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CHAPTER
5
ICE-MARGINAL TERRESTRIAL LANDSYSTEMS:
POLAR CONTINENTAL GLACIER MARGINS
Sean J. Fitzsimons
5.1 INTRODUCTION
Research on landform and sediment assemblages formed by glaciers is dominated by studies of
temperate glaciers in which sedimentary products reflect the influence of basal sliding, subglacial
sediment deformation and subglacial hydrological systems (e.g. Boulton, 1972a and b). In
contrast there have been relatively few studies of landform and sediment assemblages of polar and
polythermal glaciers (e.g. Fitzsimons, 1997a; O’Cofaigh et al., 1999). The objective of this
chapter is to synthesize recent investigations of polar continental glacier margins and move toward
a depositional model for ice-marginal environments that links our understanding of glaciology
and geomorphology. This review is based on field observations in East Antarctica (Vestfold Hills,
Bunger Hills, Larseman Hills and Windmill Islands) and in south Victoria Land (McMurdo dry
valleys and Ross Island). This chapter begins with a definition and review of the physical
conditions that control depositional processes at polar glacier margins. The review is followed by a
summary of the morphology and structure of modern polar glacier margins and associated
landforms and sediments in low- and high-relief environments. The chapter concludes with
synthesis of ideas that form the basis of a depositional model for polar continental glacier margins.
5.2 POLAR ICE-MARGINAL ENVIRONMENTS
Polar continental glaciers constitute the bulk of ice on earth (Table 5.1). Polar glaciers
constitute over 95 per cent of the glacier-covered area and over 97 per cent of ice volume
(excluding the Ross and Ronne-Filchner ice shelves). Although polar continental glaciers
dominate the earth’s glacial systems, sediment-landform associations produced by these glaciers
are the least well known and understood. The primary reason for is because there is little land
area beyond current ice margins and the glaciers are largely inaccessible. Consequently there is
little terrestrial evidence of the growth and decay of the glaciers.
Our knowledge of the geomorphology and sedimentology of large polar continental ice masses
is mainly derived from studies of small land areas that fringe Antarctica and Greenland. These
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GLACIAL LAND SYSTEMS
Area (km2) %
Volume (km3) %
Ice caps, ice fields, valley glaciers
0,680,000
4.24
0, 180,000
0.55
Greenland
1,784,694
11.06
2,620,000
7.96
East Antarctic Ice Sheet
10,153,170
63.27
26,039,200
79.11
West Antarctic Ice Sheet
1,918,170
11.96
3,262,000
9.90
Antarctic Peninsula
0,446,690
2.79
0, 227,100
0.69
Ross Ice Shelf
0,536,070
3.34
0, 229,600
0.70
Ronne-Filchner Ice Shelf
0,532,200
3.33
0, 351,900
1.07
Antarctica
Total
16,051,094
100
32,909,800
100
Table 5.1 Estimated areas and volumes of glaciers (From Williams and Ferrigno, 1993)
areas provide limited access to glacier margins and to landscapes that have experienced the
advance and retreat of glaciers. These areas, often called oases, are cold deserts characterized by
low mean annual temperatures (–10 to –20 °C), light precipitation the vast bulk of which falls
as snow, and strong winds (typically mean monthly wind speeds of 2–9 m.s–1).
The main controls on the nature and location of glacial deposition are glacier mass balance,
thermal regime, bed configuration, the properties of the material being deposited and the
climate near the ice margin (Andrews, 1975; Lawson, 1979). Studies of glacial deposits forming
at the margin of glaciers have stressed the role of thermal regime in determining the processes
involved in their deposition (Boulton, 1972b, 1975). Three different thermal boundary
conditions have been recognized in glaciers on the basis of englacial temperature gradients
(Weertman, 1961):
1. a temperature gradient that is sufficient to conduct all heat from the glacier bed, in which
case there no melting and the ice remains frozen to the bed
2. a temperature gradient that is just sufficient to conduct heat from the bed, in which case
there is an approximate balance between melting and freezing
3. a temperature gradient that is insufficient to conduct heat from the bed, in which case
there is melting and sliding.
These boundary conditions define two types of ice that are often called ‘temperate’ and ‘cold’
ice and recognizes the possibility that the state of the ice may change in space and time. When
applied to whole glaciers the scheme has yielded a three-part classification of thermal or glacidynamic basal regimes that can be identified at modern glacier margins: ‘temperate’ glaciers,
‘subpolar’ or polythermal glaciers and ‘polar’ glaciers. The geographic terminology is regrettable
because the distribution of glacier types is not simply determined by latitude. Consequently the
terms ‘wet-based’, ‘polythermal’ and ‘dry-based’ or ‘cold-based’ are preferred and are used in
this chapter. Most polar continental ice masses are of the polythermal type: where the ice is
thin, such as the ice margins, they are dry-based and where the ice is thick or flowing rapidly
the base of the ice is at pressure melting point and therefore wet-based. Thin glaciers in
particularly cold environments may be entirely dry-based.
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It has been argued that each thermal regime can be associated with diagnostic landform and
sediment assemblages (Boulton, 1972b; Boulton and Paul, 1976; Eyles et al., 1983b). Eyles et
al. recognized a polar arid sediment assemblage based primarily on the work of Shaw (1977a, b)
who examined depositional processes in the McMurdo dry valleys. The title of this sediment
assemblage ‘polar arid’ encapsulates the problem of a thermal regime-based classification
because this sediment assemblage is differentiated by the climate of the terminus area rather
than basal thermal regime. Glacier thermal regime exerts a fundamental control on glacier
behaviour by determining ice motion and erosion processes. Thermal regime is determined by
the englacial temperature gradient, which is influenced by climate and the generation of heat
close to the glacier bed. The indirect role of climate in controlling thermal regime contrasts
strongly with the direct influence of climate on depositional processes at glacier margins. The
use of both thermal regime and climate to distinguish sediment raises several interesting
questions including:
•
•
•
Can the roles of thermal regime and climate in glacier sedimentation at the terminus area
be differentiated?
If the roles of thermal regime and climate can be differentiated, which is the higher-order
control in polar continental environments?
Is glacier thermal regime a satisfactory basis for defining landform and sediment
assemblages?
5.3 ICE MARGINS IN LOW-RELIEF LANDSCAPES
5.3.1 Glacier Margins
In East Antarctica the majority of the ice margin terminates in the sea. Relatively small parts of
the ice margin terminate on land in small coastal oases. The largest oases in East Antarctica are
the Vestfold Hills and Bunger Hills (Fig. 5.1). Recent investigations of the Quaternary history
of these areas has suggested that the ice margin during the last glacial maximum was thinner
and less extensive than previously thought (Colhoun et al., 1992) and that deglaciation was
almost complete by 10,000 years BP (Fitzsimons and Domack, 1993). These conclusions are
clearly controversial as they contradict data from the Ross Embayment (Denton et al., 1989)
and marine seismic and core data in East Antarctica (Domack et al., 1991). As the mode and
pattern of ice advance and retreat have implications for the interpretation of palaeo-climate and
ice dynamics, it is vital to have appropriate depositional models for landforms and sediments.
In Vestfold Hills the edge of the continental ice sheet runs from north to south, and the
southern limit of the ice-free area is formed by the Sørsdal Glacier, which is the major outlet
glacier of the area and forms a small ice shelf. The hills consist of a complex low-relief
topography composed of valleys at and below sea level and ridges up to 158 m in altitude.
Glacial sediments and landforms are absent from most of the ice-free area and are concentrated
close to the glacier margin.
The mean annual temperature of the Vestfold Hills is –10.2°C (Schwerdtfeger, 1970) which is,
on average, warmer than Antarctic stations of similar latitude (Burton and Campbell, 1980).
Although no precipitation data are available, snowfall is light (probably <250 mm per year) and
rainfall is very rare. Melting of snow and ice is restricted to the short summer (December to
February). There is a strong diurnal component to the melt activity, which usually ceases
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Figure 5.1 Map showing the location of the Vestfold, Larsemann and Bunger hills and the
McMurdo dry valleys.
between 9 pm and 10 am when air temperatures are below or close to 0 °C and the sun has a
low angle of incidence.
In many East Antarctic coastal oases the ice-sheet margin has a complex form, and distinct
features such as an ice cliff are difficult to recognize. Consequently it is necessary to define some
terms used in this chapter: ‘ice edge’ is used to describe the terminus of a glacier where it is
sharp and easily recognizable (Fig. 5.2a) and ‘ice margin’ is used to describe an ice-terminus
that it is not clearly recognizable (Fig. 5.2b). Within an ice margin an apparent ice edge is often
recognizable as an ice cliff beyond which an area of ice-cored moraine occurs (Fig. 5.2b). The
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ICE-MARGINAL TERRESTRIAL LANDSYSTEMS: POLAR CONTINENTAL GLACIER MARGINS
width of the ice-cored moraine or debris-covered glacier is highly variable and can range from
tens of metres to several kilometres. The term outer ice edge is used to define the actual glacier
terminus where ice movement ceases (Fig. 5.2b).
The three pre-eminent characteristics of the ice margin at Vestfold Hills are its variable shape,
the presence of a large sinuous ice-cored moraine (Figs. 5.2 and 5.3), and the abundance of
large snow drifts (Fig. 5.3). The ice-sheet margin has a convex form and descends rapidly from
300 m above sea level within 2 km of the margin to approximately 100 m above sea level at the
margin. Where the ice flows into the sea, the ice margin forms 20–40 m high cliffs. On land
the margin is considerably more complex, often with multiple cliffs and snow drifts (Fig. 5.4a).
The sinuous ice-cored moraine that dominates the ice margin at Vestfold Hills is a broad,
discontinuous ridge of coarse debris, 100–300 m wide and about 20 km long that occurs inside
the ice margin (Fig. 5.3). The debris is, on average, less that 0.5 m thick but accumulations up
to 1.5 m thick occur on the sharp-crested ridges. The sinuous inner moraine contrasts with
other moraines that occur in front of ice cliffs which have sharp-crested ridges. Moraine ridges
beyond the ice margin are much higher (up to 20 m) and much shorter than the inner moraine
ridges (less than 1 km long). Most are ice-cored and unstable, as indicated by the occurrence of
numerous sediment flows, slumps and other mass movements (Fitzsimons, 1990).
(A)
(B)
Figure 5.2 Ice-margin nomenclature. A) Simple ice margin with an ice cliff and inner moraine
formed by basal ice cropping out on the glacier surface. B) Wide ice margin with an apparent
ice margin separated from an outer ice edge by numerous outcrops of basal debris.
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(A)
(B)
Figure 5.3 Oblique aerial photographs of the ice margin at Vestfold Hills. A) Looking toward
the ice margin with Sørsdal Glacier at right. Note the large sinuous inner moraine (arrowed)
and snow drift partially concealing the ice margin. The light, turbid lakes are connected to the
proglacial drainage system and the dark ones are not. The ice margin is about 10 km long. B)
Ice margin looking toward the coast. Note the deep snow drifts downstream of the inner
moraine and partly frozen lakes and fiords. The numerous dolerite dykes belie the lack of
unconsolidated sediment over the landscape.
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ICE-MARGINAL TERRESTRIAL LANDSYSTEMS: POLAR CONTINENTAL GLACIER MARGINS
In other parts of the Vestfold Hills the ice margin is buried by snow drifts and forms in a lowangle ramp. Where the margin is not buried, cliffs up to 30 m high occur near the ice-cored
moraines and at the heads of fjords (Fig. 5.4a). In these steeper sections the debris is
concentrated below the ice cliffs (Fig.s 5.4a and b) forming narrow, sharp-crested ridges up to
10 m high.
(A)
(B)
(C)
(D)
(E)
Figure 5.4 Topographic profiles of the Vestfold Hills ice margin. A) Cliffed margin with inner
moraines beyond the apparent ice edge. B) Ramp margin with two sets of inner moraines. C)
Ramp margin with a large snow wedge and two inner moraines. D) Multiple ice cliffs and snow
wedge remnants with a folded inner moraine beyond the apparent ice edge. E) Ramp margin
with numerous outcrops of basal debris.
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In the southeastern corner of the hills, where the Sørsdal Glacier forms a distinct outlet glacier,
the ice margin has a convex profile and has multiple ice cliffs (Fig. 5.4d). The slightly
deformed, basal debris zone is unconformably overlain by clean white ice. This unconformity
appears to record a former ablation surface that has been buried by ice which accumulated in
situ. A zone of ice-cored moraine with numerous sharp-crested ridges parallel to the ice edge
occurs beyond the main ice cliffs.
The structure of the ice margin at the Vestfold Hills is revealed by exposures of the basal
debris zone in ice cliffs and gullies that cross the ice margin. Deformation structures range
from relatively undeformed debris bands to intense deformation characterized by recumbent
folds and shear structures (see Fig. 4 in Fitzsimons, 1990). Deformation structures in the basal
zone of the ice cap can be divided into large-scale features, which involve the entire basal
debris zone, and small-scale features which occur within the basal zone. The most prominent
large-scale deformation structure is the upwarping of the basal debris zone to crop out on the
surface of the glacier and form a large, sinuous ice-cored moraine (Figs 5.3 and 5.4). Exposures
of basal debris reveal structures that vary from slightly deformed stratified ice (Fig. 5.5a) to
complex multi-phase folding and shearing (Fig. 5.5b). A section through an ice-cored moraine
in the southeastern corner of the hills shows that moraine ridges can form along the axes of a
series of large recumbent folds that have amplitudes over 15 m (see Fig. 4c in Fitzsimons,
1990).
Measurements of debris concentrations in the basal zone ice are consistently below 10 per cent
by volume. Debris concentration in individual bands is highly variable with most of the debris
concentrated close to the bed. Unusually high concentrations occur in rare debris lenses of
sorted fluvial sediments that have been entrained by the glacier. Most of the debris consists of
silt and sand-sized particles with larger clasts either dispersed or occurring in small lenses.
Gravel clasts are dominantly subrounded, rarely angular.
5.4 ICE-CONTACT LANDFORMS AND SEDIMENTS
Three types of well-preserved ridges can be recognized at the ice edge in East Antarctic oases:
inner moraines, ice-contact fans/screes and thrust-block moraines.
5.4.1 Inner Moraines
Inner moraines form at the margin of the ice sheet where basal debris crops out and
accumulates on the ice surface. Beyond the present ice margins older inner moraines often form
prominent end moraines. Where an ice core remains in these moraines exposures reveal large
recumbent folds with an amplitude of up to 6 m and numerous smaller sheared folds providing
evidence of intense compressive deformation within the basal debris zone close to the ice
margin (Fitzsimons, 1990, 1997a).
Exposures of the inner moraines reveal massive, matrix-supported diamictons with rare layers of
poorly sorted, sandy gravel. Pebble-fabric strengths of the diamictons measured from clast aaxes range from 0.51 to 0.81 and tend to be weaker closer to the surface of the ridges (Fig. 5.6).
Directions of maximum clustering are perpendicular to the trends of the ridges and in a few
cases oblique to the trends of the ridges (Fig. 5.6).
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(A)
(B)
Figure 5.5 Basal ice at the edge of the ice sheet in Vestfold Hills. A) Slightly deformed
stratified basal ice resting on gneiss and unconformably overlain by drift snow. B) Highly
deformed basal ice showing a series of tight sheared folds.
The diamictons are accumulations of basal debris that have cropped out on the surface of the
ice sheet and subsequently been remobilized by sediment flows. Remobilization has resulted in
relatively poorly defined directions of maximum clustering, and slight textural variation is
probably related to sorting of sediments in less viscous flows. Stronger pebble fabrics below 1 m
depth in the excavations can be interpreted as melt-out till in which the fabric of the basal
debris zone has been preserved. The formation of melt-out tills, and the preservation of basal
debris fabrics that record ice-flow direction are more likely where the sediment cover exceeds
0.5 m, after which melting slows and the debris is less likely to become saturated and flow.
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Figure 5.6 Sedimentary logs of sediments from the crests of inner moraines. The contour
interval of the Schmidt nets is two standard deviations. V1 and P1 and give the azimuth and
plunge of the principal eigenvector, S1 gives the strength of clustering about the principal
eigenvector and R shows the trend of the moraine ridge.
5.4.2 Ice-Contact Fans and Screes
Ice-contact fans and screes form sharp-crested cuspate ridge segments up to 20 m high and 500
m long. They form at ice cliffs where melting and sublimation of basal debris results in the fall
and/or flow of debris at the foot of the cliff (Fig. 5.7a). Most of these ridges have asymmetrical
profiles (Fig. 5.7a) characterized by proximal slopes between 25 and 15° and distal slopes
between 15 and 25° (Fitzsimons, 1997b).
Sediments exposed at the crests of ice-contact fans and screes show a range of sedimentary
facies, including massive and stratified gravels, horizontally laminated and cross-bedded sands,
bouldery gravels with lenses of fine-grained sediment, massive matrix-supported diamictons,
stratified diamictons and muds (Figs 5.7b and 5.8). The sediments range from moderately
sorted to very poorly sorted, but on average are moderately sorted. Particles up to 0.8 m in
diameter are common and occur in a chaotic mixture of diamicton, gravel and well-sorted and
stratified sand. Most exposures show that the sediments are well stratified with dips down the
distal slope of the moraines at angles of between 5 and 20°. The pebble fabric of diamictons
and massive gravels are transverse or oblique to the trend of the ridges (Fig. 5.8) and the
clustering about the mean axis ranges from moderate to strong (S1 0.54–0.86).
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(A)
Figure 5.7 A) An ice-contact scree
forming at the ice margin (left) and
two ice-cored ice-contact screes
adjacent to the ice margin. Note the
supraglacial stream emerging from
the contact between basal ice and
drift snow. B) Poorly sorted gravel
overlain by laminated sand and
gravel, and a clast supported diamict
exposed in the crest of the icecontact scree.
(B)
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Figure 5.8 Sedimentary logs of sediments from the crests of ice-contact screes. The contour
interval of the Schmidt nets is two standard deviations. V1 and P1 and give the azimuth and
plunge of the principal eigenvector, S1 gives the strength of clustering about the principal
eigenvector and R shows the trend of the moraine ridge.
The association of diamicton, gravel, sand and the bouldery facies suggests that both alluvial
and colluvial processes are important during the formation of the ridges (Fig. 5.7b). The
chaotic bouldery lithofacies is interpreted as the product of simultaneous accumulation of
alluvial and mass-movement deposits (i.e. large particles fall or roll into an alluvial deposits and
sediment flows).
5.4.3 Thrust-Block Moraines
Thrust-block moraines form along the lateral margins of outlet glaciers, where ice flows across
marine inlets or lakes. The ridges are up to 20 m high with proximal slopes of around 30° and distal
slopes of around 25°. As the ice core melts, large tension cracks develop along the ridge crests.
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Sediments in thrust-block moraines (Fig. 5.9a) consist of stratified diamictons (Fig. 5.9b),
massive diamictons and rare layers of horizontally laminated sands (Fig. 5.9). Many exposures
display low-angle thrust faults and sheared zones that consistently dip in an up-glacier direction
at angles of between 10 and 25°. The pebble fabric of the diamictons can be divided into a
group characterized by weak fabrics associated with stratified diamictons (S1 0.45–0.57) and a
group of stronger fabrics adjacent to low-angle faults (S1 0.67–0.85). Massive diamictons
frequently contain abundant shell fragments and stratified diamictons occasionally contain beds
of shells, some in growth position (Fitzsimons, 1997b).
The distinctive fabric, lamination, and preserved marine shell, sometimes in growth position,
suggests the diamictons are glacimarine sediments. Pebble fabrics of attenuated diamictons
(faulted and sheared) have similar strengths to deformed lodgement tills described by
Dowdeswell and Sharp (1986). The increased fabric strength is interpreted as a consequence of
attenuation by shearing either as the blocks were detached or deposited. Preservation of beds of
shells and laminations within the diamictons suggests at least some of the sediment may have
been frozen during entrainment and transportation and/or that the strain was relatively low.
Low-angle faults, together with slickensides and attenuated diamicts adjacent to the faults, show
that the glacimarine sediment has been entrained as a series of blocks with an average thickness
of about 0.5 m (Fig. 5.10). The moraines have accumulated as successively older glacimarine
sediments, were eroded from the floor of the fjord and then deposited on the distal shore.
5.5 ICE MARGINS IN HIGH-RELIEF AREAS
5.5.1 Glacier Margins
The high-relief area that is described here is the McMurdo dry valleys, which are often called
the McMurdo oasis. Glaciers of the McMurdo dry valleys can be divided into four groups:
outlet glaciers, ice shelves, piedmont glaciers and alpine glaciers. Ice from the East Antarctic ice
sheet flows through the Transantarctic Mountains to form outlet glaciers, such as Ferrar and
Mackay glaciers, which reach the coast and form small floating ice tongues. Other outlet
glaciers, such as Taylor Glacier, terminate on land. However, it could be argued that Taylor
Glacier is not strictly an outlet glacier of the East Antarctic ice sheet because it flows from a
local ice dome (Taylor Dome). North of the margin of the Ross Ice Shelf, ice streams that flow
through the Transantarctic Mountains form outlet glaciers that feed small ice shelves. The
largest ice shelf in the area is the Ross Ice Shelf, which is fed primarily by ice streams from the
West Antarctic Ice Sheet. Although the Ross Ice Shelf does not directly impinge on the dry
valleys today, during the Late Pleistocene the ice shelf grounded and flowed up the valleys.
Consequently the ice shelf had a profound impact on the geomorphology of several valleys in
the McMurdo oasis. In coastal areas of the McMurdo oasis, the slightly higher precipitation
results in broad piedmont glaciers at the seaward margins of the Victoria and Wright valleys.
Between the coastal piedmont glaciers and the inland glaciers, small alpine glaciers form a
remarkable landscape in which bare rocky slopes contrast strongly with glacier ice. Most of
these glaciers are no more than 15 km long.
Although ice margins in the McMurdo dry valleys range from gently sloping ice ramps to steep
ice margins the most common and distinctive form is a 15–20 m high ice cliff. These
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(A)
Figure 5.9 A) A series of thrustblock moraines on an island adjacent
to the margin of an outlet glacier. B)
Stratified glacimarine sediments
exposed in the crest of a thrust-block
moraine.
(B)
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Figure 5.10 Sedimentary logs of sediments from the crests of thrust-block moraines. The
contour interval of the Schmidt nets is two standard deviations. V1 and P1 and give the azimuth
and plunge of the principal eigenvector, S1 gives the strength of clustering about the principal
eigenvector and R shows the trend of the moraine ridge.
distinctive cliffs have been attributed to changes in the rheological properties of ice at around
20 m thickness (Chinn, 1991) and to a strong reduction in ablation from the foot of the cliffs
to the glacier surface (Fountain et al., 1998). Supraglacial debris is absent from most of the
glaciers and the only visible debris is often restricted to small outcrops at the foot of ice cliffs
where the basal zone of glaciers is exposed (Fig. 5.11). Although the ice margins do not have
the thick snow accumulations that characterise the low relief landscapes described above, they
are often characterised by an accumulation of ice at the foot of the cliffs produced by episodic
calving (Fig. 5.11).
All glacier margins in the McMurdo dry valleys are dry-based with basal temperatures between
–16 and –18 °C, which is very similar to the mean annual temperature (–19.8 °C at Vanda in
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(A)
Figure 5.11 Glacier margins in
the McMurdo dry valleys. A) Clean
white ice and marginal cliff
characteristic of glaciers in the dry
valleys. Note the ice apron
produced by episodic calving (Hart
Glacier, Wright Valley). B)
Stratified basal ice at the margin of
Suess Glacier.
(B)
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the Wright Valley). However, in the case of outlet glaciers such as Taylor Glacier, the ice is at
pressure melting point within a few kilometres of the margin where the glacier is thicker
(Robinson, 1984). The velocities of the glaciers are generally low. In the case of fully dry-based
glaciers surface velocities are less than 1 m.a–1 and 3 m above the bed velocities of around 250
mm.a–1 have been measured (Fitzsimons et al., 1999). The outlet glaciers move considerably
faster.
Exposures of the basal zone of the glaciers show that debris concentrations are generally low and
highly variable. In the case of Suess Glacier in the middle part of the Taylor Valley, debris
concentrations range from less than 0.1 per cent to more than 70 per cent by volume with an
average of less than 5 per cent. Debris concentrations in Taylor Glacier, which is at pressure
melting point upstream of the terminus (Robinson, 1984), are considerably greater.
Ice-marginal landforms are absent or very small at most ice margins suggesting that the coldbased glaciers are not particularly effective agents of erosion. However, a few glaciers exhibit
well-developed end moraines and have been the subjects of recent investigations. These
landforms are described below.
5.5.2 Ice-contact Landforms and Sediments
Several types of moraines are recognized at the margins of glaciers in the dry valleys although
there is considerable uncertainty about their origin. Given this uncertainty for the purposes of
this chapter they are divided into constructional and structural features.
5.5.2.1 Constructional Landforms
Constructional moraines are formed at ice margins where the glacier has been sufficiently stable
to concentrate debris, usually at the foot of an ice cliff, by ablation of the basal debris zone.
Chinn (1991) has argued that the outcrops of basal debris at the foot of ice cliffs are the
equivalent of inner moraines that are commonly seen at dome-shaped polar ice-margins.
In the McMurdo dry valleys, constructional moraines, often covered with an ice and debris
apron occur at the margins of numerous glaciers. These features are formed by debris released
from the basal zone together with sparse supraglacial debris. Shaw (1977b) has argued that
advancing glacier may override the ice and debris aprons thereby incorporating debris into the
basal zone of the glacier. This ‘apron entrainment’ mechanism is similar to processes described
in sub-polar glaciers in the Canadian arctic (Evans, 1989a).
5.5.2.2 Structural Landforms
Small structural moraines occur at the margin of several glaciers in the McMurdo dry valleys.
These features appear to consist of either sediment blocks eroded from the base of the glacier
and/or marginal sediments that have been deformed in situ (Fig. 5.12). Fitzsimons (1996a)
described moraines that formed at the margins of Suess Glacier as thrust-block moraines. This
paper posed the hypothesis that the moraines were produced by accretion of ice and debris as
the cold-based glacier margin advanced into a proglacial lake (see Fig. 4 in Fitzsimons, 1996a).
Subsequent investigations have demonstrated that the formation of moraines in this location is
more complex that initially thought. Isotopic analysis of the basal ice exposed at the foot of the
ice cliff has shown convincingly that some basal ice has formed as water froze onto the base
and/or margin of the glacier (Lorrain et al., 1999). However, excavation of a tunnel in the right
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(A)
(B)
Figure 5.12 Structural landforms at the margins of glaciers in the dry valleys. A) Moraines at
the left margin of Suess Glacier form multiple ice-cored ridges up to 10 m high. Note the
textural contrast between the small moraines in the foreground which are the product of
deformation of proglacial fluvial sediments and the coarser, larger moraines produced by
subglacial erosion. B) Ice-cored moraine at the margin of Wright Lower Glacier. Note the
sedimentary stratification in the moraine. At least part of the moraine has been formed by
deformation of the adjacent delta. The ice cliff is 18 m high.
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ICE-MARGINAL TERRESTRIAL LANDSYSTEMS: POLAR CONTINENTAL GLACIER MARGINS
side of the glacier 100 m upstream of the moraine has demonstrated that the sediment blocks
that feed the moraine have been entrained at least 100 m upstream of the glacier terminus
(Fitzsimons et al., 1999).
The new evidence shows that there is indeed support for the hypothesis that at least some of
these features are thrust-block moraines formed as ice and debris was accreted and thrust at the
glacier margin. However, the absence of the isotopic signature of ice accretion upstream of the
glacier terminus suggests that a large proportion of the debris in the moraines has been
entrained subglacially. The mechanism for erosion and detachment of the blocks of sediment
are currently unknown. The evidence for subglacial entrainment casts doubt on whether the
moraines are thrust-block moraines (sensu Kalin, 1971), which are features that are formed in a
proglacial position as the foreland of a glacier is deformed.
5.5.2.3 Glacifluvial and Glacilacustrine Landforms
Outwash surfaces are absent from the proglacial areas of most glaciers in the McMurdo dry
valleys. Their absence is a consequence of the low production of meltwater, low ephemeral
discharges in streams, and low debris concentrations in and on the glaciers. The largest stream
in the dry valleys and Antarctica is the Onyx River which flows from the Wright Lower Glacier
and Lake Brownworth into Lake Vanda. Lake Vanda like many other lakes in the dry valleys
has no outlet to the sea and water losses occur through sublimation and evaporation. Most lakes
have a 4–6 m-thick perennial ice cover although some are frozen to their beds. The majority of
these lakes receive little sediment from the glaciers again because of the low production of
meltwater, low ephemeral discharges in streams, and low debris concentrations. Even the lakes
in contact with glacier margins (Figs 5.13 and 5.14) are not strongly influenced by the presence
of the glaciers. Divers and remotely operated vehicles operating beneath the ice cover of the
lakes have revealed clear water and a lake bed covered with algae up to the contact with the
glacier cliff.
5.5.2 4 Late Pleistocene Landforms and Sediments
During the Late Pleistocene the configuration of the glaciers in the McMurdo oasis was
substantially different from the glacial systems described above. The alpine and piedmont
glaciers are thought to have receded because their precipitation source was greatly diminished
by the presence of a much larger Ross Ice Shelf. In the Taylor Valley, lacustrine strandlines and
lacustrine deltas provide evidence of a large glacial lake that occupied much of the valley (Fig.
5.13). Glacial Lake Washburn is thought to have formed as the Ross Ice Shelf thickened and
advanced into the valley. The grounded ice shelf deposited the younger Ross Sea Drift (Stuiver
et al., 1981), which extends westward into the valley as far as Canada Glacier (Fig. 5.14). The
drift consists of numerous eskers 1–5 m high and up to 2 km long, and numerous small
moraines which drape the eskers in a washboard-like structure (Fig. 5.14). The eskers and
moraines are overlain by marine sediments and lacustrine deltas.
After withdrawal of the Ross Ice Shelf and the draining of Glacial Lake Washburn, alpine
glaciers such as Canada Glacier advanced and reached their maximum positions on top of and
abutting the Ross Sea Drift (Fig. 5.14). The relationship between the alpine glaciers and the
Ross Sea Drift suggests the alpine and piedmont glaciers fluctuate out of phase with the
grounded ice in McMurdo Sound.
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Figure 5.13 Aerial photograph of Taylor Glacier (TG), an outlet adjacent to the perennially
frozen Lake Bonny. The adjacent Rhone Glacier (RG) is a small alpine glacier fringed by an
older latero-terminal moraine (RM). Numerous strandlines (S) from Glacial Lake Washburn are
evident on both sides of the valley. Several perched deltas (RD) deposited by streams that
flowed from Rhone Glacier are evident.
5.6 TOWARD A DEPOSITIONAL MODEL
Our knowledge of polar continental landform and sediment assemblages is incomplete.
Consequently a comprehensive depositional model cannot yet be assembled. However, several
elements of a model can be identified.
1. Relatively low volumes of sediment are produced by polar glaciers. Consequently landform
and sediment assemblages have modest volumes and the preservation potential of icecontact landforms is low.
2. A variety of constructional moraines form at stable ice–margins. The most common
constructional landforms are ice-contact fans and screes, which form adjacent to steep ice
margins and inner moraines (sensu Hooke, 1973a), which form where basal debris crops
out on the glacier surface.
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ICE-MARGINAL TERRESTRIAL LANDSYSTEMS: POLAR CONTINENTAL GLACIER MARGINS
Figure 5.14 Aerial photograph of Canada Glacier showing a prominent end moraine loop
(CM) and small moraines (RM) and eskers (RE) of the Ross Sea Drift deposited as the Ross Ice
Shelf grounded and advanced into the Taylor Valley.
3. Thrust-block moraines and push moraines form where cold ice contacts saturated unfrozen
sediment.
4. In some circumstances dry-based glaciers appear to be capable of bed deformation and
production of structural landforms
5. Glacifluvial landforms are generally poorly developed elements of the depositional
landscape.
Three high-order controls on the nature of polar landform and sediment assemblages are glacier
thermal regime, climate of the terminus area and the topography of the landscape. Comparison
of landform and sediment assemblages in polar maritime environments, such as Vestfold Hills,
with polar continental environments, such as the McMurdo dry valleys, suggests that the
availability of meltwater is the primary control on depositional processes in ice-marginal
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landscapes. If the summer is sufficiently warm and/or long enough for moderate quantities of
meltwater production, glacial deposits are strongly influenced by remobilization after release
from the ice. Given the critical role of meltwater, the wide climatic variability within polar
environments, and the realization that the elements of the model summarized above are not
inherently different from many other glacial environments, it seems that subdivision of
landform and sediment assemblages based on glacier thermal regime is unsatisfactory.
This review points to a striking gap in our knowledge of polar landform and sediment
assemblages. The gap in our knowledge concerns subglacial processes and resultant landformsediment assemblages. Very little is known about polar subglacial landform and sediment
assemblages because subglacial landscape elements such as streamlined forms or eskers are not
preserved in the land areas that fringe polar ice masses. The main exception is the special case of
Taylor Valley where subglacial landforms and sediments associated with an expanded and
grounded Ross Ice Shelf are preserved. It appears that the absence of subglacial landform and
sediment assemblages could be due to two factors. First, it is likely that subglacial landscapes are
eliminated or modified because of the destructive thermal transition from warm-based to coldbased marginal areas and, second, it is possible that we do not yet recognize the sedimentary
imprint of subglacial processes, particularly those associated with cold-based ice, which could be
quite subtle. The main prospects for improving our understanding of polar subglacial
landscapes are direct observations and measurements of basal processes under thick ice where
the bed is at pressure melting point (e.g. Engelhardt and Kamb, 1998) and where glaciers are
thin and dry-based (e.g. Fitzsimons et al., 1999).
Acknowledgements
This work was supported by the Australian Antarctic Science Advisory Committee and the
Marsden Fund (New Zealand). Logistical support was provided by the Australian Antarctic
Division and Antarctica New Zealand. I thank Damian Gore, Massimo Gasparon, Roland
Payne, Marcus Vandergoes, Regi Lorrain, Sarah Mager and Paul Sirota for assistance in the
field, Sarah Mager and Dr C. O’Cofaigh for critical comments on the text, and Bill Mooney for
drawing the diagrams.