TCThe CryosphereTCThe Cryosphere1994-0424Copernicus PublicationsGöttingen, Germany10.5194/tc-10-1317-2016Accelerating retreat and high-elevation thinning of glaciers in central
SpitsbergenMałeckiJakubmalecki.jk@gmail.comhttps://orcid.org/0000-0002-1338-5232Cryosphere Research Department, Adam Mickiewicz University, Poznań,
PolandJakub Małecki (malecki.jk@gmail.com)24June20161031317132925September20156November20152June201613June2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://tc.copernicus.org/articles/10/1317/2016/tc-10-1317-2016.htmlThe full text article is available as a PDF file from https://tc.copernicus.org/articles/10/1317/2016/tc-10-1317-2016.pdf
Svalbard is a heavily glacier-covered archipelago in the Arctic.
Dickson Land (DL), in the central part of the largest island, Spitsbergen, is
relatively arid and, as a result, glaciers there are relatively small and
restricted mostly to valleys and cirques. This study presents a comprehensive
analysis of glacier changes in DL based on inventories compiled from
topographic maps and digital elevation models for the Little Ice Age (LIA) maximum, the 1960s, 1990, and 2009/2011. Total glacier area has decreased by
∼ 38 % since the LIA maximum, and front retreat increased over
the study period. Recently, most of the local glaciers have been consistently
thinning in all elevation bands, in contrast to larger Svalbard ice masses
which remain closer to balance. The mean 1990–2009/2011 geodetic mass
balance of glaciers in DL is among the most negative from the Svalbard
regional means known from the literature.
Introduction
Small glaciers are natural indicators of climate, as they record even slight
oscillations via changes of their thickness, length, and area (Oerlemans,
2005). Twentieth century climate warming caused a volume loss of ice masses
on a global scale (IPCC, 2013), contributing to about half of the recent
rates of sea-level rise. Despite the relatively small area of glaciers and
ice caps, their freshwater input to sea-level rise is of similar magnitude
to that from the largest ice masses in the world: the Antarctic and Greenland
ice sheets (Radić and Hock, 2011; Gardner et al., 2013). Therefore, it is
of great importance to study the volume changes of all land ice masses in
both hemispheres.
The archipelago of Svalbard is one of the most significant arctic
repositories of terrestrial ice. Glaciers and ice caps cover 57 % of the
islands (34 × 103 km2) and have a total volume of
7 × 103 km3 (Nuth et al., 2013; Martín-Españyol
et al., 2015). It is located in close proximity to the warm West Spitsbergen
Current, and ice masses there are considered to be sensitive to changes in
climate and ocean circulation (Hagen et al., 2003). The climate record
suggests a sharp air temperature increase on Svalbard in the early 20th century,
terminating the Little Ice Age (LIA) period around the 1920s (Hagen et al.,
2003). A cooler period between the 1940s and 1960s was followed by a strongly
positive summer temperature trend, i.e. 0.7 ∘C decade-1 for
the period 1990–2010 (Førland et al., 2011; James et al., 2012; Nordli et
al., 2014). Climate warming led to volume loss of the Svalbard glaciers
(although with large spatial variability), particularly after 1990 (Hagen et
al., 2003; Kohler et al., 2007; Sobota, 2007; Nuth et al., 2007, 2010, 2013;
Moholdt et al., 2010; James et al., 2012).
Location of the study area. (a) Map of Svalbard with locations of
regions of central Spitsbergen: Dickson Land (DL), Nordenskiöld Land
(NL), and Bünsow Land (BL). (b) Map of Dickson Land and its subregions:
north (DL-N), central (DL-C), and south (DL-S). Glaciers coloured with grey
in the eastern part of DL-C are not covered by 1990 digital elevation model.
Strong climatic gradients over the archipelago are an important factor
modifying the response of Svalbard glaciers to climate change. Coastal zones
receive the highest precipitation and experience low summer temperature and
hence are heavily glacier-covered. In contrast, the interior of Spitsbergen,
the largest island of the archipelago, receives relatively low amounts of
precipitation due to its distance to the open ocean and to the surrounding
rugged terrain of Svalbard; both factors act to limit moisture transport into
the interior. As a result, this area has fewer and smaller glaciers than the
adjoining areas. Lower snow amount means earlier exposure of low albedo
surfaces and a more continental climate, with higher summer temperatures (Hagen
et al., 1993; Nuth et al., 2013; Przybylak et al., 2014). The response of
glaciers to climate change in these districts has been much more seldom
studied, probably because of their presupposed low significance in the
contribution to sea-level rise; small alpine glaciers are also
difficult to study with satellite altimetry and regional mass balance models
due to their complex relief. Detailed information on their spatiotemporal
mass balance variability could, therefore, be used to test the Svalbard-wide
modelling assessments. Moreover, research on the evolution of these small
glaciers could be of practical interest, since they surround the main
settlements of Svalbard. Their retreat may influence human activity, e.g. due
to increased water and sediment delivery from glacier basins and associated
consequences, such as floods and fjord bathymetry changes (Szczuciński et
al., 2009; Rachlewicz, 2009a; Strzelecki et al., 2015a).
One of the regions situated the furthest from maritime influences
(ca. 100 km) is the sparsely glacier-covered Dickson Land (DL). This paper
presents an inventory of the ice masses in DL and quantifies changes of their
geometry since LIA termination. This includes changes of their area and
length, as well as recent volume fluctuations, using digital elevation models
(DEMs)
obtained from aerial photogrammetry. The aim of this study is to investigate
the response of glaciers in DL to climate change, with particular focus on
their recent mass balance and its spatial variability.
Study area
The study region is located in central Spitsbergen and stretches between
78∘27′ N–79∘10′ N and
15∘16′ E–17∘07′ E. Its area is
1.48 × 103 km2 with a length of ca. 80 km in the north–south
direction and a typical width of 20–30 km. For the purpose of the
glaciological analysis, DL was divided into three subregions – south (DL-S),
central (DL-C), and north (DL-N) (Fig. 1). DL-S is the lowest elevated and is
dominated by plateau-type mountains, with summits reaching
500–600 m a.s.l., occupied by small ice fields and ice masses plastered
along gentle slopes. DL-C is the subregion with the greatest ice cover and
the largest glaciers, mostly of valley type, and summits exceeding 1000 m.
The mountains in DL-N are even slightly higher than in the central part, but
glaciers (mainly of valley and niche types) are smaller here and mostly
oriented towards the north.
The climate of DL shows strong inner-fjord, quasi-continental
characteristics, i.e. reduced precipitation and increased summer air
temperature when compared to the coastal regions. The southernmost inlet of
DL is located about 20 km north of Svalbard Lufthavn weather station (SVL,
15 m a.s.l.) near Longyearbyen. Between 1981 and 2010, the Norwegian
Meteorological Institute recorded an average annual temperature of
-5.1 ∘C at SVL, with the summer (June–August) mean of
4.9 ∘C. Annual measured precipitation was 188 mm. In DL-C daily
means of sea-level air temperature are very similar to those at SVL
(Rachlewicz and Styszyńska, 2007; Láska et al., 2012). No
meteorological stations are operating in DL-N, but the general climatic
pattern suggests it is among the driest zones in all Svalbard (Hagen et al.,
1993).
Previous glacial research performed in DL-C has focused mainly around the
impact of glacier retreat on landscape evolution (e.g. Karczewski, 1989;
Kostrzewski et al., 1989; Gibas et al., 2005; Rachlewicz et al., 2007;
Rachlewicz, 2009a, b; Ewertowski et al., 2010, 2012; Ewertowski and Tomczyk,
2015; Evans et al., 2012; Szpikowski et al., 2014; Pleskot, 2015; Strzelecki
et al., 2015a, b). More detailed glaciological investigations were performed
on Bertilbreen (e.g. Žuravlev et al., 1983; Troicki, 1988) and recently
also on Svenbreen (Małecki, 2013a, 2014, 2015). Glaciers in central and
eastern parts of DL-C are losing their mass and their fronts are retreating
(Rachlewicz et al., 2007; Małecki, 2013b; Małecki et al., 2013;
Ewertowski, 2014). Glaciers of DL-N and DL-S have not been studied yet.
DL glaciers are mostly small, and only the largest
(> 5 km2) are partly warm based (Małecki, unpublished
radar data). As a result, ice flow velocities are low; the maximum measured
on the largest glaciers is less than 12 m a-1 (Rachlewicz, 2009b),
while on smaller glaciers it is several times lower (Małecki, 2014). In
every subregion, however, surge-type glaciers are to be found. Studentbreen,
the north-eastern outlet of Frostisen ice field, surged around 1930.
Fyrisbreen advanced around 1960 (Hagen et al., 1993) and Hørbyebreen
surged probably in the late 19th or early 20th century (Małecki et al.,
2013). Also, 2009/2011 aerial imagery acquired by the Norwegian Polar
Institute (available at toposvalbard.npolar.no) shows that the
Hoegdalsbreen–Arbobreen, Manchesterbreen, and Vasskilbreen systems
are characterized by deformed (looped) flow lines and/or moraines, which may
indicate older surges.
Data and methodsGlacier boundaries
A ready-to-use Svalbard glacier inventory from the Norwegian Polar Institute
(NPI) (König et al., 2013; Nuth et al., 2013) was evaluated as a
potential data source for the purpose of this study. Due to the large
Svalbard-wide scale of this work, some difficulties were met during
preliminary geometry change analysis. Firstly, glaciers smaller than
1 km2 are not catalogued in the NPI glacier inventory. Secondly,
polygons for the 2000s, particularly of the smallest ice patches, were too
coarse to accurately reproduce their subtle decadal changes. Therefore,
glacier inventories from this paper (covering glacier extents from
Neoglacial maximum/LIA, 1960s, 1990, and 2009/2011) were prepared by the
author using the original NPI source data, i.e. maps and modified ice and
snow masks.
Glacier boundaries for the 1960s were manually digitized using ArcGIS
software from scanned and georeferenced 1 : 100 000 S100 series paper
maps, constructed by NPI from 1 : 50 000 aerial imagery taken between 1960 and
1966. The LIA area of glaciers was estimated by adding the area of their
moraine zones to the 1960s outlines, but no information was available for
their lateral extent at that time. The 1990 outlines are based on the NPI
glacier inventory (König et al., 2013; Nuth et al. 2013), but many
polygons were added or modified according to the author's experience from the
field to minimize errors of the final glacier area measurement. The most
recent outlines were taken from the NPI database (available at
data.npolar.no) as shapefiles based on 2009/2011 aerial photographs
(Norwegian Polar Institute 2014a), which proved to be very accurate during
direct field surveys.
Confluent glaciers of comparable size separated by a medial moraine were
treated as individual units, except for Ebbabreen, the largest glacier in
DL, which is historically considered as one object. Where possible, minor tributary
glaciers, which eventually separated from the main stream, were fixed as
individual glaciers in the earlier epochs as well, so area changes of a
given glacier result from ice melt-out rather than from disconnection of
former tributaries. Very small episodic snow fields and elongated
snow patches connected with main glacier bodies were excluded from the
inventory. Ice divides were fixed in time and did not account for changing
ice topography. The small ice fields of Frostisen and Jotunfonna were not
further divided into glacier basins.
Digital elevation models
As a 1990 and 2009/2011 topographic background for the analysis, 20 m
DEMs from the NPI were used (Norwegian Polar
Institute, 2014b). The 1990 DEM, which was constructed from 1 : 15 000
aerial photographs, does not cover major glaciers in eastern DL-C,
which comprise 17 % of the modern glacier area of DL (Fig. 1b), so their
elevation changes for the 1990–2009/2011 period could not be measured. Data
for the most recent DEM originate from 0.5 m resolution aerial photographs,
mainly from 2011, but the small eastern part of DL was covered by an earlier
2009 campaign. These data sources were projected into a common datum ETRS
1989 and fit onto a common grid. The universal co-registration procedure
described by Nuth and Kääb (2011) was used to accurately align the
datasets.
Calculation of glacier geometry parameters and their changes
From the modern boundaries and 2009/2011 DEM, the main morphometric
characteristics of glaciers could be extracted. These were area (A), length
(L), mean slope (S), mean aspect (α), minimum, maximum, median, and
moraine elevation (Hmin, Hmax, Hmed, and
Hmor respectively), and theoretical steady-state equilibrium line
altitude (tELA), assuming an accumulation area ratio of 0.6. The
area was measured for each polygon and epoch (Amax, A1960,
A1990, and A2011 respectively for each of the analysed epochs). S,
α, Hmin, Hmax, and Hmed were computed for each
polygon for 2009/2011. L was calculated for each epoch along the
centrelines of the 66 largest valley, niche, and cirque glaciers, excluding
irregular ice masses with no dominant flow direction, former minor tributary
glaciers that used to share front with the main glacier in their basin, and
very small glaciers with Amax < 0.5 km2. On complex
glaciers, e.g. with multiple outlets (e.g. Jotunfonna), more than one
centerline had to be used to determine the representative lengths and retreat
rates. Several parameters were used as indicators of glacier fluctuations,
including area changes (dA), length changes (dL), volume
changes over the period 1990–2009/2011 (dV), and mean elevation
change for the period 1990–2009/2011 (dH), all also given as annual
rates (dA/ dt, dL/ dt, dV/ dt, and dH/ dt
respectively). All rates of glacier change indicators were computed according
to the year of validity of geometry data.
To compute dV, elevation change pixel grids were first calculated
for each ice mass by subtracting the 1990 DEM from the 2009/2011 DEM. This is
an accurate method for determining mass change (Cox and March, 2004),
providing information about thickness changes over the entire glacier with no
need for extrapolation of mass balance values from single reference points,
as is the case with stakes used in the direct glaciological method. The
arithmetic average of elevation change pixels lying within the larger (here
1990) glacier boundary (dh‾) was then used to compute
dV using Eq. (1).
dV=dh‾×A1990.
The mean elevation change of glaciers, dH, was inferred by dividing
dV by the average area of a glacier over the period 1990–2009/2011
to account for its retreat (Eq. 2).
dH=2dV(A1990+A2011).
Near-surface glacier density changes were not considered in the conversion of
the geodetic mass balance to water equivalent (w.e.), as they were assumed to
be small when compared to climatically induced elevation changes over the
study period 1990–2009/2011. This assumption is more uncertain in the
highest zones of glaciers, where changes in firn thickness may lead to
considerable density variations. However, direct field surveys and analysis
of the available satellite images indicate that in the late summer the
highest glacier zones in DL are usually composed of glacier ice or
superimposed ice and almost no firn is present. Moreover, Kohler et
al. (2007) found a good match between the geodetic and
glaciologically measured cumulative mass balance on a small NW Spitsbergen
glacier, implying density changes may be neglected in geodetic balance
calculations on comparatively small and retreating ice masses in Svalbard.
Therefore, dH/ dt could be converted to water equivalent by
multiplication by an average ice density of 900 kg m-3.
Errors
Glacier area measurements for the 1960s epoch suffer from errors associated
with general map accuracy or misinterpretations made by cartographers, e.g.
due to the considerable extent of winter snow cover on aerial images. To
account for that, 25 m was used as a horizontal glacier polygon digitalizing
error. Each polygon was assigned a 25 m buffer with “-” and “+”
signs. Including these buffers, new areas of DL glaciers were computed and
compared to all original polygons. Differences between the new and original
values were used as an error estimate of A1960 for each glacier, with
± 6.4 % as a region-wide total, which was larger for the smaller ice
masses. Since no maps are available for the LIA maximum, LIA glacier area
estimation is based on the 1960s outlines and geomorphological mapping of
moraine zones. Such an approach assumes only frontal retreat in the period
LIA–1960s, but some lateral retreat most likely took place as well. Also,
moraine deposits of some glaciers could have either been eroded before the
aerial photogrammetry era or not formed at all. Application of a relatively
large ± 50 m buffer around the LIA outlines resulted in a total
glacier area error estimate of ± 11.5 % for that epoch. For 1990 and
2009/2011 epochs lower buffers of ± 10 m and ± 5 m were used,
resulting in glacier area uncertainty estimates of ± 3.4 and
± 2.2 % for the whole DL region. Uncertainties of length measurement
for each year were set according to the buffers described above.
Histogram of elevation differences between 2009/2011 DEM and 1990 DEM
over non-glacier-covered terrain.
Main features of the modern glacier geometry in DL: area-altitude
distribution (a), scatter plot of latitude against median glacier elevations (b), and frequency distribution of mean glacier aspects (c).
To estimate the error of dh‾ (ε), elevation
differences between the 1990 and 2009/2011 DEMs over non-glacier-covered
terrain in the whole study region were measured. Since glacier surface slopes
in DL are relatively gentle, mountain slopes steeper than 20∘ were
excluded from the analysis. The results show that an elevation difference of
over 70 % of pixels is within ± 2 m and less than 5 % are
characterized by an elevation difference of more than ± 5 m (Fig. 2).
The mean elevation difference between the two DEMs was 0.24 m, a correction
further subtracted from all obtained dh‾ values, while the
standard deviation, σ, was 2.68 m. Here, σ is used as a point
elevation difference uncertainty and is further used to compute ε
for individual glaciers. The elevation measurement error of snow-covered
surfaces was, however, expected to be larger than for rocks and vegetated
areas due to its lower radiometric contrast on aerial images. To account for
this effect, parts of glacier surfaces extending above 550 m a.s.l. (an
approximate snow line on 1990 and 2009/2011 aerial imagery) have a prescribed
error characteristic of 2σ. For each glacier, ε was then
calculated using Eq. (3):
ε=[(1-n)⋅σ]+(n⋅2σ)N,
where n is the fraction of the glacier extending above 550 m and N is
the sample size. Assuming spatial autocorrelation of elevation errors at an
order of 1000 m after Nuth et al. (2007), N becomes glacier size
in km2 rather than number of sample points. Using ε and
errors of glacier area measurements, uncertainties of dV and
dH could be assessed with conventional error propagation methods.
All errors are relatively large for the smallest ice masses and vice versa.
ResultsModern geometry of Dickson Land glaciers
In the most recent 2009/2011 inventory 152 ice masses were catalogued in DL,
all terminating on land and covering a total of 207.4 ± 4.6 km2
(14 % of the region). 110 ice masses (72 % of the population) have
areas < 1 km2 and 86 of these are smaller than 0.5 km2.
Only nine glaciers (6 %) are larger than 5 km2. The largest glaciers
are Ebbabreen (24.3 km2), Cambridgebreen–Baliollbreen system
(16.3 km2), Hørbyebreen system (15.9 km2), and Jotunfonna
(14.0 km2). North-facing glaciers (N, NW, and NE) comprise 61 % of
the population, while only 16 % of ice masses have a southern aspect (S,
SW, and SE). The mean glacier slope is 10.7∘.
DL-C is the subregion with the greatest glacier coverage (26 % or
117 km2), compared to only 8 % (39 km2) and 10 %
(51 km2) in DL-S and DL-N, respectively. The subregions also differ
significantly in their area–altitude distribution. Glacier maximum and median
elevation increases moving from south to north. DL-N contains most of the
high-elevation glacier area in DL, with a median elevation of 614 m. In
DL-C, glacier fronts reach the lowest elevations, while the glacier
hypsometry of DL-S is the flattest and contains the lowest fraction of
high-elevation areas. The median elevation of the two latter subregions is
520 m, giving an overall median elevation of glaciers in DL of 539 m and a
tELA of 504 m a.s.l. The total volume of DL ice masses, estimated
with empirical area–volume scaling parameters by Martín-Español et
al. (2015), is roughly 12 km3. The main details of glacier geometry
characteristics are depicted in Fig. 3.
Changing extent of glaciers in Dickson Land over the study periods.
* Excluding glaciers in eastern DL-C due to the lack of 1990 DEM
coverage.
Glacier area and length reduction
Since the termination of the LIA, the glaciers of DL have been continuously
losing area, in total by 38 ± 12 % (Fig. 4a; Table 1). The overall
rate of area loss was 0.49 ± 0.66 km2 a-1 in the first
epoch, which increased 4-fold to 2.01 ± 0.85 km2 a-1
after 1960 and further to 2.23 ± 0.48 km2 a-1 after 1990
(Fig. 4a). Excluding known and probable surge-type glaciers, whose areal
extent can change due to internal dynamic instability rather than in direct
response to climate, shows that increasing area loss rates are related to
climate forcing rather than to ice dynamics (Fig. 4b). The larger error bars
of dA/ dt preclude identification of any trends in that signal.
(a) Changes of the total glacier area in Dickson Land. (b) Same
as (a) but for non-surging glaciers only. (c) Average glacier length change
rates in Dickson Land and its subregions.
In contrast to dA/ dt, average length change rates dL/ dt
have smaller uncertainties. From the available temporal resolution of the
data no front advances were detected, although the surge events of Frostisen
and Fyrisbreen occurred in the first period (Hagen et al., 1993). In general,
all glaciers have been retreating since the LIA termination and the extremes
of total dL observed in DL were -46 and -3325 m. Epochs LIA–1960s and
1960s–1990 were the periods with the fastest retreat for only 26 % of
the study glaciers. In many of the latter cases, bedrock topography supported
a short-term increase in dL/ dt, e.g. due to rock sills dissecting
thinning glacier snouts into active and dead ice zones (e.g. Ebbabreen,
Frostisen, Svenbreen). The vast majority of glaciers (74 %) were
retreating at their fastest rate in the last study period 1990–2009/2011.
Glacier thinning and mass balance
A strikingly negative and consistent elevation change pattern is evident from
the 1990–2009/2011 data, including in the highest zones of glaciers all over DL
(Figs. 5 and 6). At the lowest altitudes (< 200 m a.s.l.), the
mean change rate was ca. -2 m a-1, while at the average
tELA (ca. 500 m a.s.l.) this was about -0.6 m a-1.
Positive fluctuations were observed above ca. 1000 m a.s.l., mostly in
DL-N. Some glaciers have been thinning at a very high average rate, exceeding
1 m a-1, while only a few small ice patches have been closer to
balance. Overall, the average area-weighted dH/ dt in DL was highly
negative at -0.71 ± 0.05 m a-1
(-0.64 ± 0.05 m w.e. a-1), resulting in a total volume loss
rate of 137 ± 6 × 106 m3 a-1 and a mass
balance of -0.12 ± 0.01 Gt a-1 (excluding major glaciers in
eastern DL-C due to the lack of 1990 DEM coverage). Subregional values are
given in Table 2 and indicate that the most negative specific mass balances
are found in DL-C and the least negative in DL-N.
Homogeneity of the (a) 1990–2009/2011 elevation change pattern in DL
subregions. (b) The mean pre-1990 and post-1990 elevation change rates in DL
averaged from the available data. Horizontal bars represent 1 standard
deviation. The 1960s–1990 data are compiled from Małecki (2013b) and
Małecki et al. (2013).
Scatter plots showing the relationship between mean 1990–2009/2011
glacier elevation change and median elevation of glaciers (a) and total area
change and total length change of glaciers (b).
Links between glacier change indicators and their geometry
Recent thinning rates decrease with altitude, so the highest-elevation
glaciers, mainly in DL-N, have been thinning the least, while glaciers with a
large portion of low-elevation ice (e.g. as in DL-C) had the fastest thinning
rates (Fig. 7a). Length changes are correlated with terminus altitude and
glacier length, so low-elevation fronts of long glaciers have been retreating
at the fastest rates. Relative area change was best correlated with relative
length change (Fig. 7b), glacier area, maximum elevation, and length, so large
glaciers lost the smallest fraction of their maximum extent despite
significant absolute area and length losses. In contrast to reports from many
other regions of the globe (e.g. Li and Li, 2014; Fischer et al., 2015; Paul
and Mölg, 2014), glacier aspect showed no statistical correlation with any
of the glacier change parameters, which may result from the summertime
midnight sun over Svalbard and the more balanced insolation on slopes with
northern and southern aspects, compared to midlatitudes. Pearson correlation
coefficients of glacier change parameters against other parameters and
glacier geometry variables are given in Table 3.
Pearson correlation coefficients for glacier change indicators
against other indicators and geometry parameters. Bold values indicate
statistical significance at p=0.01 level.
In agreement with earlier studies from Svalbard (Kohler et al., 2007; Nuth et
al., 2007, 2010, 2013; James et al., 2012), climate warming is anticipated to
be the main control for the observed negative glacier changes in DL. Air
temperature at the nearest meteorological station, SVL, clearly increased in
the 1920s and 1930s, as well as after 1990 (Nordli et al., 2014), which
explains the glacier retreat after the LIA maximum and in the last study
epoch, respectively. However, the clear post-1960 mass loss acceleration of
DL glaciers may not simply be explained by increased air temperature. In the
period 1960–1990 the total glacier area loss rate quadrupled (although with
large uncertainty) and front retreat rates doubled, despite the fact that the
mean multi-decadal summer air temperature was very similar to that in the
first epoch and no decrease in winter snow accumulation over Svalbard was
evident at that time (Pohjola et al., 2002; Hagen et al., 2003). In this
context, it seems likely that average summer air temperature is not the only
driver of change for small, low-activity glaciers in DL and other factors may
also play a role. These could be, for example, different response times of
glaciers or albedo feedbacks, which could modify glacier mass balance in a
nonlinear pattern, e.g. by removal of high-albedo firn from accumulation
zones, and hence increase energy absorption (Kohler et al., 2007; James et
al., 2012, Małecki 2013b).
For the majority of glaciers in DL, the post-1990 period was marked by their
fastest multi-decadal front retreat rates since the LIA maximum. This trend
is similar to that on many land-terminating glaciers of Svalbard (Jania,
1988; Lankauf, 2007; Zagórski et al., 2008; James et al., 2012; Nuth et
al., 2013) (Fig. 3). Length reduction was the main driver for glacier area
decrease (Fig. 7b), which was high in DL and amounted to 38 %, supporting
previous conclusions by Ziaja (2001) and Nuth et al. (2013) that central
Spitsbergen, with its much smaller glaciers, is losing its ice cover extent
at a relatively higher rate than maritime regions of Svalbard (e.g. 18 %
area decrease in Sørkapp Land, 1936–1991; reported by Ziaja, 2001). Area
loss rates in DL were at a similar level between 1960s–1990 and
1990–2009/2011, comparable to the results in Nuth et al. (2013), who
concluded that there was no clear trend of dA/ dt evolution over the
archipelago except for southern Spitsbergen, where area loss rates generally
decreased after 1990. In contrast, Błaszczyk et al. (2013) concluded
that there were increasing area loss rates for tidewater glaciers in Hornsund,
part of south Spitsbergen. Interestingly, ca. 800 km2 of glaciers in
Hornsund, often considered to be among the most sensitive to climate warming,
have been losing area at a rate comparable to ca. 200 km2 of small
glaciers in DL (ca. 1 km2 a-1 for the period LIA–2000s).
Clear acceleration of length loss rates indicates that glaciers in DL have
been experiencing an increasingly negative mass balance since the termination
of the LIA. This is in line with earlier studies. For seven glaciers in DL-C,
Małecki (2013b) documented mean dH/ dt of -0.49 m a-1
for the period 1960s–1990, followed by an acceleration of mass loss rate to
-0.78 m a-1 after 1990. Kohler et al. (2007) analysed
dH/ dt of two small land-terminating glaciers in Spitsbergen with
greater temporal resolution than that available for this study and concluded
there was a continuous acceleration of their thinning over the 20th century,
e.g. from dH/ dt=-0.15 m a-1 (1936–1962) to
dH/ dt=-0.69 m a-1 (2003–2005) for Midre
Lovénbreen in NW Spitsbergen. James et al. (2012) documented negative
dH/ dt for six small land-terminating glaciers all over Svalbard
since at least the 1960s and reported a post-1990 increase in mass loss rates
for four of these. Their recent dH/ dt ranged from -0.28 to
-1.21 m a-1, i.e. similar to the values observed in DL.
An important finding of this study is the observation of glacier-wide
thinning over DL up to an elevation of 1000 m a.s.l., where the average
1990–2009/2011 zero elevation change line was found. To put this into
historical context, previous analyses performed for the earlier period
1960s–1990 identified this threshold at a much lower average altitude, i.e.
at ca. 600 m a.s.l. in DL-C (Małecki, 2013b; Małecki et al., 2013)
(Fig. 6). The shift of the geodetic equilibrium suggests a recent negative
change in glacier mass balance, including former accumulation zones. This
hypothesis is supported by direct records (2011–2015) from Svenbreen (DL-C),
where negative surface mass balance has also been noted at the highest
ablation stake (625 m a.s.l.) near the glacier headwalls (Małecki,
unpublished data). On Nordenskiöldbreen, a large tidewater glacier
neighbouring DL from the east, mean 1989–2010 ELA, was modelled at
719 m a.s.l., i.e. higher than the accumulation zones of most DL glaciers
(Van Pelt et al., 2012).
Thinning at the high elevations of the study glaciers could be linked to
several factors. Firstly, there is the increased melt energy availability due
to (i) increased incoming longwave radiation from the atmosphere and
turbulent heat fluxes resulting from post-1990 summer air temperature rise;
(ii) increased energy absorption by the ice surface due to decreasing albedo
caused by firn melt-out, dust, or sediment delivery from freshly exposed
headwalls; and (iii) increased longwave emission from surrounding slopes
recently uncovered from snow and ice. Other possible explanations are related
to firn evolution, i.e. its compaction or melt-out, supporting the reduction
of internal meltwater refreezing. The last probable mechanism could be a
recent snow accumulation decrease. Data availability on winter mass balance
in DL are insufficient for such conclusions (Troicki, 1988; Małecki, 2015),
but the trend for a snow precipitation decrease after 1990 has been noted for
SVL (James et al., 2012). Glacier dynamics could also explain changes in the
glaciers' upper zones, but there are too few data to test this idea. However,
low flow velocities of DL glaciers (1–10 m a-1) suggest the minimal
importance of the dynamic component in their surface elevation changes.
High-elevation glacier thinning in DL will have important consequences for
the local cryosphere. Surge-type glaciers will not build up towards new
surges and as such could be removed from the surge cycle under present
climate conditions, as demonstrated in more detail for Hørbyebreen by
Małecki et al. (2013). This will also lead to decay of temperate ice
zones, still found beneath the largest glaciers of DL (Małecki,
unpublished data), and consequently it will influence their hydrology and
geomorphological activity and reduce ice flow dynamics, as documented for
other small glaciers in central Spitsbergen (Hodgkins et al., 1999; Lovell et
al., 2015). Eventually, given that the highest parts of glaciers in DL
typically reach 700–800 m a.s.l., the high altitude of the recent geodetic
equilibrium suggests their considerable or complete melt-out in the future,
even if the atmospheric warming trend stops. Notably, altitude had the
strongest influence on the spatial mass balance variability (Figs. 6 and 7a),
so small low-elevation glaciers were the most sensitive to climate shift.
They had the fastest front retreat rates and the most negative
dH/ dt (Fig. 7a); hence, they are likely to be the first to
disappear.
Glacier-wide surface lowering has already been triggered in some of the
world's largest ice repositories, including the Canadian Arctic Archipelago
(Gardner et al., 2011) and Patagonian ice fields (Willis et al., 2012),
causing them to significantly contribute to sea-level rise. In Svalbard, the
major ice masses are still building up their higher zones and remain closer
to balance (Moholdt et al., 2010; Nuth et al., 2010), but the process of
high-elevation thinning seems to be already widespread on smaller glaciers
across the archipelago, as documented by Kohler et al. (2007), James et
al. (2012), and this study. By the end of the 21st century, a further
3–8 ∘C warming over Svalbard is predicted by climate models
(Førland et al., 2011; Lang et al., 2015). This will eventually cause the
complete decay of the accumulation zones of Svalbard ice masses, boosting
their mass loss rates and the sea-level rise contribution from the region.
Small Spitsbergen glaciers may, therefore, be perceived as an early indicator
of the future changes of larger ice caps and ice fields.
The mass balance of glaciers in central Spitsbergen has been previously
considered by some researchers to be relatively resistant to climate change
due to the prevailing dry conditions and high hypsometry (Nuth et al., 2007).
However, at -0.71 ± 0.05 m a-1
(-0.64 ± 0.05 m w.e. a-1) the average mass balance of
glaciers in DL is among the most negative of the Svalbard regional means
reported by Nuth et al. (2010) and Moholdt et al. (2010). Previously
published occasional data from another region of central Spitsbergen,
Nordenskiöld Land, show a generally similar glacier response to climate
change and comparable mass balances to glaciers in DL (e.g. Troicki, 1988;
Ziaja and Pipała, 2007; Bælum and Benn, 2011), indicating that
observations from this study are valid for larger areas of the island's
interior. Extrapolation of the mass balance from DL to glaciers in eastern
DL-C and to neighbouring Nordenskiöld Land and Bünsow Land (Fig. 1a),
comparable in terms of climate and glacier-cover characteristics, yields an
estimate of the total mass balance of glaciers in central Spitsbergen.
Despite their negligible share of the archipelago's ice area
(ca. 800 km2 or 2 %), they contribute about 0.6 Gt a-1 to
the sea-level rise, a figure comparable to the contribution of some of the
much larger glacier regions, e.g. parts of southern or eastern Svalbard. The
total mass balance of the archipelago has been estimated to range from
-4.3 Gt a-1 (Moholdt et al., 2010) to -9.7 Gt a-1 (Nuth et
al., 2010).
Conclusions
In this study, a multi-temporal inventory and DEMs of 152
small alpine glaciers and ice patches in Dickson Land, central Spitsbergen,
were used to document their post-LIA evolution. In order to be in
balance with the present climate, their ELA should be approximately
500 m a.s.l. However, due to progressive climate warming in Svalbard, the
average ELA has increased and glaciers have been continuously losing mass for
many decades. The total ice area in Dickson Land has been declining at an
accelerating rate from 334.1 ± 38.4 km2 at the termination of the
LIA (early 20th century) to 207.4 ± 4.6 km2 in
2009/2011, corresponding to an overall 38 ± 12 % decrease.
Post-1990 area loss rate was 4.5 times higher than in the epoch LIA–1960s,
i.e. 2.23 ± 0.48 km2 a-1 vs.
0.49 ± 0.66 km2 a-1 respectively. Front retreat of
66 test glaciers has accelerated over time, i.e. from an average of
4.9 ± 0.1 m a-1 in the period from the LIA maximum to
the 1960s and 9.4 ± 0.1 m a-1 between the 1960s and 1990 to
16.4 ± 0.1 m a-1 in the last study epoch 1990–2009/2011, which
turned out to be the period of the fastest retreat for 74 % of glaciers.
The most important finding of this study is the recent rapid glacier-wide
thinning over the entire region at a mean rate of
0.71 ± 0.05 m a-1 (-0.64 ± 0.05 m w.e. a-1). The
warming climate has caused an ELA rise and a consequent increase in the
zero-elevation change line, so local glaciers have been thinning up to the
altitude of 1000 m, i.e. higher than their accumulation zones. The spatial
variability of glacier mass balance was primarily correlated with elevation,
so small low-elevation glaciers have generally been losing mass and length at
the fastest rates and are under threat of the earliest disappearance.
The Supplement related to this article is available online at doi:10.5194/tc-10-1317-2016-supplement.
Acknowledgements
The study is a contribution to the DIL*ICE project (Dickson Land Ice Masses
Evolution, RiS id 4894) supported by the Polish National Science Centre
(grant no. N306 062940) and the Institute of Geoecology and Geoinformation of
Adam Mickiewicz University in Poznań. The author sincerely appreciates
the support from Copernicus Publications and the open-access data policy of
the Norwegian Polar Institute. Constructive reviews from P. Holmlund and
J. Kohler helped to significantly improve the manuscript and are greatly
appreciated. The comments on the early manuscript and handling of the
manuscript by the editor J. O. Hagen are also acknowledged.Edited
by: J. O. Hagen
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