical accretion. Material settling from suspension requires that material be first brought into suspension. This occurs predominantly during stages of fluvial floods. These floods have a pulsational character (although less than in the terminoglacial subenvironment) so that periods with much material settling from suspension alternate with periods of less settling. This alternation is expressed in the vertical repetition of units with climbing ripples. These units are sometimes divided from each other by rippled layers, indicating that bottom transport continued while there was no settling from suspension. The reverse, which would be reflected in the presence of thick, rippled units with occasional layers showing climbing ripples, is only exceptionally present. However, rippled sands, indicating that bottom currents were the main depositional agent, are found in the proglacial bars. These bars show vertical and lateral transitions that are roughly the same: current ripples are followed by climbing ripples, then by wave ripples and parallel lamination. This picture can be explained as being the result of changing conditions when the lake becomes covered with ice during the winter. Turbidites are rare, although there are several reports of micro-scale turbidites (2-3 mm thick). These form part of the winter laminae in muddyklayey rhythmites only 2-3 cm thick. It is known that tectonic activity may trigger the micro-scale turbidity currents, but these have also been observed outside tectonically active regions (Shaw, 1977; Shaw and Archer, 1978,1979). There is discussion as to the whether or not they are true turbidites (cf. Kuenen, 1951; Lajtaj, 1967; Banerjee, 1973). The high floods that contribute so much to deposition in the proglacial fluvial facies almost no influence on the deposition of the deltaic bottom-

408

The proglacial subenvironment

sets. The sudden increase in water and sediment supply may result in unstable slopes, for instance on the proglacial deltaic foresets. If this triggers mass transport, the resulting flows tend t o either stop in the lower part of the forsets or pass through the area of the deltaic bottomsets, so as t o reach the area of the lacustrine bottomsets (Shaw and Archer, 1978,1979; Eyles and Miall, 1984). DEPOSITS OF THE PROGLACIAL LACUSTRINE FACIES (11-B-4) An abundance of lakes may form in the proglacial subenvironment. This may be a result of topographic depressions whence the water cannot escape as groundwater, due to either an impermeable sediment (e.g., a till from a previously more advanced ice sheet) or frozen subsoil (permafrost). In addition, the irregular supply of both water and debris from the proglacial braided rivers may easily result in debris barriers behind which lakes form in the dammed-off river course. The size of these lakes varies enormously: from small ponds (Fig. 248) t o lakes hundreds of

Fig. 248. Small proglacial lake in the Hornsund area (SWSvalbard). Photograph: J. Cegla.

Deposits of the proglacial lacustrine facies

409

kilometres in diameter, with depths ranging from less than one t o several tens of metres. The proglacial lakes need not be 100%glacigenic, as rivers coming from the extraglacial subenvironment may embouch in them. This, however, is a relatively rare phenomenon. The sediments formed in proglacial lakes are generally much finer those those formed in terminoglacial lakes. Relatively coarse fans and mass-transported sediments like those seen frequently in the terminoglacial lakes are rare or even absent. Three main types of deposits are found in this facies: proglacial lacustrine complexes (II-B-4-a),proglacial lake-margin deposits (II-B-4-b) and proglacial lacustrine bottomsets (II-B-4-c).

Proglacial lacustrine complexes (II-B-4-a) Several of the lakes in the proglacial subenvironment are so small that they fill with sediment within a short time. This does not give the typical distinctions between lake-margin deposits and lacustrine bottomsets a chance t o develop. For the present purpose, the 'mixtures' will be termed 'complexes'. A number of field descriptions refer t o such complexes, though never in sedimentological detail (Bartkowski, 1967; Klatkowa, 1972a; Klimek, 1972; Drozdowski, 1974, 1979, 1986; Krygowski, 1974; Kozarski and Szupryczyiiski, 1978; Kfysz, 1981; Brodzikowski, 1982, 1984). The most important difference between terminoglacial and proglacial lacustrine complexes is that the latter filled predominantly with particles settling from suspension and fine-grained debris supplied by bottom currents and by wind activity, while mass-transport only played a very minor role. Lit hofacies characteristics Most complexes consist of rippled sands, laminated sand and laminated or massive muds (Fig. 249). Relatively coarse bodies may occur where a distributuary stream embouches in the lake. The lamination is almost without exception of the low-energy type, parallel to the sedimentary surface. There are some reports of proglacial lake deposits that become finergrained towards the top. The relative frequency of current ripples shows the same tendency, with rippled units in the basal part being common, whereas ripple trains or isolated ripples are more common in the top part. There are also other, though weakly developed, regular features. The

410

The proglacial subenvironment

Fig. 249. Proglacial lacustrine complex near Potworow (Sudetes foreland, S W Poland), with current ripples, climbing ripples, horizontally laminated muds and clay drapes.

sediments near the feeding river usually consist of stratified sands that show a n increasing influx, in a downcurrent direction, of settling from suspension; this is reflected by climbing ripples. These rippled sands gradually pass into laminated muds. Such regularity is usually absent from any other point going from near the lake margin towards the centre. When no debris is supplied by a river the complexes consist almost entirely of laminated muds.

Textural characteristics Fine sand and silt dominate (Fig. 250), with medium sand and clay as common admixtures. Most of the sorting of the material has usually already taken place before the particles reach the lake; sorting is little improved once inside the lake. Settling from suspension, e.g., of wind-supplied particles, may result in grading-upward units, especially after large influxes of sediment-laden river water. Truly varve-like deposits are however rare and reversed grading even more so.

Deposits of the proglacial lacustrine facies

41 1

Fig. 250. Alternation of coarser and finer laminae, produced as a result of reEetitive influxes of sediment-laden river water in a small lake. Exposure near Bardo Slgskie (Sudetes foreland, SW Poland).

0cc urre nee While lakes can develop throughout the proglacial subenvironment, most are found in the alluvial plains (Klimek, 1972) and in the higher zones of ice-marginal streamways (Galon, l961,1968a,b; Kozarski, 1962a, 1965). In addition, they have been described from places where there is no water supply from streams (Klatkowa, 1972a,b; Szponar, 1974, 1986; Brodzikowski, 1982; Van der Meer, 1983) and where they may develop because the permafrosted subsoil prohibits infiltration and percolation of rain or meltwater. Most complexes have a sharp base, reflecting the change from subaerial t o subaqueous conditions. The upper boundaries can also be sharp but usually form gradual transitions into deltaic or fluvial deposits. These relationships are also expressed by the position of these complexes in the glacigenic sequence: they can be found throughout the proglacial interval, in most cases directly associated with fluvial and deltaic sediments.

412

The proglacial subenvironment

Depositional mechanisms There are two processes that play an important role: sedimentation from bottom currents and settling from suspension. Their relative importance depends on the water and debris supply from proglacial streams, and on the relative distance from the depositional site to the influxes. Settling from suspension may increase if there is much aeolian activity supplying particles by air. Deposition of wind-blown material on a winter ice layer may result in graded beds that differ distinctly from the 'normal' sedimentation. The energy levels involved in the sedimentation processes are, as a rule, very low. Indications of an upper flow regime are absent or very rare: the finding of a few current ripples between a n abundance of laminated muds is much more common (Fig. 251).

Fig. 251. Detail of a succession of a proglacial lacustrine complex of Drenthian age, developed in the Kleszczow graben (central Poland). Current activity was distinctly dominated overruled by settling from suspension. Height of succession shown: approx. 3.5 cm. Photograph: A. Hahszczak.

Deposits of the proglacial lacustrine facies

413

Proglacial lake-margin deposits (II-B-4-b) These deposits can be classed as a separate type because the lake in which they are formed is so large that one can differentiate between deposits formed under the influence of waves and/or the surrounding area, and deposits not affected by 'external' factors. Lakes so large t h a t such a differentiation is possible are generally also large enough that they are relatively long-lived. This implies that many of these lakes are continuations of terminoglacial lakes, and that numerous others changed in time into terminoglacial lakes (cf. Brodzikowski and Van Loon, 1980, 1983; Brodzikowski, 1984; N. Eyles et al., 1985; Drewry, 1986; C.H. Eyles, 1986; Donnelly and Harris, 1989). This shows how difficult it can be to distinguish between proglacial and terminoglacial lacustrine deposits, and particularly between the lake-margin deposits. The lake-margin deposits considered, and as defined earlier, do not include deltas, slumps and similar depositional bodies. There are other phenomena, however, that are included, such as barrier spits (cf. Hine and Boothroyd, 1978) and units with hummocky cross-stratification (Van Dijk et al., 1978; Hamblin and Walker, 1979; C.H. Eyles, 1986; N. Eyles and Clark, 1986). There may also be a n influence of drift ice, for instance in the form of pushed material near the lake shore (Pyokari, 1981).

Lit hofacies characteristics According to several authors, hummocky cross-stratification (HCS) is a diagnostic featurea. This large-scale type of stratification connected with large storm waves was studied by, among others, Dott and Bourgeois (19821, Hunter and Clifton (1982),Walker et al. (1983); Brenchley (1985) and Duke (1985). The undulatory cross-strata are subparallel, show lowangle terminations and truncations, and a r e either convex-upward (defining a hummocky, dome-like structure) or concave-upward (defining a swale; N. Eyles and Clark, 1986, apply the term 'swaley cross-stratification' t o this case). These structures are usually built of medium- t o finegrained sand; they are radially symmetrical and show wavelengths of up to four metres while their heights are less than forty centimetres (N. Eyles and Clark, 1986). The laminae in the top part of the hummocky cross strata are usually truncated and covered by a unit with symmetrical wave ripples and mud drapes. A succession studied in Ontario (N. Eyles and Clark, 1986) show thinning of the HCS units towards the top. The barrier-spit complex described by Hine and Boothroyd (1978), which is probably of partly terminoglacial and partly proglacial origin,

414

The proglacial subenvironment

was found superimposed on channel deposits formed on an outwash plain. The complex consisted of units with cross-bedded units stratified horizontally or at a low-angle towards the basin centre. The migrating swash bars contained predominantly planar beds and beds with planar tangential cross-stratification. The succession was covered by wind-blown sands. Most of the lake-margin deposits show no such special features: they consist of material deposited in the relatively shallow belt of water a t the lake margin. Such deposits consist mainly of sands, muddy sands and sandy muds, with abundant ripples of medium size (height 1-3 cm; wavelength 5-12 cm) in the sandy parts and wavy lamination in the muddy parts (Fig. 252). If wave action has not played an important role, both the sandy and the muddy units show irregular laminations; these are usually only 1-3 mm thick and extend over a few metres. It is not uncommon that these deposits have undergone numerous phases of disturbancen as a result of pushes by ice rafts. The deformations often have an imbricate structure (Fig. 253), as described by Blackwell (1965) from Harding Lake in central Alaska. The lithology of these socalled ice-shove ridges depends on the nature of the substratum; fluctuations in the water level of the lake result in lake-margin deposits commonly forming part of these ridges.

Fig. 252. Rippled sands and muds with wavy lamination, forming part of proglacial lake-margin deposits. Sudetes foreland, SW Poland.

Deposits of the proglacial lacustrine facies

415

Fig. 253. Imbricate structure of an ice-shove ridge. From: Washburn (1979).

Textural characteristics The textural characteristics of these deposits are not diagnostic. Sorting is good to moderate; prolonged wave action may result in laminae with concentrations of heavy minerals. The larger clasts, if present, are usually orientated with their ah-planes parallel to the sedimentary surface. Occurrence The distribution of these deposits over the area occupied by the lake depends on local factors such as geometry of the lake, depth profile, prevailing wind directions, etc. Atability of the water level in the lake is another important factor. However, the general configuration of these deposits is that of an irregular belt between non-subaqueous deposits on the outside and lacustrine bottomsets inside. These relationships also determine the areal relationships with other deposits and, consequently, the position in the glacigenic sequence. The thickness of the deposits varies greatly. This is only partly a result of the erosional processes that affect these deposits as soon as lowering of the water level exposes them subaerially. The supply of material from the borders or by bottom currents is equally important, as are the presence of wave activity, the size of the lake and the inclination of the marginal slopes (Blackwell, 1965; C.H. Eyles, 1986; N. Eyles and Clark, 1986; Brodzikowski and Van Loon, 1987). The thickness, however, rarely exceeds ten metres.

416

The proglacial subenvironment

The base of these deposits is sharp, as in the case of proglacial lacustrine complexes. The lateral contacts are gradual and the upper boundary is also transitional rather than abrupt. Depositional mechanisms

The same processes play a role as they do in the terminoglacial subenvironment. Bottom currents and wave action play a significant part, while reworking disturbs the general picture. For further details the reader is referred to the equivalent discussion in the section on sedimentation in the terminoglacial subenvironment. The main differences from deposition in terminoglacial lakes are a result of the commonly larger size of proglacial lakes, where there is, among others, more wind activity and, especially, a greater influence of storms. The resulting hummocky cross-stratification has already been mentioned. The precise hydrodynamic setting and evolution of HCS are not yet fully understood, but there is general agreement that the marked stratification results from strong oscillatory flows, or at least a combined oscillatory-dominant flow at the base of storm waves that have whirled up large amounts of sand. Hummocks could possibly be regarded as extremely large post-vortex ripples; individual laminae might reflect pulses of liquefaction, suspension and sedimentation of sand caused by one single storm wave or wave train (Dott and Bourgeois, 1982; N. Eyles and Clark, 1986; Greenwood and Sherman, 1986). Hummocks show a sharp base, a massive lower part, then a unit with planar lamination, followed by hummocky and draped lamination and, finally, wave ripples and mud drapes (Walker et al., 1983; N. Eyles and Clark, 1986). The sharp base and the lower massive unit might represent sand emplacement by storm-related currents; the laminated unit might represent reworking by oscilatory flow resulting from storm-induced waves; the mud drapes represent settling from suspension when t h e wind disappeared. Eyles and Clark (1986) stated that hummocky crossstratification develops above the storm wave base, while it will not survive reworking above the fair-weather wave base. HCS might therefore be considered as diagnostic for lower shoreface zones (Walker, 1984); computations suggest that HCS is formed in water 2-11 m deep, when waves have a height of 1.5-4 m and a period of 3-9 seconds. In the recent past, ice-shove ridges were found at several locations in Alaska and Arctic Canada (Dionne and Laverdihre, 1972; Hume and Shalk, 1976; Dionne, 1978), but they had been described far earlier (Gilbert, 1890; Tyrell, 1910; Hobbs, 1911; I.D. Scott, 1927). They are in

Deposits of the proglacial lacustrine facies

417

fact deformation structures, but the deformation is a result of agents that form part of the facies conditions. Two models have been developed that describe the formation of the ridges. The first model (Tyrell, 1910; Washburn, 1947; Ward, 1959) suggests that the ridges result from the grinding action of ice rafts driven against the shore by wind drag or by wind-generated waves and currents. The second concept (I.D. Scott, 1927) suggests that the ridges develop through thermal expansion of the ice cover when the temperature rises following cracking of the ice, filling of the cracks with water, and refreezing of the water in the cracks. The ice cover must then be rigid enough to act as a strut in transmitting stress, and the extent of thrusting in a given year is cumulative and depends on the number of successive periods of cracking, filling, refreezing and expanding (Hobbs, 1911). Zumberge and Wilson (1953) proposed that a rise of about 1 "C per hour in the air temperature during twelve successive hours, in combination with an ice thickness of 20 cm, already suffice to cause ice thrusting. Pessl(1969) reported that such a temperature rise needs t o continue for only six hours, and that the shoreward motion of the ice under such conditions amounts t o one metre a t the shore, a movement that requires thirty days.

Proglacial lacustrine bottomsets (II-B-4-c) This type of deposit has long been the object of much attention (Sauramo, 1923; Antews, 1925,1951; Hansen, 1940; Kuenen, 1951; Mathews, 1956; Bagnold, 1962; Fulton, 1965; Lajtai, 1966; E.H. Walker, 1967; Banerjee, 1968, 1973; Agterberg and Banerjee, 1969; Chambers, 1971; Ashley, 1972,1975; R. Gilbert, 1973,1975; Church and Gilbert, 1975; Shaw, 1975, 1977b; Merta, 1978; Shaw and Archer, 1978, 1979; Shaw et al., 1978; Brodzikowski and Van Loon, 1980, 1983, 1987; Waitt, 1980). Several studies have emphasised the point that some aspects of sedimentation in the centre of proglacial lakes are not yet fully clear (Embleton and King, 1975; Francis, 1975; Sugden and John, 1976; Edwards, 1978; Washburn, 1979; Reineck and Singh, 1980; N. Eyles, 1983; Hakanson and Jansson, 1983; N. Eyles and Miall, 1984; Drewry, 1986). The depositional conditions are however much less chaotic than those in the terminoglacial subenvironment, a fact which is reflected by the abundance of widespread, uniform deposits. Sedimentation is mainly due to settling from suspension, turbidity currents and weak bottom currents. The resulting deposits consequently have a rather monotonous character, but detailed analyses have shown that these sediments, the varvites in particular, have various origins.

418

The proglacial subenvironment

Lithofacies characteristics Silt and clay are by far the most common constituents; fine sand is also present, but usually only as minor admixtures. There are few other types of deposits with such well developed horizontal lamination and with such a n abundance of normally graded layers. These features, combined with the generally widespread occurrence of the deposit, might be considered as characteristic. When studied in detail, the graded layers often appear more complex than first expected. Units with a coarser lower part and a finer upper part commonly appear to present no true grading but to be a combination of t w o graded units with a break between the two. Subcycles may be distinguishable within one unit, In addition, the - seemingly - continuous laminae are usually of restricted extent (a few metres), and micro-scale current ripples or micro-scale tabular cross-bedding may be found a t specific levels within one unit (Fig. 254). Other units appear massive in the field, but microscopic or X-ray analysis will usually show t h a t such 'massive' units have t h e same internal structures as those just described. Still other types of units show typical successions of Bouma intervals (or the 'distal' intervals only), indicating deposition from turbidity currents.

Fig. 254. Graded proglacial lacustrine bottomsets of Drenthian age (Kleszcz6w graben, central Poland). Note the occurrence of levels with micro-scale ripples and the sandy lamina with abundant micro-scale loadcasts. Photograph: A. Hahszczak.

Deposits of the proglacial lacustrine facies

419

Coarse-grained intercalations, sometimes with a diamictic lithology, may be found (Fig. 255). Such units, which represent mass-transported deposits almost without exception, will be dealt with later (II-B-6). Cyclic sedimentation appears here under several forms. The thickest cycles the authors have observed (in the still active Kleszczow graben) were over two metres thick and consisted of material that had settled after bottom sediments had gone into suspension after an earthquake. Most cycles are much thinner (20-30 cm in the outer zone, gradually thinning to 3-4mm in the centre) and represent turbidites. The most abundant type of cycles consists of varvites, which are commonly 2-5 cm thick (Banerjee, 1973; Ashley, 1975; Brodzikowski and Van Loon, 1980, 1983, 1987; Haiuszczak, 1980, 1982), but local conditions determine the thickness.

Fig. 255. Small-scale varvites from the Kleszczdw graben, with one unit that contains somewhat larger clasts in between. This layer represents a micro-scale mass-flow deposit, probably formed after a n earthquake.

Textural characteristics The extremely good sorting of these deposits as a n entity and the abundance of grading are the most conspicuous textural characteristics. Grains of sand size usually show signs of fluvial transport, though grains indicating an aeolian origin are much more common than in the terminoglacial subenvironment.

420

The proglacial subenvironment

Fig. 256. Unstable density stratification and fluidisation structures in proglacial lacustrine bottomsets of Drenthian age (Kleszcz6w graben, central Poland).

Sand grains and - if present - coarser particles show a well developed fabric, with the ah-planes horizontal or, in the case of deformed beds, parallel t o the layer boundaries. Sedimentary deformations are found throughout these deposits (Fig. 256). Occurrence These deposits cover the central bottom of lakes, which implies that they have lateral contacts only with proglacial lake-margin deposits, occasionally with proglacial deltaic bottomsets, and exceptionally with proglacial stream deposits. All these contacts are gradual if no specific conditions existed. The base of these deposits is sharp if the lacustrine deposits rest directly on deposits of another facies; there may also be a gradual change in an upward direction, from proglacial lake-margin deposits into lacustrine bottomsets, e.g., if the lake grew and deepened gradually. As mentioned earlier, the shifting - and somewhat arbitrary - boundaries between the terminoglacial and proglacial subenvironments can also

Deposits of the proglacial lacustrine facies

42 1

result in immediate - though invisible - contacts between terminoglacial and proglacial lacustrine bottomsets. The position of these deposits in the glacigenic sequence is, in fact, determined by the position of the lake-margin deposits. The combination of these two types of deposits can be found throughout the interval with proglacial sediments, but always in the vicinity of proglacial stream deposits: fluvial currents must supply the particles t h a t settle in the centre of the lake. Part of these deposits will be found in the glacigenic sequence at the boundary between the proglacial and the extraglacial subenvironment. Depositional mechanisms The various sedimentation processes involved in the genesis of these deposits have been studied in detail and it has become clear that a variety of processes was involved, although the resulting sediments do show similarities. It was pointed out in the section on terminoglacial lacustrine bottomsets that the graded, widespread units have a partly turbiditic origin and are partly due to settling from suspension, with other processes having only a minor role (cf. De Geer, 1922; Johnston, 1922; Kindle, 1930; Kuenen, 1951; Gravenor, 1956; Ludlam, 1967; Campbell, 1968; Sturm, 1979; Sturm and Matter, 1978; Lambert and Hsu, 1979; Smith and Syvitski, 1982; Syvitski et al., 1987; Pickrill and Irwin, 1983). The theory proposed regarding turbidite fans (Mutti and Ghibaudo, 1972; Mutti and Ricci-Lucchi, 1972; Walker and Mutti, 1973; Rupke, 1975; Ricci-Lucchi, 1975; Walker, 1977,1978,1980) has been of great help with the understanding of the depositional context of turbidites in proglacial lakes (Ashley, 1975; Shaw, 1977b; Merta, 1978; Brodzikowski and Van Loon, 1983,1987; N. Eyles, 198313). Some of the processes involved in turbidity currents may be interpreted as types of intraflood transistent hydraulic surging ( a term introduced by Bretz, 1929). Abrupt rises in the local water level are propagated downcurrent along floodways and upcurrent via tributaries (Bretz et al., 1956; Bretz, 1969; Baker, 1973; Webster et al., 1976; Carson et al., 1978; Patton et al., 1979). These surges have only a n areally restricted influence (Waitt, 1980), which is the main difference from many turbidity currents. The deposits formed by intraflood transistent hydraulic surges resemble in other aspects the turbidites consisting of the B-E o r C-E intervals of the Bouma (1962) sequence, which implies that such surges may play a more important role in t h e deposition of proglacial lacustrine bottomsets than was hitherto assumed.

422

The proglacial subenvironment

Over- and interflows are very important because they may carry large amounts of debris to the lacustrine facies (Agterberg and Banerjee, 1969; Shaw, 1977; Sturm and Matter, 1978; Pharo and Carmack, 1979; Sturm, 1979; Pickrill and Irwin, 1982,1983).Large floods are also good suppliers of material that is deposited mainly by traction currents that spread out over the lake bottom rather than by turbidity currents (Merta, 1978; Shaw and Kvill, 1984; C.H. Eyles, 1986,1987; Visser et al., 1987; Shaw et al., 1989). Low-energy traction currents account for the genesis of micro-scale stratification (Fig. 257). This may occur as isolated current ripples of fine sand or silt 2-5 mm high. Low-angle cross-bedded units of the same size may also be present. If the lake bottom is covered with very cohesive material, low-energy traction currents cannot pick up and transport grains as bed load. While a current velocity of 4 cm s-l may be sufficient in the case of a sandy subsoil, a velocity of 45-60 ems-' is required if the bottom sediment is very cohesive (Orheim and Elverhoi, 1981). Currents with such a velocity can easily affect the bottom, particularly if objects are carried along, and structures such as grooves, tool marks, rill marks

Fig. 257. Micro-scale stratification in proglacial lacustrine bottomsets of Wartanian age (Kleszczow graben, central Poland).

Deposits of the proglacial aeolian facies

423

and scour-and-fill structures are then found in proglacial lacustrine bottomsets. It appears that the bottom currents do not have a continuous thickness and width; changes in these parameters affect current velocity. These changes explain the fact that local erosion may occur (scouring), while deposition (current ripples) takes place elsewhere.

DEPOSITS FROM THE PROGLACIAL AEOLIAN FACIES (11-B-5) Aeolian deposits are found much more frequently in the proglacial than in the terminoglacial subenvironment. In spite of this they constitute a much smaller fraction of the glacigenic sediments here than they do in the extraglacial subenvironment. It is remarkable that there are abundant observations of aeolian deposits in modern areas (Bird, 1951; Bout et al., 1955; Webb and McKelevey, 1959; Gilbert, 1983),particularly reports by geomorphologists, whereas pre-Pleistocene aeolian deposits from the proglacial subenvironment have not been recognised with much certainty. It must be concluded that the preservational potential of these deposits is extremely low (Fig. 258).

Fig. 258. Wind blown fine-grained cover of only 1-2 cm thick in the proglacial zone of Greenland. Rain carries most of such an aeolian cover away, changing it into fluvial material. Photograph J. Cegla.

424

The proglacial subenvironment

Proglacial aeolian deposits are not very important quantitatively. Most are thin (except local dunes) and have a restricted areal extent. Five types of deposits can be distinguished in this facies: proglacial aeolian complexes (II-B-5-a), proglacial drift sands (11-B-5-b),proglacial dunes (II-B-Ei-c),proglacial coversands (II-B-5-d) and proglacial loesses (11-B-5-e).

Proglacial aeolian complexes (II-B-5-a) Aeolian sedimentation in this subenvironment often takes place as a result of local winds that pick up material, particularly from alluvial plains, transport i t over short distances then let i t settle again. Local conditions often change so quickly that no regular aeolian deposits are formed. Wind-blown material is deposited, again eroded, redeposited, etc. This alternation of aeolian processes is very easily observable under modern circumstances, where accumulation and erosion together a r e responsible for a continuously changing topography (Fig. 259). The deposits formed during the successive stages of erosion and sedimentation sometimes have strongly varying characteristics. I t is therefore considered appropriate to distinguish them as aeolian complexes.

Fig. 259. Deflation zone i n t h e proglacial zone of Hornsund (Svalbard). A stone pavement results from the aeolian erosion of fines. The stones thus exposed act as barriers behind which wind-blown material accumulates. Photograph: J. Kida.

Deposits of the proglacial aeolian facies

425

Lithofacies characteristics The complexes consist of a mixture of one or more types of more well defined proglacial aeolian deposits. The lithology, rather than being specific, shows a variety of different characteristics (Fig. 260). Largescale cross-bedding and horizontal lamination are the most common structures, but the nature of the depositional site may also result in partial reworking by surficial waters, which results in the intercalation of levels with typical fluvial current ripples. Many of the units are usually partly truncated as a result of deflation phases. It is common in modern situations that the sedimentary surface shows abundant aeolian structures: wind ripples, lee-side accumulations, etc. Such structures are hard to find in unlithified deposits (and almost all Pleistocene deposits are unlithified), whereas these complexes are not known from older glaciations. Most material consists of fine-grained sand and silt, but both coarser and finer units may be present. The grain size within one specific layer shows much less variation.

Fig. 260. Proglacial aeolian complex (base at spade base) of possibly Wartanian age in a sand pit near Zary (western Poland). Note the variability in cross-bedding and the discontinuities. The top part has been affected by solifluction during the latest Saalian and the Vistulian.

426

The proglacial subenvironment

Textural characteristics One might expect that the aeolian character of the surface of individual grains would be a distinct textural characteristic. This is seen only in exceptional cases: transport is commonly so short that almost no aeolian texture results. Individual grains in the sediments may have undergone several phases of aeolisation; these grains may show the characteristic aeolian surface, which is especially well visible under the microscope. Sorting is usually bad in comparison with other aeolian deposits, but is must be kept in mind that the complexes in fact represent mixtures of different deposits. As a rule, sorting is found to be much better if considered only within one 'subdeposit'. Material of pebble size is not known as a constituent of these complexes. Occurrence The complexes are found distributed throughout the proglacial subenvironment, but with relatively high concentrations in the proglacial alluvial plains. They are usually of restricted thickness (up t o a few metres) and a comparatively very limited areal extent (often much less than 100 m). Most complexes rest on proglacial fluvial sediments (the base is usually very sharp) and are overlain by the same type of sediment, sometimes also by proglacial subaerial mass-transport deposits. Deposits, commonly of an aeolian nature, formed in the extraglacial subenvironment may also be found on top of the complexes. Lateral transitions occur with predominantly fluvial deposits, but also with lacustrine and mass-transport deposits and, t o a lesser degree, with deltaic deposits. The following positioning of the complexes in the glacigenic sequence is thus clear: throughout the proglacial interval, usually in association with proglacial fluvial deposits. Depositional mechanism There are three main processes rather than one specific depositional mechanism. The most characteristic process is settling from particles that were suspended in the air. Settling is a result of decreasing wind velocity. Only fine (sometimes medium) sand and finer material play a clear role in this process.

Deposits of the proglacial aeolian facies

427

The second process is settling of particles that were transported by the wind, largely in saltation. While sand is the major component in this case, finer grains may also be involved. Particle-laden wind transport over an unlithified subsoil is usually an erosive process, contributing t o the concentration of moving particles. When wind velocity decreases, a corresponding part of the particles settles. The third process, which contributes much less to aeolian deposition, is that of settling on the lee side of an obstacle. This obstacle may be a stone or a geomorphological form (a bar), but may also be a concentration of wind-blown material. Dunes can thus act as barriers behind which aeolian particles settle.

Proglacial drift sands (11-B-5-b) Drift sands are irregularly distributed, often very thin and areally limited bodies (Fig. 261). They may be found at relatively elevated sites of alluvial plains, but various reports show that they are much more

Fig. 261. Irregular drift sands in SW Svalbard. The depressions in the piedmont area are relatively wet, so that grains supplied by the wind are easily caught, then to form small accumulations. Photograph:J. Ceg+a.

428

The proglacial subenvironment

abundant outside the glacigenic environments, viz. i n the t u n d r a s (Bogacki, 1970; Hine and Boothroyd, 1978; Akerman, 1980; Ceg$a, 1982; Szczypek, 1982; Gilbert, 1983; Good and Bryant, 1985). Most of the research into these deposits was done by geomorphologists so that there are no detailed sedimentological descriptions available (Pewe, 1975; Akerman, 1980; Kida, 1981,1986; Szczypek, 1982).

Lithofacies characteristics These deposits consist predominantly of medium and fine sand. There occur silt admixtures, usually as a result of the supply of silt in the form of pellets that were transported in the same way as sand grains (Ceg+a, 1982). The deposits may be massive, horizontally laminated or crossbedded (Cegta, 1982; Good and Bryant, 1985; Kida, 1986), especially if the component units reach thicknesses of the dm-scale (Fig. 262). While the structure of drift sands in the proglacial subenvironment has not been analysed in detail, there are more general theories on aeolian deposition in cold areas (Ahlbrandt and Andrews, 1978; Kocurek and Dott, 1981; Ahlbrandt and Fryberger, 1982; Ruegg, 198310; Schwan, 1986; Fryberger et al., 1988) that indicate that both 'wet' and 'dry' stratification types are probably present (cf. Lindquist, 1988) as indistinct hori-

Fig. 262. Possibly proglacial drift sands of Wartanian age from a sand pit near mogow (SW Poland). Note the irregular cross-stratification.

Deposits of the proglacial aeolian facies

429

zontal parallel lamination and low-angle cross-lamination (Fryberger et al., 1979; Good and Bryant, 1985; Schwan, 1986). Geomorphological studies indicate that adhesion structures are abundant a t the surface of these drift sands (Bogacki, 1970; Szczypek, 1982; Kida, 1986), particularly in areas where there is a surficial drainage system or where the topsoil is wet but not fully saturated (Kocurek and Felder, 1982). Asymmetric adhesion structures, usually with a pronounced shadow tail and a downwind dipping surface seem the most common type. Accretion by avalanching over the dipping surface may account for the cross-stratification observed in thicker units. Massive units or deposits with a very vague lamination are however the more common finding (cf. Schwan, 1986; Olsen et al., 1989). Textural characteristics

Sorting depends on the local conditions but usually varies from very bad to moderately good. The surface of the grains is commonly still fairly rough. Very angular grains (broken by glacial grinding or during glaciofluvial transport) are common, but well rounded specimens, possibly the result of several phases of aeolisation, are also found. Abrasion by wind usually left few marks as short transport distances occur more frequently than a long-lasting stay in the air. Gozdzik (1979, 1981, 1986) has suggested that the badly developed aeolian characteristics of the grain surface are such a routine feature that it is the main reason why these drift sands were not recognised during microscopic analysis of 'fossil' deposits. This makes these deposits very different from aeolian extraglacial and - more specifically - desert sands as far as textural characteristics are concerned (Bagnold, 1941; Friedman, 1961; Hand, 1967; Kozarski et al., 1969; Glennie, 1970, 1972; French, 1976; Ahlbrandt, 1979; Ahlbrandt and Fryberger, 1982; Koster, 1982; Hofle, 1983; Mader, 1983; Schwan, 1987,1988). It can therefore be concluded that detailed lithological investigation of proglacial drift sands is still badly needed. Occurrence

Proglacial drift sands are usually local phenomena, which implies that they occur most frequently along proglacial streams and on 'islands' in these streams. They are also found, however, in belts around dunes or coversands. They form flat, thin, small-scale bodies of irregular shape. Their thickness rarely exceeds 0.5 m and is usually in the range of

430

The proglacial subenvironment

millimetres t o centimetres; discontinuity is therefore the rule rather than the exception (Fig. 263). The transport distance is often so short that the drift sands are found immediately next t o the deflation area where the material originated; a separating distance of more than a kilometre is rarely seen. It is remarkable that the deflation zones have a much smaller surface area than the genetically related drift sands; this explains the thin development of the latter. Akerman (1980)has reported deflation pits of a depth up to several decimetres. The preservational potential of these deposits is very low. They therefore rarely appear in the geological record as they are eroded much faster than they accumulate (Loope, 1985). However, if they are positioned in the glacigenic sequence, they are found in association with pmglacial fluvial and aeolian deposits throughout the proglacial interval of the sequence.

Fig. 263. Characteristic distribution of proglacial drift sands on Svalbard (Hornsund area): they form discontinuous elongated stripes only a few millimetres to a some centimetres thick. Photograph: J. Cegta.

Deposits of the proglacial aeolian facies

431

Depositional mechanisms Proglacial sand drifts do not form easily. They need a somewhat wet substratum or surface obstacles such as boulders, escarpments or other topographic irregularities (Cegla, 1982; Szczypek, 1982; Kida, 1986. Also see: Chepil, 1945; Cegfa, 1972; Hunter, 1973,197710,1980b;French, 1976; Clemmensen, 1978; Sarnthein,l978; Karte, 1979; Walker and Middleton, 1979; Washburn, 1979; Kocurek, 1981; Hummel and Kocurek, 1984; Schwan, 1986,1988). The grains are usually transported over or near the surface; transport as a low suspension, just above the surface, usually dominates because a somewhat wet substratum makes transport as traction or in saltation hard t o take place. Transport takes place within a short time, mainly during short pulses of increased wind velocity. Deposition is thus also pulsational and commonly takes place very rapidly. There is only limited interaction between the grains during transport as only small quantities of sediment are usually involved; this explains the lack of aeolian texture a t the grain surfaces.

Proglacial dunes (11-B-5-c) Dunes are common features in the proglacial subenvironment (Fig. 264) and are especially found in wide valleys with braided streams. They have been reported frequently from Alaska (Pewe, 1955) where they are still being formed in the Yakutat area (Miller, 19611, the delta of the Cooper river (Tar and Martin, 1914), in the area around Bristol Bay (Black, 1951), the Seward peninsula (Pewe, 1975) and in the north (Black, 1951). They occur abundantly along sandy coasts, but the largest areas with active Alaskan dunes are in the Koyukuk and Kobuk river valleys. The area in the Koyukuk river valley is only a small fragment of a much larger Pleistocene dune field now inactive in most places (Fig. 265). The active part covers approx. 65 km 2 and comprises mainly transverse dunes of up t o 60 m high and about 100 m long. The inactive part of the dunefield is twice this size; some isolated active dunes occur (Fernald and Nichols, 1953;Fernald, 1964).The dunes along the sea shores are mainly transverse foredunes and cliff-head dunes (Miller, 1961; Pewe, 1975). The Pleistocene dunes in Alaska are largely of proglacial origin. There is an area of some 780 km2 of poor to fairly well formed sand dunes against hard-rock hills, up to an elevation of 520 m, north of the Tanana river (near Big Delta, approx. 145 km southeast of Fairbanks). These dunes of Illinoian age have a parabolic ghape and a relief of 0.3-3 m that is ,/

432

The proglacial subenvironment

Fig. 264. Proglacial inland dunes in West Greenland. The dunes cover about 2 km2 and reach heights of maximally 3 m. Photograph:J. Cegla.

still undergoing change. They were formed as a result of sand supply by winds that blew northward from the widespread outwash plains and glacial floodplains of the Delta and Tanana rivers (Pewe, 1965, 1975). Some of the dunes are now covered by a loess layer 0.3-2 m thick, which is not directly genetically related. The largest dune fields in Alaska extend from the Tanana river southward over a distance of 160 km, into the drainage area of the upper Kuskowin river. The dunes are found on high terraces of the Tanana river and its tributaries, 100-300 m above sea level. This dune area, with a surface of some 15,500 km2, was mapped by Collins (unpublished) and P6we et al., 1966). The individual dunes are well shaped and have local height differences of up to 60 m. The dunes are now stabilised as they are covered by loess, preserving their dominantly longitudinal character, but parabolic forms are also present. Other dunes in Alaska date from Wisconsin and post-Wisconsin times. They have been described by Black (1951), Livingstone (19541, Rosenfeld and Hussey (1958), Fernald (1960, 1965), Williams (1962), Rickert and Tedrow (1967) and H.J. Walker (1967).

Deposits of the proglacial aeolian facies

168"

160"

162'

144'

433

136'

70"

7 0"

66"

$6'

62"

61"

58"

58"

-1 1GO"

.I-

1 162-

I

I 144-

Fig. 265. Major areas of stabilised s a n d d u n e s i n Alaska. D = small areas of sand dunes. From: Pewe (19751, compiled by Weber (1961), p a r t l y from published and partly from u n p u b l i s h e d sources. Courtesy: U S Geological Survey.

Dunes are also well known to occur in the Antarctic. Those in Victoria Valley are some 13 m high and form a desert erg (Shaw and Healy, 1981) comprising sand sheets, barchans, transverse dunes and whaleback dunes or mantles (Webb and McKelvey, 1959; Calkin and Rutford, 1974; Rains and Palmer, 1974; Washburn, 1979). The size of the wind-transported grains involved is exceptional: granules of 3 mm in diameter have been reported to have hit a tent 2 m above its base, and pebbles 19 mm in diameter have been collected from dune-ripple crests i n the Victoria Valley (Washburn, 1979). Calculations support the view of H.T.U. Smith (1966) that the high density of cold Antarctic winds is a n important parameter for the transport of coarse particles (Selby et al., 1974). Dunes have also been reported from the proglacial subenvironment in Greenland (Hobbs, 1931; Teichert, 1935; Belknap, 1941; Flint, 1948;

434

The proglacial subenvironment

Troelsen, 1952; Fristrup, 1953; Nichols, 1969). Geomorphological studies carried out by Nichols (1969) in the Inglefield land (North Greenland) indicate that dunes a few metres high occur along the rivers and coast. Fristrup (1953) reported finding larger dunes around the Sondre Stromfjord. An abundance of data regarding proglacial dunes was obtained from Arctic Canada (Mackay, 1963; Pissart, 1966; Bird, 1967; French, 1972; Pissart et al., 1977; R. Gilbert, 1980, 1982a, 1983; Good and Bryant, 1985), from Iceland (Bout, 1953; Bout et al., 1955; King, 1956; Ashwell, 1966,1972;Bogacki, 1970; Shunke, 1975; Hine and Boothroyd, 1978) and from Svalbard (Akerman, 1980; Szczypek, 1982; Kida, 1986). These studies yielded detailed data on both morphology and aeolian processes. H.T.U. Smith (1964) listed several criteria for the identification of periglacial aeolian sands, but made no distinction between the three subenvironments as the authors do. Cailleux (1973) cited three criteria that also apply to periglacial conditions, which he considers to be uniquely applicable t o aeolian transport: wind-induced gliding of stones (2-6 cm in diameter) over icy surfaces, interstratification of sand and snow, resulting in undulatory bedding, and filling of frost craks by wind-blown sands of similar age.

Lithofacies characteristics The precise lithofacies characteristics depend on the type of dune involved. Transverse dunes being the most common i n the proglacial subenvironment, will be described here most in detail. Parabolic dunes, barchans and longitudinal dunes are less common, although they may dominate in specific places. Five main types of aeolian strata can be found in the dunes, exactly as in non-glacigenic dunes. These types are: avalanche-produced sand-flow units, layers formed by grainfall due t o flow separation on the leeside of the dune or due t o saltation between small obstacles, units with crossbedding of ripple origin, strata resulting from adhesion processes, and plane-bed lamination formed as a result of very strong winds blowing over unrippled surfaces (Fryberger et al., 1983). The characteristics of these strata will be dealt with in more detail later in the relevant section in the chapter on extraglacial deposits. This order is followed as aeolian deposits are most characteristic of extraglacial conditions. Only some general aspects will be dealt with here. The most common types of strata in large dunes are those formed by sand flow and grain flow. These processes result in large-scale cosets with

Deposits of the proglacial aeolian facies

435

a distinct downwind dip. There is often found superimposed cross-bedding that results from wind ripples migrating over the dune foresets (Rubin and Hunter, 1983). Most units are tabular, although there may may also some trough-bedded sets. The latter are produced under conditions of changing wind direction, whereas the existence of tabular sets points t o fairly stable directions. The sediments in the interdune areas comprise sands with abundant wind ripples and adhesion structures. These structures are also commonly found in combination with horizontal lamination at the base of the dune bodies. The dunes themselves fairly frequently show a succession from bottom t o top with a base dominated by adhesion structures, which is followed by a unit with wind ripples that sometimes have a climbing character then follow grainfall strata with grainflow units ending the succession. The gradual succession may be interrupted by transverse scour structures resulting in the transverse migration of some material. Textural characteristics

While the dunes show good sorting, it is less good than that in dunes of desert areas or the extraglacial subenvironment. Most particles consist of rounded t o well rounded quartz grains and the larger the dune field, the better the rounding. The combination of good sorting and a high degree of rounding results in high porosity (Lindquist, 1988). Most of the grains in dunes show characteristic aeolian grain surfaces (Ahlbrandt, 1975,1979; Ahlbrandt and Andrews, 1978). Interdune areas commonly show a bimodal grain-size distribution. Fine and medium sand dominate, with an admixture of coarse sand. Akerman (1980) reported that small dunes in Svalbard also contained relatively coarse material. These dunes are exceptional in that they contain a very low percentage of fines. Akerman (1980) concluded on the basis of experiments that wind-blown ice particles and snow flakes, even more than mineral particles, make a large contribution to aeolian abrasion (Fig. 266). Occurrence

The dunes in this subenvironment are found most frequently in the 'distal' part, near the extraglacial subenvironment, particularly in fluvial plains and on river terraces. Lakes with frequently changing water levels are also good sources of sand, so that dunes may also surround such lakes. Gilbert (1983) described dunes from Arctic Canadian fjords.

436

The proglacial subenvironment

Fig. 266. Niveo-aeolian deposits in a snow cover in front of the Werenskiold glacier (Hornsund area, Svalbard). The combined transport of snow flakes, ice crystals and mineral grains results in strong abrasion of the sand grains. Photograph: J. Kida.

Dune fields that 'start' in the proglacial subenvironment may well extend to the tundra region. The absence of high vegetation facilitates dune formation and continuous dune reshaping. Dunes are obviously also formed along sandy sea shores, as in, e.g., Iceland. Hine and Boothroyd (1978)described large dune fields from the distal parts of tributaries of the Skeidararsander. Proglacial dunes are found in the glacigenic sequence, particularly in the transition zone t o extraglacial deposits. They interfinger with proglacial fluvial deposits via sharp boundaries. Lateral contacts with other proglacial deposits, mainly of fluvial and of subaerial mass-transport origin, also occur but are less frequent. Gradual transitions into proglacial coversands and loesses are also seen. Depositional mechanism The stability of the wind direction and the areal distribution of dry and wet surfaces play an important role in the formation of dunes. Detailed

Deposits of the proglacial aeolian facies

437

descriptions abound of the precise depositional mechanisms of dunes, including the processes resulting in primary and secondary structures (Stokes, 1964; McKee, 1966, 1972, 1979a,b,c; Glennie and Evamy, 1968; Glennie, 1970; Folk, 1971; I.G. Wilson, 1971, 1972a,b, 1973; McKee and Bigarella, 1972; Hunter, 1973, 1977b, 1980b, 1981; Bigarella, 1975; Brookfield, 1977, 1979; Ahlbrandt and Andrews, 1978; Ahlbrandt et al., 1978; Clemmensen, 1978; Fryberger and Ahlbrandt, 1979; Fryberger and Dean, 1979; Doe and Dott, 1980; Ahlbrandt and Fryberger, 1981, 1982; Fryberger and Shenk, 1981; Kocurek, 1981a,b, 1986,1988; Kocurek and Dott, 1981; Kocurek and Fielder, 1982; Fryberger et al., 1983, 1988; Hunter and Rubin, 1983; Rubin and Hunter, 1983; Hummel and Kocurek, 1984; Kocurek and Hunter, 1986; Kocurek and Nielson, 1986; Rubin, 1987; Hunter and Richmond, 1988; Kerr and Dott, 1988; Olsen et al., 1989). The formation of dunes requires both wind activity and the presence of obstacles. Wind-blown sand can then settle on either the lee or the stoss side (Fig. 267); larger dunes formed in this way have been termed 'shadow dunes' by Gunatilaka and Mwango (1989). At the outset, the dunes consist of massive sand on the stoss side and usually show grainfall stratification on the lee side, Once a dune has fully covered the 'core' obstacle, it can start to migrate. The foresets usually consist of large-scale tabular cross-bedding composed of grainflow strata, especially in the lower part (Hunter, 1981). Climbing ripples appear if flow separation is sufficiently developed. Large dunes are therefore often covered by a layer with abundant climbing ripples. The interdune areas are much less complicated. A wide variety of primary structures is found here, partly influenced by surficial water currents. These areas tend to be relatively wet so that adhesion lamination and adhesion ripples are common. Some material in these areas may be of niveo-aeolian origin.

Proglacial coversands (II-B-5-d) Proglacial coversands (Fig. 268) have often been described (Samuelson, 1921; Bird, 1951, 1967; Bout et al., 1955; Webb and McKelevey, 1959; French, 1972a; Selby et al., 1974; Pkiwe, 1975; Pissart et al., 1977), but details of the sedimentology are sparse. The most detailed studies in this respect are those of Hine and Boothroyd (1978) for coversands in Iceland, and of Good and Bryant (1985) for coversands from Arctic Canada. The preservational potential of these deposits appears to be extremely low, which explains their absence from 'fossil', hard-rock glacigenic

438

The proglacial subenvironment

Fig. 267. Accumulation of wind-blown sand on the stoss side of a n artificial obstacle (top) and on the lee side of stones that form part of a deflation pavement (bottom). Beach in the proglacial subenvironment of Svalbard. Photographs: J. Kida.

Deposits of the proglacial aeolian facies

439

Fig. 268. End of a proglacial coversand unit in western Greenland that reaches a thickness of 1.5 m. Photograph: J. Ceg€a.

sediments. P6we (1975) described a unit of sand sheets of Pleistocene age from Alaska, which probably represent coversands. A cap of loess layers has preserved these sands. Other proglacial sands of wind-blown origin, possibly also representing coversands, were found in the surroundings of Opole (South Poland). The sands are not purely aeolian, however, as surficial currents have left traces as current ripples (Fig. 269).

Lithofacies characteristics The proglacial sand flats described by Good and Bryant consist of medium t o fine sand. The sands are in part massive and also show some climbing stratification. Adhesion lamination is seen in the basal zone and aeolian ripples are present in the thickest parts of the still active sand sheets. Ephemeral streams are incised locally in the sands and become filled with cross-bedded aeolian sands during dry periods (Fig. 270). These characteristics agree well with the description by Hine and Boothroyd (1978). The large, shallow trough cross-stratification found in these Icelandic sand sheets may represent eroded dunes.

The proglacial subenvironment

440

Fig. 269. Coversand of Vistulian age, partly reworked by surficial currents, from the Zielona Gora region (western Poland). The 'large-scale' ripples are of a water-current origin, whereas small-scale ripples (height approx. 2 mm) are of aeolian origin. Photograph: J. Cegfa.

I

Accretion

Deflation

I

Damp sond deflalion surface I adhesion

Aeolian ripples

Sand -sheet of dry mobile sand 1 structureless or with climbing translatent stratification I

Variable wind velocity and d irec t ion

stream channels Planar bounding surfaces between units

Buried ventifacts on former deflation surfaces

Fig. 270. Characteristics of the proglacial aeolian sand flats in Arctic Canada. From: Good and Bryant (1985). Courtesy: Geografiska Annaler.

Textural characteristics None of the textural characteristics of these deposits have been described in any detail. The only information available is that the aeolian sands in

Deposits of the proglacial aeolian facies

441

the proglacial subenvironment usually do not show good sorting (if compared to the extraglacial subenvironment) and that most grains have a relatively low aeolian impact on their surfaces. Both aspects are a result of the relatively short aeolian transport distances. Fine and medium sand prevails. Silt and clay are usually absent, being transported further, and coarse sand and gravel may be present at the base in the form of a deflation pavement. Occurrence Most of these coversands are found in low-relief alluvial plains and in wide marginal valleys. They may also be found together with dunes in the close vicinity of sandy sea or lake shores. The source-related nature of the coversands is reflected in their extent: they often form relatively narrow, elongated bodies from the source area in the (prevailing) downwind direction. The deposits are usually about a few metres and sometimes less than one metre thick (Hine and Boothroyd, 1978; Good and Bryant, 1985). Parts of these deposits may be of fluvial origin as a result of surficial drainage. The basal contacts of the coversands are usually sharp, unless they rest on other aeolian deposits. The bodies often form an intermediate deposit between proglacial aeolian complexes and proglacial dunes, although they may also surround proglacial dunes on all sides. Boothroyd and Nummedal (1978) and Good and Bryant (1985) consider the aeolian sand flats as a 'lateral facies' as compared to the 'core facies' represented by glaciofluvial deposits. These two facies are very closely connected in this subenvironment. Because of the lack of 'fossil' deposits, the position of the proglacial coversands in the glacigenic sequence can only be derived at theoretically. Had the coversands been preserved, one would expect t o find them in association with proglacial dunes and proglacial fluvial deposits. Depositional mechanisms Only Good and Bryant (1985) have analysed the depositional mechanism of proglacial coversands in detail. They concluded on the basis of their research in the Sachs river plain that thin, low-angle coversands must be formed by the process described by Fryberger et al. (1979),who attributed the sedimentation on the lee side of obstacles to 'topographic deceleration'. According to Fryberger et al. (1979), the sediments thus deposited form a transition between dunes and non-aeolian deposits. However, the

442

The proglacial subenvironment

reports of Good and Bryant (1985) on Banks Island show a relationship only with fluvial sediments and not with high-angle dunes. Good and Bryant (1985) explain this by the presence on Banks Island of a permafrost and a seasonally thawed layer, which would inhibit dune formation in this area. It must be realised, however, that dunes occur frequently elsewhere in the world, including permafrosted high-latitude areas. It must therefore be concluded that factors other than those suggested by Good and Bryant (1985) also play a role, or that these authors' interpretation (the same process as described by Fryberger et al., 1979) is not entirely correct. Damp surfaces have an important effect on aeolian sedimentation, i.e. the formation of adhesion structures. The precise type and the internal stratification depend on the moisture content of the sandy substratum. Sufficient moisture slows down ripple migration so that climbing ripples are formed relatively easily. However, continued ripple migration and grain saltation were observed in the Banks Island coversands during 'adhesion plane-bed conditions' (cf. Kocurek and Felder, 1982).A parallel lamination is then formed, but whether such laminations result from water currents or from aeolian processes (Hunter, 197713, 1980b) is difficult t o reconstruct.

Fig. 271. Computer-contoured map of part of the sand flat south of the Sachs river (Canada). Width of area shown approx. 250 m. Contour intervals: 10 mm. Dotted areas: net accretion of > 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

The proglacial subenvironment

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

The proglacial subenvironment

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

Deposits of the proglacial mass-transport facies

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

The proglacial subenvironment

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.

Deposits of the proglacial mass-transport facies

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

The proglacial subenvironment

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

Deposits of the proglacial mass-transport facies

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

The proglacial subenvironment

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).

Deposits of the proglacial mass-transport facies

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

The proglacial subenvironment

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.

Deposits of the proglacial mass-transport facies

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

The proglacial subenvironment

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

The extraglacial subenvironment

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

The extraglacial subenvironment

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

The extraglacial subenvironment

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

The extraglacial subenvironment

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.

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The extraglacial subenvironment

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-

Deposits of the extraglacial aeolian facies

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.

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The extraglacial subenvironment

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.

Deposits of the extraglacial aeolian facies

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

The extraglacial subenvironment

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.

Deposits of the extraglacial aeolian facies

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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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The extraglacial subenvironment

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).

Deposits of the extraglacial aeolian facies

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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The extraglacial subenvironment

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

Deposits of the extraglacial aeolian facies

477

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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The extraglacial subenvironment

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-

Deposits of the extraglacial aeolian facies

479

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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The extraglacial subenvironment

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

Deposits of the extraglacial aeolian facies

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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The extraglacial subenvironment

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;

Deposits of the extraglacial aeolian facies

483

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

484

The extraglacial subenvironment

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

Deposits of the extraglacial aeolian facies

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

486

The extraglacial subenvironment

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.

Deposits of the extraglacial aeolian facies

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).

488

The extraglacial subenvironment

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

Deposits of the extraglacial aeolian facies

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

The extraglacial subenvironment

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

The extraglacial subenvironment

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.

Deposits of the extraglacial mass-transport facies

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

The extraglacial subenvironment

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

Deposits of the extraglacial mass-transport facies

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

The extraglacial subenvironment

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

Deposits of the extraglacial mass-transport facies

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

The extraglacial subenvironment

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

The marine terminoglacial subenvironment

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.

The marine terminoglacial subenvironment

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

The marine terminoglacial subenvironment

(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

The marine terminoglacial subenvironment

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


Molnia 119791

2-5cm 4cm

F

c5

Ma

mudldiamlct

'Llcm

F

modern

mudidiamict


Arctic Ocean lCanadian1

F T

<5 26 Ma < 0 6 Ma

mud!d#amict

Pacific Ocean fAlarkan1

mudidiamicf

) 5 - 1 14mm 40mm

T

tO6Ma

mudlrand

Rtchardr 119761

Orheim and Elverhot 119811 Elverhoi and Rosldret (19831 Earrett 119751 Earrett 119751

Oiheim and Elverhoi 119811 Clarkeral. 119801 Plpeieta!. 119731

.TRENCH

Aleutian

3 5m

Piper et a ! , I19731 Von Hurne and Kulm (19731

Fig. 299. Representative deposition rates from suspension and ice rafting in various glaciomarine settings. From: C.H. Eyles et al. (1985). Courtesy: Elsevier Science Publishers. T = thawed base glacier; F = frozen.

Lithofacies characteristics Ice-raft deposits formed in the marine terminoglacial subenvironment are commonly massive, matrix-supported diamicts (Ledbetter and Watkins, 1978). Clast-supported diamict lenses may be present, but are rare phenomena (Miall, 1983; Eyles and Eyles, 1984). The grain-size distribution and the maximum sizes of the particles depend only on the characteristics of the material that was embedded in the ice, i.e. mainly on the geology of the substratum under the continental ice cap and on the erosional forces exerted by the ice. The successive stages of sedimentation, i.e. those after icebergs have passed, and their combination with the different rates of settling, result in a more or less homogeneous sediment without distinct stratification. The most common structures found are deformation structures, results of, e.g., the impact of clasts on the soft seabottom. Typical load casts may also develop if coarse material (or a single large clast) is too heavy for the water-saturated substratum to be kept in place (Ovenshine, 1970; Powell,

506

The marine terminoglacial subenvironment

1984; McCabe et al., 1987). The sediments may also be deformed by ice grounding in the proximal zone of this subenvironment. Other deformation structures, e.g., connected with thickness gradients, are rarely seen. Investigations of current-day ice-raft deposits of the Antarctic shelf (Ross Sea) based on analyses of piston cores (Chriss and Frakes, 1972; Anderson et al., 1984) indicate that these deposits account for only some 20% of all sediments exposed a t or near the seabottom. This implies t h a t ice-raft deposits do not dominate, in spite of the frequently high sedimentation rate (ablation of the active ice is usually strong under marine conditions, so that the ice flows relatively rapidly into the sea, with frequent calving of debris-laden icebergs). The most important reason is the random sedimentation of ice-raft deposits (including dumpstones), so t h a t subaqueous slopes are formed that favour partial reworking of the sediment by subaqueous mass-transport processes. Sediments formed a t the seabottom may also be redistributed by bottom currents of nonglacigenic origin.

Fig. 300. Icebergs in the marine terminoglacial subenvironment of the Kongsfjord (Svalbard).Photograph: J.D. de Jong.

Deposits of the marine terminoglacial melting-ice facies

507

Textural characteristics Marine terminoglacial ice-raft depasits are typical diamicts, with a mixture of grain sizes ranging from clay t o boulders (Fig. 301). The nearubiguitous presence of glacial striae on the larger clasts is good proof of their glacial character (Visser and Loock, 1987). Most of the deposits have angular to subangular clasts, but any rounded and facetted clasts present in the source area will at least partly keep their shapes. Hard-rock material interpreted as marine terminoglacial ice-raft deposits sometimes shows zones with a faint gradingsuch graded units apparently occur as a kind of elongated, irregularly shaped lenses within the generally massive and structureless deposits. Sorting of the diamicts is, of course, bad; only a source area with a narrow grain-size distribution will result in more or less sorted ice-raft deposits (cf. Benn and Dawson, 1987).

Fig. 301. Marine terminoglacial ice-raft deposit, consisting predominantly of large boulders (Brazeau sandpit, South of Ottawa, Ontario, Canada).

508

The marine terminoglacial subenvironment

Occurrence It is the topography of the seabottom that largely determines the characteristics of the marine terminoglacial ice-raft deposits. Gentle slopes and wide shelves give rise t o widely extending deposits although these tend not t o be present near tunnel mouths (Anderson et al., 1984; McCabe et al., 1984, 1987). More inclined bottoms commonly result in deposition of ice-rafted material in the distal zones of this subenvironment; Miall (1983a) found the same for Huronian tillites of the Gowganda Formation. Vertical sections show frequent intercalations of ice-raft deposits and subaqueous mass-transport deposits. In a lateral direction, the ice-raft deposits gradually interfinger with mass-transport deposits, with marine glacial counterparts and with non-glacigenic sediments.

Depositional mechanisms Ice-rafting under marine (terminoglacial) conditions is identical t o the same process in a lacustrine situation. The only difference is the much higher depositional rate that can be reached (and sustained for a considerable period) under marine conditions. Marine conditions in general have a greater impact on the ice-raft deposits than do lacustrine conditions. The main reason is the difference in intensity and frequency of processes such as current activity and gravity-induced mass movement.

DEPOSITS OF THE MARINE TERMINOGLACIAL BOTTOM-CURRENT FACIES (11-D-2) This facies is characterised by relatively complex deposits, partly because sedimentation has occurred under the direct influence of streams coming from a floating ice shelf. Detailed models of this facies have been presented by Eriksson (19821, Andrews and Matsch (19831, Miall (1983a), McCabe et al. (1984, 1987) and C.H. Eyles (1987). All these authors described the situation in areas that cover both the marine glacial environment and the marine terminoglacial subenvironment as defined here. There is general agreement that tunnel-mouth sedimentation and deposition from bottom currents are very important, although masstransport deposits tend t o dominate the overall sedimentation and although there may be a continuous 'rain' of debris due to undermelting of icebergs.

Deposits of the marine terminoglacial bottom-current facies

509

In view of the above considerations it is understandable that two types of deposits are distinguished within this facies: marine terminoglacial tunnel-mouth deposits (II-D-2-c) and marine terminoglacial bottomcurrent deposits (II-D-2-d). The latter show a gradual transition into mass-transport deposits, as stressed by Eriksson (19821, Miall (1983a) and C.H. Eyles et al. (1985b). The combination of bottom-current deposits

Fig. 302. Model of the marine terminoglacial bottom-current facies. GIF = grounded ice front; TM = tunnel mouth; DF = (submarine) debris flow; TC = turbidity current; IRD = ice-raft deposition; CH = (submarine) channel; t r = traction. Partly based on C.H. Eyles (1987). 1 = nunatak; 2 = active ice; 3 = iceberg; 4 = crevasse; 5 = medial moraine; 6 = submarine fan; 7 = lobes of submarine mass-transport deposits; 8 = marine glacial melt-out complexes; 9 = marine terminoglacial ice-raft deposits; 10 = turbidity current; 11 = marine terminoglacial bottom-current deposit; 1 2 = submarine channel; 13 = seabottom; 14 = mixture of marine terminoglacial ice-raft deposits and marine terminoglacial bottom-current deposits; 15 = coarse-grained base of channel fill; 16 = fine-grained top of channel fill; 17 = partly eroded channel fill; 18 = noneroded channel fill; 19 = substratum.

510

The marine terminoglacial subenvironment

and mass-transport deposits is taken as representative of the 'proximal glacimarine environment' of C.H. Eyles et al. (1985) and of 'ice shelf sedimentation' (Andrews and Matsch, 19831, although the bottom-current deposits may - depending on regional circulation patterns of the seawater - still be found fairly far away from the margin of the ice shelf. Deposits of this facies formed in the most proximal zone of this subenvironment may, together with deposits formed by mass-transport processes and by rain-out from icebergs, build up subaqueous fans of considerable size (Fig. 302). Marine terminoglacial tunnel-mouth deposits (II-D-2-c) These deposits form an irregular mosaic of generally coarse-grained diamicts, resembling in many respects those formed under continental (lacustrine) conditions (N. Eyles, 1983b; Drewry, 1986).This resemblance is understandable because these deposits represent the general situation of subglacial streams embouching in a large body of water (Miall, 1983a; C.H. Eyles et al., 198513). The energy gradients are relatively high and strongly influence the sediment input, resulting in rapidly varying lithofacies. There are often gravelly parts in which the clasts may show a variety of fabrics (Rust and Romanelli, 1975). Cheel and Rust (1982) referred to such coarse deposits as 'uppermost proximal fan'. Channels may be found if the conditions are favourable and if the glacigenic history was appropriate (Carlson et al., 1982). Such strong currents tend to be accompanied by large sediment masses that may be left behind when the water flow slows down. This happens particularly where the tunnel-mouth deposits in the strict sense gradually pass into bottom-current deposits that are mainly formed in channels that bifurcate (Banerjee and McDonald, 1975; Powell, 1981b). The problems related to the genesis and characteristics of these deposits have been discussed in numerous reports (Aalto, 1971, 1981; Edwards, 1978; Deynoux, 1980; Domack et al., 1980; Boulton and Deynoux, 1981; Elverhoi and Roaldset, 1981; Orheim and Elverhoi, 1981; Sugden and Clapperton, 1981; Barrett and Powell, 1982; Domack, 1982, 1983; Eriksson, 1982; Anderson et al., 1983b, 1984; Armencourt, 1983; Miall, 1983a, 1985; Osterman and Andrews, 1983; Kravitz, 1983; Powell, 1983a; J.N.J. Visser, 1983a; McCabe et al., 1984, 1986, 1987; C.H. Eyles, 1985, 1986, 1987, 1988; C.H. Eyles et al., 1985a,b; Visser and Hall, 1985; Dowdeswell, 1986,1987; McCabe, 1986; Benn and Dawson, 1987; Visser et al., 1987; McCabe and Eyles, 1988; N. Eyles et al., 1989; N. Eyles and McCabe, 1989a,b; Fralick and Miall, 1989; Sharpe, 1989).

Deposits of the marine terminoglacial bottom-current facies

5 11

Lit hofacies characteristics

Slackening of the water velocity commonly causes the formation of submarine fans in front of a tunnel mouth. The most proximal parts of such fans may have different lithofacies characteristics but usually consist of coarse, clast-supported, massive diamicts (Fig. 303), with an average grain size that depends partly on the supply and partly on the substratum. Eriksson (1982) stated that conglomeratic types are common and that the coarsest material may be concentrated in channels. Andrews and Matsch (1983) mentioned the resemblance of the deposits t o the 'esker beads' described by Banerjee and McDonald (1975) and referred to the work by Borns (1980) on Norwegian forms that show distinct similarities to De Geer moraines in Sweden and Finland (Zilliacus, 1976, 1981, 1985, 1987a,b, 1990) and to so-called cross-valley moraines of Arctic Canada. The energy in the marine terminoglacial tunnel-mouth area is apparenly too high to produce the sandy, rippled intercalations that are common in the distal parts of esker beads. McCabe et al. (1984, 1987) studied the lithofacies characteristics of these deposits in detail and

Fig. 303. Coarse, diamictic marine terminoglacial tunnel-mouth deposits of Wisconsinan age from British Columbia (Canada).

512

The marine terminoglacial subenvironment

mentioned that gravelly beds coexist with massive, clast-supported diamicts. According t o N. Eyles and McCabe (1989a,b), most marine terminoglacial tunnel-mouth deposits are very coarse, clast-supported diamicts, or mixtures of pebbles and boulders that are only locally enriched in finer material. According to these authors, palaeocurrent indicators are completely absent, but the shape of the deposit and the gradual decrease in grain size in the downstream direction can, together with textural data, give indications as to flow direction (cf. Visser et al., 1987). Textural characteristics

The supply of debris generally ensures that all grain sizes are present, but the coarse fractions (gravels and even boulders) predominate. The deposits may show faint grading (Eriksson, 1982; McCabe et al., 1384; C.H. Eyles, 1987) but there may also bean upward coarsening tendency (Eriksson, 1982). The clasts usually show no preferred orientation although imbrication may be present locally. However, orientation of the particles improves with decreasing grain size and sand grains may have a pronounced orientation, with the a-axis being either parallel or perpendicular t o the flow direction. There may be sandy intercalations, probably representing the infilling of secondary channels, within the generally diamictic sediment. Sand is also present as matrix in the diamicts; fine particles are rarer because, unless they trapped, they are washed out by the strong currents. Occurrence

Marine terminoglacial tunnel-mouth deposits may form non-horizontal, elongated deposits of restricted lateral extent, much resembling (continental) eskers or glaciofluvial fans. These deposits generally interfinger with ice-raft deposits, bottomcurrent deposits and submarine mass-transport deposits; these may also alternate in a vertical direction (Eriksson, 1982). The thickness of the tunnel-mouth deposits may reach some tens of metres. Depositional mechanisms

The depositional process may be compared with that operating when pulp-like material leaves a pipe in a dammed basin. There is a sudden

Deposits of the marine terminoglacial bottom-current facies

513

drop in velocity when the material has left the pipe (or tunnel) is left behind, thus inducing rapid sedimentation of the coarse particles. McCabe et al. (1984: their Fig. 7) described excellent examples seen in Northern Ireland. Episodes with an extremely large supply of clastics (and water) commonly alternate with much longer episodes of a less abundant supply. Periods of almost zero supply are characterised by quiet water in which the finest particles may settle, but the fine-grained deposits will commonly be eroded again during a subsequent phase of high-energy currents. The erosional processes leave only little of the most proximal parts of the tunnel-mouth deposits (C.H. Eyles, 1985a,b, 1987),unless the floating ice shelf retreats so rapidly that previous tunnel-mouth deposits are no longer reached by high-energy currents. There is not a great deal known about the precise sedimentation conditions in this facies and even data concerning the modern facies are controversial (Barrett, 1975; Kellogg et al., 1979). Models are therefore based largely on hard-rock deposits from ancient glaciations (described by, among others, Young, 1978; Visser, 1982; Miall, 1983a). The identification of deposits belonging t o this facies thus remains a problem (cf. Domack, 1983), unless there are clear contacts with well developed and widespread, massive rain-out diamicts (Shaw, 1972b; Barrett, 1975; Sugden and Clapperton, 1981).

Marine terminoglacial bottom-current deposits (II-D-2-d) Channels may be formed in this facies if the meltwater currents over the bottom have a sufficiently high velocity. Such strong currents tend t o transport large masses of glacigenic material that may be left behind when the water flow slows down. This takes place especially in front of submarine fans, but deposition inside channels also takes place at the fans themselves, thus contributing to fan growth. The fans may be very long, with a gradually decreasing inclination in a downslope direction. There is no clear boundary between the fan in the area in front, so that channels starting on the fan can extend far beyond it. The great distances travelled by the debris transported by bottom currents accounts for the considerable differences that may exist between proximal and distal deposits of this type. In addition, there may be meltwater surges that produce conditions comparable t o streamfloods under continental conditions. It is therefore not surprising that complex lithofacies assemblages can be found that should all be ascribed t o this type of deposit (Fig. 304).

The marine terminoglacial subenvironment

514

Distal Sands Proximal Gravels

1

and Mud

Debris Source

c Gravel Cell

I Chute lnfill

Diamictons

Fig. 304.Model of the marine terminoglacial bottom-current facies. From: McCabe e t al. (1984). Courtesy: Journal of Sedimentary Petrology.

These deposits have been described by Eriksson (1982), Hein and Walker (1982), Miall (1983a), C.H. Eyles and N. Eyles (1984), Hein (19841, McCabe et al. (1984, 1986, 1987), C.H. Eyles (1985, 1986, 19871, C.H. Eyles et al. (1985a), J.N.J. Visser et al., 1987 and N. Eyles and McCabe (1989b), among others. Works of special interest are those by Anderson et al. (1980b), Domack (1982) and Gilbert (1983). Most authors make the point that these deposits have much in common with channel deposits formed under lacustrine conditions.

Lithofacies characteristics Channel deposits formed in the upper part of the fan usually comprise badly sorted gravels with boulders (Fig. 305). They form thick, inclined covers (Eriksson, 1982; McCabe et al., 1984), massive over their greater part, although cross-bedding and imbrication may be found in more sandy units. All sediments clearly reflect the predominantly high-energy conditions with minor periods of lower energy, exemplified by the sandy intercalations. Deposits formed somewhat lower on the fan are much sandier and tend t o show trough shapes in cross-section, due t o a relatively simple channel pattern. Truly bifurcating channels are found mainly in the lowermost part of the fan on which these channels develop (McCabe et al., 1984,1987; C.H. Eyles, 1987). Sands are the main constituent at the base of the fan and in the terminoglacial zone in front of it, (Fig. 306). Bottom currents transporting these sands may behave like meandering rivers, building levees and leaving behind fine-grained 'over-

Deposits of the marine terminoglacial bottom-current facies

0 Sandstone

Horizontal Larnnation

w

515

Flame S t r u c t u r e s Sole Marks

Fig. 305. Schematic cross-section of channels in the upper fan. Slightly modified from: Eriksson (1982).

Fig. 306. Marine terminoglacial bottom-current deposits (Brazeau sandpit, Ontario, Canada). Climbing ripples are frequent, indicating simultaneous deposition by current activity and by settling from suspension.

516

The marine terminoglacial subenvironment

bank deposits' after phases of high discharge. These 'overbank deposits' may show small-scale current ripples or fine parallel lamination. The most distal marine terminoglacial bottom-current deposits are mainly laminated sands that often consist of upward grading units. These sands show structures that can be used as palaeocurrent indicators much more commonly than do more proximal deposits (Shaw and Kvill, 1984; Shaw et al., 1989).

Textural characteristics These characteristics vary widely, depending on the distance from the source. Proximal deposits usually show bad sorting, with floating clasts without a preferred orientation. Sorting, but not rounding of the clasts, definitely improves in a downslope direction (Eriksson, 1982; McCabe et al., 1984, 1987; C.H. Eyles, 1987).The clasts show frequent glacial striae.

Occurrence These deposits can be found all over the sea bottom in front of a floating ice shelf (see Fig. 3 of C.H. Eyles et al., 1985b). They form elongated bodies with a lense-shaped cross-section. The thickness may reach about 20 m, but thicker successions are common as a result of a number of channel-fill deposits stacked on top of each other (Fig. 2 in Eriksson, 1982). The bottom-current deposits are commonly distributed over and in front of an area with a cone shape, comparable t o continental sanders, although Rust and Romanelli (1975) stated that 'subaqueous sandurs' are more common under lacustrine conditions.

Depositional mechanisms The marine terminoglacial channels are incised in the submarine fans during erosional stages that occur in periods of high current velocity and abundant meltwater supply. There are several indications that debris flows are responsible for at least some of the channel formation. Most channels are up to a few metres deep, but there are a few reports of channel infillings some 20 m thick. Deposition takes place in the channels when the outflow energy decreases. The resulting fining-upward deposits do not differ essentially from those formed in continental channels, which suggests that the same

Deposits of the marine terminoglacial deltaic facies

517

depositional mechanisms play a role. A repetition of graded successions indicates the pulsatory character of the outflows on the fan. If the currents loose much of their energy, only fine particles will be carried along and erosional phases may be absent. Channels will fill gradually in such cases and there will no longer be a recognisable glacigenic bottom current. Tractions currents may also play a role i n the deposition of these sediments. These currents are still a subject of debate. They will be discussed in some more detail in the corresponding section in the chapter on marine extraglacial bottom-current deposits (II-F-2-d). DEPOSITS OF THE MARINE TERMINOGLACIAL DELTAIC FACIES (11-D-3) Meltwater streams in the marine terminoglacial subenvironment may carry so much debris along that the sudden release into open sea (for instance after leaving a fjord), with ensuing much lower current velocity, results in fans or deltas. These deltas are not fundamentally different from the non-glacigenic deltas formed in the sea, so that terminoglacial marine deltaic top- (II-D-3-b), fore- (II-D-3-c) and bottomsets (11-D-3-d) may be distinguished. If this distinction cannot be made, the deposits are grouped together as marine terminoglacial deltaic complexes (II-D-3-a). Most deltas of this subenvironment are coarse-grained (McCabe and Eyles, 1988; N. Eyles and McCabe, 1989a) and might accordingly be termed 'fan deltas' (cf. McCabe et al., 1987; McPherson et al., 1987, 1988; Marzo and Anadon, 1988; Nemec and Steel, 1988a,b; Wood and Ethridge, 1988).There are only few detailed reports on fan deltas from glacigenic areas (J.N.J. Visser, 1983a; McCabe et al., 1987; McCabe and Eyles, 1988; Fralick and Miall, 1989), probably due t o the fact that most glacigenic deltas are found outside the marine terminoglacial subenvironment.

Marine terminoglacial deltaic complexes (II-D-3-a) These complexes are mostly a result of interrupted growth, so that top-, fore- and bottomsets cannot be differentiated. Such an interruption may be due to a shift in the course of the debris-laden stream, b u t t h e depression in which the delta started t o form may also have been so small t h a t i t filled rapidly and the stream then passed the area without leaving sediments behind shortly after the infilling had begun. Another possible cause of interrupted growth is a decrease in sediment supply, for instance by deposition upstream after development of a sedimentary trap.

518

The marine terminoglacial subenvironment

Lithofacies characteristics The complexes (Fig. 307) are mixtures of sediments with varying characteristics. Gravel may be included, but most of the material is sandy or finer. Sorting is relatively bad, although individual layers may have a rather uniform grain size (except in the case of mass-transport deposits that are often found intercalated). The complexes may show a general increase in grain size in an upward direction, but this upwardfining tendency is often not well developed. Sedimentary structures tend to be abundant, particularly currentinduced structures such as scour-and-fill structures, ripples and parallel lamination, but also small sedimentary deformations (load casts, waterescape structures, normal faults, flowage and fluidisation) are common.

Textural characteristics The scarcity of reports on these complexes means that little is known about the textural characteristics. However, there is no reason not t o

Fig. 307. Marine terminoglacial deltaic complex (sandpit near Osceola, SE of Pembroke, Ontario, Canada). Note the predominance of the foresets.

Deposits of the marine terminoglacial deltaic facies

519

assume that these characteristics will resemble those of deposits formed under continental terminoglacial conditions. This resemblance implies that rounding of the clasts will vary widely, and that the clasts will show local imbrication. A preferred orientation of flattened pebbles more or less to the sedimentary surface is t o be expected. Glacial striae may be found on the larger clasts. Occurrence The specific geographic conditions required for the formation of these complexes implies that they are rare, and occur embedded in other marine terminoglacial deposits. In the glacigenic sequence, they can thus be found throughout the marine terminoglacial interval. The thickness of these complexes is limited, usually t o less than ten metres. They may extend laterally for hundreds of metres, but tend t o frequent interfingering with other deposits (particularly those formed by mass movements), so that the precise extent is not easily determined. Depositional mechanisms The complexes are formed by the settling of material supplied by bottom currents that loose energy. Additional particles may be supplied by rainout from icebergs, but this material is only a minor fraction of the complex. Mass-transport processes, submarine slumping in particular, also contribute t o the final internal structure of these complexes.

Marine terminoglacial deltaic topsets (II-D-3-b) 'Fossil' deposits of this type were described from Canada by Fralick and Miall (1989), from the Karoo Basin (Africa) by J.N.J. Visser (1983a) and from Northern Ireland by McCabe and Eyles (1988). The topmost part of these successions had usually been eroded, so that a complete overview could not be presented. It nevertheless appeared from these studies that the topsets are well comparable with those from non-glacigenic successions (Marzo and Anadon, 1988; Nemec and Steel, 1988b; Postma and Cruickshank, 1988). Lithofacies characteristics Sand and gravel predominate in these deposits. The sands may be massive or horizontally stratified (Fig. 308), whereas the gravels are

520

The marine terminoglacial subenvironment

concentrated in channels. The deposits described by McCabe and Eyles (1988)are coarse, crudely bedded and irregular, whereas those described by Fralick and Miall (1989) are finer grained and more regular, with more (or possibly better visible) sedimentary structures such as low-angle planar stratification, ripple cross-lamination and flaser bedding. Textural characteristics The texture of these deposits has never been described in any detail, but some general aspect have been reported. An important characteristic seems to be the bad sorting of the most proximal topsets; the wide variation in grain size, which is a typical feature of fan deltas, emphasises the resemblance with deltaic topsets formed in lakes. This resemblance is also reflected in some sedimentary features such as grading and mud drapes. The more distal parts of the topsets usually show much better sorting than do the proximal parts (Fralick and Miall, 1989).

Fig. 308. Marine terminoglacial deltaic topsets of Pleistocene age. The coarse material reflects the slight drop in current velocity when the stream spread out over a larger surface area in the former Champlain Sea.

Deposits of the marine terminoglacial deltaic facies

521

The rounding of the clasts is bad to moderate, although well rounded clasts may be present, depending on the nature of the material supplied. The flattened clasts are orientated (sub)parallel to the bedding plane, but imbrication may occur locally at the base of coarse layers. Occurrence

The topsets form elongated, usually somewhat inclined bodies of restricted thickness, covering (and passing in a downstream direction into) deltaic foresets. In an upstream direction, thetopsets pass into marine terminoglacial bottom-current deposits and, sometimes, into marine terminoglacial tunnel-mouth deposits. Lateral contacts are almost exclusively with marine terminoglacial ice-raft deposits and submarine terminoglacial mass-transport deposits. In the glacigenic sequence, these deposits form the topmost part of marine terminoglacial deltas, which are usually found in the 'proximal' part. Depositional mechanisms

The mechanisms involved are not very different from those active in other deltaic topsets. Variations in the supply of material result in phases of stagnant sedimentation or even erosion, alternating with periods of relatively rapid accumulation, especially in channels.

Marine terminoglacial deltaic foresets (II-D-3-c) After they have been formed, these deposits are less affected by erosion than are the foresets. Consequently, they are found more often in the geological record and are known in much greater detail (McCabe, 1986; McCabe et al., 1987; McCabe and Eyles, 1988; N. Eyles and McCabe, 1989a; N. Eyles et al., 1989; Fralick and Miall, 1989). Most researchers emphasise the fact that the characteristics of the foresets resemble those reported for non-glacigenic topsets (e.g., by Casey and Scott, 1979; Fraser and Suttner, 1986; McPherson et al., 1987, 1988; Marzo and Anadon, 1988; Nemec and Steel, 1988a,b).

522

The marine terminoglacial subenvironment

Lithofacies characteristics The lithofacies of these deposits shows a wide range of characteristics, but gravelly/sandy associations are the most common feature. Finer-grained (predominantly sandy) examples are found if the delta developed relatively far away from the ice. The most obvious feature is the inclination of the beds (Fig. 307), usually of the order of 18-22".The inclined interfaces are planar, except for the lowermost part where they approach the underlying bottomsets tangentially. The foresets as a whole tend t o contain some 10% clay and silt, some 50% sand and some 40% gravel (with possibly isolated boulders being present). The boulders are usually found at the base of the foresets; they most probably represent dropstones set free from icebergs. Individual foresets are 10-40 cm thick and tend t o be relatively well sorted, although some diamictic layers may be present. In most cases, such diamicts do not represent deltaic accumulation in its strictest sense, but deposition by mass flows. This can be deduced from the fact that these foresets lack structures such as parallel or cross-lamination, which are frequent in the layers that were formed as a result of grain-by-grain transport.

Textural characteristics Fine-grained foresets tend to be better sorted than coarser ones, but even the latter do not usually show truly bad sorting. Sorting may be very good within any single layer and upward grading in individual layers is frequently found. Rounding of the clasts is highly variable, but seems to be somewhat better than in the topsets. This can be explained by a longer transport distance, but some rolling under the influence of wave activity also cannot be excluded. Flattened clasts are orientated with their ah-planes parallel to the inclined surface of the underlying foreset; imbrication is found only occasionally.

Occurrence These deposits form the middle part of upward coarsening deltaic sequences. The contact with the underlying bottomsets is usually gradual but remains relatively distinct, whereas the contact with the overlying topsets tends t o be much more transitional. Lateral contacts are similar in a downstream and an upstream direction and lateral contacts are

Deposits of the marine terminoglacial deltaic facies

523

commonly with marine terminoglacial mass-transport deposits, although contacts with ice-raft deposits are also known. The foresets may be several tens of metres thick, a few hundred metres wide, and have a length of the same order of magnitude. Depositional mechanisms Other t h a n the mass flows that are always connected with deltaic foresets, the mechanisms involved are, mainly avalanching of material supplied by the bottom currents, and deposition during waning flow in channels that had been incised in the foresets. Settling from suspension also plays a role, as does settling from melting icebergs, but the amount of material supplied in this way is usually far less than the amount supplied by bottom currents. It seems the rule rather than the exception that deposition from masstransport processes is even more important than the other processes (McCabe and Eyles, 1988). A massive appearance may result from gravity flows along the delta front (Postma and Roep, 1985). The absence of grading indicates that there was usually no fully turbulent flow and t h a t grains were supported largely by dispersive pressure prior to freezing. The massive, poorly sorted matrix and the lack of well developed imbrication indicate high-concentration flows with restricted clast movement (Hein, 1982). The concentration of the flows probably decreased during downslope movement, and the resulting acceleration of the flows probably was a n important factor in the progression of the delta (Postma and Roep, 1985).

Marine terminoglacial deltaic bottomsets (II-D-3-d) These deposits were investigated only as part of more general studies (Domack, 1983, 1984; Syvitski et al., 1984; Schwab et al., 1987; McCabe and Eyles, 1988; Fralick and Miall, 1989) so that only scarce details are available. It is known, however, that the density contrast, biological activity and other parameters differentiate these deposits somewhat from their counterparts formed under lacustrine conditions (Gadd, 197 1; Hoskin and Burrell, 1972; Edwards, 1978; Syvitski and Murray, 1981; Mackiewicz and Powell, 1982; Smith and Syvitski, 1982; Stewart, 1982). Lindsay (1971) suggested that the lack of cycles found in marine deltaic bottomsets is the most diagnostic difference from lacustrine bottomsets, even though some cycles have been described in prodeltaic marine sediments (Carrey and Ahmed, 1961; Powell, 1980; Gilbert, 1982). Cycles

524

The marine terminoglacial subenvironment

of this type are, however, much less clear than typical lacustrine rhythmites (Ashley, 1975; Edwards, 1978).

Lithofacies characteristics The various descriptions of these deposits indicate that there is no characteristic lithofacies. McCabe and Eyles (1988) described bottomsets of this type as consisting of thinly bedded, massive, graded and rippled sands, alternating with laminated silty clays. Domack (1984) found such deposits, which he termed 'distal meltwater fans', to consist of laminated silty sands and of pebbly sands, silts and muds. The sandy silts and silty sands were thin-bedded, showing diffuse to well developed lamination, with intraformational rip-ups of sandy silt just above the lower contacts. Other structures seen are large-scale current ripples (occasionally with overturned foreset laminae), current lineation, convolutions and waterescape structures. The pebbly muds (diamicts) were massive t o indistinctly laminated. The sands showed streaks of gravel, horizontal laminae of silt, grading, climbing ripples, water-escape structures and rhythmic alternations of sand and silt (Fig. 309). Other descriptions indicate that the frequent lamination found by Domack (1984) may be rare or at least vague, and that small-scale and large-scale sedimentary deformations may be abundant. Textural characteristics Both the activity of traction currents and settling from suspension contribute t o good sorting of these deposits, but the embedded rain-out material and mass-transport deposits have an opposite result. In spite of this, sorting is better than in the related deltaic fore- and topsets. The larger clasts, present only as a result of rain-out, show no preferred orientation. Intercalated mass flows may contain pebbles orientated parallel t o the flow lines, but these deposits belong to the marine terminoglacial mass-transport facies. Occurrence These deposits may be found in a relatively wide belt surroundig the delta in a strict sense. They gradually thin out in the off-delta direction, interfingering with ice-raft deposits, mass-transport deposits, or deposits of non-glacigenic nature. The same transitions are found in lateral directions.

Deposits of the marine terminoglacial deltaic facies

525

Fig. 309. Marine terminoglacial deltaic bottomsets of Late Pleistocene age. Surroundings of Ottawa (Ontario, Canada). The predominantly sandy nature i s possibly due to wash-out of finer particles..

The bottomsets are found as the bottom part of deltaic sequences. The transition into the overlying topsets is gradual but relatively distinct. Contact with the underlying sediments is usually much more transitional and the thickness of these deposits is therefore not easily determined. About ten metres seems to be a reasonable estimate of the maximum thickness.

Depositional mechanisms Deposits consisting of silty sand represent bottom-current or tractionload deposits, as suggested by their relatively good sorting, the presence of current ripples and lineations, grading of silt rip-ups, etc. Muds indicate quiet bottom conditions, as reflected by the absence of current structures (pebbles found in these muds are dropstones); it is likely that most of these muds were formed by settling from suspension of particles supplied by turbid overflow or interflow plumes. The relative abundance of poorly sorted sands and pebbles is due to rain-out processes, but underflow currents (in the sense used by Beates

526

The marine terminoglacial subenvironment

and Freeman, 1953) account for the larger supply of these deposits. This is proven by the abundant current structures. Underflows are rare under marine conditions because of the high specific weight of seawater (due to the salinity of approx. 3.2%), so that high concentrations ( > 39 g 1-l) of particles a r e required for the existence of underflows (Bates, 1953; Hoskin and Burrell, 1972; Gilbert, 1979). Underflow processes have been inferred under conditions of reduced salinity and abnormal concentrations of suspended sediment, such as occur near deltas formed at the mouth of a fjord (Hoskin and Burrell, 1972; Gilbert, 1982; Stewart, 1982). The absence of rhythmites from the bottomsets has been explained in several ways (Edwards, 1978). Dense underflows formed directly from inflowing meltwater have been widely recognised as playing a key role in the formation of glaciolacustrine rhythmites. Such current are, however, suppressed under marine conditions because of the low density contrast between seawater and debris-laden meltwater. DEPOSITS OF THE MARINE TERMINOGLACIAL MASS-TRANSPORT FACIES (11-D-6) This facies, the prevailing processes and the resulting deposits have been abundantly studied and described (Armstrong and Brown, 1954; Carey and Ahmad, 1961; Bjorlykke, 1967, 1969; Boltunov, 1970; Chriss and Frakes, 1972; Baker and Friedman, 1973; Deynoux and Trompette, 1976, 1981; Dreimanis, 1976a,b, 1979, 1982a, 1984a, 1987, 1989; Plafker and Addicott, 1976; Carlson and Molnia, 1977; Evenson et al., 1977; Andrews, 1978; Edwards, 1978; J.B. Anderson et al., 1979, 1980b, 1982, 1983a,b, 1984; Kurtz and Anderson, 1979; Borns, 1980; Clark e t al., 1980; Deynoux, 1980; Domack et al., 1980; Elverhoi et al., 1980, 1983; Aalto, 1981; Boulton and Deynoux, 1981; Drewry and Cooper, 1981; Elverhoi and Roaldset, 1981,1983; Hicock et al., 1981; Orheim and Elverhoi, 1981; Powell, 1981a, 1983a,b, 1984; Anderton, 1982; Barret and Powell, 1982; Birkenmayer, 1982; Carlson et al., 1982; Domack, 1982, 1983, 1984; Eriksson, 1982; Gilbert, 1982, 1983; J.N.J. Visser, 1982, 1983a,b,c; Vorren et al., 1983; J.B. Anderson, 1983; Andrews and Matsch, 1983; Armencourt, 1983; Bugge, 1983; Christie-Blick, 1983; Domack a n d Anderson, 1983; Elverhoi and Solheim, 1983; C.H. Eyles and N. Eyles, 1983b, 1984a,b, 1985; N. Eyles, 1983b, 1987; Gilbert and Church, 1983; Kravitz, 1983; Miall, 1983a, 1985; Mode et al., 1983; Molnia, 1983a,b,c; Osterman and Andrews, 1983; Schwab and Lee, 1983,1988; Bornhold and Yorath, 1984; Fortuin, 1984; Gravenor et al., 1984; Mackiewicz et al.,

Deposits of the marine terminoglacial mass-transport facies

527

1984; McCabe et al., 1984, 1986, 1987; Broster and Hicock, 1985; C.H. Eyles et al., 1985a,b; N. Eyles and Clark, 1985; Fralick, 1985; Von Brunn and Talbot, 1985; Dowdeswell, 1986, 1987; Drewry, 1986; C.H. Eyles, 1986, 1987; McCabe, 1986; Benn and Dawson, 1987; Brodzikowski and Van Loon, 1987; Schwab et al., 1987; Visser and Loock, 1987; McCabe and Eyles, 1988; N. Eyles and C.H. Eyles, 1989; N. Eyles and McCabe, 198913; N. Eyles et al., 1989; Fralick and Miall, 1989). Such wide interest is not surprising, a s this is by far the most widespread facies of the marine terminoglacial subenvironment. The interest of so many researchers is also understandable as the - predominantly diamictic - deposits of this facies result from a wide variety of processes (C.H. Eyles et al., 1985; Drewry, 1986). By definition, all mass-transport deposits in this subenvironment are formed under subaqueous conditions and thus they should all be classified as II-D-6-c (marine terminoglacial mass-transport deposits). Several models (Fig. 310) have been proposed for the facies and its deposits (Evanson et al., 1977; Stanley and Maldonado, 1981; Wright a n d Anderson, 1982 (based on Stanley, 1981); Gravenor et al., 1984; McCabe et al., 1984,1987; C.H. Eyles et al., 1985; N. Eyles and McCabe, 1989a).

Marine terminoglacial mass-transport deposits (II-D-6-c) The deposits show a variety of characteristics. One type is the so-called waterlain flow till, as understood by Evenson et al. (1977); this type is identical to the corresponding type formed under continental terminoglacial conditions (II-A-6-c). Depositional models (Dreimanis, 1979; Hicock et al., 1981) indicate that this type is formed by slumping of glacial debris on the ice margin (Fig. 62; cf. Eyles et al., 1985) and their coming to rest at the seabottom i n front of the ice shelf. Such deposits have been described by Edwards and Foyn (1981), Domack (1983) and McCabe et al. (1984). Eyles et al. (1985) attributed these sediments to their slump facies, which combines all debris-flow deposits, slumps and turbidites (cf. Cheel and Rust, 1982). A second type of deposit is formed by diamicts with rapidly varying characteristics, built up by local redeposition. These deposits were transported in the form of slumps, slides, mudflows (several types were distinguished by Middleton and Hampton, 19761, sometimes passing into turbidity currents of gradually decreasing intensity (Stanley, 1981; Stanley and Maldonado, 1981; Wright and Anderson, 1982). Local processes such as non-glacigenic traction currents and rainout from icebergs may affect the mass-transport processes and thus the resulting

528

The marine terminoglacial subenvironment

Fig. 310. Schematic model of the marine terminoglacial mass-transport facies. GIF = grounding ice front; F = (submarine) fan; S = slumping; SL = sliding; Ad = deposition by avalanching; SC = submarine canyon; IF = inner fan; MF = mid fan; OF = outer fan; BP = basin plain; tr = traction. 1 = active ice; 2 = nunatak; 3 = marginal crevasse; 4 = medial moraine; 5 = tunnel mouth; 6 = proximal part of submarine fan; 7 = distal part of submarine fan; 8 = continental slope; 9 = submarine canyon system; 10 = marine terminoglacial bottomcurrent deposits; 11 = canyon with frequent mass transport; 12 = small canyon with little mass transport; 13 = coarse-grained debris-flow deposits; 14 = bottom-current deposits on inner fan; 15 = fan built up mainly of small mass-transported units; 16 = submarine channel; 17 = distributary channel on fan; 18 = med-fan deposits; 19 = traction-current deposits; 20 = fine-grained turbidites; 21 = outer limit of turbidite fan; 22 = high-concentration turbidity current; 23 = low-concentration turbidity current; 24 = local traction; 25 = non-glacigenic sediments; 26 = substratum.

Deposits of the marine terminoglacial mass-transport facies

529

deposits. Rapid lateral variations in the characteristics of the deposits are therefore a common feature (Edwards and Foyn, 1981; Domack, 1983; McCabe et al., 1984,1986,1987; McCabe and Eyles, 1989a,b). Lithofacies characteristics

It appears from numerous descriptions (Domack, 1982,1983; Wright and Anderson, 1982; C.H. Eyles et al., 1985; Dowdeswell, 1987) that the deposits may range from almost pure gravels t o almost pure muds. Diamictic deposits (Fig. 31l),however, dominate (Von Brunn, 1987). N. Eyles and McCabe (1989a) distinguished four types of such diamicts that occur in both the proximal and the distal zones of this subenvironment. The first type consists of massive or stratified diamicts with flow structures, abundant silt and clay clasts, and local concentrations of pebbles. This type is considered to be formed as a result of reworking after an oversteepened accumulation of ice-raft deposits had become unstable. The second type also consists of massive and stratified diamicts, but shows, in addition to the characteristics of the first type, several types of

Fig. 311. Marine terminoglacial mass-transport deposits formed in the direct vicinity of the ice front (Brazeau pit, Ontario, Canada).

530

The marine terminoglacial subenvironment

grading. These deposits, which fill scours and level height differences in the seabottom, are taken t o represent several forms of debris flows. The third type, also comprising massive and stratified diamicts, is characterised by the occurrence of silt and sand stringers, rippled sand units and some winnowed layers. The bodies clearly follow channel patterns and are interpreted as being formed from high-density currents. The last type distinguished by Eyles and McCabe consists almost exclusively of stratified diamicts showing both indications of traction-current activity and abundant deformation structures (flow and drag folds, dykes, waterescape structures, load casts, etc.). These deposits are formed by lowdensity currents. The deposits are coarsest in the proximal zone of this subenvironment and gradually become finer in more distal directions. Studies carried out along the Antarctic continental slope (Anderson et al., 1979,1982,1983b; Wright and Anderson, 1982),over a distance of 600-700 km, indicate that the ensemble might be described as a turbidite fan, with massive o r poorly stratified units in the proximal part and typical turbidites, with the Bouma intervals, in the distal zone (cf. Powell, 1980; Gilbert, 1982; Domack, 1984). There are also mass-transported sediments with a breccia-like character. These diamicts tend to contain large slabs of deformed unlithified material (Morgenstern, 1967; Molnia, 1977, 1983b; Anderson et al., 1980b, l982,1983a,b, 1984; N. Eyles, 1983b; Schwab et al., 1987). Textural characteristics

As several depositional mechanisms are involved in the genesis of these sediments, no particular textural characteristics are seen. Sorting is usually bad, except in turbidites (Wright and Anderson, 1982). Vertical grading is common, particularly in turbidites (cf. Middleton and Hampton, 1976). The orientation of the larger clasts in these deposits depends on the precise transport and depositional mechanism, but most deposits formed under these conditions show a clast orientation in one direction or another (Postma, 1979, 1984a; Nemec et al., 1980, 1984; Postma and Roep, 1985; Postma, 1986; Nemec and Steel, 1988a,b; Postma et al., 1988a,b). Occurrence

These deposits cover about 80% of the marine terminoglacial subenvironment and are found usually also (buried) where other deposits cover the

Deposits of the marine terminoglacial mass-transport facies

531

seafloor. They may extend from the shelves t o the continental slopes and even to the abyssal plains (Elverhoi and Roaldset, 1981,1983; Wright and Anderson, 1982; Elverhoi and Solheim, 1983). This means that they do not extend as far as the distance icebergs may float away from a n ice shelf, but the distribution of well traceable mass-flow deposits and iceraft deposits is roughly similar. The rate of reworking depends on, among others, topographic differences. Continental slopes, but also areas where much material accumulates as a result of rain-out, are thus favourable for the genesis of these sediments. Large amounts of debris are supplied, especially in the direct vicinity of a floating ice shelf, so that mass-transport deposits may form thick successions in this area. The units gradually wedge out in distal directions. They interfinger with marine terminoglacial bottom-current deposits and marine terminoglacial ice-raft deposits.

Depositional mechanisms Several processes are involved. These are not essentially different from those active in other glacigenic or non-glacigenic (sub)environments. Debris flow, grain flow, fluidised flow, turbidity currents, slumping and sliding seem t o be the most prominent processes, but other processes (e.g., failure as a result of intraformational pressure resulting from unstable density gradients) ma prevaily locally.

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The marine proglacial subenvironment

533

THE MARINE PROGLACIAL SUBENVIRONMENT (11-E) The marine proglacial subenvironment (see also the Table on p. 130), as defined here, has no practical sedimentological meaning. To be consistent with the definition given for the continental proglacial subenvironment, one should define this subenvironment as the zone surrounding the marine terminoglacial subenvironment, not reached by mass-transported material that started flowing in the glacial environment, but still characterised by the occurrence of dead-ice bodies in the subsoil and by the presence of permafrost. There are several reasons why such a definition is purely academic rather than applicable in practice. First, subaqueous mass treansport (turbidity currents in particular) may proceed over such large distances that the extent of the terminoglacial subenvironment would be incomparably larger under marine conditions than under continental conditions (where a terminoglacial lake of 10 km is already exceptional). The continental proglacial subenvironment thus starts where the strong energy gradients of the terminoglacial subenvironment have just become much lower, whereas the marine terminoglacial subenvironment is so wide that the zone of strong energy gradient is only an inner 'shell'; the marine proglacial subenvironment would start in the middle of a wide zone with low energy gradients. A second reason for defining the continental and marine periglacial subenvironments in a different way is the nature of permafrost. This phenomenon is commonly taken as the main (if not the only) criterion for determining the extent of the periglacial zone. However, the lowermost water layers in oceans have a fairly constant temperature of about 4 "C, so that permafrost can occur under marine conditions only in a small zone within the terminoglacial subenvironment. On the other hand, it seems illogical to restrict the marine periglacial subenvironment t o the zone with permafrost, as glacigenic features (e.g., deposition from icebergs, mass movements started in the glacial environment) dominate in a much wider zone. A third reason why the existence of a marine proglacial subenvironment is of little practical significance is that this subenvironment, under continental conditions, has been defined as the zone outside the terminoglacial subenvironments where buried dead-ice bodies can still be

534

The marine proglacial subenvironment

found. By definition, Such bodies can only exist if the ice has retreated; dead-ice bodies might thus be found under the seafloor if a n ice cap retreated and the area became covered by the sea. Such a situation might occur occasionally, but buried dead-ice bodies would then melt away very rapidly. The above are considered strong enough arguments to warrant the application of other definitions under marine conditions. The marine proglacial subenvironment is, especially for the reason of buried dead-ice, taken to be of only academic interest. This might imply that the marine terminoglacial and extraglacial subenvironment would pass directly into each other (as is the case under continental conditions during many of the phases of ice advance: no buried dea-ice blocks can then be found in the subsoil), and any boundary between these two marine subenvironments would be highly hypothetical.

continental glacial environment L

marine periglacial environment

marine glacial environment ,I/

-

I A

\IT-

7 1

~

marine terminoglacial subenvironment L ~-

continental glacial environment

'A

continental slope

continental periglacial environment

I

end of marine

marine extraglacial subenvironment

I

-

IT

/

~

I

I

L -

7, 1

~

glaciomarine delta

-~

I

continental slope

1L -

.

~~

y -

marine periglacial environment

/I

,

Fig. 312. Setting of the marine terminoglacial (top) and the marine extraglacial (bottom) subenvironments, showing why the short-lived marine proglacial subenvironment has no palaeogeographic or sedimentological significance.

The marine proglacial subenvironment

535

The authors therefore deem i t most appropriate to consider the marine proglacial subenvironment o f academic significance only, without facies and deposits that would deserve attention from a practical point of view. As mentioned in the previous chapter, the marine terminoglacial subenvironment is, for the above reasons, taken to extend to the outer zone reached by mass movements of glacial origin. This subenvironment is only present if glacial ice reaches the sea (Fig. 312), so that mass flows originating in the glacial environment enter the sea immediately (without passing a continental terminoglacial zone). Icebergs can be found in this subenvironment. For the same reasons, the marine extraglacial subenvironment is, taken t o extend to the outer zone reached by mass movements of glacigenic (glacial or periglacial) origin. This subenvironment is, however, only present if glacial ice does not reach the sea (Fig. 3121, so that mass flows originating in the glacial environment enter the continental terminoglacial zone; mass flows originating from this zone flow into the sea and determine (for instance in the form of slumps or turbidity currents) the extent of this subenvironment. Icebergs are not present in this subenvironment as the ice does not reach the sea and ice calving is thus impossible.

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The marine extraglacial subenvironment

537

THE MARINE EXTRAGLACIAL SUBENVIRONMENT (II-F)AND ITS DEPOSITS According t o the definition in the last paragraph of the previous chapter (Fig. 312), this subenvironment (see also the Table on p. 130) is present in front of a continental terminoglacial (occasionallyproglacial) subenvironment, at least if the zone is still under the direct influence of glacigenic sedimentation (it must be reached by mass movements that started under glacigenic continental conditions). As the 'parent' ice does not reach the sea, icebergs are not present in this subenvironment (or they are driven to this subenvironment from a distant floating ice shelf). Permafrost may be present in a relatively narrow belt along the coast (Line11 and Tedrow, 1981). Non-glacial processes usually dominate sedimentation in this subenvironment. In contrast to the situation in the marine terminoglacial subenvironment, where the glacial ice is in direct contact with the sea, mass movements of any significant size are more an exception than the rule. Nevertheless, glacigenic deposits (predominantly submarine masstransport deposits) may cover thousands of square kilometres, particularly because non-glacigenic sedimentation mainly a result of settling from suspension at a very low rate. It follows from the above that the lack of direct contact with seawater accounts for the most obvious differences between the marine terminoglacial and extraglacial subenvironments: icebergs are absent and glacigenic mass-transport deposits are relatively rare in the latter subenvironment. Not only are these two aspects the most obvious ones from a sedimentological point of view, they are also the essential aspects. Otherwise, both subenvironments show the same characteristics: they may be reached by large meltwater streams that carry significant quantities of glacial debris, such streams may form deltas if the topography is suitable, and - as already stated - mass-transport processes are common. It is therefore logical to distinguish three facies in this subenvironment: the marine extraglacial bottom-current facies (II-F-2), the marine extraglacial deltaic facies (XI-F-3)and the marine extraglacial mass-transport facies. Non-glacigenic marine facies may be found between the glacigenic ones, with concentrations that increase with distance from the coast.

538

The marine extraglacial subenvironment

SEDIMENTATION PROCESSES IN THE MARINE EXTRAGLACIAL SUBENVIRONMENT Glacial ice h a s no direct influence on the sedimentary processes. Meltwater streams are the only glacigenic feature t h a t may affect the local energy level in this subenvironment. However, the existence of a continental terminoglacial subenvironment between the ice and this subenvironment means that relatively low energy levels are the rule in the latter subenvironment, even though relatively steep deltas may be formed where terminoglacial meltwater streams enter a sea. Extraglacial deltas (described under several names) are known from a great number of places and the processes involved have been described in detail (Andrews, 1978; Boulton and Deynoux, 1981; Andrews and Matsch, 1983). The rate of sedimentation just in front of a marine extraglacial delta has been estimated to be 10-20 cm per year, gradually decreasing to 1-2 cm per year further away (Relling and Nordseth, 1979). Mass-transport processes are most frequent along the delta fronts, but also occur elsewhere in this subenvironment if the conditions are favourable (unstable sediment accumulation, trigger, slope). Large glacigenic turbidite fans may even be formed on the abyssal plain. The non-glacigenic processes that contribute to sedimentation in this subenvironment are beyond the scope scope of the present study. They are described, sometimes together with glacigenic depositional processes, in several oceananographic studies of a more general character (e.g., Tabata and Pickard, 1957; Syvitski and Bayliss, 1980; Stigebrandt, 1981; Syvitski and Lewis, 1980; Syvitski and McDonald, 1982; Farrow et al., 1983; Kostaschuk and McCann, 1983; Schwab and Lee, 1983, 1988; Syvitski and Farrow, 1983; Syvitski et al., 1983a,b; Schwab et al., 1987). The most important work on marine glacigenic depositional processes is that of Drewry (1986), who discussed the sedimentation in fjords and estuaries, partly under conditions that are extraglacial in the sense used in the present book. The grain-size distribution under marine extraglacial conditions was studied by, for example, Relling and Nordseth (1979), who found that the largest average grain size occurs a little distance off the mouth of a (continental) terminoglacial stream, then decreases i n a n off-coast direction. Others, e.g., Gilbert (1982) found no significant decrease i n grain size in the extraglacial Maktak fjord (Baffin Island). This author ascribed the lack of lateral grading to aeolian supply, mass movements, bioturbation and widespread distribution of fine-grained sediments by active suspensions.

Deposits of the marine extraglacial bottom-current facies

539

Biogenic sedimentation seems to be more important under marine extraglacial than terminoglacial conditions (Pearson, 1980). Remnants of dead organisms, settled on the seabottom, may contribute significantly t o the overall sedimentation. Upwelling of nutrient-rich currents may favour this process but inner fjord basins that are fed by glacial meltwater may also be major sources of nutrients (Apollonio, 1973). The biogenic sediments might thus be considered as the result of a process at least partly influenced by glacigenic factors. Sharma (1979) found that biennial diatom blooms in fjords of SE Alaska extract large amounts of silica from seawater; After they have died the silicious diatoms sink to the bottom, where considerable masses may build up. Silicoflagellates and foraminifers may play comparable roles.

DEPOSITS O F THE MARINE EXTRAGLACIAL BOTTOM-CURRENT FACIES (11-F-2) This facies contains only one type of deposit: the extraglacial marine bottom-current deposits (II-F-2-d).They are formed partly inside distinct channels (as described by Eriksson, 1982; McCabe et al., 1984,1987; C.H. Eyles, 1987) and partly in not distinctly incised areas (C.H. Eyles and N. Eyles, 1984; N. Eyles et al., 1985). The facies has been described under several names, also from comparable non-glacigenic areas (Massari, 1978; Winn and Dott, 1979; Hein and Walker, 1982; Postma, 1984). The nature and the genesis of so-called 'traction-current deposits' in this facies are controversial. According t o some authors, these deposits are merely bottom currents with a relatively large amount of particles (due to a continuous and large supply), whereas others assume that the deposits are - a t least partly - of a mass-transport nature (Anderson et al., 1979; Kurz and Anderson, 1979; Hill, 1984; Piper et al., 1985). There probably coexist a variety of transport and depositional processes; a complex pattern of sediments may thus be formed, depending on which process dominates.

Marine extraglacial bottom-current deposits (II-F-2-d) Wherever channels are found in this facies, the pattern is much less complex - while there is bifurcating - than it usually is under subaerial conditions. The sedimentation within the channels can therefore be described with a relatively simple model t h a t details p a r t of the submarine fan (McCabe et al., 1984). C.H. Eyles (1987) described a wide

540

The marine extraglacial subenvironment

variety of such deposits from the Alaskan Yakataga Formation and proposed a detailed palaeogeographic model with characteristic facies associations. Other relevant work on sediments in this facies is by Eriksson (1982) on the Archean Pilbara Block (Australia) and by Walker (1975, 1978), Miall (1983a, 1985), Molnia (1983a), Clifton (19841, Hein (1984), C.H. Eyles (1986),and Visser and Loock (1987). Deposition outside the typical channels is mainly from traction currents. There is some resemblance to subaerial sheet floods, particularly on shallow shelves that are strongly affected by wind-induced waves and by strong tides (H.D. Johnson, 1978). Circumpolar bottom currents (socalled contour currents) have been reported that have velocities of 15-30 cm s-l (Gill, 19731, whereas density-driven thermohaline currents (produced by the formation of sea-ice) can reach 1-7 cm s-l (Anderson et al., 1982). Cold bottom currents may flow across the shallow shelf with velocities of up t o 40 cm sW1near the shelf break (Orheim and Elverhoi, 1981). Currents at an intermediate depth, with a velocity of up t o 57 cm s-', are known from the Arctic ocean basin, where the velocity decreases t o less than 6 cm s-l in the deep abyssal plain (Clark et al., 1980). All these data on currents prove that conditions in this facies are favourable f o r sedimentation.

Lithofacies characteristics The deposits formed in distinct channels consist mainly of coarse material; gravel does occur but sand is more common (Eriksson, 1982). Detailed studies carried out by C.H. Eyles (1987) on Middleton Island (Yakataga Formation, Alaska) have shown that grading is a common feature and massive gravels are much rarer. The massive gravels seem restricted t o the basal parts of the channel fills, and their lower boundary is often erosive. Gravels and diamicts may alternate higher in the channel-fill succession, generally with a lenticular geometry. Such successions may be up t o 30 m thick (C.H. Eyles, 1987) and be bordered by the steep walls of the U-shaped channels. The sediments become gradually less coarse in a distal direction and sedimentary structures such as (high-energy) lamination, cross-bedding and grading become more common. Reversed grading has also been observed (C.H. Eyles, 1987). A more complex pattern of gravels, sands and diamicts is seen in the deposits formed outside the channels, mainly under the influence of traction currents. Sedimentary structures, mainly small- and medium-scale trough cross-bedding, grading and reversed grading, are commonly found in the sandy parts. The more gravelly deposits tend t o be massive,

Deposits of the marine extraglacial bottom-current facies

541

although low-angle cross-bedding may be present. The diamicts deposited by the traction currents tend to be stratified as a result of phases of winnowing; silt and sand stringers are a very common occurrence within these deposits.

Textural characteristics The channel-fill sediments generally show well developed imbrication. Sorting may range from good to poor, depending on the supply and on the velocity of net deposition. The deposits formed by the traction currents also show imbrication if the clasts are the right shape and if they are large enough. Graded bedding is a common phenomenon and there is usually good sorting, especially in the thinner layers. These properties cannot be regarded as appropriate criteria for facies interpretation; comparison with the adjoining deposits is therefore generally necessary if the interpretation is t o be reliable.

Occurrence The channel-fill deposits are found primarily in the area of middle and distal submarine fans, although some channels have been reported from proximal fans (N. Eyles et al., 1982a). The deposits have a characteristic shape (generally well reconstructable in lithified deposits, if sufficiently exposed), due t o the incision into the soft substratum. The U-formed channel shape (usually a few metres deep, but can reach up t o several tens of metres) is reflected in the infilling and indicates the palaeocurrent direction. The deposits outside the channels are shaped quite differently: instead of shoestring-like bodies they form sheets, with a faint lense shape. Miall ( 1983a) has described such sheets with current ripples from the Precambrian Gowganda Formation (Fig. 313; also see Mustard and Donaldson, 1987). These sheets commonly overlie the channel-fill deposits of this facies, which seems to imply that the initial alternation of sedimentation and erosion gradually passes into a final phase without much erosion (which is comparable with the situation in a prograding deltaic facies). The deposits are usually found interfingering with deposits from the marine extraglacial mass-transport facies (II-F-6), which is consistent with the descriptions of the 'proximal glacimarine environment' given by Miall (1983a) and N. Eyles et al. (1985). There may also be contact with non-glacigenic sediments.

542

The marine extraglacial subenvironment

Fig. 313.Relatively fine-grained marine extraglacial bottom-current deposits from the Huronian Gowganda Formation (road section near Cobalt, Ontario, Canada).

In the glacigenic sequence, these deposits are found where glaciomarine sediments are in contact with continental terminoglacial sediments.

Depositional mechanisms The depositional mechanisms are relatively simple. The large amounts of glacigenic debris supplied mainly by terminoglacial streams are transported further by bottom currents that gradually build up submarine fans. As in subaerial fans, channels become incised in the fan as a result of localised flows. The proximal channels tend to be characterised by currents with a relatively high energy, resulting in more or less straight channel configurations with steep side walls. Deposition, mainly of coarse material, occurs during periods of currents with a lower energy level. The inclination of the channels usually diminishes in a down-fan direction. This change is expressed in a more irregular channel pattern and in the deposition of less coarse material.

Deposits of the marine extraglacial bottom-current facies

543

Particles are transported within the channels partly in suspension and partly as bed load. Concentration of particles in the lowermost water layers gives rise to the so-called traction currents and sometimes even to debris flows. Sudden 'freezing' of cohesive flows i n t h e channels is probably the reason why massive gravels occur in the lower parts of the channel fills (cf. Lowe, 1982; Nemec and Steel, 1984). This process also explains the occurrence of diamicts in the channels; these glacigenic diamicts thus are not tills but are comparable to the pebbly mudstones that are frequently found in marine environments where a n unstable slope occurs near points with a supply of coarse clastics (cf. Van Loon, 1970). Other coarse, gravelly beds are formed as a result of transport of cohesionless sediment, in which the clasts can move freely; sorting is generally poor due to rapid transport, short travel distance and relatively high viscosity of the flow (Middleton, 1967; Davies and Walker, 1974; Hein, 1984). The steep imbrication, up to 38" (C.H. Eyles, 1987) indicates rapid deposition (Hiscott and Middleton, 1979), possibly as a result of the pulsational character of the supply from meltwater streams. The grading found in many channel fills in the middle and distal parts of the submarine fans may be due to winnowing currents but is also due to low-concentration turbidity currents (Middleton and Hampton, 1976). The reversed grading that is sometimes found can be explained as the result of dispersive pressure or kinetic sieving (Bagnold, 1955; Middleton, 1967; Walker, 1975, 1978; Naylor, 1980). The rare cross-bedding shown by the channel-fill deposits was explained by C.H. Eyles (1987) as being the result of traction-induced structures such as dunes that develop in mid-channel positions (cf. Massari, 1978; Hiscott and Middleton, 1979; Hein, 1984; Lash, 1984). Finer deposits may be formed from waning traction currents in the areas between the channels. Diamicts with a specific grain-size distribution may thus be formed; these are the 'compound paratills' of Ruddiman and Bowles (1976) and Anderson et al. (1982). Variations i n the energy of the traction currents generate a laminated mud that is generally poor in ice-rafted material. Horizontal lamination and cross-bedding in deposits outside the channels are of the same origin as in fluvial deposits (see, for example, Sundborg, 1956; J.R.J. Allen, 1964, 1974, 1983; Selley, 1965; Barrett, 1970; Dionne, 1970; Adeleye, 1974; High and Picard, 1974; Cant and Walker, 1976; Shepard, 1978; Eriksson et al., 1979; Jones, 1979; Ruegg, 1981, 1983; P.A. Allen, 1982; P.A. Allen et al., 1983; Rust and Legun, 1983).

544

The marine extraglacial subenvironment

DEPOSITS OF THE MARINE EXTRAGLACIAL DELTAIC FACIES (11-F-3) The large amounts of debris commonly laid down at the front of a continental ice cap are partly removed by meltwater streams. These streams may embouch in the sea, where they usually form deltas (Fig. 314). While, under non-glacigenic circumstances, the boundary between the deltaic area and the sea is mainly influenced by the sediment supply (progradation of the delta when the supply is abundant and erosion when the supply is poor) under glacigenic conditions this boundary is also influenced by the oscillations of the ice cover. The implication is that marine extraglacial deltas tend to have a much less regular internal structure than do examples from more moderate climates (Andrews, 1978; Boulton and Deynoux, 1981; Andrews and Matsch, 1983). Another characteristic that allows these deltas to be differentiated from other (especially fresh-water) glacigenic deltas is the fossil content. Though macrofauna may be scarce, thorough investigation of the microfauna will nearly always evidence the cold marine character (Birkenmayer, 1982).

Fig. 314. Marine extraglacial delta (Svalbard). Photograph: J. Bierofiski.

Deposits of the marine extraglacial deltaic facies

545

Four main types of deposits can be discerned within this facies: the marine extraglacial deltaic complexes (II-F-3-a),the marine extraglacial deltaic topsets (11-F-1-b),the marine extraglacial deltaic foresets (II-F-1c) and the marine extraglacial deltaic bottomsets (11-F-1-d).

Marine extraglacial deltaic complexes (11-F-3-a) Such complexes are found where the deltas had insufficient time (or insufficient supply of particles) to develop in the classical way. As a result, such complexes usually have small dimensions, making their recognition in ancient deposits difficult. The position, in front of continental deposits, often makes even recognition of the marine character difficult but analysis of the microfossils may still reveal the true nature. The glacigenic character can usually be deduced from the occurrence of striated pebbles. There are no existing descriptions of deposits of this type dating from Pleistocene or older glaciations. The general characteristics are, however, well known from modern examples.

Lithofacies characteristics The characteristics vary greatly, depending not only on the material supplied but also on the material removed subsequently by marine processes (waves, long-shore currents, tides, etc.). Gravel may be present, but sands and finer particles are more common. Sorting may be good or bad, depending on the prevailing processes involved, but most individual layers have a fairly uniform grain size. The complexes in principle show upward grading (individual channels may be filled with upward fining sediments), but recently formed small deltas of this type do not always show a well developed upward grading. There are abundant sedimentary structures such a s scour-and-fill structures, ripples and parallel lamination. Small sedimentary deformations (load casts, water-escape structures, small normal faults, flowage and local fluidisation) are not unusual. Many of the layers and structures may have been 'decapitated' by the erosive capacity of mass movements occurring along the delta front.

Textural characteristics These characteristics are not essentially different from those of deltas formed under marine terminoglacial conditions. The clasts show a wide

546

The marine extraglacial subenvironment

range of properties, particularlyrounding and flattening. Flattened clasts have a preferred orientation (suh)parallel to the sedimentary surface, but may be imbricated locally. Glacial striae are usually present on some of the larger clasts.

Occurrence The complexes are found in a n intermediate position between glacigenic continental and marine deposits. They interfinger with bottom-current deposits, mass-transport deposits and deposits of non-glacigenic origin. Only few of the complexes are more than ten metres thick. The lateral extent may be up to a kilometre and the width is comparable.

Depositional mechanisms Avalanching of material along the delta slope is the main process that contributes t o accumulation. Deposition in channels incised in the part that might be considered as a precursor of topsets also plays a role, a s does infilling of channels in the 'foreset' part. Mass-transport processes usually make a large contribution to t h e accumulation of these deltas, but layers formed this way could also be considered as belonging t o the marine extraglacial mass-transport facies.

Marine extraglacial deltaic topsets (II-F-3-b) These deposits consist mainly of channelised bodies of sand and gravel, sometimes with intercalations of diamicts as a result of mass-flow processes. The deltaic channels commonly bifurcate. The sediments between channels may show indications of 'tidal' activity (herringbone structures). The deposits are very similar t o the deltaic topsets that are formed under continental glaciolacustrine or marine terminoglacial conditions (also see Jenness and Spencer Jenness, 1989).

Lithofacies characteristics Most topsets consist predominantly of gravel and sand. The latter may be massive or horizontally stratified, whereas the former occurs concentrated i n channels (Fig. 315), usually without a distinct i n t e r n a l structure hut occasionally showing crude stratification. Coarse-grained deposits tend to contain fewer sedimentary structures t h a n do finergrained units. Horizontal lamination, small-scale and large-scale cross-

Deposits of the marine extraglacial deltaic facies

547

lamination, occasionally accompanied by wave ripples, and flaser bedding are the most frequently found primary structures. Coarse layers overlying muddy ones usually show frequent loading structures. Textural characteristics Bad sorting of the most proximal topsets and the wide variation in grain size are the most obvious textural characteristics. Sorting usually improves with increasing distance from the coast. Most clasts are badly to moderately rounded, although there may also be well rounded clasts. If the clasts are flattened, they tend t o be orientated (sub)parallel to the bedding plane. Imbrication may, however, be found locally ,at the base of coarse layers. Occurrence The topsets, which do not necessarily extend much more in a longitudinal t h a n in a transverse section, interfinger in an upstream direction with

Fig. 315. Channel fill in possibly extraglacial deltaic topsets (Cantley quarry, North of Ottawa, Canada).

548

The marine extraglacial subenvironment

continental (usually fluvial) deposits and in a downstream direction with marine extraglacial bottom-current deposits or, sometimes, with marine extraglacial mass-transport deposits. Lateral contacts are also with bottom-current and mass-transport deposits In the glacigenic sequence, these deposits form the topmost part of marine extraglacial deltas, which are found relatively often in the marine extraglacial interval.

Depositional mechanisms

No mechanisms other than those active in deltaic topsets formed in any (sub)environment are involved. The meltwater streams entering the sea are, however, usually less cold (and thus less heavy) than those released immediately from an ice cap. Consequently, there will be relatively few underflows, so that settling from suspension (from overflows) may play a relatively important role.

Marine extraglacial deltaic foresets (II-F-3-c) These foresets develop best where the stream that embouches in the sea does not frequently shift its position. In this case there is a continuous supply of material, most of which accumulates on top of previous foreset layers. Rapid lateral changes in lithology, the presence of frequent channels, the inclined foreset stratification, the intercalations of slumps and associated deposits, etc. are all characteristic aspects. Plant remains, while very common in moderate or tropical deltaic foresets, are absent o r extremely scarce in deltaic foresets formed under these conditions.

Lithofacies characteristics Gravellykandy associations are the most common but there can also be finer-grained (predominantly sandy) examples if the continental stream embouching in the sea has travelled such a long distance that most coarse material was left behind on the continent. The beds show a distinct inclination, with planar interfaces, but this becomes almost nil at the lowermost parts. Isolated boulders, which may be considered as a type of lag deposit, can sometimes be found at the base of the foresets (Fig. 316). Individual foreset beds tend t o be fairly well sorted, but intercalations of mass-transported layers may result in an overall diamict-like appearance of the foreset body.

Deposits of the marine extraglacial deltaic facies

549

Textural characteristics

Sorting usually depends on the average grain size: fine-grained foresets tend t o be better sorted than coarser ones. Sorting may be very good within a single layer, and upward grading is frequent in individual layers. There are some reports that rounding of the clasts is somewhat better than in the foresets of marine terminoglacial deltas. This might be a result of the much longer fluvial transport that the clasts underwent, prior t o deposition at the final deltaic site. Flattened clasts show a clear preferential orientation, i.e. with the ah-planes parallel to the (inclined) sedimentary surface. Occurrence

Marine extraglacial deltas form upward coarsening sequences; the foresets constitute the middle part of such a sequence. There usually is gradual but relatively distinct contact with the finer-grained bottomsets underneath; contact with the coarser, overlying topsets tends to be much

Fig. 316. Sandy marine extraglacial deltaic foresets with pebble concentrations at the base of individual units (Cantley quarry, Ontario, Canada).

550

The marine extraglacial subenvironment

more transitional. Apart from these contacts, which are also present in lateral (upstream and downstream) directions, the foresets interfinger with mass-transported sediments and with bottom-current deposits. There is only rarely direct contact with non-glacigenic sediments. These deposits frequently reach thicknesses of several tens of metres; the width of these bodies may be a few hundred metres, occasionally a few kilometres, and the length is of the same order. Depositional mechanisms The most important depositional mechanism, apart from the activity of mass flows, is avalanching of material supplied by the meltwater streams in the topsets. Deposition in channels may also make a significant distribution in some deltas. Settling from suspension also plays a role; this role is more important here than in marine terminoglacial deltas as more overflows are present (as explained above) so that there can be a relatively large quantity of suspended particles. M a r i n e extraglacial deltaic bottomsets (11-F-3-d) The deltaic bottomsets in the marine extraglacial subenvironment are the finer-grained equivalent of the foresets; they are deposited on a generally less inclined sedimentary surface and further away from the source. They consequently appeer more regular and continuous, although this appearance is generally less evident than with their counterparts formed under glaciolacustrine conditions. Lithofacies characteristics Deltaic bottomsets are similar, whether formed under glacial or under non-glacial conditions. They consist predominantly of thinly bedded, massive, graded and rippled sands, alternating with laminated silty clays (Fig. 317). The most common sedimentary structures are current ripples, current-induced horizontal lamination, convolutions and water-escape structures. Grading, climbing ripples, and burrows a r e usually also present but are less abundant. Many of the layers have lost their original internal structure as a result of soft-sediment deformations, but the original layering tends to remain easily detectable.

Deposits of the marine extraglacial deltaic facies

55 1

Textural characteristics Traction currents and settling from suspension account for a usually good sorting of these deposits. However, intercalations of mass-transport deposits may disturb this overall picture. If these intercalated mass-transport deposits contain pebbles, there will be either a preferred orientation parallel to the flow lines or a subhorizontal orientation. Imbrication at the base of such flows (cf. Van Loon, 1970) is rare. Occurrence Marine extraglacial deltaic bottomsets are found in a wide zone in front of the main delta body. The transition, in an off-delta direction, into nonglacigenic sediments is very gradual. Interfingering with mass-transport deposits is common in all directions. These bottomsets are found as the lowermost part of deltaic sequences. The thickness is usually difficult to determine exactly because of the

Fig. 317. Fine-grained, laminated marine extraglacial deltaic bottomsets (Cantley quarry, near Ottawa, Canada).

552

The marine extraglacial subenvironment

frequent intercalations of mass-transport deposits and particularly because of the gradual transition into non-glacigenic marine sediments. If the relatively low sedimentation rate (far from the source) is taken into account, a thickness of over ten metres may probably be considered exceptional.

Depositional mechanisms While traction currents can appear to play a large role in the accumulation of these deposits, most of the sediments may be due to settling from suspension of particles supplied by over- or interflows. Poorly sorted sands a r e relatively rare, but if present, may be due t o underflow currents. DEPOSITS OF THE MARINE EXTRAGLACIAL MASS-TRANSPORT FACIES (11-F-6) This facies comprises, in analogy to facies II-D-6, only one type of deposit: the marine extraglacial mass-transport deposits (II-F-6-c).The processes involved and the deposits thus formed are not essentially different from those formed under non-glacigenic marine conditions, as can be deduced from a number of descriptions (Bjorlykke, 1967; Van Loon, 1970, 1972; Elverhoi et al., 1980; Fralick and Miall, 1989) and models (C.H. Eyles et al., 1985).

Marine extraglacial mass-transportdeposits (11-F-6-c) The deposits grouped in this class were transported over the seabottom as slumps, slides, mudflows and turbidity currents. Rapid lateral variations in the characteristics of the deposits are a common feature because the transport mechanism usually changes during the mass-movement process itself, mainly as a result of mixing with sea water, but also as a result in changes in the inclination of the slope (Domack, 1983; McCabe and Eyles, 1989a,b).

Lithofacies characteristics The mass-transport deposits in this subenvironment may be coarse, with a diamictic character, particularly in the neighbourhood of a delta front. They show a general tendency t o become gradually finer in more distal

Deposits of the marine extraglacial mass-transport facies

553

directions and the deposits found furthest away from the coast consist of fine-grained turbidites (laminites) that contain only the E-interval of the Bouma sequence. Textural characteristics

No textural characteristics can be singled out as several depositional mechanisms are involved in the genesis of these sediments. A t the foot of delta slopes, where these deposits are most abundant, sorting is usually bad because, slumping and sliding being the main processes, grain size could not be differentiated. Sorting generally increases with distance, as more water becomes mixed with the flow, changing the flow into highdensity and later even into low-density flows with a turbiditic character (Wright and Anderson, 1982). Vertical grading is common in such turbidites. Where there are slumping, mudflow and comparable transport processes, the orientation of the larger clasts in these deposits tends t o be parallel t o the flow lines (cf. Van Loon, 1983). The coarser particles that can be found at the base of proximal turbidites usually have their a h planes parallel t o the bedding plane. Occurrence

These deposits cover most of the seabottom exactly as they do in the marine terminoglacial subenvironment, and occasionally extend t o the abyssal plain. Deltas, and to some degree continental slopes also, favour the genesis of these sediments. The deposits are found to interfinger with bottom-current deposits and marine deposits of non-glacigenic origin. Depositional mechanisms

All types of subaqueous mass-transport mechanisms are active under marine extraglacial conditions, e.g., slumps and slides, cohesive debris flows, cohesionless debris flows, cohesive turbidity currents and 'normal' turbidity currents. Turbidites (Fig. 318) probably account for the most widespread deposits. Cohesive turbidity currents are also involved; their flow is turbulent, except in the final stage. The low strength of these currents allows sand grains to settle before clay particles do. Cohesive debris flows have a laminar character usually following on fully sheared laminar or turbulent flow, although they may also follow flows that were not sheared

554

The marine extraglacial subenvironment

Fig. 318. Marine extraglacial turbidite intercalated between marine extraglacial bottom-current deposits of Huronian age (road section near Cobalt, Ontario, Canada).

(e.g., slides or slumps). The flows contain a high concentration of particles and have been described from several glacigenic sites (Postma et al., 1983; C.H. Eyles, 1986).

Epilogue

555

EPILOGUE The authors had as their main objective the presentation of a systematic overview of glacigenic sediments in the context of both the depositional processes involved and the dynamics of glaciation. It is a disadvantage of this approach that, in general, only broadly applicable data can be summarised, leaving some aspects t o receive far too little attention. Still, had the analysis been more detailed, the study would have lost much of its practical applicability. It does seem justified to conclude that the glacigenic subenvironments are all so complex that they deserve far more detailed analysis. This would allow individual subenvironments to be given due attention. A more in-depth analysis of each subenvironment would have the advantage that specific problems that are outside the scope of this book could be highlighted. One of the problems regards the nature of the proglacial subenvironment. This subenvironment is usually defined in a rather vague way: 'the zone in front of glacial ice'. It is important, however, from several points of view, to differentiate between areas with and without buried dead-ice. The fact that melting of such dead-ice bodies provides meltwater, induces surface collapse (thus influencing the drainage and depositional patterns at the surface) and makes the subsoil unstable gives the differentiation a strong sedimentological, geomorphological and geotechnical interest. The present authors consider these aspects so important that the periglacial area characterised by the presence of buried dead-ice is taken here as a separate subenvironment: the proglacial subenvironment (which is of no importance under marine conditions: see the discussion on p. 533535). It should be ralised, however, t h a t dead-ice blocks can only be present if the ice front has retreated. In addition, insufficient time should have elapsed t o allow the dead-ice blocks to melt during the time-span of climatic amelioration that induced ice retreat. A succession of ice retreats and ice re-advances thus results in the local and temporary presence of a proglacial subenvironment, depending on the changes in areal extent of the ice mass with time (Fig. 319). The consequence of this approach is that the continental terminoglacial subenvironment (which frequently has a width equal t o the diameter of the largest lake directly touching the ice front: Fig. 320-A),

Epilogue

556

horizontal extent Fig. 319. Presence of the proglacial subenvironment (area with inclined lines) as a result of changes in ice extent that occurred with time. glacial environment

periglacial environment i

A

'lr

terminoglacial { subenvironment

I

\' I I

extraglacial subenvironment

\I I I

!

glacial environment >

periglacial environment

-I

I'

-? ,

\ II

terminoglacial subenvironment

extraglacial subenvironment

glacial environment

periglacial environment A

A

$Ginoglacial subenvironment ! ~'

-,,\

,

~

I

t

proglacial subenvironment 7

2

extraglacial subenvironment

4w

dead-ice blocks

Fig. 320. Models showing situations without a proglacial subenvironment (A, B) and with a proglacial subenvironment (C).

Epilogue

557

can pass directly into the continental extraglacial subenvironment (Fig. 320-A and 320-B), and thus there is no proglacial subenvironment whatsoever. However, a very wide proglacial belt may be present (Fig. 320-C) if the ice retreats rapidly or if the temperature remains so low during a slow retreat of the ice that buried dead-ice blocks hardly melt away. The problem of the nature and significance of the proglacial subenvironment may thus be solved, but a number of other - at least equally important - questions do remain to be answered and even some simple models, long considered satisfactory, now are found to need revision. This is yet another example of increasingly sophisticated methods evidencing the complexity of nature. The work presented here was based not only on field work but also on an extensive search of the literature. The findings were a disappointment in some respects: it became obvious that numerous researchers (doubtlessly, the present authors included) have interpreted each other's results incorrectly. This is partly due to incomplete descriptions, or to the application of terminology and/or symbols that are insufficiently unequivocal. A purely descriptive lithofacies analysis may, for instance, be completely accurate, but does not necessarily help with interpretation of t h e geological record. While the record of a succession of lithofacies units such as t h a t on the left side of Figure 321 - compiled on the basis of proposals by Miall (1977, 1878, 1883b, 1985) and N. Eyles (198310, 1985, 1987) - is extremely useful as a n inventory it says little about the genesis of the succession. Had lateral contacts and other data also been taken into account, specific types of deposits would have been identified (Fig. 321, middle part). The pertinent geological history would then be open to interpretation (Fig. 321, right side). The present authors therefore plead for communication, not only by means of descriptive symbols, but also on a genetic level. An important cause of misinterpretation of literature data seems language problems: western and eastern 'schools' have developed in fairly different directions, presumably because one school wa s insufficiently aware of the results the other obtained. Nomenclature, however, seems a n equally serious problem but it is hoped that this book will contribute t o a less ambiguous terminology. Finally, it seems that different interests have led t o the development and application of diverging research methods. Results obtained with different methods and under different conditions tend t o escape comparison. It is obvious, for instance, t h a t research in currently glaciated areas is quite different from that carried out in Pleistocene areas; Precambrian glacigenic sediments are investi-

558

Epilogue

gated with still other methods and purposes. The authors are convinced that the problems that still exist in this field will only be solved with a more integrated approach. Joint field work, as practised by the authors in - among others - Poland and The Netherlands, is a move i n this direction. Glacial geology will thus most probably advance enormously as a consequence of the recent political changes in Eastern Europe. The literature survey showed that three main aspects are important for a further study of glacigenic sediments. In the first place, more information is needed about the depositional processes and the resultant (net) deposits. For instance, the factors that determine the preservational potential of a specific type of deposit should be better known, as should the reason why the preservational potential of several types of deposits seems so different for deposits from different glaciations. A second aspect is the inventory of lithological and textural characteristics of the various types of deposits. These factors, which include sedi-

I-A-2-e I-A-2-d

t

area becomes covered by supraglacial fluvial deposits after melting of the ice

I-A-6-a ~

diamicts are deposited during glaciation of the area, with deposition of successively subglacial, englacial and supraglacial tills

I-B-1-a

I-C-1-a

i

_.____

ll-A-4-c

deposition in a lake directly in front of the ice

_ . _ ~ ~ _

ll-C-6-a

extraglacial fluvial sedimentation, gradually becoming more proximal

ll-C-2-d

4

non-glacigenic fluvial deposits

'

erosion surface

fluvial sedimentation not influenced by remote ice sheet

Fig. 321. Hypothetical succession of glacigenic units with symbols indicating the lithofacies (left) and the type of deposit (middle), with a reconstruction of the geological history (right).

Epilogue

559

mentary structures and grain-size compositions, have only been studied in a more or less random way, partly depending on the 'school' to which the researcher belonged. However, lithology and texture, together with facies transitions, are the basis on which genesis must be interpreted and palaeogeography must be reconstructed. Making a truly systematic inventory of these parameters is perhaps time-consuming and not of overwhelming interest, but progress in glacial geology will be held op until a n inventory of this type is made. The third aspect in need of study is modelling of the various environments, subenvironments and facies. Models abound in the literature, but each one is nearly always applicable to one specific site only. Comparison of several models is usually impossible because the various researchers focused on questions of specific rather than general interest. Only few attempts have been made to compile 'composite' models that might be used a s base for comparative studies, and only a few of these have been more o r less successful. For this reason the properties necessary as basis for geological models in general, and for models regarding glacigenic geology in particular are points that need discussion. It is interesting that both the continental glacial and the continental periglacial environment comprise one facies that is generally considered as the most characteristic. This is the subglacial melting-ice facies in the glacial environment, and the extraglacial aeolian facies in the periglacial environment. Lodgement tills and loesses are perhaps the most widely known types of deposits representing these two facies. One might suppose that the two 'characteristic' facies had been studied in most detail and would be best known. Numerous studies have indeed been devoted to these facies, the processes active there, and the resulting deposits. However, the literature survey shows that hard facts are relatively scarce and t h a t many of the interpretations a r e still higly controversial. Surprisingly, there is much less controversy regarding t h e fluvial, deltaic, lacustrine and mass-transport facies of glacigenic origin. A concerted effort by Quaternary geologists, glaciologists, gemorphologists and sedimentologists should yield a satisfactory solution for this apparent discrepancy.

This Page Intentionally Left Blank

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Subject index

669

SUBJECT INDEX aasar 10 ablation end moraines 143 ablation submarine moraine 352 ablation till 99,143,285 ablation till in situ 285 abrasion 245 acoustic facies 502 active-layer gliding 349,449,498 adhesion structures 429 aeolian complexes 119,424,463 aeolian facies 96,116,187,341,423,463 aeolisation 342 alternate bars 376 areal extent 50 Biblical Flood 3 bar 372 barrier spit complexes 332 basal tills 99,100,232,246,285 baseinova moraine 352 beach face 332 bedding characteristics 77 berm ridges 332 berm tops 332 biofacies 47 blockfall 345,349,449,498 bottom-current deposits 513,539 bottom-current facies 96,274,508,539 boulder clay 3,29 braid deltas 312 buried dead-ice 284 Cambrian glacigenic sediments 49 Carboniferous glacigenic sediments 49 catastrophic sedimentation 61 chronology 49,340 chute bars 379 classification 66 cliff-head dunes 431 climatic fluctuations 31 climbing ripples 56 coding 68,81 cohesionless debris flow 358

cohesive turbidity currents 357 229 cold thermal regime 46,137,284 collapse collapse structures 345 comminution till 232,235 complex supraglacial moraines 139 composition 70 compound paratills 543 contact characteristics 51 contour currents 540 coversands 24,119,120,437,475 creep 349 crevasse deposits 194,218,349 crevasses 121,122,218,349 377 cross-channel bars cross-valley moraines 511 crushing 223 cryoturbation 54,459 cyclic sedimentation 116,403 cyclites 63 cyclonal circulation 36 cyclothems 63 De Geer moraines 511 Diluvium 3 Dnieprovian 39 dead-ice 91,363 dead-ice blocks 134 dead-ice bodies 134 debris flow 349,449 deflation areas 342 deflation pits 430 deformation structures 53 deformation till 232,240,246 deltaic bottomsets 110,174,321,403,523,550 deltaic complexes 110,165,257,313,392,517,550 deltaic facies 164,257,312,391,517,544 deltaic foresets 110,170,318,400,521,548

670

Subject index

deltaic topsets 110,167,316,395,519,546 deltaic varves 404 depositional mechanism 60,67 desert erg 433 diagenetic till 232 diagonal bars 377 diamict 5 diamictite 5,96 diamicton 5 diamond bars 377 dislodged till 232,263 double-source turbidites 114 draping structures 210,211 drift 4 drift sands 119,187,342,427,466 dropstones 150,288 drumlin fields 7 drumlins 7,226 dump deposits 150 dumpstones 505 dune fields 43 1 dunes 119,120,431,469 earthquakes 41,43 endogenic factors 41 energy balance 33,83 a3 energy exchange energy system 29 englacial melt-out complexes 285 englacial subenvironment 203 englacial tunnels 213 erosional surfaces 465 erosive contacts 51 erratics 3 esker beads 299,511 eskers 10,103,296,304 exaration till 240,246 exceptional conditions 61 extraglacial subenvironment 459 fabric 57 facies 47,67 facies analysis 48 facies associations 58 facies interpretation 50 facies models 64 fan deltas 312,392,517 fans 110

fissuring flocculation flow lines flow till

260 358 355 125,190,198,265, 285,288,328,347,352 fluidisation structures 420 fluidised flow 345,358 fluted moraines 8,226 flutes 8 fluvial complexes 105,152,293,373 fluvial facies 150,213,252,293,367 fluvioglacial deposits 22 foliation 285 frontal crevasses 124,350 frost creep 346,349,449,496 full-pipe flow 296,217,253,255 Gilbert-type deltas 108,312 Gilbert-type fan deltas 392 gelifluction 346,349,449,497 geochemical facies 47 geometry 50 geothermal heat flux 42,45,83,197 glacial diamict 96 glacial environment 66,87 glacial lakes 22 glacial sequence 93 glacial striae 56,227 glacier-contactzone 281 glacier table 135 glacifluvial deposits 22 glacigenic facies 93 glacigenic sedimentary environments 83 glaciodeltaic facies 96,107 glaciofluvial braid plains 367 glaciofluvial facies 96,101 glaciofluvial fans 367 glaciolxustrine deposits 22,96 glaciolacustrine facies 111 glaciomarginal zone 281 glaciomarine sediments 267 glaciotectonic rafts 229 glaciotectonism 15 grain flow 358 grain size 29,52,68 grain surface 57 gravitation 83 grezes litees 492,493

Subject index

grinding 223 ground-ice slump 349 grounding ice 268 grounding line 277 grounding structures 505 Hjulstrom-type deltas 165,108,312 hematite coatings 445 high-energy conditions 311 hollow roads 444 horizontal alternating stratification 476 horizontal joints 209 hummocky cross-stratification 332,413 hyperconcentrated flow 341 ice-contact zone 281 ice dynamics 31 ice front 301 ice gouging 325 ice-marginal streamways 12,305 ice-marginal zone 281 ice-pushed ridges 330,351 ice-raft deposits 20,99,100,148,249,288,504

ice rafting 270 ice retreat 301 ice shelves 267,499 ice-shove ridges 414,415,416 ice-walled shores 499 impact crater 272 incorporation till 240 intrabasin breccia 329 intraflood transistent hydraulic surging 421 iron-pan layers 477 isostatic movements 31 joints 209 jokulhlaups 326 kaim 9 kame terraces 10,178 kames 9,178,194,304 laboratory experiments 65 lacustrine bottomsets 115,184,333,417 lacustrine complexes 115,179,261,327,409 lacustrine facies 177,261,324,408 lacustrotills 328,352 lag deposits 378 lake ice 330

671

lake-margin deposits 115, 181,330,413 landforms 5 landslides 449 lateral bars 376 lee-side till 232,263,246 liquefied flow 358 literature 3 lithofacies 47,67 load casting 54 local till 240,246 lodgement 229,231 lodgement till 99,100,232,237 loess belt 462,480 loesses 24,119,120,443,480 loessites 480 loessoid 482 longitudinal bars 376 longitudinal dunes 47 1 lowered till 139,144 marginal zone 281 marine bottom-current facies 106 marine channel deposits 105 marine glacial environment 267 marine proglacial subenvironment 533 marine terminoglacial subenvironment 499 mass movements 25 mass-transport facies 96,120,189, 218,263,278,344,447,490,526,552 meander bar 378 melting-ice facies 96,138,207,231,270,284,502 melt-ou t complexes 99,138,208,234,271,285,365 melt-out facies 96 melt out subglacial till 232 melt out till 139,144 meltwater deposits 22 meltwater streams 21,101,282 meltwater surges 277 meltwater-tunnel deposits 214,254 meltwater tunne!s 10 micro-scale turbidites 407 mineralogy 52 modified grain flow 358 moraines 6 morainization till 232

672

Subject index

mudflow 349,449 multiple retrogressive flow 349,449,498 multiple retrogressive sliding 349,449,498 needle-ice creep 349,449,497 niveo-aeolian deposits 24,436,478 13 nomenclature non-erosive contacts 51 normal sedimentation 61 nunataks 343 Ordovician glacigenic sediments 49 orientation 57 osar 10 outwash fields 293 outwash plains 293,368 overconsolidation 260 Permian glacigenic sediments 49 Precambrian glacigenic sediments 49 palaeocurrents 52,55 palaeogeography 52 palaeosoils 465 palatills 352 para-autochthonic flow till 285 parabolic dunes 431,471,473 paraesker bars 299,300,302 paralodgement tills 237 periglacial environment 66,91 permafrost 91,459 permafrost creep 346,249,449,496 petrology 52 pipe flow 213 plane-bed adhesion lamination 478 pointbars 376 potential energy 83 pradolinas 12,305,363 precipitation 36 preservational potential 30,58 primary tills 233 proglacial subenvironment 361 Quaternary 19 quasi-planar adhesion stratification 478 rain-out 271 rain-out diamicts 288 relict plug flow 358 repeating sediments 64 reversed density gradients 272 rheological characteristics 357

rhythmites 63,334,404 rigid plug flow 496 ripple-drift cross-lamination 56 rock channels 226 rockfall 349,449,498 rogen moraines 226 roundness 57 Salisbury-type deltas 108,168,312 sand drifts 342 sand flows 332 sanders 11,293,368 sandrs 11 sandurs 11 sea-ice 267,502 31 sea-level changes secondary tills 233,263 sedimentary breaks 51 sedimentary cycles 63 sedimentary facies 47,67 sedimentary structures 53,72 sediment input 32 sequences 59 shadow dunes 437 shear till 232,240,263 shear zone till 232 sheet- and streamflood deposits 105,161,308,385 sheetfloods 308 side bars 376 silt-sized particles 486 skinflow 349,449,498 sliding 349 sliding bed conditions 217 sliding bed facies 296 slippage 349 slope loesses 495 slope wash 349 slumping 349,449 slush ice 267 snow storms 24 solar irradiation 83 solifluction 349 solifluction lobe 448 source area 52 squeeze flow till 263 squeeze till 232,263 stone counts 233

Subject index

stone pavement 424 stoss-side till 232,238,246 stratified drift 4 stream deposits 105,157,305,376 streamfloods 308 strudel scours 267 subaerial debris flow 349 subaerial mass movements 26 subaerial mass-transport deposits 190,346,448,491 subaerial mass-transport processes 121 subaquatic flow till 352 subaqueous debris flows 278 subaqueous fan deltas 108 subaqueous fans 106 subaqueous floods 341 subaqueous glacilacustrine fan deposits 29 1 subaqueous high floods 105 subaqueous mass movements 26 subaqueous mass-transport deposits 198,278,352,453,526,552 subaqueous mass-transport processes 121 subaqueous sanders 275,291,299,300,302,516 subaqueous sedimentation 21 subaqueous sheetflood 403 subaqueous slides 278 subaqueous slumps 278 subenvironments 66 subglacial channels 254 subglacial chaotic debris flows 265 subglacial debris flows 265 subglacial flow till 232,263 subglacial mass-transport deposits 264 subglacial melt-out diamicts 250 subglacial slurry flows 265 subglacial subenvironment 223 subglacial undermelt diamicts 250 sublimation till 139,144,148 sublittoral deposition 332 subsidence 41 subsurface wash 349 supercooling 267 supra-aqueous deltas 108 supraglacial ablation drift 139 supraglacial moraines 132

673

supraglacial morainic drift 139 supraglacial subenvironment 131 surface wash 349 suspension-deposition stage 358 sveg till 286 swaley cross-stratification 413 swash bars 332 symbols 65 Tertiary glacigenic sediments 49 temperate glaciers 204 temperature 33 terminoglacial subenvironment 281 terminology 3,65 textural characteristics 57 thaw slumping 491 thermal expansion 330 thermal regime 34 thermosubrosion 97,207 thickness 50 tidewater glacier tongues 268,499 till 5,20,96 tillite 5,96 topographic deceleration 44 1 traction carpet 358 traction-carpet stage 358 traction-current deposits 539 transitional zone 281 transverse dunes 43 1 transverse foredunes 43 1 true kames 9 tunnel deposits 105 tunnel flow 217,253 tunnel-mouth deposits 105,154,274,296,510 tunnels 122 turbidite fans 421 turbidites 340 turbiditic fans 340 turbidity currents 114,278,339 typical loess 481 U-shaped valley 1,227 undermelt diamicton 232 undermelting 98,288 undermelt till 232 upheaval 41 uppermost proximal fan 510 varves 24,114,175,184,326,339,334

674

Subject index

vertical joints 209 warm glaciers 204 waterlaid tills 198,352 waterlain melt-out diamicts 250 waterlain tills 198,260,264,352 watermorainic deposits 264,352 water-pressure regime 252 weathered clasts 57 wet deposition 488 winds 40 xenoliths 288