CHAPTER
30
Monitoring glacier changes on
the Antarctic Peninsula
Jorge Arigony-Neto, Pedro Skvarca, Sebastián Marinsek, Matthias Braun,
Angelika Humbert, Cláudio Wilson Mendes Júnior, and Ricardo Jaña
ABSTRACT
30.1
INTRODUCTION
The Antarctic Peninsula has exhibited some of the
most spectacular changes observed in glacial systems in recent decades. The events include disintegration of ice shelves, acceleration and thinning of
glaciers, variations in the limits between glacier
facies, and retreat of glacier fronts. However, due
to the lack of both consistent systematic observations of the glacial systems and information on their
boundary conditions, it is difficult to accurately
predict the contribution of Antarctic Peninsula
glaciers to sea level rise and further responses of
these ice masses to climatic and oceanographic
changes. In this context, the activities of the
GLIMS Regional Center for the Antarctic Peninsula and its network of international collaborators
are based on the use of various types of Earth
observation imagery, mainly optical and radar
data. Although a complete glacier inventory is still
lacking, we present the results of changes in glacier
frontal positions and boundaries of glacier facies as
well as links to dynamical adjustments for various
locations in the Antarctic Peninsula’s ice masses.
Evaluation of Advanced Spaceborne Thermal
Emission and reflection Radiometer (ASTER)
digital elevation models generated for the Antarctic
Peninsula is also discussed.
The Antarctic Peninsula (AP; Fig. 30.1) has undergone rapid climatic warming during the last 50
years (Vaughan et al. 2003). As a consequence,
the glacial systems of this region have reacted with
drastic changes, such as retreat, breakup, and disintegration of ice shelves (Skvarca 1993, Rott et al.
1996, 1998, Skvarca et al. 1999, Skvarca and De
Angelis 2003, Rack and Rott 2004, Braun et al.
2009), acceleration and thinning of glaciers (De
Angelis and Skvarca 2003, Rignot et al. 2004,
Scambos et al. 2004, Pritchard and Vaughan
2007), variation in the limits between glacier facies
(Rau and Braun 2002, Arigony-Neto et al. 2007,
2009), and retreat of glacier fronts (Skvarca et al.
1995, Simões et al. 1999, Rau et al. 2004, Cook et al.
2005).
Glacial response times in the AP range widely, as
do climatic and oceanographic components of the
system, and so glacier dynamics differ considerably
in the region. Climate change is often considered a
major cause of glacier changes in the AP (Skvarca et
al. 1999, Skvarca and De Angelis 2003, Vaughan et
al. 2003). However, due to the lack of both data on
the boundary conditions of glaciers and consistent
systematic observations of glacier dynamics to feed
more complex models of glacier mass balance, it is
718
Monitoring glacier changes on the Antarctic Peninsula
Figure 30.1. Overview map of the Antarctic Peninsula showing the location of the study areas. Black boxes
correspond to areas analyzed with satellite sensor data. Glaciers discussed in the text are labeled as follows:
BS ¼ Boydell and Sjögren; DBE ¼ Dinsmoor, Bombardier, and Edgeworth; D ¼ Drygalski; HGE ¼ Hektoria, Green,
and Evans; and JCMM ¼ Jorum, Crane, Mapple, and Melville.
Regional context 719
difficult to predict accurately the contribution of AP
glaciers to changes in sea level (Rignot et al. 2004,
2005, Pritchard and Vaughan 2007) and further
responses to climatic change and variation in
oceanographic parameters (Vaughan 2006,
Arigony-Neto et al. 2007).
Work done within the framework of the GLIMS
Regional Center for the Antarctic Peninsula and its
network of international collaborators aims not
only to promote the use of optical imagery to monitor ice masses of this region, but also to investigate
the possibility of adapting products, based on
analyses of SAR data, to the GLIMS concept in
a synergistic way. In this chapter, we briefly present
some of the indicators of climatic and glaciological
changes observed in the AP, and evaluate ASTERderived products’ ability to represent the topography of AP ice masses. We also present three case
studies in which variations in dynamical parameters
of a subset of AP glaciers are measured using
methods related or complementary to GLIMS
analyses.
30.2
REGIONAL CONTEXT
30.2.1 Geologic context
The AP is vast, stretching southward from 60 S for
more than 1,500 km down to 75 S. The region is
widely glaciated, interrupted only by sparse rock
outcrops. The AP’s early development was as an
Andean-type oceanic/continental plate margin of
Gondwanaland. Construction of the plate margin—started in the Paleozoic and continued
through the Mesozoic—involved compressional
folding, uplift, and metamorphism of marine sediments (including thick sedimentary sequences shed
from Gondwanaland into the marine environment),
anatexis (partial melting of the crust), igneous intrusions, and volcanism (mainly silicic). Hence, the
AP’s basement rocks, widely exposed by uplift
and erosion, are primarily high-grade metamorphic
rocks, such as granite gneiss, and silicic igneous
rocks (Harrison et al. 1979, Birkenmaier 1994,
Millar et al. 2002).
Starting in the Late Cretaceous, marine sedimentary deposition in a back-arc basin produced
a 6,000 m thick sequence preserved on James Ross
Island (Smellie et al. 2007). The sedimentary platform sequence was uplifted and partly eroded, and
subsequently overprinted by a massive, mainly subglacial basaltic shield volcano and many subsidiary
volcanic structures over the past 9.2 million years.
Glacial erosion of the lava-capped sedimentary
edifice has produced the peculiar shape of James
Ross Island, as shown below (see Fig. 30.5 on
p. 726). Primarily basaltic, subglacial volcanism
also occurred elsewhere in the northern part of
the AP region in Late Cenozoic times, thus producing many rock deposits bearing the imprint of lava–
ice interactions, particularly on Alexander and
Vega Islands as well as James Ross Island (Smellie
et al. 1993, Hambrey et al. 2008). This late episode
of volcanism is thought to have begun beneath an
Antarctic Peninsula ice sheet and to have occurred
more recently beneath small ice caps. This activity
produced interspersed lavas, tuff cones, and glaciodeltaic, glaciofluvial, and diamictite sequences. Volcanic eruptive episodes on James Ross Island are
probably not finished, but they occur only once
every 10 5 years on average (Smellie et al. 2007),
and so eruptions are improbable anytime soon.
Active plate subduction, rifting, and volcanism
still occur in the nearby Shetland Islands region
(Lawyer 2009). In general, however, the AP is not
considered volcanically very active. AP history indicates that most rocks now undergoing glacial erosion in the AP consist of a wide variety of igneous
and high-grade metamorphic rocks.
30.2.2 Climatic context
The Antarctic Peninsula has a polar to subpolar
maritime glacial climate; as such, it experiences
considerable summer rainfall and winter snow. It
is dominated by a rugged, though not particularly
high mountain range, attaining over 2,000 m in
elevation in places but less than 1,000 m along most
of its length. The spine of the AP is high enough to
impose a strong orographic effect, with the east to
southeast side being in the precipitation shadow of
prevailing winds and precipitation sources that are
based mainly in the Bellingshausen Sea (Fig. 30.1).
The maritime glacial climate and its dominant
wet-based temperate and polythermal glaciers and
ice caps differ from the severe polar climate and
cold-based glaciers and thick ice sheets dominating
most of Antarctica. Unlike most of Antarctica,
most of the AP has a distinct summer melting
season, especially in its northern reaches. Glaciers
in the AP mainly have low supraglacial debris loads
because the ice cover is so extensive and rock outcrops are comparatively few; hence, erosion of
glacial valleys must occur mostly by subglacial
abrasion and not much at all by landsliding and
720
Monitoring glacier changes on the Antarctic Peninsula
mass wasting, bearing little resemblance to the case
in Patagonia and Alaska, for instance.
A statistically significant warming trend was
detected in the AP using different datasets, measured both locally in situ and regionally via remote
sensing. Turner et al. (2005) gave the warming rate
of the Antarctic Peninsula at Faraday/Vernadsky
Station (65 14 0 S, 64 15 0 W) as 0.056 0.043 C yr1
over the year and 0.109 0.088 C yr1 during the
winter for the period between 1951 and 2000. Moreover, radiosonde temperature observations complemented by NCEP/NCAR reanalysis data revealed a
mean annual tropospheric (850–300 hPa) warming
of 0.027 0.022 C yr1 above Faraday/Vernadsky
between 1956 and 1999 (Marshall et al. 2002).
Finally, a database of more than 500 mean annual
temperatures derived from measurements of snow
temperature at around 10 m depth and air temperature measured at meteorological stations and
automatic weather stations allowed Morris and
Vaughan (2003) to estimate the increase in temperature for the whole AP to be 0.035 0.010 C yr1
over the period 1904–2000.
In addition to atmospheric warming, a statistically significant rise in the number of precipitation
events during winter was reported from observations at Faraday/Vernadsky from 1956 to 1994
(Turner et al. 1997), resulting in the records of
precipitation per season at the end of the time series
increasing by 50%. This appears to be related to
the more frequent arrival of cyclones from the
Bellingshausen Sea (Turner et al. 1995), where this
precipitation originates (Turner et al. 1998).
Furthermore, data from ice cores (Peel 1992,
Thompson et al. 1994, Raymond et al. 1996,
Thomas et al. 2008) indicated an increase in accumulation on the plateaus of the AP in the last four
decades.
A rising trend in the duration of melt periods in
the AP was detected using two different methods.
Ridley (1993), Fahnestock et al. (2002), and
Torinesi et al. (2003) examined the duration of melt
seasons over 13 (1978–1991), 22 (1978–2000), and
20 (1980–1999) year periods, respectively, using
satellite-based passive microwave data. All three
studies found high variability in the duration of
the melt season as well as a positive trend in melt
season length, although Ridley (1993) and Fahnestock et al. (2002) analyzed the area of the AP ice
shelves only and Torinesi et al. (2003) considered
the whole AP. Vaughan (2006) used an approach
based on positive degree-days (PDDs) to predict
ablation conditions. By using temperature trends
from Morris and Vaughan (2003), annual PDDs
were calculated for the whole AP over the period
1950–2050. Results indicated a strong positive trend
in annual PDDs, mainly concentrated in the northeastern sector of the AP.
30.2.3 Summary of known
glacier dynamics
According to analyses of recent Gravity Recovery
and Climate Experiment (GRACE) data for the
period 2002–2009, which allows estimation of
glacier mass change, the northern AP has the continent’s second largest negative mass balance rate
(Chen et al. 2009). However, such low-resolution
data belie the geographic complexity and rapidity
of some glacier dynamics in the region.
The successive breakup and retreat of ice shelves
occurring in the AP over the last two decades has
received extensive attention from the general public
and media. After the retreat of Wordie Ice Shelf in
the late 1980s (Doake and Vaughan 1991, Vaughan
and Doake 1996), pronounced retreat and breakups
were also detected at the northern margins of
George VI Ice Shelf and Wilkins Ice Shelf
(Lucchitta and Rosanova 1998, Scambos et al.
2000), followed by a rapid sequence of breakup
events on the northeastern side of the peninsula:
Prince Gustav Ice Shelf and Larsen A Ice Shelf
disappeared in 1995 (Skvarca 1993, Rott et al.
1996, 1998, 2002). Larsen B Ice Shelf started calving
in 1995 and collapsed almost completely in 2002
(Skvarca et al. 1999, Rack and Rott 2004). Moreover, new breakup events and retreat were detected
in 2008 and 2009 at Wilkins Ice Shelf, when almost
40% of the ice shelf connecting the two islands
broke off (Braun et al. 2009, Humbert et al.
2010). Progressive thinning (Shepherd et al. 2003,
Skvarca et al. 2004), the formation of melt ponds
(Scambos et al. 2000, 2003), changes in structure
(Glasser and Scambos 2008, Braun et al. 2009)
and extended melt seasons (Fahnestock et al.
2002, Van den Broeke 2005) linked to the breakup of ice shelves are strong indications that some of
these events occurred in response to changing
climatic and oceanographic conditions, at least in
the case of Prince Gustav Ice Shelf and Larsen Ice
Shelf. In this context, Morris and Vaughan (2003)
proposed the 9 C isotherm of mean annual temperature as the new limit of ice shelf distribution in
2002, updating the 5 C limit previously proposed
by Reynolds (1981). Hence, further warming might
Methodology
shift the viability limit of ice shelves in the AP
farther south.
An overall trend of glacier front recession was
detected in the AP (Rau et al. 2004, Cook et al.
2005), on James Ross Island (Skvarca et al. 1995,
Skvarca and De Angelis 2003) and on the South
Shetland Islands (SSI; Park et al. 1998, Calvet et
al. 1999, Simões et al. 1999, Braun and Goßmann
2002). As suggested by Cook et al. (2005), floating
glaciers and most ice shelves may be reacting to
progressive atmospheric warming detected in the
AP (Vaughan et al. 2003, Morris and Vaughan
2003, Scambos et al. 2003). However, as they are
influenced by other external factors (e.g., oceanic
temperature and circulation), it is difficult to isolate
a clear climate cause (Rau et al. 2004). Tidewater
glaciers with their fully grounded marine termini
are influenced by a complex set of forcing mechanisms composed of atmospheric temperature and
oceanographic parameters as well as subglacial
topography (Van der Veen 2002, Benn et al.
2007), and hence their retreat rates cannot be linearly linked to short-time climate change on ice
masses (Pfeffer 2003).
Thus, for all types of glaciers in the AP, climate
change, oceanographic conditions, and glacier/ice
shelf dynamics are complexly coupled, with climate
change pretty obviously a key driving factor. However, the system has strong forcings other than
climate change. Hence, prediction of glacier and
ice shelf dynamics in the AP is hazardous at best,
and clear attribution of recent dynamics to climate
change alone is not warranted.
30.3
METHODOLOGY
30.3.1 Evaluation of ASTER-derived
DEMs for the Antarctic Peninsula
High-quality digital elevation models are only available for selected regions. However, measurement of
morphometric glacier parameters such as length,
width, area, glacier front position, basin boundaries, among others, where 3D coordinates are
needed, requires a consistent spatial frame of reference. Some methods that use spaceborne optical
sensor data to retrieve such parameters have been
tested mainly for alpine and temperate glacier
regions (Paul 2001, Paul et al. 2002). Nevertheless,
in the AP some factors related to the specific characteristics of this polar environment (e.g., frequent
high cloudiness, high reflection of snow-covered
721
surfaces, morphology of the terrain, etc.) render
application of these digital image–processing
methods difficult. Consequently, investigation, testing, and adaptation of these traditional algorithms
for application in the AP are required. With this
focus, we have investigated the use of near-infrared
stereo pairs acquired by ASTER for topographic
data generation along the AP. The ASTER
visible/near-infrared band subsystem (VNIR) offers
15 m spatial resolution (images were provided
through the GLIMS project).
An area of roughly 600 km 2 , located on the western side of the AP in the region of Marguerite Bay,
was chosen as the test site (Fig. 30.1). The area is
almost completely covered by ice and snow, having
a geographical setting inclusive of many of the ice
forms that occur in Antarctica. Topography rises
rapidly from the heads of glaciers to the central
plateau. The best independently available terrain
representation for the area is provided by the
Darmstadt University of Technology Model
(TUD DTM; Fig. 30.2) and the complementary
Base General San Martin aerial photo map (TUD
Karte; Fig. 30.2). This dataset was created from
black and white aerial photography taken by the
Bundesamt für Kartographie und Geodäsie (BKG)
in February 1989 (Wrobel et al. 2000). Generated
by means of semiautomatic processing from aerial
photo coverage, the TUD DTM has a 30 m grid cell
size and the following accuracies: 3–10 m, mountain ranges, rock areas, snow-free zones; 10–20 m,
crevasses, ice faults, structured snow-covered terrain; 50 m or more, monotonous snow-covered
areas without structures, clearly visible in the
central part of the DEM on McClary Glacier
(Fig. 30.2).
To derive a DEM for our test site, eight ASTER
L1A scenes with acceptable cloud coverage were
used from a total of 32 scenes acquired between
November 2000 and September 2005. The process
chain to derive a multitemporal median DEM (MT
DEM) was computed following the flow diagram
shown in Fig. 30.3. A breakdown of this process is
given in Sections 30.3.1.1–30.3.1.3.
30.3.1.1
DEM generation
DEMs were derived for all available scenes using a
fixed combination of parameters with 15 m output
pixel size, a 3 3-pixel correlation matrix size, no
‘‘water detection’’ or ‘‘extended correlation’’.
Terrain geometric correction for each ASTER
722
Monitoring glacier changes on the Antarctic Peninsula
Figure 30.2. (Top) Digital elevation model (IfPK TUD 1999) of the test site located in the surroundings of the
Base General San Martin, Marguerite Bay. Lambert Conformal Conic (WGS 72) cartographic projection. (Bottom)
Base General San Martin aerial photo map (TUD Karte). Sources: aerial photography (1988/1989) by Bundesamt
für Kartographie und Geodäsie (BKG, formerly IfAG); stereo modeling and DEM production by the Technical
University of Darmstadt (TUD); topographic features by the Institute of Physical Geography, University of Freiburg
(IPG) (see Wrobel et al. 2000). Figure can also be viewed in higher resolution as Online Supplement 30.1.
Methodology
723
Figure 30.3. Flow diagram for the multitemporal ASTER scenes processing approach, showing the production of
an averaged DEM by pixel-based median filtering within a layer stack.
L1B scene was performed using the DEM derived
to produce the corresponding orthorectified image.
30.3.1.2
DEM registration
A registration process was carried out in two steps
to register each ASTER-derived DEM with the
reference TUD model. Because of difficulties and
inaccuracies in image-to-image registrations based
on just two DEMs, registration of the orthorectified
L1B image with the TUD Karte image was first
performed. This step provided the GCPs and coefficients for first-degree polynomial correction applied
in each registration. Finally, image-to-map registration using the GCPs and coefficients calculated in
the previous step was carried out for each DEM.
30.3.1.3
Stacking and filtering
Co-registered DEMs were added to a layer stack.
Within this step we masked seawater areas and
resampled the data from 15 to 30 m using the
724
Monitoring glacier changes on the Antarctic Peninsula
nearest neighbor method. Finally, the median value
of stacked pixels was calculated to filter out extreme
pixel values in the derivation of the multitemporal
median DEM (MT MED model).
In order to describe errors that can occur in
ASTER-derived models we compared the altitudes
of each model with the TUD reference model pixel
by pixel. We also considered the Radarsat Antarctic
Mapping Project (RAMP) DEM (Liu et al. 2001),
which provides an independent comparison and a
measure of the magnitude and spatial distribution
of the DEM’s error in this area. RAMP models and
ASTER-derived models were each subtracted from
the TUD reference model. Map algebra operations
(i.e., subtraction and masking) were applied to generate three altitude deviation maps. The resulting
maps (Fig. 30.4) are shown in a blue–red bipolar
color schema used to represent the magnitude and
sign of altitude deviation values. In addition, Table
30.1 shows the distribution of altitudinal deviations
between different elevation models. In Fig. 30.4
blue represents an elevation of ASTER or RAMP
models higher than the corresponding TUD elevation; red represents an elevation of ASTER or
RAMP models lower than the corresponding
TUD elevation. White represents altitude deviation
less than 10 m.
Comparing the altitude deviations of ASTER
and RAMP models from the TUD DEM, we can
observe on the MT MED the preponderance of
altitude differences smaller than 20 m (Fig. 30.4;
Table 30.1). In this range, the distribution of errors
is more concentrated and they are smaller than
shown on the MED MED map. Accordingly, the
MT MED model represents the best fit with the
reference TUD model. This indicates that ASTER
DEMs can achieve higher accuracy using the multitemporal approach than using the alternative
single-scene schema.
The digital elevation model derived within
RAMP did not provide the accuracy necessary
for deriving morphometric glacier parameters as
required for the GLIMS project. This was due to
data source compilation and the scale of the RAMP
DEM (i.e., 200 m spatial resolution). Furthermore,
the accuracy parameters of the RAMP model found
at our test site were less acceptable than those
reported in the literature (Liu et al. 1999, Bamber
and Gomez-Dans 2005).
The accuracy parameters of our ASTER-derived
digital elevation models based on the doublemedian filtering scheme (MED MED) and on the
Figure 30.4. Altitude deviations between the TUD reference model versus ASTER-derived and RAMP models.
(Left) MED MED model, digital elevation model using only one ASTER scene (SC:AST_ L1A.003:2030971130,
September 19, 2005, 13:33:26). (Middle) MT MED model. (Right) RAMP DEM. Blue represents an elevation of
ASTER or RAMP models higher than the corresponding TUD elevation; red represents an elevation of ASTER or
RAMP models lower than the corresponding TUD elevation. White represents an altitude deviation less than 10 m.
Location of test site indicated in Fig. 30.2’s caption.
Case studies and special topics
725
Table 30.1. Distribution of altitude deviations between different elevation models (TUD–
RAMP, TUD—MED MED and TUD–MT MED): absolute and relative frequencies of the
classes used in Fig. 30.4. Rows in gray represent high-accuracy classes.
Deviation ranges
TUD–RAMP
TUD–MT MED
(m)
(pixels)
(%)
(pixels)
(%)
(pixels)
(%)
800 to 100
178,896
28.7
58,645
9.4
7,016
1.1
100 to 50
61,474
9.97
8,595
12.6
51,427
8.3
50 to 20
48,063
7.7
179,356
28.8
173,244
27.8
20 to 10
16,116
2.6
87,901
14.1
100,601
16.2
10 to 10
40,177
6.5
123,960
19.9
161,628
26.0
10 to 20
25,242
4.1
28,907
4.6
36,700
5.9
20 to 50
68,084
10.9
35,365
5.7
48,472
7.8
50 to 100
77,576
12.5
18,513
3.0
26,163
4.2
>100
106,715
17.1
10,862
1.7
16,934
2.7
multitemporal median scheme (MT MED) were
higher than reported in the literature for the same
type of surface (i.e., between 31 and 70 m root
mean square with maximum errors of 200 and
þ500 m; Kääb, 2005). Double-median filters, made
up of a spatial median and a multiprocess median,
were capable of suppressing artifacts in a single
scene. For each point the multitemporal median
(MT MED) used altitude values unaffected by
artifacts.
30.4
TUD–MED MED
CASE STUDIES AND
SPECIAL TOPICS
30.4.1 Monitoring glacier change in the
northeastern Antarctic Peninsula
GLIMS objectives include measurements of
changes in glacier extent with the aim of establishing a digital baseline for future comparisons. Our
monitoring of glacier change using ASTER images
covers the northeastern side of the AP extending
between 63.8 S and 65.5 S (Fig. 30.1), and comprises different types of glaciers on Vega Island
and James Ross Island, Boydell Glacier and
Sjögren Glacier calving into Prince Gustav
Channel, Dinsmoor-Bombardier-Edgeworth and
Drygalski outlet glaciers calving into Larsen A
embayment, and Hektoria-Green-Evans, Jorum,
Crane, Mapple and Mellville Glaciers discharging
presently into Larsen B embayment (Fig. 30.1). The
first usable ASTER images of the northeastern AP
were acquired in early 2001 and used since then to
document the drastic glacier retreat and ice loss in
this region until 2008–2009. Of particular importance is the monitoring of glacier retreat behind
the grounding line (GL) because of its contribution
to global sea level rise. The position of the GL
was derived by means of different InSAR interferograms (Rott et al. 2002, Rack and Rott 2004).
ASTER images were co-registered with the
Landsat ETMþ mosaic of 21-Feb-2000 covering
the northeastern Antarctic Peninsula. This mosaic
was georeferenced to characteristic ground features
using GPS. Co-registration error was of the order of
0.5 pixels (i.e., 7.5 m). All areas and glacier
retreat measurements were performed manually
by means of commercial GIS software.
30.4.2 Glaciers of Vega Island and James
Ross Island
30.4.2.1
Vega Island (VI, Fig. 30.5)
Field measurements carried out on two glaciers
with termini on land—Glaciar Bahı́a del Diablo
(GBD) and Glaciar Cabo Lamb (GCL); Fig.
726
Monitoring glacier changes on the Antarctic Peninsula
Figure 30.5. ASTER image mosaic (bands 3, 2, 1) assembled with three L1B scenes acquired 03-Mar-2009,
showing Vega Island and James Ross Island. Blue indicates glacier retreat from 1988 to 2001 derived from Landsat
TM and ASTER images. Red and green indicate, respectively, retreats and advances from 2001 to 2009 derived only
from ASTER imagery. Pink shows the area of ice shelf disintegration in Röhss Bay between 2001 and 2009. Figure
can also be viewed as Online Supplement 30.2.
30.5—reveal major thinning rates of about 1.0 m
yr1 since the early 1980s (Skvarca and De Angelis
2003). Furthermore, detailed annual mass balance
measurements of GBD carried out as a contribution
to the World Glacier Monitoring Service (WGMS)
show negative values from 1999 to 2008, with an
average of 0.31 m water equivalent. An early
ASTER image acquired on 08-Jan-2001 with the
aid of a contour map derived from kinematic
GPS was used to help define the drainage area of
GBD.
30.4.2.2
James Ross Island (JRI, Fig. 30.5)
The baseline for glacier extent on JRI had already
been established in 1975 from Kosmos KATE-200
space photos (Skvarca et al. 1995). In total, 39
outlet glaciers draining JRI were analyzed, of which
33 were tidewater calving, one was freshwater
calving into a lake, and 5 were terminating on land.
As shown by Kosmos, Landsat, and ASTER
images the retreat rate of JRI glaciers has increased
considerably from 1.8 km 2 yr1 during the period
Case studies and special topics
1975–1988 to 2.9 km 2 yr1 in 1988–2001. Glacier/
ice shelf frontal positions and their change for the
latter period were also reported by Rau et al. (2004)
and Kargel et al. (2005), though they considered
fewer glaciers. Here we extend the study from
2001 to 2009, investigating the same glaciers as in
the earlier period. Further analysis is based on
ASTER images acquired on 08-Jan-2001 and 03Mar-2009 revealing a similar decrease rate of 2.6
km 2 yr1 and an ice loss of 21 km 2 . During the
overall period 1975–2009 the total glacier area of
JRI decreased by 81.3 km 2 . An additional 101 km 2
were lost from 2001 to 2009 due to disintegration of
a small ice shelf in Röhss Bay (Fig. 30.5). Overall,
our extended record shows sustained and accelerating retreat as well as drastic ice shelf breakup
similar to that which has affected the larger ice
shelves of the peninsula’s mainland.
30.4.3 Former tributaries of Prince
Gustav Channel (PGC) Ice Shelf
30.4.3.1
Boydell Glacier and Sjögren Glacier
(B-S)
Both glaciers fed the former ice shelf within PGC,
which disintegrated in 1994/1995. The ASTER
image of 26-Sep-2001 revealed characteristic
features such as surge waves, looped moraine,
and marginal crevasses, indicating an active surging
phase of B-S glaciers (De Angelis and Skvarca
2003). Evidence for strong glacier acceleration
and significant surface lowering/thinning after the
ice shelf collapse were the remnant ice terraces
detected at glacier margins. Further retreat from
late 2001 to early 2009 (Fig. 30.6) was derived using
a co-registered series of ASTER images, which
allowed us to compute major retreats of 7.8 km
(B) and 10.8 km (S), with significant total area loss
of 69.5 km 2 behind the grounding line derived by
interferometry.
30.4.4 Former tributaries of Larsen A Ice
Shelf
30.4.4.1
Dinsmoor-Bombardier-Edgeworth
(D-B-E) glaciers
This glacier system fed the former Larsen A Ice
Shelf before its collapse in early 1995. Surge waves
and looped moraines, characteristic features of an
active surging phase, had already been detected on
Bombardier Glacier and Edgeworth Glacier (Fig.
30.7) from the ASTER image of 26-Sep-2001 (De
727
Angelis and Skvarca 2003). However, this early
surge only affected these glaciers, which have
advanced 1.6 km with net areal gain of 6.5 km 2 .
A recent ASTER image acquired on 09-Dec-2008
together with ice front positions mapped using GPS
during aerial surveys carried out along the northeastern AP in early 2007 and 2008 reveal another
surging phase of the D-B-E glacier system, which
has advanced 1.8 km. This indicates that strong
dynamic perturbations are still affecting the former
tributaries 14 years after the removal of the ice
shelf. A recent surging event has also affected
Dinsmoor Glacier, which advanced with a net areal
increase of 5.1 km 2 from 11-Mar-2007 to 09-Dec2008 (Fig. 30.7). In this period, the B-E glaciers
surged by 6.5 km 2 (i.e., the same net areal gain as
during the 2000–2001 surge). By early December
2008, the three glaciers had lost an area of 39.5
km 2 behind the GL.
30.4.4.2
Drygalski Glacier (D on Fig. 30.1)
The first available ASTER image of Drygalski
Glacier after the disintegration of Larsen A was
acquired on 22-Nov-2001. The most recent ASTER
image allowed computation of an area loss of 34.5
km 2 behind its grounding line up to 09-Dec-2008.
Three smaller glaciers also draining into the Larsen
A embayment had lost an additional 36.5 km 2
behind their grounding lines by 09-Dec-2008. Both
areas represent a contribution to sea level rise.
30.4.5 Former tributaries of Larsen B
Ice Shelf
30.4.5.1
Hektoria-Green-Evans (H-G-E)
The H-G-E glacier system has been subject to continuous retreat since early 2002 (as sequential
ASTER images reveal); about 115.1 km 2 of its area
inland of the grounding line had been lost by 25Feb-2008 (Fig. 30.8).
30.4.5.2
Jorum-Crane-Mapple-Melville
(J-C-M-M)
Farther south, five ASTER images provide information on the glacier extent of Crane, Mapple, and
Melville, the three southernmost glaciers that fed
the former Larsen B Ice Shelf and calve at present
into its embayment (Fig. 30.9). Crane Glacier with
60 km in length is the longest in the AP and
retreated 10.8 km between late 2002 and early
2008 with an area loss of 57.5 km 2 , 35.0 km 2 of
728
Monitoring glacier changes on the Antarctic Peninsula
Figure 30.6. Section of ASTER image of 28-Oct-2006 (bands 3, 2, 1) showing the retreat of B-S glaciers behind
the grounding line (GL) since 26-Sep-2001 (PGC ¼ Prince Gustav Channel). Figure can also be viewed in higher
resolution as Online Supplement 30.3.
which was behind the GL. Farther south, Mapple is
the only glacier whose ice front is still flowing outward from the grounding line, hence is not yet contributing to sea level rise, while Melville Glacier had
lost by early 2008 an area of about 1.9 km 2 behind
the GL. Jorum Glacier and two small neighboring
glaciers located north of Crane have also retreated
behind the GL (losing an area of 26.5 km 2 ).
In total, all glaciers calving into the Larsen B
embayment had lost by early 2008 about 178.5
km 2 of their area behind their grounding lines,
representing a contribution to global sea level rise.
30.4.6 Monitoring changes and breakup
events on the Wilkins Ice Shelf
30.4.6.1
Areal changes and breakups between
1986 and 2009
Wilkins Ice Shelf (WIS) is located in the southwestern part of the AP. It is currently the southern-
most ice shelf on the peninsula to show breakup
events according to the definition given by Braun
et al. (2009). We define breakup as a sudden fast
release of fragments of variable size, happening on a
timescale of hours to days; in contrast to disintegration where the complete ice shelf is lost. Calving is
seen as an ordinary process of mass loss of an ice
shelf on a timescale of months or years; while
retreat is a reduction in size on a timescale of
months or years.
WIS is confined by several islands—namely,
Alexander, Rothschild, Charcot, and Latady (Fig.
30.10). Additionally, numerous ice rises intersect
the ice shelf. The primarily remote sensing–based
analysis of Vaughan et al. (1993) suggested that the
mass balance of WIS was dominated by surface
accumulation and basal melting, and that as a
consequence it might be particularly prone to variations in atmospheric and oceanic boundary conditions. Some inflow occurs from Lewis Snowfield
and into Schubert Inlet and Haydn Inlet. However,
Case studies and special topics
729
Figure 30.7. Section of the ASTER image acquired 02-Dec-2008 (bands 3, 2, 1) showing recent surge and D-B-E
ice front fluctuations since 2001 and retreat inland behind the GL. Front positions on 11-Feb-2006, 11-Mar-2007,
and 03-Mar-2008 are from GPS surveys, while the 26-Sep-2001 and 02-Oct-2003 front positions are from ASTER.
Figure can also be viewed in higher resolution as Online Supplement 30.4.
Figure 30.8. Retreat of Hektoria-Green-Evans (H-G-E) glaciers shown on a section of the ASTER image acquired
27-Sep-2004 (bands 3, 2, 1). Figure can also be viewed in higher resolution as Online Supplement 30.5.
730
Monitoring glacier changes on the Antarctic Peninsula
Figure 30.9. Section of ASTER image acquired 25-Feb-2008 (bands 3, 2, 1). Numbers in circles indicate ice front
positions on (1) 07-Nov-2002; (2) 02-Feb-2003; (3) 13-Jan-2004; (4) 27-Sep-2004 and (5) 25-Nov-2006.
Figure can also be viewed in higher resolution as Online Supplement 30.6.
the flow from these inlets into the main ice shelf is
considerably blocked by ice rises.
While the situation between 1986 and 1990 indicated a stable ice shelf (observed in Landsat
imagery), by 1993 the northern ice front was
reported to have retreated several times by Lucchita
and Rosanova (1998). A further massive area loss
occurred during February 1998, when approximately 1,100 km 2 were lost (Scambos et al. 2000).
Up to February 2008 the extent of the ice shelf
remained almost constant and the area amounted
to about 13,000 km 2 . However, there was a considerable increase in the number of fractures arising
from tensile stresses in the vicinity of ice rises. A
review of the state of WIS by Braun et al. (2009)
revealed that several distinct breakup events in contrast to frequent calving (including the abovementioned) occurred prior to 2008. The authors
concluded that increasing basal melt rates due to
variations in the ocean regime and changes in the
properties of materials comprising WIS due to
atmospheric and ocean changes could be two possible factors responsible for reducing the integrity of
the ice shelf.
In the course of 2008, WIS underwent considerable changes, including three distinct breakup
events (Humbert and Braun 2008; Braun and Humbert 2009; Scambos et al. 2009). On February 28,
2008 the first breakup event started along the northwestern ice front between Charcot Island and
Latady Island. The signature of Envisat C-band
SAR (synthetic aperture radar) data revealed a
moist or wet surface. In contrast, during the second
breakup event on May 30/31, 2008 the bright SAR
backscatter indicated frozen conditions. The
hypothesis of meltwater from melt pools draining
into crevasses was hence ruled out as a potential
explanation. A third breakup event followed in July
Case studies and special topics
during austral winter, leaving a fragile connection
to Charcot Island only 900 m wide at its narrowest
point. The remaining ice shelf bridge finally collapsed in the first days in April 2009 (Humbert et
al. 2010), followed by destabilization of the northern ice front.
An overview of ice fronts from 1990 to the various stages is shown in Fig. 30.10. Scambos et al.
(2000) and Cook and Vaughan (2010) based their
analyses on historic maps and declassified satellite
imagery, which showed that the ice front of WIS did
not change between 1947 and 1986 in any considerable way. Table 30.2 summarizes the respective
area losses. The numbers differ slightly from the
data given by Cook and Vaughan (2010) in their
overview of Antarctic Peninsula ice shelf areal
changes as a result of differences in image dates
and potential differences in interpretation of the
heavily fractured ice shelf front. WIS has lost a total
of about 5,624 km 2 of its area since 1986, which
amounts to about 43% of its size before the onset of
breakups.
30.4.6.2
A three-step process chainfor breakups
In the past, various explanations for ice shelf
breakup and retreat have been proposed. They
cover individual factors or distinct mechanisms
leading to failure. We propose a general three-stage
process chain consisting of a cause, a trigger, and a
consequence (Fig. 30.11). It can be applied independently of the specific mechanisms of each step and
structures the breakup process along a timeline. The
primary process leads to fracture formation, the
secondary step triggers propagation of the crack
through the entire ice shelf, and the tertiary process
involves the release of icebergs. Identifying the
mechanisms involved in the first two steps is particularly in need of improvement in order to understand what lies behind breakup processes and what
their potential consequences are. A known primary
process is the formation of crevasses by flowinduced stresses, like shear margins or crevasses
in the vicinity of ice rises (Doake and Vaughan
1991, Glasser and Scambos 2008). For WIS, Braun
and Humbert (2009) proposed that fracturing was
caused by bending stresses resulting from buoyancy
forces induced by an ice thickness gradient. This
gradient arose presumably from inhomogeneous
basalt melt rates, where the seaward margin of
the ice plate experienced more basal melt than areas
farther inward. Scambos et al. (2009), on the other
hand, proposed a plate bending stress model
731
induced by buoyancy forces due to tides and the
stress boundary condition at the seaward margin,
similar to stress fields along calving fronts.
Perhaps the most prominent secondary process is
the drainage of surface melt ponds into crevasses,
which increases stress at the crack tip. Subsequently, cracks propagate vertically through the
entire ice shelf (Weertman 1973, Van der Veen
1998, Scambos et al. 2000, 2003). This secondary
process is widely acknowledged to have taken place
during the breakup of Larsen B. Scambos et al.
(2009) propose a model that replaces surface melt
water by brine for the February 2008 WIS breakup.
However, as the May and June/July 2008 breakups
occurred during austral winter, when the availability of liquid was unlikely, further secondary
processes likely exist.
Tertiary processes—such as the release of ice
fragments—may include icebergs that capsize, as
shown by MacAyeal et al. (2003) for the breakup
of Larsen B. The fast rotation of icebergs as they
capsize leads to a rapid massive area increase. For
all three breakups of WIS in 2008, the failure of the
ice bridge, and the subsequent mass loss at the
northern ice front, this process took place. This
tertiary process releases a large amount of energy
(MacAyeal et al. 2009), as the capsizing of icebergs
decreases potential energy. Guttenberg et al. (2011)
presented computations on this conceptual model
that was not tied to a specific ice shelf or observation data. They also investigated how the aspect
ratio of icebergs and its variance influence breakup
conditions. It is important to link this tertiary process to fracture patterns observed before breakup.
The failure of the ice bridge on WIS in 2009 liberated energy of the order of 10 14 J (as shown by
Humbert et al. 2010).
30.4.7 Variation of radar glacier zone
boundaries in the northeastern
Antarctic Peninsula
One of the biggest problems in building a time series
of GLIMS analyses for the AP is the lack of a
consistent spatiotemporal series of optical datasets.
An alternative is to use the database of SAR images
available for this region. This case study involves
using a time series of ERS-1/2 SAR images to ascertain the boundaries between radar glacier zones
such as the bare ice radar zone (BIRZ), wet snow
radar zone (WSRZ), frozen percolation radar zone
(FPRZ), and dry snow radar zone (DSRZ), which
can be used as proxies for GLIMS-related param-
732
Monitoring glacier changes on the Antarctic Peninsula
Figure 30.10. (Upper) Overview map of Wilkins Ice Shelf based on a Landsat mosaic from 1990 (& USGS 1990).
The features labeled are B ¼ Burgess Ice Rise, P ¼ Petrie Ice Rises, V ¼ Vere Ice Rise. The purple line indicates the
present extent, the black line the grounding line. (Lower) The situation as depicted by a TerraSAR-X ScanSAR
image on 02-Nov-2009 (& DLR 2009) superimposed on two Envisat ASAR wideswath images from 26-Jun-2009
and 30-Jun-2009 (& ESA 2009). Ice front positions on specific dates are given in color, the black line shows the
grounding line. Figure can also be viewed in higher resolution as Online Supplement 30.7.
Case studies and special topics
733
Table 30.2. Compilation of the retreat area of Wilkins Ice Shelf since 1986.
Northern ice front
Northwestern ice front
Southern ice front
Period/year
Retreat area
(km 2 )
Period/year
Retreat area
(km 2 )
Period/year
Retreat area
(km 2 )
1986–1990
97
1990–1993
57
1990–2004
196
1990–1991
655
1993–1998
0
1991–1992
0
1998–1999
20
1992–1993
544
1999–2000
87
1993–1998
0
2000–2001
0
1998–1999
1,100
2001–2003
52
1999–2008
0
2003–2004
51
2004–2009
74
2008
1,220
2004–2007
0
2008
585
Sum
948
Sum
270
2009
790
Sum
4,406
eters such as snowline position, area of the ablation
zone, etc. The concept of radar glacier zones used
here corresponds to the classification scheme proposed by Rau et al. (2000) and further discussed by
Braun et al. (2000) and Arigony-Neto et al. (2007,
2009).
The satellite data used for analysis consisted of
eight images acquired on the northeastern part of
the AP (frame 4,923, tracks 109 and 381) from 1993
to 2000 by ERS-1/2 SAR (Fig. 30.12) C-band VVpolarization instruments. Normalized backscattering coefficients (0 ) were calculated using the Basic
Envisat and ERS SAR Toolbox (BEST) from the
European Space Agency (ESA). BEST uses the
SAR calibration algorithm developed by Laur et
al. (2004), who estimated that resulting values of
0 have an accuracy of 0.4 dB. The excellent stability of SAR sensors enables direct comparison of
calibrated data (Meadows et al. 1998). To reduce
the speckle effect, a 5 5 median filter was applied.
Finally, speckle-filtered images were orthorectified
using the DEM from RAMP (Liu et al., 2001).
Image analyses were carried out automatically by
means of a knowledge-based image analysis algorithm modified from Arigony-Neto et al. (2007).
These authors developed an approach based on
classifying areas located in a 600 m buffer along
glacier centerlines. In this study, we adapted this
method so that it considered whole glacier basins.
Rules used for pixel classification were mainly
based on backscattering thresholds determined by
Rau et al. (2001) and altitude thresholds discussed
in Arigony-Neto et al. (2009) (Table 30.3). Elevation information was derived from the RAMP
DEM. Rocks and ocean areas were masked by rock
outcrops and the coastline according to the Antarctic Digital Database (ADD; SCAR 2010). Misclassified pixels usually corresponded to steep
slopes where AP plateaus break down to become
glaciers or crevasse fields inside major classes such
as the wet snow radar zone. These pixels were
eliminated using a focal majority 5 5 filter.
Fig. 30.12 shows the resulting distribution of
radar glacier zones during the time period of image
acquisitions. Analyses using images acquired on
track 381 were restricted to the area common to
track 109. In general, the boundaries of radar
glacier zones presented enormous variation in the
time period of study. This can be explained by the
fact that the development of superficial zones or
734
Monitoring glacier changes on the Antarctic Peninsula
Figure 30.11. Schematic of the three-step process during ice shelf breakups using the ice bridge on WIS as an
example. While the upper panel illustrates a vertical view of the ice bridge with initial rifts formed during July 2007,
the middle and lower panels represent cross-sections through the ice bridge along the thick black line of the upper
panel. In the middle part the dotted lines denote perforation of the complete ice plate during this process, while the
lower panel illustrates the area gain by sliver icebergs capsizing.
facies on glaciers is greatly influenced by local and
regional climatic and meteorological settings
(Braun et al. 2000, Rau et al. 2000, Arigony et al.
2007, 2009). Indeed, the boundaries and extensions
of radar glacier zones do not correspond to images
acquired in the same period of the year, as in 13Feb-1993 and 11-Feb-1997 (Fig. 30.12B and
30.12F). For example, in the image acquired on
13-Feb-1993, major extensions of the wet snow
radar zone (WSRZ) were observed in the coastal
regions of Trinity Peninsula and James Ross Island,
while the image from 11-Feb-1997 shows the
WSRZ restricted to the northwestern tip of Trinity
Peninsula. From analysis of the development of
radar glacier zones during austral summer 1996/
1997, it is possible to observe evolution of the BIRZ
from the previous year (Fig. 30.12C) in the FPRZ
after the first melt–freeze spring events (Fig.
30.12D). Subsequently, melt occurs in the snowpack covering most low-altitude areas along the
east coast of the AP and parts of James Ross Island
and Vega Island, and at the end of the summer
(Fig. 30.12H) it is possible to record the maximum
extension of the BIRZ for that balance year. Such
variations indicate that great expanses of saturated
snowpack melt during the summer. This event is
probably related to the high frequency of days with
positive temperatures recorded for the period
between 29-Oct-1996 and 18-Mar-1997 at meteorological stations located in the region (72, 30, and 74
days, respectively, for Esperanza, Marambio, and
O’Higgins Stations). Furthermore, long periods
with positive temperatures during summer 1996/
1977 are confirmed by the relatively high mean
summer temperatures recorded for Esperanza
(0.9 C), Marambio (1.46 C), and O’Higgins
(0.5 C). Skvarca and De Angelis (2003) also confirm the relationship between summer mean air
temperatures and mass balance for this region.
The small area of the BIRZ detected on 03-Oct2000 is arguably an underestimate of the total
extension of this radar glacier zone at the end of
the balance year 1999/2000, because records of
mean temperatures for austral summer 1999/2000
appeared to be higher at the meteorological stations
mentioned above.
Case studies and special topics
735
Figure 30.12. Area corresponding to footprints of ERS-1/2 SAR imagery acquired on frame 4,923, track 109 (area
1 in A) and track 381 (area 2 in A). The coastline is represented by the black continuous line and rocks are
represented by light-gray polygons (B-I ¼ thematic maps resulting from image classification; BIRZ ¼ bare ice radar
zone; WSRZ ¼ wet snow radar zone; and FPRZ ¼ frozen percolation radar zone). Figure can also be viewed in higher
resolution as Online Supplement 30.8.
736
Monitoring glacier changes on the Antarctic Peninsula
Table 30.3. Thresholds for backscattering coefficients ( 0 ) and
altitude used for the classification of radar glacier zones on the
Antarctic Peninsula.
Radar glacier zones
Dry snow radar zone
Frozen percolation radar zone
Wet snow radar zone
Bare ice radar zone
30.5
REGIONAL SYNTHESIS
The stereoscopic capability of the ASTER sensor
offers the opportunity to produce medium-scale
spatial resolution digital elevation models of the
AP compatible with the requirements of GLIMS.
Application of spaceborne optical imagery from the
AP, in particular, is often hindered by frequent
cloud cover. The incomparably large number of
(almost) cloud-free ASTER scenes for Marguerite
Bay can presumably be attributed to an atmospheric circulation pattern where depression centers
pass the northern part of the AP from west to east
and produce a lee condition westward of the central
plateau in this area. All these aspects supported the
view that the area was the best available test site
providing a suitable location against which DEMs
produced from ASTER data using our single and
multitemporal scenes approach can be compared.
What is more, ASTER images provide a digital
baseline for detection of future change in glacier
extent in one of the most rapidly changing regions
on Earth. As a result of analyzing these images we
were able to compute as of 2008–2009 the retreat
area behind the grounding lines of glaciers calving
into Prince Gustav Channel, the Larsen A embayment, and the Larsen B embayment at about
358.5 km 2 , all of which contribute to global sea
level rise.
ASTER imagery acquired at the time of and
before the February 2008 breakup of Wilkins Ice
Shelf contributed considerably to documentation
and understanding of the events. WIS is likely to
experience more mass loss, although around 8,000
km 2 are currently (2013) in a stable condition. The
first signs of the stress field adjusting to the new
load situation after the failure of the ice bridge were
already visible while this chapter was being written,
0
(dB)
Altitude H
(m)
14 > 0 > 20
H > 1,200
0 > 0 > 8
14 > 0 > 25
6 > 0 > 13
—
H < 1,200
H < 500
suggesting that a future ice front will likely meet the
worst case predicted by Braun et al. (2009). In contrast to breakups in the 1990s, the field of icebergs
has dispersed considerably, making formation of an
extensive ice melange area unlikely. Thus, albedo
and ocean will presumably respond to the 2008/
2009 series of retreats faster and stronger than they
did to breakups in the 1990s.
The dynamics of radar glacier zones in the northern AP are related to high interannual and seasonal
climate variability in this region. The variability of
surface air temperature and high frequency of days
with positive temperature recorded at meteorological stations in the northern tip of the AP caused
significant changes in the position and extension of
radar glacier zones during the period of study.
Therefore, the position and altitude of the boundaries of glacier zones can be used as proxies for
variations in temperature for regions of the AP
where no meteorological data are available.
30.6
SUMMARY AND CONCLUSIONS
DEM evaluation suggests that stereo along-track
data from Terra’s ASTER instrument can be used
to generate reasonably accurate models for the AP,
improving the spatial resolution of existing models.
A successful generation of digital elevation models
for specific sectors of the AP region will contribute
to filling current gaps in topographic data. Based on
these DEMs it should also be possible to perform
geometric correction of satellite data and use them
for subsequent semiautomatic derivation of ice
drainage catchments and additional parameters
required for GLIMS analysis.
References 737
The glaciers calving into Prince Gustav Channel,
the Larsen A embayment, and the Larsen B embayment showed a general pattern of retreat behind
their grounding lines between 2001 and 2008–
2009, immediately after a period in which some
glaciers actively surged in the early 2000s. A few
of these glaciers started to surge again between 2007
and 2008, indicating that strong dynamic perturbations are still affecting the former tributary glaciers
many years after the removal of ice shelves. Future
studies should take advantage of available sequential ASTER images to derive ice velocities, another
GLIMS objective, so that the dynamics of these
highly active calving glaciers, affected as they are
by the removal of ice shelves, can be better understood. Further analysis of glacier change in the
northeastern AP using satellite data is necessary
because of its potential impact on global sea level
rise.
The breakup events of Wilkins Ice Shelf in 2008
and 2009 are the most recent signs of its destabilization, which began in the 1990s. Development of
failure zones in the vicinity of ice rises during the
15 years prior to the events and bending stresses
induced by buoyancy forces made this ice shelf
vulnerable to extensive mass loss. Suggested reasons
for this development are increased basal melt as
well as thermal and precipitation-induced changes
in the material properties of ice.
The recorded patterns of variations in the boundaries of radar glacier zones in low-altitude areas
within the time period of the study (1993–2000),
associated with high-altitude variations in the
DSRZ detected by Arigony-Neto et al. (2009), show
that, in contrast to the Antarctic Ice Sheet, variations in climatological and glaciological conditions
on a relatively short timescale are typical for this
region. These results validate the suitability of SAR
data to derive superficial information about glaciers
to be used as GLIMS parameters in areas where
optical imagery is not available, or for providing
complementary information for temporal analyses
of glacier change.
The case studies discussed in this chapter demonstrated some of the satellite-based approaches used
nowadays to monitor the ice masses of one of the
most dynamic ice-covered areas of the planet. The
multitude of glacier types in the AP and the
different scales of analyses needed to understand
the glaciological processes happening in this region
necessitate the use of diverse satellite datasets and
ancillary data to address questions related to the
interaction among glaciers, climate, and oceans.
Furthermore, it calls attention to the need for and
potential of synergic approaches between GLIMS
activities and other satellite-based investigations of
AP glaciers.
30.7
ACKNOWLEDGMENTS
This work was partially supported under grants BR
2105/4-1/2/3 and BR 2105/8-1 from the German
Research Foundation, and grants CNPq 480701/
2008-3 and CNPq 573720/2008-8 (Brazilian
National Institute for Cryospheric Sciences) from
the Brazilian National Council for Scientific and
Technological Development. TerraSAR-X data
were acquired by the TerraSAR-X background mission entitled ‘‘Antarctic Peninsula and Antarctic Ice
Shelves’’ as well as under proposal LAN_0013 from
the German Aerospace Center (DLR). Envisat
ASAR data were kindly provided under ESA IPY
AO 4032, ERS-1/2 data were available within the
CryoSat Data AO Project 2658, and ASTER data
were provided in the context of the GLIMS project.
ASTER data courtesy of NASA/GSFC/METI/
Japan Space Systems, the U.S./Japan ASTER
Science Team, and the GLIMS project.
30.8
REFERENCES
Arigony-Neto, J., Rau, F., Saurer, H., Jaña, R., Simões,
J.C., and Vogt, S. (2007) A time series of SAR data for
monitoring changes in boundaries of glacier zones on
the Antarctic Peninsula. Annals of Glaciology, 46, 55–
60.
Arigony-Neto, J., Saurer, H., Simões, J.C., Rau, F., Jaña,
R., Vogt, S., and Gossmann, H. (2009) Spatial and
temporal changes in dry-snow line altitude on the
Antarctic Peninsula. Climatic Change, 94, 19–33.
Bamber, J., and Gomez-Dans, J.L. (2005) The accuracy
of digital elevation models of the Antarctic Continent.
Earth and Planetary Science Letters, 237, 516–523.
Benn, D.I., Warren, C.R., and Mottram, R.H. (2007)
Calving processes and the dynamics of calving glaciers.
Earth Science Reviews, 82, 143–179.
Birkenmaier, K. (1994) Evolution of the Pacific margin
of the northern Antarctic Peninsula: An overview.
Geologische Rundschau, 83(2), 309–321.
Braun, M., and Goßmann, H. (2002) Glacial changes in
the areas of Admiralty Bay and Potter Cove, King
George Island, Maritime Antarctica. In: I. Beyer and
M. Bölter (Eds.), Geoecology of Antarctic Ice-Free
Coastal Landscapes, Springer Verlag, Heidelberg,
Germany, pp. 75–89.
738
Monitoring glacier changes on the Antarctic Peninsula
Braun, M., and Humbert, A. (2009) Recent retreat of
Wilkins Ice Shelf reveals new insights in ice shelf
break-up mechanisms. Geosciences and Remote Sensing
Letters, 6(2), 263–267, doi: 10.1109/LGRS.2008.
2011925.
Braun, M., Rau, F., Saurer, H., and Goßmann, H. (2000)
Development of radar glacier zones on the King
George Island ice cap, Antarctica, during the austral
summer 1996/97 as observed in ERS-2 SAR-data.
Annals of Glaciology, 31, 357–363.
Braun, M., Humbert, A., and Moll, A. (2009) Changes of
Wilkins Ice Shelf over the past 15 years and inferences
on its stability. The Cryosphere Discussions, 3, 41–56.
Calvet, J., Garcı́a Sellés, D., and Corbera, J. (1999) Fluctuaciones de la extensión del casquete glacial de la isla
Livingston (Shetland del Sur) desde 1956 hasta 1996.
Acta Geologica Hispanica, 34(4), 365–374 [in Spanish].
Chen, J.L., Wilson, C.R., Blankenship, D., and Tapley,
B.D. (2009) Accelerated Antarctic ice loss from satellite gravity measurements. Nature Geoscience, doi:
10.1038/NGEO694.
Cook, A.J., and Vaughan, D.G. (2010) Overview of areal
changes of the ice shelves on the Antarctic Peninsula
over the past 50 years. The Cryosphere Discussions, 3,
579–630.
Cook, A.J., Fox, A.J., Vaughan, D.G., and Ferrigno,
J.G. (2005) Retreating glacier fronts on the Antarctic
Peninsula over the past half-century. Science,
308(5721), 541–544.
De Angelis, H., and Skvarca, P. (2003) Glacier surge after
ice shelf collapse. Science, 299, 1560–1562.
Doake, C.S.M., and Vaughan, D.G. (1991) Rapid disintegration of Wordie Ice Shelf in response to atmospheric warming. Nature, 350, 328–330.
Fahnestock, M.A., Abdalati, W., and Shuman, C.A.
(2002) Long melt seasons on ice shelves of the Antarctic Peninsula: An analysis using satellite-based microwave emission measurements. Annals of Glaciology, 34,
127–133.
Glasser, N.F., and Scambos, T.A. (2008) A structural
glaciological analysis of the 2002 Larsen B Ice Shelf
collapse. Journal of Glaciology, 54(184), 3–16.
Guttenberg, N., Abbot, D.S., Amundson, J.M., Burton,
J.C., Cathles, L.M., MacAyeal, D.R., and Zhang,
W.W. (2011) A computational investigation of iceberg
capsize as a driver of explosive ice-shelf disintegration.
Annals of Glaciology, 52(59), 51–59.
Hambrey, M.J., Smellie, J.L., Nelson, A.E., and Johnson,
J.S. (2008) Late Cenozoic glacier–volcano interaction
on James Ross Island and adjacent areas, Antarctic
Peninsula region. Geological Society of America Bulletin, 120(5/6), 709–731.
Harrison, C.G.A., Barron, E.J., and Hay, W.W. (1979)
Mesozoic evolution of the Antarctic Peninsula and the
southern Andes. Geology, 7(8), 374–378.
Humbert, A., and Braun, M. (2008) Wilkins Ice Shelf:
Break-up along failure zones. Journal of Glaciology,
54(188), 943–944.
Humbert, A., Gross, D., Müller, R., Braun, M., van de
Wal, R.S.W., van den Broeke, M.R., Vaughan, D.G.,
and van de Berg, W.J. (2010) Deformation and failure
of the ice bridge on Wilkins Ice Shelf, Antarctica.
Annals of Glaciology, 51(55), 49–55.
If PK TUD (1999) Aerial Photo Map—1:50.000: Base
General San Martin, Baie Marguerite, Antarctic Peninsula (produced by the Dynamische Prozesse in antarktischen Geosystemen, DYPAG, research project done
in collaboration with the Bundesamt für Kartographie
und Geodäsie, BKG, and the Institut für Physische
Geographie, IPG), Albert-Ludwigs-Universität Freiburg and supported by the Bundesministerium für
Bildung, Wissenschaft Forschung und Technologie),
Institut für Photogrammetrie und Kartographie Technische, Universität Darmstadt, Germany.
Kääb, A. (2005) Combination of SRTM3 and repeat
ASTER data for deriving alpine glacier flow velocities
in the Bhutan Himalaya. Remote Sensing of Environment, 94, 463–474.
Kargel, J.S., Abrams, M.J., Bishop, M.P., Bush, A.,
Hamilton, G., Jiskoot, H., Kääb, A., Kieffer, H.H.,
Lee, E.M., Paul, F. et al. (2005) Multispectral imaging
contributions to Global Land Ice Measurements from
Space. Remote Sensing of Environment, 99, 187–219.
Laur, H., Bally, P., Meadows, P., Sanchez, J., Schättler,
B., Lopinto, E., and Esteban, D. (2004) Derivation of
the Backscattering Coefficient 0 in ESA ERS SAR PRI
Products (ESA/ESRIN, ESTN-RS-PM-HL09, Issue 2,
Rev. 5f ), ESA European Space Research Institute
(ESRIN), Frascati, Italy, 53 pp.
Lawyer, L.A. (2009) Distributed, active extension in
Bransfield Basin, Antarctic Peninsula: Evidence from
multibeam bathymetry. Geological Society of America
Today, 6(11), 1–6.
Liu, H., Jezek, K., and Li, B. (1999) Development of
Antarctic digital elevation model by integrating cartographic and remotely sensed data: A geographic information system based approach. Journal of Geophysical
Research, 104, 23199–23213.
Liu, H., Jezek, K., Li, B., and Zhao, Z. (2001) Radarsat
Antarctic Mapping Project Digital Elevation Model
Version 2, National Snow and Ice Data Center,
Boulder, CO [digital media].
Lucchita, B.K., and Rosanova, C.E. (1998) Retreat of
northern margins of Georg VI and Wilkins Ice Shelves,
Antarctic Peninsula. Annals of Glaciology, 27, 41–46.
MacAyeal, D.R., Scambos, T.A., Hulbe, C., and Fahnestock, M.A. (2003) Catastrophic ice-shelf break-up by
an ice-shelf-fragment-capsize mechanism. Journal of
Glaciology, 49, 22–36.
References 739
MacAyeal, D.R., Okal, E.A., Aster, R.C., and Bassis,
J.N. (2009) Seismic observations of glaciogenic
ocean waves (micro-tsunamis) on icebergs and ice
shelves. Journal of Glaciology, 55(190), 193–206.
Marshall, G.J., Lagun, V., and Lachlan-Cope, T.A.
(2002) Changes in Antarctic Peninsula tropospheric
temperatures from 1956 to 1999: A synthesis of observations and reanalysis data. International Journal of
Climatology, 22(3), 291–310.
Meadows, P., Laur, H., and Schättler, B. (1998) The
calibration of ERS SAR imagery for land applications.
In: M. Borgeaud and T.D. Guyenne (Eds.), Retrieval of
Bio- and Geo-Physical Parameters from SAR Data
for Land Applications: Proceedings of the Second International Workshop, October 21–23, 1998 (ESA SP441), ESA/ESTEC, Noordwijk, The Netherlands,
pp. 35–42.
Millar, I.L., Pankhurst, R.J., and Fanning, C.M. (2002)
Basement chronology of the Antarctic Peninsula:
Recurrent magmatism and anatexis in the Palaeozoic
Gondwana Margin. Journal of the Geological Society,
159, 145–157.
Morris, E.M., and Vaughan, D.G. (2003) Spatial and
temporal variation of surface temperature on the
Antarctic Peninsula and the limit of viability of ice
shelves. In: E. Domack, A. Leventer, A. Burnett, R.
Bindschadler, P. Convey, and M. Kirby (Eds.),
Antarctic Peninsula Climate Variability: Historical
and Paleoenvironmental Perspectives (Antarctic
Research Series Vol. 79), American Geophysical
Union, Washington, D.C., pp. 61–68.
Park, B.-K., Chang, S.-K., Yoon, H.I., and Chung, H.
(1998) Recent retreat of ice cliffs, King George Island,
South Shetland Islands, Antarctic Peninsula. Annals of
Glaciology, 27, 633–635.
Paul, F. (2001) Evaluation of different methods for glacier
mapping using Landsat TM. EARSeL eProceedings, 1,
239–245.
Paul, F., Kääb, A., Maisch, M., Kellenberger, T., and
Haeberli, W. (2002) The new remote-sensing-derived
Swiss glacier inventory, I: Methods. Annals of Glaciology, 34, 355–361.
Peel, D.A. (1992) Spatial temperature and accumulation
rate variations in the Antarctic Peninsula. In: E.M.
Morris (Ed.), The Contribution of Antarctic Peninsula
Ice to Sea Kevel Rise (BAS Ice & Climate Special
Report 1), British Antarctic Survey, Cambridge,
U.K., pp. 11-15.
Pfeffer, W.T. (2003) Tidewater glaciers move at their own
pace. Nature, 426(6967), 602.
Pritchard, H.D., and Vaughan, D.G. (2007) Widespread
acceleration of tidewater glaciers on the Antarctic
Peninsula. Journal of Geophysical Research, 112,
F03S29.
Rack, W., and Rott, H. (2004) Pattern of retreat and
disintegration of Larsen B Ice Shelf, Antarctic Peninsula. Annals of Glaciology, 39, 505–510.
Rau, F., and Braun, M. (2002) The regional distribution
of the dry snow zone on the Antarctic Peninsula north
of 70 South. Annals of Glaciology, 34, 95–100.
Rau, F., Braun, M., Saurer, H., Goßmann, H., Kothe,
G., Weber, F., Ebel, M., and Beppler, D. (2000)
Monitoring multi-year snow cover dynamics on the
Antarctic Peninsula. Polarforschung, 67(1/2), 27–40.
Rau, F., Braun, M., Friedrich, M., Weber, F., and
Goßmann, H. (2001) Radar glacier zones and their
boundaries as indicators of glacier mass balance and
climatic variability. EARSeL eProceedings, 1(1), 317–
327.
Rau, F., Mauz, F., De Angelis, H., Jaña, R., ArigonyNeto, J., Skvarca, P., Vogt, S., Saurer, H., and
Goßmann, H. (2004) Variations of glacier frontal positions on the Northern Antarctic Peninsula. Annals of
Glaciology, 39, 525–530.
Raymond, C., Weertman, B., Thompson, L., MosleyThompson, E., Peel, D., and Mulvaney, R. (1996)
Geometry, motion and mass balance of Dyer Plateau,
Antarctica. Journal of Glaciology, 42, 510–518.
Reynolds, J.M. (1981) The distribution of mean annual
temperatures in the Antarctic Peninsula. British Antarctic Survey Bulletin, 54, 123–133.
Ridley, J. (1993) Surface melting on Antarctic Peninsula
ice shelves detected by passive microwave sensors.
Geophysical Research Letters, 20(23), 2639–2642.
Rignot, E., Casassa, G., Gogineni, P., Krabill, W.,
Rivera, A., and Thomas, R. (2004) Accelerated ice
discharge from the Antarctic Peninsula following the
collapse of Larsen B ice shelf. Geophysical Research
Letters, 31(18), L18401.
Rignot, E., Casassa, G., Gogineni, S., Kanagaratnam, P.,
Krabill, W., Pritchard, H., Rivera, A., Thomas, R.,
Turner, J., and Vaughan, D.G. (2005) Recent ice loss
from the Fleming and other glaciers, Wordie Bay, West
Antarctic Peninsula. Geophysical Research Letters,
32(7), L07502.
Rott, H., Skvarca, P., and Nagler, T. (1996) Rapid collapse of northern Larsen Ice Shelf, Antarctica. Science,
271, 788–792.
Rott, H., Rack, W., Nagler, T., and Skvarca, P. (1998)
Climatically induced retreat and collapse of northern
Larsen Ice Shelf, Antarctic Peninsula. Annals of
Glaciology, 27, 86–92.
Rott, H., Rack, W., Skvarca, P., and De Angelis, H.
(2002) Northern Larsen Ice Shelf: Further retreat after
the collapse. Annals of Glaciology, 34, 277–282.
Scambos, T.A., Hulbe, C., Fahnestock, M., and
Bohlander, J. (2000) The link between climate warming
and break-up of ice shelves in the Antarctic Peninsula.
Journal of Glaciology, 46(154), 516–530.
Scambos, T.A., Hulbe, C., and Fahnestock, M. (2003)
Climate-induced ice shelf disintegration in the Antarctic Peninsula. In: E. Domack, A. Leventer, A. Burnett,
R. Bindschadler, P. Convey, and M. Kirby (Eds.),
Antarctic Peninsula Climate Variability: Historical
740
Monitoring glacier changes on the Antarctic Peninsula
and Paleoenvironmental Perspectives (Antarctic
Research Series Vol. 79), American Geophysical
Union, Washington, D.C., pp. 77–92.
Scambos, T.A., Bohlander, J.A., Shuman, C.A., and
Skvarca, P. (2004) Glacier acceleration and thinning
after ice shelf collapse in the Larsen B embayment,
Antarctica. Geophysical Research Letters, 31(18), doi:
10.1029/2004GL020670.
Scambos, T.A., Fricker, H.A., Liu, C.-C., Bohlander, J.,
Fastook, J., Sargent, A., Massom, R., and Wu, A.-M.
(2009) Ice shelf disintegration by plate bending and
hydro-fracture: Satellite observations and model
results of the 2008 Wilkins ice shelf break-ups. Earth
and Planetary Science Letters, 280(1/4), 51-60.
SCAR (2010) Antarctic Digital Database. Available at
http://www.add.scar.org [Scientific Committee on
Antarctic Research].
Shepherd, A., Wingham, D., Payne, T., and Skvarca, P.
(2003) Larsen Ice Shelf has progressively thinned.
Science, 302(5646), 856–859.
Simões, J.C., Bremer, U.F., Aquino, F.E., and Ferron,
F.A. (1999) Morphology and variations of glacial
drainage basins in King George Island ice field,
Antarctica. Annals of Glaciology, 29, 220–224.
Skvarca, P. (1993) Fast recession of the northern Larsen
Ice Shelf monitored by space images. Annals of Glaciology, 17, 317–321.
Skvarca, P., and De Angelis, H. (2003) Impact assessment
of regional climatic warming on glaciers and ice shelves
of the northeastern Antarctic Peninsula. In: E.W.
Domack, A. Leventer, A. Burnett, P. Convey, M.
Kirby, and R. Bindschadler (Eds.), Antarctic Peninsula
Climate Variability: A Historical and Paleoenvironmental Perspective (Antarctic Research Series Vol.
79). American Geophysical Union, Washington,
D.C., pp. 69–78.
Skvarca, P., Rott, H., and Nagler, T. (1995) Satellite
imagery, a baseline for glacier variation study on James
Ross Island, Antarctica. Annals of Glaciology, 21, 291–
296.
Skvarca, P., Rack, W., Rott, H., and Ibarzábal y
Donángelo, T. (1999) Climatic trend, retreat and disintegration of ice shelves on the Antarctic Peninsula:
An overview. Polar Research, 18(2), 151–157.
Skvarca, P., De Angelis, H., and Ermolin, E. (2004) Mass
balance of ‘‘Glaciar Bahı́a del Diablo’’, Vega Island,
Antarctic Peninsula. Annals of Glaciology, 39, 209–213.
Smellie, J.L., Hole, M.J., and Nell, P.A.R. (1993) Late
Miocene valley-confined subglacial volcanism in
northern Alexander Island, Antarctic Peninsula.
Bulletin of Volcanology, 55(4), 273–288.
Smellie, J.L., Nelson, A.E., Johnson, J.S., McIntosh,
W.C., Esser, R., Gudmundsson, M.T., Hambrey,
M.J., and van Wyk de Vries, B. (2007) Six Million
Years of Environmental (Glacial–Interglacial) Conditions Preserved in Volcanic Lithofacies of the James
Ross Island Volcanic Group, Northern Antarctic Penin-
sula (USGS Open-File 1047, Extended Abstract 208),
U.S. Geological Survey, Reston, VA and the National
Academies, Washington, D.C.
Thomas, E.R., Marshall, G.J., and McConnell, J.R.
(2008) A doubling in snow accumulation in the western
Antarctic Peninsula since 1850. Geophysical Research
Letters, 35, L01706, doi: 10.1029/2007GL032529.
Thompson L.G., Peel, D.A., Mosley-Thompson, E.,
Mulvaney, R., Dai, J., Lin, P.N., Davis, M.E., and
Raymond, C.F. (1994) Climate since ad 1510 on Dyer
Plateau, Antarctic Peninsula: Evidence for recent
climate change. Annals of Glaciology, 20, 420–426.
Torinesi, O., Fily, M., and Genthon, C. (2003) Interannual variability and trend of the Antarctic Ice Sheet
summer melting period from 20 years of spaceborne
microwave data. Journal of Climate, 16, 1047–1060.
Turner, J., Lachlan-Cope, T.A., Thomas, J.P., and
Colwell, S.R. (1995) Synoptic origins of precipitation
over the Antarctic Peninsula. Antarctic Science, 7(3),
327–337.
Turner, J., Colwell, S.R., and Harangozo, S. (1997)
Variability of precipitation over the coastal western
Antarctic Peninsula from synoptic observations. Journal of Geophysical Research, 102(D12), 13999–14007.
Turner, J., Leonard, S., Lachlan-Cope, T.A., and
Marshall, G.J. (1998) Understanding Antarctic Peninsula precipitation distribution and variability using a
numerical weather prediction model. Annals of Glaciology, 27, 591–596.
Turner, J., Colwell, S.R., Marshall, G.J., Lachlan-Cope,
T.A., Carleton, A.M., Jones, P.D., Lagun, V., Reid,
P.A., and Iagovkina, S. (2005) Antarctic climate
change during the last 50 years. International Journal
of Climatology, 25, 279–294.
Van den Broeke, M. (2005) Strong surface melting
preceded collapse of Antarctic Peninsula ice shelf. Geophysical Research Letters, 32, L12815, doi: 10.1029/
2005GL023247.
Van der Veen, C.J. (1998) Fracture mechanics approach
to penetration of bottom crevasses on glaciers. Cold
Regions Science and Technology, 27, 213–223.
Van der Veen, C.J. (2002) Calving glaciers. Progress in
Physical Geography, 26, 96–122.
Vaughan, D.G. (2006) Recent trends in melting conditions on the Antarctic Peninsula and their implications for ice-sheet mass balance and sea level. Arctic,
Antarctic, and Alpine Research, 38(1), 147–152.
Vaughan, D.G., Mantripp, D.R., Sievers, J., and Doake,
C.S.M. (1993) Synthesis of remote sensing data on
Wilkins Ice Shelf, Antarctica. Annals of Glaciology,
17, 211–218.
Vaughan, D.G., and Doake, C.S.M. (1996) Recent atmospheric warming and retreat of ice shelves on the
Antarctic Peninsula. Nature, 379(6563), 328–331.
References 741
Vaughan, D.G., Marshall, G.J., Connolley, W.M.,
Parkinson, C., Mulvaney, R., Hodgson, D.A., King,
J.C., Puldsey, C.J., and Turner, J. (2003) Recent
rapid regional climate warming on the Antarctic
Peninsula. Climatic Change, 60, 243–274.
Weertman, J. (1973) Can a Water-filled Crevasse Reach
the Bottom Surface of a Glacier? (IAHS Publication
No. 95), International Association of Hydrological
Sciences, Rennes, France, pp. 139–145.
Wrobel, B.P., Walter, H., Friehl, M., Hoppe, U.,
Schlüter, M., and Steineck, D. (2000) A topographical
data set of the glacier region at San Martin, Marguerite
Bay, Antarctic Peninsula, generated by digital photogrammetry. Polarforschung, 67(1/2), 53–63.