Arctic, Antarctic, and Alpine Research, Vol. 44, No. 3, 2012, pp. 359–367
Geometric Changes in a Tidewater Glacier in Svalbard
during its Surge Cycle
Tavi Murray*‡
Timothy D. James*
Yuri Macheret†
Ivan Lavrentiev†
Andrey Glazovsky† and
Helena Sykes*
*Glaciology Group, Swansea University,
Swansea, SA2 8PP, U.K.
†Institute of Geography, Russian
Academy of Sciences, Staromonetny per.,
29, Moscow 119017, Russia
‡Corresponding author:
t.murray@swansea.ac.uk
Abstract
Fridtjovbreen, Svalbard, is a partially tidewater-terminating glacier that started a 7-year
surge during the 1990s. Flow peaked during 1996 and no surge front was apparent. We
use two pre-surge (1969 and 1990) and a post-surge (2005) digital elevation models
(DEMs) together with a bed DEM to quantify volume changes and iceberg calving during
the surge, calculate the changes in glacier hypsometry, and investigate the surge trigger.
Between 1969 and 1990, the glacier lost 5% of its volume, retreated 530 m and thinned
by up to 60 m in the lower elevations while thickening by up to 20 m in its higher
elevations. During the surge, the reservoir zone thinned by up to 118 m and the receiving
zone thickened by ⬃140 m. Fridtjovbreen’s ice divide moved ⬃500 m, incorporating
extra ice into its catchment. Despite this volume gain, during 1990–2005 the glacier lost
⬃10% of its volume through iceberg calving and 7% through surface melt. The surge
occurred in a climate of decreasing overall ice volume, so we need to revise the notion
that surging is triggered by a return to an original geometry, and we suggest Fridtjovbreen’s
surge was triggered by increasing shear stresses primarily caused by increases in surface
slope.
DOI: http://dx.doi.org/10.1657/1938-4246-44.3.359
Introduction
A surge-type glacier cycles between long periods of slow flow
(quiescence) and short periods of fast flow (surge), during which
ice flow speed may accelerate by over an order of magnitude (e.g.,
Cuffey and Paterson, 2010). During the quiescent phase, ice flow
rates are slower than the balance velocity, and mass builds up in
the upper part of the glacier; simultaneously, the lower part of the
glacier thins and the glacier consequently steepens overall. The
front margin of the glacier often also retreats. During the surge,
fast flow results in transport of ice downglacier at rates above the
balance velocity, often accompanied by an advance of the glacier
front margin. Thus a surge-type glacier is viewed as cycling between two end member geometries, with the extremes occurring
immediately pre- and post-surge. In his description of this dynamic
and geometric cycle, Raymond (1987) stated that thickness changes
during quiescence ‘‘reverse the thickness changes of the surge and
gradually return the glacier to near its presurge state’’; van der
Veen (1999) stated that surging starts when a ‘‘glacier reaches a
certain thickness’’, and Cuffey and Paterson (2010) when a ‘‘glacier attains some critical profile.’’ There thus seems to be considerable agreement in the literature that we should expect glaciers to
return to a pre-surge geometry prior to a subsequent surge.
Remote sensing techniques, such as satellite radar interferometry and cross-correlation tracking, have led to considerable knowledge of the changing ice dynamics through the active phase of
glacier surging in Svalbard (Luckman et al., 2002; Strozzi et al.,
2002; Murray et al., 2003a, 2003b; Mansell et al., 2012). Synthesis
of remote sensing results with those from field methods suggests
two types of surges occur in Svalbard (Murray et al., 2003b). During surges of land-terminating glaciers, typified by the surge of
Bakaninbreen (Murray et al., 1998), fast flow appears to initiate
in the reservoir zone and a surge front propagates downglacier, in
䉷 2012 Regents of the University of Colorado
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a manner similar to that observed at Variegated Glacier, Alaska
(Kamb et al., 1985). In contrast, during surges of tidewater-terminating glaciers, fast flow appears to initiate low down and spread
both up and downglacier (Pritchard et al., 2005). In Svalbard, this
type of surge is typified by the surge of Monacobreen (Luckman
et al., 2002; Murray et al., 2003b). The two types of surge form
distinct crevasse and structural patterns because the passage of a
surge front downglacier is characterized initially by compression
(typically forming longitudinal crevasses) followed by extension
(forming transverse crevasses) as the front passes (Murray et al.,
1998). If no surge front forms or in regions of the glacier that
experience only extensive flow regimes only transverse crevasses
will occur.
Remote sensing techniques also allow mapping of the changing morphology of a glacier surface in 3-D during a glacier surge.
Such studies have been undertaken in a variety of environments
including Greenland (Jiskoot et al., 2001; Pritchard et al., 2003)
and Alaska (Muskett et al., 2008, 2009; Shugar et al., 2010). In
Svalbard an analysis of repeated ASTER digital elevation models
(DEMs) was used to suggest a long initiation phase, with mass
being transferred downglacier slowly causing deflation of the surface at high elevations prior to the fast flow event (Sund et al.,
2009). In this paper we use photogrammetry and airborne lidar
to produce three surface DEMs covering the full surface area of
Fridtjovbreen, Svalbard. Fridtjovbreen is a partially tidewaterterminating glacier at which fast flow during the surge started low
down on the glacier (Murray et al., 2003a). Our DEMs cover both
late quiescence and postsurge, and therefore for the first time for
a Svalbard surge, bracket the full active phase. We use these DEMs
to quantify the redistribution of ice and iceberg calving, and to
calculate the changes in glacier hypsometry and to investigate the
surge trigger.
TAVI MURRAY ET AL. / 359
FIGURE 1. Shaded relief DEM from 2005 of Fridtjovbreen (FJ)
and Grønfjordbreen (G) from lidar data. Contours are spaced at
50 m intervals, heavy white contours at 100 m. Coordinates are
UTM coordinates (zone 33X) in meters (ellipsoid WGS84). Black
line delineates the 2005 glacier boundary. Dashed white line is profile shown in Figure 4. Insert shows location of FJ within Svalbard.
Additional margin positions are shown from 1990 and 1969 (aerial
photographs), 1997, 1998/1999 from maps and reconstructed 1860/
1861 margin from Lønne (2006).
Study Area
Fridtjovbreen, Svalbard (Fig. 1; 77⬚50′N, 14⬚26′E) is a 13km-long surge-type glacier with an area of ⬃38 km2. The glacier
flows southwards from an ice divide with the non-surge-type glacier, Grønfjordbreen. Fridtjovbreen flows towards van Mijenfjord,
which trends east-west, approximately perpendicular to Fridtjovbreen’s flow. Fridtjovbreen varies in width from ⬃5 km to ⬃1.5
km near its front margin. Radio-echo sounding shows the glacier
is polythermal over most of its length. A cold surface layer of
thickness ⬃100 m overlies warm ice up to 220 m thick (Macheret
and Glazovsky, 2000). This thermal structure was confirmed by
two deep boreholes (Kotlyakov and Macheret, 1987). Fridtjovbreen
is one of a small number of glaciers in Svalbard where multiple
surges have been observed (Mansell et al., 2012): one occurred ca.
1860–1861 (Hagen et al., 1993), a possible surge event in ca. 1898
360 / ARCTIC, ANTARCTIC, AND ALPINE RESEARCH
(see discussion below), and one which began in 1995 (Murray et
al., 2003a).
During 1858 Fridtjovbreen was visited by Torell’s expedition
and was clearly in quiescence, with its terminus situated inland of
Fridtjovhamna (harbor), separated from the fjord by a ‘‘marsh’’
(Liljequist, 1993). However, by winter 1860–1861 Nordenskiold’s
expedition found that the ‘‘once small glacier had expanded across
the moraines before it in the harbour’s interior; it had then filled
the entire harbour and even passed beyond’’ (Liljequist, 1993). At
that time the terminus almost reached Axelsøya, a rock bar at the
mouth of van Mijenfjord (Hjelle et al., 1986). We interpret the
glacier to have surged ca. 1860–1861 from these descriptions (see
inferred marginal position from Lønne [2006] on Fig. 1).
Six years later in 1866, maps produced by Nordenskiold show
the glacier terminating well inland (Glasser et al., 1998). A subsequent map of 1898 by Kyellstriöm and Hamberg again shows the
glacier beyond the confines of Fridtjovhamna (in Nathorst, 1900).
However, by 1909, the glacier surface was described as smooth
and uncrevassed (Glasser et al., 1998). This report of the front
margin again outside the harbor in 1898 suggests a previously unreported surge event may have occurred at the end of the 19th century.
Subsequently for the majority of the 20th century, the glacier
continued to retreat, displaying a pattern of low-elevation thinning
and high-elevation thickening typical of a surge-type glacier in
quiescence. For example, between 1936 and 1988, the front margin
retreated by ⬃1.3–2 km, and thinned by ⬃100 m (Glazovsky et
al., 1991). Over the same time period, the glacier thickened at
elevations higher than 350 m a.s.l., with the northwestern (NW)
tributary thickening by ⬃40 m (0.8 m aⳮ1) and the northeastern
(NE) tributary by ⬃10 m (0.2 m aⳮ1). During the early 1990s, the
western side of the front margin was tidewater-terminating, forming
a cliff, whereas the eastern side terminated on land (Musial, 1994).
The majority of the ice surface was crevasse-free.
The dynamics of the 1990s surge was studied using satellite
radar interferometry (Murray et al., 2003a). The glacier underwent
a slow acceleration between 1991 and June 1995, with flow rates
exceeding the balance velocity throughout. Fast flow started in the
lower third of the glacier (Murray et al., 2003a). The trunk and
NW tributaries moved faster than the NE tributary. The glacier
surface changed, with intermittent lakes and increased transverse
crevassing (Musial, 1994). The glacier front remained static until
at least March 1994; however, by October 1995 the glacier had
advanced 350 m. Between June 1995 and February 1996, the glacier
accelerated rapidly, reaching a measured maximum ice flow speed
of 2.5 m dⳮ1. By May 1997, the glacier front had advanced ⬃2.8
km from its pre-surge location at peak rates of 4.2 m dⳮ1 and its
margin became intensely chaotically crevassed. The next measurement in October 1997 showed flow rates had begun to decrease
and the margin had begun to retreat (Murray et al., 2003a), however,
Lønne (2006) reports subsequent re-advance of ⬃650 m with maximum frontal extent in 1998/1999. The maximum frontal position
was less advanced than that of 1860–1861 (Fig. 1). Bamber et al.
(2005) reported thinning of ⳮ0.26 m aⳮ1 along a single flowline
on Fridtjovbreen between 1996 and 2002, but these measurements
were made after the main advance had occurred and are not comparable either in technique or timeframe with the measurements we
present in this paper.
Field observations suggested that the fast-flow phase of the
1990s surge started near the terminus (A. J. Hodson, personal communication, 2002) and propagated up-glacier, initially to the NW
tributary in agreement with flow speeds. The initiation zone for
fast flow is corroborated by surface crevasses on the upper glacier,
which were exclusively transverse (Murray et al., 2003a), suggesting only extension had occurred in this region as well as the remote
sensing observations described above. Drawdown in the reservoir
zone was visually estimated to be ⬃10 m (A. J. Hodson, personal
communication, 2002), although in this paper we show it was actually much greater. No surge front was observed during the fast
flow phase in the field, on satellite images or on interferograms, and
the 1990s surge of Fridtjovbreen appears to fit the characteristics of
a tidewater-terminating glacier surge.
Methods
SURFACE DEMS
A lidar DEM was collected in July 2005 over the Fridtjovbreen
and Grønfjordbreen catchments by the U.K. Natural Environment
Research Council’s (NERC) Airborne Research and Survey Facility (ARSF) using an airborne Optech Airborne Laser Terrain Mapper (ALTM) 3033 system. This instrument records the timing of
the first pulse, last pulse, and 8-bit laser intensity data for each
observation. A Global Positioning System (GPS) base station was
run at Finsterwalderbreen, ⬃30–45 km from Fridtjovbreen, and
these data were combined with: (i) the laser range data; (ii) GPS
mounted on the aircraft; and (iii) Inertial Navigation System (INS)
measurement of aircraft attitude (pitch, roll, and yaw) in post-processing to produce a data cloud of (X, Y, Z, intensity) data. The
lidar data set consisted of over 150 million data points with an
average spatial resolution of ⬃1 point per 1.67 m2. The data are
of sufficient quality and spatial resolution that both surface crevasses and supraglacial channels can be identified on the glacier
surface (Fig. 1).
Topographic data were also derived from 1969 panchromatic
(1:20,000) and 1990 color infrared (1:50,000) vertical aerial photographs from the Norwegian Polar Institute, which were captured
using calibrated metric cameras. All of the images were collected in
summer (July or August). Photogrammetric-quality scanning was
undertaken at resolutions of 12.5 and 10 m for the 1969 and
1990 photographs, respectively. The photogrammetric models were
controlled using 67 ground control points located on geologically
stable surfaces extracted from the lidar data (James et al., 2006;
Barrand et al., 2009). Where the automated stereo-image correlation failed (i.e. in steep or snow-covered areas), points were collected manually wherever surface texture allowed. Both photogrammetric DEMs were collected with a point spacing of 20 m.
The three surface DEMs bracket the fast flow phase of Fridtjovbreen’s 1990s surge. The period between the 1969 and 1990
photogrammetric DEMs covers 21 years of late quiescence,
whereas the period between 1990 and 2005 covers the active phase
of the surge.
Volume changes between epochs were calculated by differencing the surface DEMs. To isolate ice-covered areas, ice mask
polygons were constructed for each epoch from the aerial photography. The relevant masks were combined to capture the largest ice
area for each period. However, since Fridtjovbreen is tidewater-
terminating, the volume of ice stored or lost between successive
ice margin positions also requires knowledge of the fjord bottom
bathymetry, which is discussed below.
The hypsometry of a glacier is the area-elevation distribution.
Hypsometry is controlled by factors such as valley shape and topographic relief, and is thought to be a factor that may contribute to
whether a glacier has normal, stable flow or exhibits unstable, surge
behavior. In Alaska a glacier that is ‘bottom heavy’ has been suggested to be more likely to be of surge-type (Wilbur, 1988); however, no such relationship was found in Svalbard (Jiskoot et al.,
2000). We therefore calculate and present the hypsometry of Fridtjovbreen both before and after its surge.
SURFACE DEM AND VOLUME CHANGE ERROR ASSESSMENT
The remote nature of Fridtjovbreen meant it was not possible
to collect ground-truth data at the glacier simultaneously with lidar
data collection, but such data were acquired at several more accessible sites around Svalbard. These data included points close to the
main GPS base station (i.e., over the airport runway), as well as
two sites further afield with baseline lengths of similar distance as
the Fridtjovbreen site. Those data were used to provide a robust
assessment of the quality of the lidar data collection and postprocessing. The results show mean elevation differences on glaciated regions of 0.14 m (Barrand et al., 2009).
The quality of the photogrammetrically derived surface DEMs
was assessed in two ways. First, the root-mean-square (RMS) errors
of the block adjustments provided an indication of the degree to
which the block adjustments conformed to measured parameters
and their specified precisions (Table 1). Second, a small area of
each DEM (3500 pixels), which was geologically stable, and had
similar contrast and slope to the glacier surface, was compared to
the 2005 lidar DEM. This approach provided a means of assessing
the quality of the image correlation stage of the photogrammetry
in the absence of local ground-truth data. The results show mean
errors of 1 m or less with standard deviations better than 1.4 m
(Table 2). Errors in the difference DEMs can be calculated using
standard error theory, which gives ⬃1.8 m and ⬃1.4 m in the pre
and post-surge difference DEMs, respectively. These errors are
evaluated for good contrast areas, and are the values we use in the
area of the glacier which consists of bare ice in the photographs.
In regions of the difference DEM where the contrast is lower due
to snow cover we take a more conservative error estimate of 2 m.
These estimates are then propagated to estimate errors in volume
change which are quoted in Table 3.
BED DEM
We combined fjord sounding depths, photogrammetric heights
of terrestrial margin areas, and radio-echo sounding (RES) data
TABLE 1
Block adjustment root-mean-square (RMS) errors (in meters).
RMS error X
RMS error Y
RMS error Z
Image (pixels)
1969
1990
2.7
3.4
0.7
1.3
1.3
1.3
0.4
1.1
TAVI MURRAY ET AL. / 361
BASAL SHEAR STRESSES
TABLE 2
Error statistics for the comparison of small regions of each photogrammetric digital elevation model (DEM) with the 2005 lidar
DEM. Regions were selected to have stable characteristics and similar slopes to the glacier surface (in meters).
Minimum
Maximum
Number of pixels
Mean
Std Dev
2005–1969
2005–1990
ⳮ5.8 m
9.2 m
3232
1.0 m
1.2 m
ⳮ6.1 m
7.5 m
3232
ⳮ0.3 m
1.4 m
We extracted profiles from the DEM and bed on an approximate flowline along the NE tributary (profile location shown in
Fig. 1) to the front margin. These surface and bed profiles allow
calculation of the basal shear stress using the method of Kamb
and Echelmeyer (1986) which takes into account the effects of
longitudinal stress gradients. Basal shear stress B is calculated
from:
B(x) ⳱
1
2艎h1/n
冕
Ⳮⴥ
L(x′)h1/n(x′) expⳮ|x′ⳮx|/艎dx′,
(1)
ⳮⴥ
(Fig. 2, part a) to create a bed DEM for the glacier. These data
sets were interpolated into a 200 m DEM using an inverse distance
weighting algorithm within ArcGIS (Fig. 2, part b).
The sounding depths were collected in 1984 and covered much
of the fjord region between the 1990 and 2005 calving fronts of
Fridtjovbreen (L. Christiansen, personal communication, 2007).
The sounding depths are affected by systematic error related to the
unknown height of the tide at the time of measurement, but this is
small in Svalbard (⬃2 m tidal range; Norwegian Hydrographic
Service, 1990). Radio-echo sounding (RES) at 2–13 MHz was
undertaken at Fridtjovbreen in 1988 (Glazovsky et al., 1991). The
survey consisted of 197 soundings over the relatively uncrevassed
regions of the glacier (Fig. 2, part a). These soundings were
converted to depths below surface using the results of a commonmidpoint survey, which determined the mean velocity of electromagnetic wave (e.m.) propagation to be 0.169 Ⳳ 0.002 m nsⳮ1,
and then to absolute elevation using the 1990 photogrammetric
surface elevations.
The uncertainty in the radar depths are likely to be dominated
by (i) error in the e.m. velocity used, especially given that Fridtjovbreen is polythermal and therefore e.m. propagation velocity will
vary spatially; and (ii) by picking error, because of the low frequency of the radar transmission. The likely error is ⬃Ⳳ10 m.
where 艎 is the longitudinal coupling length (i.e. the length over
which the effects of changes in ice thickness and surface slope are
averaged), h is the ice thickness, L is the local slope stress, n is
the exponent in Glen’s flow law, which we take equal to 3, and x
is the distance coordinate along the glacier centerline. The ratio 艎/h
lies in the range 1.5–2.5 for valley glaciers (Kamb and Echelmeyer,
1986), thus we use 艎 equal to twice the ice thickness.
Results
The new bed DEM for Fridtjovbreen (Fig. 2) covers ⬃85%
of the glacier bed of which ⬃5% lies below sea level. The crevasse
pattern and the steep shape of the calving front suggests that much
of the area where depths are unknown also lies below sea level.
Individual basins of up to 50 m below sea level were revealed in
similar locations as reported in Glazovsky et al. (1991), but overall
the bed is less deep than was provisionally reported in that publication.
During the pre-surge period (Fig. 3, part a), the glacier retreated by ⬃530 m (25 m aⳮ1) and thinned by up to 60 m (2.9 m
aⳮ1). The line delineating mass loss near the front margin from
mass gain is at an elevation of ⬃310 m a.s.l. At elevations lower
than this, the mean thinning rate was 0.7 Ⳳ 0.1 m aⳮ1 (Table 3
summarizes the changes between the epochs), while at elevations
above this the glacier thickened by up to 20 m (0.95 m aⳮ1) with
a mean thickening rate of 0.04 Ⳳ 0.1 m aⳮ1. However, much of
TABLE 3
Summary of changes between DEMs and error estimates. In the early epoch we use the term ‘‘reservoir zone’’ to refer to the high elevation
region above ⬃310 m a.s.l., which shows areas of minor mass gain, and receiving zone for the part of the glacier lying at elevations lower
than this altitude, which is losing mass (Fig. 3, part a). In the later epoch (surge) we use the terms ‘‘reservoir zone’’ and ‘‘receiving zone’’
to describe the regions that lose and gain mass, respectively. The line delineating these zones during the surge is at an elevation of
⬃204 m a.s.l. (Fig. 3, part b). See text for discussion of areas A–C and Figure 3 for their locations.
Pre-surge (1969–1990)
Reservoir zone
Receiving zone (excluding area A)
Area A
Total pre-surge
Surge (1990–2005)
Reservoir zone
Receiving zone (excluding area B and C)
Area B
Area C
Receiving zone total
Total surge
Area (m2)
Mean thickness change (m)
Volume change (m3)
24.8
17.2
0.6
42.6
⳯
⳯
⳯
⳯
106
106
106
106
1Ⳳ3
ⳮ15 Ⳳ 2
ⳮ77 Ⳳ 2
20
ⳮ270
ⳮ50
ⳮ300
⳯
⳯
⳯
⳯
106
106
106
106
Ⳳ
Ⳳ
Ⳳ
Ⳳ
60 ⳯ 106
30 ⳯ 106
1 ⳯ 106
70 ⳯ 106
34.4
7.0
2.0
3.0
12.1
46.5
⳯
⳯
⳯
⳯
⳯
⳯
106
106
106
106
106
106
ⳮ46
36
119
56
2
1
1
1
ⳮ1600
260
244
170
670
ⳮ930
⳯
⳯
⳯
⳯
⳯
⳯
106
106
106
106
106
106
Ⳳ
Ⳳ
Ⳳ
Ⳳ
Ⳳ
Ⳳ
70 ⳯ 106
10 ⳯ 106
3 ⳯ 106
4 ⳯ 106
10 ⳯ 106
70 ⳯ 106
362 / ARCTIC, ANTARCTIC, AND ALPINE RESEARCH
Ⳳ
Ⳳ
Ⳳ
Ⳳ
FIGURE 2. (a) Data locations (red points from radio echo-sounding and lidar/photogrammetry; blue points from bathymetric survey) for
(b) bed DEM for Fridtjovbreen in m a.s.l. Contours are spaced at 50 m intervals. Coordinates are UTM coordinates (zone 33X) in meters
(ellipsoid WGS84). Main text discusses data sources used in producing this DEM. White area shows region where no data exist because
of heavy surface crevassing. Green dashed line is 1990 margin, blue dotted line is 1969 margin, and red line is 2005 margin.
the NW tributary lost mass over this period even at the highest
elevations, and mass gain was largely limited to the upper and
central parts of the NE tributary (Fig. 3, part a): a total volume of
2 Ⳳ 6 ⳯ 107 m3 was added to the glacier’s reservoir zone. Ice
margin retreat between 1969 and 1990 meant the glacier lost area
of 9.4 ⳯ 105 m2. This retreat showed the glacier had been partially
land- and partially tidewater-terminating. The exposed bed above
sea level could be measured from the 1990 DEM and had maximum
elevation of ⬃65 m. Fjord depths where the glacier was tidewaterterminating between the two margin positions are reasonably well
known (Fig. 2), and range between 0 and 39 m below sea level.
Combining the bed and surface topography in this frontal region
shows 5.0 ⳯ 107 m3 of ice were lost due to marginal retreat over
the fjord (region A in Fig. 3, part a) together with 27 ⳯ 107 m3
of ice from the rest of the lower glacier. Overall the glacier lost
volume equal to 30 Ⳳ 7 ⳯ 107 m3 during these 21 years, which
is ⬃5% of the total glacier volume, equivalent to 0.33 m aⳮ1 averaged over the whole glacier surface.
The glacier front was ⬃1800 m further advanced in 2005 than
in 1990 (Figs. 1 and 3, part b), although some retreat had already
occurred and, as discussed above, this position does not represent
the maximum extent during the surge. The reservoir zone thinned
by a maximum of 118 m (mean 46 Ⳳ 2 m), and the receiving zone
had thickened by a maximum of 140 m (mean of 36 Ⳳ 1 m).
These thickness change values exclude the area of the ice advance
between the 1990 and 2005 margin, where the average ice elevation
is 110 m a.s.l. Combining the surface and bed DEMs, the volume
stored by ice advancing in areas with elevations below sea level
(area B in Fig. 3, part b) and above sea level (area C) was 41 ⳯
107 m3 (Table 3 summarizes the changes and gives error estimates).
Overall the glacier lost 93 ⳯ 107 m3 of ice during this 15 year
period, equivalent to 17.2% of the glacier volume, or a thinning
rate of 1.33 m aⳮ1 averaged over the entire glacier surface. The
line delineating mass loss (reservoir zone) and mass gain (receiving
zone) lay at ⬃204 m a.s.l. (Fig. 3, part b), somewhat lower on the
glacier than the equivalent line in the earlier epoch (Fig. 3, part a),
and much better defined. At the highest elevations in the NE tributary, the catchment of Fridtjovbreen expanded during the surge to
capture some ice that was previously within the Grønfjordbreen
catchment (compare Fig. 3, parts a and b). This change in the size
of the glaciers’ catchments can also be seen in topographic profiles
(Fig. 4, part a). These profiles highlight the changes in the surgetype Fridtjovbreen in contrast to its neighbor Grønfjordbreen.
To estimate the volume of icebergs calved during the surge
we compare the overall thinning rates of Fridtjovbreen (ⳮ0.33 m
aⳮ1) and Grønfjordbreen (ⳮ0.74 m aⳮ1) in the first pre-surge
epoch (1969–1990) and assume that the ratio of the thinning rates
due to melt between the glaciers in the second epoch (1990–2005)
TAVI MURRAY ET AL. / 363
FIGURE 3. (a) Pre-surge difference DEM (1969 to 1990). Line shows approximate delineation between reservoir zone and receiving zone,
although not all areas within the reservoir zone thickened. Maximum thickening ⬃20 m occurred in the NE tributary. Area A is region
where glacial retreat exposed fjord water, so the volume change in this area is less well known because of limitations of the bed DEM. (b)
Post-surge difference DEM (1990 to 2005). The DEM shows the effects of the glacier surge. Line shows delineation between reservoir zone
and receiving zone (note that this is at a different location and elevation to the equivalent line in (a); see text for details). Area B is region
where advance was into fjord water, so the volume change in this area is less well known because of limitations of the bed DEM. Areas
C are regions where advance was over sediments or the tributary glacier Sagabreen (on west side). In both parts of this figure coordinates
are UTM coordinates (zone 33X) in meters (ellipsoid WGS84). Contours of elevation change are at 10 m spacing, contours at 50 m spacing
are thicker: red for mass loss, blue for mass gain. Zero change contour is black dashed line.
is the same. Using this ratio and the thinning rate of Grønfjordbreen
in the second epoch (ⳮ0.53 m aⳮ1), we estimate the mass loss of
Fridtjovbreen due to melting was ⬃37 ⳯ 107 m3. Thus ⬃56 ⳯
107 m3 (equivalent to a thinning rate of 0.81 m aⳮ1 averaged over
the whole glacier surface) was most likely calved as icebergs into
the fjord. This volume of calved ice represents ⬃10.4% of the total
glacier volume.
The greatest changes in basal shear stress occur during the
surge (Fig. 4, part b), between the 1990 and 2005 epochs, when
basal shear stresses decrease in the reservoir zone and increase in
the receiving zone due mainly to the large thickness changes. In
the earlier epoch, during the buildup to fast flow (1969–1990),
shear stresses increase over most of the lower glacier whereas
stresses in the upper glacier show no change or decreasing shear
stress during this period.
Figure 5 shows the hypsometry of Fridtjovbreen calculated at
each of the DEM epochs. The hypsometric curves have essentially
the same shape in both 1969 and 1990, reflecting a small amount
of mass gain at elevations mainly between 450 and 500 m a.s.l. on
364 / ARCTIC, ANTARCTIC, AND ALPINE RESEARCH
the NE tributary; however, there is a small mass gain at all elevation
bands above this in addition, so there is no support for this period
representing a period of mass transport downglacier as described
by Sund et al. (2009). Rather the eastern side of the glacier thickened slightly and the western side thinned (Fig. 3, part a).
The surge caused transfer of mass downglacier, resulting in
a large areal increase in the region of the glacier between 200 and
400 m a.s.l. In 1969, 61% of the glacier’s area lay at elevations
above the 1969–1990 measured reservoir-receiving line altitude,
and by 1990 this had increased to 64%. After the surge, however,
only 43% of the glacier’s area lay above this altitude. Furthermore,
comparison with measurements on neighboring non-surge-type glaciers show the equilibrium line altitude (ELA) is much higher during this period. For example, on Slakbreen the ELA has been at
around 750 m a.s.l. since 1961 (Kohler et al., 2007). It is normally
assumed that for a glacier to be in equilibrium, its accumulation
area should comprise around 70% of the total glacier area (Cuffey
and Paterson, 2010), so that overall mass loss from Fridtjovbreen
in both epochs is not unexpected.
FIGURE 4. Centerline profiles from south to north along Fridtjovbreen and Grønfjordbreen. Location shown in Figure 1. (a)
Surface and bed topography. (b) Basal shear stress calculated for
Fridtjovbreen only as there is no bed DEM for Grønfjordbreen;
see text for details of calculation.
In reporting the volume change of the glacier during the surge
we have not made any correction for the increase in crevassing on
the glacier during the surge. In a study of Sortebr+, east Greenland,
Jiskoot et al. (2001) applied a volume correction in crevassed areas
of ⬃10% to allow for this. The surface of Fridtjovbreen did become
significantly crevassed during the surge (Murray et al., 2003a),
which decreases the volume of ice stored in the post-surge glacier;
however, the ice surface during a surge in Svalbard is less crevassed
than occurs at Sortebr+, and any volume correction needed is unlikely to be as large as this. If we were to apply such a correction,
this would increase the calculated volume of ice calved and hence
rates of volume loss during the surge event. We have further assumed that the release of subglacially stored water does not contribute to volume change significantly. This could be an important
consideration if the glacier was fast flowing during any of the DEM
collection epochs, as measurements at other glaciers show the ice
surface can drop significantly at the end of the fast flow phase
(Kamb et al., 1985). However, during 1990, flow rates at Fridtjovbreen had not accelerated, and by 2005, flow rates had dropped
significantly and the margin was already in retreat, so this is unlikely to be a significant effect.
Discussion
FIGURE 5. (a) Hypsometry for Fridtjovbreen at the different time
epochs. A* and Z* are dimensionless area and elevation, respectively (e.g., Furbish and Andrews, 1984). A* is the cumulative area
above a particular elevation. Z* is 0 at the glacier margin and 1
at the highest elevations. (b) Fraction of area of glacier in each
elevation band in each time epoch.
An implication of the standard description of glacier surging
is that surge behavior is a dynamic perturbation either side of equilibrium (flow being alternately too fast, too slow), and a glacier is
assumed to return at the end of quiescence to its pre-surge geometry
(see for example, discussions in van der Veen [1999], Benn and
Evans [2010], and Cuffey and Paterson [2010]). However, in many
surge clusters in the world, glacier mass balance is known to be
negative. In western Svalbard, for example, Kohler et al. (2007)
have shown that mass loss has been occurring at an accelerating
TAVI MURRAY ET AL. / 365
rate since the 1930s, and mass balance is likely to have been negative since the end of the Little Ice Age. Surges are common in this
area even though the glaciers clearly cannot be returning to their
original geometry after each event.
At Variegated Glacier, Alaska, surging appears to occur triggered by increases in glacier thickness (Eisen et al., 2001). However, at Fridtjovbreen crevassing on the glacier was first reported
close to the front margin (A. J. Hodson, personal communication,
2002), where considerable thinning of the glacier was occurring.
Remote sensing observations of flow during the early part of the
surge show that fast flow started in the lower third of the glacier,
and although the location cannot be better determined due to the
limitations of SAR interferometry, fast flow clearly did not start
at high elevation (Murray et al., 2003a). Furthermore, at the highest
elevations, fast flow occurred first on the NW tributary, with the
NE tributary, which is the part that thickened most in our measurements (the last 21 years of quiescence; Fig. 3, part a), speeding up
only later in the surge (Murray et al., 2003a). Thus, on Fridtjovbreen
it appears that the surge did not start in the upper part of the glacier,
and certainly not in the region of greatest recent thickening (although note that the earlier measurements for 1936–1988 reported
by Glazovsky et al. [1991] show the NW tributary thickening most).
Kohler et al. (2007) reported that the glaciers and small ice cap
they measured have been thinning at all elevations to 700 m a.s.l.
since 1977. This is equivalent to the highest elevation on Fridtjovbreen (Fig. 5). Furthermore, in 1987–1988 the measured net mass
balance on Fridtjovbreen was ⳮ0.37 m w.e. (Glazovsky et al.,
1991). It therefore seems unlikely that the fast flow during Fridtjovbreen’s surge was triggered by thickness changes, and increases in
shear stress due to steepening (Fig. 4, part b) are more likely to be
the cause. Steepening occurred mainly due to thinning in the lower
glacier and retreat of the ice front rather than thickening of the
upper glacier (Fig. 4 shows both Fridtjovbreen and Grønfjordbreen
steepening, despite the latter thinning at all elevations). This pattern
of steepening in profile of the lower glacier causing increased basal
shear stresses (Fig. 4, part b) is consistent with the surge initiating
in this region.
Dowdeswell et al. (1995) suggested that climate warming
since the end of the Little Ice Age has reduced the number of
glaciers that surge in the Svalbard archipelago due to two effects.
The first is that reductions in mass balance will lengthen the time
taken for a glacier to return to its pre-surge geometry. The second
is that glacier geometries and regional climate changes favor transition to glaciers comprised entirely of cold-ice in Svalbard. This
latter is almost certain to prevent a glacier being capable of surging,
as sliding and sediment deformation are negligible beneath coldice, and no cold-based glacier has ever been observed to surge.
However, as discussed above it is almost certain that most surgetype glaciers in Svalbard are not returning to their pre-surge geometry. If we are correct in our assertion that the surge at Fridtjovbreen
was triggered by increasing shear stresses, changes in surface slope
during the quiescent phase may allow glaciers in Svalbard to continue to surge despite their negative mass balances unless thinning
causes them to become cold-based, or perhaps they retreat and
become land-terminating. Indeed, the interval between successive
surges could decrease as glaciers retreat if they also consequently
steepen.
366 / ARCTIC, ANTARCTIC, AND ALPINE RESEARCH
Conclusions
We have used geodetic methods to produce detailed surface
DEMs at three time epochs that bracket fast flow during a surge
event at Fridtjovbreen in central Spitsbergen. We used these in
combination with a bed DEM to measure changes in ice volume,
hypsometry, basal shear stress, and the volume of icebergs calved.
The glacier lost ⬃5% of its volume during 21 years of quiescence
despite some thickening in the reservoir zone. During the surge
the glacier lost ⬃17.3% of its volume, with ⬃10.4% being calved
as icebergs and the remaining ⬃6.8% running off as meltwater.
The overall volume of Fridtjovbreen has almost certainly been
decreasing at an accelerating rate over the period we have studied.
Despite this reduction in overall volume, the glacier surged in the
1990s. We therefore need to revise the notion that surging in this
region is triggered by a return to an original geometry. In western
Svalbard surges are superimposed on a more than 30 year trend of
accelerating mass loss at elevations to at least 700 m a.s.l., so that
it is unlikely that surges are triggered by the glacier reaching a
certain thickness. We suggest that as a glacier progressively thins
and shortens, changes in surface slope and consequent increases
in basal shear stress trigger the surge.
Acknowledgments
This research was supported by the U.K. Natural Environment
Research Council (NERC) (NE/B505262/1), the NERC Airborne
Remote Survey Facility (ARSF), NERC Geophysical Equipment
Facility (GEF), and the School of the Environment and Society,
Swansea University. The Russian component was supported by
Program 14 from the Branch of Earth Sciences of Russian Academy
of Sciences. We thank the Norwegian Polar Institute and Adrian
Fox (British Antarctic Survey) for their support in the field and
Applied Imagery for supplying Swansea Glaciology Group with
QT Modeler. Tim James is currently funded by C3W (the Climate
Change Consortium of Wales).
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