Landforms of Cold Climates

J.L.
Davies
An Introduction to Systematic G eo m o r p h o Io g y Volume Three

T h is is a n o th er volum e in th e series, An
In tro d u ctio n to System atic G eom orphology.
It is concerned w ith the landscapes produced
w here w ater exists com m only in solid form
— as ground ice, as snow, or as glacial ice.
A lthough th e presen t d istrib u tio n of glaciers,
snow banks, a n d frozen gro u n d w ater is re la
tively lim ited, these p h enom ena were m uch
m ore extensively d istrib u te d d u rin g the
Pleistocene ice ages and they have left their
m ark on the landscapes of alm ost all parts
of th e tem p erate w orld. It is im possible to
u n d e rstan d th e landscapes of m uch of so u th
eastern A ustralia, New Zealand, E urope, and
n o rth e rn N o rth A m erica w ith o u t taking in to
consideration th e p a rts played by glacial and
periglacial processes d u rin g the Pleistocene.
Since W orld W ar II th ere has been a

g reat upsurge of interest in the phenom ena
associate w ith ice sheets, th e tu n d ra lands,
a n d h igh m o u n ta in areas. M uch of the
work carried o u t has been inspired by the
difficulties of p lan n in g a n d executing engi
neerin g works in regions w here snow and
ice are p revalent, a n d some of th e results
of th is recent work in the Arctic and
A ntarctic are in co rp o rated in this volum e.
W ell illu stra te d w ith half-tone plates,
m aps, a n d diagram s, L a n d fo rm s o f Cold
Clim ates has been designed p a rticu la rly for
schools and universities, b u t should interest
a m uch w ider audience.
Jacket design by R o b in Wallace-Crabbe
P rin te d in Australia
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An Introduction to Systematic Geomorphology
VOLUME THREE
LANDFORMS
OF COLD CLIMATES
J. L. DAVIES
Reader in Geography
University of Tasmania
1969
AUSTRALIA N N A T IO N A L U N IV ER SITY PRESS

CANBERRA
First published 1969
Printed in Australia
This book is copyright in all countries subscribing
to the Berne Convention. Reproduction, in whole or
in part, without written permission of the publishers,
is forbidden.
Registered in Australia for transmission
by post as a book.
SBN 7081 0155 0

Library of Congress Catalog Card no. 69-12378
National Library of Australia reg. no. AUS 68-1293
I N T R O D U C T I O N TO THE SERIES
T h is series is conceived as a systematic geomorphology at

university level. It will have a role also in high school
education and it is hoped the books will appeal as well to
m any in the com m unity at large who find an interest in
the why and wherefore of the natural scenery around them.
T h e point of view adopted by the authors is that the
central them es of geom orphology are the characterisation,
origin, and evolution of landform s. T h e study of processes
that make landscapes is properly a part of geomorphology,
b u t w ithin the present fram ew ork process will be dealt with
only in so far as it elucidates the natu re and history of the
landform s u n d er discussion. C ertain other fields such as
subm arine geomorphology and a survey of general principles
and m ethods are also not covered in the volumes as yet
planned. Some knowledge of the elem ents of geology is
presum ed.
F our volumes will approach landform s as parts of systems
in which the interacting processes are alm ost completely
m otored by solar energy. In hum id climates (Volume One)
rivers dom inate the systems. Fluvial action, operating
differently in some ways, is largely responsible for the land
scapes of deserts and savanas also (Volume Tw o), though
winds can become preponderant in some deserts. In cold
climates, snow, glacier ice, and ground ice come to the fore
in m orphogenesis (Volume T hree) . O n coasts (Volume
Four) , waves, currents, and w ind are the prim e agents in
the com plex of processes fashioning the edge of the land.
T h re e fu rth er volumes will consider the parts played
passively by the attributes of the e a rth ’s crust and actively
by processes deriving energy from its interior. U nder
structural landform s (Volume F iv e ), features im m ediately
consequent on earth m ovem ents and those resulting from
tectonic and lithologic guidance of denudation are con
sidered. Landform s directly the product of volcanic activity
v
VI Introduction to the Series
and those created by erosion w orking on volcanic m aterials
are sufficiently distinctive to w arrant separate treatm ent
(Volume Six) . T h o u g h karst is undoubtedly delim ited
lithologically, it is fashioned by a special com bination of
processes centred on solution so that the seventh volume
partakes also of the character of the first group of volumes.
J. N. Jennings
General Editor
PREFACE
Like others in this series, the present book is intended

m ainly as a text at university level, b u t it is hoped that it
may be of use in advanced classes in schools and also to
anyone who is interested in landscapes at high altitudes and
high latitudes. A lthough w ritten prim arily with A ustralian
university students in m ind and m aking use of many
examples and illustrations from southeastern A ustralia and
New Zealand, I have attem pted to give a balanced world
picture.
I would like to thank the following for contributing
photographs used in the plates: M r V. C. Browne, Dr A. B.
Costin, Dr J. D. Ives, Professor J. Ross Mackay, M r J. A.
Peterson, Professor T roy L. Pewe, M r J. G. Speight, the
Surveyor-General of T asm ania, the T asm anian D epartm ent
of Film Production, the A ntarctic Division of the A ustralian
D epartm ent of Supply, the New Zealand Geological Survey,
the New Zealand N ational Publicity Studios, and Aerofilms
Ltd, London. I am also indebted to M r G. van de Geer
who took some of the photographs and drew or redrew all
the line illustrations. Most of all I wish to acknowledge
my debt to the editor of the series, M r J. N. Jennings, for
his encouragem ent and constructive criticism at all stages.
J. L. D.
vii
CONTENTS
page
Introduction to the Series, J. N. Jennings v
Preface vii
I INTRODUCTION 1
Glacial climates 3
The proglacial zone 7
Periglacial climates 9
Nivation climates 15
Comparison of glacial and periglacial limits 16
II PERIGLACIAL PROCESSES 18
Ground ice 18
Frost sorting 21
Weathering processes 24
Mass movement processes 28
Fluvial processes 35
Aeolian processes 37
Cryoplanation 38
III MASS MOVEMENT LANDFORMS 40

Slope forms 40
Tor forms 47
Terrace forms 49
IV OTHER PERIGLACIAL LANDFORMS 54

Mound-like forms 54
Thermokarst 60
Tundra lakes 62
Tundra bogs 63
Valley forms 65
Aeolian forms 69
IX
x Contents
page
V NIVATION PROCESSES ANDLANDFORMS 70
Snow-melt processes 70
Snow movement processes 73
Depositional features 75
Nivation landscapes 77
VI GLACIERS 78
Glacier forms 79
Glacier economies 83
Glacier movement 88
Glacier types 96
Glacier types and glacial geomorphology 98
VII GLACIAL PROCESSES 101
Corrasion 102
Transportation 108
Deposition 111
Meltwater processes 116
VIII PROGLACIAL PROCESSES 117
Glacilluvial processes 117
Drainage derangement 122
Aeolian processes 127
IX GLACIATED MOUNTAIN INTERFLUVES 131
Cirques 131
Upland dissection by cirques 141
Glacial cols 145
The end product of mountainglaciation 147
Ice sheet glaciation of mountains 148
X GLACIATED MOUNTAIN VALLEYS 149
Cross profile of troughs 149
Long profile of troughs 156
XI GLACIATED PLAINS 168
Zones of erosion and deposition 168
Erosional landforms 169
Depositional landforms 173
XII CONCLUSION 183
Bibliography 186
Index 193
FIGURES
page
1 Present-day cold m orphogenic zones of th e n o rth e rn hem isphere 2
2 C old m orphogenic zones of th e n o rth e rn hem isphere d u rin g the
last Pleistocene glaciation 2
3 A ltitu d e of th e snow line d u rin g the late Pleistocene glaciation
of T asm ania 4
4 A pproxim ate variatio n of th e regional snow line w ith la titu d e 6
5 D iagram m atic representation of the influence of aspect on the
h e ig h t of the orographic snow line in the southern hem isphere 7
6 T h e proglacial zone in central E urope d u rin g the last glaciation 8
7 P resent-day d istrib u tio n of perm afrost in th e n o rth e rn hem isphere 10
8 Present-day cold m orphogenic zones of th e so u th ern hem isphere 10
9 T h e th ree process systems of cold clim ate landform evolution 16
10 Involutions at the gravel-shale contact in a g round section in
central M ontana 22
11 Sorted stone circles changing to sorted stone stripes on a sm ooth,
northeasterly, convex slope in Spitsbergen 23
12 Four ways in w hich periglacial creep m ay occur 29
13 O rie n ta tio n of stones in turf-banked terraces in th e M t Kosciusko
region of New South W ales 34
14 Spatial variation in orien tatio n of boulders in channelled p e ri
glacial solifluction deposits — the n o rth e rn blockstream of Mt
Barrow , T asm ania 36
15 T h re e types of talus slope 41
16 T h e dam m ing of Lake Fenton, T asm ania, by block glacis 42
17 N a tu ral scale profiles of talus slopes on M t B arrow , T asm ania,
rock glaciers in the Alaska Range, and a blockstream an d blockfield
on M t Barrow 44
18 T h e blockstream s of M t W ellington, T asm an ia 45
19 T h re e locational types of tor 47
20 Two-cycle evolution of tors 48
21 D iagram m atic com parison of three terrace forms 51
22 V ertical sections th ro u g h two earth hum m ocks 55
23 Sections th ro u g h ground-ice m ounds 56
24 D iagram m atic section th ro u g h East G reenland-type pingo 57
25 D iagram m atic represen tatio n of evolution of M ackenzie-type pingo 58
26 Section th ro u g h fossil pingo in the N eth erlan d s 59
27 A beaded stream and therm okarst ravine in Alaska 60
28 E volution of a thaw lake in the zone of tu n d ra ice wedges 61
Xi
xii Figures
page
29 Some tundra lakes in Alaska 62
30 Two surface patterns produced by string bogs 64
31 Evolution of string bogs 65
32 Diagrammatic section across the Tea Tree valley, southeastern
Tasmania, showing two alluvial fills forming terraces 66
33 Three ways in which asymmetrical valleys may be produced 67
34 Four ways in which snowpatches may transport 73
35 Comparison of debris tongues and cones 75
36 Cirque glaciers around Nautgarstind, Aust-Jotunheimen, Norway 79
37 Drangajökoll, a plateau glacier in northwestern Iceland 80
38 A dendritic system of mountain glaciers compared with a transec
tion system 82
39 Section through part of the Antarctic ice sheet at about longitude
95°E 82
40 The Greenland ice sheet 84
41 Diagrammatic representation of the economy of a glacier 86
42 Area-height curves of four valley glaciers 87
43 Firn basins and glacier snouts of two Swiss glaciers 88
44 Calculated isopleths of cross-sectional velocity in feet per year of
the Saskatchewan glacier 89
45 Calculated flowlines and velocity-depth profiles in a cirque glacier,
the Vesl-skautbreen, Norway 90
46 Calculated flowlines and velocity-depth profiles in a valley glacier,
the Saskatchewan Glacier 90
47 Flow zones in a glacier according to Nye (1952) 93
48 Presumed ice flow directions in the cross-section of an idealised
ice sheet 94
49 Stages in the retreat of the Pleistocene Scandinavian ice sheet 100
50 The stoss and lee effect 106
51 Particle size analysis of a Pleistocene till from the upper Mersey
valley, Tasmania 106
52 The movement of surface moraine through a cirque glacier 109
53 Tributary ice streams and associated morainal systems 110
54 Terminus of the Biafo Glacier in the Karakorum Himalaya, with
associated deposits 118
55 Formation of käme terraces 119
56 Meander terraces of the upper Mersey River, Tasmania, formed
by postglacial downcutting into till and outwash 122
57 Subglacial drainage channels cut into till about 70 miles north
west of Schefferville, Quebec 125
58 Proglacial drainage channels in relation to the edge of an ice sheet 126
59 Loess deposits of New Zealand 129
60 Some representative cirque profiles 132
61 Cirques and troughs in the Du Cane Range, Central Tasmania 132
Figures xiii
page
62 T h e D enison R ange in west central T asm an ia 134
63 D iagram m atic relationships of the headw all gap to a ranrlkluft
and bergschrund 136
64 Location of cirques in the Snowy M ountains of New South W ales 142
65 C irques in the A rth u r R ange of southw estern T asm an ia 143
66 T h ree types of glacial col 146
67 P art of F iordland, New Zealand 150
68 Some characteristic features of a glacial trough 151
69 Four circum stances of trough cross profile developm ent 153
70 C om ponents of the long profile of a glacial tro u g h 156
71 Long profile of the Lake Seal-Broad Valley glacial troug h at Mt
Field, T asm ania 157
72 Long profiles of the glaciated South Esk valley and the n e ig h b o u r
ing unglaciated Prosen valley 160
73 T h e Lake Vera rock basin, Frenchm ans Cap N ational Park,
T asm ania 162
74 Lake St Clair, cen tral T asm an ia 164
75 T h e Charles Sound fiord, South Island, New Zealand 166
76 Zones of p re d o m in an t erosion an d p red o m in an t deposition re su lt
ing from the Pleistocene L au ren tid e and Scandinavian ice sheets 169
77 Small rock basin lake on an ice-eroded p lain, C entral Plateau,
T asm ania 171
78 Some stream lined glacial landform s 172
79 Knob and basin topography on an end m oraine n e ar V oltaire,
N orth Dakota, U.S.A. 174
80 Section of a stream lined till p lain w ith d ru m lin s near M iddletow n,
N.Y., U.S.A. 176
81 H airpin-type end m oraine of a sm all Pleistocene glacier n o rth of
M t C anopus, A rth u r Range, southw estern T asm an ia 178
82 Profile of short glacial valley and m in ia tu re outw ash plain, foot
of F ran k lan d R ange, southw estern T asm an ia 180
83 O ne m ode of esker form ation 181
84 D istrib u tio n of Pleistocene cold m orphogenic zones in A ustralia 185
PLATES
facing
page
1 M in iatu re nonsorted polygons a t a b o u t 1500 m etres on th e Ben
L om ond p late au , T asm ania '6
2 Sorted stone circles on a dry lake bed in Alaska 16
3 Section th ro u g h a steep fossil solifluction slope in dolerite on
M t W ellington, T asm ania ^
4 Two-cycle tor n ear T id b in b illa , A ustralian C apital T errito ry 17
5 Edge of fossil blockstream , M t W ellington, T asm ania 32
6 Rock glaciers in the Alaska R ange, Alaska 32
7 Large-scale rock-cut terraces in Alaska 33
8 T u rf-b a n k e d terrace, Snowy M ountains, New South W ales 33
9 S lum ping caused by m elting of ground w ith a high ice content,
n o rth e rn C anada 64
10 Collapsed M ackenzie-type pingo on old lake floor, Mackenzie
delta, northw est C anada 64
11 T h aw p it form ed by m elting of g round ice in loess near Fairbanks,
Alaska 65
12 String bog, Alaska 65
13 Actively developing circular nivation hollow near sum m it of
Frenchm ans Cap, T asm ania 80
14 T ransverse snow patch occupying the rear of a rock-cut altiplana-
tion terrace near the sum m it of M t W ellington, T asm ania 80
15 C om pound talus slope w ith snow-covered rockfall chutes, C entral
P lateau, T asm a n ia 81
16 Fan-type avalanche boulder tongue, Baffin Island 81
17 C irque glacier on Baffin Island 96
18 T h e T asm an G lacier, New Zealand, from Anzac Peak 96
19 Edge of th e A ntarctic ice sheet w ith n unataks n e ar Mawson 97
20 Shelf ice and icebergs on the coast of A ntarctica 97
21 An o u tle t glacier a t the edge of the A ntarctic ice sheet 112
22 F'irnfields of th e Southern Alps, New Z ealand 112
23 Sum m er conditions on glacier tongues descending from a tr a n
section system on Baffin Island 113
24 Crevasse systems an d surface m oraines on the D art Glacier, Otago
Alps, New Z ealand 113
25 A blation m oraine covering the snout of th e T asm an Glacier, New
Z ealand 128
XV
XVI Plates
facing
Page
26 Hummocky surface of käme terrace near Ben Dhu, South Island,
New Zealand 128
27 Valley train and lakes enclosed by end moraines 129
28 Subglacial drainage channels inherited from Pleistocene ice sheet,
central Quebec 129
29 Fretted upland of Federation Peak block, Tasmania 144
30 Scalloped upland of the Snowy Range, Tasmania 144
31 Asymmetrical ridge produced by intersection of cirque headwall and
upland surface, Frenchmans Cap, Tasmania 145
32 Small rock basin cirque on eastern slope of Mt Murchison,
Tasmania 145
33 Mt Aspiring, New Zealand 160
34 Col de la Rousse, an inosculation col in the French Alps 160
35 Stepped, asymmetrical glaciated valley, West Coast Range,
Tasmania 161
36 Milford Sound Fiord, New Zealand, looking seaward 161
37 Vertical air photograph of part of an ice-eroded plain on the
Central Plateau, Tasmania 184
38 Multiple end moraines, central Baffin Island 184
39 Drumlins forming islands in Strangford Lough, Northern Ireland 185
40 Esker ridge in northern Canada 185
I
INTRODUCTION
From a geom orphological point of view the cold lands of

the w orld may be thought of as those where landscape
evolution has been influenced by two im portant groups of
processes to which the adjectives glacial and periglacial are
most generally applied. In the glacial system the m ajor role
is played by glacial ice as an agent of erosion: in the peri
glacial system the m ajor role is played by frost-activated
processes of mass m ovem ent which are able to transport
m aterial over slopes at abnorm ally low angles. In both
systems rock w eathering by frost action is an auxiliary
process of im portance. Realisation of the effect of glaciers,
m oving bodies of perm anent ice, in producing a special and
characteristic suite of landform s, came as early as the first
half of the n in eteen th century, and an account of the rise
and developm ent of the glacial theory is given by Chorley,
D unn, and Beckinsale (1964, ch. 13) . In contrast, periglacial
geom orphology is essentially of tw entieth century origin
and, although some recognition of the potential effect of
frost in encouraging mass m ovem ent came relatively early
(Fisher, 1866), its basic concepts have been form ulated since
1900. In particular the great extension of engineering works
that has taken place since W orld W ar II in the arctic and
subarctic regions of both N orth A m erica and Eurasia has
led to fundam ental advances in our knowledge of the action
and effect of frost processes.
At present the cold lands of the w orld are restricted in
extent (Fig. 1). O nly A ntarctica and G reenland represent
extensive areas being actively glaciated: only the treeless
tun d ra lands of the n o rth ern hem isphere provide examples
of the widespread, co n tinuing evolution of landscapes in
1
2
Introduction 3
periglacial regimes. O utside these areas both systems of land
reduction are confined to m ore or less isolated highland
regions of varying extent and vertical am plitude. However,
there are other parts of the world — extensive areas of
N orth America, Europe, and n o rth ern Asia, smaller areas of
South America, Africa, A ustralia, and New Zealand — which
have been cold lands d u rin g Pleistocene times of clim atic
refrigeration, and here the effects of form er glacial and peri
glacial activity are still discernible in the landscape (Fig.
2) . In these parts of the world, landscape evolution has to
be explained in terms of successive phases dom inated by
different climatically controlled systems. T h u s the T as
m anian highlands and the Snowy M ountains of New South
W ales owe some of their character to a long period of
preglacial denudation (the effects of which in this case are
rather p roblem atical), some of their character to the action
of glacial and periglacial processes in Pleistocene times, and
some to a com paratively short period of postglacial attack
by the processes ordinarily operating in tem perate hum id
climates. Such polygenic landscapes, in which the effects of
one m orphogenic regim e are superim posed on those of
another, probably comprise the great m ajority of world
landscapes. C otton (1958) discussed a particular application
of the concept of alternating m orphogenic regimes in New
Zealand.
Glacial climates
Initiation of the glacial system depends upon the generation
of glacial ice, and this in tu rn depends upon the existence
of perm anent snow. Areas where snow may lie year in and
year out are bounded by the regional or climatic snowline.
In any particidar place they will either be at higher latitudes
or higher altitudes than the snowline. T h e position of the
regional snowline depends essentially upon w inter snowfall,
which m ainly determ ines the am ount of snow accum ulating,
and upon sum m er tem peratures, which determ ine the
am ount of snow wasting or ablating: the line lies where
accum ulation and ablation are equal. In practice the
boundary is m arked by w hat has been term ed the orographic
4 Landforms of Cold Climates
snowline and this is because other factors such as topography,
aspect, and lithology may give rise to locally increased
accum ulation or decreased ablation, and thus cause the zone
of perm anent snow to deviate into lower altitudes or lower
latitudes.
O n a small scale the nature of the regional snowline can
be illustrated from the evidence left by Pleistocene glaciers
in Tasm ania. In Fig. 3 an attem pt has been m ade to plot
isopleths for the height of the regional snowline at the time
of greatest intensity of glaciation in the late Pleistocene by
deducing the lowest altitude in each highland mass at which
\ \ \
\ \ \
3 Altitude of the snowline during the late Pleistocene glaciation of

Tasmania. The isopleths represent the lowest levels (in feet) at which
glacial ice is thought to have formed.
Introduction 5
glacial ice appears to have been generated. In fact, of course,
such a procedure should be m ore correctly described as
picking o u t regional trends in the orographic snowline, and
the plane surface represented by the true regional snowline
w ould be parallel b u t somewhat higher than that repre
sented by the isopleths, b u t it serves to illustrate the way
in which a balance betw een snowfall and tem perature
determ ines the lim its of perm anent snow. In the west of
Tasm ania glacial ice developed at heights of about 600 m
in the F rankland and A rth u r Ranges b u t the snowline rose
spectacularly eastward so that it was at over 1200 m on the
C entral Plateau and on Ben Lom ond in the northeast. T his
trend reflects the present-day precipitation gradient and
implies that, in the w etter west, perm anent snow lay and
glacial ice could form at m uch lower altitudes and signific
antly higher tem peratures than in the drier east. It will
be noticed that the snowline isopleths trend northw est to
southeast rather than n orth to south as in the case of
present-day isohyets. T h is seems to be because in the
Pleistocene, as now, most snow m ust have come w ith south
westerly winds after the crossing of a cold front, b u t it may
also reflect higher sum m er tem peratures in the no rth ern
part of the island, whicli w ould tend to raise the snowline
in this direction.
A sim ilar behaviour of the regional snowline b u t on a
larger scale may be deduced from an inspection of Fig. 2,
where it may be seen that it was not the coldest areas of the
world which were glaciated in the Pleistocene, b u t the
wettest, coldest areas. T h u s the regional snowline rose east
ward from Scandinavia and relatively little of Siberia was
glaciated, presum ably because, although colder, it was also
drier. T h e present-day snowline appears to lie at about
4500 m on the equator bu t rises towards the drier latitudes
of the subtropical high pressure belts and then descends
rapidly to the w etter latitudes of tem perate frontal activity
(Fig. 4) .
A corollary of this dual control of snowline position is
that some areas lie w ithin the snowline prim arily because
they are particularly snowy (for instance parts of the South
Island of New Z ealan d ), while other areas carry perm anent
snow prim arily because they have cold summ ers (for instance
parts of A ntarctica) . T h e re is an im portant distinction in
resulting glacier regimes betw een these two sorts of areas,
and this will be discussed in C hapter VI.
Latitude
4 Approximate variation of the regional snowline with latitude. This

can only be a rough guide, mainly because of longitudinal variation in
precipitation within individual land masses (latitudinal precipitation averages
from Trewartha, 1954).
T h e nature of the orographic snowline can also be

illustrated from the results of the T asm anian glaciation for,
where a m ountain mass has a northerly trend, then in
variably there is evidence that perm anent snow lay lower
on the east-facing side and, where it has an easterly trend,
there is often evidence that the snowline was lower on the
south-facing side. In a region dom inated by westerly air
streams the w ind tends to move fallen snow from west-facing
slopes on to the m ore sheltered eastern sides and the re
sulting greater accum ulation on the lee slope lowers the
snowdine there. A t the same time, in the southern hem i
sphere, slopes with a northerly aspect receive m ore insolation
than those facing south. Snow lying on south-facing slopes
therefore suffers less ablation and the local snowline also
tends to be lower there. These relationships are represented
diagram m atically in Fig. 5. A m ajor exception occurs where
the preglacial relief is itself asymmetrical. In such a case
the snowline may be lower on one side because slopes are
Introduction 7
gentler and provide a better supply of depressions in which
snow is able to lie. A good exam ple of such a case is pro
vided by the Pleistocene glaciation of the M t W ilhelm
area of New G uinea where, because of its nearness to the
equator, insolation cannot be an im portant factor b u t
asymmetry of the divide favoured snow accum ulation on the
southern side (Reiner, 1960) .
Insolation
Prevailing
5 Diagrammatic representation of the influence of aspect on the height

of the orographic snowline in the southern hemisphere. The snowline is
normally lowest in the southeast of any highland block.
W herever an asymmetry in the orographic snowline

occurs it is likely to be of considerable significance from a
geom orphic point of view. N ot only will it lead to
asymmetry of glacial sculpture in the highland area con
cerned b u t it will lead to bigger and longer glaciers
em anating from the slopes where the snowline is lowest, so
that forms of glacial deposition are likely to be m ore in
evidence below them.
The proglacial zone

Given sufficient accum ulation w ithin the snowline, glaciers
are able to extend to varying distances outside it, b u t their
direct influence ends along a sharp line — the line of
m axim um ice advance. T his is not always easy to reconstruct
in the case of past glaciation since the m orphologic and
stratigraphic effect of some glaciers becomes weak towards
their outer limits, but, however difficult it may be, the
possibility of defining a line is always present. O utside this
line the glacial influence is extended by three m ajor groups
of phenom ena — fluvial, lacustrine, and aeolian.
Ice limit, ^ f
latest glaciation / ) *1
Glacial drainage
channels
— j «*-
6 The proglacial zone in central Europe during the last glaciation. It

is marked by abandoned glacial drainage channels and wind-blown deposits
termed loess (adapted from Flint, 1957).
M eltw ater issuing from the glacier has m ore or less

im portant erosional and depositional effects, sometimes for
great distances. Such effects are term ed glacifluvial (some
times fluvioglacial) . T h e m eltw ater may become ponded
between the glacier mass and higher ground or the glacier
may dam back the original preglacial drainage. In either
event proglacial lakes may form. T h e outwash deposits left
by m eltw ater provide surfaces over which wind can operate
so that finer sedim ent particles may be rem oved and lodged
elsewhere.
Introduction 9
T h e zone in which these proglacial phenom ena are to be
found will clearly vary considerably in w idth betw een one
glacial com plex and another and also w ithin a single glacial
complex. T h is is particularly to be expected in the case of
m ountain glaciers lying w ithin irregular relief. W hereas
proglacial lakes by definition are in contact with the glacier,
glacifluvial outwash may be carried long distances and wind
may transport m aterial to areas well rem oved from those
that have suffered the direct effects of glaciation. Fig. 6
illustrates the relationship betw een glacial and proglacial
zones at the fullest extent of the last glaciation in central
Europe.
Peri glacial climates

T h e term periglacial has been considered unsatisfactory
because it is apt to give the m isleading impression that areas
to which that adjective can be applied are peripheral to
glaciers and th at the processes involved are related to those
of glaciation. U nfortunately no generally agreed alternative
term has emerged. In this book the term periglacial is used
in a clim atic sense — that is in the sense of a climate
approaching a glacial one — and such a usage now seems to
be generally accepted (see, for instance, the review by Dylik,
1964) . Clim ates in which periglacial processes operate are
those in which alternate freezing and thaw ing of ground-
water com m only occurs. T h e frequency and am plitude of
freeze-thaw cycles may vary and give rise to different types
of periglacial processes or different intensities of operation,
b u t some m inim um occurrence of such cycles appears
necessary. In any event it is clear that the clim atic req u ire
m ents for periglacial activity bear no intim ate relationship
to those for glacial action and as a result it is not surprising
to find that, at any point in time, areas experiencing peri
glacial processes may be far rem oved from contem porary
glaciers. A good exam ple is provided by M acquarie Island
where about a thousand miles of ocean separate a land
surface undergoing m odification by periglacial action from
the nearest glacier.
10
Introduction 11
Periglacial limits. Periglacial regions with perm anently
frozen ground, or permafrost, that undergoes partial thaw
ing seasonally, are relatively easy to dehne, although even
in this case the outer boundaries of discontinuous and
sporadic perm afrost are hard to plot (Fig. 7) . Some peri
glacial processes, however, are not dependent on the
presence of perm afrost and their outer lim it is particularly
difficult to draw, for, as one goes into lower latitudes or
lower altitudes, these processes become progressively less
effective and increasingly subservient to those of other
climates. T ro ll (1944) suggested the outer lim it of frost
patterned ground (p. 22) as a suitable boundary. Brochu
(1964) suggested that regions with a periglacial climate
should be considered as all those where at least ten annual
freeze-thaw cycles occur. Such a criterion w ould extend the
periglacial dom ain into regions norm ally considered as
tem perate hum id and even tropical arid.
It seems im portant to distinguish betw een regions where
some freeze-thaw processes occur and those where the whole
or m ajor part of the landscape system depends on such
processes. T h u s m echanical w eathering of rocks by frost is
an expectable feature of several m orphogenic systems: the
most im portant aspect of the periglacial system is the
presence of frost-induced mass m ovem ent and such mass
m ovem ent only attains real significance in landform evolu
tion in regions where trees are com pletely or com paratively
absent. T h is is partly because of the effect of forest cover
in m odifying the micro-climate by reducing the incidence
and am plitude of freeze-thaw cycles, and partly because of
the inh ib itin g effect of root systems. T rees also reduce the
w ind effect which is considered by most authors as an
im portant ancillary in the periglacial regime. Conversely,
establishm ent of the trees themselves is in h ib ited by frost
churning of the soil. T h e most realistic and most easily
defined lim it to the periglacial dom ain therefore is probably
the treeline. T h e basic work of T ro ll (1944, 1947, 1948)
im plies a general coincidence betw een periglacial landform
evolution and unforested regions. Perhaps the m ajor
anom aly is created by the presence of areas of perm afrost
beneath the forests of continental Siberia and N orth
America. M uch of this perm afrost may be fossil but, as it
thaws, it creates ‘d ru n k en forests’, where the trees tilt at
a variety of angles, and it also produces characteristic peri-
glacial landforms. T h e best agreem ent betw een periglacial
lim its and tree lim its probably occurs in high altitude
regions where perm afrost is absent. In A ustralia, for
instance, active mass m ovem ent due to freeze-thaw takes
place only above the treeline in New South Wales, Victoria,
and Tasm ania.
If the delineation of present-day periglacial zones presents
problems, these are m agnified still fu rth er when the question
of identifying form er zonal lim its arises. Partly because of
the short period of tim e w ithin which fossil periglacial
phenom ena have been recognised, partly because of the
relative lack of exploration in this direction, and partly
because of difficulties in in terp retatio n of the evidence that
has been found, m ajor uncertainties exist as to the exact
extent to which the periglacial influence spread d u rin g the
cold phases of the Pleistocene. Some indication is given by
Figs. 1 and 2 which show lim its for the present-day tun d ra
together with deduced lim its for the tu ndra d u rin g the last
glacial stage. It m ust be emphasised though that the late
Pleistocene lim its are m erely reconstructions based on the
best available evidence and this evidence is very varied in
its am ount and reliability. It m ust be rem em bered too that,
as pointed out by Kessler (1925) , clim atic conditions in the
m odern and Pleistocene tu n d ra cannot have been strictly
the same, since the seasonal length of day and night would
have been different in the different latitudes.
It is now generally recognised that, in the past, periglacial
processes operated at m uch lower levels on the highlands of
southeastern A ustralia than they do now. On the highlands
of southern New South W ales they descended to at least
1000 m and possibly 700 m (Galloway, 1965), while in
T asm ania the general lower lim it was about 450 m, going
down to 300 m or even lower in places (Davies, 1965) .
One difficulty in placing lim its arises from the fact that
different rock types react differently to frost w eathering and
frost-induced mass m ovem ent — they vary in their readiness
to be mobilised by frost. Because of this, periglacial condi-
Introduction 13
tions may appear to have extended nearer to sea level on
some rock types than on others.
Periglacial subdivisions. Ju st as it is im portant to bear

in m ind variations betw een wetter, w arm er glacial climates
and colder, d rier ones, so it is necessary to rem em ber that
there is a great range of conditions w ithin the overall peri
glacial climate. It is basic to present work in periglacial
geom orphology th at there are contrasting kinds of frost
clim ate and that these tend to encourage different processes
giving rise to different groups of landform s (T roll, 1944) .
T h e most useful classification of periglacial climates is
probably that of T ric a rt (1950, 1963), and this is given
below.
A. Dry with severe winters (associated with perm afrost,

little or negligible ru n n in g water, and strong w ind
action)
B. H u m id w ith m arked winters

1. Arctic variety (associated with less perm afrost, many
m ore freeze-thaw cycles than A, and m uch m ore snow,
which provides protection from w ind in w inter and
ab u n d an t snowm elt in summer)
2. M ountain variety (associated with frequent frost
action, m ore ru n n in g water, and little w ind effect
because of higher precipitation and m ore irregular
topography)
C. Cold w ith little seasonal tem perature change

1. H igh latitu d e island variety
2. Low latitude m ountain variety (both associated with
abu n d an t short-period frost cycles penetrating only
shallow depths and w ith reduced w ind effectiveness
because of high hum idity) .
Some im portant contrasts emerge. Perm afrost is charac
teristic of A, of irregular and infrequent occurrence in B,
and absent in C. Parallel w ith this, A experiences seasonal,
14 Landjorm s of Cold Climates
high am plitude frost cycles, penetrating to great depths,
while C experiences diurnal, low am plitude frost cycles
affecting only shallow depths. Again, B is interm ediate.
A nother significant dichotom y from a landform point of
view lies betw een A in which w ind is im portant b u t ru n n in g
water is not and B and C in which the reverse tends to be
the case.
O ther writers have draw n attention to the possibility of
distinguishing roughly concentric zones in which particular
assemblages of landform s occur, b u t these zones tend to
follow vegetational rath er than clim atic lines (cf. P olunin,
1951). In the n o rth ern hem isphere, Büdel (1948) m apped
a high latitude frost rubble zone and a m ore southerly
tundra zone, in which periglacial landform s are developing
in somewhat different ways (Figs. 1 and 2) . F u rth e r south
still he indicated a less clearly definable boreal forest zone,
into m uch of which perm afrost extends, b u t where typical
periglacial w eathering and mass m ovem ent processes are
considerably reduced. In the southern hem isphere, deglaci-
ated A ntarctica and some of the sub-antarctic islands clearly
fall w ithin the frost rubble zone, but, because of the
disposition of land and sea, there is no true equivalent of
the tundra and boreal forest zones (Fig. 8) . Some of the
subantarctic islands, such as M acquarie Island, share many
characteristics of the tu n d ra zone, b u t their extrem ely
oceanic position and the absence of perm afrost prevent their
being homologous.
T h e frost ru b b le zone comprises the true polar deserts
in which there is m uch bare rock and stony ground, vege
tation is sparse or absent, and precipitation rarely averages
m ore than the equivalent of 125 m m of rainfall a year, so
that there is relatively little snow cover. T h e zone coincides
m ore or less with the arid zone of Corbel (1961) and the
drier part of T ric a rt’s A climate. T h e extrem e paucity of
soil and vegetation is due not only to the severity of the
cold, dry clim ate b u t also to the fact that, with possible
m inor exceptions, these regions have been covered by glacial
ice for m uch of the Pleistocene. In A ntarctica recolonisa
tion by vegetation awaits the arrival of a suitable flora. It
is these arid periglacial landscapes that most closely resem ble
Introduction 15
the arid landscapes of the tropics and, where steep rock
slopes encourage some degree of slopewash and gullying,
the resem blance may be very close indeed.
T h e tundra zone is generally covered by vegetation and
has a more hum id climate: however, there is great variation
in precipitation and the zone falls w ithin part of T ric a rt’s
A climate as well as lying m ainly in his B 1 category. In
general the w etter areas have been glaciated whereas the
d rier have not. T h e presence of a cover of coherent soil
and vegetation as com pared with the frost ru b b le zone leads
to variation in process effectiveness and landform evolution.
T h e divisions suggested by Büdel may also be applied on
an altitudinal basis and R app and R udberg (1960) have
used such a scheme in describing recent periglacial
phenom ena in Sweden.
N ivation climates
Interm ediate betw een the glacial and periglacial systems of
landscape evolution is a third, less im portant, group of
processes which may be placed under the general heading
of nivation. In this third system sedim ent transport is
carried out, not by mass m ovem ent as in the periglacial
system, nor by glaciers as in the glacial system, b u t by snow
m ovem ent or snow-melt runoff. Snowpatches are able to
erode by a com bination of freeze-thaw w eathering and snow
or snow-melt transport. T h e system is truly interm ediate
in nature because it is not separated by any clear dividing
line from either of the two m ajor systems. T hus, as the
proportion of liquid to solid increases, it is possible to
envisage a continuum from periglacial solifluction, where
m eltw ater is an im portant ingredient, right through to what
m ight be considered true snow-melt transport. It is also
possible, as the snow becomes m ore compact and changes
by various stages to glacial ice, to envisage another con
tinuum from transport by m oving snow to that by a flowing
glacier. T h e system is truly interm ediate in the landscape
since it demands periodic m elting of snowpatches. These
tend to lie near the snowline b u t on either side of it, so
that they overlap into both glacial and periglacial domains.
N ivation is particularly characteristic of cold, hum id
climates where snow is abundant, b u t also requires some
am ount of pre-existing relief to provide depressions w ithin
which the snow can accum ulate.
GLACIAL
■SS,V 4
NIVATIONAL
Frost \ PERIGLACIAL
weathering
9 The three process systems of cold climate landform evolution. The

glacial system is dominated by glacial ice, the nivational system by snow,
and the periglacial system by ground ice, but the systems grade into one
another so that they are not separable by sharp boundaries.
Comparison of glacial and periglacial limits

T h e w idth of the periglacial zone at any tim e is always
greater in dry climates than in wet. T h is is because, while
glaciation is strongly influenced by the availability of
m oisture as well as by low tem peratures, periglacial activity
is m uch less dependent on precipitation am ount and m uch
m ore dependent on the incidence of freeze-thaw cycles. In
any circum stance where the snowline rises or falls in re
sponse to a m arked precipitation gradient the periglacial
zone will clearly be wider where the snowline is higher.
1 Miniature nonsorted polygons at about 1500 metres on the Ben
Lomond plateau, Tasmania. The depressions are wallaby tracks and the
glove is about 20 cm long.
3 Section through a steep fossil solifiuction slope in dolerite on M t
W ellington, Tasm ania. N ote the characteristic inversion of soil profile
and alignm ent of boulders dow nslope (G. van de Geer)
4 Two-cycle tor near T idbinbilla, A ustralian C apital T erritory (J. N.

Jennings)
Introduction 17
T h is is well illustrated on a local scale from Pleistocene
T asm ania where, as was indicated in Fig. 3, the regional
snowline rose from west to east. A lthough it is unlikely
th at the lower lim it of periglacial action was absolutely
horizontal, all available evidence supports the view that it
was m uch less precipitation dependent than the snowline.
In the drier east, on Ben Lom ond, where the snowline was
at over 1200 m, glaciers did not extend below 900 m, b u t
the periglacial lim it was about 450 to 600 m lower. T h e re
was thus a wide zone w ithin which periglacial processes
could operate. In the w etter west, on the T yndall Range,
where the snowline lay at about 750 m, glaciers came down
to 75 m, well below the periglacial lim its and probably into
the forest zone. H ere there was no room for periglacial
processes to operate.
O n a sm aller scale than this, it follows that periglacial
action is likely to achieve greater expression on those slopes
of any individual highland massif where the snowline lies
at a higher level because of such factors as aspect or orienta
tion. O n a larger scale it will be evident from a com parison
of Figs. 1 and 2 that great variation in w idth of both present
and form er periglacial zones occur on a w orld scale. In drier
eastern Siberia, for instance, the periglacial zone has always
been of greater w idth than in w etter E uropean Russia and
Scandinavia.
A nother corollary is that m any cold dry areas have
escaped glaciation com pletely and periglacial processes have
been able to operate over a long period of time, while in
cold wet areas that have been glaciated the periglacial cycle
dates only from deglaciation. Again this can be illustrated
on different scales. Periglacial landform s are m uch better
developed on the relatively dry 1250 m high, unglaciated
plateau of M t Barrow in northeastern Tasm ania than they
are on w etter glaciated plateaus of sim ilar altitude fu rth er
west. Periglacial processes have a m uch longer history on
the drier, unglaciated lowlands of the Alaskan arctic coast
than they do on the w etter, glaciated coastlands of Labrador.
II
PERIGLACIAL PROCESSES
There are some geomorphic processes that may be con

sidered endemic to the perigiacial domain, or at least may
be thought of as highly characteristic of it. Some processes
on the other hand are equally represented in other morpho-
genic regions, and may or may not take on a particular
character under frost conditions. In this chapter all these
types of processes will be considered in an attempt to give
an overall picture of the main ways in which landforms
evolve in perigiacial climates. Chapters III and IV deal with
the landforms themselves in their characteristic assemblages.
Ground ice
The most consistent and typical feature of perigiacial
regions is the presence of ice in the ground. Such ice may
be present from one year to the next, in which case it is
termed permafrost; it may appear and disappear seasonally,
often in the form of ice lenses; it may be only of nightly
occurrence and produce little more than a crust at or near
the surface of the soil. When water freezes its volume in
creases by about 9 per cent, so that ground with a large
moisture content will increase in volume proportionately
more than ground with a small moisture content. But the
effect is much magnified by the growth of ice crystals, which
attract remaining liquid water and so cause progressive
segregation of groundwater into favoured areas (Taber,
1929, 1930). In unconsolidated sediments this results in local
expansion and contraction so that there is upward heaving
of some sections relative to others. On a large scale frost
18
Periglacial Processes 19
heave may produce structures capable of developing into
landforms and in this way ground ice may be considered in
part as a structure-forming phenomenon.
Permafrost. Permanently frozen ground (‘perennial tjäle’

in some European literature) underlies large sections of the
continental surfaces, forming in both consolidated and un
consolidated materials. Its distribution in the northern
hemisphere, as shown in Fig. 7, is particularly extensive;
in the southern hemisphere continuous permafrost is found
in Antarctica but elsewhere is of only sporadic occurrence
in some higher mountains. In the continuous zone, perma
frost is commonly about 300 m thick but attains 600
m in parts of Siberia. Thicknesses decrease progressively
towards the sporadic zone where they are usually less
than 30 m (Black, 1954). Permafrost is normally absent
under large rivers and lakes and in any case its upper limit
is depressed in such situations. Its continuity and thickness
are very much influenced by such factors as vegetation,
aspect, and snow cover, all of which cause local variations in
the thermal regime of the ground. The Russian word talik is
being increasingly used for gaps or windows in the perma
frost.
W ithin the permafrost there is great variation in the
water content. If the pore space is completely filled, a hard
impermeable rock is formed in which supersaturation is
indicated by granules, veins, or lenses of ice. If they are
undersaturated and the pore space is incompletely filled,
unconsolidated sediments are likely to be left incoherent
and friable. Liquid water may exist even within the perma
frost mass, since its freezing point is considerably lowered
under pressure, and, in permafrost hydrology, sub-, intra-
and supra-permafrost water is distinguished. Artesian type
systems sometimes arise in which water is moved upward
through taliks by hydrostatic pressure. A similar upward
movement can result when progressive freezing in a down
ward direction towards the impermeable permafrost layer
sets up pressures in unfrozen water and materials between.
Washburn (1950) has termed these cryostatic pressures and
has noted that they may be created in non-permafrost areas
if an alternative im perm eable substratum , such as bedrock,
is present. W hen w ater reaches the surface in this way it
comm only forms an ice mass, called in Russia a naled.
Strong perm afrost springs in Siberia may produce enorm ous
naleds m any square miles in area; weak springs or those
involving translation of sedim ents can produce dom ed
structures above the ground surface, developing landform s
at different scales.
Ice wedges. Segregated ground ice tends to form at right

angles to the direction of heat flow and so is usually m ore
or less parallel to the ground surface. In horizontally bedded
deposits this effect is reinforced by pre-existing structure.
Such lens-like segregations may be large and easily visible
or they may be q u ite gneissic in appearance and not readily
observable. T h e m ost outstanding exception to this rule is
that of the ice wedges, which have been found active in
polar regions and have left sedim ent-filled casts in some
tem perate lands. T h e thicker, upper ends of the wedges
comm only intersect the surface in a m ore or less polygonal
pattern although T a b e r (1943) has m aintained that such
outcropping is a result of subsequent denudation. T w o m ain
hypotheses of origin have been p u t forward. T h a t most
comm only held relates wedging to the freezing of seasonally
thaw ed water in a system of cracks produced by frost desic
cation (see for instance Cailleux and T aylor, 1954) : the
other envisages tensional jo in tin g as a result of differential
frost heave and has been proposed notably by T a b e r (1943) .
Both may well be correct in different instances or even in
the same instance. Ice wedge systems produce polygonal
trenching and are im portant in the developm ent of therm o
karst (Figs. 27 and 28, pp. 60, 61). B ut in the fossil state their
greatest significance to geom orphology may be as indicators
of past severely cold climates acting on the ground surface
on which they are now found (Pewe, 1966) .
Non-perennial ground ice. Seasonal ground ice (‘seasonal

tjäle’) naturally forms on a m uch sm aller scale than does
perm afrost, b u t thin, norm ally lenticular, ice segregations
may occur. T h e location of such ice segregations is very
m uch influenced by particle size since they form m uch m ore
readily in clays than in sands. T a b e r (1929) reported the
results of experim ents in which segregation took place
readily in well sorted m aterials of less than 1 m icron in
diam eter, whereas favourable conditions were needed for
m aterials averaging about O'002 mm, and very fine sand w ith
a m axim um diam eter of O'07 m m gave only the faintest
evidence of segregation. Seasonal ground ice derives its
greatest im portance from the way in which differential
expansion and contraction or frost heave, repeated season
ally, induces sorting of waste particles and their m ovem ent
in lateral, upw ard and dow nw ard directions (Corte, 1966).
L ateral m ovem ent and sorting are also the m ajor effects
of superficial ice segregations form ed diurnally; these are
too shallow in location for the vertical m ovem ents they cause
to be of direct significance. T h e most typical form of super
ficial ice segregation is th at of needle ice (pipkrakes),
composed of colum nar ice crystals oriented norm al to the
cooling surface and capable of lifting earth particles, stones
and so forth several centim etres above the general ground
surface. According to Beskow (1935) needle ice can be u p
to about 20 cm in height. H e differentiated betw een a
com pact type in which the needles are contiguous, a porous
type in which interstices are present, and a discontinuous
type where ice forms only un d er favourable sites. T h e
form ation of discontinuous needle ice un d er stones and
other large particles can be an im p o rtan t condition for size
sorting.
Frost sorting
T h e strong m obility induced in soil particles and soil mois
ture by freeze-thaw processes causes notable churning of
the waste m antle to take place. Evidence for this may be
provided by involutions, visible in sections through frost-
contorted m aterials and involving the in terpenetration of
sub-spheroidal or ‘dum b-bell’ shaped masses of different
grain size (Fig. 10) . It also leads to sorting of particles in
both vertical and horizontal directions. T h a t buried stones
may be b ro u g h t to the surface and eventually ejected is well
22 Landform s of Cold Climates
known. T h e process was explained by H am burg (1915)
and re-outlined by T a b e r (1943) as involving uplift in
cohesion w ith freezing soil, followed by lim ited settling due
to m ovem ent of fines into the void created beneath. T h e
most easily lifted stones are those that are wedge-shaped
downw ard and those that lie norm al to the surface and are
tabular or elongated. T a b u la r or elongated frost-thrust
blocks stand above the surface in characteristic perpendicular
fashion and are a useful indication of freeze-thaw action.
According to T a b e r (1943), dow nw ard m ovem ent of
boulders of high specific gravity also may occur when ice
segregation near the surface releases sufficient thaw w ater
to supersaturate fine-grained soil. T h is may become so fluid
as to perm it boulders to settle downward.
Humus-rich horizon
Gravel
Shale
1 Metre
10 Involutions at the gravel-shale contact in a ground section in central

Montana (redrawn from Schafer, 1949)
Patterned ground. Processes leading to horizontal sorting

produce striking surface patterns of a symm etrical nature
that are well know n from both high m ountain and polar
regions and may cover vast areas of ground in the latter.
T hey have attracted a m ultiplicity of names, num erous
hypotheses and a large literature, conveniently summarised
by Washburn (1956) . Such patterned ground was grouped
by Washburn into five main forms. Circles and polygons
are sufficiently descriptive; nets signify a mesh that is inter
mediate between these two and cannot satisfactorily be
placed in either. All three usually occur on flattish ground.
Steps are found on steeper gradients where gravity assists
in their formation and are loop-shaped with the open end
normally upslope and the closed end usually banked. They
have often been referred to as garlands and are transitional
to stripes, which are characteristic of still steeper slopes and
form in essentially parallel fashion down the steepest
gradient (Fig. 11). Each of these forms may be subdivided
into sorted and nonsorted varieties. In sorted kinds there is
normally a gradation from fine particles in the centre to
coarser particles on the rim, but rare exceptions have been
found in which the reverse is the case: in nonsorted kinds
no such gradation is apparent and vegetation may play a
significant part in outlining the pattern. The distinction
between sorted and nonsorted kinds is essentially the same
as was made by Troll (1944) between patterns in hetero
geneous and homogeneous material.
11 Sorted stone circles changing to sorted stone stripes on a smooth,

northeasterly, convex slope in Spitsbergen (after Biidel, 1960). Note that
in this case the intermediate stone ‘garlands’ are open-ended downslope.
Patterned ground occurs with and without permafrost,

although, as emphasised by Troll (1944), there is a tendency
for larger patterns to be associated with polar regions. In
the Bunger ice window in Antarctica, polygons with a
diameter of about 3 m may be observed (Markov, 1960),
whereas polygons and nets on high mountain plateaus in
Tasmania rarely measure more than 30 cm across. It is
certain, however, that this is not the only factor involved in
size variation, and understanding of the exact processes
generating these patterns remains much less than perfect.
Washburn (1956) listed nineteen hypotheses that have been
put forward to explain patterned ground and concluded
that contraction due to drying and low temperatures, local
differential heaving, the operation of cryostatic pressures,
ejection of stones towards freezing surfaces, eluviation of
fines, and the operation of gravity were all likely processes.
He thought it probable that groups of these processes
operate together or serially and that similar forms can arise
from quite different processes.
The detailed study of patterned ground is probably more
appropriate to pedology than to geomorphology since it
cannot be considered to contribute important landforms.
Its significance in geomorphology seems to be twofold. In
the first place, some of the patterns are apparently related
to things which may properly be called landforms and which
are discussed in the following chapters. Thus steps, as they
become more elongated laterally, seem to grade into terraces,
while the hummocky forms of nonsorted circles, nets, and
polygons, perhaps explicable in terms of cryostatic pressure,
may grade upward into a variety of successively bigger pro
tuberances. In the second place, discovery of fossil-patterned
ground, as in the case of fossil ice wedges and involutions,
tells something of the climatic conditions formerly operating
above the ground surface concerned and which may have
been concerned in the development of that surface. This
second consideration will grow in importance when more is
known of the processes involved and it is possible to classify
patterns genetically as well as descriptively.
Weathering processes
Frost weathering. Frost achieves maximum significance
as an agent of rock weathering in periglacial conditions.
This is because of the relative abundance and intensity of
freeze-thaw cycles and the relative absence of a soil and
vegetation cover. Some moisture is necessary and laboratory
experim ents have shown th at frost is ineffective in perfectly
dry conditions (W im an, 1963) . W ater percolates into the
rock mass and exerts outw ard pressure by expansion and
growth of ice crystals on freezing. In rocks w ith well-m arked
jo in t and bedding planes, w ater enters along these fissures
and the resulting process may best be described as frost
wedging: in rocks which are m ore or less porous, water
enters along vastly m ore num erous and m ore m inute
avenues of entry and freezing causes disintegration or even
bursting or splitting of the rock so that frost shattering
seems the appropriate term . If the rock is a weakly cem ented
clastic rock, or coarsely grained crystalline rock, granular
disintegration is the rule: strongly cem ented and fine grained
rocks may suffer fracture th at is quite in d ependent of struc
ture and texture. T ric a rt (1956) separated m acrogelivation,
or frost w eathering along structural or textural lines of
weakness, from m icrogelivation in which fracture has no
such obvious control. Frost shattering may continue in
particles as small as clay, and M cDowall (1960) has shown
experim entally that clay m ineral particles as small as 0.001
m m in equivalent spherical diam eter can be split by sub
jecting them to rapid, num erous, low am plitude freeze-thaw
cycles. T h e process was m ost rap id in weakly bonded
m inerals of tabular habit.
A relationship betw een porosity of a rock and its suscepti
bility to freeze-thaw w eathering has been dem onstrated in
the laboratory (W im an, 1963) and is also in accord with
m any field observations. In wet climates calcareous rocks
are particularly liable to frost shatter because m any types
are porous and in others planes of weakness have been
opened u p by solution. Rock faces in especially wet situa
tions will also undergo m ore rapid frost destruction. In
general, the frequency of freeze-thaw cycles, the supply of
water, and the ability of the rock to absorb m oisture appear
the m ost im portant factors influencing the rate of frost
w eathering. It is n o t yet clear w hether in freq u en t severe
frosts are m ore potent than m ore frequent less severe frosts,
since the experim ents of T ric a rt (1956) and W im an (1963)
produced apparently conflicting results. In natu re long
period, high am plitude, freeze-thaw cycles are associated with
dry climates so that any effect of frost type is likely to be
masked by the moisture availability factor.
The type of frost cycle may have an effect on the
character of the weathering product. Thus Eakin (1916),
working in Alaska, observed that basalts currently produced
clay as a result of frost weathering, but had disintegrated
into fields of blocks during previous periods of more severe
frost. Hopkins (1949) reported in similar fashion. Corbel
(1961) in a discussion of periglacial morphology in the
Arctic, suggested that oceanic climates with high frequency,
low amplitude, frost cycles produced relatively few large
weathering fragments but an abundance of clay, whereas
continental climates with low frequency, high amplitude,
frost cycles caused deep frost cleavage and consequent large
amounts of angular boulders. Experimental work may not
throw much light on this problem because of the difficulty
of taking the time factor into account and again, in nature,
the moisture availability factor may be critical and difficult
to measure. Taber (1943) emphasised the importance of
slow persistent freezing and continuing water supply in
encouraging ice crystal growth.
Other mechanical weathering. Taber has also suggested

that wetting and drying, producing solution and crystallisa
tion of minerals, may be the next most efficient agent of
rock disintegration in cold climates and that this may be
accentuated by temperature change. Another process that
may be important in periglacial regions is that of pressure
release weathering or dilation jointing. Plutonic rocks that
have solidified under pressure and even some consolidated
sedimentary rocks are susceptible to expansion and cracking
more or less parallel to the surface, when overlying rock
masses are removed. This may be expected especially in
areas that have been heavily denuded by glaciation and may
produce exfoliated sheets of rock vulnerable to further
breakdown by frost.
Chemical weathering. Because of low temperatures, the

relative absence of liquid water, and the slow rate of plant
growth and decomposition, chemical weathering is of
distinctly secondary im portance in periglacial conditions,
b u t may be m ore potent than was once thought. It has long
been realised that the process of carbonation does n o t de
crease significantly in cold climates because lack of heat
and m oisture is counterbalanced by the high rate of absorp
tion of carbon dioxide by w ater at low tem peratures, so that
snowm elt is a notably aggressive agent of solution in
carbonate-rich rocks. O ther chemical processes have gener
ally been assumed to be of very lim ited effect, b u t recent
work, particularly in Russia, suggests that some revision of
this idea is necessary. In periodically frozen soils, where
red istrib u tio n of the w ater content takes place on freezing,
the onset of thaw may b rin g liq u id water in contact w ith
dry m ineral particles at greater depths than was once
thought. N ot only does this increase the potential physical
effects of w etting and drying b u t it has been suggested that
the interaction of water w ith dry m inerals is accom panied
by chemical exchange reactions that are actually intensified
as tem peratures drop fu rth er below norm al freezing point
(Tyutyunov, 1964) . Such a conclusion is of interest as
indicating another way in which ground ice may help to
produce w eathering particles in the sm aller grade sizes.
Weathering products. T h e great preponderance of m ech

anical w eathering in periglacial conditions leads to the
abundance of angular fragm ents that is so characteristic of
frost ru b b le zones in polar regions and on the higher parts of
m ountains, b u t the association betw een frost and angularity
m ust not be taken too far. Frost shatter close to the surfaces
of stones and boulders com m only causes spalling and the
freeing of outer particles so that ro u n d in g is effected.
F urtherm ore, angularity may be caused by other w eathering
processes and also by features of original rock structure and
com position.
T h e preponderance of m echanical disintegration un d er
frost conditions also reduces the depth to which rock
w eathering can take place. It is generally assumed that in
cold climates most w eathering occurs w ithin a few centi
m etres of the surface and very little below about a m etre,
so that deep residual soils cannot therefore be formed. Such
a conclusion may need m odification if previously m entioned
suggestions about deeper w eathering in perm afrost regions
become fully substantiated.
Mass movement processes

Processes of mass m ovem ent of unconsolidated sedim ents
or the w eathered m antle play a particularly significant
part in periglacial landscape evolution. T h e abundance
of soil m oisture at certain periods of the year, the
com parative absence of vegetation — especially of plants
w ith vigorous rooting systems — and above all the action
of frost, are m ajor factors in bringing this about. Very many
term s have been coined and used by different workers for
different types of mass m ovem ent and the term inology of
periglacial types is notably unsettled. O ne difficulty has
been that such term inology has sometimes been approached
purely from a periglacial point of view and w ithout regard
for geomorphology as a whole. Sharpe (1938) was one of
the few who have attem pted an overall classification into
which periglacial processes can be fitted, an d in the discus
sion that follows his term s have been used as far as possible.
It m ust be borne in m ind that, as Sharpe indicated, no
well-m arked boundaries exist betw een some of the processes
defined, and some grade into others. T h re e broad groups
of processes will be considered — those of slow flowage,
rap id flowage, and sliding.
Slow flowage. T h e two m ain types are creep, in which

individual particles or groups of particles tend to move
separately, and solifluction, in which the particles tend to
travel as a mass. In the periglacial environm ent creep is
activated m ainly by frost heave, frost cracking, and by
needle ice (Fig. 12). W hen groundw ater freezes, the re
sultant heave tends to be perpendicular to the surface, but
when thaw occurs particles tend to settle perpendicular to
the horizontal. In the same way needle ice lifts particles
perpendicular to the surface, whereas gravity pulls them
downslope on settling. Cracks and fissures created by selec
tive abstraction of w ater on freezing tend to be filled from
the upslope direction. In these ways a steady downslope
m ovem ent of the waste m antle is induced, sometimes over
very low angle slopes. Soil creep takes place in m aterials
that are not fully saturated, so is characteristic of coarser
grain sizes and of slopes n o t un d erlain by im perm eable
substrata. As a result it is not so typical of perm afrost areas
as solifluction, b u t is very im portant on m o u n tain slopes in
lower latitudes. Rock creep is a term used by Sharpe to
refer to the steady dow nw ard m ovem ent of large blocks or
boulders resting on the ground surface and m ade unstable
by frost heave of the surface or its lubrication by rain or
m elt wash.
12 Four ways in which periglacial creep may occur (mainly after Sharpe,
1938). A and B: frost heave is perpendicular to the surface but gravity
pulls particles downhill on thawing; C: frost cracks tend to be filled from
upslope when thaw occurs; D: the products of frost weathering tend to fall
downhill.
As m aterial in the waste m antle becomes coarser, soil

creep grades into talus creep, in which pebbles or boulders
are involved and fines are absent. T h e slow downslope
m ovem ent of this m aterial is speeded in periglacial condi
tions by freezing and thaw ing of interstitial ice and by
reduction in friction as a result of the abundance of ice and
snow. ‘Black ice’ or frozen rain coating boulders may be
an im p o rtan t agent in this respect. In tu rn talus creep
grades to rock glacier creep, in which the p art played by
interstitial ice becomes m ore im p o rtan t still and rates of
m ovem ent are increased. In the case of rock glaciers the
boulder supply usually appears better, so that there is m ore
pressure of m aterial from above. Rock glacier creep also
differs in being m ore linear and channelled in nature. Some
rock glaciers have been shown to derive from the wasting
of true glaciers, w hen ice flow has ceased and an abundance
of interstitial ice rem ains in a very blocky till, b u t others
cannot have originated in this way.
Solifluction is a term proposed by Andersson (1906) for
‘the slow flowing from higher to lower ground of masses of
waste saturated w ith w ater’. It differs from creep in that
the waste particles move as a mass and req u ire saturation
so that, in its pure form, it represents true flow. Solifluction
is assisted by an abundance of soil m oisture and so it is
encouraged by snow and ice m elt in cold, wet climates. It
is also assisted by an im perm eable substratum and so is
encouraged by frozen ground and particularly by perm afrost.
Because of this it tends to be characteristic of high latitudes
b u t is not excluded from other regions where the necessary
conditions exist. T h e saturation req u irem en t means that
solifluction most comm only involves the finer grades of
m aterial, b u t larger particles and even boulders of consider
able size may be carried along m ore or less in suspension.
Flow may be channelled into fairly well defined streams or
may occur in enorm ous sheets covering whole m ountain
slopes, and such sheets are often lobed along their lower
edge.
Both periglacial solifluction and periglacial creep are
typically in te rm itte n t in operation. In any one case their
effectiveness depends on the occurrence of conditions that
may be present only for certain seasons, and even indeed
for a few days or weeks. It is possible for solifluction con
ditions and creep conditions to follow each other in the
same waste mass so that both processes may contribute to
its m ovem ent. Because of the difficulty or perhaps u n
desirability of always separating the two processes, some
term seems req u ired which will include them both. Many
authors therefore use ‘periglacial solifluction’ in a wider
sense to indicate all slow flowage phenom ena in frost con
ditions. T h e term ‘congeliturbation’ was suggested by Bryan
(1946) b u t has not gained favour, perhaps because of its
length. Edelm an, Florschütz, and Jesw eit (1936) suggested
‘cryoturbation’ for the kind of mass m ovem ent resulting
from seasonal and diurnal freezing and thaw ing of the
ground but, in subsequent years, this too has been used in
both wider and m ore restricted senses — sometimes in
opposition to solifluction, sometimes to include it.
Rapid flowage. Flow at higher velocities requires super

saturation and so is transitional from slow flowage to stream
or glacier flow. T h e two types of im portance in periglacial
conditions are mudflow in which fine particles are trans
ported and debris avalanche m ovem ent in which coarse
m aterial is involved. Both have been associated traditionally
w ith other climates b u t are com m on in cold wet highland
areas, and mudflow may be m ore im portant in high latitudes
than has been generally recognised (W ashburn, 1947).
R app (1960a) has particularly deprecated a tendency to
confuse mudflow w ith solifluction from which it differs in
its greater speed, greater w ater content, and invariable
concentration in channels. It is therefore a m ore im portant
m ethod of linear transport, and, although it relies on an
abundance of fines, may move m uch coarser m aterial by
various forms of traction. T h e process of suff osion, described
by Nikiforoff (1928) and illustrated by Paterson (1940), is
a sort of subsurface mudflow, occurring where thawed, super
saturated m aterial moves from beneath a m ore coherent
vegetation-fixed surface mat.
Debris avalanches also involve the rapid linear m ovem ent
of m aterial un d er conditions of superabundant surface
water, b u t in this case the m aterial is coarser and largely
too coarse to be carried in suspension. M uch of it rolls or
is dragged or even slides, leaving a typical elongated scar
down the slope. As slopes steepen and w ater content gets
less, debris avalanches grade into debris slides, while as
water becomes snow they grade into the dirty snow
avalanches of R app (1959) . T hese latter are regarded here
as nivation features and dealt w ith in C hapter V.
Sliding. Sharpe (1938) defined landslide as ‘the percept

ible downw ard sliding or falling of a relatively dry mass of
earth, rock or m ixture of the tw o’. Usually there is sufficient
water present to aid in lubricating the slip surface. Several
workers in cold lands have noted the association of solifluc
tion with slumping — lim ited sliding along a rotary slip
plane. Washburn (1947) suggested that slumps, and also
possibly debris slides in which no rotary movement is in
volved, initiated solifluction on some slopes of Victoria
Island in the Canadian Arctic. However, since slumping
and sliding are themselves commonly triggered off by the
removal of material lower down, it seems probable on
theoretical grounds at least that it is often solifluction which
is the initiating process. Rock slides and rock falls are
important in cold lands on steeper slopes, particularly where
there has been previous glaciation resulting in oversteepen
ing, and where cliff sapping takes place because of the
existence of a cap of relatively frost-resistant rock.
Rate of movement. The speed with which these various

processes of mass movement operate in periglacial climates
varies principally with the size and other characteristics of
the material being moved and also with the angle of slope.
In recent years, however, enough measurements have been
taken to enable some initial generalisations to be made (for
example, Michaud and Cailleux, 1950; Smith, 1960; Rapp,
1960a; Williams, 1962; Caine, 1963) . It is becoming clear
that, with few exceptions, slow flowage implies rates of less
than about 30 cm a year. The exceptions are sections of
rock glaciers where flows of up to about 3 m a year have
been recorded. By far the greatest amount of movement
is at rates considerably less than these maxima. Normal
rates for periglacial creep and periglacial solifluction seem
to be up to about 5 to 8 cm a year: those for talus creep
are up to about 12 to 20 cm. Particularly where creep
processes are involved, there is abundant evidence that
larger particles move at a notably greater rate than smaller
ones, because of their greater susceptibility to frost heave.
Rapid flowage implies much greater velocities and Rapp
(1960a) measured mudflow at between 30 and 60 cm per
second.
Sorting and orientation. Material transported by all

forms of periglacial mass movement is typically unsorted
and may be difficult to distinguish when fossil from the
deposits of glaciers (see p. 111). Some horizontal sorting
5 Edge of fossil blockstream, M t Wellington, Tasmania. The steepest
surface slope is about 9 degrees (G. van de Geer)
6 Roch glaciers in the Alaska Range, Alaska. These show little of the
transverse ridging and toe bulging characteristic of many (T. L. Pewe)
7 Large-scale rock-cut (altiplanation) terraces in Alaska (T. L. Pewe)
S Turf-banked terrace, Snowy Mountains, New South Wales (A. /> .
Costin)
may occur as a result of frost patterning, so that sorted steps
and, more particularly, sorted stripes are found in relation
to creep and solifluction deposits. In some polar regions the
pattern of stripes may cover vast areas and demonstrate
spectacularly the direction of debris movement.
Some degree of surface sorting may arise from the tend'
ency of larger stones and boulders to move further than
smaller ones when rolling because of their greater kinetic
energy and bigger diameter. This fall sorting is well known
in the case of talus deposits but may occur on gentler slopes
where frost heaving creates instability in surface particles.
Periglacial slope deposits may also show some degree of
vertical sorting and such material has been termed stratified
slope waste by Dylik (1960) . Included here are the grezes
litees and eboulis ordonnes of French workers. In sections
through these deposits, layers of coarser material can be seen
to alternate with bands of finer particles, and all layers are
more or less parallel with the surface. Stratified slope waste
occurs in many parts of the Tasmanian highlands and is
particularly evident when the component material is shaley
gravel of siltstone origin. The gravel is usually well bedded
parallel to the bands and to the surface. In close association
may be found deposits derived from dolerite in which
bouldery bands alternate with layers of fines. The cause of
the stratification is not known, but it may be related to
seasonal or longer period alternations in climate leading to
changes in the efficiency of creep and solifluction. In some
instances changes in the size of available material may be
involved and a thin periodic cover of snow or ice may assist
in producing the well marked and characteristic layering.
The block material in slow flowage deposits displays
orientation characteristics that have been reviewed particu
larly by Lundqvist (1949). In solifluction earths, boulders
and stones tend to lie with their long axes in the downslope
direction — that is in the direction of flow (Fig. 13). How
ever, where movement is arrested, the particles tend to turn
at right angles to this direction. Thus in a lobe of solifluc
tion material, larger particles in the tread section tend to
be oriented downslope, whereas those in the bank or riser
tend to lie across the slope.
34
10%
Stony soil
S to n y laye r
12-18
Peri glacial Processes 35
D etailed m easurem ents m ade on T asm anian slow flowage
deposits by Caine showed th at there is a strong tendency
for blocks to be aligned in the direction of the local slope,
and that this tendency increases dow nw ard from the source
area b u t decreases again as the lower edge of the deposit is
reached. C aine’s results suggested that downslope orienta
tion may be stronger in the m ore channelled deposits (Fig.
14), and analysis of the fabric of unchannelled deposits
showed a lack of orientation such as m ight be expected in
talus accum ulations.
Fluvial processes
A lthough clearly subservient in most cases to the processes
of mass wasting, ru n n in g w ater rem ains a significant agent
of erosion in periglacial landscapes. T h is is particularly so
in those frost climates where precipation is high, where
slopes are steep, or where rock types com bine w ith other
circumstances to in h ib it the form ation of a waste m antle.
U nder any or all of these conditions sheet wash may be an
im p o rtan t or even dom inant m ethod of transport. Pissart
(1966) has emphasised the essentially antagonistic nature
of the action of mass m ovem ent and th at of ru n n in g w ater
un der periglacial conditions. Solifluction tends to prevent
runoff from excavating channels, while runoff, by cutting
channels and d raining the soil, tends to lim it the action of
solifluction.
A good exam ple of the dom inance of ru n n in g water is
provided by southwest T asm ania, where steep slopes on
quartz m etam orphic rocks seem to have been vulnerable to
frost w eathering in the Pleistocene b u t heavy precipitation
caused rap id dow nw ard rem oval of the frost shatter debris
so that a regolith could n o t form. As in the case of this
T asm anian exam ple, the operation of slopewash is norm ally
associated w ith the form ation of fan deposits, sometimes of
great bulk. In the absence of soil and vegetation such pro
cesses can be im portant even u n d er dry conditions and are
13 Orientation of stones in solifluction lobes (turf-banked terraces) in

the Mt Kosciusko region of New South Wales (redrawn from Costin and
others, 1967) (opposite)
36
described by Büdel (1948) as characteristic of parts of his
frost ru b b le zone.
Stream flow in cold clim ate lands also varies in im port
ance w ith local circumstances, b u t alm ost everywhere is
strongly influenced by the great quantities of detritus being
produced by rapid rock w eathering and its efficient move
m ent over slopes of alm ost all possible gradients. Of
im portance too is the coarse and unsorted nature of much
of this m aterial and seasonal changes in availability of
liq u id water. T h e general result is that streams are com
m only braided and perform relatively little vertical incision.
T h e association of periglacial conditions w ith braiding and
aggradation is especially well seen in peripheral areas where
alternations of h o st climates w ith w arm er hum id conditions
has given rise to systems of fill terraces in the valleys. It
should be noted, however, that Büdel (1963, reviewed by
Cotton, 1963) believed th at relatively rap id dow ncutting by
streams may take place in the frost ru b b le zone, apparently
due to rapid frost w eathering of the stream beds.
Aeolian processes
T h e im portance of w ind as an agent of erosion increases in
the periglacial environm ent and particularly in the frost
rubble zone of the highest latitudes and altitudes. In part
this is because of the relative absence of both vegetation
and snow cover, which not only im plies m ore bare ground
bu t also stronger m icro-clim atic winds at the ground surface.
In part it is due to the seasonal locking-up of w ater in its
solid form so that an effective state of aridity arises. In
part it is due to the existence of extensive sources of fine
m aterial capable of being transported by wind. T h e two
most im portant of these sources are the deposits of braided
stream beds and those of waste sheets where frost churning
and heaving continually b ring fresh supplies of fine particles
to the surface. Needle ice has been reported to be especially
efficient in preparing particles for w ind removal.
Sand may be moved from source areas and accum ulate
in dunes or un d u latin g sheets. Nicolls (1958) ascribed
fossil inland dunes in unglaciated parts of T asm ania to wind
winnowing from the braided beds of streams draining catch
ments in which periglacial mass movement operated during
the Pleistocene. Finer material, predominantly silt sized,
may be carried in suspension and deposited as sheets of
loess. Sands and loesses of true periglacial origin may be
difficult or impossible to distinguish from those of glacial
sources and which originate from glacial outwash sheets in
the great majority of cases. Reference to these is made in
Chapter IX. A proportion of this aeolian material may be
disseminated and incorporated into other deposits. Cailleux
(1942), whose work was reviewed by Wright (1946), identi
fied sand fractions which he deduced to be of periglacial
origin in European Pleistocene sediments. He obtained
higher percentages of aeolian sand grains from a belt lying
between the loess deposits and the outer limit of the ice
sheets, a belt associated with evidence for most intense
former frost action and in which fossil ventifacts have been
discovered in particularly large numbers. Such ventifacts
and associated wind eroded surfaces are well known from
the present-day polar regions. Their form and evolution
have been reviewed and illustrated by Sharp (1949).
Cryoplanation
Bryan (1946) suggested the term cryoplanation for the whole
system of land reduction occurring in a periglacial climate,
and the idea that such a morphogenic system can be identi
fied has been developed notably by Troll (1948) and Peltier
(1950). In a general review, Birot (1968) compared the
postulated periglacial system with those proposed for other
climates. There seems little doubt that a periglacial system
can be defined and that its component processes are those
outlined in this chapter, but it is worth noting the general
similarity of the cryoplanation system to the pediplanation
system proposed for arid lands. Both are characterised by
much bare ground, a dominance of mechanical weathering
giving an abundance of coarse, generally angular, material,
a relative absence or irregularity in occurrence of running
water that forms braided patterns and aggrades, distributes,
and corrades laterally rather than incises. In both, wind is a
significant auxiliary agent of erosion. In both, the landscape
as a whole evolves by slope retreat rather than by the down
cutting action of rivers, so that the marine base level does
not exert such an overriding control as it does in temperate
humid lands. Cairnes (1912) suggested the term ‘equiplana-
tion’ for a system in which there is no net loss of material
but merely a redistribution within the landscape. Cryoplana-
tion and pediplanation are ‘backwasting’ systems, contrasting
with the ‘downwasting’ system of peneplanation, and, in
both, large areas may undergo equiplanation.
In spite of these overall similarities, certain differences
remain. The most important of these concerns the greater
part played by mass movement in slope evolution in peri
glacial climates and the greater part played by slopewash
in arid climates. Except where a waste mantle is not present,
the predominance of solifluction and creep in frost con
ditions leads to characteristic convex upper slopes and a
general absence of pediments that contrasts markedly with
arid conditions. Furthermore the ability of periglacial creep
and solifluction to operate on slopes of very low angles
means that slope retreat is accompanied by slope reduc
tion, and this is perhaps in even greater contrast with the
conditions of arid lands.
Recognition of a periglacial system of landscape evolu
tion does not necessarily imply recognition of its wide
significance in moulding the earth’s surface. Parts of Alaska
and Siberia and perhaps of the Tibetan highlands have
undergone continuous reduction by frost processes through
out the Quaternary at least, but very large sections of the
present zone of active peri glaciation have only emerged from
beneath glacial ice in late Pleistocene times. Furthermore,
those regions where fossil periglacial processes have been
inferred are now dominated by other climatic systems. Both
area and time have therefore been limited. It is now
becoming increasingly evident also that periglacial processes
are not as rapid as many earlier writers thought, and in
particular that the movement of waste by creep and solifluc
tion is extremely slow. All this has led many to think that,
even if the concept of a cryoplanation system is a valid one,
it has been of little practical significance in the large scale
evolution of landscapes.
Ill
MASS M O V E M E N T L A N D F O R M S
Since the m ovem ent of waste m aterial over low angle slopes
is probably the most characteristic and exclusive feature of
landform evolution in periglacial regions, it is n o t surprising
that the most characteristic landform s are those owing most
to mass m ovem ent. Biidel (1948) thought that mass move
m ent gains its most extrem e expression in the frost ru b b le
zone, where great sheets of m obile m aterial com m only show
downslope stripey patterns on their bare, stony surface. In
some contrast the vegetation of the tu n d ra zone has an
in h ib itin g effect and the ground becomes alternately stable
and unstable as first vegetation and then mass m ovem ent
take control. T h is tends to produce terracing at rig h t angles
to the downslope direction. Using the term in its w ider
sense, Biidel has distinguished in this way betw een w hat
may be translated as free and impeded solifluction.
T h e various types of periglacial mass m ovem ent do not
depend on the existence of perm afrost, b u t it is clear that
their efficiency is increased by its presence. It follows that
landform s to be described in this chapter may occur
throughout the periglacial realm , b u t they are likely to be
better developed when associated with perm afrost.
Slope forms
Talus slopes. T alu s slopes occur in all climates wherever
free faces are found b u t are perhaps especially evident under
periglacial conditions where strong frost w eathering pro
duces a com parative abundance of source m aterial. In
glaciated m ountain areas the action of the glaciers in
oversteepening slopes has also encouraged talus form ation.
40
Mass M ovement Landforms 41
R a p p (1960b) suggested that there are three basic talus
forms (Fig. 15) . A simple tains slope occurs beneath an
extensive free face which is retreating uniform ly along its
length. If the m ountain wall is dissected into rock-fall
chutes or rock-fall funnels, a talus cone forms in front of
each chute or funnel. T h ird ly , if the talus cones grow
together laterally they form a compound talus slope.
'I
15 Three types of talus slope. A: simple talus slope; B: talus cone;

C: compound talus slope (redrawn from Rapp, 1960b).
A lthough the overall profile of talus slopes is generally

rectilinear, considerable m odification may take place through
processes shifting and rem oving the constituent m aterial.
In periglacial environm ents talus creep and rock glacier
creep are particularly im portant in this, b u t snow avalanches
and snowm elt runoff also help. In an exam ination of some
dolerite talus slopes in T asm ania, Caine (1967a) con
cluded that sliding of whole groups of blocks subsequent
to their initial em placem ent was responsible for some smaller
convexities and concavities showing u p on his survey profiles.
Solifluction slopes. T h e most widespread effect of peri

glacial mass m ovem ent is to produce extensive low angle
slopes — generally at less than 15 degrees — partly erosional
and partly depositional in nature. These were called soli
fluction slopes by Eakin (1916) and the term has been
widely used since. T h e use of such a genetic term appears
undesirable because of the way in which ‘solifluction’ may
be used in both w ider and m ore restricted senses, b u t no
accepted alternative has yet appeared. W here such slopes
lie below a free face, they may head in intervening talus,
as described by Rapp (1960a) . Where no free face is in
volved a smooth convexo-concave slope is developed. In
high latitude regions, solifluction slopes cover vast areas: in
high altitude areas they tend to form individual debris
aprons, the size of which depends on the vertical and hori
zontal extent of the mountain mass. Sometimes the debris
is channelled into well-marked streams of material.
A common effect of the evolution of solifluction slopes
under periglacial conditions is to feed waste material into
pre-existing valleys. Where stream flow is insufficient to
remove this material the valleys may be filled or even
obliterated, a noticeable effect in the upper sections of
catchment areas in highland Tasmania and New Zealand
(see for instance, Cotton and Te Punga, 1955). If material
is fed relatively rapidly into the valley at a point below the
headwater section, then drainage derangement may occur
with the formation of new stream alignments, marshes, and
lakes (Fig. 16) . Such effects are particularly important
where overall stream gradients are low, as for instance in
some of the periglacially modified dolerite hill country in
Tasmania. Where stream flow and transporting power are
|^ L . RAYNER
NICHOLLS
Vl Mile
1/2 Kilometre
Contours in feet
16 The damming of Lake Fenton, Tasmania, by block glacis. At the

lower end of the lake these combined to form an incipient blockstream.
Lakes Rayner and Nicholls occupy a two-storeyed cirque.
Mass Movement Landforms 43
sufficient the debris is carried down-valley and spread over
the floor so that only partial fill is achieved.
Blockfields. Solifluction slopes on which many of the

individual particles exceed about 30 cm in diameter can be
distinguished as blockfields since they produce a very striking
kind of ground surface. Although the term blockfield is
normally applied to accumulations of matrix-free boulders,
such accumulations, at least in Australia, are commonly
continuous with vegetated fields of blocks with interstitial
fines. In many instances, particularly in the polar frost
rubble zones, the blocks have never been associated with a
matrix; but in other instances their bareness is a secondary
feature resulting from removal of previously existing finer
material. Blockfields that descend short distances but have
a wide cross-slope extent may be termed block glacis. Those
that are more extensive in a downslope direction and are
more or less channelled may be termed blockstreams (Figs.
17 and 18).
Movement in matrix-free blockfields appears to take place
by creep, but where a matrix is present true solifluction
occurs, the blocks being carried along in the flowing mass
of fines. Fabric orientation studies of the blockstreams of
Tasmania carried out by Caine (Fig. 14) indicate that they
moved in this latter way, although the matrix has now been
partly removed.
Origin of the debris. The material incorporated in

solifluction slopes and blockfields may originate in different
ways. In the simplest case it is produced by weathering
under periglacial conditions, mainly by frost shatter, so that
a one-cycle system of denudation results. This sort of
system is the one to be expected in the frost rubble zone
and in such cases the blockfield blocks are usually markedly
angular and matrix-free.
The more complicated case is where the debris is pro
vided by readily mobilised bedrock so that two cycles are
involved in denudation — one to produce the debris and
another to remove it. The bedrock may be readily mobilised
because it is unconsolidated; a very important case is when
44
TALUS SLOPES
100 M e tre s
ROCK GLACIERS
3 0 0 M e tre s
BLOCKSTREAM
3 0 0 M e tre s
BLOCKFIELD
3 0 0 M etres
17 N a tu r a l scale p ro file s o f ta lu s slopes on M t B a rro w , T a sm a n ia (a fte r

C aine , 1967a), ro c k g lacie rs in th e A la s k a R a n g e ( a fte r W a h r a ftig a n d C ox,
1959), a n d a b lo c k s tre a m a n d b lo c k fie ld on M t B a rro w (a fte r C aine, 1967a).
P a rtic le size o f surface m a te ria l is a b o u t th e same in each case.
Northern
Summit
/ Dead \
( Island ' --------- Edge of blockstream
.............Toe of blockstream
Unvegetated blockstream
Blockfield saddle
— “ “ Edge of summit plateau

Southern *
Summit tor
Summit
Plateau
IM ile Directic
ongl
1 Kilometre
18 The blockstreams of Mt Wellington, Tasmania
it is composed of glacial till. Through much of the wetter

part of the tundra zone, the material being redistributed
by mass movement has been deposited previously by glaciers.
A striking example in Tasmania is provided on Ben
Lomond where glacial moraines have been redistributed
and translated into periglacial blockstreams (Davies, 1967) .
The bedrock may also be readily mobilised because it
has been extensively pre-weathered: this is especially likely
to be the case where fossil slopes and blockstreams are
concerned. On the dolerite highlands of Tasmania, deep
chemical weathering on old land surfaces in the Tertiary
46 Landjorms of Cold Climates
provided a great mass of clay and of weathered joint blocks,
or core stones, which was moved by periglacial creep and
solifluction during the Pleistocene. This weathering has
been both exogenic (by percolation of meteoric water down
joints) and endogenic (by pneumatolysis). In these areas
many blockstreams are oriented along joint and fault con
trolled lineations that are zones of especially intensive
weathering.
Block material involved in a two-cycle system is likely to
be more rounded than that resulting from frost weathering,
especially when, as is most often to be expected, it has been
derived from core stones or till boulders. With rare excep
tions, such as in the case of very blocky till, it can also be
expected to have a matrix of fines, although this may have
been partly or wholly removed by running water, either
subsequently or even concurrently with mass movement.
When previously weathered material is involved, an inver
sion of the former soil profile commonly occurs: finer
material in the upper horizons is removed and deposited
first so that it lies beneath the coarser particles removed and
deposited later. In this way fossil blockfields on dolerite
country in Tasmania usually show in section a layer of
transported core stones overlying sand, silt, and clay.
Rock glaciers. Rock glaciers resemble unvegetated and

matrix-free blockstreams in that they consist of rivers of
blocks, and so the two forms may be hard to distinguish in
the fossil state. Rock glacier blocks move in a matrix of
ice by rock glacier creep. The two commonest sources of
supply for the ice matrix may be stagnant glacial ice in the
dying stages of a glacier or freezing spring water issuing
from near the head of the rock glacier. This latter is the
chrystocrene of Tyrrell (1910) which seems to be a kind
of naled. Some rock glaciers then may be direct descendants
of true glaciers, whereas others may be independent of
glaciation. Their surface is characteristically ridged in the
longitudinal direction but, near the toe, the ridges run
across the surface. The strongly ridged and bulging convex
toes of rock glaciers may be due to relatively higher rates
of movement than in other forms arising from slow flowage.
In a study of Alaskan rock glaciers, Wahrhaftig and Cox
(1959) refer to previous literature.
Talent (1965) has suggested that streams of basalt and
rhyodacite boulders in the highlands of eastern Victoria
may be small fossil rock glaciers originally related to the
formation of chrystocrenes, and that their partial vegetation
is due to subsequent weathering. Although the general
movement of blockstreams is a result of soil creep and
solifluction, it is probable that some parts of some block-
streams are able to move by rock glacier creep. If this is
so, intermediate forms may occur and this may make difficult
the classification of the Victorian features and apparently
similar ones described by Jennings (1956) from the head of
the River T um ut in New South Wales.
Tor forms
Mass movement in periglacial areas frequently results in the
isolation of free-standing masses of bare bedrock known as
tors, and these features have been the subject of considerable
Summit tor
Break of slope
Valleyside tor
19 Three locational types of tor
recent discussion (Linton, 1955, 1964; Pullan, 1959; Palmer

and Radley, 1961; Palmer and Neilson, 1962). They occur
most commonly on crests and ridgetops, at convex breaks
of slope and on valley sides (Fig. 19), and result from
differential weathering and mass movement leaving more
coherent masses as residuals. In many cases tor location
seems to reflect more massive local jointing or isolation
within a system of lineaments along which weathering has
proceeded more rapidly.
20 Two-cycle evolution or tors. A and B: preliminary deep weathering;

C: subsequent periglacial mass movement (redrawn from Linton, 1955).
Origin. It seems possible to distinguish one-cycle tors,

in which weathering and mass movement proceed simul
taneously, from two-cycle tors, where weathering has pre
ceded mass movement and may have resulted from different
climatic conditions (Fig. 20). Two-cycle tors on the Monaro
plateau in New South Wales (Costin, 1950) and in Tasmania
(Davies, 1967) were produced by deep chemical weathering
of rock followed by periglacial mass movement in Pleisto
cene cold periods, in the manner more fully delineated by
Mass M ovement Landforms 49
L inton (1955). T hey are associated w ith two-cycle block-
fields. U nder periglacial conditions one-cycle tors result
from differential frost w eathering and concom itant removal
of the debris by frost-influenced mass m ovem ent. Such
features are likely to be largely confined to the present-day
frost ru b b le zone. Some small tors in northeastern T asm ania
are probably of even m ore com plex origin and appear to
result from frost m odification of two-cycle proto-tors.
Shape. Periglacial tors vary considerably in size and form.

Size depends largely on the nature of the factors that isolated
them in the first place and also presum ably on the stage
reached in their subsequent attenuation. T h e ir height
reflects the depth of w eathering and the depth to which
rem oval of m aterial has proceeded. One-cycle tors and two-
cycle tors form ed in rock resistant to chemical tveathering
are usually angular in form: other two-cycle tors may show
considerable rounding and, if they are associated w ith deep
rotting, may consist of piles of core stones, sometimes
incorporating perched boulders or rocking stones. T h e very
smallest tors may consist m erely of single jo in t blocks
projecting i?i situ and sometimes form ing fields of notable
extent.
T errace forms
Slopes which have been subject to periglacial processes of
mass m ovem ent are often m arked by terraces of varying
am plitude, which may be depositional or erosional in
character. Depositional terraces may take the form of
individual lobes or fronts of m oving m aterial form ing a
scarplet on the slope and, if these are m ultiple, a system of
terraces will result. T h e phenom enon appears due to down-
slope changes in the efficiency of creep and soli Auction that
produce a slotving of the rate of m ovem ent. A change in
gradient, a change in the efficiency of frost heave, and the
retarding effect of vegetation are am ong possible causes of
such slowing. Depositional terraces can be divided into
turf-hanked (or nonsorted) and stone-banked (or sorted) .
Turf-banked terraces. It is impossible to say where the
steps of W ashburn (1956) end and terraces begin since the
latter appear only a m ore elongated form of the form er.
T urf-banked terraces are often lobate b u t may also extend
in linear fashion. T hey may vary in size from m iniature
forms, commonly referred to as wrinkles, to those a m etre
or m ore in height and several m etres in w idth. T h e smaller
varieties seem to residt from a greater rate of m ovem ent
than in the case of the bigger types. V egetation cover is
thickest on the scarp, and thins, often to nothing, towards
the in n er edge of the step. T h e m aterial of which they are
composed is not sorted b u t is norm ally in the finer grade
sizes. Sections through the terraces show buried layers of
hum us and peaty soils. Most writers have stressed the
retarding effect of vegetation in the form ation of such
terraces and T a b e r (1943) thought that they grew by freeze-
thaw creep, the slow accum ulation of detritus behind mats
of vegetation and the incorporation of plant rem ains in the
soil. O n the other hand, W illiam s (1957) has stressed the
im portance of downslope variation in frost heave as a factor
in developm ent. H e com m ented further on a tendency to
forward bulging and eventual collapse of the scarp or riser.
Costin and others (1967) have described turf-banked
terraces in the Kosciusko area of New South W ales where
they occur on slopes between 4 and 25 degrees. These
terraces are mostly lobate in form and frequently overlap
and intersect. Stones in the terrace m aterial have their long
axes oriented downslope (Fig. 13) and dip parallel to the
general hill slope. Local variation in soil particle size and
the effect of w ind and vegetation are thought to be im por
tant in influencing terrace form, particularly the incidence,
orientation, and shape of the lobes.
Stone-banked terraces. T h e nonsorted forms, in which

vegetation appears to play an im portant part, grade into
sorted or stone-banked forms, in which the riser scarp is
form ed by a bank of large stones or boulders. T h e coarse
m aterial at the scarp grades to finer m aterial at the inner
edge of the tread. T h e vegetation also becomes denser in
this direction and therefore varies in the reverse sense to
Mass M ovem ent Landform s 51
that on turf-banked terraces. T h e stone-banked terraces
appear to be form ed by larger m aterial travelling m ore
quickly downslope than the fines under the influence of
frost-m otivated creep and being checked by a decrease in
slope, by vegetation, or by increasing m utual contact. T a b e r
thought that, when small, the resulting em bankm ent m ight
move in a mass, b u t that it then became stable with age
and grew in place, degrading eventually because of frost
w eathering of the blocks m aking up the scarp.
STONE - BANKED
TERRACES
TURF - BANKED
TERRACES
Humic lenses
R O C K -C U T
(a lt ip la n ati on)
TERRACES
Arrows indicate direction of scarp movement
21 Diagrammatic comparison of three terrace forms
Rock-cut terraces. Some terraces are in part or even

m ainly erosional and th eir evolution poses a somewhat
different problem . Benches may be cut into a wide variety
of rocks — from unconsolidated sedim ents like tills to
igneous rocks like granite and dolerite. A sub-horizontal
plane of weakness and vertical variation in susceptibility of
rocks to frost w eathering appear to assist the process in many
cases. W aters (1962) described terraces cut into dolerite on
slopes of 6 to 15 degrees in W est Spitsbergen and concluded
that they were being form ed by retreat of vertical risers,
m ainly under the influence of frost sapping, and mass move
m ent dowmslope of the w eathered m aterial. In this way
their mode of evolution represents the m ultiple small-scale
occurrence of a process operating comm only on a large scale,
especially in the frost rubble zone. Some dow nw earing of
the treads may occur, b u t W aters thought this clearly sub
sidiary to backwearing of the riser scarps. G radually the
terraces lose their individuality and the final product in the
Spitsbergen case appears to be a debris-covered b u t rock-cut
slope, concave in profile b u t w ith gradients of less than 6
degrees. Similar terraces, less clearly m arked and on lower
gradients, appear to occur on the sum m it plateau of M t
W ellington in T asm ania.
T erm inology relating to terraces has become confused.
Eakin (1916) suggested and used the term ‘a ltip la n a tio n ’
terrace for features that he described and explained in terms
of stone-banked terraces, and most subsequent workers have
equated them in this way, b u t his term has been used by
some writers to designate rock-cut benches. T h e Russian
term ‘goletz’ terrace seems to have been used originally to
describe rock-cut terraces, though often w ith debris-m antled
scarps, and here again different workers have used it in
different senses. Some of the confusion probably stems from
the way in which rock-cut terrace scarps may become covered
w ith stones m oving downw ard from the tread above, so that,
w ithout excavation, it is very difficult to tell w hether the
terraces are being cut or built. T h e vital difference is that
in banked, or depositional, terraces each terrace grows at
the expense of the one below, because if the scarps move at
all they advance. In contrast, the cut terraces extend at the
expense of the one above because they develop by scarp
retreat. In the one case any eventual smooth slope m ust
develop from above; in the other it m ust develop from
below.
According to Alexandre (1958) and Pissart (1966) the
tendency for the form ation of conspicuous rock terraces in
Mass M o ve m en t Landforms 53
periglacial clim ates is m ainly due to a greater capacity for
mass m ovem ent on gentler slopes with a gradient of less
than about 10 degrees than on steeper slopes w ith gradients
of more than about 20 degrees. In tu rn this difference in
capacity may be a result of the m ore rapid drying out of the
m ore steeply inclined slopes so that solifluction is inhibited.
IV
O T H E R PER IG LACIAL LANDFORMS
Landscape features discussed in C hapter III may evolve w ith

or w ithout perm afrost. W ith differing nuances of form they
may occur both in high polar regions and at height in
tem perate and tropical regions. T h e rem aining landform s
to be discussed in this chapter are very closely tied to perm a
frost and, with few exceptions, cannot evolve w ithout it.
T hey are therefore characteristic of polar regions and are
rare or absent elsewhere.
Mound-like forms
T h e surface of the tu n d ra zone is widely characterised by
hum m ocky and pitted relief, most of it on a m icro scale
with an am plitude of no m ore than about a m etre, b u t in
some cases rising to several metres. A great n um ber of
names has been given in various languages to these features,
and their origin is still im perfectly understood. However,
it seems possible to distinguish two m ain, structurally
different m ound forms of m inor am plitude, here called
earth hummocks and ground-ice mounds, and a third form
of greater size, the pingo, which is itself divisible into at
least two sub-types.
Earth hummocks. T hese are m icro-relief forms, norm ally

occurring in fields, often of great expanse. T h ey may there
fore appear to constitute an extrem e form of unsorted net.
T h e individual hum m ocks are knob-like, vegetation-covered
eminences, rarely m ore than about 30 cm in height.
A pparently sim ilar features are called ‘th u fu rs’ or ‘buttes
gazonnees’ by E uropean writers. A lthough their external
54
Other Periglacial Landforms 55
22 Vertical sections through two earth hummocks. The developing soil

profile horizons A, B, and C have been redistributed by frost pressure from
the sides (redrawn from Rudberg, 1958).
appearance and origin may owe som ething to the tussock

h ab it of some comm on tu n d ra plants, such as sedges, sections
through these m ounds show an updom ing of m ineral soil
(Fig. 22) . T his was explained by H opkins and Sigafoos
(1951) in terms of the m ore effective freezing of poorly
insulated m ineral soils betw een the tussocks. T h e frozen
inter-tussock soils were thought to expand and move later
ally into the thawed zone beneath the tussocks and thus
force them upward. Some earth hum m ocks may be of
somewhat different origin and owe their elevation to the
upw ard m ovem ent of fines u n d er hydrostatic or other forms
of cryostatic pressure: they may thus bear a relationship to
subsurface involutions.
Ground-ice mounds. In this group of features the m ound

itself contains an ice body, and updom ing is due directly
to the form ation and growth of the ice (Fig. 23) . T h e
structures to which they are related have been term ed
‘hydrolaccoliths’, particularly in Russia, because of the
apparent analogy to some forms of intrusive vulcanism.
Ground-ice m ounds have been described and illustrated by
Sharp (1942b), who drew attention to their essentially
cyclic nature. As growth proceeds a stage is reached where
the insulating cover becomes so extensively cracked that loss
of height by m elting begins to exceed additions by freezing
so that decay and collapse occur. Collapse produces a
jum ble of m aterial which, when m ore or less vegetated,
forms an irregular m icro-relief of knobs and depressions.
T h e cycle may be seasonal or cover tens of years, and growing
and collapsing m ounds may occur in juxtaposition.
In poorly drained, boggy areas, peat m ounds may form
(Fig. 23 B) . These vary considerably in plan and may be
quite elongated. T ro ll (1944) has attrib u ted such features
to accentuation of vegetation differences by corresponding
differences in the rapidity of freezing. In particular snow,
tending to lie deeper and longer on lower sections, causes
elevated areas to freeze m ore rapidly and encourages ice
segregation. T h e m ound is thus lifted by an accession of
ice as in the case of ground-ice m ounds generally. Such peat
m ounds are called ‘palsen’ (singular ‘palse’) in European
writings and the term paisa bog is used in Canada for a
well-known type of tu n d ra landscape dom inated by peat
m ounds (for exam ple Cook, 1961), b u t ‘palsen’ has also
been used for ground-ice m ounds generally.
0 5 10 Feet
2 Metres
Peat m | Ice | _ | W ater Sg"ave|nd fe ) A Sands,one

23 Sections through ground-ice mounds. A: a ground-ice mound (redrawn
from Sharp, 1942b); B: a peat mound or palse (redrawn from Chemekov,
1959).
As the volume and pressure of w ater increases there is
probably a com plete gradation from ground-ice m ounds,
through the ‘water blisters’ of Nikiforoff (1928), to naleds.
24 Diagrammatic section through East Greenland-type pingo (redrawn

from Müller, 1959)
Pingos. T h e m uch larger m ounds, know n as pingos, may

reach heights of about 50 m and are generally of pseudo-
volcanic form. T h e ir truncated cone shape is often associ
ated with a crater in which a lake may be present and
transverse fissures or radial clefts are common. T hey are
associated with high latitude areas where perm afrost is
relatively thin and may form in both unconsolidated and
consolidated sediments. T h e most suitable substratum
appears to be alluvial sand: the least suitable to be crystal
line rock. An extensive study of pingos was published by
M üller (1959), who made a genetic distinction betw een
pingos of the East Greenland type and those of the
Mackenzie type.
East G reenland type pingos are associated with the expan
sion or new form ation of taliks, or gaps in the perm afrost,
and the ascension of subperm afrost w ater and gases u n d e r
hydrostatic pressure (Fig. 24 ). T h is produces massive ice
form ation and updom ing b u t also a central zone of weakness
in which crater form ation is evident and from which fissures
- Water Pingo ice
Water - saturated Ground water

Quaternary barrier
Drifted material Direction of forces
25 Diagrammatic representation of evolution of Mackenzie-type pingo

(according to Müller, 1959). Sequence is from top to bottom.
tend to radiate. Because the system is an open one, rep eti
tive form ation is norm al and new m ounds may appear
w ithin the rem ains of older ones: associated groups are also
likely to occur.
M ackenzie type pingos are associated with the contraction
and extinction of taliks, and, since such a sequence is most
com m only caused by the disappearance of a lake, these
pingos are usually found on old lake floors (Fig. 25) . As
the perm afrost closes, the increase in volum e associated with
freezing causes an upw ard ascent of water and sedim ents
in a com paratively narrow vertical channel. G row th ceases
once the talik has been elim inated. In contrast to the other
type of pingo, the system is a closed one so that reactivation
and group association are unlikely.
300
M etres
m'Zr dhä?
26 Section through fossil pingo in the Netherlands (redrawn from
Maarleveld and van den Toorn, 1955)
W hereas the disintegration of Mackenzie type pingos is

due to external processes — the m elting of the ice body from
the outside and mass m ovem ent down the slopes — that of
the East G reenland type is m ore complex, since the effects
of internal processes are also involved. According to M üller
the fossil rem ains of M ackenzie type pingos are therefore
likely to be m ore regular in form. Fossil pingos whicli
E uropean workers believe they have identified (for exam ple
M aarleveld and van den T o o rn , 1955; Pissart, 1963b) take
the form of m ore or less circular depressions, often peat-
filled and sometimes with a definable ram part (Fig. 26) .
R am part form ation is apparently due in large part to the
action of mass m ovem ent down the sides of the pingo while
active and in decay. T h is may well be significant in the
case of pingos, which rarely grow faster than about 30 cm
a year and may have a life cycle of hundreds or even th o u
sands of years. R am parts appear to be missing from the
collapsed rem ains of ground-ice m ounds, where the m uch
shorter life cycle and the gentler side slopes m ake significant
mass m ovem ent unlikely.
50 Feet
10 Metres
Contours in
feet
27 A beaded stream in thermokarst ravine in Alaska (redrawn from

Anderson and Hussey, 1963). The straight unbroken lines represent the
surface alignment of ice wedges.
Thermokarst
T h e term thermokarst is used to describe landform s due to
subsidence following the thaw ing of ground ice. T his sub
sidence may result simply from differential thaw ing of a
relatively even surface over perm afrost or it may involve
conditions where ice segregation has caused lateral sub
surface m ovem ent of w ater and sometimes m ineral particles.
T h e literature on therm okarst has been reviewed by A nder
son and Hussey (1963) and, in describing Alaskan examples,
they particularly stress the significance of ice wedge thawing.
O ther Periglacial Landform s 61
T h e thaw ing of ice wedge polygons produces a system of
therm okarst m ounds and thaw pits or thaw depressions, such
as that illustrated by Rockie (1942). T h e m ounds occupy
the centres of the polygons and the pits develop at the in ter
section of the wedges: such pits may contain ice wedge
intersection pools. In itiation of a drainage system along the
lines of ice wedge orientation can lead to the form ation of
therm okarst ravines, which, because they tend to connect a
series of intersection pools, are likely to be occupied by
headed streams (Fig. 27) . However, thaw depressions are
not necessarily related to ice wedges and may occur wherever
there is locally deep thaw. Com m on locations are where the
vegetation m at has been disrupted by frost heave or hum an
interference.
St ro n g e s t summer w in ds
Fine - g r ai n e d Bodies of
soil c l e a r ice
0 M e tr e s 25
28 Evolution of a thaiu lake in the zone of tundra ice wedges. The lake
migrates in the direction of strongest summer winds (redrawn from Hopkins,
1949).
Thaw lakes. T h aw depressions tend to be self-perpetuat

ing and self-enlarging because they usually become filled
with water which further accelerates thaw. E nlargem ent of
depressions in this way may produce thaw lakes, the develop
m ent of which has been treated by H opkins (1949) . These
in turn become enlarged by thaw ing and caving of their
banks, since thaw proceeds most rapidly at w ater level and
undercutting of the margins is norm al (Fig. 28) . W hen a
m inim um size is reached — about 30 m diam eter according
to Hopkins — wave erosion may accelerate the process. In
this case enlargem ent may proceed in one particular direc
tion depending on the incidence of wave-generating winds.
If, as often happens, the opposite side of the lake progrades
by mass m ovem ent and the advance of vegetation, then the
lake may m igrate. M any lakes eventually drain and their
flat floors are recolonised by vegetation and perm afrost.
Thaw dolines. T haw dolines (thaw sinks of H opkins,

1949) differ from drained thaw lakes in having hum m ocky
floors and walls d raining inw ard to a single steep-walled
linear cleft filled w ith open rubble. H opkins thought that
they originated as thaw lakes and that the clefts represented
thawed out ice wedges through which the original lake had
been drained.
2 Miles
2 Kilometres
29 Some tundra lakes in Alaska (redrawn from Black and Barksdale,

1949). The lakes are oriented at right-angles to wave-producing winds
from the ENE.
Tundra lakes
Lakes are extrem ely num erous in tu n d ra regions. M any of
those in unconsolidated m aterials are thaw lakes, b u t a large
num ber in both hard and soft rock areas results from form er
glaciation. M any m ore result from shifting river courses.
T h e great north-flowing rivers of arctic N orth America and
Eurasia thaw in a downstream direction so that, in spring,
great volumes of snow-melt are dam m ed back to cause wide
spread inundation near their m ouths. Such flooding seems
to encourage instability of river courses and a legacy of lakes
originating in old m eanders and braided channels.
Very many tu n d ra lakes resting on unconsolidated
m aterials display striking orientations, which may rem ain
uniform over large areas (Fig. 29) . In spite of early
O ther Periglacial Landform s 63
assum ptions to the contrary, it has been shown that it is
their m inor axes which lie parallel to the wave-generating
winds, and the processes giving rise to their plan form have
been discussed by Carson and Hussey (1962) and by Rex
(1961) .
Seasonally frozen lakes in the tu ndra and elsewhere may
form ice-push ramparts around their edges. These are low
circum ferential ridges composed of the m aterial m aking up
the lake shore and form ed as a result of lateral shoving by
the seasonally form ing ice. Such ram parts are known from
some of the lakes on the C entral Plateau in Tasm ania,
where they are m ade up of dolerite boulders. T h e existence
of m ore than one centre of ice form ation may give rise to
segm entation of the lake by ram parts b u ilt from m ultiple
directions.
T undra bogs
T h e poorly organised drainage and abundance of lakes
through m uch of the tu n d ra zone lead to the form ation of
extensive bogs, which may dom inate the landscape not only
because of the large areas they occupy bu t also because of
their frequently striking surface patterns. Peat m ound bogs
or paisa bogs have already been m entioned. These are
associated w ith perm afrost and generally lie polew ard of a
second characteristic type called string bogs.
String bogs. T ro ll (1944) pointed out that the general

distribution of string bogs in both Eurasia and N orth
America is just outside the lim its of continuous perm afrost
and just north of the treeline. However, H enoch (1960)
among others has described w hat appear to be the same
features m uch further n orth in the C anadian Arctic, so
that the distrib u tio n of paisa and string bogs may not be as
zonal as was once thought. T h e surface of string bogs is
patterned by long string-like ridges or dikes betw een which
lies open water, or sometimes marsh vegetation (Fig. 30) .
T h e m aterial in the ridges appears to be dom inantly silty.
Many string bogs lie on slight slopes and in these cases the
ridges tend to lie across the slope so that they dam a gently
inclined staircase of pools. O n flatter ground the patterns
may become m ore complex, and H enoch suggested the term
‘fingerprint bog’ for some of these.
30 Two surface patterns produced by string bogs (redrawn from Henoch,

1960)
T h e origin of string bogs is not clearly known and they

are probably caused by several processes that may operate
conjointly or separately in different instances. Some w riters
suspect thrusting by w inter freezing of the pond ice, b u t
where a slope is present there seems an evident relationship
to the terrace forms produced by mass m ovem ent, and
H enoch suggested that the im peding of solifluction by the
pond ice may be involved. Corbel (1961) asserted that the
patterns were produced by w ind action and likened the
ridges to barchan fields. In particular he noted that their
9 Slumping caused by melting of ground with a high ice content,
northern Canada. T h e frozen sedim ent contains much segregated ice (J.
lloss Mac hay)
77 T/irnc pit formed by melting of ground ice in loess near Fairbanks,
Alaska. T h e opening at the bottom connects with a tunnel (T. L. Pewe)
12 String bog, Alaska (T. L. Pewe)

Other Peri glacial Landforms 65
w indw ard face displays evidence of w ind erosion. A lthough
w ind may sometimes or even always be concerned in the
evolution of string bogs, their great diversity of orientation
over q u ite short distances suggests that it cannot be the
m ajor controlling agent.
31 Evolution of string bogs (as postulated by Schenk, 1966). The move

ment of thawed material has been from left to right in the diagram.
A m ore recent review by Schenk (I960) concluded that

string bogs are produced by the collapse of perm afrost, when
the underflow of m oving w ater and m ud tilts the still frozen
upper layers so that a succession of form erly horizontal
surfaces dips down in the direction from which flow
emanates (Fig. 31) .
Valley forms
Periglacial processes tend towards the form ation of flat-
bottom ed valleys, generally because of the dom inance of
mass m ovem ent and its relatively rapid action over potenti
ally low slopes. In the first place this causes the feeding
into the river systems of extraordinary quantities of poorly
sorted and often coarse m aterial so that aggradation tends
to occur and the depth of the valley is decreased (Fig. 32) .
In the second place it causes the com paratively quick retreat
of valley sides so that the w idth of the valley is increased.
Even under Biidel’s hypothesis of excessive incision in the
frost rubble zone referred to earlier, a flat floor is still
m aintained, because the conditions which he believed led
to dow ncutting are also favourable to valley slope retreat.
.Iffy. ^mrrlf ..
£
32 Diagrammatic section across the Tea Tree valley, southeastern Tas

mania, showing two alluvial fills forming terraces. The older fill is believed
to result from a phase of periglacial mass movement on surrounding
interfluves (based on Goede, 1965).
Asymmetrical valleys. Since mass m ovem ent on valley

sides is particularly im portant in influencing valley form
under periglacial conditions, it follows that any variation
in effectiveness betw een the two sides is likely to lead to
asymmetry. T h is is particularly to be expected in the case
of small rivers and streams which are m ore strongly governed
in behaviour by slope developm ent processes. T h e occur
rence of asymmetrical valleys in such circum stances and
incapable of being explained by structural control has been
suggested by a n um ber of writers. T hey have been described
from present-day periglacial regions, for instance by Shosta
k o v ich (1927), and especially by E uropean w riters from
areas that have experienced frost conditions d u rin g the
Pleistocene. A recent discussion of asymmetrical valleys in
a part of southern England by O llier and Thom asson (1957)
provides references to m uch of this literature. T h e ir dis
cussion particularly points to the large num ber of ways in
33 Three ways in which asymmetrical valleys may be produced. A:
asymmetry develops by downcutting accompanied by increased mass move
ment on one slope; B and C: one slope is reduced by excessive mass
movement after downcutting has ceased (redrawn from Ollier and Thomasson,
1957).
which asymm etry may develop, m any of them unrelated to

periglaciation (Fig. 33) .
W here mass m ovem ent through freeze-thaw is the con
trolling factor, w ind may have a significant effect by piling
snow on one valley side rath er than the other. A m ore
far-reaching effect, however, is produced by differential
insolation causing m ore intensely heated slopes to thaw out
m ore often and thus induce faster mass m ovem ent on their
surfaces. If a certain am ount of stream dow ncutting is
adm itted this will cause a m igration of the stream course in
the direction of the m ore rapidly retreating side, thus
leaving beh in d w hat am ounts to a gentle slip-off slope. In
apparent sympathy w ith this it is usually the south- and
west-facing slopes of asymmetrical valleys in western Europe
that are steeper. T h e insolation factor may also work
through its influence on the plant cover and, in any case,
is more likely to be effective in higher latitudes where the
sun’s angle of incidence is less.
Wash-cut slopes. T h e effect of perm afrost in acting as an

im perm eable layer p rohibiting percolation at times of thaw,
and thus aiding solifluction and other mass m ovem ent pro
cesses, has frequently been m entioned. W here waste m aterial
is in short supply and particularly where slopes are bare,
the same circum stance leads to m axim um runoff of water
over the frozen rock even where the rock may norm ally be
relatively perm eable. Lack of vegetation and m inim al
evaporation such as occur in high latitudes may still further
increase the discharge ratio so that, even though water may
be scarce and virtually restricted in appearance to the
annual sum m er snow-melt, its effect on bare slopes may be
considerable.
M ortensen (1930) described the processes of slope cutting
under these conditions in Spitsbergen and they have been
more recently emphasised by B iidel(1948). C ut slopes are
concave and lie at angles of betw een about 15 and 40
degrees. T hey are covered by small sub-parallel rills and
gullies rarely m ore than about 30 cm deep and a m etre wide.
Eventually, greater concentration of run-off along certain
favourable lineam ents may lead to the form ation of deep
clefts so that the slope is cut into well-marked facets.
Dry valleys. T h e in term itten t nature of water runoff and

the occurrence of frozen substrata are both factors leading
to the comm on occurrence of dry valleys in periglacial
regions. These may be on a very small scale and be little
m ore than gullies, b u t can also be large enough to deserve
the name of true valleys. In regions where perm afrost was
present during the Pleistocene bu t has now disappeared,
dry valleys on perm eable rocks such as limestones have been
attrib u ted to cutting under frozen ground conditions. T his
explanation was early offered in the case of dry valleys in
the English chalk and, although more likely explanations
have now been suggested in this particidar instance, the
possibility of explaining sim ilar occurrences in this way yet
remains.
Major rivers. T h e rem arks made above refer essentially
to in term itten t streams and the smallest rivers. Large rivers
w ith massive transporting power and able to in h ib it perm a
frost by their very presence seem little affected by factors
pecidiar to the periglacial province, unless, as m entioned
earlier, they How poleward.
Aeolian forms
T h e part played by wind in m odifying the form of arctic
lakes and bogs in the tu n d ra zone has already been m en
tioned, but, as noted in C hapter II, it is in the frost rubble
zone that it achieves greatest effectiveness both in high
latitudes and at high altitudes. Erosional forms are relatively
scarce and, ap art from some small sculptured and faceted
forms, virtually non-existent in consolidated rock. Deflation
hollows and trenches occur in loose vegetation-free m aterials
and are know n even from vegetated country. Boye (1950),
for instance, described and figured deflation trenches in
vegetation-covered ground in G reenland, and rather sim ilar
things occur near crests in the Snowy M ountains of New
South Wales, where, however, they are likely to have been
initiated by grazing and burning.
M uch the greatest effect of w ind is to sort and redeposit
the detritus produced by frost w eathering and frost heave
and lying readily exposed, particularly in the high latitude
deserts. T h e resultant landform s are very sim ilar to those
of the deserts of warm er climates: the finest particles are
exported in suspension and deposited as loess; the sand
sized particles are moved by saltation and grouped into dune
fields; the coarsest particles are left behind as lag deposits
to form stone pavements.
V
N I V A T I O N P R O C E S S E S AND
LANDFORMS
T h e term nivation was introduced by M atthes (1900) to cover

the geom orphic effect of snowpatches accum ulating in pre
existing depressions. Snow causes sedim ent-transport in two
ways: first when it melts and runs off as water, and second
in its solid form through varieties of snow slide and snow
creep. M atthes thought in terms of transport by m eltw ater
but, since it now appears that the same snowpatch may move
m aterial in both ways, it seems desirable to group both
under the same general process heading. However, it should
be borne in m ind that transport by snow-melt is often closely
related to periglacial solifluction and transport by snow
m ovem ent may be transitional to transport by glacial ice.
Snow-melt processes
Snowpatch erosion. Most writers have thought of snowpatch
erosion in terms of frost w eathering of rock and the trans
port of resulting debris by snow-melt runoff, and it may be
taken that these are the essential processes involved, even
though others may operate in particular instances. W. V.
Lewis (1939) described how freeze-thaw of water, derived
largely from periodic m elting of the snowpatch, results in
w eathering of the adjacent rock, and how the weathered
particles are rem oved by m eltw ater ru n n in g beneath the
snow and out from its lower edge. Because snow is a good
insulator the most effective w eathering takes place at the
edges of the patch, but, as these change position with
seasonal growth and dw indling, the position of m axim um
effectiveness also changes.
70
Nivation Processes and Landforms 71
In this way the snowpatch gradually digs into the ground
surface so that the initial depression is enlarged and becomes
a nivation hollow. E nlargem ent of the depression allows
m ore snow to accum ulate and allows it to lie longer, so that
the rate of erosion is accelerated. Vegetation is increasingly
inh ib ited by the m ore effective snow cover and this causes
further exposure of the surface, not only to frost w eathering
and snow-melt b u t also to w ind during the sum m er period.
T h e steeper, barer, sides of the hollow are attacked more
effectively than the floor, which m ust in any case have a
forw ard slope, so that headw ard erosion is dom inant and a
cliff-like slope tends to form at the rear.
If the hollow forms in solid rock, streams and sheets of
m eltw ater may carry away detritus by traction and even in
suspension, b u t where snowpatch erosion takes place in
unconsolidated m aterials or w ith a thick soil cover the
sequence of events is m ore complex. In such a case m elt
water tends to percolate and give rise to solifluction in the
resulting saturated m antle. T h e association of snowpatches
with solifluction further downslope is well known and is
illustrated, for example, by W illiam s (1957) .
Lewis divided nivation hollows and their associated snow-
patches into three m ajor groups — transverse, longitudinal,
and circular — and this distinction is an im portant one
because each is associated w ith other land form types.
Transverse snowpatches. Transverse snowpatches lie

across drainage lines and the resulting hollows are ledge
like. T hey are favoured by sub-horizontal planes of weak
ness in the bedrock or by any preceding process giving rise
to transverse depressions. A close relationship undoubtedly
exists betw een transverse snowpatches and the terrace forms
produced by periglacial mass m ovem ent. W aters (1962)
noted that rock-cut benches in dolerite studied in Spits
bergen, though prim arily due to differential w eathering
and mass m ovement, were also influenced by the accum ula
tion of snow. Pissart (1963a) thought that turf-banked
terraces formed on glacial d rift in W ales were initiated by
transverse snowpatches b u t modified later by mass move
m ent. It seems certain that m any other sorts of relationship
betw een transverse nivation hollows and altiplanation
terraces will eventually be shown to exist.
Longitudinal snowpatches. L ongitudinal snowpatches

form along the direction of drainage and are thus commonly
associated with actual stream courses. T h e ir floor is some
times gullied and their effect often is to m odify a small
pre-existing valley. As the overall slope of their floor
becomes steeper they probably grade into avalanche chutes:
with further snow accum ulation they may also become a
locus for the developm ent of niche glaciers (p. 79).
Circular snowpatches. T h e circular snowpatch is perhaps

not as common as the other types, b u t is basically m ore
im portant, because it may grow into a large form, term ed
a nivation cirque, which is a progenitor of the true glacial
cirque treated in C hapter IX. Because circular snowpatches
are sheltered from the w ind on three sides, they are m ore
efficient in accum ulation. T hey also offer a greater extent
of backwall in relation to their size than do the other types
and this too helps them to erode m ore efficiently. W ell-
developed small examples, such as that shown in Plate 13,
may be not m uch m ore than 10 feet in diam eter, b u t the
largest nivation cirques may be hundreds of feet across and
very difficult to distinguish from glacial features. R itchie
and Jennings (1956) reinterpreted supposed glacial land-
forms in the Grey M are Range of New South W ales in
term s of nivation.
Other snow-melt features. T h e effectiveness of snow-melt

runoff as a gullying and slope-cutting agent and its particular
potency in relation to carbonate-rich rocks has already been
referred to in the discussion of periglacial processes and
landforms. So also has its significance in in itiatin g and
accelerating solifluction. In all highland areas where snow
falls, the m eltw ater, representing as it does an unusually
concentrated burst of surface flow, becomes an erosion agent
of potential im portance, especially perhaps because of its
role in distributing the products of w eathering.
Snow movement processes
T h e operation of snow in its solid form may be divided
conveniently into slow sliding and creeping on the one hand
and rapid sliding and falling on the other.
34 Four ways in which snowpatches may transport. A: saturation of

subjacent unconsolidated material, perhaps aided by weight of snow, causes
soliffuction beneath the snowpatch and the construction of frontal terraces;
B: snow-melt runoff on a hard, impermeable and perhaps steeper surface
moves finer material away from the snowpatch; C: material falling on to
the snow surface from above slides down to form a protalus rampart;
D: in a deep snowpatch snow pressure causes basal sliding and the move
ment of loose particles over a hard surface.
Slviv sliding and creeping. A growing volum e of evidence,

sum m arised by Costin and others (1964), suggests that snow
may move forward by mass creep in which an elem ent of
basal sliding is involved. Presum ably there are m inim um
conditions of snow thickness and slope in which such move
m ent can take place, b u t little is known of them as yet. On
M t T w ynam in the Snowy M ountains of New South Wales,
Costin and his co-workers studied the results of m ovem ent
in a snowpatch that has been observed to survive through
some summ ers and is estim ated to have a m axim um thick
ness of about 40 m. T hey recorded the gradual forward
m ovem ent beneath the snow of m arked stones up to 86 x
77 x 30 cm in size. Some of the stones moved uphill for
short distances. Associated with m ovem ent over the grano-
diorite rock were scratched abrasion tracks and white rock
flour. Similar striations had previously been noted in rela
tion to a snowpatch on M t La Perouse in T asm ania by A.
N. Lewis (1925) .
Com plete or com parative absence of vegetation appears
to be an im portant factor in influencing the ability of
creeping snow to move rock particles, bu t both may norm ally
be associated w ith deep, almost perennial drifts. In the case
of the M t Tw ynam snowpatch the significance of preceding
glaciation in providing a large expanse of bare rock w ithin
the depression has been emphasised.
In spite of deficiencies in knowledge it seems clear that
snow creep may play a significant part in the form ation of
some nivation hollows, although it would appear impossible
for it to operate except in conjunction with snow-melt
transport, and it is therefore likely to be auxiliary to the
m ore widely recognised processes of nivation described
earlier.
Rapid sliding and falling. T h e m ore rapid processes of

snow m ovem ent are im m easurably better known, if only
because they are frequently spectacular and may involve
loss of hum an life. Yet there seems increasing agreem ent
that they have not been given the attention they deserve
by geomorphologists. Fast sliding of snow takes place in
avalanches of which many types are know n to exist. T h e
type that is of m ajor im portance from a geom orphic stand
point and in which the snow moves along masses of subjacent
earth m aterial is generally known as a ground avalanche,
bu t has been term ed dirty snow avalanche by R app (1959),
partly to avoid confusion with the debris avalanches of
Sharpe (1958) . D irty snow avalanches norm ally occur on
steep slopes in late spring or early sum m er when m elting
causes large masses of wet snow to slide at depth. As earth
debris is incorporated and the path of the avalanche
becomes rougher, the anterior sections tend to develop a
rolling m otion, while the whole mass moves at very high
speeds. T h e avalanches move down regular paths which,
by repeated use, take on the form of large gullies or
avalanche chutes.
Avalanche chutes may bear a superficial resem blance to
gullies produced by snow-melt and to rockfall chutes, bu t
are usually wider and on a larger scale (Fig. 35) . T hey may
also display evidence of rock polishing and scratching and
differ in their associated deposits. Avalanche chutes, rockfall
chutes, and snow-melt gullies are all im portant features
giving rise to the ‘buttressed’ appearance of m any steep
slopes in m ountain regions. T h e chutes and gullies form
along m ajor jo in t axes and other lines of weakness leaving
buttresses of sounder rock between.
In contrast to the essentially linear effect of snow
avalanches is that produced by the falling of snow cornices
which may form where a sharply convex break of slope
occurs. Peterson (1966) a ttrib u ted the asymmetrical form
of crests on some T asm anian m ountains to the form ation
of cornices on the leew ard edge and the tearing away of
rock m aterial when they fell.
35 Comparison of debris tongues and cones; A: a talus cone; B: an

alluvial cone; C: an avalanche boulder tongue; D: a rockslide tongue
(redrawn from Rapp, 1959)
Depositional features
So far in this chapter only the erosional effects of snow have
been considered. T h e related deposits have been relatively
little described and are com m only dispersed in such a way
as not to give rise to characteristic landforms. N otable
exceptions are the protalus ramparts associated with some
snowpatches and the boulder tongues associated with
avalanches.
Protalus ramparts. W here w eathered m aterial from cliffs
behind a snowpatch falls on to the snow surface, it may
slide over the surface and come to rest along the outer edge
to form a ridge-like feature term ed a protalus ram part. T h e
m aterial m ust be of a bouldery nature and incapable of
being moved by m eltw ater issuing from the snowpatch. It
is generally distinguishable from glacier deposits by its
sim pler form and structure and by the presence only of
subaerially w eathered rock particles.
A special kind of protalus ram part is the avalanche ram
part described by M arshall (1912) from the Fiordland region
of New Zealand. Snow avalanches produce aprons of snow
at the bottom of steep slopes and rock debris falling on to
this slides over and comes to rest as a ram part. Small lakes
may form behind the ram part dam.
Avalanche boulder tongues. T h e accum ulations of rock

debris resulting from dirty snow avalanches have been
described by R app (1959) as avalanche boulder tongues.
Such tongues are m arkedly concave in long prohle and may
extend for some distance out on to flat valley floors, some
times even partly clim bing the opposite side of the valley.
T h e ir cross prohle is m ore or less flat, although some show
som ething of a fan form. T h e m aterial in the tongues is
not sorted b u t there is often a tendency for larger boulders
to lie close to the edges.
R app com pared the appearance of boulder tongues to
that of sim ilar types of debris accum ulation (Fig. 35). T alus
cones may be distinguished am ong other things by their
relatively straight prohle and m arked fan shape; alluvial
cones, such as m ight occur at the foot of snow-melt gullies,
by their generally round-edged fan shape and tendency for
finer debris to occur towards the base; rockslide tongues by
their very rough and uneven surface and general evidence
of having been form ed in one catastrophic m ovem ent.
Characteristic of the surface of avalanche boulder tongues
are avalanche debris tails, which are small ridges of finer
m aterial lying in the lee of upslope boulders tending to
protect them from avalanche transport. T hey resem ble in
m iniature the ‘crag and tail’ features produced by glacial
action (p. 172).
N ivation Processes and Landforms 77
Nivation landscapes
Perhaps because nivation processes and landform s are so
closely related in many ways to glacial and periglacial
features, there has been little attem pt to characterise or
delineate nivation systems and nivation landscapes as a
whole, b u t in w etter m ountain regions the nivation lim it
may extend substantially lower than the periglacial lim it
in sympathy w ith the depression of the snowline in such
places. T h e higher parts of m ountains in high rainfall areas
of western T asm ania, although currently unaffected by
periglacial processes, are being significantly altered by the
action of seasonal snow.
W here nivation has an opportunity to proceed long
enough to m odify the landscape on its own, it may be
expected to give rise m ainly to shallow scalloping at higher
levels, but at lower levels to linear furrowing, particularly
through the developm ent of snow-melt gullies and avalanche
chutes.
VI
GLACIERS
W hen snow lies through the sum m er m onths, so that the

accum ulations of successive w inters are superim posed, its
density increases. In the m ain this is brought about by
com paction of ice particles under the weight of the upper
layers and also, especially in m ore tem perate climates, by
percolation and refreezing of m eltw ater. W hen the snow
reaches a density of about O’55 it is term ed fu n . T h e
French word neve, originally synonymous w ith firn, is now
used in its technical sense to denote the area occupied by
hrn and not the m aterial itself. However, it is probably
better not to use it at all and so avoid possible confusion:
firn field seems a suitable alternative.
F u rth er consolidation of the firn leads to the form ation
of glacial ice, in which the ice crystals comprise a relatively
im perm eable mass w ith a density close to 0 ‘90. A density
figure of 0 84 is usually taken to indicate the dividing line
between firn and ice. Bodies of glacial ice are called
glaciers and their study is included in the science of
glaciology. A hlm ann (1948) and Sharp (1960) provided
im portant reviews of progress in general glaciology and
McCall (1960) and M eier (1960) reported w hat are perhaps
especially revealing studies of individual glaciers. T h e book
of Hobbs (1911) rem ains a useful descriptive com pendium
of existing glaciers and the Journal of Glaciology is the
m ajor English language periodical covering aspects of their
nature and behaviour.
T h e geom orphic significance of glaciers is twofold. In the
first place they are landform s in their own right, even though
relatively m obile and ephem eral. In the second place they
are agents of erosion and influence the m oulding of the
78
Glaciers 79
landscape which underlies or is near to them . T h is chapter
attem pts not so m uch to summ arise the present state of
knowledge on glaciers as to indicate those characteristics
of glaciers which seem to be im portant to a proper
understanding of the ways in which they modify landscapes.
Glacier forms
Smaller forms. T h e form which a glacier takes depends in
the first instance on the m orphology of the gathering
grounds for snow and, broadly speaking, two m ain cases may
be envisaged. W here the pre-existing topography is dissected,
and especially where the snowline is relatively low, glacial
ice is generated in depressions on slopes — norm ally the
nivation hollows described in C hapter V. T h e most
elem ental forms are probably the hanging glaciers or niche
glaciers, appearing to adhere precariously to sloping hollows
on scarps and valley sides (Groom, 1959, for e x a m p le).
T h e most typical developed form, however, is the cirque
glacier, which occupies a m ore or less arm chair-shaped
depression and is often related genetically to the circular
snowpatch (Fig. 36) . Sometimes, in very large depressions,
'/2 K ilom etre
C o n to u rs in m etres
36 Cirque glaciers around Nautgarstind, Aust-Jotunheimen, Norway

m ore than one area of ice generation develops and a com
pound cirque glacier forms, the continuous surface of which
tends to disguise separate ice bodies with potentially
different characteristics.
66°15'
800-
5 Miles
5 K ilom etres
C o n to u rs in m etres
23*30'
57 Drangajökoll, a plateau glacier in northwestern Iceland. The ice

surface is strongly influenced by the underlying topography.
13 Actively developing circular nivation hollow at about 1500 metres
near summit of Frenchmans Cap, Tasmania (J. A. Peterson)
14 Transverse snowpatch occupying the rear of a rock-cut altiplanation

terrace probably developed during colder Pleistocene times near the dolerite
capped summit of Mt Wellington, Tasmania
J
15 Compound talus slope with snow-covered rockfall chutes cut into a
dolerite scarp along major structural lineaments so as to produce buttresses,
Walls of Jerusalem, Central Plateau, Tasmania (J. A. Peterson)
16 Tan-type avalanche boulder tongue, Baffin Island (J. A. Peterson)
- • „ / W .
Glaciers 81
The other main case is that of relatively undissected
highlands where the ice tends to generate on upland plains
in pancake-like masses termed plateau glaciers. Whereas the
ice surface in niche and cirque glaciers slopes markedly in
one major direction and ice movement is in this direction,
the surface of plateau glaciers slopes more or less radially
outward and ice movement too may be in all directions.
But they are not so divorced from the influence of
topography as the ice sheets to be mentioned later (Fig. 37).
A well-known antithesis in present day ice bodies is that
between the cirque-type glaciers of the thoroughly dissected
Swiss Alps and the plateau-type glaciers occurring on the
relatively undissected surfaces of some of the Norwegian
fjelds. The contrast can be illustrated also from Pleistocene
Tasmania where plateau glaciers formed on the undissected
Central Plateau and on the Ben Lomond mesa, while cirque
glaciers were the important initial forms in the much more
rugged country further west. In this case the contrast was
accentuated by the fact that the snowline was higher on the
central and eastern plateaus so that the only possible gather
ing grounds were on the plateau surfaces themselves. In
the west, on the other hand, where the snowline was much
lower (Fig. 3, p. 4), ice was able to form further down
on the sides of the ranges. The situation in the Snowy
Mountains of New South Wales was rather similar, with
cirque glaciers developing on the flanks of a well-dissected
highland.
Larger forms. Ice from both cirque-type and plateau-type

glaciers may spill into pre-existing valleys to form valley
glaciers, veritable streams of ice moving downward out of
the highlands. If these coalesce in their source regions the
whole body straddles the mountain divides and is termed a
transection glacier: if they coalesce in their lower ends they
form a dendritic glacier (Fig. 38). A massive glacier formed
by ice streams coalescing when they emerge from highlands
on to lower ground is called a piedmont glacier.
Such ice streams, in which movement is to all intents and
purposes unidirectional, contrast with ice sheets, which tend
to be roughly circular and in which movement is charac-
teristically radial. The smallest type of ice sheet is the
plateau glacier, but this may grow through an intermediate
stage, often referred to as a glacier cap, to a full-scale
continental glacier such as that at present covering
Antarctica.
38 A dendritic system of mountain glaciers (A) compared with a tran

section system (B)
The large continental ice sheets of Greenland and

Antarctica, although they may be thought of in one sense
as huge single glaciers, in fact comprise a great number of
identifiable smaller glaciers, especially around the periphery
where their outlet streams have many of the characteristics
of valley glaciers. Where they enter the sea they may form
glacier tongues afloat or ice shelves.
Vostok I Vostok
Mirnyy
-1 0 0 0 '
Kilometres
39 Section through part of the Antarctic ice sheet at about longitude

95°E (redrawn from Shumskiy, 1959). The ice surface is independent of
underlying topography.
Continental glaciers are extremely independent of the

basement relief (Fig. 39) and centres of accumulation tend
to be determined by their location in respect of large-scale
Glaciers 83
meteorological systems rather than by any effect of under
lying topography. As the ice sheet waxes and wanes so the
centres of accumulation may shift and directions of ice flow
may change. Flint (1957) has discussed possible changes of
this nature in the Pleistocene ice sheets of Europe and
North America.
Ice sheets may arise from the swamping by ice of a high
land area originally occupied by cirque and valley glacier
systems, and the large continental glaciers have certainly
gone through a very complicated history of development.
The realisation that dissected highlands which have been
covered by ice sheets have gone through a series of stages
of occupation by glaciers is important to an interpretation
of their morphology. Thus the thoroughly dissected Du
Cane Range in Tasmania was at one stage swamped by
part of the glacier cap covering the west-central section of
the island in the late Pleistocene; yet it also shows clear
evidence of having been occupied by cirque and valley
glaciers. In such a case it is necessary to envisage a sort of
‘sequent occupance’ — an advancing hemicycle from a cirque
glacier stage through a cirque and valley glacier stage to an
ice cap stage, then a retreating hemicycle in a reverse
sequence.
In spite of these and other complications it is usual and
useful in glacial geomorphology to distinguish between
continental glaciation, where the landscape has been covered
by large ice sheets, and mountain or alpine glaciation, where
the dominant agents have been cirque and valley glaciers.
The large continental ice sheets are and have been much
less influenced in behaviour by pre-existing topography than
have mountain glacier systems.
Glacier economies
The ice body comprising a glacier suffers addition and
subtraction by processes of accumulation and ablation
respectively. Although other forms of precipitation may be
significant, accumulation is usually in the form of snow
which falls or is driven into the gathering grounds forming
the upper part of the glacier. Ablation may be through
84
\ Yih
° ' / \\ii
H ii I I
I'M I |
Contours in metres
250 Miles
250 Kilometres
40 The Greenland ice sheet. This continental glacier rises to two major
domes and, around most of its perimeter, produces a complex system of
ou tlet glaciers strongly influenced by the coastal mountains.
Glaciers 85
m elting, by direct evaporation or sublim ation, or by the
breaking off of unaltered segments of ice such as when a
glacier ‘calves’ to produce ice avalanches on land or icebergs
in water. A hlm ann (1948), who discussed the relationship
of accum ulation and ablation in a n u m b er of glaciers around
the N orth A tlantic, used the term ‘glacier regim e’ to refer
to the m aterial balance of a glacier, b u t F lin t (1957) has
suggested that glacier economy is m ore appropriate in that
it leaves glacier regim e to be used in the w ider geomorpho-
logical sense in which it is used for instance in ‘river
regim e’. T h u s glacier regime in the w ider sense includes
not only the economy b u t also meteorology, rates of flow,
and fluctuations.
W hen accum ulation exceeds ablation the volum e of the
glacier obviously increases, and its outer lim its generally
advance: this is a positive economy. Conversely in a negative
economy ablation exceeds accum ulation and the glacial
edges retreat. If over a period of years accum ulation and
ablation are approxim ately equal then the glacier is in
eq u ilib riu m and its volum e and lim its rem ain stationary.
Surface or superglacial ablation increases steadily towards
the lower lim its of the glacier and takes place m ainly as a
result of direct insolation and conduction from w arm er over-
lying air. M elting is proportionately m ore significant in
w arm er w etter regions such as New Zealand, whereas sub
lim ation tends to be m ore im portant un d er colder, drier
conditions such as are found in A ntarctica. However, even
where sublim ation is relatively im portant its total effect is
small. W indiness is a factor conducive to both sublim ation
and m elting by conduction. W here m elting dom inates,
m eltw ater streams may appear on the glacier surface,
although these tend to percolate beneath the ice or, in the
case of valley glaciers, move to the valley side as the snout
is approached. Englacial and subglacial ablation takes place
apparently as a result of heat generated by internal friction
or because of m eltw ater percolating from above. M elt
streams may start in the subglacial or englacial zones or they
may percolate from the surface. T h e m ost im portant direct
geom orphic effect of ablation is the in itiatio n of these m elt
water streams, which may become significant agents of
erosion in their own rig h t and may modify or supplem ent
in varying degree the effects of the glacial ice itself.
It is im portant to keep in m ind that glaciers with a
negative economy shrink by th inning as well as by recession
of the snout. T h in n in g is particularly im p o rtan t in that it
leads to the form ation of masses of passive or even dead ice
which produce characteristic landforms.
41 Diagrammatic representation of the economy of a glacier. The rate

of ice transport is greatest around the firn limit which approximately
divides the zone of accumulation from the zone of ablation and where net
gain and net loss are at zero.
Glacier zones. O n alm ost all glacier surfaces it is possible

to distinguish a zone of net accum ulation from a zone of
net ablation (Fig. 4 1 ). T h e exceptions include some
A ntarctic glaciers where there appears to be no significant
zone of ablation at all and some glaciers where there is no
firn, perhaps because they are reconstructed glaciers supplied
by ice avalanches from glaciers higher up, or because, as
may be the case in Baffin Land, the glacier has survived
from a form er tim e of active accum ulation by having
extrem ely low rates of ablation. W here zones of accum ula
tion and ablation may be identified they are separated by
the firn lim it, which is the lower lim it of the firn in sum m er
and lies somewhat lower than the local snowline because it
is being carried along by the glacier and because the rela
tively poor heat-conducting qualities of the underlying ice
Glaciers 87
allow snow to persist in higher air temperatures than would
normally be the case. In summer the two zones have a
markedly different appearance. Above the firn limit the
glacier surface is white: below it is a bluey-grey colour. In
the case of narrow valley glaciers the surface tends to be
concave in cross-section in the zone of accumulation but
convex in the zone of ablation, because of the effect of
radiation and conduction from the valley sides in speeding
up the rate of melting and also because of the greater
median velocity. On a valley glacier the two zones are often
distinguished as the firn basin and glacier tongue.
Glacier snout Firnfield ----- ►
Interval of altitude
42 Area-height curves of four valley glaciers (redrawn from Ahlmann,
1948)
There is a general relationship between the size of the

accumulation and ablation zones, so that glaciers with larger
zones of accumulation may be expected to have larger zones
of ablation and thus extend further below the firn line. This
relationship may be obscured by variations in regime
between one glacier and another and also by differences in
the character of the overall slope on which the ice body is
distributed. Ahlmann (1948), for instance, illustrated four
different types of area-height curve in the case of valley
glaciers (Fig. 42). But in dealing with closely distributed
mountain glaciers in a single highland mass it is often
possible to relate the length of the glacier snout to the size
of the firn basin so that neighbouring cirques of differing
capacity may be seen to be associated with valley glaciers
of differing length (Fig. 43) .
Even neighbouring glaciers may differ significantly in
their economy at any one time. Because of differences in
size, aspect, and altitude of accumulation and ablation zones,
negative and positive economies may occur within the same
general area so that some glaciers are advancing while others
are retreating. At present most of the world’s glaciers are
retreating but some are advancing.
2700m
. ' . 1650 m
5 Km ’ •
J '. .
1550m ' ■
43 Firn basins and glacier snouts of two Swiss glaciers. The glaciers
have a similar aspect and are both divided approximately at the 2700
metre contour; but the glacier with the greater area of accumulation extends
further towards sea level.
Glacier movement
Ice moves from the region of accumulation to the region of
ablation and its movement is dominated by gravity, but it
Glaciers 89
is important to note that motion is in a down-ice rather
than a down-slope direction. That is to say the glacier
virtually always moves in the direction in which its surface
slopes. In the case of ice streams this may give rise to uphill
movement at the base over short distances and in the case
of large ice sheets such movement may be of much greater
magnitude and of much more geomorphic importance.
Other things being equal the velocity of movement is
greatest around the firn limit because the volume of ice
increases steadily down to this point and then decreases
steadily below it (Fig. 41) . It is here therefore that maxi
mum ice transfer downstream becomes necessary and,
assuming that the available cross-section remains constant,
this must be achieved through maximum velocity. Changes
in the character of the channel — its cross-sectional area or
gradient — give rise to other variations of velocity super
imposed upon the general one; and in a valley glacier when
the valley narrows or its floor steepens the ice velocity is
locally increased.
Lines of e q u al velocity a t intervals
44 Calculated isopleths of cross-sectional velocity in feet per year of

the Saskatchewan Glacier (redrawn fro?n Meier , 1960)
In cross-section, ice velocity in a valley glacier is greatest

at the surface and in the centre line, decreasing relatively
sharply as the valley sides and bottom are approached (Fig.
44) . Measurements taken at the surface of the glacier may
therefore give some indication of the speed of transport of
superglacial rock debris but will give an exaggerated idea
of the rate at which the glacier is m oving in contact w ith
the rock surface and of the rate at which it may be trans
porting m aterial near its base.
’ Calculated velocity -
depth profile
Calculated flowlines
45 Calculated flowlines and velocity-depth profiles in a cirque glacier,

the Vesl-skautbreen, Norway (redrawn from McCall, 1960)
h Calcul»lated
/ velocit;
ty - depth
profile
46 Calculated flowlines and velocity-depth profiles in a valley glacier,

the Saskatchewan Glacier (redrawn from Meier, 1960)
Components of m ovement. T w o com ponents of move

m ent may be identified. T h e first of these is basal slip, in
which the ice slides over the rock surface w ithout necessarily
being internally deform ed, and the second is internal flow,
which takes place through m utual displacem ent of ice con
stituents, probably at m any different scales. A t different
points in different glaciers the proportion of m ovem ent
taking place in these two ways may vary considerably. A
glance at the velocity curves shown in Figs. 45 and 46 will
indicate that there is relatively little differential m ovem ent
in the upper layers of ice and that these are largely being
carried along by the ice beneath. T h e velocity at the surface
is the sum of the velocities of basal slip and internal flow.
From a geom orphic point of view it is im portant to note
again that the rate of surface transport depends on this
Glaciers 91
surface velocity, whereas the rate at which the glacier moves
over its bed — which in turn should influence its power of
erosion and basal transport — depends essentially on the rate
of basal slip alone. A glacier or section of a glacier in which
little movement takes place by basal slip is not therefore
likely to be carrying out much erosion.
The theory of basal sliding propounded by Weertman
(1957, 1960) has been generally supported by field and
laboratory evidence presented by Kamb and La Chapelle
(1964) . The major mechanism is thought to be connected
with regelation, the freezing and thawing of ice as a result
of changes of pressure. In glaciers where the temperature
of the basal ice is close to normal freezing point, meltwater
may occur because the pressure of the ice mass reduces the
actual freezing point. Any release of pressure will cause the
freezing point to lift and the water then freezes. The exist
ence of quite tiny irregularities on the rock surface may give
rise to a regelation process in which wrater produced by
melting on the upstream high pressure side of the irregu
larity flows around the protuberance and re freezes on the
downstream low pressure side. The thin zone of refrozen
ice, only a centimetre or two thick, in which this process
takes place may be called the regelation layer, and it is con
tinuing displacement in this layer which carries the main
ice mass along and is mainly responsible for the sliding
phenomenon. It seems likely that observed changes in ice
velocity in some glaciers between summer and winter may
be a result of the greater availability of meltwater at the
base of the ice during the warmer months. The thickness
of water involved needs only to be of the order of half a
millimetre.
Regelation slip is favoured by irregularities of minimal
size: where larger obstacles are involved, what Weertman
termed a stress concentration mechanism is thought to come
into play. It is envisaged that stress concentrations existing
near the obstruction lead to locally increased rates of plastic
flow and the obstacle is therefore bypassed.
Of the two mechanisms suggested as producing com
ponents of basal slip, it seems clear that regelation slip is
the fundamental one. Both theory and the very limited
amount of observational evidence available suggest that
basal slip is absent in very cold glaciers where regelation
cannot occur.
Several hypotheses have been put forward to explain
internal flow in glaciers and they are probably not mutually
exclusive. The major ones have been concisely reviewed
by Sharp (1960) . Current thought suggests that the most
important process is that of intragranular yielding and re
crystallisation in which gliding takes place along the basal
crystallographic planes of the ice crystals. This has been
shown to occur in laboratory experiments (Glen, 1962)
and is supported by the strong preferred orientation demon
strated by ice crystals within actual glaciers. Crystal form
is thought to maintain itself in the face of this distortion
by a progressive process of recrystallisation. Crystals with
planes parallel to the direction of ice flow tend to grow at
the expense of those which are not parallel: this produces
the preferred orientation which promotes internal gliding
and also increases crystal size.
Other suggested modes of flow include intergranular
adjustment, the mutual displacement of the ice grains them
selves, and regelation, involving melting under pressure and
recrystallisation in the downstream direction. Neither are
now thought to be generally significant but may be im
portant in the firn field. Another way in which glaciers are
known to move, particularly near obstructions or in their
terminal sections, is by internal slip along shear planes —
a sort of thrust faulting. This may not contribute much to
the total movement of a particular glacier but may prove
of rather more significance from a geomorphic point of view.
Extending and compressing flow. Nye (1952) put forward

the concept of extending and compressing flow in glaciers
and this concept has received considerable support since
then. Extending flow occurs where ice velocities are increas
ing downstream, as where the bed slope steepens, while
compressing flow occurs where velocities are decreasing, as
where the gradient becomes less. It is important to note
that it is the glacier which is extended and compressed and
not the ice. Changes in the volume of ice passing through
Glaciers 93
a particular section also have an effect on velocity and there
fore on the type of flow. In the zone of accumulation
velocities are normally increasing and therefore extending
flow is to be expected: conversely, velocities are normally
decreasing progressively in the zone of ablation, and com
pressing flow may be expected towards the terminus of the
glacier (Fig. 47) .
Nye also concluded that compressing flow would be
associated with an upward component of ice movement and
upward pointing internal slip planes, whereas in extending
flow the reverse would be the case.
47 Flow zones in a glacier according to Nye (1952). Extending flow

occurs where velocity is increasing as in the zone of accumulation and on
release from obstruction. Compressing flow occurs where velocity is decreas
ing as in the zone of ablation and behind an obstruction. The straight
lines indicate postulated directions of slip planes.
Surges. There are many recorded instances of bulges of

increased ice thickness moving down through the body of a
valley glacier at speeds greater than the normal velocity of
the glacier, rather like river floods. These glacier surges
result in a relatively sudden advance of the glacier snout.
In some cases they have been related to exceptional short
periods of accumulation, such as when there has been con
siderable avalanche activity in the firn basin. In other cases
no obvious initiating factor has suggested itself and it is
thought that surges may result from the existence of some
threshold amount of accumulation in a given basin, above
which sudden periodic evacuation of ice occurs.
Flow directions. It was noted above that the direction of

flow in plan deviates little from that of the steepest slope
on the ice surface. Even within valley glaciers it has been
shown (Meier, 1960, for instance) that a small sideways
com ponent of flow may exist in the zone of ablation where
cross profiles are norm ally convex. In the zone of accum ula
tion on the other h and there may be a tendency for an
inw ard com ponent to occur.
T h e direction of flow in elevation has long been deduced
to have a dow nw ard com ponent in the zone of accum ulation
and an upw ard com ponent in the zone of ablation (Reid,
1896) . A t the firn lim it it is m ore or less parallel to the
surface. M easurem ents m ade on glaciers of different types
appear to support this deduction in general, although it may
be, as pointed ou t by Sharp (1960), that the upw ard move
m ent near the snout is usually upw ard in relation to the
glacier surface rath er than to the horizontal. T h a t such
upw ard m ovem ent occurs is well know n from the way in
which rock debris carried w ithin the glacier emerges at the
surface near the snout and gives it a characteristic dirty
appearance. It also accords w ith N ye’s postulated slip plane
directions in compressing flow.
48 Presumed ice flow direction in the cross-section of an idealised ice

sheet. Indicated depression of the subjacent crust follows calculations made
by Weertman (1961).
Flow directions in large ice sheets are m ore difficult to

reconstruct b u t it is generally assumed that the centre is
occupied by slowly subsiding firn and the m ore plastic ice
flows out radially (Fig. 48). Surface gradients on ice sheets
are extrem ely low as a rule and the picture may be
com plicated by the occurrence of m ore than one region of
accum ulation.
Crevasses. Cracks appear in the ice because of differen

tial m ovem ent w ithin the glacier body and these may be
Glaciers 95
enlarged to form open crevasses reaching several feet in
w idth. T h e ir size and orien tatio n is related to the tensional
stresses w hich give rise to them , b u t their depth is rarely
m ore than about 50 m because of the tendency to closure
by the m ore plastic and m ore compressed lower ice. H ow
ever, in spite of the fact that most crevasses extend for
lim ited distances dow nw ard from the surface, there appear
to be some which form for lim ited distances upw ard from
the base. N ot m uch is know n of basal crevasses b u t they
may be located principally near the glacier margins.
F lin t (1957) recognised five m ain groups of crevasses.
Transverse crevasses form on the outside of bends and where
the bedslope steepens so that extending flow occurs. T h e
m ost spectacular expression of these is at an ice f a l l where
the glacier moves over an exceptionally steep section of its
bed. Longitudinal crevasses occur where the glacier suddenly
becomes less confined and is able to spread laterally, and
these are related to radial crevasses which form where an
expanded foot glacier emerges on to a piedm ont. Marginal
crevasses are in h eren t in the differential flow in the glacier
cross-section and form chevron patterns in which the apex
is upstream . Bergschrunds are crevasses form ing in the hrn
field a short distance in from the rear edge of the glacier.
A lthough the detailed study of crevasse systems is pri
m arily a m atter for those interested in the physics of ice,
the existence of such systems is of geom orphic im portance
in a n u m b er of ways. Crevasses provide routes along which
englacial w ater moves, and dow nw ard stream ing water may
cause local enlargem ent in to a ro u n d hole know n as a
glacier mill or m oulin. T hey also form gaps into which rock
debris penetrates. T h u s m aterial falling on to the glacier
from above may accum ulate in surface crevasses and in some
circumstances m aterial beneath the glacier appears to be
m oved u p in to basal crevasses. As will be discussed later,
transverse crevasses and bergschrunds have been invoked as
co n trib u tin g factors in the location of zones of excessive
erosion. T o a lim ited extent therefore the patterns of
glacier crevasses may be reflected in patterns of erosion and
deposition.
Glacier types
Glaciers may be classified according to their form, as has
been outlined earlier, and the major factor in this is the
preglacial morphology. But, they may also be classified in
other ways so as to bring out differences which are largely
climatic in origin and virtually independent of the previous
landscape. Ahlmann (1948) summarised his geophysical
classification of glaciers as follows:
I. Temperate type — in which there is relatively rapid

change of snow to glacial ice, predominantly through exten
sive melting and recrystallisation. Relatively little firn is
therefore present at any one time in the accumulation zone.
Throughout the profile the temperature of the ice is close
to melting point, except for some surface freezing in winter.
II. Polar type — in which snow changes relatively slowly
to glacial ice, predominantly or entirely through compaction.
Great depths of firn exist with temperatures below freezing
point. A subdivision may be made into:
(a) High-polar type with deeper firn and little or no
melting even in summer.
(b) Sub-polar type with shallower firn (to about 20 m)
and surface melting for short periods during summer.
It should be noted that it is possible, but unusual, for
polar-type glaciers to occur in temperate latitudes and for
temperate-type glaciers to occur in high latitudes. It more
commonly happens that a large and complex glacier differs
in type in different sections of its mass and this may make
detailed classification difficult; but from a geomorphological
point of view the broad differences suggested remain of
great importance.
Dynamic classification. It is also possible to classify

glaciers according to their degree of dynamic activity, and
Ahlmann (1948) referred to active, passive, and dead
glaciers. In this sense a dead glacier is one in which no
transfer of ice from accumulation zone to ablation zone is
taking place. Any movement is due solely to the slope of
17 Cirque glacier on Baffin Island with well-defined ice cored end
moraines (]. D. Ives)
IS The Tasman Glacier, New Zealand, from Anzac Peak (N.Z. Geological
Survey photo by D. L. Homer)
19 Edge of the Antarctic ice sheet with nunataks near Mawson (ANARE
photo by J. Bechervaise)
20 Shelf ice and icebergs on the coast of Antarctica (ANARE photo by

Royal Australian Air Force)
Glaciers 97
the bed and is in the nature of a settling process. While
it should be possible to define a dead glacier in such an
absolute fashion, active and passive glaciers can only be
distinguished in a relative sense. It will be obvious, however,
that, from the point of view of landform evolution, the
distinction is an important one.
The rate of flow in glaciers varies considerably. It may
be of the order of several metres a day in some glaciers but
little more than about a metre a year in others. It has been
calculated that snow accumulating in the centre of Antarc
tica may take something like 50,000 years to reach the sea
in some directions. That velocity varies with gradient and
ice thickness and may change significantly within a single
glacier, has already been noted. The most important reason
for the difference in velocity between glaciers is the varia
tion in the actual rate of turnover of ice — that is to say
the absolute amount of accumulation and ablation which
takes place. Glaciers such as those of New Zealand, which
exist largely because of high rates of precipitation, also have
high rates of ablation because summer temperatures in these
latitudes are relatively high. There is consequently a high
rate of movement of ice through the glacier bodies. In
contrast, glaciers such as those of Antarctica, which exist
primarily because of the low summer temperatures, also
have very low rates of accumulation. The rate of movement
is therefore low.
This distinction is independent of whether the regime
is positive or negative, for even where the regime is negative
and the glacial snout is actually retreating there may still
be high absolute rates of accumulation and ablation and
consequently a rapid transfer of ice through the system. A
long continued negative regime, however, will result in
thinning of the glacier and some reduction in velocity will
eventually occur as a result.
In general terms it could be expected that temperate-
type glaciers will tend to be active ones and polar-type
glaciers tend to be passive. This is because the climatic
factors affecting the geophysical distinctions are much the
same as those affecting the dynamic differences. But other
non-climatic factors prevent a complete correlation.
Glacier types and glacial geomorphology
T h e recognition that different glaciers vary in their physical
and dynam ic attributes is of very great im portance in the
study of glacial landform s, for it means that processes
operating in the case of one glacier may not necessarily
operate in the case of another, or at the least that they may
operate in different ways or w ith different degrees of in te n
sity. It means also that not only is it necessary to have some
idea of the types of present-day glaciers b u t it is necessary
to m ake at least an inspired guess as to the types of extinct
glaciers which produced glacial landform features d uring
the Pleistocene. T h e large continental ice sheets of Pleisto
cene Europe and N orth America, for instance, have no true
counterparts today b u t they were m uch m ore like the
southern part of the G reenland ice sheet than like the
glaciers of Antarctica.
Knowledge of the attributes — and particularly the flow
attributes — of different glacier types is as yet too ru d i
m entary to allow of m ore than initial deductions, b u t
enough is probably know n to allow some broad generalisa
tions. T h e most extrem e contrast is betw een active
tem perate-type glaciers and passive high-polar-type glaciers.
In the former, ice velocities are high, m eltw ater is abundant,
and a large p art in ice m ovem ent is played by basal slip:
in the latter, ice velocities are low, m eltw ater is rare or
absent and so is basal slip as a com ponent of m ovem ent.
Both glacial and glacifluvial rates of erosion and transport
are therefore likely to be higher — probably m uch higher
— in the first case than the second. Between these two
extrem es all sorts of grades of variation are possible.
T h e existence of great variations in the am ount of geo-
m orphic work done by different glaciers or by different parts
of the same glacier at different times is the reason why it may
appear desirable sometimes to use the term glacierisation
to indicate the covering of a landscape by ice. In possible
distinction, glaciation implies some m odification of the land
surface. It is probably the m ajor reason too for the existence
of different schools of thought in relation to the efficiency
of glacial erosion and in particular for the argum ents
Glaciers 99
betw een ‘erosionists’ and ‘protectionists’. It seems logical to
assume th at a region occupied by passive glaciers and which
is being glacierised rath er than glaciated is in the m ain
being protected from erosion, since this w ould presum ably
proceed m ore rapidly in a fluvial regime. O n the other
h an d it is abundantly clear from the evidence they have left
b ehind th at the most active glaciers are capable of very great
rates of denudation indeed.
W ith in the A ustralasian region it seems that the glaciers
of w estern T asm ania, existing in conditions of high pre
cipitation and low snowlines, were of an active tem perate
type sim ilar to the present-day glaciers of the South Island
of New Zealand. T h e glaciers of central and northeastern
Tasm ania, and also possibly of the Kosciusko region of New
South Wales, are likely to have had a notably lower rate of
ice turnover since they existed in d rier regions and, in
T asm ania at least, their significantly higher snowline m ust
have m eant lower rates of ablation. T h e difference in extent
of glacial m odification of the landscape betw een western
T asm ania and the other areas is probably partly due to the
difference in glacier type, although the effect of differences
in the length of the period of glaciation and the relationship
of the snowline to topography m ust also have played a part.
In contrast, m uch the greater proportion of the present
A ntarctic ice mass is of passive high-polar type. Ice velocities
are very low and m eltw ater relatively scarce even in
sum m er. It may be expected that very little, if any, basal
slip is occurring beneath these glaciers, and reports of fossil
pattern ed ground being uncovered by retreating ice fronts
(Stephenson, 1961) support the view that, in some areas
at least, the rate of landscape m odification is very low.
In addition to there being geographical differences in
glacial type there will also be historical ones; that is to say
a given glacier may change in type d u rin g its life span. T h e
large continental ice sheets in particular vary in very com
plex fashion through their history and through their extent,
so that the landform s which they leave behind are com m only
related to particular phases of their existence, and particular
sections of the area glaciated. T h e picture is m ade m ore
com plex by the characteristic asymmetry of expansion and
contraction brought about by clim atic variations over the
area occupied by these very big glaciers (Fig. 49) .
200 M
Years before
present
2 0 0 Km
49 Stages in the retreat of the Pleistocene Scandinavian ice sheet.

Asymmetry of development is characteristic of continental glaciers.
As ice sheets wax and wane, they themselves do m uch to

change the climate of the region in which they exist. It
has been widely thought that the largest continental glaciers
of the past may have created their own m ore or less
perm anent high pressure areas and reduced their rates of
accum ulation so as to impose a lim it on their m axim um size.
VII
GLACIAL PROCESSES
Because of the opacity of glacier ice, it is relatively difficult

to watch geom orphic work being carried ou t by a glacier.
Both the erosional and depositional results of such work can
be seen, b u t com paratively few direct observations have been
m ade, by people entering n atu ral ice caves and crevasses
and artificial tunnels, of the processes whereby these results
are attained. F urtherm ore, whereas the work of streams,
waves, and winds can be readily studied in laboratory
models, glacial processes are difficult to sim ulate because of
the m uch greater scale problem s involved.
Several w riters have com m ented that observations of
work done suggest conclusions differing from those to be
draw n from observations of work being done. M any geo-
m orphological studies show clearly that glaciers can be
extrem ely effective agents of erosion whereas many, if not
most, glaciological studies seem to indicate that the am ount
of rock destruction to be expected from glacial ice is very
little. T h e apparent contradiction may be a result of u n d er
estim ation of the tim e factor: it may also have come about
because present-day glaciers mostly have a negative economy
and their geom orphically im p o rtan t work is done at other
phases of their life history. It is very likely to be due, at
least in part, to the small am ount of observation which has
been carried out into active glacial erosion. In any event
m any of the ideas about glacial processes sum m arised in this
chapter are derived less from actual observation than from
a com parison betw een w hat glaciers may be seen to have
done and w hat they are considered theoretically likely to
do. Hypotheses m entioned here are those which appear to
be currently held: reference should be m ade to Cotton
(1947) and, particularly, Charlesw orth (1957) for historical
101
discussion of ideas no longer widely favoured or proven
incorrect.
Glacial processes in geom orphology can be discussed con
veniently under the headings of corrasion, transportation,
and deposition b u t it m ust be rem em bered that these are
the positive effects of ice action. G laciation may also be
thought of as having a negative effect. It has already been
noted that, by their bulkiness and extensive ground cover
age, glaciers in h ib it other forms of subaerial denudation,
and it is im portant therefore to think of the lim its of
glaciation not only as the outer lim its of glacial action b u t
also as the in n er lim its of periglacial and nivational action
in particular.
Corrasion
As far as the ensuing landform s are concerned the most
outstanding characteristic of glacial corrasion is its dual
natu re and the contrast betw een the process of abrasion on
the one hand and that of plucking or quarrying on the other.
Abrasion. Glacial abrasion is a grinding process con

trolled by m ovem ent of the base of the glacier against the
rock surface. Because glacial ice is of considerably lower
density and hardness than rock it seems certain th at it is
the rock particles held in the base of the ice that are the
operative tools, and extensive abrasion m ust be associated
w ith the existence of a large load of coarse basal m aterial.
However, it may be assumed that the load cannot be too
large or deposition is likely to occur instead: some optim um
condition seems necessary. T h e size of transported particles
and the fineness of the grain in the bedrock are im portant
factors governing the character of abrasion. In particular,
the am ount of abrasion is generally proportional to the area
which individual particles present to ice pressure from
above, and McCall (1960) calculated that it w ould be
independent of ice depth w hen depths are greater than
22 m. Increasing ice velocities have no significant effect
on abrasive forces: however, they m ust be im portant in
controlling the rate of supply of fresh rock tools.
Glacial Processes 103
Polishing of surfaces seems to be associated with an
abundance of silt grains, but coarser material may etch
striations on the rock surface and also grooves which are
enlarged and deepened striations resulting from repeated
channelling of basal debris along the same line. Some rocks
appear too hard to take striations: others are too susceptible
to chemical weathering to retain them for long after the
glacial cover has been removed. Striations are rare in
glaciated Tasmania for instance, apparently because, of the
two commonest rock types, the quartz metamorphics were
too resistant and the dolerite has weathered excessively in
the post-glacial period. Where striations are numerous they
afford valuable clues to the direction of former ice move
ment. On some rock types crescentic marks develop trans
verse to the direction of ice movement. These are supposedly
due to sudden and perhaps jerky local changes in friction
between the ice and hard brittle rocks: they include
crescentic gouges, concave up and downstream, and crescentic
fractures concave downstream (Flint, 1957).
In spite of the tendency to produce these small-scale
irregularities the overall effect of abrasion is to smooth and
round rock surfaces — a process termed mamillation. Such
smoothing is especially evident on surfaces facing upstream.
The precise result will obviously depend to some extent on
rock type, but the rounding process is characteristic of a
wide range of rocks (Linton, 1962). The glaciated edges
of dolerite-capped plateaux in Tasmania show clear evidence
of rounding by the passage of ice, and Jennings and Ahmad
(1957) used this as a major clue in the reconstruction of
former ice movements.
Some crystalline rocks, however, may produce a pseudo-
mammillated appearance and granites in particular are well
known to display such a rounded form under a variety of
conditions of denudation. Many apparently mamillated
surfaces in glaciated districts may in fact owe much of their
appearance to preglacial rather than glacial processes. Caine
(1967b) has suggested that the extensively mamillated
surfaces of Ben Lomond in Tasmania may represent largely
the basal surface of weathering in the dolerite, exposed
when the ice scraped away the overlying regolith. Such
surfaces may not have suffered intensive glacial abrasion in
spite of their smoothed and rounded appearance. Similar
surfaces may in fact be seen emerging from beneath peri-
glacially redistributed waste on unglaciated highlands such
as Mt Barrow and Mt Wellington.
Plucking. The plucking or quarrying process involves

the pulling away of relatively large particles of rock and
their incorporation into the base of the glacier. Of necessity
this occurs mainly on downstream facing slopes and con
siderable doubt exists as to its precise nature, since tunnel
ling has shown that, at least in some cases, the ice is not all
in direct contact with the rock surface at such places. The
existence of planes of weakness in the rock material seems
a prerequisite for plucking, and well-jointed rocks are
especially well suited to its operation. Previous weathering
along the joint planes is likely to facilitate the process still
further.
One important facet of the plucking process may simply
be the drag force exerted by the ice passing over a cliffed
face and thus causing semi-loosened rock to move down
stream. That this may occur even at a very large scale is
suggested by the existence beneath glaciated dolerite scarps
in Tasmania of enormous ‘hinge blocks’ originally isolated
by deep weathering along major lineaments and subsequently
pulled away from the scarp face apparently by ice moving
over it from above.
However, not all plucked surfaces display such an obvious
association with previous joint block isolation, and other
mechanisms may need to be invoked. W. V. Lewis (1954)
suggested that rock particles might be loosened by the forma
tion of dilation joints parallel to the surface where rocks
formed under pressure are unloaded by erosion. The
phenomenon has been described, for instance, by Jahns
(1943) in the case of granites. In this way glacial erosion
may be to some extent self-perpetuating since, as material
is removed, exfoliation prepares further loosened particles
for incorporation into the body of the glacier. Such a pro
cess is likely to be more effective in the case of glacial rather
than fluvial erosion because glaciers are able to remove
com paratively large volumes of m aterial relatively quickly
over a broad front.
T h e way in which plucked m aterial is actually taken into
the glacier and so rem oved is linked w ith the process of
regelation described earlier. Supercooled w ater at the base
of the glacier, freezing w hen a local decrease in ice pressure
raises the freezing point, may incorporate loose particles
which it surrounds into the m ain ice mass. Kamb and
La C hapelle (1964) noted that the regelation layer which
they observed at the base of an ice tunnel was heavily loaded
w ith debris in com parison w ith the ice above and McCall
(1960) recorded th at the lowest debris-laden layer of ice in
a Norw egian cirque glacier was form ed directly from re
freezing of w ater and not from com paction of firn.
Several workers have suggested that regelation may be
responsible also for loosening of rock particles when super
cooled w ater freezes in rock crevices.
T h e volume of m aterial rem oved by plucking is norm ally
very m uch greater than that rem oved by abrasion and this
is particularly tru e in the case of w ell-jointed rocks, and also
probably in the earlier phases of glaciation when greater
quantities of rock particles loosened by preglacial w eathering
are available.
Stoss and lee effect. T h e dom inance of abrasion on u p

stream slopes and of plucking on dow nstream slopes pro
duces a very characteristic appearance in hard rock country
that has experienced glaciation, especially of course where
plucking has been a significant com ponent of erosion (Fig.
50). If one looks in the direction in which the ice travelled,
the smoothed abraded stoss slopes are visible: if one looks
in the direction from which the ice came, the rough plucked
lee slopes are visible. Such stoss and lee topography occurs
on a m u ltitude of scales and can appear either in plan or
in elevation. In the dolerite highlands of T asm ania, where
the m ajor am ount of glacial erosion took place by the pluck
ing of jo in t blocks, the identification and m apping of stoss
and lee effects is the most im portant way in which directions
of form er ice m ovem ent can be reconstructed (Derbyshire
and others, 1965) .
Fine Coarse
particles particles -
produced produced
50 The stoss and lee effect. Both this and the bimodal nature of till
arise from the characteristic dichotomy in glacial corrosion.
Production of bimodal sediments. The commonly bi

modal character of particle size in sediments produced by
glacial corrasion is a direct result of the dichotomy in the
corrasion processes (Fig. 51). Abrasion produces fine
6 6
Diameter in millimetres
51 Particle size analysis of a Pleistocene till from the upper Mersey

valley, Tasmania (redrawn from Spry, 1958)
material, commonly of clay size. This is the ‘glacial flour’
that gives a typically milky appearance to meltwater streams.
Plucking on the other hand produces much larger particles
— usually pebbles or boulders. A characteristic product of
the combined operation of both processes is the type of
deposit known, particularly in the British Isles, as ‘boulder
clay’.
Efficiency of glacial corrasion. It is well known that the

efficiency of glacial corrasion appears to vary enormously,
not only between one glaciated area and another but also
within the area covered by the same glacier. Some of these
variations are obviously due to differences in lithology and
in the degree of previous weathering that has taken place,
since these affect the erodibility of the bed and the nature
of the basal material forming the rock tools. Others are
due to the type of glacier involved — its degree of activity
and the extent to which it moves in contact with its bed by
basal slip. Yet other variations seem to be associated with
particular zones within the glacier-covered area and especi
ally zones of differing ice thicknesses and ice velocities.
Some zones of maximum velocity are associated with changes
in topography such as constrictions and steepening of slopes:
these can be expected to remain substantially unchanged
throughout the period of glaciation. Some are related to
the regime of the glacier itself and will change position as
the glacier grows and shrinks. Most important of these is
the position of the firn limit, which is the zone of maximum
ice transport and approximates to the zone of maximum
velocity if topographic factors are equal. This will move
back and forth along the course of the glacier as ice volumes
change, but its position at times of prolonged glacier
equilibrium is likely to be a favoured location for corrasion.
Subglacial chemical weathering. In contrast to the de

duced effectiveness of some forms of mechanical weathering,
chemical weathering beneath glaciers is generally thought
to be minimal except on limestone terrain. The prevalence
of temperatures at about 0°C and the general absence of
organic acids are im portant factors in in h ib itin g chemical
action, b u t some forms of w eathering — w etting and drying
for exam ple — may be m ore im portant than has been
realised.
Transportation
T h e m aterial transported by a glacier is derived from two
m ain sources. Some of it results from corrasion and is in
corporated in the first place in the basal part of the glacier:
the rem ainder results from w eathering and mass m ovem ent
on exposed surfaces above the glacier and falls on to the
ice surface where it is carried along rath er as on a conveyor
belt. In the case of cirque and valley glaciers the am ount
of m aterial com ing to the glacier from above may be very
large and exceed that com ing from below. Conversely, in
the case of plateau glaciers and larger ice sheets, only the
occasional island of rock, or nunatak, will provide a source
of m aterial from above, and virtually all m aterial carried
will have been corraded by the glacier itself.
Glacier transportation induces little or no sorting in the
m aterial being transported.
Material from below. M aterial corraded by the glacier

and taken into its basal section forms a ground moraine. It
rarely appears to travel great distances and m uch of it is
lost by attritio n or lodgem ent fairly soon. O nly near the
snout of the glacier does it tend to work its way upw ard
towards the ice surface as a result of the dom inant direction
of ice m ovem ent in the zone of rapid ablation. It can be
deduced too that the rate of travel of basal m aterial will
be relatively slow since its velocity will closely approxim ate
that of the glacier sliding over its bed, which is always less,
b u t not necessarily very much less, than the velocity of the
ice surface.
T h e re is evidence that fine saturated basal m aterial of
sufficient plasticity may be squeezed up into fissures in the
bottom of glaciers by the weight of ice on either side. Such
plastic flow may take place into basal crevasses or into
tunnels opened by m eltw ater streams.
M aterial from above. Rock m aterial m oving on to the
glacier surface from valley sides or nunataks is generally a
result of w eathering and nivational processes on the exposed
rock surfaces. T h is m aterial accum ulates on the glacier
surface adjacent to the rock wall in the form of a lateral
moraine. T h e re may be an accession of debris eroded along
the periphery of the glacier b u t in general the size of the
lateral m oraine is a function of the rate of supply of m aterial
com ing from above and the velocity of the glacier.
-----Flow lines
__ Ablation surfaces computed
at 10 year intervals
, . . Debris patch at 10 year
intervals
Metres
52 The movement of surface moraine through a cirque glacier (redrawn

from McCall, 1960)
W hen two glaciers join, their adjacent lateral m oraines

coalesce to produce one m edian moraine, so that glaciers
w ith a n um ber of tributaries carry an equal n u m b er of
m edian m oraines. T h e position of the m edian m oraine
depends on the relative volum e and velocity of the joining
glaciers. W here inset or superim posed tributaries term inate,
a transverse m oraine may form and com plex looped patterns
may eventuate w here glaciers of different sizes and different
regimes meet.
Both lateral and m edian m oraines are visible only on the
glacier tongue and norm ally have subsurface extensions so
Superimposed
Juxtaposed
Englacial
53 Tributary ice streams and associated morainal systems (based on

Sharp, 1960)
th at they are often, perhaps norm ally, contiguous w ith the

ground m oraine beneath. In the glacial zone of accum ula
tion, surface m oraines are covered by firn and m ovem ent
occurs along the flowlines in the ice, which here may be
expected to be dow nw ard in relation to the glacier surface.
Figure 52 shows how rock debris falling on to the upper part
of a cirque glacier moves dow nw ard into the ice body and
then subsequently emerges near the snout. In this way
trib u tary glaciers may rem ain defined by well-marked
m orainal septa for long distances past their point of junction.
Figure 53 illustrates the side-by-side, inset and superim posed
conditions distinguished by Sharp (1960) .
Because of the higher ice velocities near the surface of a
glacier it is to be expected that surface m oraines will move
at a significantly faster rate than ground m oraines: m edian
m oraines may be expected to move fastest of all because they
are located directly in the position of greatest ice velocity
(Fig. 44). Valley glaciers that have received num erous
tributaries charged w ith surface m oraines will therefore be
transporting exceptionally large quantities of m aterial rela
tively rapidly along their centre lines of flow. It is glaciers
such as these (for instance in New Zealand) that display
excessively dirty term inal zones as their m edian m oraines
spread, coalesce, and rise tow ard the glacier snout.
Deposition
M aterial directly deposited by glacial ice is term ed till and,
since the processes of glacial deposition are as little inducive
to sorting as those of glacial transportation, it is charac
teristically nonsorted. Basal till, derived in the m ain from
glacial erosion, is norm ally bim odal because of the dual
n atu re of the corrasion processes. Drift is an all-inclusive
term sometimes used for all those deposits resulting from
glaciation — not only till b u t m aterial em placed by icebergs,
m eltw ater, and so forth. T h e word moraine has been used
on occasions as a synonym of till b u t is norm ally used to
indicate a body of m aterial, generally com prising till b u t
com m only w ith some other forms of d rift incorporated. It
is in this second sense that it is used here. T h e term s till
and drift, therefore, refer to sediments, and the term
m oraine to a body of sedim ent which as previously indicated
may be in transport, b u t when deposited becomes a land-
form. Some m oraines — lateral and ground m oraines —
may exist as both m oving and deposited forms. M edian
m oraines are alm ost always m oving and are rarely identifi
able as deposits. O th er m oraines, such as the end m oraines
dealt w ith below, only exist as depositional forms.
For a long tim e, and certainly since the review by
C ham berlin (1894), it has been custom ary to envisage
glacial deposition as occurring in three m ain ways, which
may be term ed dumping, lodging, and pushing, and the
broad distinction betw een these three rem ains im portant
from a geom orphological point of view.
Dumping. When glacial ice melts, the rock material
which it carried is let down on to the ground surface
beneath. In the case of a stable glacier, in which the position
of the terminus is more or less stationary, dumping takes
place at the edge of the ice so that a ridge-like body — the
end moraine — is constructed parallel to the glacier margin.
If the glacier has a negative economy and its terminus is re
treating, dumping will take place progressively rearward so
that a sheet of ablation till is deposited. If glacier retreat
should cease for a time, then another end moraine may be
constructed. It is usual to distinguish the moraine ridges
marking the outer limit of glaciation as terminal moraines
and those marking such temporary halts as recessional
moraines, but in practice, in field studies of formerly
glaciated terrain, it is sometimes difficult to be sure of the
status of a particular ridge, and some ‘recessional’ moraines
may in fact result from readvances. In the Broad River
valley at Mt Field National Park in Tasmania a number of
recessional moraines can be identified, but ground moraine
extends past the furthest one and there is no apparent sign
of a terminal moraine marking the outer limit of ice
advance.
Till deposited by dumping is likely to be loose and un
compacted and is often largely devoid of fines, because these
have been removed by the meltwater associated with
ablation. Because the ice at the moment of deposition is
virtually stagnant and has no significant forward velocity,
streamlined structures and forms are likely to be absent:
material is emplaced by having settled more or less vertically
downwards.
The process of dumping may be associated with the pro
duction of depressions termed kettles. When blocks of
wasting ice become detached and incorporated into a mass
of drift their subsequent melting causes subsidence if the
ice was buried or, more simply, a void if it was not. In
either event depressions result, the size of which may vary
considerably although they are usually rounded and rela
tively shallow. Kettling occurs in sediments deposited by
meltwater as well as in till.
In some circumstances ablation till, which incorporates
21 An outlet glacier at the edge of the Antarctic ice sheet descending
into a frozen lake in the frost rubble zone near Mawson (ANARE photo
by P. G. Laic)
22 Pirnfie Ids of the Southern Alps, Neiv Zealand, feeding the Franz
Joseph Glacier (V. C. Browne)
25 S u m m er c o n d itio n s on g la c ie r tongues descending fro m a tra n se ctio n
system on B a ffin Is la n d . N o te th e f r n lim it , s u p e rg la c ia l, m a rg in a l, a n d
te r m in a l d ra in a g e stream s a n d m u lt ip le h a ir p in - ty p e e n d m o ra in e ( ] . D . Ives)
24 Crevasse systems a n d surface m o raine s on th e D a rt G la cie r, O tago

A lp s , N ew Z e a la n d (N .Z . G e o lo g ic a l S urvey)
a large am ount of saturated fines, may flow off the glacier
surface instead of being dum ped in the usual way. Such
ftow till, as it has been called by H artshorn (1958) , is likely
to overlie or become interstratified with ice contact sedi
m ents at the side of the glacier.
Lodging. It has already been noted that m aterial carried
in the basal section of glaciers appears to move only rela
tively short distances. M uch of it is lodged in rock crevices
or depressions or in the d rift which has accum ulated pre
viously. C um ulative lodgem ent may produce till sheets of
considerable depths, given enough m aterial and enough
time.
O nly very local m elting under pressure is involved in
the freeing of particles for lodgem ent, and the ice mass itself
continues to travel over the lodged particles. Lodgem ent
till therefore is likely to be m uch m ore compact and to
contain m ore fines than ablation till produced by dum ping.
Since it is deposited by actively moving ice it is also likely
to display stream lining of structure and form. A fissile
structure is often induced in the fines and is due apparently
to the successive plastering of very thin layers of till.
Pushing. An advancing ice front may bulldoze loose

m aterial in its path to form a push moraine. Such m aterial
itself may be of glacial origin and end m oraines are often
rew orked in this way; b u t it may also derive from a great
variety of proglacial and even non-glacial sources. T h e
character of the m aterial emplaced by pushing will therefore
vary widely, but typical push m oraines seem to display
asymmetrical cross profiles and internal fold and fault
structures related to the deform ation suffered.
Overriding. A lim it is set to pushing by the size of the

debris mass against which the ice is advancing. T h u s push
m oraines are rarely much m ore than about 80 m in height.
Masses of greater height cause the glacier to ride over them .
T h ru s t structures due to overriding may also be produced
in other d rift deposits and even in unconsolidated sedim ents
of non-glacial origin. Low fold ridges transverse to the
direction of ice m ovem ent have been widely described from
n orthern Germ any and the N etherlands. Such structures
need to be distinguished from those of true tectonic origin
and also from contortions due to periglacial processes. Some
thrust structures in end m oraine complexes are held to be
inherited from the disposition of m oraine m aterial along
thrust planes in the ice prior to its m elting (Slater, 1926) .
Relationship of processes. T h e results of different deposi-

tional processes can be found in close association. End
m oraines may display the effects of dum ping, b u t also of
lodging and pushing, and a great variety of sequences is
possible. G round m oraines may result from dum ping and
lodging, and here the comm on sequence is predictable from
the conditions giving rise to the two processes. Lodging
may occur with both positive and negative economies but
is especially induced by the conditions of glacier advance.
D um ping to produce a till sheet is essentially associated
w ith glacial retreat and the very last phases of glacial activity
in a particular area. A blation till may therefore be expected
to overlie lodgem ent till where the two occur in ju x ta
position. A blation till is also likely to be th in n er than
lodgem ent till because it consists only of the m aterial being
carried by the ice at the m om ent of dum ping. As indicated
earlier, lodging can continue for very long periods so that
lodgem ent tills may attain great thicknesses. Most ground
m oraines and particularly those of the great continental ice
sheets contain m uch m ore sedim ent deposited by lodging
than by dum ping.
Till fabric. Boulders and stones carried along in flowing
ice tend to lie w ith their long axes parallel to the direction
of flow. W hen such m aterial is deposited this orientation
may be preserved so that it is revealed by detailed exam ina
tion of the till fabric. Holm es (1941) has described m ethods
of exam ining and analysing the fabric of tills, and the
mechanism of orientation has been discussed by Glen,
Donner, and W est (1957). As m ight be expected, lodge
m ent tills display particularly good alignm ent of constituent
stones, with a large proportion oriented along the flow axis
at the time of deposition. In ablation tills the fabric is not
nearly so well organised and only some of the larger pebbles
and boulders retain flow directions, the greater num ber of
constituents having been reoriented by the dow nw ard
settling which accompanies dum ping. In practice then it is
the fabric of lodgem ent till in ground m oraine which may
be profitably studied w ith a view to determ ining directions
of form er ice m ovem ent. Such study has proved especially
useful where continental glaciation has produced successive
tills laid dow n by ice travelling in different directions.
T h e re is an obvious analogy betw een the orientation of
stones in a till fabric and that in the fabric of solifluction
earths (Fig. 13, p. 34). In both, a great num ber of larger
constituents may be oriented parallel to flow direction at
the tim e of deposition and a secondary transverse o rien ta
tion, perhaps due to rolling, may appear. T h e im portant
distinction is probably that in the periglacial deposits the
preferred orientations are always parallel or transverse to
the slope of the land, whereas in the glacial deposits this
will only be so when the direction of ice flow coincided w ith
the steepest gradient.
T ill stones. A study of the individual till stones may be

of geomorphological significance in helping to elucidate the
provenance of the body of sedim ent in which they are found.
Stones in superglacial till are little altered by transportation
b u t basal till stones tend to become rounded, presum ably
as a result of rotation while being held by the ice. A large-
scale study reported by Holm es (1960) suggested strongly
that an ovoid form is the expectable final result. A small
proportion of stones in basal till may become faceted or
soled by being ground against bedrock while being carried
in the base of the glacier. Some stones may become crushed
betw een larger ones when sufficient ice pressure is present:
others may be fractured through pressure against bedrock
and a small proportion may show striations due to a sim ilar
cause. In large part the extent to which faceting, crushing,
fracturing, and scratching may take place is obviously a
function of lithology.
Some of the largest boulders may lie free on the surface
of the till or even on bedrock, where they are sometimes
dum ped in positions of such instability that they are easily
rocked. Free boulders reaching several thousand tons in
weight become m inor landform s in their own right.
W hen a till stone is carried on to bedrock of different
type, it becomes an erratic and, when its place of origin is
known, an indicator. T h e identification and plotting of
erratics and indicators have long been used in country of
varied lithology for the reconstruction of form er ice move
ments, and earlier accounts of Pleistocene glaciation in the
Snowy M ountains of New South W ales relied to some extent
on such evidence. In the dolerite-capped highlands of T as
m ania, uniform ity of bedrock over relatively large distances
has very m uch reduced the value of this potential technique
for extensive areas of glaciated land. In order to argue
closely from the occurrence of erratics and indicators, a
thorough and com plete knowledge of lithological d istrib u
tion is necessary and this is often difficult in glaciated terrain
because of the existence of d rift covers. It is also necessary
to distinguish and bear in m ind other processes likely to
give rise to erratics and, in particular, rafting by floating
ice and periglacial mass m ovem ent over low angle slopes.
W hen a num ber of indicators can be traced back to one
outcrop it is possible to plot an indicator fan, sometimes
called a boulder train.
M eltw ater processes

In tem perate-type glaciers, superglacial, englacial, and sub
glacial m eltw ater is comm only present in association with
the ice, and du rin g the retreat phase m ust also be associated
with those of polar type. Glacial deposition is thus often
associated with the presence of m eltw ater, which may modify
the glacial deposits or cause them to be interm ingled or
interstratified with wash m aterial. A nd so till grades into
water-laid sedim ents in a variety of ways and it is in such
circumstances that the term ‘d rift’ probably attains its
m axim um degree of usefulness. Furtherm ore glacial
erosion is amplified or modified by the action of subglacial
m eltw ater streams. However, the most im portant effects of
m eltw ater are in the proglacial zone, and discussion of them
is left to the next chapter.
V III
PRO G LA C IA L PROCESSES
T h e effects of glaciation are not lim ited to those created by

the action of the glaciers themselves, for water and w ind
may extend these effects for considerable distances beyond
the lim its of ice advance. These extended effects are referred
to here as proglacial, but, in discussing the processes in
volved, it is convenient to include certain ice contact
phenom ena, which are not strictly proglacial b u t yet owe
their existence to the same group of agencies. M uch the
most im p o rtan t of these agencies is the m eltw ater em anating
from the glaciers.
Glacifluvial processes
All glaciers release m eltw ater, b u t the volume and tim ing
of release varies enorm ously with the type of glacier. It has
already been noted that tem perate-type glaciers with high
rates of accum ulation and ablation tend to produce very
large am ounts of m eltw ater at all stages of their history,
even when ice m argins are m oving forward with a positive
economy. In the case of polar-type glaciers m eltw ater may
be virtually absent as ice volumes grow, and may only
become significant in the retreating hemicycle. T h e size of
the glacier and the rate at which it finally dwindles are also
likely to be significant in determ ining the volume and
velocity of w ater m oving away from the ice.
By com parison with rivers generally, m eltw ater streams
are typified by great seasonal, and even diurnal, variations
in flow resulting from changes in the rate of glacial ablation.
T hey are also characterised by the abundant, poorly sorted
and incoherent nature of their sedim ent load. Fahnestock
117
(1963) has given a detailed account of the m orphology and
hydrology of one such stream in western N orth America.
M any of the large rivers of the world begin life as m elt
w ater streams and many m ore had such an origin in the
past. T h e Snowy River in New South W ales and the
D erw ent, Forth, and Mersey in Tasm ania are some A ustra
lian exam ples of streams which had a glacier source in the
Pleistocene.
I_ __I S to g n o n t ice
M oraine w ith ridges
Lakes a n d stream s
54 Terminus of the Biafo Glacier in the Karakorum Himalaya, with

associated deposits (redrawn from Hewitt, 1967)
Ice contact streams. M eltwater streams may begin by

flowing over the glacier surface or w ithin the body of the
ice, in either of which cases they have little effect on the
succeeding landscape. However, when they flow over bedrock
at the base of the ice and when they flow along the ice
m argin, again in contact with bedrock, their effect may be
im portant. Subglacial streams may drill potholes and cut
channels beneath the ice: in some tem perate-type valley
glaciers such erosion may be quantitatively significant and
it is thought to have occurred extensively in the retreat
phases of the large Pleistocene ice sheets of the northern
hem isphere (M annerfelt, 1945, for instance) . Judging by
the degree of preservation of such features as valley steps it
may not be of great im portance in hard rock areas, but
softer m aterials may be considerably incised, as in the case
of the tunnel valleys cut beneath the Pleistocene ice sheet
in D enm ark and north ern Germany.
If deposition occurs instead of erosion, then the aggraded
Proglacial Processes 119
bed, bounded by the walls of the ice tunnel, becomes a
sinuous ridge of alluvium when the glacier melts, so
producing one type of esker (p. 181).
Ice m argin m eltw ater streams, ru n n in g along the side of
a valley glacier or the edge of an ice sheet, have a sim ilar
dual potential. T h ey may cut rock terraces, notches, or
gorges, sometimes abandoned on the disappearance of the
glacier b u t sometimes incorporated into the postglacial
drainage system. If, on the other hand, the streams aggrade,
they b u ild terraces which may survive the period of glacia
tion and are term ed käme terraces (Fig. 55) . T h e term
käme or käme hum m ock itself is used for any isolated
m ound of d rift m aterial deposited by w ater in contact with
the ice.
55 Formation of käme terraces. A: lateral terraces are built by marginal

meltwater streams; B: disappearance of the glacier causes collapse of the
glacifluvial deposits (redrawn from Flint, 1957)
Ice contact deposits differ from outwash sedim ents in

their generally greater range of particle size, often showing
ab ru p t changes betw een coarser and finer m aterials. T hey
also m ore com m only incorporate masses of till and, perhaps
most characteristically of all, their structure is apt to show
evidence of deform ation by the collapse resulting from
rem oval of the ice contact. T h u s collapse structures are
likely to be found along the ice contact face of käm e terraces
as well as w ithin individual käme hummocks.
Outwash streams. Outwash streams are proglacial streams
in the strict sense and may flow for hundreds of miles away
from the ice front. Many of the river systems of Europe
and N orth A m erica and their associated valley landform s
have been fundam entally affected by their function as m elt
water outlets of the great Pleistocene ice sheets. On a
smaller scale in n o rth ern Tasm ania, pebbly fills extending
to the m ouths of the Mersey and Forth rivers appear to be
associated w ith a tim e when the channels of these rivers
carried outwash from a small inland ice cap.
It is in the im m ediate vicinity of the glacier, however,
that the m orphological effect of outwash streams is of m ajor
significance, and because of the very heavy load carried by
these streams this effect is essentially a depositional one.
T h e word outwash itself is norm ally used for the deposits
so laid down.
Streams issuing from a glacier may carry m orainic
m aterial from well back along the course of the ice and in
this way they reduce the potential am ount of m aterial
which the glacier itself may deposit. Furtherm ore they erode
m aterial which the glacier has deposited, particularly in the
form of end moraines. So where, as in tem perate-type
glaciers, m eltw ater is abundant, this is a potent factor in
hib itin g the form ation and preservation of end moraines.
Outwash sedim ents differ from till in being stratified and
relatively well sorted. T hey still retain a wide range of size
grades — from boulders to sand — but, unless trapped by
obstructions, silts and clays are norm ally missing, having
been exported further downstream or deflated by w ind and
in many cases losing their identity by m ingling w ith other
sediments. Studies that have been made of outwash particles
suggest that they become increasingly rounded relatively
quickly and lose any facets and striations they may have
possessed, b u t lithology is clearly an im portant factor in this
regard.
Essentially outwash is laid down in a series of alluvial
fans. These may be simple and related to one particular
point of discharge from the glacier, b u t individual outwash
fans or cones rarely occur in nature because as the glacier
retreats new er fans are added to older ones in a complex
way. In the case of an ice sheet, where there are m ultiple
points of discharge, the fans also coalesce sideways to form
an ontwash sheet or outwash apron: in the case of a valley
glacier w ith lim ited outlets and channelling valley walls,
the fans form an elongate body of sedim ent term ed a valley
train.
As m ight be expected, vertical sections through outwash
m aterial tend to show a stratification of thin foreset beds
with particle size decreasing locally downstream . T h e sur
face form of the deposits comm only displays convexities
related to their fan-like natu re and kettle holes may occur,
bu t the m ost obvious surface features are a result of the
braided stream systems that develop. Braiding is encouraged
by the abundance of load m aterial, its poorly sorted nature,
and its incoherence, resulting largely from lack of fines
(Krigström , 1962) .
O utw ash may be laid down du rin g glacial advance and
the am ount of this can be considerable in the case of some
tem perate-type glaciers with high rates of ablation. In such
a case it will be destroyed or overlain by till. It seems clear,
however, that, even where advance outwash is present, the
m ore im p o rtan t am ount of outwash accum ulates d uring
recession and it is recessional outwash which is of geom orphic
rath er than stratigraphic im portance.
Both the total bulk of outwash and its bulk relative to
that of till vary principally with lithology, the glacial regime,
and tim e. Lithology is m ainly im portant in influencing the
proportion of till that will be evident as outwash. T hus,
on siliceous rocks in western Tasm ania, where tills have a
large content of pebbles and coarse sand, a relatively big
proportion of this m aterial is retained in the outwash
deposits. O n the other hand, on basic rocks in central
T asm ania where the tills are typically boulder clays, only
the boulder com ponent is retained in the outwash and the
clay is rem oved downstream . Glacifluvial deposits are
correspondingly smaller. T h e glacial regim e will tend to
affect both the absolute and relative am ount of outwash,
because glaciers w ith high accum ulation and ablation rates
will tend to produce more till and also m ore outwash since
Terrace heights in feet
10-16
1 9 -2 7
56 Meander terraces of the upper Mersey River, Tasmania, formed by

postglacial downcutting into till and outwash (redrawn from Spry, 1958)
both their rates of erosion and the am ount of m eltw ater

they produce are likely to be higher.
Several processes tend to give rise to terracing in the o u t
wash. A lternations of cutting and filling occur due to
changes in discharge to load relationships and these tend to
produce paired terraces of varying extent and am plitude.
W hen the ice finally disappears and a purely fluvial drainage
system takes over, postglacial streams are likely to cut non-
paired m eander terraces as they incise into the outwash. A
good example is provided by the terraces of the m iddle
Mersey in Tasm ania described by Spry (1958) and illustrated
in Fig. 56.
In reconstructing the chronology of form er glaciations,
a study of outwash deposits may reveal as m uch if not m ore
than a study of tills and m uch of the evidence for m ultiple
glaciation in New Zealand derives from the identification
and exam ination of outwash deposits of different ages.
Drainage derangement
In addition to producing m eltw ater with its attendant effects,
glaciers derange other drainage systems by im pinging upon
the meltwater outlets of other glaciers and upon rivers flow
ing from neighbouring unglaciated uplands. Such derange
ment is likely to be particularly drastic in the case of a large
ice sheet expanding towards higher ground, but important
local effects may be produced even by valley glacier tongues
and especially by those related to large gathering grounds
and able to push exceptionally great distances outside the
snowline. Derangement may be caused by the ice body
itself but also by the till and outwash deposits associated
with it.
Basically there are two aspects of such derangement,
although these are closely interconnected. Diversion causes
glacier-margin streams to cut channels that may or may not
become permanent: damming causes proglacial lakes, the
overflow from which also cuts new channels or modifies old
ones.
Because the character of the derangement depends essen
tially on the relationship between the ice front and the
pre-existing topography a great range of occurrences is
possible. In Pleistocene Europe the southern edge of the
Scandinavian ice sheet intercepted rivers flowing northward
from the Alpine lands. Relief along the glacier margins
was so slight and bedrock so erodible that escape of river
water westward was possible and a succession of glacier-
margin streams cut a series of east-west channels as the
position of the ice edge changed. These are the ‘Urstrom
täler’ of northern Germany and Poland. In Pleistocene
eastern North America, the Laurentide ice sheet, retreating
northward from a drainage divide, also intercepted north
ward flowing streams, but because of greater relief and less
erodible rocks these were not able to escape by lateral
diversion and a great system of proglacial lakes was pro
duced, eventually to become the Great Lakes of today. In
the case of these North American ice-dammed lakes, the
divide to the south was not a very high one and overflow
streams were able to breach it in several places.
Valley glacier systems produce effects on a local rather
than regional scale and a common location for ice-dammed
lakes is in valleys tributary to that carrying the glacier itself.
Lakes may occupy such tributary valleys and be contained
simply by the edge of the valley glacier. In some cases, as
in that of the L inda Valley of western T asm ania, described
by Ahm ad, Bartlett, and G reen (1959), a lake may be
created by a diffluent branch of the valley glacier advancing
into the trib u tary valley.
Effects of channel cutting. D rainage channels or spillways
cut as a result of glacial interference may be abandoned on
deglacierisation and present the appearance of anom alous
dry valleys and cols, sometimes floored with alluvial m aterials
and sometim es not. In many cases, however, the channels
cut in glacial times become incorporated into the postglacial
drainage systems so that the latter often show resulting dis
cordances w ith relief and structure and a high incidence of
m arkedly ‘underfit’ streams. In some instances whole
drainage divides have been moved great distances in this
way.
Effects of lake formation. A lthough some proglacial lakes
may survive the period of glaciation in m odified form, most
of them disappear. W here the outlet was across bedrock
and the w ater level rem ained stationary for significant
lengths of tim e the position of the form er lakes is m arked
in the landscape by old shorelines, either erosional or
depositional. Lakes from which the overflow channel was
across ice or in contact with ice are not likely to have left
m uch evidence of form er shorelines. T h e lake floors may
give rise to exceptionally well-developed flats or lacustrine
plains, often poorly drained.
T h e sedim ents accum ulating in proglacial lakes are pre
dom inantly silts and clays and these are often rhythm ically
lam inated so that bands of coarser m aterial alternate with
bands of finer particles. Such bands, which may vary from
a few m illim etres to a few centim etres in thickness, are
term ed varves, and each pair is thought to represent the
deposits of one year. O n m elting in spring the coarser silts
are laid down, b u t finer silts and clays are kept in suspension
through the sum m er and do not settle un til the accession
of m eltw ater ceases in autum n and winter. As the cycle is
repeated a particularly sharp break in sedim entation occurs
betw een the finer m aterial deposited in w inter and the first
coarser deposits of the succeeding spring. Varves have been
widely used for chronological reconstruction and the classic
work is that of De G eer in Sweden. From a geom orpholo-
gical point of view, the chief significance of varved clays is
as evidence of form er conditions giving rise to a proglacial
lake. W ith the disappearance of the ice wall form ing part
of the lake boundary and w ith the erosion of m uch of the
sedim ent of the lake floor, rem nants of varved clays may be
the only clues to the lake’s previous existence.
C o n t o u r s in feet
57 Subglacial drainage channels cut into till about 70 miles northwest

of Schefferville, Quebec. Large subglacial channels run downslope and are
joined by smaller submarginal channels running more or less parallel to
the contours and, presumably, to the former edge of the vanished Pleistocene
ice sheet (redrawn from Ives, I960).
In some lakes deltaic deposition may take place. T h e

deltas tend to be jetty-like or lobate and to residt from
accum ulation opposite particular distributaries, b u t in some
circumstances the sedim ents build along a broad front
between the ice edge and lake border producing an ice-
m argin terrace superficially resem bling a käme terrace
(Flint, 1929).
Im portant geom orphic results may follow from the
sudden evacuation of a proglacial lake. Such outbursts
occur in m odern instances and may be assumed to have been
comm on in the Pleistocene. T h e rapid release of large
quantities of water in this way may account for some features
difficult to explain by continuous processes.
58 Proglacial drainage channels in relation to the edge of an ice sheet
Classification of glacial drainage channels. Derbyshire

(1962) has presented the following classification of glacial
drainage channels, modified from M annerfelt (1945) and,
with the exception of subglacial forms, this is illustrated
diagram m atically in Fig. 58.
Proglacial
1. Channels which fall consequent upon local slope
(a) Overflow channels
(b) Col gullies
(c) T erm in al channels
2. Channels which fall inconsequently
(a) M arginal channels and benches
(b) E xtram arginal channels
Subglacial
1. C hannels which fall consequent upon local slope
(a) Subglacial chutes
(b) Subglacial col gullies
(c) Subglacial channel systems
2. C hannels which fall inconsequently

(a) Subm arginal channels and benches
Subglacial chutes occur alone and fall precipitously down-

slope, in contrast to col gullies which cut transversely
through ridges and to channel systems which incorporate
well-developed accordant tributaries. Subm arginal channels
are form ed below and roughly parallel to the ice edge: they
are usually m uch deeper and steeper than strictly m arginal
channels, although the two types may be continuous, so
form ing one type of composite channel.
Aeolian processes
W ind becomes an im portant agent of erosion in the pro-
glacial zone by its action on the finer fraction of the outwash
sediments. A lthough clays and a high proportion of silts
are carried further by outwash streams and do not form a
significant part of these sediments, sands and some of the
silts are trapped between coarser particles and so form
veneers on spool bars or other exposed sections of a braided
stream bed. W hen water levels fall these may be exposed to
w ind action, particularly as the w inter half of the year is
likely to be the period when water levels are lowest and
w ind velocities highest.
T h e sedim ents and depositional landform s resulting from
the relocation of glacial sands and silts are very sim ilar to
those derived from similarly sized particles of periglacial
origin described in C hapter II. As indicated there, peri
glacial sands and loesses are derived in the m ain from frost
splitting and frost churning of bare ground and from the
beds of braided streams formed under periglacial conditions.
Glacial sands and loesses are understood to be derived in
the m ain from glacial outwash deposits. T h e term ‘peri-
glacial loess’ is comm only used to describe sedim ents
originating in glacial outwash b u t this is not in accord w ith
the sense in which ‘periglacial’ is used in this book.
Sand movement. Sands are moved predom inantly by

saltation and are therefore relocated only relatively short
distances from the source area. T hey may often rem ain on
a section of the outwash feature in which they originated.
O n the other hand very extentive sand sheets of glacial
origin are known from N orth Am erica and Europe. T h e
farthest travelled sand tends to form thin sheets with very
little topographic expression, b u t well m arked fossil dune
areas may occur closer to the source. A lthough other forms
may be found, the characteristic dune is the parabolic or
U-shaped dune, which is concave upw ind and results from
the interaction of w ind and vegetation. In rather m ore arid
regions such as the G reat Plains of N orth America, desert
type barchans and longitudinal dunes appear to have
form ed du rin g the Pleistocene, b u t in m oister regions the
plant-stabilised parabolic dune appears the rule.
Glacial loess. T h e term ‘loess’ is used here in its broadest

sense and as it was defined by F lint (1957) — ‘a sedim ent,
comm only nonstratified and comm only unconsolidated,
composed dom inantly of silt-sized particles, ordinarily with
accessory clay and sand, and deposited prim arily by the
w ind’. T h e predom inant particle size and the sheet-like
m anner in which loess is deposited imply that it is trans
ported essentially in suspension. It is generally recognised,
however, that some loess has been retransported and re
deposited by mass m ovem ent and ru n n in g water so that in
its im m ediate provenance it may be colluvial or alluvial.
T h e extensive loess sheets of the South Island of New
Zealand may be taken as exem plifying such sediments,
although the compact, relatively im perm eable and non-
calcareous character of New Zealand loess would exclude it
from some m ore restrictive definitions of the term (Raeside,
1964) . In the South Island, the loess forms blankets varying
between about 80 cm and 8 m in thickness, covering river
25 Ablation moraine covering the snout of the Tasman Glacier, New
Zealand. Note lateral moraine, end moraine ridges, marginal and terminal
meltxvater streams, and valley train. The course of a major subglacial
drainage channel is marked by four prominent melt clolines (N.Z. Geological
Survey photo by N. S. Beatus)
26 Hummocky surface of käme terrace near Ben Dhu, South Island,

New Zealand (J. G. Speight)
27 Valley train and lakes enclosed by end moraines at the termini of
the Classen Glacier (left) and the Godley Glacier, New Zealand (N.Z.
Geological Survey photo by N. S. Beatus)
2,S' Subglacial drainage channels inherited from Pleistocene ice sheet,

central Quebec (]. D. Ives)
terraces and undulating low hill country and plastered
against slopes of greater am plitude. It is thickest towards
river floodplains and towards the present coast and this
suggests sources on the braided beds of proglacial streams
and the outer surfaces of the enorm ous outwash fans of the
C anterbury Plains currently below sea level b u t exposed
59 Loess deposits of New Zealand. Note their relationship to the area

glaciated in the Pleistocene and the distribution of probable glacial outwash
now submerged on the continental shelf (based mainly on Raeside, 1964)
I
during the low water phases of Pleistocene glacials (Fig. 59) .
At least six layers of loess have been identified from New
Zealand, these being thought to have originated during three
stadia of each of the last two glacial ages (Raeside, 1964) .
The sheet-like form of loess deposits prevents their giving
rise to distinctive depositional landforms: their effect is
rather to even out pre-existing relief forms. However, the
cohesive nature of loess results in the ready production of
cliffs and steep-walled gullies when it is subsequently
eroded. Some loess surfaces show a dimpling which may be
related to unevenness of deposition or subsidence due to
underground drainage or solution.
IX
GLACIATED MOUNTAIN
INTERFLUVES
T h e effect of glaciation upon the form of interfluve areas

in highland country depends in the first place on the
character of the interfluves themselves. If these are broad
and flattish they will encourage the developm ent of plateau
glaciers and even m inor ice caps, so that the resulting land-
forms will resem ble those of the glacially eroded plains
described in C hapter XI. If, on the other hand, they are
narrow er and m ore dissected they will tend to prom ote the
evolution of cirque glaciers. It is w ith the second of these
two cases that this chapter is concerned.
Cirques
Cirques are usually described as arm chair-like hollows,
com prising a floor which is norm ally concave and a headwall
which is steep and semi-circular or horseshoe shaped in plan.
O ften the floor shows a reverse slope at the cirque exit —
so form ing a threshold — and the headwall usually meets
the floor in a sharp break of slope. In some cirques the
headwall and floor m eet in a m ore continuous curve, b u t a
lesser break of slope is found separating the higher steeper
section of the headwall from a lower less steep section: this
is the schrund line. T h ere is in fact great variation of cirque
form, n o t only because of variations in the form of the
glacier which occupied them b u t also because of variations
in the m orphology, attitude, and lithology of the bedrock
in which they are carved. Some representative long profiles
of cirques are shown in Fig. 60.
131
Head wall Schrund line

Threshold
60 Some representative cirque profiles
Origin. It is difficult to envisage glacial cirques develop

ing from other than the snowpatch hollows or nivation
cirques described in C hapter V, since the cirque glacier m ust
have been preceded by first a seasonal and then a perennial
snowbank. If this is so, then the first fashioning of the
glacial cirque m ust have proceeded by frost shattering and
Du Cane
1 Mile
Contours in feet
1 Kilometre
61 The Long Lake, Narcissus, and Cephissus valleys in the Du Cane

Range, central Tasmania. These head in valley-head cirques which grade
directly into the upper ends of glacial troughs without any intervening
trough headwall. The Du Cane Gap is a transftuence col (p. 145).
Glaciated M ountain Interfluves 133
transport of shattered particles by snow-melt or snow-slide.
G radually as the snow w ent through its various changes to
glacial ice, cirque glacier processes took over.
Broadly speaking cirques can be grouped into two m ain
types according to their place of origin. T h e first group is
th at of valley-head cirques. These are form ed at the heads
of stream systems so that they have an im m ediate relation
ship to the pre-existing drainage pattern. T h e second group
— sometimes term ed hanging cirques — is independent of
the previous stream system and is found cut into m ountain
slopes in a variety of situations. Valley-head cirques tend to
be bigger than hanging cirques and some of them are very
big indeed (Fig. 61) . T hey also tend to be m ore rounded
in plan, because the location of hanging cirques is com
m only controlled by rock structure and this often induces
a ledge-like disposition.
O rientation. Factors affecting the orientation of cirques

were ou tlin ed in C hapter I. In the westerly w ind belt of
the southern hemisphere, cirques are bigger and more
num erous on the southeasterly side of highland massifs,
because there is m ore accum ulation of snow on the eastern
lee sides and less ablation on the shaded southern sides
(Fig. 62).
E volution. T h e evolution of cirques can be considered

u n d er two heads — the developm ent and retreat of the
headwall and the developm ent and low ering of the floor.
It is generally considered that cirque enlargem ent takes
place predom inantly by recession of the headwall so that
it has the effect of cutting m ore or less horizontally into
the m ountain mass. T h is is inferred from the general
topographic relationships of existing cirques and from
observations of the relative efficiency of downw earing and
backwearing processes. T h e re is norm ally a strong contrast
betw een the roughened appearance of the surface of the
headwall and the smoothed surface of the floor, reflecting
the contrast between the processes concerned in developing
the two surfaces, b u t this is not always easy to see. T h e
lower part of the headwall is often covered by postglacial
134
425
1 M ile
1 K ilo m e tre C o nto urs in fe e t

talus and the floor is com m only obscured by d rift deposits,
or, in the case of some cirques, covered by a lake.
T h e cirque floor is produced by retreat of the headwall,
b u t its form is due to corrasion by the cirque glacier. T h e
relatively smooth surface of the floor and the frequent exist
ence of reverse slopes suggest that abrasion is the m ore
im p o rtan t operative process, b u t it is possible that plucking
is significant in the early stages of cirque deepening, at least
u n til the base of preglacial w eathering is reached. In
ad dition W . V. Lewis (1954) has discussed the possibility
that some am ount of dilation jo in tin g may occur as down-
w earing progressively unloads the sounder rock beneath. If
this occurs then this m aterial too m ust be rem oved by a
plucking process.
T h e overdeepening that is characteristic of m any cirques,
so that a reverse slope occurs headw ard from the threshold,
is not properly understood, b u t appears to result from a
n u m b er of circumstances. Velocity m easurem ents suggest
that the rate of sliding over the bed decreases towards the
term inus of the glacier and this is in accord w ith general
expectation from discussion in C hapter V. A lthough changes
in velocity do not appear to have any significant effect on
the forces which the basal ice can exert, they will clearly
influence the rate of rem oval of particles freed by both
abrasion and plucking and also influence the rate of supply
of fresh rock tools. T h ere is evidence from many cirque
glaciers of the accum ulation of debris under the snout so
that the term inal ice is overriding the upstream side of an
end m oraine (McCall, 1960). T h is too m ust give rise to a
decrease in the rate of bed low ering towards the term inus.
T h e most im portant factor, however, may be that of the flow
lines in the glacier (Fig. 45, p. 90), for these imply that
in sum the glacier is undergoing a rotational m ovem ent, and
it is easy to imagine that concavity in the floor has arisen
in sympathy w ith this.
62 The Denison Range in west central Tasmania. This shows the

asymmetry of cirque development typical of snow accumulation with a
strong westerly air stream. The cirques are all on the eastern lee side of
the mountain range, whereas the western slopes show evidence of smoothing
by periglacial processes (opposite).
O verdeepening in some cirques may be encouraged by
their location in outcrops of weaker rock, and Peterson
(1966), for instance, thought that several cirques in the
Frenchm ans Cap massif in Tasm ania were located where
dolom ite outcrops in juxtaposition with m etam orphosed
quartzose rocks.
Headwall retreat. T h e retreat of cirque headwalls appears

to be associated with processes taking place at the headwall
gap, which is a norm al appurtenance of cirque glaciers and
exists in discontinuous fashion between the upper glacier
and the headwall. It is usually no m ore than a m etre or
two in width and may or may not be connected to a surface
gap, or randkluft. Headwall gaps may also obtain continuous
connection with the surface through an ice crevasse close
to the rear of the glacier term ed a bergsehrund (Fig. 63) .
Randkluft
Bergsehrund
Headwall
gap
Headwall
gap
63 Diagrammatic relationships of the headwall gap to a randkluft and

bergsehrund. Headivall gaps are not necessarily connected with the surface
in such icays.
According to McCall (1960) the headwall gap occurs where

the headwall has a greater slope than the direction of ice
m ovem ent, while its lower lim it is determ ined by the depth
at which the weight of overlying firn and ice is too great
for the gap to be supported.
T h e bergsehrund hypothesis of headwall sapping was
form ulated by W. D. Johnson (1899, 1904) and envisaged
alternate freezing and thaw ing of water on the headwall at
the base of a bergschrund, the frost shattered m aterial being
incorporated into the glacier and then rem oved by it.
Objections to the hypothesis on the grounds that all cirque
glaciers do not have bergschrunds now appear unfounded
since it is likely that all such glaciers have headwall gaps
even if these are not always open to the surface. More
serious objections concern the m easured am plitude of
tem perature variation in bergschrunds and the physical
impossibility, because of ice plasticity, of bergschrunds being
deep enough to cover the entire height of some headwalls.
Battle (1960) reviewed the results of continuous tem perature
recording in bergschrunds of different depths and concluded
that, particularly in deep bergschrunds, the am plitude of
tem perature change is very low — generally between 0 and
1 5 or 2° C — and the rate of change is also very low.
It seems probable then that, instead of invoking alter
nate freeze and thaw as in the original Johnson hypothesis,
it is necessary to envisage the freezing of liq u id water
coming into the headwall gap from outside. Such a sugges
tion was m ade independently by W. V. Lewis (1938) and
Nussbaum (1938), both of whom observed that m eltw ater
and rainw ater ru n n in g down the headwall at certain times
m ight freeze in rock crevices and become a potent w eather
ing agent. A nother potential source of liquid water is
groundw ater seeping from the headwall, and it has been
thought (C ham berlin and C ham berlin, 1911) that the level
at which such water issues may be im portant in influencing
the position of the schrund line. T h e existence of a berg
schrund or any kind of headwall gap would not appear
necessary where the freezing of groundw ater is envisaged:
nor would the downward percolation of rain and m eltw ater
require m ore than the narrow est of gaps between ice and
rock. W hile there seems little doubt that frost w eathering
of rock is im portant in headwall retreat and that some sort
of headwall gap is likely to be of assistance in this, it now
seems very unlikely that a bergschrund is as essential as was
envisaged in W. D. Johnson’s original hypothesis.
In some of the earlier discussion on mechanisms of
cirque headwall retreat the issue seems to have been
obscured to some extent by an assum ption, either im plicit
or explicit, that freezing and thaw ing are necessary for frost
heave and frost shattering. As a result attention was focused
on ways in which freeze-thaw cycles could occur. In fact,
as discussed earlier and as Penner (1961) has emphasised,
the presence of a thaw ing phase is not necessary as long as
there is a supply of liquid water to begin with. It is the
process of freezing with ice crystal growth and water
segregation that provides the operative force.
In order that the relative steepness of headwalls should
be m aintained it has been widely thought that some form
of sapping or u n derm ining m ust take place. C otton (1947)
observed that m any w riters seem to have proceeded ‘on the
tacit assum ption that cirques were sapped to their present
dimensions (with walls in some cases 2000 or 3000 feet high)
by glaciers or nevees of insignificant thickness resting on the
cirque floors below these great cliffs, as glacier rem nants still
do in some cases’. T h e assum ption is illustrated in the
diagrams of D. W. Johnson (1941) . However, it seems highly
unlikely that the form of cirque headwalls is determ ined by
processes operating when the glacier itself is at its least
potent. Such hypotheses seem to have been put forward
because of the assumed necessity to bring conditions for
alternate freeze-thaw down to the zone of sapping near the
base of the headwall. If, as previously indicated, freezing
alone is required, then the problem is reduced to that of
suggesting ways in which liquid water may accum ulate at
the base of the headwall and in which ice pressures can be
reduced sufficiently for ready freezing to take place at depth.
T h e latter req u irem en t is readily m et along the headwall
gap and below this to some extent by the fact that the ice
is m oving away from the rock wall. As McCall (1960)
suggested, an accum ulation of free water at the base of the
headwall, from whatever source, is not difficult to envisage
and this may well be the fundam ental reason for sapping.
It should be kept in m ind, however, th at the necessity
to invoke und ercu ttin g ought not to rest solely on com
parisons between the headwalls of fresh glacier cirques and
those of empty cirques which become degraded by subaerial
weathering, mass m ovem ent, and gullying. Downwasting of
deglacierised headwalls is brought about largely by talus
and fan accum ulations at their base, since this builds out
the lower slope while the up p er slope continues to retreat.
T h e re is no doubt that the cirque glacier encourages
parallel retreat of headwalls by rem oving the debris falling
from them and thereby allow ing greater equality of attack
by w eathering over their surface. As in the parallel case of
m any sea cliffs, this is an im p o rtan t factor in m aintaining
scarp steepness. T h e basal ice in the cirque glacier studied
by M cCall (1960) appeared to have originated from water
which had penetrated the headwall gap, frozen at the bottom
and become incorporated into the glacier at the same time
as the rock debris with which it was heavily laden.
Observations on the walls of N orwegian cirques made
by Battey in 1960 suggested that headw ard erosion is largely
controlled by the appearance of dilation joints due to
spontaneous expansion of the rock following release from
compression. In his view the walls ‘retreat roughly parallel
to themselves by the spalling off of sheets from the free face’.
It seems likely that basal sapping, dilation jointing, and
rem oval of basal debris are all processes co n tributing to the
parallel retreat of cirque headwalls. It is im portant to
rem em ber though that the history of a particular cirque
may extend through m ore than one period of glaciation
and therefore through at least one interglacial period. If
this is so, its form may owe som ething to a very wide range
of processes, not all of which are strictly glacial.
O verridden cirques. M any cirques initiated in the early

stages of an ice cap glaciation subsequently become swamped
with ice and over-ridden, usually from the rear. Independent
cirque glaciers may be re-established in the w aning phases
of the glaciation. W here overriding has occurred it may be
expected that an appreciable am ount of erosion has taken
place by glacial plucking.
Levels of cirque cutting. T h e idea that headwall reces

sion by frost shattering is an im portant factor in cirque
developm ent led some earlier w riters to a concept of
preferred levels of cirque cutting which are clim atically
controlled. O ne of the foremost exponents of this view
was T aylor (1926) who thought it possible to reconstruct
m ultiple Pleistocene snowlines from the evidence of in d i
vidual cirque floor levels. Few workers w ould now use the
evidence of cirque floor levels in such close fashion and,
as discussed above, cirque developm ent is not necessarily
correlated with a great num ber of freeze-thaw cycles, even
though cirque in itiation through nivation may be. T h e re
m ust clearly be lim its to the zone in which cirque develop
m ent can proceed. O ne of these is set by the snowline,
outside which there can be no cirque glacier: the other, less
easily defined, may be set by tem peratures too low for liquid
water to be present in significant quantities, for, where
polar-type glaciers are produced, cirque cutting seems likely
to be inhibited: b u t between these broad lim its, develop
m ent may go on at a variety of levels.
Two-storeyed cirques. O ne circumstance which led A. N.

Lewis (1922) to proceed with the use of m u ltip le cirque
levels in Tasm ania as a guide to m ultiple snow line positions
was the frequent occurrence of ‘two-storeyed’ cirques where
a higher cirque lies in tandem above a lower. T hese occur
in glaciated m ountains in many parts of the w orld w ith a
frequency to suggest that they are an expectable concom itant
of cirque production. T h e processes giving rise to them
are probably several. In some cases they may be due to a
change in the position of the snowline so that the two
cirques are not contem porary but of different generations,
as A. N. Lewis postulated. W here they are contem porary
C otton (1947) has suggested that ‘reconstructed glaciers’
may have been form ed in lower niches from avalanche ice
descending from higher cirques. In T asm anian examples
the existence of well-marked pseudo-bedding planes in
dolerite often seems to provide m ultiple horizons along
which cirques may develop, and it is possible that structure
is a m ajor control in the case of many two-storeyed cirques.
In m any other T asm anian examples listed by Lewis the
lower cirque is probably not a true cirque b u t represents
an exaggerated valley step or trough headwall (see p. 157).
Cirque deposits. In some cirques where the glacier snout
rem ained for a considerable tim e near the cirque lip, well-
m arked end m oraines were deposited. In other cases where
the cirque glacier fed descending valley glaciers and the
retrea t phase was a rapid one, very little in the way of an
end m oraine is to be seen. W here present, the end m oraine
may contain n o t only m aterial deposited by the glacier and
dum ped and lodged at its term inus b u t also talus m aterial
w eathered from the headwall above the glacier which has
slid over the glacier surface. In this way some end m oraines
may also be in part protalus ram parts.
Glacial deposition on the floor is norm ally small and
often dom inated by hum m ocky ablation m oraine produced
by the dying glacier. However, on cirque floors with reverse
slopes deposition is com plicated by the developm ent of
a lake, so that silt and peat accum ulate sometimes to
considerable depth.
Cirque lakes. A bandoned cirques comm only contain

lakes, b u t in some no lake has ever form ed while in others
the lake has had a relatively short history and been o b liter
ated by infilling and erosion at the lip. W here lakes do
occur some are dam m ed behind an end m oraine, some are
contained w ithin a rock basin, while others are the result
of both erosion and deposition by the glacier. D rift con
tained lakes are usually shallow, b u t lakes in overdeepened
rock floors on to which little debris has fallen from above
may be m any tens of m etres in depth.
Upland dissection by cirques

T h e fullest significance of cirque developm ent lies in the
way in which it influences landform evolution as a whole.
T w o basic topographic relationships of cirques to uplands
have long been recognised and were fully described and
illustrated by H obbs (1911). T h e first is the scalloped
upland in which cirques appear to be cutting headward into
a flattish or smoothly rounded upper surface (Fig. 64) : the
second is the fretted upland in which the divides betw een
A/Mt Twynam
Mt Townsend
Mt (
Kosciusko A
2 Miles
2 Kilometres Contours in feet

/A / f\
64 Location of cirques in the Snowy Mountains of New South Wales
(based on Galloway, 1963). This is basically a scalloped upland with marked
east-west asymmetry.
cirques are narrow and steep (Fig. 65) . T h e two cases

have comm only been associated in sequential developm ent
w ith the assum ption that the second is a derivative of the
first. H obbs described such a sequence and Davis (1912)
fitted it to his concept of a cycle of erosion so that the
scalloped upland represented a ‘youthful stage’ and the
fretted upland a ‘m ature stage’. T h e re seem to be no
theoretical difficulties in postulating such a sequence and
Glaciated Mountain Interfluves 143
Contours in feet
'/2 Kilometre
65 Cirques in the Arthur Range of southwestern Tasmania. They have

produced a fretted upland in ruhich the cirques are separated by glacial
horns and aretes. Inosculation cols are developing notably to the west of
Lake Dione (B) and between Lakes Saturn (A) and Hyperion (C). Lakes
Dione and Saturn appear to occupy a two-storeyed cirque. Immediately
southwest of Lake Ganymede (D) is a diffluence col through which ice from
the Ganymede Glacier appears once to have spilled. Compare Fig. 66.
its actual occurrence seems dem onstrable from the existence

of both types w ithin the same highland mass. However,
there are undoubtedly instances where the fretted upland
has not developed from a pre-existing scalloped upland b u t
has resulted from a sharpening of a preglacial landscape
with already narrow interfluves. In Tasm ania, for example,
it is notable that typical scalloped and fretted topography
is to be found am ong the tabular dolerite-capped highlands
whereas the quartz m etam orphic highlands with their more
com plex and more steeply dipping structures and greater
degree of pre-dissection show only what m ust be term ed
fretting. T h ere seems little doubt that in the first case a
sequence of glacial forms is represented but that in the
second no true scalloped stage preceded the present
condition.
Sequence of cirque cutting. T h e sequence of dissection

by cirques as described by H obbs (1911) envisaged an early
stage in which the preglacial surface was dom inant and
cirques began their invasion by scalloping from the edges.
W here the cirques lay well w ithin the upland mass the
flanks of the highland m ight be dissected by glaciated valleys
heading in the cirques, and this produces a ‘grooved’ or
‘channelled’ upland in the term inology of Hobbs (but see
p. 170). H e recognised that, because headwall recession in
cirques takes place along a m ore or less sem icircular front,
the cirques tend to enlarge laterally as well as to the rear.
Such a circum stance m ight destroy the upper part of the
preglacial surface at a greater rate than the part on the
flanks between the down-cirque valleys. Flat-topped spurs
w ould then result and these w ould eventually be reduced
to pyram idal ‘m onum ents’ which C otton (1947) suggested
should be grouped under the m ore general heading of
tinds. A tin d may be defined as a residual peak isolated by
glacial erosion on the edge of a highland mass: it also
comm only originates where an ice cap flowing over the edge
of a plateau divides around and progressively isolates a
nunatak. M t Ida on the east shore of Lake St C lair in
central T asm ania may provide an exam ple produced by
both mechanisms, since its early separation may have been
influenced by cirque developm ent on either side, whereas
it has m ore recently been sculptured by ice com ing off the
plateau behind and dividing around it (Fig. 74, p. 164).
C ontinued enlargem ent of the cirques is envisaged as
eventually destroying the pre-existing upland surface so that
a characteristic fretted upland or ‘a lp in e’ landscape results
in which the divides betw een cirques are formed by razor-
29 Fretted upland of Federation Peak block, Tasmania, with two-storeyed
cirque (Tasmanian Department of Film Production)
30 Scalloped upland of the Snowy Range, Tasmania, with moraine-

dammed cirque lake
31 Asymmetrical ridge produced by intersection of cirque headwall and
upland surface, Frenchmans Cap, Tasmania (J. A. Peterson)
32 Small rock basin cirque with well-developed abraded threshold on

eastern slope of Mt Murchison, Tasmania (J. A. Peterson)
edged ridges or aretes, term ed ‘comb ridges’ by Hobbs. T h e
form of aretes varies w ith the circumstances of cirque in ter
section and also w ith lithology and rock structure, b u t they
are characteristically pinnacled and gullied so as to have a
serrate appearance. W here m ore than two cirques intersect,
a m ore or less pyram idal peak or horn is norm ally produced
at the ju n ctio n of aretes, reflecting the fact that lowering
of the skyline proceeds first at the points where cirques first
intersect or inosculate and last at the points where the aretes
m eet. It w ould seem then th at the degree of prom inence
of h o rn peaks in an alpine landscape depends essentially
upon the ex ten t of lowering of the adjoining aretes.
In some cases, headwall intersection may occur laterally
before it does in the back-to-back direction, thus form ing
a cirque terrace. T h e T a rn Shelf at M t Field in Tasm ania
may be one such feature b u t they do not appear to be
common, perhaps because they dem and a high degree of
coincidence in cirque floor height in order to reach their
ideal developm ent.
Glacial cols
An im p o rtan t feature of interfluves in glaciated uplands is
the presence of gaps or cols, some of which may be in part
or in the m ain of preglacial origin, b u t all of which owe
their eventual form to glacial action (Fig. 66) .
Inosculation cols. T h e sim plest form of col occurs when

an arete is lowered at the intersection of two cirque head-
walls. Because the essential cirque shape is a bowl-like one,
intersection comm only produces a com paratively regular
hyperbolic curve w hen the col is seen in elevation. Such
cols tend to retain the sharp divide in herited from the
parent cirques, b u t if m ovem ent of glacial ice takes place
through them they suffer rap id m odification. T hey may
then be term ed diffluence or transfluence cols.
Diffluence and transfluence cols. Diffluence is downstream

bifurcation of an ice stream so that the direction of glacier
flow is split: transfluence is a special case of diffluence in
which one of the divided ice streams invades another glacier.
W hen an inosculation col is formed, ice from a thicker or
higher cirque glacier may spill into the th in n er or lower
one. A transfluence col arises and its floor is progressively
sm oothed and ro u n d ed by ice erosion to produce w hat is
essentially a perched and very short glacial trough.
66 Three types of glacial col. Glacier A has a diffluent branch and

also a transfluent branch becoming tributary to glacier B. Inosculation
cols are produced by headward erosion of a cirque glacier C and those
feeding glacier A.
Diffluence cols occur along interfluves separating glacial

troughs where valley glaciers overtop the divides and flow
into another valley. T h e breaching of divides in these
ways by both cirque and valley glaciers may be of great
significance in subsequent landscape history, particularly if
they originate at or are lowered to a level at which the
postglacial drainage is affected. W idespread river capture
and divide m igration may result, some of the best examples
of which have been described from Scotland by L inton
( 1951) .
Glaciated Mountain Interfluves 147
The end product of mountain glaciation
Davis (1912) projected the sequence of cirque-induced land
scapes so that the ‘youthful’ scalloped upland and ‘m atu re’
fretted upland were succeeded by a ‘post-m ature’ stage in
which the horns and aretes were still further lowered to
produce a m uch m ore subdued landscape. Such a final stage
has never been dem onstrated satisfactorily in nature and
even as a theoretical abstraction it is probably not sound,
since a reduction in height of cirque walls would lead to a
decrease in the volum e of firn and consequently to a glacial
retreat. Most w riters on glacial landform s seem to have
assumed, even if only tacitly, that the fretted upland rep re
sents the end stage of any glacial sequence, and it may be
that it approaches a condition of landscape equilibrium .
In m any parts of the world, and notably in the European
Alps and the W estern C ordillera of N orth America, height
accordance of glaciated peaks in fretted uplands has been
noted over wide areas. A summ ary of European observations
and discussions is given by Charlesw orth (1957, pp. 324-8).
Such alpine sum m it plains (gipfelflur of the E uropean
Alps) have often been in terp reted as dissected peneplains,
their accordance reflecting the height of an earlier plateau
surface out of which they were cut. In a small way and where
plateau rem nants are still present this can probably be
dem onstrated, as in the Du Cane Range of T asm ania where
several glaciated peaks appear to reflect the general height
of the Central Plateau, of which they are dissected frag
ments. But it seems doubtful that gipfelflur extending for
many hundreds of miles can have arisen in this way, and
much more likely that the tendency to sum m it accordance
is an in b u ilt one which is in h eren t in the processes control
ling the form of the summits. R ichter (1906) supposed that
the altitude of the higher cirque floors was the operative
factor since this could be expected to control in large degree
the height of aretes and horns above, b u t Daly (1905)
followed Dawson (1896) in invoking interglacial clim atic
conditions and particularly the height of the treeline. Daly’s
hypothesis, reviewed and amplified by T hom pson (1962),
involves a consideration of the evolution of the upper slopes
of glacial troughs (see later, p. 154). Meanwhile it provides
a salutary reminder that glaciated landscapes, and especially
those which are at present deglaciated, have evolved not
only through Pleistocene glacial ages but through inter
glacials as well, so that their form is rarely explicable solely
in terms of glacial processes. The cirques, aretes, horns, cols
and other features of glaciated mountain interfluves where
ice is no longer active can be expected to have inherited
something from periglacial and nivational processes preced
ing and supplanting those of the glacial regime itself.
Ice sheet glaciation of mountains

When the volume of ice in a mountain glacier system
becomes so great that interfluves are swamped, an ice sheet
stage may take over and introduce a further phase of skyline
modification. Broadly speaking the major effect is to wear
down and smooth the aretes and horns of the fretted upland,
although the effect on the scalloped upland is probably not
very great. It is generally assumed that relatively little
lowering of relief takes place by ice sheet erosion of moun
tains, but Linton (1963) has presented evidence suggesting
that, in the case of long continued erosion by large ice
sheets, significant lowering or even elimination of interfluves
may take place.
X
G L A C I A T E D M O U N T A I N VALLEYS
T h a t section of a pre-existing valley occupied and modified

by a valley glacier is term ed a glacial trough, b u t the
relationships of such troughs are m any and varied. L inton
(1963) distinguished betw een those of Alpine type, which
occupy a preglacial drainage system and originate in cirques,
and those of Icelandic type, also taking advantage of fluvially
d eterm ined routes b u t descending b o m a plateau glacier
or larger ice cap into the valleys dissecting the highland
mass. H e noted that some troughs, which he term ed com
posite, did not conform to the preglacial drainage system
b u t developed their own routes by varying degrees and
com binations of diffluence and transfluence. In some
instances they are able to impose a pattern which is radial
from the centre of greatest ice accum ulation, an exam ple
being the Fiordland region of New Zealand (Fig. 67) . In
the case of some glacial troughs, in particular on the m argins
of ice sheets, the ice is intrusive and moves up instead of
down the pre-existing river valley.
It is probably most convenient to discuss the charac
teristics of glacial troughs by reference first to their cross
profile and then to their long profile.
Cross profile of troughs

Glacial troughs are traditionally described as being U-shaped
in cross-section when allowance is m ade for superficial
deposits which often mask the bedrock surface. T hey m ight
m ore accurately be com pared to a ro unded V in which the
steepness of the sides and the radius of curvature of the
149
bottom vary in relation to a n um ber of factors, which in
clude the am ount of glacial erosion and the form and
67 Part of Fiordland, New Zealand. Note the relationship of piedmont

lakes such as Poteriteri, Hauroko, and Manapouri to the coastal fiords:
also the development of a radial pattern of glacial troughs.
lithology of the pre-existing valley. Svensson (1959) con

cluded from the study of a large num ber of cross profiles
that they norm ally approxim ated to parabolas.
Glaciated Mountain Valleys 151
The production of this cross profile implies lateral, but
not necessarily vertical, corrasion by the ice and there is
general agreement that this is occasioned by the much
greater cross-sectional area of an ice stream when compared
with a water stream draining the same catchment. The
glacial trough must be thought of as the bed of the glacier
and its sides are therefore analogous to river banks and not
to the sides of river valleys. If the analogy may be carried
further, the sharp convex break of slope at the top of the
river bank is often represented by a similar break of slope
at the upper limit of the glacial trough.
Truncated spurs. The lateral corrasion that becomes

necessary when a valley glacier accommodates itself in a
pre-existing river valley is accompanied by a reduction in
68 Some characteristic features of a glacial trough. A former tributary

glacier is likely to be associated with a glaciated hanging valley and a
valley step.
the sinuosity of the valley floor and a truncation of the

spurs formed between the points of arrival of tributary
valleys (Fig. 68). This shearing or truncating of spurs is
the major reason for the trough-like or trench-like appear
ance of glaciated valleys, an appearance which allows an
observer to see the floor of the valley along m uch greater
distances than is norm ally possible in the case of a river.
T ru n c ate d spurs are m arked by facets of roughly tri
angular or trapezoidal shape, although it should be noted
th at in m any cases facet production may be partly a result
of postglacial dow ncutting by streams trib u tary to that
re-occupying the m ain glacial trough.
In some instances the spur may not be w orn away com
pletely and the basal portion rem ains as a reduced and
m am illated rem nant projecting into the trough floor.
Some of these have been described as being semi-detached
or isolated from the trough wall itself so th at they are knob
like in form. T h e reasons for their isolation are not known
b u t it has been suggested that it may be brought about by
incision of lateral m eltw ater streams.
Ice-eroded spurs projecting into the m ain valley at the
junction with a trib u tary glacier have been term ed bastions.
A bastion is thought to occur because the tributary ice
pushes the m ajor ice stream away from the valley side, which
consequently suffers less than norm al reduction.
A lthough m any trough cross profiles show a high degree
of truncation of structure this is by no m eans always the
case, and structural benches occur in situations where pre
vious subaerial erosion or even glacial erosion has been
guided by subhorizontal bedrock trends. A good example,
cited by Cotton (1947), is that of the ‘groove and bench
terraces’ described by Park (1909) from some glacial troughs
in New Zealand.
Trough shoulders. T h e upper lim it of the trough wall

is term ed the trough edge and this, as previously noted, is
norm ally m arked by a sharp, convex break of slope. In the
case of Icelandic-type troughs the ground above the trough
edge is often part of the plateau surface and the break of
slope is m ore or less rectangular, as it is in the upper Mersey
trough in central Tasm ania. In Alpine-type troughs on the
other hand there is often an interm ediate slope, the trough
shoulder, betw een the steeper trough wall below and w hat
is often a steeper slope above. T h e ‘a lp ’ slopes of Switzer
land provide the classic exam ple of such shoulders.
Glaciated M ountain Valleys 153
A n u m b er of explanations have been p u t forw ard for
trough shoulders (Fig. 69). T h e simplest and perhaps the
m ost fundam entally applicable is that they represent that
p art of the original valley side which has not been over
steepened by glacial erosion because the glacier has n o t fully
occupied the available valley. In this conception the trough
edge coincides w ith the lateral lim it of the glacier, or at
least its lateral lim it d u rin g phases of greatest erosional
activity. T h e contrast in gradient betw een the trough wall
and the trough shoulder may be accentuated in another
way, because while the trough wall is being oversteepened
by the glacier the trough shoulder can be expected to be
undergoing a reduction of gradient by processes of nivation
and mass m ovem ent.
Trough
shoulder
Trough
69 Four circumstances of trough cross profile development. An Icelandic

type trough (C) contrasts with three Alpine type troughs (A, B, D). In A
the preglacial valley-in-valley form results in pronounced trough shoulders.
In B the trough edge is a result merely of glacial occupation of a V-shaped
valley. D illustrates the possible development of trough shoulders by glacier
overdeepening and oversteepening in glacials followed by lowering of the
upper slopes by periglacial mass movement in interglacials.
Some trough shoulders may be inherited from the pre

glacial relief and result from a valley-in-valley form produced
by in term itten t river incision. It has been suggested that
others reflect a history of m ultiple glaciation in which older
trough walls have been degraded and partly resteepened
by later glacial action. These cyclic views are not widely
held b u t the inheritance of riverine valley-in-valley forms is
strongly suggested by some European studies (Louis, 1962) .
Many shoulders are undoubtedly the result of structural
influences, and Cotton (1947) has pointed out that some of
the most striking examples from the European Alps and
from western North America clearly have a relation to
structure. However, it is not always clear whether structure
has caused the development of shoulders or has merely
determined their location.
Trough shoulders are a frequent site of cirque develop
ment, and the occurrence of a number of cirque glaciers
side by side along a trough shoulder suggested to some
earlier workers that the shoulder itself was the result of
cirque cutting along a horizon of preferred development
(cf. p. 145). Most, if not all, authorities today, however,
would regard cirque location and development as a
secondary feature resulting from the favourable conditions
for snow accumulation along the shoulders.
In all the explanations suggested so far it has been
assumed that trough shoulders are developed during glacia
tion, but a good case (recently summarised by Thompson
(1962)) can be made out for an origin which is interglacial,
at least in part. Thus it may be argued that shoulders are
produced essentially by mass wasting and particularly by
periglacial mass movement, so that the trough edge must
be interpreted as the general lower limit of periglacial action
in interglacial times. The frequency with which the modern
interglacial treeline coincides with the trough edge has been
pointed out. The headward intersection of slopes developed
above the trough edge has been put forward as an explana
tion for the frequent accordance of alpine summits over
wide areas (p. 147).
The lower portion of some trough shoulders shows
obvious evidence of glaciation and may be markedly
abraded. Such an ice-scoured surface is termed Schliffbord
in German and its upper limit Schliffgrenze, and there
appear to be no satisfactory alternative English terms,
although Lawrence (1950) used the term trimline to describe
the upper limit of glacial action in a trough. Where a
trough shoulder is in the nature of a preglacial inheritance
the schliffbord is easily explained, but where the shoulder
is thought to be the result of glacial oversteepening of the
valley sides the schliffbord would seem to result from a
relatively short phase or phases of overflow when the ice
spilled over the trough edge, smoothing and abrading the
shoulder without removing significant quantities of rock
material.
Hanging valleys. An expectable result of the oversteepen

ing of valley sides by glaciers and the relatively rapid
evolution of shoulders is the production of discordant stream
junctions and hanging valleys (Fig. 68). T hat these are a
common feature of existing glacial troughs is a matter of
observation, but by no means all are due to tributary
streams being left to cascade over the trough edge when
their lower courses have been removed by glacial excava
tion. Many hanging valleys occur where tributary glaciers
made a discordant junction with the main ice stream because
smaller glaciers tend to dig less deeply than the larger
glaciers into which they become incorporated. Some hang
ing valleys lie below hanging cirques which may or may not
be related in altitude to trough shoulders. It is possible
therefore to distinguish between non-glaciated and glaciated
hanging valleys. Glaciated hanging valleys may have an
associated bastion as described on p. 152.
Although glaciation is clearly conducive to the subsequent
appearance of hanging valleys, particularly because of the
overdeepening or oversteepening of troughs, it should be
borne in mind that discordant stream junctions may occur
in non-glaciated country for a variety of reasons and their
presence by itself is no infallible criterion of the former
existence of glaciers.
Erosional modifications. Meltwater streams are liable to

modify the cross profiles of troughs in varying extent. Ice
margin streams may cut rock benches into the trough wall
or channels in the trough shoulder. Subglacial melt streams
may cut minor valleys into the floor of troughs below the
level reached by ice action. In postglacial time greater
modifications arise from the action of rivers cutting into the
trough floor and incising the trough shoulders.
Depositional modifications. Particularly in the period
im m ediately following deglaciation, depositional effects are
m uch m ore im p o rtan t than erosional effects in modifying
what is usually regarded as the typical trough cross profile.
T h u s the m arkedly flat floor of many glacial troughs is due
to bodies of ground m oraine and outwash, sometimes of
enorm ous volum e and often incised so as to form a series of
terraces. Im m ediately on the retreat of the glaciers, lateral
m oraines are left against the lower trough walls, often, as
in recently deglaciated sections of Swiss valleys, term inating
upw ard along a strikingly sharp line; b u t this is a short
lived phase. L ateral m oraines are usually dispersed by mass
m ovem ent and swamped by colluvial m aterial from above,
which often builds into prom inent screes along the foot of
the trough walls. H anging tributaries commonly form
alluvial fans and cones, which in tu rn influence the course
of the stream occupying the trough floor.
Trough
headwall
Tread
Reverse
slope tread
basin
70 Components of the long profile of a glacial trough
L ong profile of troughs

T h e long profile of glacial troughs is characterised by the
occurrence of relatively sharp convex breaks of slope or
valley steps, which in series give rise to a glacial stairway;
by the presence of occasional reverse slopes in the bedrock,
sometimes occupied by lakes b u t often obscured by drift;
and by overall and local tendencies for greater vertical
corrasion upstream than would occur in the case of a river
valley (Fig. 70) . These three characteristics are closely
in terrelated and are responsible for the ‘down-at-the-heel’
appearance of so many glacial profiles to which W . D. Jo h n
son (1904) drew attention. T h e scale at which such charac
teristics may be observed varies enorm ously: they may give
rise to q u ite m in iatu re features and to others extending for
m any miles.
/ Shelf
Lake Seal
Lake Webster
3000-
71 Long profile of the Lake Seal-Broad Valley glacial trough at Mt

Field, Tasmania. The Tarn Shelf is a cirque terrace separated by a trough
headwall from Lake Seal which in turn is separated by a valley step from
Lake Webster.
Valley steps. M any valley steps appear to represent the

exaggeration by glacial action of a pre-existing convexity
and may thus represent m odifications of nickpoints in the
original river profile, b u t others are determ ined in location
by additional factors. Some are clearly associated w ith ice
junctions and seem to result from the sudden overdeepening
necessitated by the influx of an additional volum e of ice
into the m ain channel. T h e ju n ctio n of a trib u tary glacier
w ith the m ain tru n k glacier is therefore a favoured place
for step form ation. In m any troughs, especially those of
alpine type, the most spectacular step is the trough headwall,
located betw een the head of the trough and a group of
cirques from which ice is fed into the m ain stream. T ro u g h
headwalls are usually absent where the trough is related to
one m ajor cirque (Fig. 61), and are most notable where a
large volum e of ice is produced in a relatively large upper
area of accum ulation and then has to be accom m odated in
a relatively n arrow trench below. A good exam ple is pro
vided by the trough headwall beh in d Lake Seal in the M t
Field massif of T asm ania (Fig. 71) where ice from the wide
collecting area of the T a rn Shelf was concentrated into the
head of what later became the Lake Seal-Broad River glacial
trough.
Some valley steps appear to be controlled by lithology
and to have originated where a sharp discontinuity in bed
rock caused a sudden change in the ability of the ice to
erode. In most of these cases resistance to plucking rather
than resistance to abrasion is probably involved and changes
in the nature of the jo in t system have often been suggested
as being particularly likely to be involved. T h u s the step
may be located where joints are especially widely spaced
and where closely jointed rock downvalley has been very
vulnerable to plucking. A t the same tim e it should be borne
in m ind that such changes in bedrock w ould have rather
sim ilar effects in inducing differential dow ncutting by
rivers, and accum ulation of fluvial nickpoints at hard rock
bars is a well-known phenom enon. In itiatio n of valley steps
by river action is thus always to be suspected.
A lthough there may be some general agreem ent now on
factors influencing the location of steps, the processes giving
rise to the typical step itself are by no means properly u n d er
stood. M any glacial valley steps consist simply of a riser and
a tread, with the tread following the general slope of the
trough floor. In many others the tread has a reverse slope
and is separated from the riser by a raised rock bar known
as a riegel (Fig. 70) . It seems probable that the develop
m ent of steps is closely connected w ith the dual nature of
glacial corrasion, and to the m uch greater quantitative
im portance of plucking when com pared with abrasion as an
agent of rock rem oval. Because the steeper sections of the
valley long profile present the most favourable locations for
glacial plucking, they are also likely to suffer the greatest
am ount of erosion. Such plucking appears to m aintain or
even to steepen still further the initial slope and the effect
is that any pre-existing convexity in the profile becomes
exaggerated and m ore step-like in appearance. A riegel then
is essentially a stoss and lee feature in which the lee slope
has a notably greater am plitude than the stoss slope.
W hile this may be a general explanation of step develop
m ent in glacial troughs, it is doubtful w hether it constitutes j
a full explanation since, am ong other things, it does not
account properly for the steepness and often cliff-like appear
ance of the riser. Very m any risers are not cliff-like and few
approach the degree of steepness represented in exaggerated
text-book diagrams such as Fig. 70: yet there is general
agreem ent that they are of exceptional declivity, and the
need to account for this has engaged the special attention
of m ost of those who have w ritten on the problem of steps.
Analogies have been made, notably by C ham berlin and
C h am berlin (1911), betw een the step riser and the cirque
headw all, and there is a parallelism betw een explanations
given for both. Transverse crevasses, form ed at the ice fall
as the glacier moves over the convexity, have been invoked
to allow freeze-thaw action on the riser, just as the berg-
schrund hypothesis was erected to account for frost shattering
on the cirque headwall. Transverse crevasses rarely extend
rig h t through to the base of the glacier and are no t now
m uch favoured as a possible agent, b u t there is probably
a general opinion that some form of sapping takes place
a ro u n d the foot of the riser and that this contributes to its
steepness. Such sapping is likely to be due to freezing of
w ater b u t it is difficult to invoke w ater percolation b o m
above as in the case of cirque headwalls. T h e most likely
explanation may concern the significant variations in ice
pressure which m ust take place as the glacier moves over
different segments of the step and the variations in the
freezing point of subglacial w ater which m ust occur as a
result (see especially Holmes, 1944; W . V. Lewis, 1948).
T h e relatively th in extended ice passing over a step is
associated w ith a reduction of pressure and consequently
the possibility of freezing of super-cooled water at the riser.
T h e sapping or un d ercu ttin g may be associated with
accum ulation of water at the base of the riser as may be
the case w ith cirque headwalls. W . V. Lewis (1948) also
pointed to the possibility of descending m eltw ater streams
being able to keep open the base of some crevasses and
allow entry of w ater and air from above to subglacial gaps.
D ilation jointing due to spontaneous expansion of rocks
form erly subject to compression, b u t from which confining
loads have been rem oved by erosion, may also aid in riser
recession and. evolution as it appears to do in the case of
cirque headwalls. The frequent association of stepping with
massive crystalline rocks seems to be in accord with such
an idea.
4000
3000
8 2000
u-
1000
0
"20 "is "77
Miles
72 Long profiles of the glaciated South Esk valley and the neighbouring
unglaciated Prosen valley. Also shown are their respective interfluves and
the snowline suggested by local levels of cirque cutting. The South Esk
trough has been cut downward notably more in its upper section than has
the Prosen valley (redrawn from Linton, 1963).
Down-at-the-heel effect. The down-at-the-heel effect may

also be largely a result of the great relative efficiency of
plucking in terms of the volume of rock removed and seems
to be detectable not only in individual segments of the
trough, such as below valley steps, but also when one looks
at the entire profile. One of the most striking things about
most glacial troughs is the relatively great depth of their
upper sections below the general level of the landscape,
apparently as a result of the great volume of erosion
accomplished in these upper sections as compared with parts
of the trough lower down. The effect is that the glacial long
profile tends to be steeper in its upper section and less steep
in its lower section than the corresponding fluvial profile.
Linton (1963) has redrawn attention to this phenomenon
and provided examples from a number of regions (Fig. 72) .
His conclusion that it involves the removal by glaciers of
great depths of bedrock seems an inescapable one and his
suggestion that dilation jointing brought about by the
successive removal of rock masses may be important in
preparing the subjacent rock for removal is in line with
discussion above on the evolution of cirque headwalls and
valley steps.
Two broad factors may contribute most towards down-
at-the-heel effects. At both local and regional scales the
33 Mt Aspiring, New Zealand, a glacial horn flanked by cirque glaciers
(N.Z. National Publicity Studios)
34 Col de la Rousse. An inosculation col between cirques in the French

Alps near Chateau Queyras (]. N. Jennings)
35 Stepped, asymmetrical glaciated valley, West Coast R ange, Tasmania,
with lateral moraine (J. A. Peterson)
36 M ilford Sound Fiord, N ew Zealand, looking seaward. Trough

shoulders, truncated spurs, and a well-marked hanging valley are visible.
T h e glacial trough of Sinbacl v a lle y on the left of the photograph heads
in a large valley-head cirque (National Publicity Studios, Wellington)
particular effectiveness of glacial plucking with its strong
quantitative importance on steeper downstream slopes must
lead to heavily increased erosion in the upper steeper
sections of troughs as it appears to do where the trough is
steepened locally. Headward erosion of troughs has been
envisaged by numerous writers and extensively and explicitly
by W. D. Johnson (1904) and Taylor (1914). It is clearly
inherent in any recession of valley steps and trough
headwalls.
The second factor, which is independent of the first, is
the location of the firn limit which, at the time of greatest
glacier erosion, commonly lies at or somewhat above the
point of greatest inflexion in the long profile curve (e.g.
Fig. 72). It has already been noted (Fig. 41) that greatest
ice velocity and greatest ice thickness is likely to occur
around the firn limit. Other things being equal this is
likely also to be the zone of greatest erosion or more
specifically the zone of greatest potential debris transport.
It is not difficult to imagine exceptionally large amounts
of rock, prepared perhaps by dilation jointing, being re
moved very effectively from this zone in a way which the
geomorphic evidence seems to require.
Reverse slopes. Often — but not necessarily — local down-

at-the-heel sections of the trough are associated with the
formation of reverse slopes and therefore rock basins. It is
probable that, where reverse slopes occur below steps, they
may be associated with a rotational movement of the ice
rather similar to that which can occur in cirque glaciers.
Hobbs (1911) suggested that, if processes operating on the
riser of a valley step were similar to those operating on a
cirque headwall, then they may well have a downward as well
as a backward component, and this might explain reverse
slopes, at least in some instances. It seems not unlikely that
any sapping processes, localised at the base of the riser and
dependent on an exceptional supply of water, operate
downward as well as backward.
Where reverse slopes occur away from a step they are
associated, perhaps invariably, with a reduction in the power
of the glacier to erode. In some cases this may be due to
changes in lithology, b u t m uch m ore comm only it is due
to some increase in the effective w idth of the glacier which
causes a reduction in ice thickness and velocity. T hus
favoured locations for reversed slopes are where some con
striction is rem oved and the glacier is able to spread out
Isobath interval
10 feet
73 The Lake Vera rock basin, Frenchmans Cap National Park, Tasmania.
This was excavated by a valley glacier descending a step immediately to
the southwest and moving northeastward (redrawn from Peterson, 1966).
laterally, and where diffluence occurs and part of the ice

stream is abstracted. Just as a valley step or sudden deepen
ing of the trough may occur when a trib u tary glacier joins,
so a reverse slope or sudden shallowing of the trough is
likely to occur when a distributary glacier leaves.
W hatever the in itiating factor it is the ability of the
glacier to transport upslope that is crucial.
Rock basins. T h e occurrence of reverse slopes and rock

basins is not always im m ediately apparent in the deglaciated
trough because m any are filled w ith d rift deposits and may
only be detected by subsurface exploration. O thers are
m arked by the form ation of lakes, usually elongated because
of their alignm ent along the floor of the trough. T hey may
be called trough lakes and long profiles of their floors often
disclose a down-at-the-heel effect (Fig. 73) . T h e term
paternoster lakes has been used for a series of m ore rounded,
beadlike lakes lying in less elongated depressions or resulting
from the partial postglacial filling of form er trough lakes.
All these lakes are likely to be com plex in origin and to
involve some degree of dam m ing by drift.
It follows from the discussion of reverse slopes that basins
in troughs are likely to occur at the foot of steps or to be
associated w ith locally increased erosion by the form er
glacier. A lthough this latter case may be due to a local
bedrock especially susceptible to erosion, it is m uch m ore
likely to result from a constriction in trough w idth which
resulted in an increase in ice thickness and ice velocity. As
an apparent result m any trough lakes seem to be positioned
at particularly narrow sections of the trough in which they
lie.
Piedm ont lakes. A generally larger and considerably

deeper type of trough lake, form ed where a valley glacier is
about to emerge on to a piedm ont plain, may be distin
guished as a piedm ont lake. A good exam ple is provided
by Lake St Clair in Tasm ania (Fig. 74) which is 14*5 km
long, L 5 km wide, and over 170 m deep. A lthough to some
extent it is m oraine damm ed, by far the greater proportion
of this depth is the result of excavation by a large valley
glacier com ing down the Narcissus trough to the north.
Lake St C lair lies in that p art of the trough constricted
betw een M t Olym pus to the west and the edge of the
C entral Plateau to the east and at its southern end is a great
series of festooned end m oraines, lying on the edge of the
piedm ont plain, and m arking a n um ber of closely spaced
recessional positions of the glacier snout.
T h is assemblage of features is found in piedm ont lakes
in many other parts of the w orld and may be interpreted
as follows. C onstriction of the trough tends towards the
164
\ i Jv
74 Lake St Clair, central Tasmania. This piedmont lake lies where a

large glacier coming from the north emerged on to more open country and
produced a long series of end moraines. M t Ida is a glacial tind (redrawn
from Derbyshire, 1963).
formation of a trough lake, but overdeepening is increased
by the fact that the terminus of the glacier lies immediately
below the eventual position of the piedmont lake for con
siderable lengths of time. This is indicated by the large
series of closely spaced recessional moraines and is brought
about by the sudden emergence of the valley glacier on to
the piedmont. Escape from constriction in such an expanded
foot glacier means that advance or retreat of the ice edge
in response to the prevailing glacier economy becomes very
much slower. Because of the pronounced upward move
ment of ice and its enclosed debris near the snout of a
glacier, a static position of the snout is always conducive to
excessive vertical corrasion upstream.
In many regions, such as at the junction of the Italian
Alps with the North Italian Plain, a number of valley
glaciers emerge on to the piedmont virtually side by side.
Piedmont lakes tend therefore to form along a well marked
line as they do along the eastern edge of the Norwegian
highlands.
Fiords. There has been considerable discussion on the

origin of fiords, but much of it has concerned the origin of
the initial valleys occupied by glaciers and from which they
were produced rather than the origin of the fiords them
selves. Thus fiord patterns often show the marked influence
of structures such as fault systems and fold axes in their
plan, but so do river valleys in general and, since it is known
that glacial troughs are normally modified river valleys, it
seems reasonable to assume that the factors affecting sub
aerial drainage patterns will be reflected in the disposition
of the troughs from which fiords are eventually produced.
Many authors, not the least of whom was Cotton (1947),
have realised that fiords have a good deal in common with
piedmont lakes and that in the main they may be thought
of as piedmont basins which extend into the sea. Like pied
mont lakes they tend to form along a line where the valley
glaciers emerged on to a piedmont, but in this case the
piedmont is beneath the surface of the sea. Like the pied
mont lake, fiords are steep-sided, enclosed basins with
bedrock thresholds often, if not usually, capped by submerged
Charles \
~ \ Sound
2 Miles
3 Kilometres Depth in fathoms
75 The Charles Sound fiord , South Island, New Zealand. Note the
characteristic enclosed basins and threshold.
moraine systems near their mouths (Fig. 75) . The geo

graphical distribution of both forms also suggests a strong
genetic relationship, for in Scandinavia, British Columbia,
Patagonia, and New Zealand glaciated highlands can be
found bounded by piedmont lakes on the landward side
and by fiords on the seaward side (Fig. 67). It is difficult
to believe that they can be other than two expressions of
what are essentially the same processes and circumstances.
The great depth of many existing fiords — over 1300
metres in Norwegian and Patagonian examples — in com
parison with the amount of sea level change to be expected
from eustatic and local isostatic movement suggests that
most of them were excavated below water. The typical fiord
is not a drowned feature in the sense that the typical ria is
a drowned river valley. In the one case the glacier was able
to erode well below sea level, to a depth of nine-tenths of
its own thickness, and subsequent drowning has increased
soundings in the resulting fiord by only a hundred metres
or so: in the other case the river was able to erode only to
the glacial sea level and the depth of the postglacial ria
owes everything to subsequent drowning. Whereas a ria
coast may generally be taken as providing evidence of
submergence, a fiord coast may not.
Erosional and depositional modifications. As in the case

of the cross profile, the long profile of glacial troughs is
modified, especially in detail, by meltwater erosion which
tends to destroy features such as valley steps and by
deposition of valley trains which tend to obscure steps and
rock basins. In general, erosion is more important in the
upper, steeper, part of the trough and deposition in the
lower, more gently graded, section. The more frequent
occurrence of valley steps in the upper lengths of troughs,
as noted for instance by W. V. Lewis (1948), may be in
part due to their being masked lower down. What appear
to be drowned steps have been identified on the floor of
trough lakes, and many are doubtless to be found beneath
outwash mantles.
XI
G L A C I A T E D P LAI NS
G laciated plains may be upland plateau surfaces occupied

by small ice caps or they may be low land surfaces over
which the large continental ice sheets have spread. In either
case the ice will have been relatively little influenced in
behaviour by the pre-existing topography and only to a
m inor extent will it have been channelled along well-marked
avenues of m ovem ent. T h e direction of ice m ovem ent at a
given point will have depended largely on its relationship
to centres of flow, and since centres of flow in ice sheets may
change w ith the developm ent of the glacier, evidence of ice
m ovem ent from different directions at different times is
m ore likely to be forthcom ing from a point on a glaciated
plain than it is from a point on glaciated terrain set w ithin
the confines of a m ountain system. Some of the landform s
of glaciated plains are to be found virtually only in this
environm ent: others are found com m only in the m ountain
environm ent as well b u t are less num erous or less well
developed. In any event it is m ore convenient to group
them under the present chapter heading for purposes of
discussion.
Zones of erosion and deposition

W ith in glaciated plains it is usually possible to distinguish
an inner zone of predom inant erosion occurring around the
centre of form er ice flow from an outer zone of predom inant
deposition. It has already been noted that a sim ilar distinc
tion can often be m ade w ithin a glacial trough, b u t in a
plains environm ent the distinction is norm ally m uch clearer
and m uch m ore im portant in term s of its effect on the total
168
Glaciated Plains 169
landscape. It may be m ade in the case of the largest ice
sheets as well as the smallest plateau glaciers. O n the largest
scale in N orth Am erica and E urope the eroded plains of
the C anadian and Baltic Shield areas contrast with the drift-
covered plains of n o rth ern U.S.A. and the Germano-Polish
lowlands (Fig. 76) . O n a small scale, Jennings and A hm ad
(1957) have described zones of predom inant erosion and
deposition on the C entral Plateau of Tasm ania — the only
glaciated plain of any extent in Australia.
In some sections of the zone of deposition — notably in
parts of Europe — not only the landscape b u t the land itself
may owe its existence to the form ation of m oraines, and
Lam plugh (1920) estim ated that one-eleventh of England
and W ales was created in this way.
Predominant
deposition
16 Zones of predominant erosion and predominant deposition resulting

from the Pleistocene Laurentide and Scandinavian ice sheets.
Erosional landforms
Ice-eroded plains. It is in the in n er areas of glaciated plains
that ice-eroded surfaces gain their greatest expression. H ere
fields of bare rock showing the m yriad repetition of m ore or
less pronounced stoss and lee form occur over potentially
vast areas. Although the appearance of such surfaces suggests
widespread abrasion and plucking by glacier ice it seems
probable that in most instances the glacier does little m ore
than remove the regolith and reveal the basal plane to which
previous w eathering has proceeded. T h is w ould certainly
appear to be the case on the dolerite-capped C entral Plateau
and Ben Lom ond plateau in T asm ania, and most workers
now agree that, even on the continental scale, very little
low ering of relief by glacial corrasion itself is involved in
the process of ice scouring. Thus the lack of relief on
glacially-eroded plains in Finland and Canada is the result
of very long periods of subaerial denudation preceding
glaciation. It is very probable that in many such areas
glaciation has increased rather than decreased the amplitude
of relief by incising along lineaments of exceptionally deep
preglacial weathering.
The bare rock surfaces of ice-eroded plains show up
sharply the structure of the underlying rock, just as do the
bare rock surfaces in deserts, and well-marked trend lines
or lineaments are very much in evidence, particularly when
photographed or seen from the air. The orientation of
ridges, grooves, and basins may be influenced over large
areas by two main factors. The first of these is the orienta
tion of the structural lineaments and this is normally much
the more important. Masses of sounder rock remain as
abraded hills or ridges while shear zones, joint planes, and
the like tend to be gouged out and perhaps partly refilled
with drift materials. The second factor is that of the direc
tion of ice movement which will superimpose stoss and lee
and directional depositional forms on the larger and more
important structural features. The most striking orienta
tions of relief occur where structural and ice movement
trends coincide, and in such cases very pronounced ridging
and grooving may be produced since, other things being
equal, ice seems to corrade more deeply where lines of
weakness lie parallel to its flow. Von Engeln (1935) used
the term ‘grooved upland’ for a region where ice sheet flow
is locally concentrated along pre-existing linear depressions
so as to cause overdeepening, but it must be noted that this
is not the same feature as the ‘grooved upland’ of Hobbs
(see p. 144).
Rock basins. The production on ice-eroded plains of

basins surrounded by bedrock and of many more formed
partly by ice gouging and partly by till deposition leads to
a multiplicity of lakes. Where the lakes are mainly erosional
in origin their outline tends to reflect structural influences:
the composite lakes may be expected to show the effect of
both structural and ice depositional factors. Jennings and
Ahmad (1957) examined and mapped the orientation re
lationships of several thousand lakes within the ice-eroded
plain of the Central Plateau of Tasmania and showed clearly
the effect of structural lineaments on the orientation of
rock-formed shores (Fig. 77) . They noted that this effect
was most pronounced where the ice flow coincided in
direction with the structural trend.
77 Small rock basin lake on an ice-eroded plain, Central Plateau, Tas

mania. The maps show plan relationships with direction of ice movement
and structural lineaments (unbroken lines) (redrawn from Jennings and
Ahmad, 1957).
Roches moutonnees. The ice-eroded plain is the most

favoured location for glaciated knobs displaying streamlined
form and to which the French term roche moutonnee
appears to have become firmly fixed (Fig. 78). In such
circumstances great fields of roches moutonnees may cover
extensive areas, although the feature may occur individually,
of course, in any glacially-eroded country. The characteristic
roche moutonnee presents a streamlined appearance with
smoothed, more gently sloping, upstream end and sides and
a steeper lee side sometimes smooth but characteristically
plucked. These features are best developed where the rock
mass is more or less elliptical, but in practice the shape
varies very much, principally with the structure and pre
existing form. Size is also very variable from quite miniature
examples to large hills.
In some instances the relatively steep lee side of the
glaciated knob is covered by a streamlined tail of glacial
debris. Such crag and tail forms probably show a notably
greater variation in the amount of abrasion visible on their
stoss slopes than do typical roches moutonnees.
78 Some streamlined glacial landforms. The ice movement is presumed

to have been from left to right in each case. A and B: roches moutonnees;
C and D: crag and tail; E and F: drumlins.
Some glaciated knobs have been worn down from pre

existing eminences by glacial erosion: in the case of the
larger ice-shorn hills this is clearly to be expected. However,
many smaller ones, including many roches moutonnees,
have been exhumed by glacial gouging along the lines of
weakness which delineate them and may therefore owe both
their initial isolation and their shape to the action of ice.
T h is is especially likely to be the case on ice-eroded plains
where, as previously noted, relief may owe m uch to the
uncovering of the preglacial w eathering base.
Depositional landforms
T h e extensive developm ent of depositional landform s of
glaciated plains is restricted to lowlands in the northern
hem isphere over which the great ice sheets of the Pleisto
cene extended. T h e great drift-covered plains areas of N orth
A m erica and Eurasia comprise large-scale assemblages of
such forms, well illustrated, for example, by the glacial maps
of C anada (Geological Association of Canada, 1958) and
Sweden (Atlas over Sverige} Stockholm, 1953 c o n tin u in g ).
Landform s prim arily or entirely produced from till occur
in conjunction w ith others produced from outwash: both
historical and geographical relationships may be complex.
In essence, however, it is the last group of processes to have
operated which dom inates the succeeding landscape in any
one area. In some where there has been little or no glaci-
fluvial accum ulation the m ajor resulting landform is the
till plain: in others where the last significant phase of
deposition was carried ou t by m eltw ater the corresponding
landform is the outwash plain.
Till plains. T h e surface of till plains is form ed in the

m ain of ground moraines, the chief characteristic of which
is their general m onotony and evenness of form. T h e ir
effect is to fill pre-existing valleys and depressions and to
plaster hills thinly so that the general am plitude of relief
is m uch reduced. Ver Steeg (1933) from the logs of 2800
borings over 120,000 sq km of C entral O hio found that the
average thickness over buried uplands was about 15 m as
com pared with about 60 m over buried valleys. T h e thickest
drift encountered was 230 m b u t depths of over 300 m are
known elsewhere.
A part from where they are broken by the projection of
higher areas of bedrock, till plains display a surface of very
low undulations often referred to as sag and swell, suggesting
some degree of smoothing on deposition but no amount of
streamlining in a particular orientation. The absence of
streamlining suggests that the smoothing was not carried
out primarily by the ice but results from subsequent re
distribution of clays by mass movement and also perhaps
by flowage in a supersaturated state. From what may
perhaps be categorised as this expectable mean, the surface
can vary in the one direction to knob and basin topography,
showing minimum evidence of smoothing and streamlining,
and in the other direction to fluted moraine and drumlins,
which show the maximum development of these characters.
19 Knob and basin topography on an end moraine near Voltaire, North

Dakota, U.S.A. The edge of the constructing ice sheet lay approximately
WNW to ESE.
Knob and basin topography. Extremely flat or gently

undulating till plains are generally underlain by clays in
which larger particles are relatively absent. If the till is
pebbly or bouldery, a knob and basin topography is likely
to develop because the bigger particles have a larger angle
of rest and are not so likely to be redistributed postglacially
as are the clays (Thornbury, 1954) . But knob and basin
topography is also associated with the deposition of ablation
moraine in which till has been let down more or less
vertically from stagnant ice. The unevenness of the surface
in this case reflects the varying amount of till held in
different parts of the glacier mass before melting and the
inclusion of blocks of wasting ice in the deposited mass,
which blocks eventually melt to form kettles.
W hat is here grouped under the general heading of knob
and basin topography may vary in appearance according to
the extent to which the knobs or the basins dominate. The
most characteristic knob and basin in which the two com
ponents are equally balanced forms a landscape of chaotic
appearance in which numerous small hills are interspersed
with numerous swamp- or pond-filled depressions. A grada
tion is possible to pitted till plains in which the knob
element is more or less lacking, and in the reverse direction
to hummocky moraine in which the knobs are the more
important component. Derbyshire (1963) described and
illustrated hummocky moraine from areas of very gentle
slope in the Lake St Clair district of Tasmania and noted
that the knobs became more conical with a higher boulder
content. He inclined to the view that these moraines
originated within and upon thin masses of motionless ice and
thought they could be distinguished from the hummocky
moraines discussed by Hoppe (1952), which this latter
worker thought originated by the squeezing up of water
logged ground moraine into basal crevasses of thick, almost
motionless, ice.
In summary then it appears likely that knob and basin
type moraine is created in several ways but that the most
important factors are the presence of pebbly or bouldery
till and slowly moving or stagnant ice.
Fluted moraine and drumlins. Smoothed and streamlined

forms apparently oriented in the direction of ice flow at the
time of their formation have always attracted considerable
interest, not only because of their more obvious contribu
tion to the landscape but also because of their importance
as indicators of glacier movement. Drumlins are smooth,
oval hills or hillocks of till variously likened in shape to an
inverted spoon or the top half of an egg split along its long
axis. Where there is a steeper, blunter end, it normally points
in the upstream direction; any more gently sloping, more
pointed end faces in the downstream direction. Although
there is some variation on this typical form, all drum lins
are characterised by their strikingly smooth outlines. D rum
lins norm ally occur in fields of considerable extent and
relatively rarely as individual features. T h ey may in con
sequence often m erge so that they occupy virtually the whole
of the till plain. In other instances they are separated by
flatter areas of till or outwash. Chorley (1959) and Reed,
Galvin, and M iller (1963) have discussed the form and
relationships of drum lins.
'ft Mile
y2 Kilometre
Contours in feet
80 Section of a streamlined till plain with drumlins near Middletown,

N.Y., U.S.A. The direction of ice movement was from the northeast.
T h e factors and processes involved in d ru m lin form ation

are not know n w ith certainty in spite of considerable
speculation (see especially Charlesw orth (1957, pp. 389-
403)) . A lthough it has been argued that they are of
erosional origin and have been carved out of older ground
m oraine, it now seems virtually certain that they are essen
tially depositional and were form ed at the same tim e as the
ground m oraine with which they are norm ally continuous.
It has been widely thought their developm ent may be
triggered by some sort of traction blockage such as an
accum ulation of rath er large-size m aterial at the base of the
ice, and the occurrence of rock cores of varying size in some
drum lins has encouraged this view. It is not known to what
extent rock nuclei are present, b u t sufficient instances have
been described to make it evident that a com plete transition
may be traced from uncored drum lins through rock drum-
lins, in which there is little m ore than a veneer of till, to
crag and tail and roches m outonnees (Fig. 78) . W hereas
the depositional d ru m lin norm ally has its steeper end facing
upstream the erosional roche m outonnee displays the reverse
characteristic.
T h e typical dru m lin also grades through longer and m ore
elongated forms to w hat may be term ed fluted moraine. In
the Narcissus glacial trough in central Tasm ania, D erby
shire (1963) described flutings averaging betw een 4 and 8
feet in height with individual crests which can be followed
for nearly two miles. Features of sim ilar form b u t on a
m uch larger scale are know n from Canada (Flint, 1957).
T h e apparently com plete gradation from flutings to d ru m
lins suggests that all these forms share com m on origins, at
least in part, b u t H oppe and Schytt (1953) have suggested
that some smaller flutings may have been caused by the
squeezing up un d er pressure of plastic till into tunnels
scratched into the base of the glacier by large boulders or
rock projections.
End moraines. It is probably true to say that end

m oraines are a m uch m ore im p o rtan t geom orphic feature
on glaciated plains than in m o u n tain regions. In the m ain
this is because end m oraines constructed across a glacial
trough are particularly vulnerable to rem oval by glacial
m eltw ater and subsequent river action, whereas those b u ilt
at the periphery of ice sheets are m ore likely to survive,
especially where they are left perched on broad interfluve
areas. In addition to this channelling factor, the existence
and bulk of end m oraines is influenced by the am ount of
m eltw ater, the am ount and character of the m aterial de
posited by the glacier, and the length of tim e du rin g which
the position of the glacier snout was stable. As C otton
(1947) has stressed, the bulk of the end m oraine is no guide
to the am ount of erosive work carried ou t by the glacier.
In the case of tem perate-type glaciers in New Zealand, such
as the Tasm an, the comm on absence of defined end m oraines
.
is clearly associated with abundant, well-channelled melt
water (Speight, 1940), and where end moraines have been
built they have often been eroded or buried in outwash
material. On the other hand the absence of recognisable
end moraines from the glaciated Central Plateau in Tas
mania remarked on by Jennings and Ahmad (1957) is more
likely to be due to there never having been a stage when
the ice cap margins were sufficiently stationary, or perhaps
in part to destruction by subsequent periglacial mass move
ment. Drift sheets not bordered by end moraines are said
to be attenuated.
In the case of mountain glaciers, end moraines are likely
to be fairly simple ridges, often looped so that in extreme
cases they assume a hairpin plan (Fig. 81) . The end
44 3 444 445
Contours in feet |/2 Kilometre
81 Hairpin-type end moraine of a small Pleistocene glacier north of

Mt Canopus, Arthur Range, southwestern Tasmania.
moraines of continental ice sheets, however, are very much
larger and more complex features extending for hundreds
of kilometres in the form of belts of hills many kilometres
wide. With retreat and readvance many of these have
assumed considerable complexity of structure, but morpho
logically they have been most influenced perhaps by the
predominance of dumping as the process of glacial deposi
tion and by the incorporation of glacifluvial outwash. The
extensive dumping of till from dead ice at the glacier margin
is mainly responsible for the typically chaotic internal relief
of these end moraine belts, with their multitudinous small
hills and undrained depressions holding lakes and swamps.
Kettling resulting from the temporary survival of detached
masses of dead ice is a frequent cause of pitting, and where
kames are frequent the complex is sometimes known as a
käme moraine.
The end moraines of glaciated plains, like those associated
with cirque and valley glaciers, are often looped in plan in
response to the commonly lobed edge of the ice sheet by
which they were formed. This in turn marks an increased
response of thinner marginal ice to avenues and barriers
posed by the underlying topography.
Outwash plains. The most extensive outwash plains are

associated with former continental ice sheets as in northern
Germany and with piedmont areas such as Bavaria and the
Canterbury Plains of New Zealand where the outwash from
a number of valley glaciers has emerged and coalesced on
to an extensive lowland. In the latter case there is a clear
tendency for the outwash plain to be made up of a number
of cones or fans and, even along the borders of ice sheets,
the outwash tends to head in embayments opposite gaps in
the end moraine belt. Typical of these gaps are the tunnel
valleys of Denmark and northern Germany. Modern out
wash plains are difficult to find because the edges of the
existing large ice sheets lie more or less along the coast. The
best currently forming examples are probably the ‘sandur’
plains of Iceland, and their surface, gently sloping but
marked with numerous shifting braided stream courses,
gives the best indication of how the fossil plains must once
have looked (Krigström, 1962) . T h e Icelandic sandurs have
an overall gradient of about 1:200 to 1:250 b u t steepen to
about 1:60 near the glacier edge so th at they are concave
in long profile. T h e overall gradient of the C anterbury
Plains in New Zealand is about 1:120. A m iniature outwash
plain along the foot of the Frankland R ange in southw estern
Tasm ania has an overall gradient of about 1:100 b u t changes
from about 1:50 near its inner edge to about 1:200 near
its outer edge (Fig. 82) . Variations in gradient seem to be
related in the m ain to the grain size of the dom inant
m aterial involved.
V.E. = 2-5
2000-
Outwash
82 Profile of short glacial valley and miniature outwash plain, foot of

Frankland Range, southwestern Tasmania. Note the contrast between the
glacial and glacifluvial sections of the profile.
Because the constituent fans of outwash plains form at

successive stages of glacier retreat, shallow depressions
parallel to the ice edge may be produced betw een each row
of fans.
M inor surface relief, resulting from braided stream
patterns, seems largely to be obliterated d u rin g postglacial
colonisation by vegetation and may be detectable only by
reference to aerial photographs. Pits produced by kettling
appear to survive m ore successfully and pitted outwash
plains are know n particularly from N orth Am erica. In these
cases the outwash m aterial was presum ably laid dow n over
stagnant ice, or m inor icebergs carried along by m eltw ater
were b uried beneath the sands and gravels.
T h e building of an outwash plain, like the laying down
of a till plain, represents a m ajor phase of landscape
aggradation. The Canterbury Plains outwash is around
500 m in thickness and observation along the edges of the
Icelandic glaciers show that the surface of the sandurs is
many metres above the level of the base of the ice. One
result is that postglacial incision is commonplace, especially
in the case of piedmont outwash, and extensive terrace
systems result.
83 One mode of esker formation. Deposition in a subglacial drainage

tunnel is followed by collapse of the glacifluvial sediments when the ice
disappears (redrawn from Schou, 1949).
Eskers. Long, sinuous ridges of glacifluvial material

lying roughly parallel to the direction of former ice flow
are usually termed eskers, although the term originally had
a wider meaning and has been used by Charlesworth (1957)
to include käme forms as well. Charlesworth and some
others have used the term ‘os' for what is here called an
esker but, in spite of possible confusion when earlier
literature is used, the general custom is to use ‘käme’ for
hummocky outwash forms and ‘esker’ for linear ones.
Eskers vary in size and the biggest may be almost 50 m high.
They also vary in length and degree of sinuosity. In some
the ridge broadens here and there and the esker is thus
beaded. Some extend upslope for parts of their course.
The considerable discussion on the origin of eskers has
not been helped by the fact that there are few known
examples being presently constructed. Those that have been
described support the most comm only held idea that they
are form ed by deposition in subglacial m eltw ater tunnels
so that they represent the aggraded beds of streams left
standing as ridges by the disappearance of the ice (Fig. 83).
T h e ir ability to extend upslope is thus attrib u tab le to
hydrostatic pressure in the enclosed subglacial tunnel. Some
eskers appear to have been bu ilt as a series of small deltaic
cones added one to the other as the ice front receded. In
other cases a subaerial origin in ice ravines cut into a
shallow section of the glacier seems likely. Some shorter
ridges superficially like eskers appear to have originated as
‘crevasse-fillings’ let down from dead or extrem ely passive ice
(Flint, 1928).
T h e variety of form and composition shown by eskers
encourages the view that they are of varying origin, b u t in
all cases it seems clear that they are form ed du rin g the
retreat phase of the glacier d uring a negative regim e w ith
very low ice velocities.
Geographically, eskers are highly characteristic of large
glaciated plains and are found over vast areas of Canada
and Scandinavia. It is notable, however, that they are found
in the zone of predom inant erosion rath er than in the zone
of predom inant deposition, and this is in accord with
their being characteristic of the dying phases of the large
continental ice sheets.
XII
CONCLUSION
W ith in the broad held of investigation of cold climate

landform s and the processes which have given rise to them,
study of different aspects has developed unequally and often
in unrelated fashion. It has been noted (Chapter I) that
glacial studies developed well before the significance of
periglacial processes was generally realised. T h e relative
neglect of the periglacial system has m eant that m uch earlier
work on cold clim ate landform s either ignored periglacial
processes com pletely or attem pted to fit them into a glacial
system. Even today there is not as m uch interrelating of
the glacial and periglacial systems as m ight be desirable.
O ne reason other than the historical one suggested above
may have been the relative lack of understanding of nivation
processes, which in many ways form a bridge betw een the
two m ajor systems. Sometimes they have been included in
a glacial context, at other times in a periglacial context,
and this underlines their essential n ature as a link.
In fact cold clim ate geom orphic processes have a funda
m ental unity stem m ing from their dependence on the
presence of w ater in solid form as ground ice, snow, or
glacial ice, and there is a physical continuum from peri
glacial processes through nivational to glacial, so that it is
impossible to draw a clear m eaningful line. N ot only do
the m ajor systems grade into one another b u t so do many
of the constituent processes, such as those of mass m ovem ent.
T h is co ntinuum has historical and geographical aspects
as well, for the succession of clim atic oscillations through
the Q uaternary has m eant that different processes have
succeeded each other in tim e at the same location and have
changed location at different times. Closer studies reveal,
183
184
Ws W'tm 4‘
; V 5 i- * V
I hg.
. - - '
■
I *.
;
.$ 4
37 Vertical air photograph of part of an ice-eroded plain on the Central

Plateau, Tasmania. Ice movement was from a northerly direction (Lands
and Surveys Department, Hobart)
3<V Multiple end moraines, central Baffin Island. The moraines are up
to about 25 metres high and are thought to harre been formed largely
beneath the surface of a former ice-dammed lake
39 Drumlins fo rm ing islands in Strangford L o ugh, N o rth e rn Ireland
(Aerofilms Ltd)
40 Esker ridge in northern Canada (J. D. Ives)

Conclusion 185
for instance, th at periglacial processes prepare the way for
glacial ones and in tu rn modify the results of glaciation.
In the tem perate hum id regions of the world there re
m ains considerable doubt as to the extent of the landscape
legacy of Pleistocene periglacial phenom ena. In N orth
A m erica the periglacial zone d u rin g glaciations seems to
have been a very narrow one (Fig. 2) and relatively little
residual effect has been postulated; b u t E uropean workers
have identified w hat they believe to be relic periglacial
features on a grand scale. Some of this E uropean work may
have gone too far in ascribing to form er periglaciation
things which are explicable in term s of the ‘norm al’
tem perate h u m id system, b u t there is no do u b t that many
features — such as the convexo-concave slope — which have
traditionally been thought of as typical of ‘no rm al’ land
scapes are in reality largely derived from frost action in the
Pleistocene.
In subhum id areas of southeastern A ustralia a further
problem is that of distinguishing the effects of past clim atic
swings to frigidity from swings to aridity. Both increased
cold and increased drought appear to have the effect of
dim inishing vegetation cover on interfluves and result in
accelerated mass wasting and valley alluviation. Of the two
phases of alluviation shown in Fig. 32 (p. 66), one is
believed to be periglacial in origin and the other to be
a ttrib u tab le to a rath er d rier clim atic phase in the mid-
Recent. O n a w ider scale landscape evolution in the peri
glacial system has m any sim ilarities to landscape evolution
in the arid system, and in the colder, d rier parts of the
w orld distinguishing the in h erited effects of one from the
effects of the other may pose considerable problems.
84 Distribution of Pleistocene cold climate morphogenic zones in Aus

tralia. Glaciated areas (black) are probably reasonably accurate but the
periglacial zone (shaded) is more poorly known. The area shown in main
land Australia is the conservative estimate of Galloway (1965) (opposite).
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INDEX
(Bold figures indicate plate numbers)
Ahlmann, H. W., 78, 85, 87, 96 Blockfields, 43, 44, 49

Ahmad, N., Bartlett, H. A., and Block glacis, 42, 43
Green, D. H., 124 Blockstreams, 36, 42-7, 5
Alaska, 17, 26, 39, 47, 60, 62 Boreal forest zone, 14
Alaska Range, 44 Boulder clay, 107, 121
Alexandre, J., 52 Boulder train, 116
Alluvial fans, 35, 75, 76, 120, 139, Boye, M., 69
151, 156 Braided streams, 121, 127, 129, 179,
Alp slopes, 152 180
Alpine glaciation, 83 Break of slope tor, 47
Alpine summit plains, 147, 148, 154 British Columbia, 166
Altiplanation, 51, 52 British Isles, 107, 169
Anderson, G. S. and Hussey, K. M., Broad River, 157, 158
60 Brochu, M., 11
Andersson, J. G., 30 Bryan, K., 30, 38
Antarctica, 6, 10, 14, 19, 23, 82, 85, Biidel, J., 2, 14, 15, 23, 37, 40, 66,
86, 97-9 68
Aretes, 143, 145, 147, 148 Bunger ice window, 23
Arthur Range, 5, 143, 178 Buttes gazonnees, 54
Avalanche boulder tongues, 75, 16
Avalanche chutes, 72, 74, 75, 77 Cailleux, A., 38
Avalanche debris tails, 76 Cailleux, A. and Taylor, G., 20
Avalanche ramparts, 76 Caine, N., 32, 35, 36, 41, 44, 103
Avalanches, 74-7, 93; dirty snow, Cairnes, D. D., 39
31, 74, 76; ground, 74 Calcareous rocks, 25, 27, 68, 107
Canterbury Plains, 130, 179-81
Baffin Island, 86 Carbonation, 27, 72
Barchans, 128 Carson, C. E. and Hussey, K. M., 63
Basal slip of glaciers, 90-2, 98, 99, Central Plateau (Tasmania), 5, 63,
107, 108 81, 147, 163, 169, 171, 178
Basalts, 26, 47 Chalk, 68
Bastions, 152, 155 Chamberlin, T. C., Ill
Battey, M. H., 139 Chamberlin, T. C. and Chamberlin,
Battle, W. R. B„ 137 R. T„ 137, 159
Beaded stream, 60, 61 Channelled upland, 144
Ben Lomond (Tasmania), 5, 17, 45, Charles Sound, 166
81, 103, 169 Charlesworth, J. K., 101, 147, 176,
Bergschrunds, 95, 136, 137, 159 181
Beskow, G., 21 Chemekov, J. F., 56
Biafo Glacier, 118 Chorley, R. J., 176
Birot, P., 38 Chorley, R. J., Dunn, A. J., and
Black, R. F., 10, 19 Beckinsale, R. P., 1
Black, R. F. and Barksdale, W. L., Chrystocrenes, 46, 47
62 Circles: nonsorted, 23, 24; sorted,
Blockfield saddles, 45 23, 2
193
194 Index
Cirques: deposits, 141, 30; evolu Deltas, 125
tion of, 133, 139, 144, 159, 160; Denison Range, 135
glacial, 72, 131, 149, 156, 157, Denmark, 118, 179
30-3; hanging, 133, 155; lakes, Derbyshire, E„ 126, 175, 177
135, 141, 29, 30; levels, 139, 140, Derbyshire, E., Banks, M. R.,
147, 154; nivation, 72, 132; Davies, J. L., and Jennings, J.
orientation, 133, 134, 142; over N., 105
ridden, 139; terrace, 145, 157; Derwent River, 118
thresholds, 131, 132, 135, 32; two- Diffluent glaciers, 124, 145, 146, 162
storeyed, 42, 140, 143, 29; valley- Dilation jointing, 26, 104, 135, 139,
head, 132, 133, 36 159-61
Classen Glacier, 27 Divides, effect of glaciation on, 124,
Col gully, 126 146, 149
Cols: dilfluence, 143, 145, 146; Dolerite, 33, 41, 42, 45, 46, 51, 52,
inosculation, 143, 145, 146, 34; 63, 71, 103-5, 116, 140, 143
transfluence, 132, 145, 146 Dolomite, 136
Comb ridges, 145 Drainage derangement, 42, 122, 127
Compressing flow in glaciers, 92-4 Drangajökoll, 80
Congeliturbation, 30 Drift, definition of, 111
Continental glaciation, 83 Drumlins, 172, 174, 177, 39; rock,
Cook, F. A., 56 177
Corbel, J., 14, 26, 64 D runken forests, 12
Corte, A. E., 21 Du Cane Range, 83, 132, 147
Costin, A. B., 48 Dumping by glaciers, 111, 115, 179
Costin, A. B., Jennings, J. N.,
Black, H. P., and Thom , B. G., Eakin, H. M„ 26, 41, 52
50, 73 E arth hummocks, 54, 55
Costin, A. B., Thom , B. G., Wim- Eboulis ordonnes, 33
bush, D. J., and Stuiver, M., 35 Edelman, C. H., Florschütz, F., and
Cotton, C. A., 3, 37, 101, 138, 140, Jesweit, J., 30
144, 152, 154, 165, 177 Englacial ablation, 85, 95
Cotton, C. A. and T e Punga, M. Equiplanation, 39
T., 42 Erratics, 116
Crag and tail, 76, 172, 177 Eskers, 119, 181-2, 40; beaded, 181
Creep, 28-30, 32, 33, 43, 46, 49-51 Extending flow in glaciers, 92, 93,
Crescentic fractures, 103 95
Crescentic gouges, 103 Extram arginal channels, 126
Crevasses, 94, 95, 182, 24; basal,
Fahnestock, R. K., 117
95, 108, 175; longitudinal, 95;
Fall sorting, 33
marginal, 95; radial, 95; trans
verse, 95, 159 Fingerprint bog, 64
Cryoplanation, 38, 39 Finland, 170
Cryoturbation, 30 Fiordland, 76, 149, 150
Fiords, 150, 165-7, 36
Daly, R. A., 147 Firn, 78, 94, 96, 105, 110, 136, 147
D art Glacier, 24 Firn basin, 87, 88, 93
Davies, J. L., 45, 48 Firn field, 78, 87, 95, 22
Davis, W. M., 142, 147 Firn limit, 86, 89, 94, 107, 161, 23
Dawson, G. M., 147 Fisher, O., 1
Debris aprons, 42 Flint, R. F„ 8, 85, 103, 119, 125,
Debris avalanches, 31, 74 128, 177, 182
Debris slides, 31 Flowtill, 113
Deflation hollows, 69 Fluvioglacial, 8
De Geer, G., 125 Forth River, 118, 120
Landforms of Cold Climates 195
Frankland Range, 5, 180 Glaciers: active, 96-9; cirque, 79-
Franz Joseph Glacier, 22 83, 90, 105, 108, 109, 131, 133,
Freeze-thaw cycles, 9, 11, 13, 14, 16, 136-9, 161, 179, 17, 33; com
24-6, 137, 138, 140 pound cirque, 80; compressing
Frenchmans Cap, 136, 162, 31 flow in, 92-4; continental, 82;
Fretted uplands, 141-4, 147, 148, 29 dead, 96, 112; dendritic, 81, 82;
Frost: cracking, 20, 28, 29; heave, dum ping by, 111, 115, 179;
18, 19, 21, 28, 29, 32, 33, 49, 50, dynamic classification of, 96-9;
69; mobilisation, 12, 43; shatter expanded foot, 95, 165; extend
ing, 25, 43, 132, 137; sorting, 21; ing flow in, 92, 93, 95; form of,
wedging, 25 79-83; geophysical classification
Frost rubble zone, 2, 10, 14, 15, 27, of, 96; hanging, 79; ice cap, 82,
37, 40, 43, 49, 52, 69, 21 83, 120, 144, 149, 168, 178; ice
sheet, 81-4, 89, 94, 98-100, 108,
Galloway, R. W„ 12, 142, 185 114, 118-26, 148, 168-82, 19, 21;
Germany, 114, 118, 123, 179 internal flow, 90, 92; internal
Gipfelflur, 147, 148, 154 slip, 92-4; niche, 72, 79, 81; o u t
Glacial: abrasion, 102-7, 135, 154, let, 82, 21; overriding by, 113;
155, 158-61; accumulation and passive, 96-9; piedm ont, 81;
ablation, 3, 83, 97, 121; calving, plateau, 80, 81, 108, 149, 169;
83; corrasion, 101-8; deposition, polar type, 96-9, 116, 117; push
111-16; plucking, 102, 104-6, 135, ing by, 111, 113, 114; recon
139, 158-61; protection, 99, 102; structed, 86, 140; relation to
quarrying, 102; transport, 108-11 snowpatches, 15, 132; temperate
Glacial climates, 3, 16, 17 type, 96-9, 116-21; transection,
Glacial drainage channels, 8, 117-27 81, 82, 23; valley, 81-3, 87-9, 93,
Glacial flour, 107 108, 111, 118-24, 141, 162-5, 179
Glacial horns, 143, 145, 147, 148, 32 Glacierisation, 98
Glacial ice: generation of, 3, 78, Glacifiuvial effects, 8, 9, 98, 117-27,
79, 96; stagnant or dead, 46, 86, 156, 167, 177, 178
96, 112, 118, 175, 179-82 Glacifiuvial outwash, 9, 38, 119-23,
Glacial meltwater, 8, 85, 91, 98, 127-9, 156, 167, 173, 177-82
107, 111, 112, 116-27 Glen, J. W „ 92
Glacial monuments, 144 Glen, J. W., Donner, J. S., and
Glacial stairways, 156 West, R. G., 114
Glacial system, 1, 3, 16, 183
Godley Glacier, 27
Glacial troughs, 132, 146, 148-67,
Goede, A., 66
177, 35, 36; Alpine type, 149,
152, 153, 157; composite, 149; Granite, 51, 103, 104
cross profile of, 149-56; Icelandic Granodiorite, 73
type, 149, 152, 153; intrusive, 149; Great Lakes, 123
long profile of, 156, 167; radial Great Plains, 128
pattern of, 149, 150 Greenland, 69, 83, 84, 98
Glaciated areas: Pleistocene, 2, 3, Grey Mare Range, 72
185; present, 1, 2, 10 Grezes litees, 33
Glacier: economy, 83-8; movement, Groom, G. E., 79
88-99, 107, 110, 135; regime, 85, Grooved upland, 144, 170
87, 107, 121; surges, 93; zones, Grooves, 103
86-8 Ground ice, 18-21, 60; perm anent,
Glacier cap, 82 11, 12, 14, 18-20, 40, 54-65, 68,
Glacier margin streams, 123 69; seasonal, 18, 20, 21; diurnal,
Glacier mills, 95 21
Glacier tongues, 87, 88, 109, 23 Ground ice mounds, 54-7, 60
I
196 Index
Groundwater: movement with p er Johnson, D. W., 138
mafrost, 19; segregation, 18, 27, Johnson, W. D., 136, 137, 157, 161
28
Kamb, B. and La Chapelle, E., 91,
Hamburg, A., 22 105
Hanging valleys, 151, 155, 156, 36 Käme, 119, 181
H artshorn, J. H., 113 Käme hummock, 119
Headwall gap, 136-9 Käme moraines, 179
Henoch, W. E. S., 63, 64 Käme terraces, 118, 119, 125, 26
Hewitt, K., 118 Kessler, P., 12
Hobbs, W. H., 78, 141, 142, 144, Kettles, 112, 121, 175, 179, 180
145, 161, 170 Knob and basin, 174, 175
Holmes, C. D„ 114, 115, 159 Krigström, A., 121, 180
Hopkins, D. M., 26, 61, 62
Hopkins, D. M. and Sigafoos, R. Labrador, 17
G„ 55 Lacustrine plains, 124
Hoppe, G., 175 Lake St Clair, 144, 163, 164, 175
Hoppe, G. and Schytt, V., 177 Lamplugh, G. W., 169
Hydrolaccoliths, 55 Landslides, 31
Hydrostatic effects of permafrost, Laurentide ice sheet, 123, 169
19, 55, 57 Lawrence, D. B., 154
Lewis, A. N., 74, 140
Ice avalanches, 85, 86, 140 Lewis, W. V., 70, 71, 104, 135, 137,
Icebergs, 85, 111, 116, 180, 20 159, 167
Ice contact deposits, 119 Linda Valley, 124
Ice contact streams, 117-27 Linton, D. L„ 47-9, 103, 148, 149,
Ice-eroded plains, 169-73, 37 160
Ice fall, 95, 159 Lodgement by glaciers, 108, 111,
Iceland, 179, 180 113-15
Ice lenses, 19, 20 Lodgement till, 114, 115
Ice push ram parts, 63 Loess, 8, 38, 69, 127-30, 11
Ice segregation, 18, 20, 21, 60, 138, Longitudinal dunes, 128
9 Louis, H„ 154
Ice shelves, 82, 20 Lundqvist, G., 33
Ice stream: inset, 110; juxtaposed,
110 Maarleveld, G. C. and van den
Ice wedges, 20, 24, 60, 62 Toorn, J. C., 59
Ice wedge intersection pools, 61 McCall, J. G., 78, 90, 102, 105, 109,
Ice wedge polygons, 20, 61, 10 135, 136, 138, 139
Indicator fans, 116 McDowall, I. C„ 25
Indicators, 116 Macquarie Island, 9, 14
Intergranular adjustm ent, 92 Macrogelivation, 25
Internal flow of glaciers, 90, 92 Mamillation, 103
Internal slip in glaciers, 92-4 Mannerfelt, C. M., 118, 126
Intragranular yielding and re Marginal channels, 126, 155, 23, 25
crystallisation, 92 Marshall, P., 76
Involutions, 21, 22, 24, 55, 114 Markov, K. K., 23
Italian Alps, 165 Mass movement, 1, 11, 12, 16, 28-
Ives, J. D., 125 35, 38-53, 64, 67, 71, 104, 108,
116, 138, 153-6, 174, 178; rates
Jahns, R. H., 104 of, 32; sorting and orientation,
Jennings, J. N., 47 32-6
Jennings, J. N. and Ahmad, N., Matthes, F. E., 70
103, 169-71, 178 Meier, M. F., 78, 89, 90, 94
Meltwater tunnels, 108, 119 Nivation, 15, 16, 31, 70-7, 102, 109,
Mersey River, 118, 120, 122 140, 148, 153
Mersey valley, 106, 152 Nivation cirques, 72, 132
Michaud, J. and Cailleux, A., 32 Nivation climates, 15, 16
Microgelivation, 25 Nivation hollows, 71, 79, 13
M ilford Sound, 36 Nivational system, 16, 77, 183
Monaro Plateau, 48 Norwegian fjelds, 79, 81, 165, 167
Moraines, 45, 111; ablation, 112-14, Nunataks, 108, 109, 19
141, 25; attenuated, 178; end, Nussbaum, F., 137
111-14, 120, 135, 141, 163, 164, Nye, J. F., 92, 93
166, 174, 177-9, 17, 23, 25, 27,
38; englacial, 110; fluted, 174-7; Ohio, 173
ground, 108, 110-15, 156, 173-7; Ollier, C. D. and Thomasson, A.
hummocky, 175; käme, 179; J., 66, 67
lateral, 109-11, 156, 24, 25, 35; Orientation: of mass movement
median, 109-11, 24; push, 113; m aterial, 32-6, 43, 50, 115; of till
recessional, 112, 185; terminal, stones, 114, 115
112; transverse, 109 Os, 181
Mortensen, A., 68 Outwash: apron, 121; plains, 173,
Mound forms, 20, 54-60 179-81; sheets, 121; streams, 117,
Mt Aspiring, 33 120-3
Mt Barrow, 17, 36, 44, 104 Overflow channel, 124, 126
Mt Canopus, 178 Overriding by glaciers, 113
Mt Field, 112, 145, 157 Palmer, J. and Neilson, R. A., 47
Mt Ida, 144, 164 Palmer, J. and Radley, J., 47
Mt Kosciusko, 35, 50, 99, 142 Paisa bog, 56, 63
Mt La Perouse, 74 Palse, 56
Mt Olympus, 163 Parabolic dunes, 128
Mt Twynam, 73, 74, 142 Park, J., 152
Mt W ellington, 45, 52, 104, 14 Patagonia, 166, 167
Mt Wilhelm, 7 Paternoster lakes, 163
M ountain glaciation, 83 Paterson, T. T., 31
Mudflow, 31, 32 Patterned ground, 11, 22-4, 33
Müller, F„ 57-9 Peat mounds, 56, 63
Pediplanation, 38, 39
Naled, 20, 46, 57 Peltier, L., 38
Narcissus valley, 132, 163, 177 Penner, E., 138
N autgarstind, 79 Periglacial areas: Pleistocene, 2, 3,
Needle ice, 21, 28, 29, 37 10, 12, 39, 184, 185; present, 1,
Negative glacial economy, 85, 86, 2, 3, 10, 12
88, 97, 101, 112, 114, 182 Periglacial climates, 9-17
Netherlands, 114 Periglacial system, 1, 9, 16, 18-39,
Nets: nonsorted, 23, 24, 54; sorted, 183-5
23 Permafrost, 11, 12, 14, 18-20, 40,
Neve, 78 54-65, 68, 69, 9; distribution, 10,
New South Wales, periglacial limits, 11, 19; thickness, 19
12, 184 Peterson, J. A., 75, 136, 162
New Zealand, 3, 5, 42, 76, 85, 97, Pewe, T . L., 20
99, 111, 122, 128-30, 149-52, 166, Piedm ont lakes, 163-6
177, 179, 180 Pingo, 54, 57-60; East Greenland
Nickpoints, 157, 158 type, 57-9; fossil, 59, 60; Mac
Nicolls, K. D„ 37 kenzie type, 57-9, 10
Nikiforoff, C., 31, 57 Pipkrakes, 21
198 Index
Pissart, A., 35, 52, 59, 71 Sag and swell, 174
Pneumatolysis, 46 Sand dunes, 37, 38, 69, 127, 128
Poland, 123 Sandurs, 179-81
Polunin, N., 14 Saskatchewan Glacier, 89, 90
Polygenic landscapes, 3, 39 Scalloped uplands, 141-4, 147, 148,
Polygons: nonsorted, 23, 24, 1; 30
sorted, 23 Scandinavia, 5, 17, 182
Positive glacial economy, 85, 88, 97, Scandinavian ice sheet, 8, 100, 123,
169
114, 117
Proglacial lakes, 8, 9, 118, 123-6 Schafer, J. P., 22
Schenk, E., 65
Proglacial zone, 7-9, 116
Schliffbord, 154
Prosen River, 160
Schliffgrenze, 154
Protalus ramparts, 73, 75, 76, 141
Schou, A., 181
Pullan, R. A., 47 Schrund line, 131, 132, 137
Pushing by glaciers, 111, 113, 114 Sharp, R. P., 38, 55, 56, 78, 92, 94,
110
Quartzite, 103, 136, 144 Sharpe, C. F. S., 28, 29, 31, 74
Sheet wash, 35, 39, 68
Raeside, J. D., 128-30 Shostakovitch, W. B., 66
Randkluft, 136 Shumskiy, P. A., 82
Rapp, A., 31, 32, 41, 42, 74-6 Siberia, 5, 11, 17, 19, 20, 39
Rapp, A. and Rudberg, S., 15 Slater, G., 114
Reed, B„ Galvin, C. J., and Miller, Slumping, 31, 32, 9
J. P., 176 Smith, J., 32
Regelation, 91, 92, 105 Snow avalanches, 31, 41
Regelation layer, 91, 105 Snow cornices, 75
Reid, H. F., 94 Snow creep, 70, 73, 74
Reiner, E., 7 Snowline: climatic, 3; effect of
Rex, R. W., 63 aspect on, 6, 7; location of, 4-7,
Reverse slopes in troughs, 156, 161 16, 17, 166; orographic, 3-7;
Rhyodacite, 47 regional, 3-5; relation to glacia
Rias, compared with fiords, 167 tion, 78-81, 99, 123, 140, 160
Richter, E., 147 Snow-melt gullies, 75-7
Riegel, 156, 157 Snow-melt runoff, 15, 41, 62, 68,
Ritchie, A. S. and Jennings, J. N., 70, 72-4
72 Snow-melt transport, 16, 70, 71, 73,
River terraces: aggradational, 37, 133
66, 120; cut in outwash, 122, 156, Snow movement, 15, 70, 73
181 Snowpatches, 15, 70-4, 132; circular,
Roches moutonnees, 171, 172, 177 71, 72; longitudinal, 71, 72;
Rock basin lakes, 141, 156, 161, transverse, 71, 72
162, 167, 170, 171, 32 Snow slide transport, 16, 70, 73, 133
Rock creep, 29 Snowy Mountains, 3, 69, 73, 81,
Rockfall chutes, 41, 75, 15 116, 142
Rockfall funnels, 41 Snowy River, 118
Rock falls, 32 Soil creep, 29
Rock glacier creep, 29, 32, 41, 46, Solifluction, 15, 28, 30-5, 41, 43,
47 46, 49, 53, 64, 68, 70-3
Rock glaciers, 29, 44, 46, 47, 6 Solifluction slopes, 41-3, 3
Rock slides, 32 Sorting: by frost action, 21-4, 32,
Rockslide tongues, 75, 76 33; by glacial outwash, 120; by
Rudberg, S., 55 glaciers, 108, 111; by nivation, 76
South Esk River, 160 Terminal channels, 126, 23, 25
Speight, R., 178 Terraces, 24, 49-53, 64, 71; alti-
Spitsbergen, 23, 52, 68, 71 planation, 52; goletz, 52; groove
Spry, A., 106, 122 and bench, 152; rock-cut, 5T3,
Stephenson, P. J., 99 71, 7, 14; stone-banked, 49-52;
Steps: nonsorted, 23, 24, 50; sorted, turf-banked, 49-52, 71, 8
23, 24, 33, 50 Thaw depressions, 61
Stone garlands, 23 Thaw dolines, 62
Stone pavements, 69 Thaw pits, 61, 11
Stoss and lee, 105, 106, 158, 169-72 Thaw sinks, 62
Stratified slope waste, 33 Thermokarst, 20, 60-2; mounds, 61;
Stream flow in periglacial condi ravines, 60, 61
tions, 37, 42, 62, 66-9 Thompson, W. F., 147, 154
Striations, glacial, 103, 115, 120; Thornbury, W. D., 174
snow slide, 73, 74 Thufurs, 54
String bogs, 63-5, 12 Tibetan highlands, 39
Stripes: nonsorted, 23; sorted, 23, Till, 30, 32, 45, 46, 51, 111-16, 119-
33, 40 25, 173-9; ablation, 112-14, 141,
Subglacial ablation, 85 174; bimodal character of, 106,
Subglacial channels, 125, 127, 155, 111; fabric, 114, 115; lodgement,
28 114, 115; pitted, 175; plains, 173-
Subglacial chutes, 127 9; stones, 46, 114
Subglacial col gullies, 127 Tinds, 144, 164
Subglacial streams, 118, 125, 181, Tjäle, 19, 20
182,25 Tors, 45, 47-9, 4
Submarginal channels, 125, 127 Transfluence, 132, 146
Suffosion, 31 Treeline: effect on glacial land
Summit tor, 47 scape, 147, 154; relation to peri
Superglacial ablation, 85 glacial limits, 11
Superimposed ice stream, 110 Trewartha, G. T., 6
Surface ablation, 85 Tributary glaciers, 109-11, 151, 152,
Svensson, H., 150 155, 157, 162
Sweden, 15, 125, 173 Tricart, J., 13, 14, 25
Swiss Alps, 81, 88 Trimline, 154
Troll, C„ 11, 13, 23, 56, 63
Taber, S., 18, 20-2, 26, 50, 51 Trough edge, 151-4
Taliks, 19, 57, 59 Trough headwall, 132, 140, 156,
Talent, J. A., 47 157, 161
Talus, 33, 35, 40, 139, 141, 156 Trough lakes, 163, 165, 167
Talus cones, 41, 75, 76 Trough shoulders, 152-4, 36
Talus creep, 29, 32, 41 Trough walls, 152, 153
Talus slopes, 40, 41, 44; compound, Truncated spurs, 151, 152, 36
41, 15; simple, 41 Tumut River, 47
Tarn Shelf, 145, 157, 158 Tundra, 2, 10, 14, 15, 45, 54, 62-5
Tasman Glacier, 177, 18, 25 Tundra bogs, 56, 63-5
Tasmania: periglacial limits, 12, Tundra lakes, 59, 61-3
17; Pleistocene snowline, 4, 17 Tunnel valleys, 118, 179
Tasmanian highlands, 3, 23, 33, 35, Tvndall Range, 17
41, 42, 45, 46, 48, 49, 75, 77, 81, Tyrrell, J. B„ 46
99, 103-5, 116, 121, 143 Tyutyunov, I. A., 27
Taylor, G., 140, 161
Tea Tree valley, 66 Urstromtäler, 8, 123
200 Index
Valleys: association w ith snow- W eathering: by chemical processes,
patches, 72; asymmetrical, 66-8; 26, 27, 49, 103, 107, 108; by frost,
dry, 68, 124; in periglacial con 1, 11, 12, 15, 16, 24-6, 29, 37, 40,
ditions, 35, 37, 42, 43, 65-9, 185 46, 49, 51, 69-71, 137; by pressure
Valleyside tors, 47 release, 26; by w etting and dry
Valley steps, 118, 140, 151, 156-62, ing, 26, 27, 108
167, 35 W eathering products, 26-8
Valley trains, 121, 167, 25, 27 W eertman, J., 91, 94
Varves, 124, 125 Williams, P. J., 32, 50, 71
V'entifacts, 38, 69 W iman, S., 25
Ver Steeg, K., 173 W ind: effect in glacial system, 8,
Vesl-skautbreen, 90 85, 120, 127-30; effect in niva-
Victoria, periglacial limits in, 12, tional system, 71; effect in p eri
184 glacial system, 11, 37, 61, 62, 64,
Victoria Island, 32 67, 69
Victorian highlands, 47 W right, H. E„ 38
Von Engeln, O. D., 170
Wahraftig, C. and Cox, A., 44, 47 Zone of glacial ablation, 86-8, 93,
Wales, 71 94, 108
W ashburn, A. L., 19, 24, 31, 32, 50 Zone of glacial accumulation, 86-8,
Wash-cut slopes, 68 93, 94, 96, 110
W ater blisters, 57 Zones of predom inant erosion and
Waters, R. S., 52, 71 deposition, 168, 169, 182
T h e tex t of this book is set in 11/12 p o in t Baskerville type,

an d p rin te d on 85 gsm B u rn ie E nglish F in ish paper by T h e
D om inion Press of N o rth B lackburn, V ictoria
Dr. J. L. Davies is a graduate of the Uni
versity of Wales and the University of
Birmingham. For the last fifteen years he
has worked in Tasmania where he is now
Reader in Geography at the University of
Tasmania in Hobart. His special interest
in the effects of climate on processes of
landform development have led among other
things to investigation of the legacies of
Pleistocene glacial and periglacial conditions
in the island.
An Introduction to
Systematic Geomorphology
Seven volumes are at present planned for
this series, to be published between 1968
and 1971 (provisional publication dates are
shown in parentheses)
1 Humid Landforms
I. Douglas (1971)
2 Desert and Savana Landforms
J. A. Mabbutt (1969)
J. L. Davies (1969)
4 Coasts E. C. F. Bird (1968)
5 Structural Landforms
C. R. Twidale (1969)
6 Volcanic Landforms
C. D. Ollier (1969)
7 Karst J. N. Jennings (1970)
Further details are available from the
A.N.U. Press, P.O. Box 4, Canberra, A.C.T.
2600, Australia.

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J.L.

An Introduction to Systematic G eo m o r p h o Io g y Volume Three

Since W orld W ar II th ere has been a

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T h is series is conceived as a systematic geomorphology at

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III MASS MOVEMENT LANDFORMS 40

IV OTHER PERIGLACIAL LANDFORMS 54

From a geom orphological point of view the cold lands of

3 Altitude of the snowline during the late Pleistocene glaciation of

4 Approximate variation of the regional snowline with latitude. This

T h e nature of the orographic snowline can also be

5 Diagrammatic representation of the influence of aspect on the height

W herever an asymmetry in the orographic snowline

The proglacial zone

6 The proglacial zone in central Europe during the last glaciation. It

M eltw ater issuing from the glacier has m ore or less

Peri glacial climates

Periglacial subdivisions. Ju st as it is im portant to bear

A. Dry with severe winters (associated with perm afrost,

B. H u m id w ith m arked winters

C. Cold w ith little seasonal tem perature change

9 The three process systems of cold climate landform evolution. The

Comparison of glacial and periglacial limits

4 Two-cycle tor near T idbinbilla, A ustralian C apital T erritory (J. N.

There are some geomorphic processes that may be con­

Permafrost. Permanently frozen ground (‘perennial tjäle’

Ice wedges. Segregated ground ice tends to form at right

Non-perennial ground ice. Seasonal ground ice (‘seasonal

10 Involutions at the gravel-shale contact in a ground section in central

Patterned ground. Processes leading to horizontal sorting

11 Sorted stone circles changing to sorted stone stripes on a smooth,

Patterned ground occurs with and without permafrost,

Other mechanical weathering. Taber has also suggested

Chemical weathering. Because of low temperatures, the

Weathering products. T h e great preponderance of m ech­

Mass movement processes

Slow flowage. T h e two m ain types are creep, in which

As m aterial in the waste m antle becomes coarser, soil

Rapid flowage. Flow at higher velocities requires super­

Sliding. Sharpe (1938) defined landslide as ‘the percept­

Rate of movement. The speed with which these various

Sorting and orientation. Material transported by all

S Turf-banked terrace, Snowy Mountains, New South Wales (A. /> .

13 Orientation of stones in solifluction lobes (turf-banked terraces) in

15 Three types of talus slope. A: simple talus slope; B: talus cone;

A lthough the overall profile of talus slopes is generally

Solifluction slopes. T h e most widespread effect of peri­

16 The damming of Lake Fenton, Tasmania, by block glacis. At the

Blockfields. Solifluction slopes on which many of the

Origin of the debris. The material incorporated in

17 N a tu r a l scale p ro file s o f ta lu s slopes on M t B a rro w , T a sm a n ia (a fte r

— “ “ Edge of summit plateau

18 The blockstreams of Mt Wellington, Tasmania

it is composed of glacial till. Through much of the wetter

Rock glaciers. Rock glaciers resemble unvegetated and

19 Three locational types of tor

recent discussion (Linton, 1955, 1964; Pullan, 1959; Palmer

There are some geomorphic processes that may be con

Weathering products. T h e great preponderance of m ech

Rapid flowage. Flow at higher velocities requires super

Sliding. Sharpe (1938) defined landslide as ‘the percept

Solifluction slopes. T h e most widespread effect of peri

Thaw lakes. T h aw depressions tend to be self-perpetuat

31 Evolution of string bogs (as postulated by Schenk, 1966). The move

32 Diagrammatic section across the Tea Tree valley, southeastern Tas

38 A dendritic system of mountain glaciers (A) compared with a tran

Components of m ovement. T w o com ponents of move