Downwasting of the Tasman Glacier, South Island, New Zealand

New Zealand Journal of Geology and Geophysics, 1995, Vol.38: 1-16
0028-8306/95/3801-0001 $2.50/0 © The Royal Society of New Zealand 1995
Downwasting of the Tasman Glacier, South Island, New Zealand:
changes in the terminus region between 1971 and 1993
MANFRED P. HOCHSTEIN
DAVID CLARIDGE*
Department of Geology
University of Auckland
Private Bag 92019
Auckland, New Zealand
STUART A. HENRYS
ALEX PYNE
Research School of Earth Sciences
Victoria University of Wellington
P.O. Box 600
Wellington, New Zealand
DAVID C. NOBES
STEPHEN F. LEARY
Department of Geology
University of Canterbury
Private Bag 4800
Christchurch, New Zealand
* Present address: Clear Communications Ltd., Private Bag,
Auckland, New Zealand.
Abstract Downwasting has altered the morphology of
the terminus region of the Tasman Glacier between 1971
and 1993. Rapid melting began in the late 1960s in a few
isolated melt ponds in the centre and in a small elongated
lakelet at the eastern lateral moraine. These ponds and lakes
grew rapidly in size during the 1970s and coalesced to form
a large melt lake by about 1990. This melting has led to a
disintegration of the entire terminus region, and now occurs
as far as 3 km upstream from the old terminus. The main
front of the glacier has retreated c. 1.5 km since 1982. The
breaking up of the glacier has been accelerated by the onset
of iceberg calving—a process which probably started in
1991. The icebergs can have volumes of several millions of
cubic metres before they break up into smaller ice masses
that melt slowly during the summer. A temperature survey
has shown that the melt lake is almost isothermal (0.30.5°C). A poorly understood convection mechanism prevents
suspended silt from settling and causes the uniform grey
colour of the lake (here called "Tasman Lake").
Gravity surveys in 1971/72 and in 1982 revealed that
the average thickness of the glacier was between 150 and
200 m over the large (almost 2 km2) area now occupied by
the melt lake. The bottom level of the glacier was close to
600 m a.s.l.; this level has been confirmed by recent radar
094024
Received 13 June 1994; accepted 29 September 1994
soundings and bathymetric surveys. The present lake level
stands at 727 m a.s.l. The surveys demonstrate how the
terminus region of the largest New Zealand glacier has
disintegrated over the past 22 years.
Keywords temperate glacier; downwasting; glacier breakup; bathymetry; isothermal lake; calving; icebergs; ice
thickness; gravity surveys; radar soundings
INTRODUCTION
The Tasman Glacier in the Mt Cook National Park is the
largest compound valley glacier in New Zealand. Its main
trunk is 28 km long, between 1 and 2 km wide, and covers
an area of c. 55 km2 (220 km2 if tributary glaciers are
included). The glacier descends from an altitude of 2400 m
to c. 730 m near the terminus (Fig. 1). Geodetic measurements have shown that the level of the Tasman Glacier has
decreased ever since the glacier was surveyed for the first
time in 1883 (v. Lendenfeld 1884) and in 1890 (Broderick
1891). The decrease is well documented for profiles crossing
the glacier near the old Ball Hut (Harper 1934; Skinner
1964). In 1958, mass losses by ablation dominated any
gains by accumulation (Goldthwait & McKellar 1962). Later
detailed studies by Anderton (1975) showed that net
accumulation occurred in 1972 only above 1600 m altitude
whereas net ablation could be observed over the remaining
part, covering c. 70% of the total length of the glacier.
Continuous decrease in the level of a glacier in the net
ablation zone is referred to as downwasting.
Both net ablation and melting at the bottom of a glacier
contribute to downwasting, which is controlled by longterm climatic changes (Posamentier 1977). To obtain
information about the downwasting process, several surveys
of the Tasman Glacier have been made since 1971. As
downwasting leads to a reduction in cross-section of the
glacier, the studies included measurement of ice thickness
using gravity and seismic surveys. The first survey in 1971/
72 was undertaken by staff of Geophysics Division,
Department of Scientific and Industrial Research (DSIR)
Wellington, establishing baseline monitoring sections across
the glacier near the old Malte Brun Hut, the old Ball Hut,
and near the terminus (Hochstein & Broadbent 1971). Results
from the 1971/72 survey were published by Broadbent
(1974), who found that the total glacier volume was of the
order of 1 x 1010 m3. The Ball Hut and terminus sections
were resurveyed in 1982 by a University of Auckland team
(Claridge 1983). The same sections were again surveyed
during the summer of 1993 by a group from the Universities
of Auckland and Canterbury (Christchurch) and Victoria
University (Wellington). These surveys showed that
considerable changes have occurred during the past 20
years, the most impressive of these having taken place in
the terminus region where some catastrophic melting has
New Zealand Journal of Geology and Geophysics, 1995, Vol. 3
North Island,
NEW
ZEALAND,
2
3
4
5
LEGEND
Glacier boundary
Catchment boundary
Ablation moraine on
glacier ice
Lateral moraine
AREA SHOWN
IN FIGS. 2 - 4 .
_70
o
Outwash gravels
Regional gravity field
(mgal)
Reduced Bouguer anomaly
(mgal) on outcropping rocks
Contour Interval on glacier is in metres
Fig. 1 Tasman Glacier and tributary glaciers in 1972 (based on a modified map by Anderton 1975). The area covered by Fig. 2 4 is boxed. Also shown is the regional gravity field at stations on outcrops where the gravitational effect of the masses in the
valleys has been reduced (modified after Claridge 1983).
Hochstein et al.—Downwasting of Tasman Glacier
occurred during the last decade on a scale not previously
reported for temperate glaciers in the Southern Hemisphere.
Although downwasting and terminal retreat have been
observed from about 1935 onwards at many other, smaller
valley glaciers in the South Island of New Zealand (Suggate
1950; Harrington 1952; Odell 1960; Sara 1968; Gellatly
1985; Bishop & Forsyth 1988), none of these changes has
been as great as those caused by rapid melting of the Tasman
Glacier in the terminus region since 1982. These changes
are described in more detail in this paper. An analysis of
downwasting along the Ball Hut and Malte Brun Hut sections
will be presented separately.
THE TERMINUS OF THE TASMAN GLACIER
(1883-1964)
The Tasman Glacier was mapped in detail in 1883 by v.
Lendenfeld (1884); his 1:80 000 map contains elevation
contours of the glacier based on numerous altimeter traverses
tied into trig stations of known height. The lower reaches of
the glacier and the terminus were also mapped at a 1:31 680
scale by Broderick in 1890 and 1906. If one compares the
older maps with the first photogrammetric map of the area
(New Zealand Department of Lands and Survey 1972, based
on 1964 air photos), one finds that the position of the
glacier front did not change between 1890 and 1964, although
d minor advance between 15 and 45 m might have occurred
between 1890 and 1906 (Broderick 1906). Some changes,
however, occurred in front of the terminal moraine. Up to
1883, the "mouth" of the glacier, called "Gletschertor" by
v. Lendenfeld, was near the southeastern corner of the end
moraine, that is, in a position similar to that occupied by a
smaller meltstream in 1971/72 (Fig. 2). Between 1883 and
1890 the mouth changed to the southwestern corner
(Broderick 1891). This new meltstream became the "Tasman
River".
The maps by v. Lendenfeld (1884) and Broderick (1891)
show that, in the terminus region, the glacier was completely
covered by ablation till; clear ice was visible only at the
mouth and c. 6 km upstream from the terminal moraine. An
earlier sketch map of the glacier (v. Haast 1866) was found
to be unreliable (Broderick 1897) and did not allow an
assessment of whether movement had affected the terminus
between 1866 and 1890. Stations in the terminal region
with elevations measured by Broderick in 1890 are shown
in Fig. 2; the data listed indicate that the elevation of the
glacier was between 870 and 890 m along an east-west
profile c. 2.5 km upstream from the terminus. The 900 m
contour across the glacier in v. Lendenfeld's map lies
c. 500 m upstream from Broderick's east-west profile; the
800 m contour of the 1883 survey occurred c. 1000 m
downstream of Broderick's profile. However, the contours
in v. Lendenfeld's map are based on only a few barometric
heights and field sketches; the map has been reproduced by
Gellatly (1985).
Little is known about changes in the terminus region
between 1890 and 1971. Harper (1934) remarked that, in
1890, the level of the glacier was at the same height as the
top of the eastern lateral moraine across the Murchison
Valley. In December 1933, Harper found that the glacier
was c. 20 m below the top of the lateral moraine in the same
area. Both Harper (1934) and Rose (1937) commented that
downwasting of the Tasman and the nearby Murchison
Glaciers might have started as late as 1913/14. Some
monitoring of the terminus using photographs from fixed
centres was begun in 1957 by Ian McKellar, who detected
no significant changes in the terminus up to 1971 (I.
McKellar pers. comm. March 1972). These studies were
summarised by Goldthwait & McKellar (1962) when they
stated: "Downwasting at the terminal has been of the order
of 50 m in this period [i.e. from 1890 to 1960], and yet
comparison of the present day situation with old surveys
and photographs shows that there has been little horizontal
movement in the position of the terminal, and the ice front
occupies much the same position as it did in 1900."
It has been inferred that all the New Zealand glaciers
advanced to some extent during the period 1885-95
(Burrows & Greenland 1979) during an 80 year climatological period with lower mean annual temperatures lasting
from about 1865 to 1945 (Salinger et al. 1993). The lowest
temperatures during that period occurred around 1900.
Significant warming occurred from about 1945 onwards
(Salinger et al. 1993) and continues at the present time. It
can be inferred that the Tasman Glacier reached a maximum
level along its entire course during the period 1890-1914,
and that its volume has decreased continuously since that
time.
FIRST SURVEY (1971/72) (Fig. 2)
During the first survey of the terminus region from
November 1971 to July 1972, 21 gravity stations were
occupied along a transverse profile A-A' and an axial
profile B - B ' (Fig. 2). All stations were surveyed by
tacheometry (mean error in height ±0.3 m) and tied to a
base station (entrance to a bulldozer shed whose foundation
still exists today). The elevation of the base was not known
at that time, and the observed gravity anomalies were
interpreted in terms of "modified" Bouguer anomalies
assuming that the effect of any regional field could be
neglected. For logistic reasons, seismic measurements could
not be made on the glacier close to the terminus at that time.
However, seismic measurements over the outwash gravels
showed that there was no concealed old ice body in front of
the terminus {vp of saturated, presumably compacted, gravels
was 2.8 km/s). Resistivity soundings along the seismic line
confirmed this (Broadbent 1974). The seismic surveys were
made along a 1.9 km long profile which started c. 400 m
south from the base and ended near point B' in Fig. 2. The
studies showed that the outwash gravels are c. 400 m thick
in the centre of the valley. Assuming a similar structure for
the infill below the glacier, a maximum ice thickness of
220 m was obtained by Broadbent (1974) at the centre of
line A-A', as indicated by a maximum residual anomaly of
-9.5 mgal (-95 x 10"6 N/kg).
In 1981, the elevation of the base station became available
(720.8 m). All survey data from the 1971/72 survey could
therefore be expressed in terms of absolute heights.
Comparison of these data with contours in the 1972
topographic map (based on 1964 photos) showed that the
level of the glacier along line A-A' had probably decreased
between 1964 and 1971 on average by c. 6 m. The elevation
contours were slightly adjusted, and the smoothed contours
shown in Fig. 2 reflect approximately the 1971 surface of
the glacier in the terminal region. The 1971/72 survey data
also showed that the level of one melt pond near line A-A'
New Zealand Journal of Geology and Geophysics, 1995, Vol. .
Survey station 1890 (elevation in metres)
Younger moraines exposed in 1972
Melt pond on glacier (grey water)
Exposed ice cliff
Pond level (a.s.l.)
Perched lake (blue-green water)
Ball Hut Road
Gravity station (July 1972)
Smoothed elevation contour covering glacier
mouth1 of glacier
2282
2283
Fig. 2 Topography of the terminus region of the Tasman Glacier during the summer of 1971/72. The glacier contours have been
smoothed using data from the S79 Lands and Survey map (1972) based on 1964/65 air photos. Tacheometric survey data (1971/72)
were used to reconstruct the 1971/72 surface. The grid co-ordinates refer to those of the DOSLI Infomap 260-H36 (1992). The position
of the melt holes was taken from a 1971 air photo reproduced by Kirkbride (1993). Spot heights on the glacier and the moraine were
taken from Broderick's map 60.T. (1891); B stands for barometric height; all others were surveyed using a theodolite.
Hochstein et al.—Downwasting of Tasman Glacier
was 730 ± 1 m, similar to the level of the mouth of the
Tasman River. The melt ponds were much smaller craters
in 1964 (New Zealand Department of Lands and Survey
photo 3724/24), when they could have been described as
large "moulins" although there was little water in the craters
at that time. By 1971 the craters had grown in size and were
filled with grey-coloured meltwater (i.e. they had become
melt ponds). Air photos taken in April 1971 (Kirkbride
1993) indicate an average diameter of 100 m for the ponds.
Their size as shown in Fig. 2 is only approximate. A smaller
meltstream had breached the end moraine in the southeastern
corner (see Fig. 2).
SECOND SURVEY (1982) (Fig. 3)
The morphology of the terminus region had changed when
the second survey was made during the summer of 1981/82.
The melt ponds had increased in size to 0.56 x 106 m2 (see
Fig. 3). Over 70 gravity stations were established on the
glacier and were tied in to the base station (only half of
these stations are shown on Fig. 3). Allowing for intermediate
stations and local assessment of surrounding terrain, there
were sufficient data available to construct the elevation
contours of the terminal region shown in Fig. 3 (mean error
in station height was still ±0.3 m). The surface level of most
of the melt lakes lying south of line A-A' was surveyed,
and these levels were almost the same as that of the Tasman
River outlet, namely 728 ± 1 m. These melt lakes were
therefore hydraulically connected with the outlet by a
network of fractures and fissures. All lakes south of A-A'
were filled with grey-coloured meltwater containing
suspended silt. Two small lakes to the north of line A-A', at
higher elevation, contained grey-green water, and were
assumed to be perched shallow lakes; two other perched
lakes occur within the western lateral moraine ("Blue Lakes"
in Fig. 2).
A bathymetric survey using a small boat and a 100 m
leadline showed that the water depth of most of the melt
lakes was between 20 and 50 m. In one lake, depths >100 m
were found at two sites c. 80 m apart, indicating that vertical
channels probably extended to the bottom of the glacier. In
the summer, the surface water temperature near the ice
cliffs was as high as 10°C, decreasing to 3°C at 0.3 m depth.
The warmer surface layer enhanced melting, resulting in 13 m deep horizontal cuts (grooves) causing slumping of the
overlapping ice cliffs (Claridge 1983). Comparison of glacier
levels along lines A-A' and B-B' showed that downwasting
between 1972 and 1982 had been irregular. Consistent
downwasting of 12 ± 3 m is indicated for the axial profile
B-B', whereas decreases between 5 and 45 m are indicated
for the transverse profile A-A'.
As the effect of terrain had been carefully assessed
using an accurate method by Olivier & Simard (1981), the
1982 gravity data were of high quality. For interpretation,
the gravitational effect of subglacial till was assumed to be
similar to that in Broadbent's model. In the centre of profile
A-A', the ice thickness was found to be c. 160 m (i.e.
bottom level of glacier close to 600 m a.s.l.). A 3-D model
showed that thick ice (>100 m) occurred c. 500 m upstream
of the terminus. The ice thickness was confirmed by DCresistivity surveys which pointed to a thickness close to
230 m c. 2 km upstream at station A110 and 100 m (max.
thickness) at station A104 c. 0.3 km upstream from the
terminus (Fig. 3). The ice resistivities were about 11 and 5.5
x 106 £2m, respectively (i.e. falling in the lower range of
those obtained on other temperate glaciers; Rothlisberger &
Vogtli 1967); a significant anisotropy (25% near Al 10 and
10% near A104) was observed, with the resistivity in flow
direction attaining a minimum value.
Because the whole glacier is covered in the terminus
region by a 1-2 m thick layer of ablation till, which is fairly
coarse, it was impossible to define accurately the ice edge at
the terminus. DC-resistivity traverses with small spacings
(100 m) were used with success to map the ice-end moraine
contact (Claridge 1983). Attempts to map the glacier bottom
with seismic methods, using a 12-channel, portable SIErefraction unit, failed; the measurements only provided the
trivial result that ice was coherent beneath the till at four
profiles (vp of glacier ice = 3650-3700 m/s).
THIRD SURVEY (1993) (Fig. 4)
In February 1986, a photogrammetric survey of the Mt
Cook area was repeated by the New Zealand Lands and
Survey Department; the air photos indicate that the surface
of the melt lakes had increased and covered 1.07 x 106 m2,
a twofold increase since 1982. The new Mt Cook topographic
map was published just before the third survey (New Zealand
DOSLI 1992). It was now apparent that some catastrophic
melting was occurring in the terminus region. A glaciological
study of the terminus region had been made in the meantime
(Kirkbride 1989); ice flow measurements by Kirkbride near
the centre of profile A-A' showed that the glacier was still
moving slowly at a rate of 3.3 m/yr at that locality.
Several new methods were introduced in 1993 to assess
the structure of the terminus region, namely the groundpenetrating radar method, and detailed limnological surveys
(water depth, temperature, and water velocity). Standard
tacheometric surveys were used to determine the position
of stations on the glacier (with reference to the 1971 base
station), and a small satellite navigation system (Ensign
GPS) was used to determine positions on the lake (controlled
by observations with a land-based theodolite and compass
bearings).
Radar soundings
Earlier studies had shown that the thickness of ice caps and
thick cold ice can be measured with radar (radio echo)
reflections (Harrison 1973). Because of relatively strong
attentuation of the signal and scattering effects of ice with
englacial debris, the method provides less clear results if
used on temperate glaciers and permafrost (e.g., Goodman
etal. 1975; King etal. 1987). Recent developments, however,
have shown that ground-penetrating radar can be used to
map temperate glaciers and ice bodies (Robinson et al.
1993) and that ice thicknesses of temperate glaciers based
on radio echo-soundings are, in general, accurate to within
5% of the actual ice thickness (Haeberli & Fisch 1984).
For our surveys in February 1993 we used an EKKOIV
ground-penetrating radar system employing 25 MHz
antennas. The equipment was tested first in the terminus
region of the nearby Miiller Glacier, where good reflections
from the bottom of the glacier (max. depth 140 m) were
obtained (Leary 1993). The survey was then shifted to the
terminus region of the Tasman Glacier. For logistic reasons
we could only take measurements along one 500 m long
New Zealand Journal of Geology and Geophysics, 1995, Vol.
2282
2283
Fig. 3 Topography of the terminus region of the Tasman Glacier at the end of the 1982 summer. The glacier contours are based on ;i
tacheometric survey (Claridge 1983). The outlines of the larger melt lakes were constructed from air photos and ground surveys.
Hochstein et al.—Downwasting of Tasman Glacier
Bathymetric contour (in metres) of
Tasman Lake
Boat track with continuous soundings
Spot sounding of water depth
Bathymetric station (temperature and
velocity logs) see Fig. 7.
Station with GPS fix
Gravity station (February 1993)
Radar sounding profile on ice
0
SCALE
500 m
TS Radar
Profile
\-_>
2282
2283
Fig. 4 Topography and bathymetry of the terminus region of the Tasman Glacier at the end of the 1993 summer. The glacier contours
have been smoothed using data from the H36 DOSLI map (1992) based on 1986 air photos. Tacheometric survey data (1993) were used
lo reconstruct the 1993 surface. The bathymetric survey lines are controlled by satellite fixes; details of the margins of Tasman Lake are
based on 1993 air photos.
New Zealand Journal of Geology and Geophysics, 1995, Vol. 38
profile near the western lateral moraine (profile TT in Fig.
4) and another 300 m long north-south trending profile (TS
in Fig. 4) over the glacier snout. Separations between the
antennas were typically 2 m, although a few common
midpoint (CMP) profiles and "step-out" profiles were also
observed. Analysis of the traveltime of direct and reflected
radar waves showed that the speed of these signals in ice is
c. 0.135 m/ns at shallow depth; it increases to 0.16 m/ns at
greater depths.
A processed section of the TT profile is shown in Fig. 5
(taken from Leary 1993). This section shows that old ice
underlies the whole profile; the observed ice thickness has
been incorporated in the cross-section of the A-A' line
shown in Fig. 8. Since the ablation till in the western half of
the TT profile is covered by sparse vegetation (moss,
lichens), it had been assumed in the past that there was no
thick ice beneath this area. Interpretation of gravity anomalies
over the western part of profile A-A' by Broadbent (1974)
and Claridge (1983) were based on this assumption. Our
radar soundings showed that the assumption had to be
revised. The sounding results of the TS profile were less
clear although the bottom of the glacier can be recognised
along certain segments of the profile. The section of the
snout shown in Fig. 9 was taken from a processed section of
Leary (1993). A summary of our radar soundings has been
presented recently (Nobes et al. 1994).
Tasman Lake surveys
A reconnaissance survey in February 1993, using a small
boat and a 100 m long leadline, had shown that the water
depths of the greater part of the melt lake were >100 m. It
was therefore likely that the glacier had melted down to its
base. A separate bathymetric survey was mounted in April
1993 to investigate the structure of the large melt lake
(Tasman Lake*). Since it was possible that bottom melting
was enhanced by a large subglacial flow of meltwater, an
Inter-Ocean current meter was used to measure temperature
and horizontal flow at various depths in the lake.
Water depths were measured using a pair of 28 kHz and
200 kHz acoustic transducers (CRAME 800 Echo Sounder).
A depth sounding section for the bathymetric profile 4,
which is close to line A-A' in Fig. 4, is shown in Fig. 6A;
another sounding section (bathymetry profile 11, close to
line B-B') is shown in Fig. 6B. It can be seen from Fig. 6B
that the ice cliffs at the northern edge of the lake are almost
vertical, and dead ice at the southern end of the profile dips
smoothly c. 35° beneath the lake. At the western edge of the
profile in Fig. 6A, the ice slope dips c. 30° to the east.
Leadline soundings showed that in certain depressions a
thin layer ( e l m ) of silt sludge had already accumulated at
the bottom of the lake.
The bathymetric profiles and spot soundings were used
to compile a bathymetric map of the terminal lake (Fig. 4);
the outlines of the lake were taken from a vertical photo
(May 1993; L. Homer pers. comm.). Integration of the data
in Fig. 4 showed that in April 1993 the Tasman Lake
occupied an area of c. 1.95 x 106 m2. Analysis of the
tacheometric survey showed that the lake level was 727 ± 1
m a.s.l. from February to April 1993.
*New name proposed to the New Zealand Geographic Board.
The temperature-depth profiles taken at six bathymetric
stations are shown in Fig. 7; the temperatures displayed are
the mean of down and up logging runs at 1 m intervals. The
data indicate that the lake is almost isothermal, wild
temperatures between 0.3 and 0.5°C (the instrument;A
accuracy was ±0.05°C). A 0.1 °C colder layer is indicate;!
for the bottom section at stations 3 and 5, which were
closest to the northern vertical ice cliffs (Fig. 4). Measure ments showed that the direction of the horizontal component
of flow in the lake was essentially random; the speed varied
at the lake bottom between 0.05 m/s at station 5 to 0.25 nVs
at station 3. Similar values were observed at the other
stations. The vertical component could not be measured
with the equipment.
An oblique air photo taken in August 1993 (R. Bellringer
pers. comm.) has shown that the melt lake froze over during
winter except for three large patches of open water up to
100 m in extent along the northern ice cliffs between stations
3 and 5. It is therefore likely that mixing of the lake is
driven by subglacial meltwater which enters the lake near
the northern ice cliffs. Since the density of water is a
maximum at 4°C, the slightly colder temperature of the
bottom layer would result in minor buoyancy forces causiri ^
secular convection.
SECTIONS OF THE GLACIER (TERMINUS
REGION)
By April 1993 we had collected sufficient data to reconstrm t
the actual shape of the glacier in the terminus region. The
shape of the bottom of the glacier is constrained by the
radar soundings and the recent bathymetric survey. Putting
all interpretation models together in the form of an integrated
model produced the sections shown in Fig. 8 (line A-A )
and Fig. 9 (line B-B'). To obtain representative residual
Bouguer anomalies, we subtracted the regional field showa
in Fig. 1 (modified after Claridge 1983) from the observed
Bouguer anomaly values. This field represents "absolute"
Bouguer anomalies at stations on solid rocks, allowing for
the "edge effect" of nearby glaciers.
Using all available data, the theoretical gravitational
effect of the glacier (1982 surface) was computed for lines
A-A' and B-B' using the algorithm of Barnett (1976). The
results are shown in the upper half of Fig. 8 and 9. It can be
seen that 2-D models produce theoretical effects (ice only)
that are too small, causing a difference of almost 2 mgal (20
(xN/kg) near the intersection of the two lines (stations A108
and K106). The 2-D model over the "snout" (right-hand
side of section in Fig. 9) also produces anomalies which are
too small. When a 3-D model was used, these inconsistencies
disappeared. The difference between computed 3-D (ice)
anomalies and observed residual anomalies reflects the effect
of low-density gravels and glacial till beneath the glacier.
The two curves in Fig. 8 indicate an asymmetric distribution
of these deposits for the cross-sectional profile A-A', wiih
the thickest deposits occurring beneath the western half of
the section. The upstream divergence of the two anomaly
curves in Fig. 9 points to an increasing thickness or
decreasing density of subglacial deposits in that direction; it
might also reflect poor control of the regional field in the
terminus region. Further studies of the regional field are
required before the structure of the subglacial deposits
beneath the terminus region can be presented.
T
c
Position (m)
250
300
350
400
450
500
538
n_
ft
5"
-250
re
500
1
^-750
1000
fro
-1250
1500
o
-1750
2000
Processed
West
•o- 3
Position (m)
50
I
100
I
9'acier surface
150
200
250
300
350
400
I
East
450
-780
^^
-760
=
E
-740
(O
lodgement till
Lake level
(726.8 m)
y -
:
•
•
•
•
•
—
•
•
•
•
.
:
•
•
.
•
-
•
-
;
.
"
;
•
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»
.
;
•>' - v ' ^ S ; .
'••.
lodgement till
\
N
-720
Tasman glacier
<D
-640
Interpretation
Fig. 5 Processed radar sounding section TT; for locality see Fig. 4. The distance between recording sites is 2 m. The bottom reflection comes from the ice-bottom till interface (max. thickness
110 m).
10
New Zealand Journal of Geology and Geophysics, 1995, Vol. 38
Bathymetry LINE 4 (Tasman Lake), 24.4 1993
14:16.9
I*.: I.
200 Ml,
T
W 0
500
250
750
1250
approx. distance (m)
Fig. 6A Bathymetric sounding profile of Line 4 (W-E) across Tasman Lake (for location see Fig. 4). The profile in the upper half was
recorded with a 200 kHz transducer, and the profile in the lower half with a 28 kHz transducer (copy of original record).
-< Fig. 6B Bathymetric sounding profile of Line 11 (S-N) acn ss
Tasman Lake (copy of original record).
Bathymetry LINE 11 (Tasman Lake), 24.4.1993
15:54.7
15:59.5
16:06.3
RAPID MELTING AND ASSOCIATED
DISINTEGRATION OF THE GLACIER TERMINUS
500
approx. distance (rn)
r
750
1000
Long-term changes in the level of the glacier near the
terminus region can be assessed if the spot heights of
Broderick's survey in 1890 are compared with the smoothed
contours of the glacier in 1971 (Fig. 2). This shows that the
level decreased at an average rate of 1.1 m/yr during that
time. This rate, however, probably contains an opposing
component associated with a short-term increase in crosssectional ice flux and, hence, increase in level during the
period 1885-95 when other New Zealand glaciers advanced.
Significant downwasting probably began only from 1914
onwards (Harper 1934; Rose 1937).
Between 1964 and 1971 the downwasting rate was
probably 0.85 m/yr, reaching a value of 1.2 m/yr between
1972 and 1982 (these rates refer to the area near the
intersection of profiles A-A' and B-B' in Fig. 2). The
surface profiles in Fig. 9 indicate that downwasting at the
terminal moraine was insignificant until about 1972. An
increase in the rate of downwasting from about 1960 onwai Is
has been reported for the lower reaches of another laige
valley glacier in the South Island, the Dart Glacier (Bisri ip
& Forsyth 1988), although downwasting there has stopped
more recently (G. Bishop pers. comm. 1993).
The 1-2 m thick layer of ablation till in the terminus
region shielded the Tasman Glacier from melting by
Hochstein et al.—Downwasting of Tasman Glacier
Fig. 7
Temperature-depth
logging profiles of Tasman Lake
al six bathymetric stations (for
station locality see Fig. 4). The
error bars represent the difference
in temperature between down- and
upward logs. Temperatures are
plotted at 1 m depth intervals;
surface temperature is not shown.
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insolation; glacier downwasting upstream from the terminal
moraine was presumably the result of a continuous decrease
in cross-sectional ice flow. The shape of the 800 m contour
in v. Lendenfeld's map (1884) suggests that there were two
slowly moving ice streams, one on the eastern side, probably
fed by the Upper Tasman Glacier, and a westerly stream fed
mainly by the Hochstetter Glacier (see Fig. 1). Large
sinkholes had developed along the margins of these two ice
streams: v. Lendenfeld (1884) mapped numerous sinkholes
in the lower 4 km of the glacier, and a number of circular
features are also shown in Broderick's map (1891). These
sinkholes, called "Dolinen" by v. Lendenfeld, were
characteristic features of the terminal region, and occur in
all older photos of this region (Kirkbride 1993).
Before 1960, the western ice stream became stagnant,
and sinkholes developed mainly along the margins of the
eastern ice stream (see Fig. 2). After 1960, sinkholes to the
south of line A-A' (Fig. 2) filled with meltwater, the bottom
of these holes having intersected the coherent intraglacial
water table at c. 730 m elevation. Melting was now enhanced
by convection in vertical channels and tubes, as described
by Rothlisberger (1972) and Shreve (1972), and water in
the ponds attained a characteristic grey colour caused by
mixed, suspended silt particles. Other sinkholes further
upstream also filled with meltwater but the water often
remained unmixed, attaining a grey-green colour (perched
water table). Rapid melting of the exposed bare ice (due to
insolation) and bottom melting by convection in vertical
tubes caused the rapid increase in size of these ponds noticed
between 1971 and 1982 (see Fig. 2, 3). Two of these deep
tubes only 80 m apart were detected by our first bathymetric
survey in 1982; the location of one with a water depth
>100 m is shown in Fig. 3. The lateral growth of the melt
ponds was studied in detail by Kirkbride (1993) during
1986/87, who detected pronounced seasonal growth.
Another process probably caused the elongated melt
lakes at the contact between the glacier and eastern lateral
moraine. A 550 m long and c. 40 m wide lake had already
New Zealand Journal of Geology and Geophysics, 1995, Vol. !8
12
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Fig. 8 Upper part: Theoretical and observed residual gravity anomalies along line A-A' (Tasman Glacier); for locality see Fig. 4. The
theoretical anomalies (gravitational effect of ice only) are based on the section of the glacier shown in the lower part and a modified ice
isopach map by Claridge (1983) bounded by the 1982 surface.
Lower part: Cross-section of the Tasman Glacier along line A-A'. The bottom was constructed using results from bathymetric and
radar soundings. The level of the glacier and lateral moraines are from tacheometric survey data.
been formed by the summer of 1971/72 (Fig. 2). It is
possible that melting in this lake was induced by surface
water, which infiltrated from the east, through the lateral
moraine that blocks the Murchison Valley (Fig. 1).
When the rapidly growing, quasi-circular melt lakes in
the western part of the terminus region reached the elongated
melt lakes in the eastern half (about 1988/89, according to
Kirkbride 1993) a new ablation mechanism, namely calving
of icebergs, began to operate. This enhanced the breaking
up process of the glacier. In February 1993 we noticed that
six large icebergs (up to 100 m long) were floating in the
terminal lake. Their position changed from day to day
depending on the prevailing wind direction. One of the six
icebergs, however, did not move—it had become grounded
near the outlet, where the water depth was 50 m (Fig. 4).
Triangulation from land stations showed that the top of this
iceberg stood 4.7 m above the lake level, pointing to a
maximum submerged thickness of 50 m (assuming
hydrostatic equilibrium for the ice body with a density of
0.916 x 103 kg/m3). Its volume could therefore be assessed,
and it was found to be c. 200 000 m3.
Another, much larger ice block at the calving front was
noticed in 1993. This tabular block covered an area of
c. 86 000 m2 and was part of a promontory of the northern
ice cliffs (Fig. 4). A large, well-defined crack, 2^1- m wide,
had appeared by April 1993 and could be entered by boat on
the eastern side. Since the bathymetric survey had shown
that the ice block has vertical sides to the south, and using
the geometry shown in Fig. 9, it could be inferred that the
volume of this ice block was 12.5 x 106 m3 (i.e. about onetenth of the volume of the present lake).
Calving of icebergs is controlled by the buoyancy ratio
(pw/pi)d!/7j, where d = water depth, h = ice thickness, p w and
Pj = density of water and ice, respectively. For the large ice
block at the calving front in the Tasman Lake (see Fig. 9). d
= 125 m and h = 140 m at the southern cliff. Hence, the
buoyancy ratio is about 0.97, and the large ice mass should
be grounded. However, its average density could be slightly
less than 0.916 x 103 kg/m3, as indicated by air bubbles
enclosed in the matrix of smaller icebergs floating in the
lake, and the error in d and h is of the order of a few metres.
Therefore, the buoyancy ratio could be >1, which implies
Hochstein et al.—Downwasting of Tasman Glacier
13
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of Fig. 8 also apply.
that the ice block is afloat although it might touch the
bottom moraine in a few places. We therefore use the term
"iceberg" for this large block of ice. An air photo taken in
August 1993 showed that the large iceberg was still in the
same position as in April 1993. The large crack along the
northern margin had widened to a V-shaped valley, and 5-7
east-west striking fissures with en echelon structure had
developed within this 320 m wide ice block. It is therefore
possible that this iceberg will disintegrate by further calving.
Another important parameter, which reflects melting in
the terminus region, namely the outflow rate of the Tasman
River, is still not yet known. From the drifting speed of a
small boat, and a bathymetric section of the outlet area, we
estimated in February 1982 that the outflow rate was of the
order of 50 m3/s. This flow rate, however, changes with
season and the degree of ongoing melting. The observed
changes in lake surface and in volume of meltwater of the
terminal lakes between 1972 and 1993 are shown in Fig. 10.
These point to an almost linear increase in surface area with
time since 1982.
POSSIBLE NEAR-FUTURE CHANGES
If the disintegration of the terminus region continues at the
present rate, some significant changes are indicated which
are associated with the damming of the lake outlet,
downwasting and slumping of the lateral moraines, and
melting of dead ice beneath the old terminus.
In 1993, when we launched a boat in a small bay that
lies c. 450 m north from the lake outlet (Fig. 4), we noticed
silt terraces standing 1-3 m above the lake level. Obviously,
the lake level had changed during the last few years. The
observation of a grounded iceberg near the outlet might
explain such changes as being the result of the damming
effect of icebergs in that area. Changes in lake level might
also be caused by seasonal meltwater flows.
Melting of the glacier over a distance of almost 3 km has
weakened the support of the eastern lateral moraine (Fig.
4). Enhanced slumping of the moraine can be expected in
this area. Such slumping has already occurred near line AA'. Surface level data show that the top of the moraine has
New Zealand Journal of Geology and Geophysics, 1995, Vol ; 8
14
1968
70
</
74
78
82
Time (year)
90
1992
Fig. 10 Total area of melt lakes and lake volumina for the
period 1972-93. The graph covers all ponds and lakes in the
terminus region (Tasman Glacier) shown in Fig. 2-4.
decreased in level by c. 8 m since 1982 (Fig. 8). After
completion of the study, the Murchison River broke through
the eastern moraine in January 1994 and now discharges
into the Tasman Lake.
The section in Fig. 9 indicates that the terminus of the
Tasman Glacier has retreated by c. 1.3-1.6 km as a result of
rapid melting between 1971 and 1993. The new terminus is
represented by the ice cliffs (calving front) forming the
northern margin of Tasman Lake. When the dead ice near
the old terminus disappears, some additional drainage of
the lake may take place via existing gaps in the old terminal
moraine, which could lead to flood surges. In view of the
possible developments outlined here, further, more detailed
monitoring of the terminal region is desirable. Our old
monitoring schedule with 10-11 year intervals is no longer
suitable for such surveillance.
SUMMARY AND DISCUSSION
Sporadic observations of the terminus region between 1900
and 1972 indicate that the Tasman Glacier decreased in
thickness at a rate of the order of 1 m/yr. More detailed
monitoring began in the summer of 1971/72 and showed
that the terminus region began to disintegrate from 1972
onwards. This was caused by rapid melting which started
from a few melt ponds and an elongated marginal melt lake;
all lakes were hydrologically connected by a network of
intraglacial fissures and channels. The level of the
interconnected ponds and lakes was controlled by the level
of the outlet into the Tasman River; this decreased slightly
from 730 m a.s.l. in 1971/72 to 727 m in 1993.
Downwasting became uneven between 1972 and 1982,
although along an axial strip it was still rather uniform (1.2
m/yr). It is likely that the 1-2 m thick layer of coarse
ablation till that already covered the glacier at lower altitudes
before 1890 reduced ablation. The total surface area of the
melt ponds and lakes has increased at a rate of c. 125 000
m2/yr since 1982. By 1990, these features had coalesced to
form one large, coherent lake (Tasman Lake) which covered
an area of 1.95 km2 in April 1993. A bathymetric survey
showed that the glacier had melted down to its bottom,
which lies near 600 m a.s.l. The glacier's retreat w.is
accelerated in 1991 by iceberg calving.
The thickness of the glacier was assessed during our
earlier surveys from residual gravity anomalies. The 19 2
and 1982 gravity models of an east-west section of tl e
glacier c. 1.5 km upstream from the terminus indicated a
bottom level of 580 m (Broadbent 1974) and 600 m (Claridi e
1983), respectively. Considering the various uncertain tics
in these interpretations, the agreement with the actual level
of 600 m, as found in 1993 from bathymetric and rad.tr
sounding surveys, is remarkably close and confirms that the
glacier thickness can be assessed reliably in steep terrain by
gravity surveys (Kanasewich 1963; Klingele & Kahle 1977).
A detailed structure of the ice-bottom moraine contact,
however, cannot be obtained with this method. Neither
could it be obtained from seismic reflection surveys as the
1-2 m thick debris layer on top of the glacier prevented us
from obtaining good reflection data. An important finding
of the 1993 survey, therefore, was that radar signals can
penetrate the debris layer, and that radar reflections of hie h
resolution can be obtained from the glacier bottom (Nobcs
et al. 1994).
Detailed information about the fine structure of the
bottom comes from bathymetric soundings, which show that
the bottom of the glacier is quite irregular. This is a surprising
result because it was expected that movement of the glacier
would have created a smooth subsurface. Analysis of seism i c
and gravity data near the terminus has shown that a thick
bottom moraine exists and that it consists of compacted
gravels and till that are c. 400 m thick near the terminus.
Another important result of our study is the discovery i if
the anomalous hydrological setting of the terminus region,
where a coherent intraglacial water table was extant in
1972, before smaller melt ponds started to develop. The
hydraulic gradient was small (<1 x 103), pointing lo
anomalously high permeability in the glacier. In other
temperate glaciers, hydraulic gradients are at least 1-2 orders
of magnitude greater (Hantz & Lliboutry 1983), even in the
presence of intraglacial meltwater channels (Rothlisberger
1972). The effective permeability of these glaciers appears
to be at least an order of magnitude lower than that indicated
for the lower Tasman Glacier. We do not know what caused
this high permeability but it can be inferred that it is an
important parameter which controlled subsurface melting
once melt ponds intersected the water table. This caused
meltwater convection, which was maintained by the small
density difference between water in the ponds (max. density
at T = 4°C) and colder water in deeper melt channels i T
close to 0°C). The fact that silt remained in suspension, us
indicated by the uniform grey colour of the water in the
melt ponds, confirms that convection did occur. Melting of
the bare ice exposed in the walls of ponds and lakes wits
enhanced by insolation, but surface melting was minor
compared with melting of the ice below the water table.
Ice calving began in 1991 after most of the ponds and
smaller lakes had coalesced to form one large melt lake
(Tasman Lake). Calving dynamics are poorly understood
and seem to differ between temperate and cold glaciers, and
also between glaciers with grounded and floating termini
(Hughes 1992; Warren 1992). Calving is common where a
fast-moving glacier discharges into the sea (in Greenland
and Alaska, for example), or into deep terminal lakes
(Patagonia). The Tasman Glacier, however, is almost
stagnant at its calving front (<4 m/yr).
Hochstein et al.—Downwasting of Tasman Glacier
15
Although we cannot yet classify the calving dynamics
of the Tasman Glacier, some conditions which facilitate
calving of a stagnant temperate glacier can be described.
Our studies indicate that calving can commence once most
of the ice at the bottom of the terminal lake has melted. For
calving to occur, the buoyancy ratio (p^lp^d/h has to be >1
at the calving front. A coherent intraglacial water table with
a small hydraulic gradient, as occurs in the lower part of the
Tasman Glacier, is probably a prerequisite for calving in
this setting.
Rapid melting, formation of a large terminal lake, and
subsequent calving occur not only at the Tasman Glacier
but have also been reported for smaller glaciers in the Mt
Cook National Park (Kirkbride 1993), although the onset of
rapid melting differed from glacier to glacier. By 1950,
large melt lakes had developed at the foot of the Classen
and Grey Glaciers, c. 22 km to the northwest of the
Murchison Terminus (Fig. 1). The terminus of the Godley
Glacier, which borders the Grey, disintegrated after the
coalescence of several melt lakes in the mid 1970s. Calving
has been reported since 1961 for the Classen Glacier; the
Godley Glacier has retreated, mainly by calving, since 1974
(Kirkbride 1993). The terminus of the Tasman Glacier had
disintegrated by 1989, with calving starting in 1991.
Glaciological Section and the Central Laboratories of the New
Zealand Ministry of Works. The 1982 and 1993 surveys were
supported by grants from the University of Auckland Grants
Committee. Members and rangers of the Mt Cook National Park
Board assisted during each survey. Sierra software at VUW was
used to process the 1993 radar data. W. Haeberli (ETH Zurich)
and C. R. Warren (University of Edinburgh) read a draft of the
manuscript and provided valuable comments.
At the present time, accelerated growth of preglacial
lakes also occurs at three other glaciers in the vicinity of the
Tasman Glacier, namely the Murchison (Fig. 1), Hooker,
and Miiller Glaciers (c. 7 km to the west of the Tasman
terminus). During visits in 1982 we found that several melt
ponds, similar in size to those seen in 1972 on the lower
Tasman Glacier, had just formed. Most of the ponds are
shown in small maps by Gellatly (1985). By 1993, the
ponds had become terminal melt lakes, which are c. 0.7 km
long at the foot of the Hooker and Miiller Glaciers. Calving
has not been observed here yet.
The ongoing, rapid distintegration of the whole terminus
region of the Tasman Glacier is ultimately a result of climatological changes. These began in about 1945, when mean
annual temperatures began to rise in New Zealand (Salinger
et al. 1993). However, the mean temperature is only one of
the many parameters affecting the behaviour of glaciers
(Hay & Fitzharris 1988) because timing and the nature of
glacier retreat are also modified by other, non-climatic
factors (Warren 1992). The staggered development of melt
lakes and calving of glaciers in the Mt Cook National Park
is good evidence for the influence of non-climatic factors.
Broderick, T. N. 1891: Map of the Tasman and Murchison Glaciers
(1 inch to 40 chains). (Topo 60.T.). Held by Department
of Survey and Land Information (DOSLI), Christchurch.
A CKNO WLEDGMENTS
The surveys described were made possible by the assistance of
many helpers, volunteers, students, and professionals. In
appreciation of their work, the names of those mentioned in our
field books are listed:
1971/72 party: P. W. Anderton (Joint Leader), R. A. Atkins, H.
Bibby, M. Broadbent, L. Carrington, F. Davey, B. Hochstein, M.
P. Hochstein (Joint Leader), D. Innes, R. Jenkins, C. O'Reilly,
arid R. Williams.
1982 party: R. Bellringer, G. Bulte, G. Caldwell, D. Claridge, S.
Davidge, S. A. Henrys, M. P. Hochstein (Leader), S. Rawson, D.
J Robertson, H.P. Schmidt, and A. Sutherland.
1993party: R. Bellringer, S. A. Henrys, M. P. Hochstein (Leader),
S. F. Leary, D. C. Nobes, A. Pyne, and M. Watson.
The 1971/72 survey was sponsored by Geophysics Division
(DSIR) and supported by Antarctic Division (DSIR), the
REFERENCES
Anderton, P. W. 1975: Tasman Glacier 1971-73. Hydrological
research: annual report 33. Wellington, Ministry of Works
and Development. 28 p.
Barnett, C. T. 1976: Theoretical modelling of the magnetic and
gravitational fields of an arbitrarily-shaped threedimensional body. Geophysics 41: 1353-1364.
Bishop, G.; Forsyth, J. 1988: Vanishing ice. An introduction to
glaciers based on a study of the Dart Glacier. John Mclndoe
and New Zealand Geological Survey, Department of
Scientific and Industrial Research. 56 p.
Broadbent, M. 1974: Seismic and gravity surveys on the Tasman
Glacier 1971-72. Geophysics Division report 91.
Wellington, Department of Scientific and Industrial
Research. 43 p.
Broderick, T.N.I 897: Letter to Lands and Survey Chief Surveyor.
Letterbook S43 86/21. Held by Department of Survey and
Land Information (DOSLI), Christchurch.
Broderick, T. N. 1906: Glacier movements. Journal of the House
of Representatives of New Zealand, session II, volume I,
C-1A, appendix 2: 16—17.
Burrows, C. J.; Greenland, D. E. 1979: An analysis of the evidence
for climatic change in New Zealand in the last thousand
years: evidence from diverse natural phenomena and from
instrumental records. Journal of the Royal Society of New
Zealand 9: 321-373.
Claridge, D. 1983: A geophysical study of the termini of the
Mount Cook National Park glaciers. Unpublished M.Sc.
thesis, lodged in the Library, University of Auckland.
Gellatly, A. F. 1985: Historical records of glacier fluctuations in
Mt. Cook National Park, New Zealand: a century of change.
The geophysical journal 151: 86-99.
Goldthwait, R. P.; McKellar, I. C. 1962: New Zealand glaciology.
In: Antarctic research. American Geophysical Union
publication 1036, geophysical monograph 7: 209-316.
Goodman, R. H.; Clarke, G. H. C ; Jarvis, G. T.; Collins, S. C ;
Metcalfe, R. 1975: Radio soundings on Trapridge Glacier,
Yukon Territory, Canada. Journal of glaciology 14:
79-84.
v. Haast, I. 1866: The glaciers on Mount Cook and Mount Tasman
(sketchmap 1 inch : 4 miles). Wellington, Alexander
Turnbull Library.
Haeberli, W.; Fisch, W. 1984: Electrical resistivity soundings of
glacier beds: a test study on Grubengletscher, Wallis,
Swiss Alps. Journal of glaciology 30: 373-376.
Hantz, D.; Lliboutry, L. 1983: Waterways, ice permeability, and
water pressures at Glacier D'Argentiere, French Alps.
Journal of glaciology 29: 227-239.
Harper, A. R. 1934: Glacier retreat. New Zealand alpine journal
5:68-71.
Harrington, H. J. 1952: Glacier wasting and retreat in the Southern
Alps of New Zealand. Journal of glaciology 2: 140-145.
16
New Zealand Journal of Geology and Geophysics, 1995, Vol. 38
Harrison, C. H. 1973: Radio-echo sounding of horizontal layers in
ice. Journal of glaciology 12: 383-397.
Hay, J. E.; Fitzharris, B. B. 1988: A comparison of the energybalance and bulk-aerodynamic approaches for estimating
glacier melt. Journal of glaciology 34: 1-9.
Hochstein, M. P.; Broadbent, M. 1971: How thick is the Tasman
Glacier? Department of Scientific and Industrial Research
newsletter, December: 18-21.
Hughes, T. 1992: Theoretical calving rates from glaciers along
ice walls grounded in water of variable depths. Journal of
glaciology 38: 282-294.
Kanasewich, E. R. 1963: Gravity measurements on the Athabasca
Glacier, Alberta, Canada. Journal of glaciology 4:
617-630.
King, L.; Fisch, W.; Haeberli, W.; Waechter, H. P. 1987:
Comparison of resistivity and radio-echo soundings on
rock glacier permafrost. Zeitschrift fur Gletscher-kunde
und Glazialgeologie 23:11-91.
Kirkbride, M. P. 1989: The influence of sediment budget on
geomorphic activity of the Tasman Glacier, Mount Cook
National Park, New Zealand. Unpublished Ph.D. thesis,
lodged in the Library, University of Canterbury,
Christchurch.
Kirkbride, M. P. 1993: The temporal significance of transitions
from melting to calving termini of glaciers in the central
Southern Alps of New Zealand. The Holocene 3: 232-240.
Klingele, E.; Kahle, H. G. 1977: Gravity profiling as a technique
for determining the thickness of glacier ice. Pure and
applied geophysics 115: 989-998.
Leary, S. F. 1993: Ground-penetrating radar profiles of the Tasman
and Mueller Glaciers. Unpublished B.Sc. (Hons) thesis,
lodged in the Library, University of Canterbury,
Christchurch.
v. Lendenfeld, R. 1884: Der Tasman-Gletscher und seine
Umrandung. Petermanns Mitteilungen, Erganzungsband
XVI, No. 75, 80 p.
New Zealand Department of Lands and Survey 1972: Mt Cook
topographic map S79 (1:63,360). Wellington, New Zealand
Government Printer.
New Zealand Department of Survey and Land Information
(DOSLI) 1992: Mt Cook topogaphic map 260-H36
(1:50,000). Wellington.
Nobes, D. C ; Leary, S. F.; Hochstein, M. P.; Henrys, S. A. 199 i:
Ground-penetrating radar profiles of rubble-coven d
temperate glaciers: Results from the Tasman and Mud I r
Glaciers of the Southern Alps of New Zealand. Proceedin :is
SEG 1994 annual meeting, Los Angeles (extendi d
abstract).
Odell, N. E. 1960: The mountains and glaciers of New Zealan 1.
Journal of glaciology 3: 739-741.
Olivier, R. J.; Simard, R. G. 1981: Improvement of the con c
prism model for terrain correction in rugged topograpli,.
Geophysics 46: 1054-1056.
Posamentier, H. W. 1977: A new climatic model for gla;: :r
behaviour in the Austrian Alps. Journal of glaciology I <:
57-65.
Robinson, S. D.; Moorman, B. J.; Judge, A. S.; Dallimore, S. R.
1993: The characterization of massive ground ice at Ya^ a
Lake, Northwest Territories, using radar stratigrap y
techniques. In: Current research, part B. Geological Surv< v
of Canada paper 93-1B: 23-32.
Rose, J. H. 1937: Recent retreat of Murchison Glacier, tfv-v
Zealand journal of science and technology 19: 453^t5'i
Rothlisberger, H. 1972: Water pressure in intra- and subglaci .1
channels. Journal of glaciology 11: 177-203.
Rothlisberger, H.; Vogtli, K. 1967: Recent D.C. resistivity
soundings on Swiss glaciers. Journal of glaciology ':>:
607-621.
Salinger, M. J.; Hay, J.; McGann, R.; Fitzharris, B. 1993: Sout iwest Pacific temperatures: diurnal and seasonal trend-..
Geophysical research letters 20: 935-938.
Sara, W. A. 1968: Franz Josef and Fox Glaciers 1951-1967. Ne v
Zealand journal of geology and geophysics 11: 768-78'i.
Shreve, R. L. 1972: Movement of water in glaciers. Journal f
glaciology 11: 205-214.
Skinner, B. E. 1964: Measurements of twentieth century ice Sc -s
on the Tasman Glacier, New Zealand. New Zealandjourr:id
of geology and geophysics 7: 796-803.
Suggate, R. P. 1950: Franz Josef and other glaciers of the Southei n
Alps, New Zealand. Journal of glaciology 1: 422-429.
Warren, C.R.I 992: Iceberg calving and the glacioclimatic record.
Progress in physical geography 16: 253-282.
Warren, C. R.; Sugden, D. E. 1993: The Patagonian icefields; a
glaciological review. Arctic and alpine research 2>:
316-331.