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Is there a single surge mechanism? Contrasts in dynamics between
glacier surges in Svalbard and other regions
Tavi Murray
School of Geography, University of Leeds, Leeds, UK
Tazio Strozzi
1
and Adrian Luckman
Department of Geography, University of Wales, Swansea, UK
Hester Jiskoot
2
and Panos Christakos
School of Geography, University of Leeds, Leeds, UK
Received 31 March 2002; revised 24 October 2002; accepted 30 December 2002; published 9 May 2003.
[1]During the 1990s, Monacobreen, a 40-km-long tidewater glacier in Svalbard,
underwent a major surge. We mapped the surge dynamics using ERS synthetic aperture
radar images, differential dual-azimuth interferometry and intensity correlation tracking. A
series of 11 three-dimensional (3-D) velocity maps covering the period 1991–1997 show a
months-long initiation and years-long termination to the surge, with no indication of a
surge front travelling downglacier. During the surge, the front of the glacier advanced
2 km, the velocity and derived strain rate increased by more than an order of magnitude,
and maximum ice flow rates measured during 1994 were 5md
1
. The spatial pattern of
both velocity and strain rate was remarkably consistent and must therefore be controlled
by spatially fixed processes operating at the glacier bed. We combine these results with
those published in the literature to construct a typical Svalbard glacier surge cycle and
compare this to surge dynamics of glaciers from other cluster regions, especially those of
Variegated Glacier in Alaska. The strong contrast in dynamics suggests that there exist at
least two distinct surge mechanisms. INDEX TERMS:1863 Hydrology: Snow and ice (1827); 0933
Exploration Geophysics: Remote sensing; 9315 Information Related to Geographic Region: Arctic region;
1827 Hydrology: Glaciology (1863); KEYWORDS:glacier surging, glacier dynamics, interferometry, Svalbard,
synthetic aperture radar (SAR), remote sensing
Citation: Murray, T., T. Strozzi, A. Luckman, H. Jiskoot, and P. Christakos, Is there a single surge mechanism? Contrasts in
dynamics between glacier surges in Svalbard and other regions, J. Geophys. Res.,108(B5), 2237, doi:10.1029/2002JB001906, 2003.
1. Introduction
[2] A surge-type glacier experiences cyclic flow that
alternates between decades of slow flow and shorter periods
of flow that is typically 10–1000 times faster. Less than 1%
of Earth’s glaciers are believed to surge [Jiskoot et al.,
2000], but surge-type glaciers are of great importance in the
understanding of glacier dynamics. The two-phase flow
regime shows that profound changes occur in processes
and conditions beneath the glacier that allow it to switch
between fast and slow flow despite only small changes in
driving stress.
[3] In order to comprehend the mechanisms of glacier
surging it is important to search for systematic differences in
surge dynamics between regions. Many surge-type glaciers
are situated in remote locations, and a surge may be
completely missed or be well under way before it is
recognized. Archives of remotely sensed data provide a
valuable resource for studying glacier dynamics and allow
long-term studies that would logistically be difficult in situ.
Both surface topography and surface displacement can be
measured using differential satellite radar interferometry
(SRI) [Kwok and Fahnestock, 1996]. Dual-azimuth imagery
from ascending and descending passes of the satellite in
conjunction with a digital elevation model (DEM) can be
used to resolve the 3-D ice velocity field [Mohr et al., 1998;
Joughin et al., 1998]. However, because glacier surging is
often associated with extreme ice velocities, SAR interfer-
ometry can be unsuitable for mapping ice dynamics. For
example, Fatland and Lingle [1998] used interferometry to
study the effects of the Bering Glacier surge on the Bagley
icefield reservoir zone, but interferometric coherence was
lost on the glacier itself because the ice was moving and
deforming so rapidly. Similarly, studies on Storstrømmen in
northeast Greenland [Mohr et al., 1998] and Sortebræ, east
Greenland [Murray et al., 2002], were able to use interfer-
ometry only in the quiescent phase. However, the relatively
low ice velocities and long duration that characterize surges
JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. B5, 2237, doi:10.1029/2002JB001906, 2003
1
Also at Gamma Remote Sensing, Muri, Switzerland.
2
Now at Department of Geography, University of Lethbridge, Calgary,
Alberta, Canada.
Copyright 2003 by the American Geophysical Union.
0148-0227/03/2002JB001906$09.00
EPM 3 -1
in Svalbard [Dowdeswell et al., 1991; Murray et al., 1998]
makes them amenable to study using interferometry [Dow-
deswell et al., 1999; Luckman et al., 2002].
[4] In this paper, we use dual-azimuth differential radar
interferometry to derive the flow dynamics of Mona-
cobreen, a surge-type glacier in northern Svalbard, through
seven years of its active phase. A preliminary analysis of
the interferometric data has been presented by Luckman et
al. [2002]. In this paper we also present the results of using
intensity correlation tracking [Strozzietal., 2002] to
determine the ice dynamics in regions of the glacier where
flow rates exceed that measurable using interferometry,
together with a glaciological interpretation of the data. The
combination of these techniques allows us to determine
glacier surface dynamics during the surge in unprecedented
spatial and temporal detail which, together with published
results, allows us to construct a typical Svalbard surge
cycle. We compare this cycle to surge dynamics of glaciers
from other cluster regions, especially those of Variegated
Glacier in Alaska. The strong contrast in surge dynamics
leads us to suggest that there exist at least two distinct
surge mechanisms.
2. Location and Characteristics of Monacobreen
[5] Monacobreen (7924
0
N, 1234
0
E) is a 40-km-long,
tidewater glacier (Figure 1) that descends northward from
an elevation of 1250 m on the Isachsenfonna ice cap and
terminates in Liefdefjord. During the past century most
glaciers in Svalbard have been in retreat from their maximal
extent during the Little Ice Age, and aerial photographs
acquired by Norsk Polarinstitutt (NP) show that the margin
of Monacobreen retreated between 0.75 and 1.45 km
between 1966 and 1990 (NP photographs S66V-4286 and
S90 6568). The onset of heavy crevassing first noticed in
May 1992 (J. O. Hagen, personal communication, 2001),
together with marginal advance noted by comparing 1993
Landsat imagery with 1990 aerial photography, suggests
that Monacobreen had begun to surge between 1990 and
1992.
[6] Radio echo sounding (RES) can be used to infer
glacier thermal regime. RES suggests that Monacobreen is
polythermal because an internal reflecting horizon (IRH) at
adepthof120 to 200 m obscured any bed return,
suggesting that the glacier has a warm-based upper region
[Bamber, 1987].
[7] Monacobreen overlies two geological formations,
which are separated by a fault running approximately
NNW-SSE [Hjelle and Lauritzen, 1982]. The eastern part
of the glacier, including the entire width of the front, is
underlain by coarse conglomerates and sandstones. The
western parts are underlain by a combination of granitic
bands, gneisses, and schists. All of these geologic forma-
tions show various degrees of metamorphism and tecton-
ism. Monacobreen terminates in a fjord and its terminus is
probably underlain in part by Quaternary marine deposits.
3. Methods
3.1. Differential Interferometry
[8] We applied the usual procedure for differential inter-
ferometry [e.g., Joughin et al., 1996a, 1996b] using either
triplets from the 3-day repeat phases, or two pairs from the
tandem phase of ERS operation. In order to maximize phase
coherence between pairs of ERS SAR images, we avoided
periods with melting or precipitation and periods of strong
winds following recent snow (Table 1). To help phase
unwrapping of the differential topography only interfero-
grams, we used a digital elevation model (DEM) derived
from mapping by NP (digital versions of 1:100,000 map
Figure 1. Location of Monacobreen (box in insert) and of
meteorological station at Ny A
˚lesund (cross) in northern
Spitsbergen. The locations of the long profile and cross
profiles (X1–X6) are also shown. Profiles were extracted in
the direction of arrows. The long profile is taken along an
approximate flow line from 0 km at the 1995 margin to a
point on the plateau and has a total length of 33.5 km. A, B,
and C are the locations of points used in Figure 9 and are at
km 10.3, km 20.0, and km 30.6, respectively. Background
backscatter intensity image is from 28/29 December 1995.
An IRH was seen on a radio echo sounding [Bamber, 1987]
upglacier of X1, suggesting that this portion of the glacier
was warm based.
EPM 3 -2 MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM?
sheets A5, B5, and B6). We also employed this DEM to
determine ground control points beyond the margins of
Monacobreen and for baseline refinement. Six interfero-
metric DEMs for Monacobreen were produced (Table 1),
with an estimated vertical accuracy of only 20 to 40 m due
to the rugged topography and relatively low resolution of
the map-derived DEM used for ground control.
[9] In order to remove the topographic phase signal and
produce velocity-only interferograms we used the interfero-
metrically derived DEM closest in time to each interfero-
gram (Table 1). To help unwrap fringes, we used the
displacement maps from one date as a model of velocity
from other dates. Most velocity interferograms could be
phase unwrapped from a known zero reference (a nunatak in
the upper part of Monacobreen) to 7 km from glacier
front. In all eleven velocity maps were produced (Table 1).
3.2. Dual-Azimuth Processing
[10] In order to determine the full 3-D surface flow
direction, data from three look directions are required
[Joughin et al., 1998; Mohr et al., 1998]. ERS SAR
provides two look directions from the ascending and
descending passes of the satellite, and data from these can
be combined using dual-azimuth processing to produce 2-D
velocity fields. In the absence of in situ measurements, the
vertical component of flow can be estimated from a DEM
by assuming flow occurs parallel to the ice surface [e.g.,
Reeh et al., 1999]. This assumption gives us the third
dimension of the velocity field: the potential error associ-
ated with this assumption is discussed below.
[11] Data suitable for dual azimuth processing were only
available in winter 1995/1996 (Table 1). In order to com-
pute surface displacement maps for the other dates we
assumed the ice flow direction did not change during the
surge. In order to verify this assumption, we used coherence
tracking (explained below) [e.g., Rott et al., 1998; Strozzi et
al., 2002] between descending image pairs to estimate the
2-D horizontal component of the flow direction. This
analysis showed no significant change in flow direction
during the surge, albeit at a reduced resolution. Further-
more, two dual-azimuth snapshots during the surge of
Fridtjovbreen, Svalbard produced using the same techni-
ques show a mean absolute change in flow direction of only
1.2over 23 months despite an increase in velocity by 4 to 5
times [Murray et al., 2003].
3.3. Intensity and Coherence Tracking
[12] Surge velocities close to the front of Monacobreen
exceed those for which interferometric phase can be use-
fully interpreted. Therefore we applied a technique based on
tracking the optimum correlation between image patches
[Rott et al., 1998; Michel and Rignot, 1999; Murray et al.,
2002; Strozzi et al., 2002]. This technique can be applied
successfully only where displacements between images are
larger than 1 m and derives displacement information at a
much coarser resolution than interferometry. The procedure
can use either the optimization of intensity correlation or the
optimization of phase coherence within image patches.
[13] Intensity correlation tracking was used to derive
surface velocities from Monacobreen near its front where
displacement rates and chaotic surface change limited
interferometric coherence. Coherence tracking was used in
unwrapping the January and March 1994 data to tie a
disconnected island of coherence to a velocity reference.
Assuming an estimation error of 0.05 pixel in range and
azimuth, the displacement error from the tracking techni-
Table 1. Dates and Characteristics of Images and Interferometric Combinations Used to Determine the Flow
Evolution of the Surge
Image Dates Orbit Frame
Perpendicular
Baseline, m
Topography
Baseline,
a
m
Date of Resulting
Velocity Map
Descending Pass Images
4 – 7 Sept. 1991 E-708/E1-751 1989
b
184 #Sept. 1991
13 – 16 Oct. 1991 E1-1267/E1-1310 1989
b
82 160 Oct. 1991
16 – 19 Oct. 1991 E1-1310/E1-1353 1989
b
242
15 – 18 Nov. 1991 E1-1740/E1-1783 1989
b
163 "Nov. 1991
10 – 13 Feb. 1992 E1-2987/E1-3030 1989 87 #Feb. 1992
23 –26 March 1992 E1-3589/E1-3632 1989 74 137 March 1992
26 – 29 March 1992 E1-3632/E1-3675 1989 63
12 – 15 Jan. 1994 E1-13037/E1-13080 1971
b
177 93 Jan. 1994
15 – 18 Jan. 1994 E1-13080/E1-13123 1971
b
84
28 – 31 March 1994 E1-14112/E1-14155 1974
b
141 "March 1994
1 – 2 June 1995 E1-20280/E2-607 1981 2#June 1995
28 – 29 Dec. 1995 E1-23286/E2-316 1981 38 206 Dec. 1995
1 – 2 Feb. 1996 E1-23787/E2-4114 1981 168
30 April to 1 May 1996 E1-25061/E2-5388 1980
b
35 41 May 1996
c
3 – 4 May 1996 E1-25104/E2-5431 1976
b
6
Ascending Pass Images
26 – 27 Dec. 1995 E1-23262/E2-3589 1613
b
99 33 used with Dec. 1995
30 – 31 Jan. 1996 E1-23763-E2-4090 1613
b
132 descending images to
derive flow direction
8 – 9 Oct. 1997 E1-32595/E2-12922 1608
b
94 "Oct. 1997
a
Topography baseline is the equivalent perpendicular baseline for the topography component of the differential combination of
interferograms.
b
Indicates (and arrows) for a particular interferogram which other pair was used to remove the topographic signal.
c
Larger errors are expected in the velocity for May 1996, when the interferograms used in differential processing were formed
from images from different satellite tracks.
MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM? EPM 3 -3
ques is less than 1.2 m d
1
for tandem data and less than
0.4 m d
1
for SAR data with a 3-day repeat cycle [Strozzi et
al., 2002].
3.4. Potential Sources of Error
[14] There are no in situ field measurements of velocity
that can be used as ground truth for our SRI velocity
measurements. Joughin et al. [1998] and Fatland and
Lingle [1994] provide discussion of the sources of error in
3-D estimation of velocity using SRI. These sources include
systematic errors arising from changes in atmospheric water
content, errors in the estimates of the baseline (the perpen-
dicular distance between satellite positions at the time of
acquisition), unwrapping errors, and errors in the DEM, as
well as those arising from failure of the assumptions used in
processing [Goldstein, 1995; Rignot et al., 1996; Joughin et
al., 1999]. Changes in atmospheric water content are
expected to be small during high latitude winter, so that
the resulting uncertainty in phase is expected to be at most
±0.25 cycle [Gray et al., 1997]. This phase change is
equivalent to ±0.005 and ±0.02 m d
1
in the ground range
direction for 3-day and tandem data respectively. In this
study, baselines were calculated from precision orbit models
except for the ERS-1 1997 ascending data when ERS-1
precision orbits were not available. Uncertainty from this
source is expected to be less than ±0.01 m d
1
. Unwrapping
faults will lead to integer multiples of 2perrors in unwrap-
ped phase. These will lead to errors in velocity in multiples
of ±0.02 and ±0.07 m d
1
for 3-day and tandem data,
respectively, in the ground range direction. The magnitude
of the systematic errors in the flow rate can be assessed by
the maximum apparent surface displacement of rock areas
other than those used as a zero velocity reference. After
masking areas of ice and water using data from digital maps
(NP 1:100,000 map sheets A5, B5, and B6), mean displace-
ment values were 0.1 m d
1
for the 1997 ascending
mode data when precision orbits were not available, and
0.05 m d
1
for all other dates.
[15] The processing sequence we have used relies on
three assumptions:
[16] 1. Differential processing assumes that both the sur-
face topography and the displacement rate remain constant
between the images used. A change in topography between
interferograms will lead to a corresponding velocity error.
For a perpendicular baseline of 50 m, a relatively large
change in topography of 50 m between interferograms will
lead to an error of only 0.007 m d
1
[Joughin et al., 1996b],
which can be considered to be negligible in this study. If the
velocity has changed between the pairs of scenes used to
form differential interferograms, this will result in errors in
the separation of topography and velocity. As long as the
velocity varies systematically, the resulting error in velocity
cannot be greater than the change in velocity [Luckman et
al., 2002]. In general, the resulting error will be smaller for
triplets of images separated by 3 days and larger for tandem
data.
[17] 2. The assumption of surface parallel flow has been
commonly used to derive the vertical component of flow for
valley and outlet glaciers [e.g., Rignot et al., 1996; Vachon
et al., 1996; Mohr et al., 1998; Rott et al., 1998; Mattar et
al., 1998; Joughin et al., 1999]. However, this assumption
has limitations because flow is in general submergent in the
accumulation zone and emergent in the ablation zone [Reeh
et al., 1999]. In steady state, these vertical components are
equal to the local mass balance. Surge-type glaciers exist in
a state perpetually out of balance with the local climate. The
vertical component of flow can be accounted for if the ice
thickness is known [Reeh et al., 1999], which unfortunately
is not the case for Monacobreen. On the quiescent-phase
surge-type Black Rapids Glacier, Alaska, the surface paral-
lel assumption was shown to result in errors of up to 20%
or 0.03 m d
–1
in the magnitude of 3-D velocity [Rabus
and Fatland, 2000]. These errors could be largely removed
by using a seasonally adjusted vertical velocity of 0.002 to
0.003 m d
1
representing mass gain in the reservoir zone.
Our study is undertaken in the active phase when vertical
velocities could both be higher and have greater temporal
variation.
[18] In order to make some estimate of the likely error
resulting from the assumption of surface parallel flow, we
use vertical velocity measurements made during the active
phase of Bakaninbreen, Svalbard. Over the entire surge
duration of between 5 and 10 years at Bakaninbreen, ice
thinning in the reservoir zone was 15 m and thickening in
the receiving zone was 20 to 60 m [Murray et al., 1998], so
the rate of thickening or thinning varied from 0.008 to
+0.03 m d
1
. The ERS SAR configuration is much more
sensitive to vertical than horizontal displacements because
of the steep look angle. If we assume that the glacier surface
is flat, a vertical displacement d
v
would be misinterpreted as
a horizontal displacement of d
h
dh¼dv
tan qdcos jfjd
;ð1Þ
where j
d
and j
f
are the orientation angle of the SAR
observation, and ice flow directions respectively, and q
d
is
the incidence angle of the SAR observation. Vertical
velocity equivalent to values from Bakaninbreen would
result in a horizontal velocity error of up to ±0.07 m d
1
if the flow direction coincides with the satellite look
direction. The error will increase as the angle between the
incidence angle and flow direction increases. For an angle
between the flow direction and look direction of 30,the
error is ±0.08 m d
1
, and as this angle increases to 75the
error increases to about ±0.27 m d
1
.
[19] 3. As discussed earlier, because of the scarcity of
ascending scenes in the data archive, we assumed that the
flow direction did not vary during the surge. The magnitude
rof the three-dimensional displacement vector can be
computed by
r¼rd
sin qdcos ufcos jfjd
þcos qdsin uf
;ð2Þ
where r
d
is the magnitude of the displacement in the look
direction of the SAR. In order to make an estimate of error,
we again assume the glacier is flat, i.e., u
f
is zero. By
representing with jthe difference (j
f
j
d
) between flow
and look directions we obtain
r¼rd
sin qdcos j:ð3Þ
EPM 3 -4 MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM?
The error with respect to the difference between the flow
and look direction is therefore
@r¼rd
sin qd
sin j
cos2j@f;ð4Þ
and the relative error is
@r
r¼tan jðÞ@j:ð5Þ
Using equation (5), an estimate of the effect of a change in
the ice flow direction during the surge can be made. For a
5change in flow direction, the error is 5% when the
angle between the flow direction and look direction is 30.
As this angle increases to 75, the error increases to about
30%.
[20] Because the error in the calculated velocity increases
rapidly with increasing angle between ice flow direction and
look direction (discussion points 2 and 3 above), all data
within ±15of perpendicular to the look direction were
masked from the 3-D velocity maps when only descending
data were available.
4. Results
[21] The 34 ERS SAR scenes used in this study (Table 1)
allowed the computation of 18 interferograms (e.g., Figure 2).
The fringe visibility on these interferograms is strongly
related to the ice surface displacement field. For example, a
strip of ice from a tributary glacier on the eastern side of
Monacobreen was coherent right to the glacier front in the
early surge interferograms (arrow on Figure 2a), becoming
activated by the surge between March 1992 and January 1994
(Figures 2b and 2c). Clearly visible on the interferograms are
shear margins at each side of the glacier with spatially high
rates of change of displacement manifested as a high fringe
rate (Figures 2a–2c). In October 1997, the glacier was
coherent over virtually its whole surface suggesting that the
flow rate had dropped dramatically (Figure 2d).
[22] As discussed above, the direction of ice flow at all
times was assumed to be the same as the flow field for
winter 1995–1996, which was the only time for which dual-
azimuth data were available (Table 1 and Figure 3). Figure 4
shows eleven temporal snapshots of 3-D surface velocity on
Monacobreen between 1991 and 1997 superimposed on the
SAR intensity images, which have been calibrated to be
directly intercomparable in brightness. Any residual bright-
ness differences in these images will result from differing
local conditions between the acquisitions chosen (e.g.,
changes in snow water content or surface roughness). The
greatly increased brightness, particularly in 1994 (Figure 4),
almost certainly reflects increased surface roughness due to
intense crevassing at this time. Glacier velocity increased
progressively from 1991 to 1994 and then decreased,
showing an apparent temporary increase over the whole
glacier between December 1995 and May 1996 (Figure 4).
It should be noted that greater errors are expected in the
May measurement because the interferograms used in
differential processing were formed from images from
different satellite tracks (Table 1). Surface velocity measure-
ments are available only where coherence levels permitted
phase unwrapping (Figure 4). Velocity values are more
extensive during and after 1995 when tandem (1-day repeat)
data were available (Table 1), and the glacier was beginning
to slow down. This increase in coverage results because the
interferometric phase coherence is increased by decreases in
either velocity or repeat-pass delay.
[23] The spatial differences in the velocity fields are most
easily compared using profiles oriented upglacier (Figure 5)
and cross-glacier (Figure 6). The main velocity increase
Figure 2. Selected geocoded interferograms over Mon-
acobreen. Baselines are given in Table 1. These interfero-
grams contain information on both topography and
displacement in the look direction of the sensor. Hue
indicates interferometric phase with color intensity modu-
lated by the SAR backscatter. Interferograms formed from
scenes in June 1995 and May 1996 have a very short
topographic baseline and are therefore virtually insensitive
to topography. In northern Svalbard, the look direction is
(a–c) 126for descending data and (d) 238for ascending
data at an incidence angle of 23to the vertical.
MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM? EPM 3 -5
occurred uniformly, initially in the lower part of the glacier,
with the velocity more than doubling between September
1991 and March 1992 (Figure 5a). This initial increase in
velocity predominantly affected the bend in the glacier and
areas downglacier of km 22. (This notation refers to the
location on the flow line shown in Figure 1 with km 0 being
the December 1995 margin; for example, km 22 refers to a
location 22 km upglacier of km 0). The velocity in the upper
part of the glacier did not increase substantially until
January 1994 (Figure 5a). There was no indication of a
surge front or activation wave travelling downglacier, and if
this occurred, it did so before September 1991. The cross
profiles (Figure 6) showed generally flat cross sections.
Only X6, located farthest upglacier, showed a parabolic
cross section. Finally, while the velocity increased dramat-
ically during the surge, the pattern of velocity, especially the
locations of local maxima and minima, remained remark-
ably consistent throughout the surge.
[24] We could not use interferometry to map the velocity
over the entire glacier surface because coherence was lost
close to the sides and at the front. The problem was most
acute when the velocity was greatest (Figures 2 and 4).
Intensity correlation tracking shows that the velocity
increased very rapidly toward the glacier front (approxi-
mately downglacier of km 8.5) to velocities over 5md
1
(Figure 7).
[25] The longitudinal velocity profiles allow calculation
of strain rates (Figure 8). During the surge, strain rates on
Monacobreen increased to 0.001 d
1
(0.37 yr
1
), which is
greater than the rates of <0.1 yr
1
typical of most non-
surging glaciers [Kamb et al., 1985]. While strain rates
increased as the velocity increased, in general, many loca-
tions on the flow line that were locally extensile or com-
pressive remained consistent in space (Figure 8).
[26] From these results, we can also investigate the
temporal evolution of the glacier surge. Figure 9a shows
how the velocity at three locations on the glacier surface (A,
B, and C in Figure 1) changed during the six years of
observation. These three sample points were chosen as they
have valid measurements from every differential observa-
tion and as they are representative of the whole glacier. The
velocity increased steadily between September 1991 and
March 1992 with a remarkably constant rate of change
during this period. We have therefore fit straight lines to the
data by least squares regression, and for these lines the
coefficients of determination (R
2
) vary between 0.97 and
1.00. Between January 1994 and October 1997, the velocity
dropped, again approximately linearly, and although the R
2
values are lower, they are still remarkably high (0.93 to
0.97). The gradient of these lines gives the acceleration or
deceleration of the glacier at each location. This graph
shows that the switch between surge acceleration and
deceleration occurred during the period when there are no
data available (after the second ice phase of ERS-1 and
before the tandem mission), and strongly suggests that
maximum flow rates occurred during 1993.
[27] This temporal analysis is extended to every point on
the flow line in Figure 9b, which shows the acceleration,
deceleration and associated coefficients of determination.
During surge acceleration (September 1991 to March 1992),
the coefficient of determination was close to 1.0 down-
glacier from km 28 suggesting that the velocity was
increasing systematically. Upglacier from km28, the line-
arity of the relationship breaks down. The rate of acceler-
ation increases from km 28 to km 18, downglacier from
which point the pattern of acceleration peaks and troughs
corresponds to the pattern of velocity peaks and troughs:
regions of higher velocity experience greater velocity
increases than regions of lower velocity. During surge
deceleration (January 1994 to October 1997), the coefficient
of determination was lower but was still greater than 0.8
downglacier of km 27. In general, the deceleration rate
Figure 3. The 3-D velocity derived from winter 1995–
1996 dual-azimuth differential interferometry shown as
color (hue) modulated by SAR backscatter (intensity). Ice
flow direction for all other times was taken from these
results. Velocity arrows are displayed every 750 m. Surface
velocity measurements are available only where coherence
levels permitted phase unwrapping. Background image is
SAR backscatter.
EPM 3 -6 MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM?
Figure 4. Plan view velocity maps from September 1991 to October 1997 showing 24 x 45 km area of
Monacobreen and its surroundings. Where surface displacement rates have been retrieved, hue indicates
velocity (see color scale) with color intensity modulated by the SAR backscatter. Background image is
SAR backscatter, calibrated so that intensity is directly comparable throughout the sequence.
MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM? EPM 3 -7
increases steadily downglacier and the deceleration rate was
approximately one third the acceleration rate.
[28] SAR intensity images at approximately annual inter-
vals were used to track the terminus position (Figure 10).
The glacier advanced 2 km after September 1991 and was
at maximum extent in April 1996. Between April 1996 and
October 1997 the central section retreated more than 250 m,
while the western and eastern sections still advanced 100
m. Between October 1997 and September 1998 the entire
front margin was in retreat. The maximum measured
advance rate of the glacier terminus was during winter
1991–1992, and this preceded the velocity peak in January
1994 (Figures 5 and 10). Terminus retreat started once the
ice flow velocity fell to 2md
1
close to the glacier front.
5. Discussion
5.1. Velocity During Quiescence
[29] An approximation of the velocity during the quies-
cent phase on Monacobreen can be made by assuming
glacier flow is due solely to ice creep. In this case, the ice
creep velocity u
q
may be approximated by
uq¼tn
dH2A=nþ1ðÞ;ð6Þ
where t
d
is the basal shear stress, n= 3 and A=2.410
15
s
1
kPa
3
at 2C are flow law parameters, and His the ice
thickness [Paterson, 1994, p. 251]. If we assume the basal
shear stress is equal to the gravitational driving stress then
td¼FrigH sin a;ð7Þ
where Fis a valley shape factor, r
i
is the density of ice, gis
the gravitational constant, and ais the ice surface slope. We
do not have any ice thickness measurements for Monacob-
reen as the RES study failed to penetrate to the glacier bed
[Bamber, 1987]. An estimate, however, can be made using
an empirical relationship between glacier thickness and
surface area, S,
H¼33 ln Sþ25;ð8Þ
which is valid for Svalbard outlet glaciers with surface areas
greater than 1 km
2
[Hagen et al., 1993]. Monacobreen has a
surface area of 408 km
2
[Hagen et al., 1993] and hence a
predicted thickness of 225 m. The shape factor, F,is
therefore 1[Paterson, 1994, p. 269]. The average surface
slope is 2, so the gravitational driving stress is 69 kPa
(equation (7)) and, using equation 6 with the full range of
ice densities quoted by Paterson [1994, p. 9], the creep
velocity is 2.2 to 3.0 m yr
1
(0.006 to 0.008 m d
1
).
[30] These values are significantly less than the lowest
velocity we measured on Monacobreen. However, they
are comparable to those measured on Svalbard surge-type
glaciers during quiescence such as Bakaninbreen (0.2 to
0.4 m yr
1
)[Murray et al., 1998], Finsterwalderbreen
(1to5myr
1
)[Nuttall et al., 1997], Kongsvegen
(2.9 m yr
1
)[Melvold and Hagen, 1998], and Sef-
strømbreen (2myr
1
)[Liestøl et al., 1980]. The
calculation suggests that the increase in velocity between
the quiescent and active phase at Monacobreen is 100 to
1000 times. The very large contrast in velocity between the
calculated creep velocity and our measured velocities sug-
gest that basal motion (sliding or sediment deformation)
dominates flow during the active phase of glacier surging.
5.2. Characteristics of the Surge of Monacobreen
5.2.1. Acceleration and Deceleration Phases
[31] Figure 9a shows two phases during Monacobreen’s
surge: namely a months-long acceleration phase (September
1991 to March 1992), and longer and more progressive
deceleration phase. We do not have data from this study to
observe transitions between either of these states and quies-
cence, nor between acceleration and deceleration.
Figure 5. Time series of velocity from interferometry along long profile of Monacobreen (location in
Figure 1). (a) Progressive velocity increase from September 1991 to January 1994 and (b) progressive
velocity decrease from January 1994 to October 1997, with a temporary possible velocity increase
between December 1995 and May 1996. Gaps in the data indicate regions that were incoherent and could
not be unwrapped in the interferometric image at the location of the transect or where the flow direction
approached an angle perpendicular to the line of sight. X1–X6 indicate the locations of the cross profiles
presented in Figure 6.
EPM 3 -8 MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM?
[32] Our observations span only those months with low
surface temperatures because of the need to maintain
coherence. This sampling strategy means that any velocity
variations that occur during the summer months may not
be observed. However, the increase of velocity between
September 1991 and March 1992 is monotonic (Figure 9).
The decrease between January 1994 and October 1997 is
broken only by a temporary apparent increase in flow rate
between December 1995 and May 1996. Prior to the May
1996 images, surface temperatures had been below zero
(mean –5.4C) for the previous 19 days (data collected in
Ny A
˚lesund by Norske Meteorlogiske Institutt), so this
variation is unlikely to have been driven by surface melt.
We believe that short-lived or seasonal velocity variations
are subdued compared to the overall velocity cycle of
acceleration and deceleration we have revealed.
5.2.2. Spatial Pattern of Velocity During Active Phase
[33] We have no evidence of a surge front propagating
downglacier on Monacobreen. Such a surge front would be
expected to be marked by closely spaced fringes in inter-
ferograms or by incoherence because of the high spatial rate
of change in velocity. A surge front could have reached the
glacier front before our observations began, but this seems
unlikely given the relatively low ice velocities measured.
Rather, the surge appeared to start simultaneously over the
entire lower region of the glacier and then to propagate
upglacier. Two other studies of Svalbard surges also report
no surge front. At both Osbornebreen [Rolstad et al., 1997]
and Fridtjovbreen [Murray et al., 2003] the surge is reported
to have started at the glacier terminus and to have propa-
gated upglacier. In contrast, Bakaninbreen had a clear surge
front [Murray et al., 1998], and at Usherbreen crevassing
started in the upper part of the glacier and propagated
downglacier [Hagen, 1987]. We note that those glaciers
where a surge front has been reported are land terminating,
whereas glaciers without reported surge fronts terminate in
water deep enough to calve at their termini (Table 2).
5.2.3. Static Nature of Velocity Pattern in Space
[34] Surge velocities on Monacobreen were as much as 5.0
md
1
(Figures 4–6), and in general increased downglacier,
especially close to the glacier margin (Figures 5 and 7). On
other Svalbard glaciers reported surge velocity magnitudes
are similar (Table 2). We noted earlier that the local pattern of
both velocity and strain rate at Monacobreen are remarkably
consistent throughout the surge (Figures 5 and 8), so that
whatever controls the spatial distribution of velocity must
remain fixed throughout the surge phase. Furthermore, as the
flow rate increases the difference between velocity highs and
lows also increases. In other words, those regions that are
‘‘sticky spots’’ at low velocity became relatively stickier, and
the ‘‘slippery spots’’ became more slippery as the velocity
increased. The spatial pattern of surge velocity has been
observed for only a very few glaciers in Svalbard. At
Osbornebreen, tracking of crevasses on Landsat and SPOT
imagery during 1988 showed generally extensile flow with
velocities varying between 0.0 to 6.0 m d
1
[Rolstad et al.,
1997].
[35] Controls on glacier flow fall into two groups, those
which would be expected to change during a surge (e.g.,
thermal structure, basal sediment composition, the effect of
tributary glaciers, structure of the basal water system and
input of surface water), and those that would remain
constant (e.g., bedrock features and valley shape). The
wavelength of spatial variation in velocity and strain rate
is much less than variations in overall valley shape (com-
pare Figures 1, 5, and 8). Furthermore, the variations in
Figure 6. Time series of cross glacier velocity profiles
(locations in Figure 1). (left) Increasing velocities (Septem-
ber 1991 to January 94) and (right) the decreasing velocities
(March 1994 to October 1997). Transects were chosen
approximately perpendicular to the flow line and extend,
where possible, from valley side to valley side. Dark gray
zones indicate rock, and light gray zones indicate a different
flow unit or tributary. Gaps in the data indicate regions that
were incoherent and could not be unwrapped in the
interferometric image or where the flow direction ap-
proached an angle perpendicular to the line of sight. The
dotted line indicates the intersection with the long profile.
MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM? EPM 3 -9
velocity and strain rate appear to be regularly spaced,
whereas variations in valley shape are irregularly spaced.
We therefore believe that the main controls on the small-
scale glacier velocity are basal bedrock features.
5.2.4. Surge Duration
[36] We do not believe that we captured the full surge
with these SRI observations. This assertion is based on the
following: (1) there is no change of shape in the velocity
cross profiles with time, and hence no change in the
dominant flow mechanism over our observation period;
(2) there is no evidence of a switch on or off in either the
flow rate or the acceleration or deceleration; (3) the
estimated quiescent-phase velocity is much lower than
our observed velocities; and (4) the glacier was already
advancing at the start of our observations. We therefore
believe that the active phase was longer than the 73
months (more than six years) of our observations. Extrap-
olation of the linear trends in Figure 9a to the calculated
quiescent velocities suggests that the surge started between
early January and late May 1991 and ended between
November 1998 and January 2002. In early May 1991,
there were no significant crevasses on Monacobreen,
which suggests that strain rates were quite low at this
time, whereas the following year it was heavily crevassed
(J. O. Hagen, personal communication, 2001). Therefore
our estimate of the active phase duration is 89 to 133
months (7 to 11 years), which is in keeping with obser-
vations of active phases (3 –10 years) on other Svalbard
glaciers [Dowdeswell et al., 1991].
5.3. A Typical Svalbard Surge Cycle
[37] In Figure 11a we have used the results from this
study, and results from the literature to reconstruct glacier
dynamics during a ‘‘typical Svalbard glacier surge cycle.’’
Overall cycle lengths in the archipelago are thought to be
50–500 years [Dowdeswell et al., 1991] but are not well
Figure 7. Time series of velocity from differential SAR interferometry (solid line) and intensity
correlation tracking (open circles). A, acceleration phase; D, deceleration phase. January 1994 is the
fastest measured flow rate. Intensity correlation tracking data are displayed from km 25 to the ice margin
(variable over the years) and give velocity data where SRI measurements are missing due to regions of
poor coherence or where the flow direction approached an angle perpendicular to the line of sight.
EPM 3 -10 MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM?
known as only five glaciers have been observed to surge
twice. Observed surge intervals have been between 40
(Tunabreen) and 133 years (Fridtjovbreen) [Hagen et al.,
1993]. It is likely that the duration of the average surge
cycle in Svalbard is longer than these intervals. The surge
starts with a years-long period of steady acceleration, which
was measured at Fridtjovbreen using similar methods to this
study [Murray et al., 2003]. This is followed by a months-
long period of relatively rapid acceleration that we see at
both Monacobreen (Figure 9) and Fridtjovbreen [Murray et
al., 2003]. The length of the active phase is typically 3 – 10
years [Dowdeswell et al., 1991], and was probably between
7 and 11 years at Monacobreen. The end of the fast flow
phase is very gradual, with velocity decreasing over a years-
long period (Figure 9). The characteristic of a much longer
and more progressive termination on Svalbard glaciers is
confirmed by data from Kongsvegen, where flow rates 20
years after the surge started were still 1md
1
, whereas
after 45 years they were 0.008 m d
1
[Melvold and
Hagen, 1998].
5.4. Contrasts With Surges of Glaciers in Other
Regions, Especially Those of Variegated Glacier
[38] The best studied glacier surge is probably the 1982 –
1983 surge of Variegated Glacier in Alaska [e.g., Kamb et
al., 1985] (Table 2), and the dynamics of this surge are
Figure 8. Profiles of longitudinal strain rate along long
profile of Monacobreen for surge acceleration (A) and
deceleration (D). Profile location is shown in Figure 1. The
dashed and numbered lines 1 to 8 show the location of
prominent compressional strain rate anomalies in early
surge. Most anomalies remain throughout the surge,
although amplitudes change, e.g., 2, 3, 6, 7, 8; others switch
from compressional to extensile, e.g., 4, 5; others switch off
and reoccur later, e.g., 1.
Figure 9. (a) Surface velocity as a function of time for
positions A, B, and C on the glacier (Figure 1). Gray areas
indicate times when data suitable for SRI were not
available. Straight-line fits show the linearity of the rate
of change of velocity for both the acceleration and the
deceleration phases of the surge. (b) Rate of change of
velocity with respect to time and associated coefficient of
determination (R
2
) for the entire flow line shown in Figure
1. Gray indicates surge acceleration (September 1991 to
March 1992). Black indicates surge deceleration (January
1994 to October 1997). See text for full explanation.
MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM? EPM 3 -11
summarized in Figure 11b. Its surge was characterized by
extreme basal water pressures and a highly tortuous, high
volume basal water system that had low water throughflow
rates, postulated to be a linked cavity system [Kamb et al.,
1985; Kamb, 1987]. The surge occurred in two phases, the
first beginning in the winter of 1982 in the upper part of
Variegated Glacier, which accelerated to a maximum in
late June 1982 before slowing abruptly over a few hours.
The velocity then decreased until August/September,
although there were major pulses of increased motion
superimposed on this decelerating trend. The second phase
began in the upper part of the glacier in October 1982 and
propagated progressively downglacier activating the lower
region of the glacier. Surge velocities were 10 to 50 m
d
1
and strain rates at the surge front were 0.2 d
1
.
Evidence for the nature of the basal water system during
the surge came from dye tracing [Kamb et al., 1985].
During the surge, dye transit velocities were slow (0.025
ms
1
) and the dye appeared at the margin at a number of
outlet streams. After the surge, the dye appeared rapidly
(transported at a velocity of 0.7 m s
1
) at a single outlet
stream.
[39] The surge terminated extremely rapidly over a few
hours on 4 July 1983, and by September the whole glacier
was moving at less than presurge velocities. The termina-
tion was marked by an outburst flood of turbid water, an
abrupt drop in the pressure in the basal water system, and a
drop in the glacier surface of 0.1 m. This very rapid
termination is inferred to have resulted from the collapse of
basal cavities and formation of an efficient tunnel based
drainage system. Similarly rapid surge terminations (days)
were seen at Medvezhiy Glacier, Pamirs [Osipova and
Tsvetkov, 1991] and West Fork Glacier, Alaska [Harrison
Table 2. Comparison of the Surge of Monacobreen With Other Surges
a
Glacier
Length,
km
Area,
km
2
Surge
Date
Surge
Duration, years
Quiescence
Duration, years
Surge
Advance, km
Drawdown,
m
Thickening,
m
Basal Shear
Stress, kPa
Quiescent
Velocity, m d
1
Typical Surge
Velocity, m d
1
Surge Front
Propagation,
md
1
Source
Variegated 20 29 five times in
20th century
2 16 – 26 2– 5 50 110 160 – 180 0.1 – 1.0 50 23 – 80 1
Finsterwalder 11 45 1900s 1.5 50 100 40 – 95 0.002 – 0.03 2
Fridtjov
b
13 49 1860s, 1990s 4+ 130 >2.5 105 0.08– 0.3 2.0 – 3.3 no front 3, 4
Usher 13 69 1980s 7 1.5 40 >100 33 – 61 1.5 – 4.3 possible
front
5, 6
Bodley
b
16 87 1970 – 1980s 4 – 13 2.7 10 – 15 10– 60 0.05 – 0.5 2.0 – 2.6 5, 7
Bakanin 17 61 1980 – 1990s 5 – 10 >85 0 15+ 20 – 60 52 – 66 0.001 1 – 3 0.9 – 4.5 8
Osborne
b
20 152 1980s 4+ 2 100 100 1.2 – 6.0 no front 5, 9
Sefstro¨m
b
23 155 1880s 6.5 0.005 1.3 10
Kongsvegen 26 102 before 1948 1.5 – 2.0 55 – 85 0.008 >1.0 11
Tuna
b
35 203 1930, 1970 50 1.5 12
Monaco
b
42 408 1990s 6+ 2 70 0.005– 0.008 0.5 – 5.0 no front
Hinlopen
b
68 1250 1970s 4+ 3 14 – 16 6
a
All glaciers except Variegated Glacier are located in Svalbard.
b
Glaciers terminate in water sufficiently deep to calve at their terminus.
Data sources: 1, Kamb et al. [1985], Raymond et al. [1987], and Wilbur [1988]; 2, Nuttall et al. [1997]; 3, A.-M. Nuttall (personal communication, 2000); 4, Murray et al. [2002b]; 5, Dowdeswell et al. [1991]; 6, Hagen
[1987]; 7, Dowdeswell and Collin [1990]; 8, Murray et al. [1998]; 9, Rolstad et al. [1997]; 10, Boulton et al. [1996]; 11, Melvold and Hagen [1998]; 12, Hodgkins and Dowdeswell [1994].
Figure 10. Position of the glacier terminus between
September 1991 and September 1998 obtained from
tracking six points along a 3.5 km west to east transect.
Displacements are estimated to be accurate to ±1 pixel or
about ±30 m.
EPM 3 -12 MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM?
et al., 1994]: all are in stark contrast to the prolonged
terminations of surging observed in Svalbard. At West Fork
Glacier, termination was associated with floods of turbid
water. Large outburst floods also marked the termination of
fast flow of Bering Glacier, Alaska [Molnia, 1994], and
water pressures during the surge were extremely high, as
evidenced by the presence of pressurized basally derived
water in crevasses [Herzfeld and Mayer, 1997]. Turbid
water was also seen in supraglacial lakes and marginal
crevasses during the surge of Peters Glacier, Alaska [Echel-
meyer et al., 1987]. The surge of Sortebræ, east Greenland,
was characterized by very high ice velocities, extreme basal
water pressures, and a rapid termination [Murray et al.,
2002]. All these surges had ice dynamics similar to those of
Variegated Glacier.
[40] Figure 11 summarizes the major contrasts in dynam-
ics between the surges of Svalbard glaciers and those of
Variegated Glacier. Flow velocities on Variegated Glacier
were an order of magnitude higher and strain rates two
orders of magnitude higher than those measured on Mon-
acobreen. On Variegated Glacier, there would appear to be
two states: surging, when the glacier is underlain by a linked
cavity hydraulic system, and quiescent, when the glacier is
underlain by a channelled water system. These states are in
contrast to the three states inferred for Monacobreen.
Variegated Glacier switches rapidly between these two
states, although the switch from slow to fast flow (weeks
to months) is less rapid than the switch from fast to slow
flow (two days) (Figure 11a). In contrast, on Monacobreen
acceleration was more rapid than deceleration. During the
surge of Variegated, the flow rate was highly variable,
whereas on Monacobreen we have no evidence for signifi-
cant short-lived velocity variations.
5.5. Evidence Against the Linked Cavity
Mechanism of Surging Operating in Svalbard
[41] The linked cavity mechanism provides a good model
for surging at many Alaskan glaciers. We do not believe
however that the linked cavity mechanism is applicable to
the surge of Monacobreen or other Svalbard surges for the
following reasons:
[42] 1. On Variegated Glacier, the glacier stopped almost
instantaneously, coinciding with the release of a large
amount of turbid and therefore probably basal water. Kamb
[1987] explained the termination of glacier surging of
Variegated as resulting from the collapse of a linked cavity
water system into a channeled water system. The very slow
surge termination on Monacobreen and other Svalbard
glaciers suggests that the termination mechanism in the
archipelago is fundamentally different from that of Varie-
gated Glacier.
[43] 2. There are no reports of basal water in marginal
lakes or water-filled crevasses during surges of Svalbard
glaciers. Such lakes or water-filled crevasses are used to
support assertions of basal water systems that are large
volume and high water pressure (i.e., have characteristics
similar to linked cavity systems) beneath surging glaciers in
other regions [Lingle et al., 1994; Reeh et al., 1994; Jiskoot
et al., 2001].
[44] 3. There is no evidence of short-lived, large-scale
velocity variations during Svalbard surges. Such velocity
variations appear to be characteristic of the surge of Varie-
gated Glacier. Any such velocity variations are subdued on
Svalbard glaciers.
[45] 4. The linked cavity theory predicts that glaciers with
low slope are more likely to be of surge-type [Kamb, 1987;
Clarke, 1991]. Statistical studies have shown that in Sval-
bard, surge-type glaciers tend to be steeper than normal
glaciers [Jiskoot et al., 1998].
[46] Given the strong evidence for the linked cavity
mechanism of surging for Variegated Glacier and that
against its operation in Svalbard, we suggest that there exist
at least two different surge mechanisms.
5.6. Possible Surge Mechanism in Svalbard
[47] Our observations of Monacobreen allow inferences
to be made about basal conditions during the surge event.
The glacier is polythermal and was most likely warm based
over the majority of its length during our observations. The
glacier is probably soft bedded, but bedrock features
control its large-scale velocity pattern. These basal con-
ditions are probably common beneath Svalbard surge-type
glaciers. Statistical studies have shown that such glaciers
tend to be polythermal and to overlie sedimentary bedrock
[Hamilton and Dowdeswell, 1996; Jiskoot et al., 1998,
2000]. The results of these studies, together with field
observations of polythermal regimes [Murray et al., 2000;
Smith et al., 2002] and deforming sedimentary beds [e.g.,
Porter and Murray, 2001] beneath Svalbard surge-type
glaciers, are thought to support a thermally regulated soft
bed surge mechanism in Svalbard. Similar basal conditions
also occur beneath Trapridge Glacier, Yukon Territory [e.g.,
Figure 11. (a) Schematic summary of the dynamics of
glacier surging on Monacobreen and other Svalbard
glaciers. Bold lines are the data from this study, the dotted
line showing the early slow (years-long) part of the
acceleration phase is taken from a related study of
Fridtjovbreen [Murray et al., 2003], and K1 and K2 are
velocity measurements 20 years and 45 years after the
start of the surge of Kongsvegen [Melvold and Hagen,
1998]. (b) Variegated Glacier, a surge thought to be
promoted by linked cavity formation and terminated by
their collapse (data from Kamb et al. [1985]). Quiescence is
characterized by seasonal cycles in velocity with faster flow
in summer than winter [Raymond and Harrison, 1988] and
minisurges [Kamb and Engelhardt, 1987], neither of which
are shown on this figure. The initiation of the surge of
Variegated Glacier was less abrupt than the termination,
whereas the opposite is true on Monacobreen.
MURRAY ET AL.: IS THERE A SINGLE SURGE MECHANISM? EPM 3 -13
Clarke et al., 1984; Clarke and Blake, 1991]; however, no
observations have yet been made at this glacier in its active
phase.
[48]Fowler et al. [2001] present a model of thermally
regulated surging over a soft bed that we believe may be
applicable to our observations at Monacobreen as well as
other Svalbard glaciers: (1) At the start of the surge cycle
(quiescence) the glacier is cold based or polythermal. In the
former case, ice builds up until the pressure melting point is
reached over some portion of the glacier bed. (2) Once the
pressure melting point is reached, basal meltwater is pro-
duced and leads to elevated pore water pressures and
weakening of the underlying till. (3) These increased pore
water pressures cause deformation and hence dilation of the
till, increasing water storage and further weakening it. (4)
Accelerated deformation of the basal till produces frictional
heating, which results in further melting and raises pore
water pressures at the glacier bed. (5) The positive feedback
between basal motion and meltwater production continues
resulting in rapid basal motion. (6) This basal motion
continues causing ice thinning until increased heat loss
results in refreezing at the glacier bed.
[49] The critical model parameters are the thickness and
permeability of unfrozen till [Fowler et al., 2001], which
together govern the efficiency of meltwater evacuation from
the bed, and thus control basal water pressure. If the
thickness and permeability are both low, basal drainage is
impeded, effective pressure can reach zero and surging can
occur. If the till thickness and permeability are close to, but
above threshold values, dynamic oscillations can occur, but
without a very rapid flow phase. The rate of surge onset is
primarily controlled by the thickness of unfrozen till,
because this controls the volume of meltwater required to
reach overburden pressure, while the rate of ice flow during
a surge is controlled by the bed roughness or by friction at
the glacier sides.
[50] Once the basal water pressure rises and the effective
pressure reaches zero, the water is conceived of as being
stored in ‘‘blisters’’ formed at the ice-bed interface [Fowler
et al., 2001]. If basal water pressure is high enough to cause
decoupling of the glacier from its bed or the opening of
faults through the ice, discharge of water from these blisters
will occur. The blisters are not supposed to play a signifi-
cant role in promoting glacier flow, although discharge from
them is significant during surge termination. At a certain ice
thickness, blister discharge becomes sufficiently high to
reestablish normal effective pressures. In general, the rate
of surge termination is controlled by ice thinning, which in
turn is controlled by the rate of transport of ice downglacier
(i.e., slow surges should have slow terminations).
[51] In this model, the fronts between surging and non-
surging ice are seen as propagating both up and downglacier
as a ‘‘thermal activation wave’’, i.e., the transition between
cold and warm ice [cf. Murray et al., 2000]. Two types of
surge are suggested by the Fowler et al. [2001] model. In
the first, the propagation of the activation wave is slow
relative to the ice speed, and this causes a wall of ice to
propagate downglacier (as at Bakaninbreen and Usher-
breen). In other words, the thermal activation front is also
the surge front. If the propagation of the activation wave is
faster than ice flow, the surge is gentler. Lack of a surge
front (as observed at Monacobreen) can result if there is no
restriction to flow at the margin. This can occur if the
glacier base is at the melting point all the way to the
terminus, and is even more likely if the glacier terminates
in tidewater sufficiently deep to calve [c.f., Hodgkins and
Dowdeswell, 1994]. Hence the model may be able to
explain the differences observed between tidewater and
land terminating surging glaciers in Svalbard.
[52] We believe that our observations started somewhere
between steps 2 and 4 of the cycle described above and
continue into step 6 but not into full quiescence. Notable
aspects of the model with respect to our observations are
that (1) the rate of flow during the surge is controlled by
the form drag at the bed (this is in keeping with our
inference that sticky spots at Monacobreen result from
bedrock obstructions and that these become increasingly
important as the flow rate increases) and (2) the model can
explain the lack of an observed surge front, the progressive
surge initiation, prolonged active phase duration and slow
termination.
6. Conclusions
[53] Using a combination of differential interferometry
and intensity correlation tracking we have elucidated, in
unprecedented detail, the spatial and temporal development
of glacier dynamics during a surge of Monacobreen. We
infer a three-phase surge cycle: acceleration, deceleration
and quiescence. The 6-year SRI sequence of velocity
measurements covers part of the acceleration and deceler-
ation phases. Surge accceleration was approximately three
times more rapid than deceleration. Surge dynamics at
Monacobreen are in marked contrast to the observed surge
of Variegated Glacier in Alaska [Kamb et al., 1985], which
started relatively progressively and terminated over a period
of only a few hours. The contrast in dynamics of Mon-
acobreen and other Svalbard glaciers with Variegated Gla-
cier make it very unlikely that the same surge mechanisms
are involved. We therefore hypothesize that there exist at
least two markedly different types of glacier surges.
[54]Acknowledgments. ESA provided the ERS images under project
AO3-283. Meteorological data were kindly supplied by the Norske
Meteorlogiske Institutt and digital map data by NP. This project is funded
by the UK Natural Environment Research Council (GST/02/2192 and F14/
6/37). We thank U. Wegmu¨ller and D. Vaughan for their contributions to the
analysis and discussion, B. Kulessa and A. Fowler for comments on the text,
and J.O. Hagen for drawing our attention to the surge of Monacobreen. The
referees (R. LeB. Hooke and J.O. Hagen) and associate editor (T. Pfeffer)
have worked hard to improve this paper, for which we thank them.
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