The Dynamics and Mass
Budget of Arctic Glaciers
Extended abstracts
Workshop and GLACIODYN (IPY) meeting,
29 - 31 January 2008, Obergurgl (Austria)
IASC Working group on Arctic Glaciology

Institute for Marine and Atmospheric Research Utrecht
Utrecht University, The Netherlands

The Dynamics and Mass Budgetof Arctic Glaciers
Extended abstracts
Workshop and GLACIODYN (IPY) Meeting,
29 - 31 January 2008, Obergurgl (Austria)
IASC Working Group on Arctic Glaciology
Organized by J. Oerlemans and C.H. Tijm-Reijmer

Institute for Marine and Atmospheric Research Utrecht
Utrecht University, The Netherlands
1ISBN: 978-90-393-4854-3
2

CONTENTS
Preface...............................................................................................7
J. Oerlemans

Program .............................................................................................9
List of participants..........................................................................13
Abstracts .........................................................................................16
Front types of Svalbard tidewater glaciers ...................................17
M. Błaszczyk and J.A. Jania

Relations between atmospheric circulation and glaciers’ massbalance elements on Southern Svalbard (Hansbreen as an
example).......................................................................................20
T. Budzik, M. Grabiec, D. Puczko and K. Migała

Monitoring the Mass Balance of Ice Caps in the Canadian Arctic24
D. Burgess, R. Koerner, M. Demuth, L. Gray, N. Short and G. Cogley

Terrerstrial photogrammetry for measurements of the calving rateof Kronebreen, Svalbard ..............................................................27
A. Chapuis, C. Rolstad, and I. Maalen-Johansen

Surface mass-balance modeling of Nordenskiöldbreen, Svalbard.....................................................................................................28
M. den Ouden, C.H. Reijmer and J. Oerlemans

Iceberg calving flux and mass balance of the Austfonna ice cap on
Nordaustlandet, Svalbard .............................................................32
J.A. Dowdeswell, T.J. Benham, T. Strozzi and J.O. Hagen

Application of a numerical ice sheet model on Austfonna, Svalbard- model preparations and input data.............................................33
T. Dunse, T.V. Schuler and J.O. Hagen

Greenland surface mass balance using a regional atmosphericclimate model ...............................................................................36
J. Ettema, M. van den Broeke, E. van Meijgaard, J. Bamber and R.
Gladstone

Changes of geometry of Renardbreen (Spitsbergen) and evolutionof morphology of its forefield during the last century ....................40
G. Gajek, M. Grabiec, D. Puczko and J. Jania

Some results of radio-echo soundings of frontal parts andforefields of selected glaciers in Svalbard and N Scandinavia .....46
M. Grabiec, W. Dobiński, D. Puczko, P. Dolnicki

Glacier geometry change in the forlandsundet area (NW
Spitsbergen) using remote sensing data ......................................50
M. Grześ, M. Król and I. Sobota
3

Recent changes in surface mass balance of the Austfonna ice
cap, Svalbard ...............................................................................53
J.O. Hagen, T. Eiken, T.V. Schuler, T. Dunse, K. Melvold, G. Moholdt, A.
Taurisano, O. Brandt and J. Kohler

The role of basal fractures in storing and routing water ...............57
J. Harper, J. Bradford, N. Humphrey, T. Pfeffer and T. Meierbachtol

Meltwater runoff and firn densification in the percolation facies,
Greenland Ice Sheet ....................................................................60
J. Harper, N. Humphrey, T. Pfeffer, J. Brown, D. Schuler and D. Sturgis

Glaciodyn activities on and under Engabreen ..............................62
M. Jackson and T. Haug

Natural radionuclides in glacier meltwater as useful tracers ........65
A. Kies, A. Nawrot, H. Surbeck and Z. Tosheva

MODIS albedo and regional mass balance of Svalbard glaciers .69
J. Kohler

Calving dynamics of Jakobshavn Isbrae ......................................70
Martin Lühti

Mass balance of the Prince of Wales Ice-field, Ellesmere Island,
Nunavut, Canada .........................................................................71
D. Mair, D. Burgess, M. Sharp, S. Marshall, F. Cawkwell, J. Dowdeswell and
T. Benham

Ablation and vertical gradients of air temperature – a study from
Hans Glacier, SW Spitsbergen.....................................................74
K. Migała, B. Luks, D. Puczko, S. Sikora, A. Drzeniecka–Osiadacz, M.
Grabiec and P. Głowacki

Internal Structure of Ariebreen, Spitsbergen, from radio-echosounding data ...............................................................................78
F. Navarro, M. Grabiec, D. Puczko, U. Jonsell and A. Nawrot

Polar Panoramas: Images worth a thousand megabytes.............82
M. Nolan

Rapid surface elevation losses of Spitsbergen glaciers (19361990-2005) ...................................................................................86
C. Nuth, G. Moholdt, J. Kohler, J.O. Hagen and H. Faste-Aas

Ice volume changes of Ariebreen, Spitsbergen, during 1936-19902007 .............................................................................................89
M. Petlicki, J. Lapazaran, F. Navarro, P. Glowacki and F. Machío

Altimetry of Nordenskiöldbreen and Vestfonna 1991 - 2007 Preliminary Results ......................................................................93
V. Pohjola, J. Hedfors, J. Moore, B. Sjögren, L. Kolondra, C. Reijmer, E.
Isaksson, and J.O. Hagen

4

Mass balance and velocity observations on Nordenskiöldbreen,
Svalbard .......................................................................................97
C. Reijmer, M. den Ouden, V. Pohjola, B. Sjögren and J. Oerlemans

Spatio-temporal variability of snow melt on Svalbard derived fromspaceborne scatterometer data (Qscat) .....................................101
G. Rotschky, J. Kohler, E. Isaksson

IPY Outreach, Education, and Polar Science Days ...................103
R. Salmon

A Five-Year Record of Summer Melt on Eurasian Arctic Ice Caps...................................................................................................104
M. Sharp and L. Wang

Regional distribution of snow accumulation on North-Western
Spitsbergen glaciers, Svalbard...................................................105
I. Sobota and M Grześ

Identifying causes of peripheral thinning of the Greenland Ice
Sheet ..........................................................................................109
A. Sole, T. Payne, J. Bamber, P. Nienow and W. Krabill

Relations between horizontal strain and ice elevation change...113
M. Stober

Recent surge activity in southern Spitsbergen ...........................117
1

M. Sund , T. Eiken, J.O. Hagen, A. Kääb, T.V. Schuler, K. Müller, G.
Moholdt and C. Nuth

Glaciological and meteorological observations in Suntar-Khayata
Range, Eastern Siberia, in 2004-2007 .......................................122
S. Takahashi, K. Sugiura, T. Kameda, H. Enomoto, Y. Kononov and M.D.
Ananicheva

Preliminary results of GIS-based solar radiation model forHornsund area, SW Spitsbergen................................................126
M. Szymanowski, M. Kryza, K. Migała, P. Sobolewski and L. Kolondra

Observation and modelling of the surface energy balance in theablation zone of the West Greenland ice sheet..........................129
M. van den Broeke, P. Smeets, J. Ettema, R. van de Wal, J. Oerlemans

Large and rapid variability in the velocity in the ablation zone of theGreenland ice sheet ...................................................................133
R.S.W. v.d. Wal, W. Boot, M.R. van den Broeke, P. Smeets, C.H. Reijmer,
J.J.A. Donker and J. Oerlemans

5

6

PREFACEThe 2008 annual workshop of the Working Group on Arctic Glaciology was
combined again with a GLACIODYN workshop. It was held in Obergurgl, Austria,
29-31 January 2008.
GLACIODYN is a lead IPY project dealing with Arctic glaciers. It comprises aninternationally-coordinated effort to study the dynamics of Arctic glaciers and todevelop better tools to predict the response of these glaciers to climate warming.
The key elements of this effort are (i) to make better use of observational
techniques to assess the detailed dynamics of a key set of glaciers, and (ii) to
develop models that can be used to aggregate data and that are sufficiently robustto have predictive power. A set of target glaciers have been identified for intensive
observations (in situ and from space) for the period 2007-2010. This set covers awide range of climatic/geographical settings and takes maximum advantage ofprior long-term studies.
Over 58 participants presented 36 oral contributions and 16 posters on a wide
range of subjects. Previous reports of the workshop appear to be very popular, so it
was decided to make a book of extended abstracts once more. A considerable
number of the workshop presentations are summarized here, and I want to thankthe contributors for their efforts to produce clear and well-illustrated texts.
IASC generously provided financial means for 7 young scientists to participate in
the workshop and to cover the costs of the congress facilities. Further support wasreceived from IMAU (Utrecht University) and the Netherlands IPY Programme(NWO), covering the printing of this book.
Hopefully, the long series of workshops we had until now will be continued in the
future. The next workshop is scheduled for February 2009, and will be hosted by
the Canadian GLACIODYN partners. General information about past and future
meetings can be found at:
http://www.phys.uu.nl/%7Ewwwimau/research/ice_climate/iasc_wag/
Once more, Carleen Tijm-Reijmer has been the key person in the organization of
the workshop and the production of the abstract book. Carleen, thank you verymuch for all the work you have done to make these workshops a success!

Johannes Oerlemans
Coordinator of GLACIODYN

78PROGRAM
Tuesday 29
Chair: Johannes Oerlemans• 09:00 – 09:10 Johannes Oerlemans: Welcome• 09:10 – 09.30 Douglas Mair: The mass balance of the Prince of Wales
Icefield, Ellesmere Island, Canada• 09.30 – 09:50 David Burgess: Monitoring the mass balance of ice caps andglaciers in the Canadian High Arctic• 09.50 – 10:10 Joel Harper, John Bradford, Neil Humphrey and Tad Pfeffer:
The role of englacial fractures in storing and routing water• 10:10 – 10:30 J. Jania, L. Kolondra, G.K.C. Clarke, D. Puczko, T. Budzik and
E. Majchrowska: Subglacial water outflows and related elements of dynamics ofHansbreen – tidewater glacier in south Svalbard
10:30 – 10:50

Coffee break• 10:50 – 11:10 F. Navarro, M. Grabiec, D. Puczko, U. Jonsell, A. Nawrot:
Internal structure of Ariebreen, Spitsbergen, from radio-echo sounding data.
• 11:10 – 11:30 Bartek Luks, K. Migała, D. Puczko and M. Grabiec: Local
climate and vertical gradients of glaciological processes• 11:30 – 11:50 K. Migała, M. Szymanowski and M. Kryza: Spatial distribution ofpotential and total radiation on the glaciers in the Hornsund area, SW
Spitsbergen• 11:50 – 12:10 Jack Kohler: MODIS albedo and regional mass balance of
Svalbard glaciers.
12:10 – 14:00

LUNCH

Chair: Jason Box• 14:00 – 14:20 Christopher Nuth: On the fast-flowing GLACIODYN target
glacier Kronebreen/Holtedahlfonna• 14:20 – 14:40 Monica Sund: Recent surge activity in southern Spitsbergen• 14:40 – 15:00 C. Reijmer, M. den Ouden, V. Pohjola, J. Oerlemans: Massbalance and velocity observations on Nordenskiöldbreen, Svalbard• 15:00 – 15:20 Veijo Pohjola: Altimetry of Nordenskiöldbreen - changes
between 1991 and 2007
15:20 – 15:50

Coffee break• 15:50 – 16:10 M. Petlicki, J. Lapazaran, F. Navarro, P. Glowacki, F. Machío:
Ice volume changes of Ariebreen, Spitsbergen, during 1936-1990-2007.
• 16:10 – 16:30 Gerit Rotschky, J. Kohler, E. Isaksson: Spatio-temporalvariability of snow melt on Svalbard derived from spaceborne scatterometerdata (Qscat)• 16:30 – 16:50 Matt Nolan: Polar panoramas: Images worth a thousand
megabytes• 16:50 – 17:10 Rhian Salmon: IPY outreach, education, and polar science days
• 17:10 – 18:00 POSTER SESSION
19:00

DINNER
9

Wednesday 30
Chair: Regine Hock• 09:00 – 09:20 M.A.G. den Ouden, C.H. Reijmer and J. Oerlemans: Surface
mass-balance modeling of Nordenskiöldbreen, Svalbard• 09:20 – 09:40 J.O. Hagen: Mass balance of Austfonna Ice cap• 09:40 – 10.00 Thomas Schuler: The accumulation pattern of Austfonna,
Svalbard• 10.00 – 10:20 J.A. Dowdeswell, T.J. Benham, T. Strozzi, J.O. Hagen: Iceberg
calving flux and mass balance of the Austfonna ice cap on Nordaustlandet,
Svalbard
10:20 – 10:50

Coffee break• 10:50 – 11:10 Martin Sharp and Libo Wang: A five-year record of summer
melt on Eurasian Arctic ice caps
• 11:10 – 11:30 Shuhei Takahashi, Konosuke Sugiura, Takao Kameda, Hiroyuki
Enomoto, Yury Kononov and Maria D. Ananicheva: Glaciological andmeteorological observations in Suntar-Khayata Range, Eastern Siberia, in
2004-2007• 11:30 – 11:50 Andrey Glazovskiy and Yury Macheret: RES studies of glaciers
on Franz-Josef Land
• 11:50 – 12:10 Jason Box: Greenland outlet glacier dynamics from repeat
photography
12:10

LUNCH

19:00

DINNER• 20:30

Meeting of national delegates of the Working Group on
Arctic Glaciology of IASC

Thursday 31
Chair: Martin Sharp• 08:30 – 08:50 Michiel van den Broeke, Paul Smeets, Janneke
Ettema, Roderik van de Wal: Observed and calculated melt in the ablation zoneof the West Greenland ice sheet• 08:50 – 09:10 Janneke Ettema, Michiel van den Broeke and Erik van
Meijgaard: Greenland surface mass balance using a regional atmosphericclimate model• 09:10 – 09.30 Jay Zwally: Mass balance of the Greenland Ice Sheet: A recent
increase in rate of ice loss Observed by ICESat and GRACE• 09.30 – 09:50 Andrew Sole: Investigating causes of peripheral thinning of theGreenland Ice Sheet: Constraint on dynamic thinning?
• 09:50 – 10:10 R.S.W. v.d. Wal, W. Boot, M.R. van den Broeke, P. Smeets,
C.H. Reijmer, J.J.A. Donker and J. Oerlemans: Large and rapid variability in thevelocity in the ablation zone of the Greenland ice sheet• 10:10 – 10:30 Manfred Stober: Relations between horizontal strain andelevation change at the Swiss Camp area (West Greenland)
10:30 – 11:00

Coffee break
10Chair: Jon Ove Hagen• 11:00 – 11:20 Martin Lüthi: Calving dynamics of Jakobshavn Isbrae• 11:20 – 11:40 Tad Pfeffer: Limits on Greenland dynamic discharge during thenext century• 11:40 – 12:00 Regine Hock: Parameterization of longwave incoming radiation
on Storglaciaren, Sweden• 12:00 – 12:20 Miriam Jackson: Glaciodyn activities in Svartisen subglaciallaboratory• 12:20 – 12:40 Antoine Kies, A. Nawrot, H. Surbeck, Z. Tosheva, J. Jania:
Natural Radionuclides in Glacier meltwater as useful tracers• 12:40 – 13:00 M. Blaszczyk and J. Jania: Front types of Svalbard tidewater
glaciers
13:00

LUNCH

19:00

DINNER

11

POSTERS• T. Budzik, M. Grabiec, T. Niedzwiedz, D. Puczko and K. Migala: Relationships
between atmospheric circulation and glaciers´ mass balance elements on
Southern Svalbard• T. Budzik and D. Ignatiuk: Short-term variations of the ablation and its sensivity
to global radiation recorded on Hans Glacier, SW Spitsberegn, Svalbard• Anne Chapuis, Cecilie Rolstad, and Ivar Maalen-Johansen: Terrerstrial
photogrammetry for measurements of the calving rate of Kronebreen, Svalbard• Thorben Dunse, Thomas V. Schuler and Jon Ove Hagen: Application of anumerical ice sheet model on Austfonna, Svalbard - model preparations andinput data
• G. Gajek, M. Grabiec, D. Puczko and J. Jania: Changes of geometry ofRenardbreen (Spitsbergen) and evolution of morphology of its forefield during
the last century• M. Grabiec, D. Puczko and W. Dobinski: Some results of radio echo soundingsof frontal parts and forefields of selected subpolar glaciers in Svalbard and N
Scandinavia• M. Grześ, M. Król and I. Sobota: Glacier geometry chang in the Forlandsundet
area using remote sensing data
• Joel Harper, Neil Humphrey and Tad Pfeffer: Meltwater runoff and firndensification in the Percolation Facies, Greenland Ice Sheet• Antoine Kies, A. Navrot, H. Surbeck, Z. Tosheva, J. Jania: Natural
Radionuclides in Glacier meltwater as useful tracers• L. Kolondra, M. Blaszczyk, J. Jania and P. Glowacki: New satellite map of theKinnvika Station surroundings, NW Nordaustandet• Geir Moholdt and Christopher Nuth: On the use of ICEsat measurements on
Svalbard glaciers.
• Faezeh Nick: Modelling the behaviour of Greenland outlet glaciers• Christopher Nuth, G. Moholdt, J. Kohler, H.F. Aas and J.O. Hagen: Rapid
surface elevation losses of Spitsbergen glaciers [1936-1990-2005]• W. Schöner, A. Richter, A.I. Sharov and Bernhard Hynek: The Austrian
research focus within IPY in Greenland and Franz Josef Land
• I. Sobota and M. Grześ: Regional distribution of snow accumulation on northwestern Spitsbergen glaciers, Svalbard• Dirk van As: A long-term network of automatic weather and ice monitoring
stations in the ablation zone of the Greenland Ice Sheet

12

LIST OF PARTICIPANTS
1.
2.
3.
4.
5.
6.
7.
8.
9.
10.
11.
12.
13.
14.
15.
16.
17.
18.
19.
20.
21.

Andreas Ahlstrøm (apa@geus.dk)
Geological Survey of Denmark and Greenland, Copenhagen, Denmark
Kumiko Azuma (kumiko@pmg.nipr.ac.jp)
National Institute of Polar Research, Tokyo, Japan
Malgorzata Blaszczyk (mbl4@wp.pl)
Faculty of Earth Sciences, University of Silesia, Poland
Jason Box (box.11@osu.edu)
Byrd Polar Research Center, Ohio State University, Columbus, Ohio, USA
Tomek Budzik (tomasz.budzik@us.edu.pl)
Faculty of Earth Sciences, University of Silesia, Poland
David Burgess (David.Burgess@NRCan.gc.ca)
Natural Resources Canada, Otawa, Canada
Anne Chapius (anne.chapuis@umb.no)
Department of Mathematical Sciences and Technology, Norwegian Universityof Life Sciences, Norway
Marianne den Ouden (M.A.G.denOuden@phys.uu.nl)
Institute for Marine and Atmospheric Research, Utrecht University,
Netherlands
Evelyn Dowdeswell (e.k.dowdeswell@bris.ac.uk)
Bristol Glaciological Centre, Bristol, UK
Scott Polar Research Institute, University of Cambridge, Cambridge, UK
Julian Dowdeswell (jd16@cam.ac.uk)
Scott Polar Research Institute, University of Cambridge, Cambridge, UK
Thorben Dunse (thorben.dunse@geo.uio.no)
Department of Geosciences, University of Oslo, Oslo, Norway
Janneke Ettema (j.ettema@phys.uu.nl)
Institute for Marine and Atmospheric Research, Utrecht University,
Netherlands
Grzesiek Gajek (gajcy@o2.pl)
M. Curie-Sklodowska University, Poland
Andrey Glazovsky (icemass@yandex.ru)
Institute of Geography, Russian Academy of Sciences, Moscow, Russia
Piotr Glowacki (glowacki@igf.edu.pl)
Institute of Geophysics, Polish Academy of Science, Warszawa, Poland
Mariusz Grabiec (mariusz.grabiec@us.edu.pl)
Faculty of Earth Sciences, University of Silesia, Poland
Jon Ove Hagen (j.o.m.hagen@geo.uio.no)
Department of Geosciences, University of Oslo, Oslo, Norway
Joel Harper (joel@mso.umt.edu)
Dept. of Geosciences, Univ. of Montana, Missoula, USA
Regine Hock (regine.hock@gi.alaska.edu)
Geophysical Institute, University of Alaska, Fairbanks, Alaska
Bernhard Hynek (bernhard.hynek@zamg.ac.at)
Dep. of Climatology, Central Institute for Meteorology and Geodynamics,
Vienna, Austria
Miriam Jackson (mja@nve.no)13

Norwegian Water Resources and Energy Directorate, Oslo, Norway
22. Jacek Jania (jjania@us.edu.pl)
Faculty of Earth Sciences, University of Silesia, Poland
23. Antoine Kies (antoine.kies@uni.lu)
University of Luxembourg, Luxembourg
24. Jack Kohler (jack.kohler@npolar.no)
Norwegian Polar Institute, Tromsø, Norway
25. Michael Kuhn (Michael.Kuhn@uibk.ac.at)
Institute of Meteorology and Geophysics, University of Innsbruck, Austria
26. Martin Lühti (luethi@vaw.baug.ethz.ch)
Laboratory of Hydraulics, Hydrology and Glaciology, Swiss Federal Institute ofTechnology (ETH), Zurich, Switzerland
27. Bartek Luks
Institute of Geophysics, Polish Academy of Science, Warszawa, Poland
28. Yuri Macheret (macheret@gol.ru)
Institute of Geography, Russian Academy of Sciences, Moscow, Russia
29. Douglas Mair (d.mair@abdn.ac.uk)
School of Geosciences, University of Aberdeen, Aberdeen, UK
30. Krzystof Migala (migalak@meteo.uni.wroc.pl)
University of Wroclaw, Institute of Geography & Regional Development,
Poland31. Heinz Miller (hmiller@AWI-Bremerhaven.de)
Alfred Wegener Institute for polar and marine research, Bremerhaven,
Germany
32. Francisco Navarro (fnv@mat.upm.es)
Madrid Polytechnical University, Spain
33. Faezeh Nick (Faezeh.Nick@durham.ac.uk)
Department of Geography, University of Durham, Durham, UK
34. Matt Nolan (fnman@uaf.edu)
Institute of Northern Engineering, University of Alaska, Fairbanks, Alaska
35. Christopher Nuth (christopher.nuth@geo.uio.no)
Department of Geosciences, University of Oslo, Oslo, Norway
36. Friedrich Obleitner (Friedrich.Obleitner@uibk.ac.at)
Institute of Meteorology and Geophysics, University of Innsbruck, Austria
37. Hans Oerlemans (J.Oerlemans@phys.uu.nl)
Institute for Marine and Atmospheric Research, Utrecht University,
Netherlands
38. Michal Petlicki (petlicki@igf.edu.pl)
Institute of Geophysics, Polish Academy of Sciences, Poland
39. Tad Pfeffer (pfeffer@tintin.colorado.edu)
University of Colorado, Boulder, USA
40. Veijo Pohjola (veijo.pohjola@geo.uu.se)
Department of Earth Sciences, Uppsala University, Uppsala, Swede
41. Dariusz Puczko (dp@igf.edu.pl)
Institute of Geophysics, Polish Academy of Science, Warszawa, Poland
42. Gerit Rotschky (Gerit.Rotschky@npolar.no)
Norwegian Polar Institute, Tromsø, Norway
43. Rhian Salmon (ipy.ras@gmail.com)
IPY Education and Outreach Coordinator1444. Elisabeth Schlosser (Elisabeth.Schlosser@uibk.ac.at)
Institute of Meteorology and Geophysics, University of Innsbruck, Austria
45. Thomas Schuler (t.v.schuler@geo.uio.no)
Department of Geosciences, University of Oslo, Oslo, Norway
46. Martin Sharp (martin.sharp@ualberta.ca)
Earth and Atmospheric Sciences, University of Alberta, Edmonton, Canada
47. Paul Smeets (c.j.p.p.smeets@phys.uu.nl)
Institute for Marine and Atmospheric Research, Utrecht University,
Netherlands
48. Irek Sobota (irso@geo.uni.torun.pl)
Dept. of Cryology and Polar Research, N.Copernicus University, Torun,
Poland49. Andrew Sole (A.J.Sole@bristol.ac.uk)
Bristol Glaciological Centre, Bristol, UK
50. Manfred Stober (manfred.stober@hft-stuttgart.de)
Stuttgart University of Applied Sciences, Stuttgart, Germany
51. Monica Sund (Monica.Sund@unis.no)
University Centre in Svalbard, Longyearbyen, Norway
52. Shuhei Takahashi (shuhei@mail.kitami-it.ac.jp)
Kitami Institute of Technology, Japan
53. Carleen Tijm-Reijmer (C.H.Tijm-Reijmer@phys.uu.nl)
Institute for Marine and Atmospheric Research, Utrecht University,
Netherlands
54. Zornitza Tosheva
University of Luxembourg, Luxembourg
55. Dirk van As (dva@geus.dk)
Geological Survey of Denmark and Greenland, Copenhagen, Denmark
56. Roderik van de Wal (R.S.W.vandeWal@phys.uu.nl)
Institute for Marine and Atmospheric Research, Utrecht University,
Netherlands
57. Michiel van den Broeke (m.r.vandenbroeke@phys.uu.nl)
Institute for Marine and Atmospheric Research, Utrecht University,
Netherlands
58. Jay Zwally Zwally (zwally@icesat2.gsfc.nasa.gov)
NASA Goddard SFC, USA

15

ABSTRACTS

16

FRONT TYPES OF SVALBARD TIDEWATER GLACIERS
MAŁGORZATA BŁASZCZYK AND JACEK A. JANIA
Faculty of Earth Sciences, University of Silesia, Poland

We have classified all tidewater glaciers in Svalbard with respect to their dynamics,
taking into account patterns of crevasses examined on satellite images and aerial
photos. They were used as a proxy for the determination of glacier movement
velocity.
Information on glacier velocity close to its terminus is crucial for calculation of thecalving flux. The amount of direct data on the Svalbard tidewater glaciers motion
velocity is very limited. Moreover, glacier flow velocity varies on different timescales (diurnal, interdiurnal, seasonal and interannual). Therefore, some of thedata (e.g. InSAR) were not representative for an annual ice flux. Due to lack ofrepeated ground survey data close to glaciers’ terminus, a feature tracking
technique using the sequential high-resolution ASTER imagery from 2000 – 2006
was applied to determine the surface velocities near several glaciers fronts. The
annual surface velocity of 10 glaciers was derived by measuring displacements ofsurface features (crevasses, moraine elements) visible on images acquired in oneyear period. An error of feature tracking velocity measurements is estimated as ±30-1ma . For 16 glaciers very small crevassing areas were noted and no motion of theglaciers’ front was observed on ASTER images acquired in one or even two year
interval (within the above accuracy).

I
II
III
IV

Table 1. Categories of front types of Svalbard tidewater glaciers
Types of tidewater glaciers fronts
Length of crevassing
Glacier velocity-1zone [m][ma ]very slow or stagnant glaciers
0 - 300
10 ± 5slow glaciers≥ 300 - 1000
70 ± 30
fast flowing glaciers≥1000
200 ± 50
surging glaciers or fast flowing ice streams>>1000700± ?

Evaluation of the dynamic status of tidewater glaciers was based on several
characteristics such as: length of glacier, length of crevassed zone, area of glacier,area of crevassed zone and length of cliff acquired from ASTER imagery. Allavailable data (from research and literature) on glaciers’ velocity were analyzed incomparison with the above features. The best correlation (r = 0.71) was foundbetween the length of the crevassing zone and the glacier velocity. This parameter
was determined as the main criterion for estimation of tidewater glaciers velocity.
Four categories of front types of Svalbard tidewater glaciers were distinguished:I.
very slow or stagnant glaciers (Figure 1a)II.
slow glaciers (Figure 1b)
III.
fast flowing glaciers (Figure 1c)IV.
surging glaciers or fast flowing ice streams (Figure 1d).
17

Length of crevassing zone and an average glacier velocity near the front assumed
for each of the group is presented in Table 1.

Figure 1. Examples of front types of Svalbard tidewater glaciers: a) Recherchebreen - very
slow or stagnant glacier; b) Aavatsmarkbreen - slow glacier; c) Osbornebreen - fast flowing
glacier; d) Perseibreen - surging glacier.

Complete length of tidewater glaciers’ cliffs was measured using geocoded ASTER
images and used as data for the calving intensity assessment. The total length is859.1 km for the whole Svalbard, while Dowdeswell (1989) reported 1028 km. It
shows a reduction of tidewater glacier seaward margin by 16% in c. 40 years. The
length of ice cliffs of the main islands of the archipelago is shown in Table 2.
One can calculate calving ice flux from all Svalbard glaciers when having at one’s
disposal: length of the ice cliffs, average glaciers’ velocities and their ice thickness
near terminus. Due to very sparse data on sea water depth close to ice cliffs, anaverage value of 100 m for ice thickness was used in these calculations.
ASTER images also provide a good overview of glacier fronts’ fluctuations. Front
position changes of about 30 glaciers were measured and, as a result, an average18

-1

retreat rate by 20 – 40 ma was suggested. After having compiled the available
information, the total calving intensity of Svalbard glaciers (without Kvitøya)
3 -1
3 -1
amounts to 6.7 – 8.2 km a ± 1.5 km a .
Table 2. Length of ice cliffs of the main islands of Svalbard: L1 – from Dowdeswell (1989);
L2 – data derived from ASTER satellite images (2000-2006).
IslandLength of ice cliffs
L1 [km]L2 [km]Spitsbergen
484
388.4
Nordaustlandet
306
272.3
Edgeøya79
68.7
Prins Karls Forland
17
8.9
Barentsøya
23
8.8
Storøya13
12Kvitøya106
100Sum1028 km
859.1 km

Examination of ASTER images from 2000-2006 for the whole Svalbard has shownthat 163 glaciers appear as the tidewater type. In comparison to the Hagen et al.
(1993) inventory, 14 glaciers have retreated from the sea to land during the periodof c. 30-40 years. Basing upon the literature, 54 tidewater glaciers have been
considered as surge-type, which makes up 33% of all tidewater glaciers. Taking
into account other evidence visible on ASTER images (e.g. badly crevassed glaciersurface, distinct lateral shear zones, folded medial moraines, etc.), the number oftidewater glaciers classified as a surge-type may reach even 43%. It is much morethan 13% of the whole population of Svalbard glaciers suggested by Jiskoot et al.
(1998) and closer to preliminary estimates proposed by Jania (1988).

Acknowledgements.
Studies presented in the abstract were financed by the Ministry of Science andHigher Education, Republic of Poland under terms of special research grant No.
IPY-269/2006 (PL-GLACIODYN).

References:
Dowdeswell J.A., 1989. On the nature of Svalbard icebergs. Journal of Glaciology, 35(120),
224 - 234.
Hagen J.O., Liestøl O., Roland E., Jørgensen T., 1993. Glacier Atlas of Svalbard and Jan
Mayen. Norsk Polarinstitutt, Meddelelser, 129, Oslo, 141 pp.
Jania J., 1988. Dynamiczne procesy glacjalne na poludniowym Spitsbergenie w swietle
badan fotointerpretacyjnych i fotogrametrycznych [Dynamic glacier processes insouthern Spitsbergen in the light of photointerpreration and photogrammetric research English summary]. Wydawnictwo Uniwersytetu Slaskiego, Katowice 1988, 258 pp.
Jiskoot H., Boyle P., Murray T., 1998. The incidence of glacier surging in Svalbard: Evidence
from multivariate statistics. Computers & Geosciences, 24(4), 387 - 399.

19

RELATIONS BETWEEN ATMOSPHERIC CIRCULATION AND
GLACIERS’ MASS BALANCE ELEMENTS ON SOUTHERN SVALBARD(HANSBREEN AS AN EXAMPLE)
T. BUDZIK1, M. GRABIEC2, D. PUCZKO3 AND K. MIGAŁA4
1

University of Silesia, Faculty of Earth Sciences, Department of Climatology, Poland
University of Silesia, Faculty of Earth Sciences, Department of Geomorphology, Poland
3
Institute of Geophysics Polish Academy of Sciences, Department of Polar and Marine
Research, Poland
4
University of Wroclaw, Institute of Geography & Regional Development, Department ofMeteorology & Climatology, Poland
2

Objectives
The atmospheric circulation is one of the most significant elements which shapeboth the weather and the climate. It plays an important role in the heat andhumidity transfer within the Arctic, which is due to geographical location anddifferentiated light conditions in a yearly course (the polar day and night). As aconsequence, circulation conditions shall have a significant influence uponelements of glaciers' mass balance. The aim of this work is to indicate relations
between the mass balance of Southern Spitsbergen glaciers and the atmosphericcirculation with Hansbreen as the example.

Figure 1. Location of Hansbreen

20Methods
Relations between the discussed factors occurring in one summer season as wellas in the years 1989-2007 have been analyzed. Types of atmospheric circulation
have been drawn up on the basis of the calendar of circulation types for
Spitsbergen for the years 1951 - 2007 (Niedzwiedz 1997, 2006). The synoptic
conditions have been divided into 2 groups of anticyclonic and cyclonic dependingon the direction of inflow of air masses over Svalbard: N, NE, E, SE, S, SW, W,
NW. Additionally, two conditions of stagnation have been distinguished: Ca - thecentre of high pressure over Svalbard and Ka - a wedge of high pressure, and two
situations of variable advection. These situations are: Cc - in case of emergence ofthe low pressure centre over Svalbard and Bc - the emergence of cyclonic furrow.
Also, a type of circulation marked as “x” has been distinguished separately for baric
col and situations when the weather situation cannot be defined explicitly.
On the basis of circulation types westerly zonal index (W), southerly meridional
index (S) and cyclonicity index (C) have been defined for accumulation andablation seasons separately. The S index express intensity of latitudinal circulation.
Positive values of the index indicate prevailing inflow of the air from the southern
sector and diminishing one from the northern sector. The W index defines theintensity of western zonal circulation (positive values) and eastern one (negativevalues).
The study of mass balance elements relates to Hansbreen (South Spitsbergen).
2
Hansbreen (an area of 56 km ) is a polythermal glacier terminating in the sea(Figure 1). This glacier's mass balance has been constantly monitored since 1988.
Short-term changes in surface elevation in the summer season of 2007 were
defined by a SR50 Campbell Sonic Ranging Sensor. The device was located at
Hansbreen at the elevation of 210 meters a.s.l. In addition, meteorological datafrom the Polish Polar Station in Hornsund and automatic weather station (AWS) at
Hansbreen were used in the study.

Figure2.
Courseofcumulative change of surface
elevation [m] and averagesurface elevation change [cm]on Hansbreen in summer
2007.

Results
The glacier's surface elevation in summer 2007 was decreasing at a considerable
rate till the 8th of August (Figure 2). The ablation was stopped in the second part ofthe season. These stages are connected with types of circulation dominating in21

those two periods. In the first half of the season the advection with northern andsouthern component (Nc, NEc, E and SEc) as well as anticyclonic non-advective
types dominated. The circulation conditions after the 8th of August can becharacterized with the domination of cyclonic types - chilly and wet ones (Nc andNEc).
The average daily changes in the elevation of glacier's surface can differ inparticular types of circulation (Table 1). The most effective melting is accompanied
by inflows of the air of Ea and SEa type (3.6 cm per day on average) and Ec, SEc
type (3.4 cm per day on average). Humid and chilly air masses of Wc and NWc
type, on the other hand, that generate much rainfall can cause an increase in theelevation of glacier's surface even in summer (1.0 cm per day on average).
Table 1. Statistical characteristics of summer 2007 in circulation types
Circulation type
Na+ Ea+ Sa+ Wa+ Ca+ Nc+ Ec+ Sc+ Wc+
NEa SEa SWa NWa Ka NEc SEc SWc NWc
Frequency [%]
5.7
6.6
0.8
2.5 13.9 29.5 15.6 4.1
5.7
Average teperature
4.4
4.3
3.0
2.4
3.3
2.5
4.3
3.6
2.9
[°C]
∑PDD [°C]31.0 34.5 3.0
7.2 55.8 94.3 82.1 18.2 20.4
Total precipitation
0.0
0.6
0.4
3.1 24.1 15.2 17.3 42.4 26.5[mm]
Average surface-2.9 -3.6 -3.0 -0.3 -2.2 -1.2 -3.4 -1.8
1.0
elevation change [cm]Cumulative change of -20.0 -29.0 -3.0 -1.0 -38.0 -42.0 -64.0 -9.0
7.0
surface elevation [cm]

Cc+
Bc
13.1
2.5

x
2.5
2.5

39.5
25.7

7.6
0.0

-1.1

-1.3

-18.0 -4.0

Extreme ablation event recorded on the 1st of August was caused by advection of
SEc type lasting for several days (Figure 2). A shift of advection to NWc provokeda precipitation and positive changes in the glacier's elevation.
Relations between the glacier's balance and atmospheric circulation in a long-term
perspective were considered for ablation and accumulation season separately(Table 2). Important statistical relations (p = 0.01) were noted between the summer
balance and the frequency of occurrence of circulation of Ca + Ka type (r = -0.61).
It reflects the increase in ablation along with the increasing frequency ofanticyclonic, cloudless and minimum precipitation weather as well as lessened
activity of cyclonic advection from W and NW. The advection of air masses of Wc
and NWc type (r = 0.60) cause ablation decrease. Due to the fact that the most of
frequency of occurrence of those circulation types fall in summer (14 and 6 %percent respectively), they belong to the key weather shaping factors oversouthern Spitsbergen.
Table 2. Correlation coefficients (p=0.01) between Hansbreen mass balance elements andfrequency of circulation types/indexes occurrence in period 1989-2007.
Balance Na+ Ea+ Sa+ Wa+ Ca+ Nc+ Ec+ Sc+ Wc+ Cc+x
SWCNEa SEa SWa NWa Ka NEc SEc SWc NWc Bc
Winter
-0.19 0.04 -0.10 -0.65 -0.20 -0.16 0.26 0.14 0.35 0.15 -0.01 0.45 -0.15 0.33
Summer -0.35 -0.49 0.05 0.04 -0.61 0.26 0.03 0.40 0.60 0.18 0.28 -0.05 0.50 0.57

22

The winter balance, in turn, shows the closest relation with the advection of Wa +NWa (r = -0.65). The increase in frequency of such advection results first of all in
reduction of winter precipitation. The frequency of occurrence of this advection islow (5% on average). Thus it shall be treated only as an indicative element and not
a dominating one.
Within circulation indices, significant statistical relations were observed only among
the summer balance and indexes of C and W (Table 2). Seasons with dominating
eastern advection and anticyclonic conditions can be characterized with higher
surface ablation of glacier.

Conclusions
Relations between elements of Hansbreen mass balance and atmosphericcirculation are complex and do not show very strict dependence.
The anticyclonic wedge (Ka 10.4%) is the most frequent synoptic situation on
Svalbard (Niedzwiedz 2007). The second one is circulation of Ec (9.9%) type and
Ec. The climate of Svalbard is modelled by an intense cyclonic activity during 56%
of days in a year.
One can indicate a set of circulation types which influence the mass balance ofdistinguished summer and winter seasons. Very negative values of the summer
balance are generated by circulation types with accompanying radiation weather,i.e. Ca and Ka and advection from the eastern sector. Ablation is limited in acondition of inflow of cyclonic masses from the northern and western sector. The
winter balance is conditioned mainly by availability of snow supply. The exemplary
type of circulation is the advection of Wa + NWa type, with air mass of continental
features: dry, cold and with low precipitation. Frequent emerging of this air masscauses reduced winter accumulation.
Relations between the dynamics of mass balance and circulation types have acomplex nature. The preliminary results are indicative only. For example: anincrease in C cyclonic indicator with a simultaneous decrease in W indicator value
in the ablation season of 2001 can suggest very negative values of the summer
balance at that period of time.

Acknowledgements
We would like to thank professor Garry Clarke of the University of British Columbia
and professor Tadeusz Niedzwiedz of University of Silesia for making
meteorological data available.

Bibliography
Niedzwiedz T., 1997, Long-term variability of the atmospheric circulation indices overSpitsbergen and its influence on the air temperature. Problemy Klimatologii Polarnej, 7:19-40.
Niedzwiedz T., 2006, The main forms of atmospheric circulation above Spitsbergen(December 1950 – september 2006). Problemy Klimatologii Polarnej, 16: 91-105.

23

MONITORING THE MASS BALANCECANADIAN ARCTICOF

ICE CAPSIN THEDAVID BURGESS1, ROY KOERNER1, MICHEAL DEMUTH1, LAURENCE GRAY1,NAOMI SHORT1 AND GRAHAM COGLEY2
1
2Natural Resources Canada, Ottawa
Trent University, Canada

Mass balance monitoring of ice caps and glaciers in the Canadian high Arctic
commenced in 1959. Measurements were initially obtained from the Meighen ice
cap by Keith Arnold of Energy, Mines, and Resources (now Natural Resources
Canada) and from the White Glacier, Axel Heiberg Island, by Dr. Fritz Muller of
McGill University, Quebec. Mass balance surveys started on the Devon Island ice
cap by Dr. Koerner and S. Ekman of the Arctic Institute of North America in 1961
with a fourth site installed on the South Melville ice cap in 1963. A fifth site on the
Agassiz ice cap was added to the network in 1976 (Please refer to:
http://pathways.geosemantica.net/WSHome.aspx?ws=NGP_SECG%26locale=en-CA
for base-maps of the Arctic monitoring sites and individual site layouts). Morerecently, mass balance monitoring sites have been established in the lower Arctic
regions on the Penny ice cap, Baffin Island (maintained by Christian Zdanowicz[Christian.Zdanowicz@NRCan.gc.ca] – NRCan) and in the Mackenzie Mountains
of Nahanni National Park, Northwest Territories (maintained by Mike Demuth[Mike.Demuth@NRCan.gc.ca]–NRCan). With respect to the high Arctic monitoringsites, Dr. Koerner eventually became responsible for maintenance of all sites
except for the White Glacier, which has been maintained by researchers from Trent
University, Ontario since 1983.
Each of the 4 monitoring sites maintained by NRCan (Devon, Meighen, Melville,and Agassiz) consists of a network of mass balance poles and at least 1 automatic1
weather station . Winter accumulation is estimated from snow depth and density
values taken from each pole location, while summer balance is estimated as thedifference between the previous winter balance and the previous annual firn layer.
Percolation trays are used in the summit region of the Agassiz ice cap in order toestimate internal accumulation due to refreezing of meltwater within the nearsurface snowpack or firn. Automatic weather stations (AWS) were installed during
the late 1980’s and provide records of change in height of the snow pack, and airtemperature at hourly intervals between annual site visits. These records highlightthe timing of significant temperature (melt onset, intensity, and freeze-up) andsnowfall events throughout the year, and provide evidence from which themagnitude and timing of internal accumulation events are inferred.
In-situ observations reveal a significantly negative mass balance of all high Arctic
monitoring sites since observations began, becoming increasingly negative since
the late 1980’s (Figure 1). The South Melville ice cap exhibits the most negative-1
cumulative balance of all sites monitored, having thinned by ~18 cm a and losing
3
approximately 0.45 km of ice over the past >40 years. The Meighen ice cap has
24

-1

experienced a significantly lower thinning rate of ~10 cm a , equating to a total
3
volume loss of ~0.3 km since 1960. The White Glacier and the Northwest Devon-1
sector have been thinning by 13 and 8 cm a respectively. While preliminary
results indicate that the in-situ data can be used to accurately quantify volume
changes of small stagnant ice caps (Koerner and Burgess, in preparation), thespatially limited nature of these data, and their lack of accounting for mass loss due
to ice-berg calving, make them less representative of mass changes across themuch larger dynamic ice caps.

Figure 1. Cumulative net mass balance
for 5 ice caps/glaciers in the Canadianhigh Arctic. Mass balance values are
spatially
averaged
across100m
elevational bands at the basin-wide
scale for the Devon ice cap and White
Glacier measurements, and at the icecap wide scale for the Meighen andMelville observations. Cumulative massbalance measurements along theAgassiztransect
represent
2dimensional values only.

As part of NRCan's Glacier-Climate Observing System (GCOS), efforts are
currently underway to augment in-situ observations with remote sensing data inorder to provide a more complete estimate of mass balance across all land ice in
the Queen Elizabeth Islands. One component of mass balance that will be includedin GCOS is annual calving flux from tidewater glaciers that drain the main ice caps
in the QEI. Surface velocities of 12 tidewater glaciers draining the Prince of Wales,
North Ellesmere, and Axel Heiberg ice caps have been mapped on an annual
basis since 2000, and on the Devon ice cap since 2005, using radar speckle
tracking techniques as described in Short and Gray (2005). Annual velocity
measurements coincide with ice thickness profiles flown by NASA in 2005 acrossthe terminus of 5 major outlet glaciers that drain the Prince of Wales ice field andthe Devon Island ice cap. Both datasets (ice thickness and surface velocity) are
required to quantify flux through a glacier cross-section thus, future campaigns
aimed at acquiring ice thickness profiles across the remaining flux gates are a highpriority for this program. The second component of mass balance that we plan tomonitor with remote sensing technology is the spatial distribution of waterequivalent snow mass and near surface firn across land ice in the QEI. Preliminary
results from the CryoSat-2 Cal/Val experiments over the Devon Island ice cap
reveal that these layers can be tracked continuously over large distances fromASIRAS (Figure 2), which is an airborne version of the Ku band radar altimeter that25

Elevation (m a.s.l.)

Depth Below Surface (cm)will be onboard CryoSat-2. Over regions where the average density of the near-3
surface firn is relatively low (ie < ~500 kg m ), annual firn layers up to ~4 m below
the ice cap surface can be identified. These measurements provide important
insight into the topographic, elevational, and distance-from-moisture-source
controls on snow distribution, and provide spatial continuity to point measurements
collected in the field. Incorporating remote sensing techniques with traditional insitu observations will thus provide a more complete estimate of mass balance ofland ice in the Arctic regions.
0
100
200
300
400
0

10

20

30

40

50

60

40

50

60

Distance (km)
2000
1800
1600
1400
1200
0

10

20

30

Distance (km)Figure 2. Upper frame displays a section of radar backscatter from the ASIRAS airborne
radar altimeter (13 Ghz) acquired across the Devon ice cap in May, 2006. Above ~1800 m
a.s.l., subsurface firn layers were tracked continuously over horizontal distances of >20 km
to depths of ~300 cm below the surface. Below ~1400 m a.s.l., boundaries between annual
layers were obscured by the effects of wind and downward percolation of meltwater within
the snowpack.

In-situ mass balance measurements acquired from high Arctic ice caps andglaciers to date provide an invaluable climate record from a region that is predicted
to experience significant warming over the next century (IPCC, 2007). The
development of an operational mass balance monitoring system through GCOS,
based on integration of in-situ and remote sensing observations, will provide
knowledge critical for quantifying freshwater flux to the ocean, and forunderstanding the future impact of climate change on the state of health of land ice
across the Canadian high Arctic.

References
IPCC. 2007. Climate Change 2007: The Physical Science Basis: Contribution of Working
Group 1 to the Fourth Assessment Report of the Intergovernmental Panel on Climate
Change, S. Solomon et al., Eds., Cambridge University Press, New York, 2007.
Short, N.H. and A.L. Gray, 2005. Glacier dynamics in the Canadian High Arctic from
RADARSAT-1 Speckle Tracking. Can. J. Remote Sensing, Vol. 31, No. 3, pp. 225–239.

26

TERRERSTRIAL PHOTOGRAMMETRY FOR MEASUREMENTS OF THE
CALVING RATE OF KRONEBREEN, SVALBARD
ANNE CHAPUIS, CECILIE ROLSTAD, AND IVAR MAALEN-JOHANSEN
Department of Mathematical Sciences and Technology, Norwegian University of Life
Sciences (UMB).

Abstract
Calving is a large component in the mass budgets of tide water glaciers.
Measurements of calving rates are essential for the determination of a calving lawthat can be used in a mass balance model of the glacier. Changes in climate mayalter the ice dynamics and the melt water production, thus the calving rates maychange rapidly. Despite their importance, calving processes are not well
represented in numerical ice dynamic models, partially due to the difficulties anddangers of field data collection. Measurements of calving must be conducted in asafe distance to the calving front, and in this project we have applied terrestrial
photogrammetry to the front of Kronebreen, Svalbard, in August-September 2007.
Seven repeated stereo pairs were photographed from Nielsen-fjellet mountain side
over a four-days period. Markers were placed on the strand and measured withGPS to georeference the images. The digital images are processed
photogrammetrically for velocity measurements and volume estimates of thecalving ice loss. The velocity data are compared to ground based interferometric-1
radar measurements conducted in the same period, showing a velocity of 2.7 md .

27

SURFACE MASS-BALANCE MODELING OF NORDENSKIÖLDBREEN,SVALBARD
MARIANNE DEN OUDEN, C.H. REIJMER AND J. OERLEMANS
Institute for Marine and Atmospheric Research, Utrecht University, the Netherlands

Introduction
Evidence shows that increased meltwater production on a glacier or ice sheet may
result in higher ice-flow velocities (e.g. Iken and Bindschadler (1986), Zwally andothers (2002)). This mechanism is an especially interesting feature on Arctic
glaciers, as this area experiences large climatic changes (ACIA, 2005). To
investigate this relation on seasonal and inter-annual time scales, on the one handdetailed temporal information on the flow velocity is needed, while on the other
hand the amount of meltwater on a glacier has to be estimated. Finally, the routefrom the surface to the bed is of interest.
On the Arctic glacier Nordenskiöldbreen (Svalbard) the flow velocities aremeasured by means of a stand-alone GPS technique. Furthermore, the net massbalance is measured using stakes, and one Sonic Ranger (SR) providing acontinuous record of the change in surface height. However, measurements offer apoor spatial resolution. Thus, to gain more insight in the spatial variations of melt
distribution, we present a mass-balance model, based on the equations describedin Klok and Oerlemans (2002), and (preliminary) results for the year 2006-2007.

Figure 1. Map of the outlet glacier
Nordenskiöldbreen.
The
red
dots
represent the mass-balance stakes, theblue dot gives the location of the SR andthe green dot marks the location of the
GPS reference station.

28

Location
Nordenskiöldbreen is located on central Spitsbergen (Figure 1). It has an area of2
242 km , and the length of its main streamline is ±26 km (Hagen and others, 1993).
Its depth varies from ±600 m at 7 km from the terminus to ±100 m at the top,
1200 m a.s.l. (V. Pohjola, personal comm.). The glacier flows with an averagevelocity of ±50m/a (GPS). Currently, Nordenskiöldbreen is changing from atidewater glacier into a glacier ending on land.

MB modelFollowing Klok and Oerlemans (2002), the mass-balance model is based on thefollowing equation for the specific mass balance:
M=

€





∫ QL + P + QL dt,
m

L

m

S

(1)

where Qm/Lm is the mass lost through melt, P is the mass gained from precipitation
and QL /LS is the mass lost through sublimation. To obtain the energy available formelt, the model calculates the six components of the total surface energy flux,
namely incoming and reflected solar radiation (Sin and Sout), incoming and outgoing
longwave radiation (Lin and Lout), sensible turbulent heat (QH) and latent turbulent
heat (QL). In case the surface temperature is not at the melting point, or the flux is
directed away from the surface, the energy (G) is used to heat or cool the surface
layer, otherwise the energy is available for melting (Qm):(2)

Sin − Sout + Lin − Lout + QH + QL = QM + G
thThe model runs for the period March 30 2006 until March 21
corresponds to the period of continuous MB observations.
€st2007, which

Input
As input to the model we use synoptic weather data from two stations located in
the vicinity of Longyearbyen (see Table 1).
The air temperature at different heights is obtained by using a constant lapse rateof 6.5 K/km, based on measured temperatures. The cloud fraction is assumed
constant with height. Lastly, the precipitation is determined so that its annual sum
at a specific elevation corresponds with the observed snow pack at the beginningof the melt season.
The record of the SR and stake mass-balance measurements serve as comparison
for the model output. The SR is located at an altitude of ±430 m a.s.l., while the 7
stakes are placed at 100m altitude intervals, ranging from ±150 m a.s.l. to ±700 m
a.s.l. (Figure 1).
Parameter
Air temperature, Ta
Relative humidity, RH
Air pressure, p
Cloud fraction, nPrecipitation, P

Table 1. Details input data
StationDistance
Adventdalen±60 kmAdventdalen±60 kmAdventdalen±60 kmLYB Airport±60 kmLYB Airport±60 km29

Altitude
75 m a.s.l
75 m a.s.l
75 m a.s.l
28 m a.s.l
28 m a.s.lTime interval
10 min10 min10 min
6 hour12 hourResults and comparison to observations
In a general sense the model performs reasonably well. At the altitude of the SR for
example, the snow pack disappears at approximately the right moment (±10 days
too early in the model).
To validate the model, its output is compared to observations (Figure 2). In the firstgraph the model places the onset of melt too early. Another noticeable feature is
the discrepancy between observed and modelled precipitation. In the beginning themodel misses a snow event altogether, while in the second half of the year there isa mismatch in timing. A likely reason is the large distance between the source ofthe precipitation input and the target glacier.
In Figure 2b the model captures the net annual mass balance below the ELA(located at ±600 m a.s.l.) well. Nevertheless, it underestimates the mass balance inhigher parts. Note that the observations in Hagen and others (2005) span the
period 1986-1996. In recent years the mass balance has probably increased (V.
Pohjola, pers. comm.), enlarging the difference even more. This discrepancy iseither caused by an overestimation of melt, an underestimation of precipitation
and/or the absence of refreezing.

Figure 2. a) Modelled change in surface height (green line) compared to observations with aSR(blue line). Please note that the instrument detached from its housing around day 165. b)
Modelled mass balance (green line) compared to mass-balance observations for 2006-2007(blue dots) and from Hagen and others, 2005 (red dots).

Outlook• Fine tune the model, including an extensive sensitivity analysis• Run the model with different input data (Temperature series from GPS,RACMO)• Extend the model to a 2D topography• Include a refreezing routine

Acknowledgements
We thank the NPI for providing a DEM. Also, we are grateful for weather data
supplied by NMI and UNIS.
30

References
ACIA, 2005. Arctic Climate Impact Assessment Report. C. Symon, L. Arris and B. Heal, eds.
Cambridge University Press, Cambridge, UK, pp 146.
Hagen, J.O., O. Liestøl, E. Roland, and T. Jørgensen. 1993. Glacier atlas of Svalbard andJan Mayen. Norsk Polarinst., Oslo.
Hagen, J.O., T. Eiken, J. Kohler and K. Melvold. 2005. Geometry changes on Svalbard
glaciers: mass-balance or dynamic response?. Ann. Glaciol., 42: 255-261.
Iken, A. and R.A. Bindschadler. 1986. Combined measurements of subglacial waterpressure and surface velocity of findelengletscher, Switzerland: conclusions about
drainage system and sliding mechansim. J. Glaciol., 32(110):101–119.
Klok, E.J. and J. Oerlemans. 2002. Model study of the spatial distribution of the energy andmass balance of Morteratschgletscher, Switzerland. J. Glaciol., 48(163), 505–518.
Zwally H.J., W. Abdalati, T. Herring, K. Larson, J. Saba, and K. Steffen. 2002. Surface
melt-induced acceleration of greenland ice-sheet flow. Science, 297:218–222.

31

ICEBERG CALVING FLUX AND MASS BALANCE OF
AUSTFONNA ICE CAP ON NORDAUSTLANDET, SVALBARD

THE

J.A. DOWDESWELL1, T.J. BENHAM1, T. STROZZI2 AND J.O. HAGEN3
1

Scott Polar Research Institute, University of Cambridge, Cambridge, UK
Gamma Remote Sensing, Gümligen, Switzerland
3
Department of Geosciences, University of Oslo, Oslo, Norway
2

Abstract
Satellite radar interferometry, 60 MHz airborne ice-penetrating radar data, andvisible-band satellite imagery were used to calculate the velocity structure, ice
thickness, and changing ice-marginal extent of Austfonna (8,120 km2 and 2,500
3km ), the largest ice cap in the Eurasian Arctic. Ice-cap motion is less than about-1
10 m yr , except where faster-flowing curvilinear features with velocities of several-1tens to over 200 m yr . Most drainage basins of Austfonna have undergone icemarginal retreat over the past few decades at an average of a few tens of metres
per year. Integrating margin change around the whole ice cap gives a total area
2
-1
loss of about 10 km yr . Iceberg flux from the marine margins of Austfonna is
3-1
about 2.5 ± 05 km yr (w.e.), about 45% of the total calving flux from the whole
Svalbard archipelago. When mass loss by iceberg production is taken into account,3
the total mass balance of Austfonna is negative, by between about 2.5 and 4.5 km-1yr . This iceberg flux represents about 33 ± 5% of total annual mass loss from
Austfonna, with the remainder through surface ablation. Iceberg flux should be
included in calculations of the total mass balance of the many large Arctic ice caps,
including those located in the Russian and Canadian Arctic that end in tidewater.
The neglect of this term has led to underestimates of mass loss from these ice
caps and, thus, to underestimates of the contribution of Arctic ice caps to global
sea-level rise.

32

APPLICATION OF A NUMERICAL ICE SHEET MODEL ON
AUSTFONNA, SVALBARD - MODEL PREPARATIONS AND INPUTDATA
THORBEN DUNSE1, THOMAS V. SCHULER1,2 AND JON OVE HAGEN1
1
2

Department of Geosciences, University of Oslo, Norway
Norwegian Water Recourses and Energy Directorate, Oslo, Norway
2

Covering an area of 8120 km , Austfonna on Nordaustlandet, Svalbard is among
the largest arctic ice caps outside Greenland (Figure 1). It has a polythermal
regime with a simple dome-shaped topography and well defined drainage basins,
several of them known to be of surge-type.
Surface elevation measurements since the mid 1990s suggest thickening in the1interior and thinning at the margins . There is still considerable uncertainty about towhat degree this changes are driven by surface processes (accumulation and
2,3,4
ablation) or ice dynamics (possible build-up towards surge activity) . To better
understand both surface processes and glacier dynamics, we complement ongoingglacier observations with numerical modelling.
We use the Simulation Code for Polythermal Ice Sheets (SICOPOLIS), developed
5
by Ralf Greve . The code accounts for different mechanical behaviour of cold andtemperate ice regions, separated by a phase-transition surface, called the coldtemperate transition surface, and for which jump conditions of mass, momentum
6and energy must be fulfilled . It is based on the shallow-ice approximation anduses finite-difference discretization.

Figure 1. Map of Svalbard with the red frame indicating the position of the model domain
Austfonna on Nordaustlandet.

33

Figure 2. Colour-coded bedrock topography of the model domain overlain by a mesh of theAustfonna ice cap’s current surface topography. The white line in the colour code represents
the sea level and the grey line the 250-m contour line, while the black line is encircling theice-cap boundary. Note that Vestfonna is shown as bedrock, since ice thickness data of thisice cap is not yet included.

Modelling ice dynamics requires accurate description of surface and bedrock
topography (Figure 2) and of surface mass balance. We present a basic data set,
7,8,9
which is based on previous studies
and that enables us to run SICOPOLIS on
Austfonna. Further data collection will progressively refine this data set and serve
to validate the model.
Given the required input data the model computes ice extent and thickness,
velocity and temperature fields, water content within temperate ice regions, age ofthe ice and the displacement of the lithosphere.
Preliminary computations show that the model produces high velocity units that4
coincide with subglacial troughs and flow units in an interferrogram (Figure 3).

Acknowledgments
This work is supported by the Norwegian Research Council (IPY, GLACIODYN),
the CryoSat Calibration and Validation project (ESA contract no. C90118), and theSvalbard Science Forum (Arktisstipend). Thanks go to Ralf Greve who generously
supplied the model code and to Julian Dowdeswell for making the results from the
1983-RES survey over Austfonna available.

ab

Figure 3. Preliminary model results of the surface velocity field overlying contours of theinitial surface topography (a). ERS-1 SAR interferogram superimposed on Landsat image
4(b), modified from Dowdeswell .

34

References[1] Bamber, J., W. Krabill, V. Raper, and J. Dowdeswell. 2004. Anomalous recent growth ofa part of a large Arctic ice cap: Austfonna, Svalbard. Geophysical Research Letters, 31,
L12402-1 – L12402-4.
[2] Bevan, S., A. Luckman, and T. Murray. 2007. Positive mass balance during the late 20th
century on Austfonna, Svalbard revealed using satellite interferometry. Annals of
Glaciology, 46, 117-122.
[3] Hagen, J.O., T. Eiken, J. Kohler, K. Melvold. 2005. Geometry changes on Svalbard
glaciers: mass-balance or dynamic response? Annals of Glaciology, 42, 255 - 261.
[4] Dowdeswell, J. A., B. Unwin, A.-M. Nuttall and D. J. Wingham. 1999. Velocity structure,
flow instability and mass flux on a large Arctic ice cap from satellite radar interferometry.
Earth and Planetary Science Letters 167, 131-140.
[5] Greve, R. 1995. Thermomechanisches Verhalten polythermer Eisschilde - Theorie,
Analytik, Numerik. Doctoral dissertation, Department of Mechanics, Darmstadt Universityof Technology, Germany, 226 pp. Shaker Verlag, Aachen, Germany.
[6] Greve, R. 1997. A continuum-mechanical formulation for shallow polythermal ice sheets.
Philosophical Transactions of the Royal Society London A 355 (1726), 921-974.
[7] Dowdeswell, J. A., D. J. Drewry, A. P. R. Cooper and M. R. Gorman. 1986. Digital
mapping of the Nordaustlandet ice caps from airborne geophysical investigations. Annals
of Glaciology 8, 51-58.
[8] Schuler, T.V., E. Loe, A. Taurisano, T. Eiken T., J.O. Hagen and J. Kohler. 2007.
Calibrating a surface mass balance model for the Austfonna ice cap, Svalbard. Annals of
Glaciology, 46, 241-248.
[9] Taurisano, A., T.V. Schuler, J.O. Hagen, T. Eiken T., E. Loe, K. Melvold and J. Kohler.
2007. The distribution of snow accumulation across Austfonna (Svalbard): direct
measurements and modelling. Polar Research 26 (1), 7-13.

35

GREENLAND SURFACE MASS BALANCE USING A REGIONALATMOSPHERIC CLIMATE MODELJANNEKE ETTEMA1, MICHIEL VAN DEN BROEKE1, ERIK VAN MEIJGAARD2JONATHAN BAMBER3 AND RUPERT GLADSTONE3
1

Utrecht University, Institute for Marine and Atmospheric Research, The Netherlands
Royal Netherlands Meteorological Institute (KNMI)
3Bristol University, School for Geographical Sciences, UK
2Introduction
As part of the RAPID Climate Change program we investigate the effects of rapid
climate changes on the surface mass balance and freshwater contributions of theGreenland ice sheet. For this purpose we use the Regional Atmospheric Climate
Model (RACMO2.1/GRN). The period September 1989 to December 2005 issimulated to validate the model output with other regional climate model studies
and direct mass balance measurements along the K-transect in West-Greenland.
This leads to an assessment of the spatial and temporal variability in the surface
mass balance (SMB) and its subcomponents.

Figure 1. Left: RACMO2.1/GRN model domain for Greenland with dashed height contours
every 250 m. Dark blue area is land ice, green tundra and light blue sea or sea ice. Right:
annual surface mass balance for 2004 in mm w.e.; values have no meaning over land
points.

Model setup
RACMO2.1/GRN is adapted from the second version of the Regional Atmospheric
Climate Model (RACMO). It consists of the atmospheric dynamics of the HighResolution Limited Area Model (HIRLAM) and the physical processes package
from the European Centre for Medium-Range Weather Forecasts (ECMWF). To
36

better represent the conditions on Greenland the multi-layer snow parameterizationof Bougamont et al (2005) is implemented. This scheme computes subsurface andsnow metamorphism processes like ice and snow melt, percolation and refreezingof meltwater, snow densification and ageing within the upper 25 meter of firn.
Feedback between the snow surface and the overlying atmosphere is based on askin temperature formulation and the surface albedo as function of the snow
density of the first firn layer (Greuell and Konzelman, 1994).
The horizontal resolution of RACMO2.1/GRN is about 11 km. The model has 40
hybrid levels in the vertical, of which the lowest is about 10 m above the surface.
ERA-40 fields force the model at lateral boundaries complemented with operational
analyses until December 2005. Height data of the digital elevation model ofBamber et al. (2001) is used for an accurate topographic representation as shownin Figure 1a.

Model evaluation
The surface mass balance in the model is defined as the total precipitation (solid
plus liquid) minus meltwater production and sublimation. The runoff component in
the model is incorrectly parameterized, which makes this data unusable for explicit
SMB calculation. Therefore, we assume that the meltwater production is anacceptable first order estimate of the runoff. This may lead to an overestimation of
the mass loss since part of the meltwater will refreeze in the firn layer.
The pattern of the annual SMB of 2004 shown in Figure 1b agrees with otherestimates based on observations (Zwally and Giovinetto, 2001) and models(Fettweis, 2007; Box et al., 2006). The simulated minimum and maximum values
are more extreme than in previous estimates. Most mass is gained northeast of
Baffin bay and the southeastern coastal area. These areas are associated withsteep slopes that enhance solid precipitation rates. The regions with the mostnegative mass balance occur along the western margin south of Jacobshavn
glacier and the northern ice margin where most melt occurs. Due to the lowtemperatures in the latter area, most of the meltwater will refreeze and actually notcontribute to mass loss.

Figure 2. Left: Time series of the annual total precipitation for period 1990-2005. Right: thesame plot as a for the annual surface mass balance

37

The RACMO2.1/GRN annual precipitation rates for the period 1990-2005 areplotted in Figure 2a, together with estimates based on the MAR (Fettweis, 2007)
and the polar MM5 model (Box et al., 2006). RACMO2.1/GRN shows an enhanced
precipitation rate mainly due to its horizontal resolution of 11 km, where the MAR
model is running at 25 km and MM5 at 24 km resolution. The higher precipitation
rates in RACMO2.1/GRN lead to a more positive SMB than the other models as
seen in Figure 2b. The simulated interannual variability of the different SMB
components is consistent with the other models.
In-situ data helps us to determine if the larger precipitation rates simulated are
realistic or not. Figure 3 displays the simulated and observed (Van de Wal et al.,2005; complemented by pers. comm. Van de Wal) surface mass balance along the
K-transect located at 67°N on the western margin in the ablation zone. For the year
2004, the simulated ablation is slightly weaker close to the ice margin due to an
overestimation of winter snowfall, which prevents the darker glacier ice to surface
early in the melt season. Overestimation further from the margin may be the resultof the assumption that all meltwater produced runs off, while at these locations
about half of the meltwater refreezes in the firn. Overall, the modeled values staywithin the range of the measurements, which is promising for further validation of
the model with other observational datasets.

Figure 3. Surface mass balance as function of the distance to the ice margin. Dots are themeasurements are for the period 1990-2007; black line is the mean over the entire period;
the green line for the year 2004. The blue line is RACMO2.1/GRN for 2004.

Conclusions and outlook
RACMO2.1/GRN is able to simulate a realistic spatial and temporal SMB for theGreenland ice sheet. Compared to other model studies it simulates enhanced
precipitation rates, mainly on the steep southeastern slopes. Even with theassumption made that all meltwater produces runs off, the total SMB is morepositive than in other estimates.
38

The stake measurements show a steep gradient in SMB with distance from themargin, which is reasonably represented in RACMO2.1/GRN. Resolving the runoff
problem by removing more efficiently the water from the ice surface and refreezing
the meltwater in the firn pack will further improve the model agreement.
More validation of the model results is planned for the near future using available
automatic weather stations data, ice core accumulation datasets and coastal
precipitation gauge data.

References
Andreas, E.L. A theory for the scalar roughness and the scalar transfer coefficients oversnow and sea ice. Boundary-Layer Meteorol., 38(1-2):159–184, 1987.
Bamber, J.L., S. Ekholm, and W.B. Krabill. A new, high resolution digital elevation model ofGreenland fully validated with airborne laser data. J. Geophys. Res., 106:33773–33780,
2001.
Bougamont, M., J. L. Bamber, and W. Greuell. Development and test of a surface massbalance model for the greenland ice sheet. J. Geophys. Res., 110(F04018), 2005.
doi:10.1029/2005JF000348.
Box, J.E., D.H. Bromwich, B.A. Veenhuis, L-S. Bai, J.C. Stroeve, J.C. Rogers, K. Steffen, T.
Haran, and S-H. Wang. Greenland ice sheet surface mass balance variability (19882004) from calibrated polar MM5 output. J. Climate, 19(12):2783–2800, 2006.
Fettweis, X. Reconstruction of the 1979-2006 Greenland ice sheet surface mass balanceusing the regional climate model MAR. The Cryosphere Discuss, 1:123–168, 2007.
Greuell, W. and T. Konzelman. Numerical modelling of the energy balance and the englacial
temperature of the Greenland ice sheet. Calculations for the ETH-camp location (WestGreenland, 1155 m a.s.l.). Global and Planetary Change, 9(1-2):91–114, 1994.
doi:10.1016/0921-8181(94)90010-8.
Van de Wal, R.S.W., W. Greuell, M.R. van den Broeke, C.H. Reijmer, and J. Oerlemans.
Surface mass-balance observations and automatic weather statiuon data along atransect near kangerlussuaq, west greenland. Ann. Glaciol., 42:311–316, 2005.
Zwally, H.J. and M.B. Giovinetto. Balance mass flux and ice velocity across the equilibriumline in drainage systems of greenland. J. Geophys. Res., 106, 2001.

39

CHANGES OF GEOMETRY OF RENARDBREEN (SPITSBERGEN)
AND EVOLUTION OF MORPHOLOGY OF ITS FOREFIELD DURING THE
LAST CENTURY
G. GAJEK1, M. GRABIEC2, D. PUCZKO3 AND J. JANIA2 " \L 3 \N
1

M. Curie-Sklodowska University, Poland
University of Silesia, Poland
3
Institute of Geophysics, Polish Academy of Sciences
2

Introduction
Warming of the climate during the last century caused changes in geometry(volume and extent) of the Arctic glaciers. The process has gained in speed in the
recent decades (ACIA, 2005). Such intensive changes can also be observed on
Svalbard, in tidewater glaciers. Special cases are the glaciers that – due togeometry changes – have retreated from the sea to land. The presented project
covers transformation of tidewater glaciers into land-based ones, on the example
of Renardbreen (northwestern Wedel Jarlsberg Land, southern Spitsbergen). The
main purpose of the project is to create a model of glaciological andgeomorphological processes during such a transformation. This poster presents
the geometry changes of the glacier throughout the 20th century, as well as thegeomorphological consequences of the recession on its forefield.

Figure 1. Changes of the Renardbreen glacier area and terminus position during lastcentury.

40

The glacier and its environmental setting
Renardbreen is located in the central area of the northwestern Wedel Jarlsberg
Land (Figure 1). It is surrounded from the west, north and south by the mountains
reaching 831 m a.s.l. It is the largest glacier in the area. Its drainage basin covers
22
the area of 39 km , while the area of the glacier is 28 km (in 2006). The glacier is
8 km long and its width ranges from 2.5 to 6 km. The glacier surface is exposed
towards north-east. The highest parts of the glacier reach 720 m a.s.l. and its front
descends almost to the sea level. The tongue of Renardbreen is supplied by six
firn fields, amphitheatrically located at 450-550 m a.s.l. The glacier is relatively
poorly crevassed. Minor crevasses can be seen in the southern part of theterminus; major crevasses are located in the break of the glacier on 200-250 ma.s.l. In the early decades of the 20th century, the glacier tongue descended
towards Recherche Fjord. It is on average 3.5 km wide and it gets slightly wider
towards mountain range forefield. In its northern part the ice wall currently lies ontransformed, rocky raised marine terraces on the level of several to 40 m a.s.l. andin its southern part it descends almost to the sea level, while it used to reach
Josephbukta bay. The glacier snout is flat-protuberant. At present the glacier does
not reach the shoreline, although in late 1980s calving of the tongue was observedin the narrow part of the active cliff.
Renardbreen glacier is sub- and supraglacially drained. Subglacial channels
function predominantly in the southern and central part of the glacier tongue. The
main outflow is located in the central part of the tongue and has a form of a glaciermouth that is 8 m wide and 3-5 m high. The mouth usually activates as late as thesecond decade of July. The glacier is also hydrologically active during winter, theevidence of which are naled ice covers in its forefield. In the middle of the summer
3
season the total outflow from the glacier reaches approximately 5 m /s
(Bartoszewski et al., 2006).
Up till 2005, the glacier was not systematically observed from the glaciologicalpoint of view. Located nearby is a building, formerly a part of the Calypsobyen
mining camp – currently it serves as a base for seasonal expeditions of scientists
from Maria Curie-Skłodowska University.

Geometry changes
Evaluation of the geometry changes of Renardbreen was possible thanks to usingarchival maps, aerial photos, ASTER satellite imagery and GPS measurements.
The materials used are as follows:• Topographic map in 1:100 000 scale, issued by Norwegian Polar Institute,based on oblique aerial photos taken in 1936;• Photogeological map of Renardbreen, Scottbreen and Blomlibreen forefield in
1:10 000 scale, based on aerial photos from 1960 (Szczęsny et al. 1989);• Orthophoto map of the northwestern part of Wedel Jarlsberg Land in 1:25 000
scale, based on aerial photos from 1990 (Zagórski 2005);• GPS measurements carried out using the kinetic method in 2005-2007.
Furthermore, publications on geomorphological research of the region (Reder
1993, 1996) were referred to.
At the end of the 19th century, the glacier wall was reaching the frontal moraine
rampart, formed during the Little Ice Age. In 1896, the glacier ended with an ice cliff41

in the Recherche fjord, in the area now occupied by Josephbukta bay (Figure 1).
Between 1896 and 1936, the cliff recessed by 400 m to the west, which averages
to a recession speed of 10 m per year. In 1936, Josephbukta was a small bay,surrounded by an ice cliff from the south, west and north. Between 1936 and 1960
the bay quickly expanded, mainly towards west and north. The northern part of theglacier tongue, neighbouring the moraine ramparts and separated from the watersof the bay, was undergoing a relatively slow frontal recession. At that time,recession of the glacier reached 250-300 m on land and up to 400-500 m by thebay, averaging 16 m per year. The northern part of the tongue recessed at an
average speed of 12.5 m per year, while the southern part – at 20.8 m per year,
the latter was mainly due to glacier calving.
In 1960, the shape of the ice cliff was irregular, featuring small embayments andtips. It is most probably the result of the influence of the bedrock on the calving
frequency and its low speed on shallow waters. At that time, bay expansion into theglacier was hardly noticeable, the ice cliff moved slightly towards the west.
Recession of the glacier between 1960 and 1990 totalled to only 292 m, with its
pace dropping to 9.7 m per year. However, the northern part, rather little dynamic
before, started retreating much faster. During those thirty years, the glacier wall
recessed by 684 meters, averaging to 22.8 m per year. Field observations carriedout in the early 1990s led – among other things – to the discovery of preserved dry
marginal river channels, parallel to the northwestern edge of the glacier wall, andindicated that the recession pace of the on-land part of the glacier wall was about
20 m per year (Reder 1993, 1996).
Table 1. Changes of the Renardbreen glacier area and terminus position during last century
1936 1960
1990
2005
on-land bay
on-land bay on-land
bay
part
part
part
Changes of the
0
300
500
984
792
1245
898
Renardbreen terminus
position (metres)changes of the Renardbreen
0
-2,1
-3,4
-4,7
2
area (km )2
Renardbreen area (km )
33
30,9
29,6
28,3

In the spring of 2005, the glacier front extent was measured using differential GPS
method. In comparison to 1990, the glacier extent in the northern land based parthas decreased by 261 m, and in the southern part (by the bay) – by 106 m. It
averages to 17.4 m per year and 7.0 m per year respectively. GPS differential
measurements carried out in the spring of 2006 proved that the glacier terminus
recessed on average by 12 m when compared to 2005. The recession speed
decrease at that time may be connected with the change in the position of theglacier tongue, which retreated to a higher hypsometric level (Table 1).
The glacier front recession is accompanied by the changes in ice thickness.
Elevation measurements have been taken in 2005 along two longitudinal profiles in
the northern and southern part of the glacier. In the frontal part, glacier height42

reduction reached 25-30 m when compared to 1990, which averages to 1.6-2.0 m
per year.

Geomorphology of the glacier forefield and its changes
At the time when the glacier descended to the sea, the marginal zone consited of
ice-core frontal and lateral moraines and areas of sub-marine glacigenic deposits.
The geomorphology of the marginal zone was registered as the photogeological
map in 1960 (Szczęsny at al., 1960). Routes of sandur outflow, a few esker ridges
and – next to the frontal moraine – kame terrace developed in the central zone.
The characteristic form of the forefield is a vast naled ice, upon which fluvioglacial
forms are developing. The sediments were deposited by the northern subglacial
lateral torrent. They constituted almost 20% of the whole marginal zone at thattime. The southern part of the glacier snout had a form of an ice cliff. Sub-marine
deposition was related to the ground moraine and the material transported by two
subglacial rivers – the central river and the lateral southern river (bedload andsuspended matter of turbid subglacial waters).

Figure 2. Northwesternpart of Wedel Jarlsberg
Land (Spitsbergen,
Svalbard, Norway).
Orthophoto 1:25 000,
Zagórski 2005, [ed.] K.
Pękala and H. F. Aas.

43

As a result of the recession of the glacier, its marginal zone expanded significantly
by 1990 (Figure 2. Zagórski 2005). Distribution of forms and sediments depends on
the topography of the bedrock, i.e. the raised marine terraces, remodelled byglacial processes. Apart from the rock-knobs covered with a thin layer of groundand fluted moraine, large areas are occupied by fluvioglacial sediments, formed by
the central outflow among others. The statistics of the area of the marginal zone
occupied by particular landforms are as follows: moraines (ablation moraine,
terminal and lateral moraine, ground and fluted moraine) occupied 56% of theRenardbreen forefield in 1990, while the fluvioglacial forms (young outwash plains,
old outwash plains, kame terraces, delta and others) – 44%.

Conclusions
Recession of the cliff part of the ice wall in the period of intensive calving was
faster in comparison to the extent change of the on-land part of the ice wall. Whenthe ice cliff retracted into a shallow bay and the calving slowly died away, therecession rate slowed down from 20 m per year to 10 m per year recently.
Transportation of mineral deposits by the subglacial rivers resulted in forming adelta and shallowing of the bay.
Phase 1. The glacier terminates with an ice cliff in the sea. All of the fluvioglacial
and ground moraine sediments are deposited in the bay. The bay is being
shallowed directly in front of the terminus. The delta emerges and is built up.
Phase 2. Intensive frontal recession takes place. Glacier terminus gets separated
from the sea water by the sediments. The ice cliff withers. Ground moraine material
is no longer deposited in the bay. The material is transported by subglacial rivers(from the central outflow and minor marginal outflows). Fluvioglacial forms start todominate on the forefield.
Phase 3. As a result of intensive ablation and predominance of negative massbalance, the recessing glacier uncovers larger area of internal marginal zone,
between the ice-moraine ridges and the contemporary edge of the ice wall, in thearea of which fluvioglacial accumulation forms start to prevail. The marginal rivergains length, and thus the volume of material accumulated in the bay isdecreasing. The delta is built up at a much slower rate.
Renardbreen seems to be a good example of the evolution of frontal and marginalzone of Spitsbergen glaciers retreating to the shallow seas. Quantitative analysis ofmineral mass transfer from the glacier to its forefield and to the seathe is plannedin the next stage of the project.

Acknowledgements
The authors would like to thank all the colleagues of the 27th and 28th Polar
Expedition to Polish Polar Station on Spitsbergen for their assistance in conducting
the research. The special thanks go to Professor Marian Harasimiuk for his
valuable comments and advice. The study was financed by the Ministry of Scienceand Higher Education, Poland (grants No. PBZ-KBN-108/P04/2004 and special
grant No. IPY/296/2006)

44

References
ACIA 2005: Arctic Climate Impact Assessment - Scientific Report, Cambridge University
Press, 1042 pp.
Bartoszeksi, S., A. Gluza and K. Siwek, 2006. Selected problems of shaping of
meteorological and hydrological conditions in the northwestern part of Wedel Jarlsberg
Land (Spitsbergen) – The state and changes of natural environment of the northwesternpart of Wedel Jarlsberg Land (Spitsbergen) in the conditions of climate change andanthropopressure. 20 Years of Polar Research of the Institute of Earth Sciences UMCS
on Spitsbergen, UMCS, Lublin, 53-60.
Reder J. 1993. Contemporary deglaciation of the Recherche and Renard Glaciers in theBellsund region (Western Spitsbergen). – XX Polar Symposium, Lublin: 421-427.
Reder J. 1996. Evolution of marginal zones during a continued glacial retreat innorthwestern Wedel Jarlsberg Land, Spitsbergen. – Polish Polar Res. 17(1-2): 61-84.
Szczesny, R., J. Dzierzek, M. Harasimiuk, J. Nitychoruk, K. Pekala and J. RepelewskaPekalowa, 1989. Photogeological map of the Renardbreen, Scottbreen and Blomlibreen
forefield (Wedel Jarlsberg Land, Spitsbergen, 1:10 000). – UMCS Geographical
Expeditions to Spitsbergen. Wyd. Geol., Warszawa.

45

SOME RESULTS OF RADIO-ECHO SOUNDINGS OF FRONTAL PARTS
AND FOREFIELDS OF SELECTED GLACIERS IN SVALBARD AND N
SCANDINAVIAMARIUSZ GRABIEC1, WOJCIECH DOBIŃSKI1, DARIUSZ PUCZKO2, PIOTRDOLNICKI1
1
2

Faculty of Earth Sciences, University of Silesia, Sosnowiec, Poland,Institute of Geophysics, Polish Academy of Sciences, Warsaw, Poland

Objectives
This work presents results of GPR surveys performed in a transition zone between
glaciers and their forefields on selected locations in Svalbard and Northern
Scandinavia. The purpose of the study is to identify relations between glacial andperiglacial/paraglacial processes which occur at the glacier termini and in thenearest proglacial area, affected by permafrost (Dobinski 2006).

Figure 1. Location of GPR surveys.

Study area
Research was performed in Northern Sweden and the High-Arctic (Svalbard) inAugust and September 2007 (Figure 1).
• Storglaciären in Northern Sweden is a polythermal valley-glacier partly frozen to2
the bed with area of about 3.1 km (Pettersson et al 2003).
• Werenskioldbreen (S Svalbard) is a valley-type, polythermal glacier terminating2
on the land. The surface area of the glacier is about 28 km and the length ca. 7
km. Folding of medial moraine shows that glacier surge took place in the past.
2• Ariebreen in Hornsund area is a small (0.4 km ), entirely cold valley-glacier(Machio et al. 2007).
2• Hansbreen (56 km ) is an outlet glacier calving to Hornsund, with forefield only
partly developed on land. It is a typical subpolar glacier with cold ice layer.
46

Methods
Ground penetrating radar (GPR) was used for data collection for the study. GPR isan effective tool for identifying the sedimetological structure of subsurfaces of thelithosphere (Neal 2004). Particularly good survey depth has been found inglaciological research (Jansson et al 2000). At periglacial areas affected by icereach permafrost this method has been also applied (Moorman at al 2003). In ourwork the research was performed in glacial as well as in periglacial environments.
Surveys were performed with using ground-penetrating radar Ramac/GPR CUII,
equipped with unshielded 200 Mhz antennas. GPR surveys were conducted along
profiles, which started on frontal part of the glacier and finished on its forefield.
GPS receiver was used for recording position and distance of survey profile.
Radexplorer software and filters: DC removal, time-zero adjustment, backgroundremoval, amplitude correction, trace edit, topography and radio-wave velocity
models were used for data processing and analysis.
Radio-wave velocity in the glacier ice and in the forefield sediments was calculated
based on common mid-point (CMP) surveys. In the ice radio-wave velocity of 16
cm/ns has been calculated as a typical value. On the forefield affected bypermafrost the recorded velocity value was of about 10 cm/ns, whereas publishedvalues of velocity for permafrost are slightly higher: 11-15 cm/ns (RAMAC/GPR
Operating Manual).
Temperature measurements made with Pt100 thermistors gave indirect information
about buried ice existence.

Results
Analyzed GPR profiles show some common features as well as some differences.
In all cases contact between the glacier and its bed was detected and it is visible
very clearly (Figure 2). The permafrost table is also clear-cut noticeable. The depthof permafrost active layer is going to be thinner in the direction towards the glacierfront. Thickness of this layer is variable and reaches its maximum of about 2.8 m at
Storglaciären forefield (Figure 2.IV). On the forefield of Svalbard glaciers the active
layer is no more than 2 meters deep (Figure 2.I-III).
On the presented graphs ice cores and other types of buried ice have been also
found. In the frontal/lateral moraine of Hansbreen an ice-core is very likely (Figure2.I). Its thickness might be more than 5 meters. Ice-cored moraines are common in
vicinity of the measurement site. Summer ground temperature courses on themoraine are also typical, as for places where massive ice was detected. Groundtemperature increase is remarkably slower than in tundra with no ice body beneath.
Similar features in Ariebreen frontal moraine can be observed (Figure 2.III).
On Werenskioldbreen forefield, in general, the permafrost table is equal to thedepth of the buried ice (Figure 2.II). Thickness of this layer varies between 1.5 and2.5 m. Existence of massive ice covered by fluvioglacial sediments can be directly
verified in profiles cut out by the proglacial river. The thickness of the ice core/ice
layer is up to 6 m and sharply decreases in the direction towards glacier terminus,
47

which is covered by the debris cover. It seems to be possible that the glacier front
has connection with not active ice on the forefield, as it is shown in Etzelmüller andHagen (2005). Close to the front of Werenskioldbreen, in places where the ice issufficiently thin, big outburst of subglacial water might occur. Massive ice-bodies
were not detected on the forefield of Storgläciaren (Figure 2.IV). In this location on
the radargram also several englacial debris bands are visible. These bands areparallel to the foliation in the ice and hence have an arcuate shape.

Figure 2. Examples of GPR sections and interpretation on selected glaciers and theirforefields (uncorrected for topography): I – Hansbreen, II – Werenskioldbreen, III –Ariebreen, IV - Storglaciären.

Conclusions• Ground penetrating radar with 200 Mhz antennas is the effective tool for
research in glacial and periglacial environments especially for active layer,buried ice and glacier ice detection.
• The general scheme of overlapping glacial and periglacial environments occurs
in different parts of Arctic with some varieties.
• The occurrence of buried ice and ice-cored moraines on glacier forefields ismore likely in High Arctic than in lower latitudes. The thickness of buried ice
bodies detected on Svalbard may reach 6 m or even more.

48

• The active layer is thinning towards glaciers' fronts and the depth of permafrost
table varies from 2.8 m in Low Arctic to 1.5 m on Svalbard.
• The transition of ice from the glacial into the periglacial zone is gradual and theglacier ice becomes an element of the periglacial environment by losing its
movement steadily.
• In permafrost affected areas application of other geophysical methods i.e.
electroresistivity tomography can give more complete insight in the distributionof permafrost.

Acknowledgements
This research has been supported from finances for science in 2007 - 2010 as theresearch grants of the Polish Ministry of Science and Higher Education(IPY/269/2006; IPY/272/2006 and N30605232/3405) and the EU ATANS (FP6506004) grant.

References
Dobinski W 2006: Ice and environment: A terminological discussion. Earth Science Reviews
79: 229-240
Etzelmüller and Hagen 2005. Glacier – permafrost interaction in Arctic and alpine mountainenvironments with examples from southern Norway and Svalbard. In: Harris C., &Murton J.B., (eds.) 2005. Cryospheric systems: Glaciers and Permafrost. GeologicalSociety, London, Special Publications, 242, 11-27
Jansson P., Näslund J-O., Pettersson R., Richardson- Näslund C., Holmlund P., 2000.
Debris entrainment and polythermal structure in the terminus of Storglaciären. In: DebrisCovered Glaciers (Proceedings of a workshop held at Seattle, Washington, USA,
September 2000).IAHS Publ. no. 264, 2000. 143-151
Machio F., Lapazaran J., Dolnicki P., Petlicki M., Głowacki P., Nawarro F., 2004. Preliminary
results of radio-echo sounding at Ariebreen, Hornsund Spitsbergen. In: The Dynamics
and Mass Budget of Arctic Glaciers. Extended abstracts. Workshop and GLACIODYN(IPY) Planning Meeting, 15 - 18 January 2007, Pontresina (Switzerland), 64 – 66
Moorman B.J., Robinson S.D., Brugess M.M. 2003. Imaging periglacial conditions withground-penetrating radar. Permafrost and Periglacial Process. 14. 319-329.
Neal A., 2004. Ground-penetrating radar and its use in sedimentology: principles, problems
and progress. Earth-Science Reviews 66: 261–330
Pettersson, R., P. Jansson and P. Holmlund. 2003. Cold surface layer thinning on
Storglaciären, Sweden, observed by repeated ground penetrating radar surveys. J.
Geophys. Res., 108(F1), 6004. (10.1029/2003JF000024.)
RAMAC/GPR Operating Manual 1997. Version 2.28. Malå Geoscience. Appendix 1.

49

GLACIER GEOMETRY CHANGE IN THE FORLANDSUNDET AREA(NW SPITSBERGEN) USING REMOTE SENSING DATAMAREK GRZEŚ, MICHAŁ KRÓL AND IRENEUSZ SOBOTA
Department of Cryology and Polar Research, Institute of Geography, N. Copernicus
University, Poland

Fluctuations of glacier ranges are one of the basic elements in the study of theirgeometry changes. The studies of the fluctuation of the glaciers in the
Forlandsundet area (NW Spitsbergen) were made in three main regions: Kaffiøyra,St. Jonsfjorden and Prins Karls Forland. The investigations refer to mountain andcalving type glaciers.
The main investigative aim of this article is to estimate and interpret changes ofglacier ranges in the Forlandsundet area in the XX and XXI centuries by remotesensing data and field measurements.
We used ASTER ortho-images and a digital elevation model derived from 3N and3B bands to determine the fluctuations of the glaciers. The data analysis confirmedththat most of the glaciers in this area are in a recession phase (Figure 1). In the 21
century the rate of recession increased. The most quickly retreating are the calving
type glaciers - Aavatsmarkbreen, Konowbreen, Osbornebreen and Dahlbreen. The
front of the Aavatsmarkbreen retreated from 2000 to 2006 max. about 700 m (1002m/y). Between 2000 and 2006, the glacier area reduced by over 1.6 (km /y). The
glaciers terminating in the sea were characterized by a distinct retreat at theth
beginning of 21 century. The six small valley glaciers in Kaffiøyra in this period
retreated several m/year. Maximum glacier front recession in the period 2000-2006
we noted on the Osbornebreen (700-800 m/y). The retreat and mass loss noted on
almost all these glaciers have been observed on entire Spitsbergen. The recession(retreat and mass loss) is triggered by changes in climatic conditions. Due tovarious factors of the topoclimatic - geomorphologic nature, the recession has
taken a diverse course with respect to particular glaciers.
Moreover in this article we present the results of soundings profiles made during
the summer seasons 2004, 2005 and 2006 on chosen calving type glaciers. Basedon submarine relief evidence and marginal zones we made an attempt todetermine old zones of glaciers (at glacial episode 3.0-2.5 ka BP and Late Vistulian
13-10 ka BP).
One of the key issues while modelling subaqual glacial relief is the dynamics of theice cliffs. The research carried out by Brown et al. (1982), Boulton (1986) and Jania(1986) showed that ice cliffs can undergo considerable fluctuations in short periodsof time (such as a hydrological year or a balance year), irrespective of the generalrecessive tendency characteristic for most Spitsbergen glaciers. Moreover, theseoscillations are much larger than the rate of the multi-year recession. This meansthat during a balance year, mainly in May and June, a glacier advances as a result50

of ‘using’ the masses of snow delivered in winter; it also has minimal rate of ice cliff
calving. Winter advance may result in pushing glacial and fluvioglacial deposits up,
thus forming annual small push moraines. It often happens that the distance
between these moraines equals the mean annual recession. Accumulation oflarger push moraines, both subareal and subaqual ones, is predominantly
connected with intensive dynamics of movement and surge (Jania 1988).

Figure 1. The retreat of glaciers in Forlandsundet region (Spitsbergen). Numbers means the
maximum recession of a front of glaciers in the period 2000-2006.

Using the forms of the sea bottom and of the land marginal zones the author tried
to delimit the old ranges of the Aavatsmarkbreen and Dahlbreen. As the first results
show, the subaqual relief of the sea forefield of the selected glaciers was createdby pushing both glacial and marine material up by the anchored zone of a glacierwhich ends in the sea. In the presented profiles perpendicular to the ice cliffs forms
both seasonal and multi-annual oscillations are visible. Seasonal forms (Fig. 2) arefrom 3 to 10 m high; their genesis is probably connected with the winter re-advance51

of the ice front. Such a mechanism has already been described in the literature by
Boulton (1986) and Jania (1988). Marsz (1987) states a longer stagnation of an ice
cliff with an underwater bottom leads to the formation of a subaqual frontal moraine
at the foot of the cliff.

Figure 2. Annually push moraines in the forefield of Aavatsmarkbreen formed between 2004
and 2001 H- height of form, W- width of form, 2004- ice cliff range in 2004.

The oldest range of Dahlbreen, recorded in the sea bottom relief, is probably
connected with the glacial episode of 3.0-2.5 ka BP. It is, however, smaller than the
one delimited by Szupryczyński (1983). A trace of the Little Ice Age is clearlyvisible in the bottom relief. It corresponds with the deposits of that age left on theland (Grześ et al. 2007, Lankauf 2002). Morphogenetic analysis of the subaqual
forms of the sea bottoms where the studied glaciers end needs further research.
However, their sequence as well as the fact that they correspond with the oldranges of glaciers proves their glacial genesis. Echo sounders’ results of thesurging Aavatsmarkbreen are also interesting.

References
Boulton G.S., 1986. Push-moraines and glacier-contact fans in marine and terrestrial
environments, Sedimentology, 33, 677-698.
Brown C.S., Meier M.F., Post A. 1982. Calving speed of Alaska tidewater glaciers, withapplication to Columbia Glacier. U.S. Geological Survey Professional Paper 1044-9612,
C1-C13.
Grześ M., Król M., Sobota I. 2007. Subaqual recordings of the changes in the range ofglaciers in the Forlandsundet region (NW Spitsbergen). Landform Analysis, 5, 148-150.
Jania J. 1986. Dynamika czół spitsbergeńskich lodowców uchodzących do morza(Dynamics of the Spitsbergen tidewater glaciers fronts - Engl. summary), Geographia,
Studia et dissertationes, vol. 9, Uniwersytet Śląski, Katowice.
Jania J. 1988. Dynamiczne procesy glacjalne na południowym Spitsbergenie (w świetlebadań fotointerpretacyjnych i fotogrametrycznych), Prace Naukowe Uniwersytetu
Śląskiego w Katowicach Nr 955, Katowice. pp. 257.
Lankauf K.R. 2002. Recesja lodowców rejonu Kaffiøyry (Ziemia Oskara II-Spitsbergen) w
XX wieku, IG i PZ PAN, Warszawa, pp. 221.
Marsz A. 1987. Brzegi lodowe, Wydawnictwo Polskiej Akademii Nauk, Wrocław, pp. 117.
Szupryczyński J. 1983. Some problems of the Quaternary on Spitsbergen, [in:] Studies in
Quaternary Geomorphology, Geo Books, Cambridge, pp. 149-156.

52

RECENT CHANGES IN SURFACEAUSTFONNA ICE CAP, SVALBARD

MASSBALANCEOFTHE

JON OVE HAGEN1, TROND EIKEN1, THOMAS V. SCHULER1, THORBEN DUNSE1,KJETIL MELVOLD1, GEIR MOHOLDT1, ANDREA TAURISANO2, OLA BRANDT2 AND
JACK KOHLER2
1
2

Department of Geosciences, University of Oslo, Norway
Norwegian Polar Institute, Tromsø, Norway
2

Austfonna (79.5°N, 25°E) at 8120 km is by far the largest ice cap in Svalbard andone of the largest in the Arctic. The ice cap has a maximum elevation of about 800
m a.s.l., a thickness up to about 560 m (Dowdeswell and others, 1986) and a
relatively simple geometry, characterized by one main dome feeding a number ofdrainage basins. Austfonna is polythermal, and probably much of the bed is at thepressure melting point. Most of its boundary is calving and several of the outlet
glaciers are of surge-type (Dowdeswell, 1986; Hagen and others, 1993) (Figure 1).
For example, when an outlet from Austfonna, Bråsvellbreen, surged in 1936, thefront advanced c. 10 km out into the sea over a width of about 10 km.

Figure 1. Austfonna and the stake net marked with dots. To the right typical riming on thestakes is shown.

Austfonna is a target glacier for the IPY Glaciodyn project and also selected by
ESA as a CalVal site for CryoSat II. A number of investigations have been carriedout since 2004, these include GPS-elevation profiles, GPR measurements of thesnow thickness and snow distribution, operation of two Automatic Weather
Stations, shallow cores, snow pits and mass balance stakes. In spring 2004, 20
mass balance stakes were deployed across the ice cap (Figure 1). This network of
stakes has been maintained since then. The stakes have been remeasured once a
year, during the annual visits in springs 2005, 2006 and 2007. During the spring53

measurements the snow thickness at each stake is measured and thus we get theheight of the stakes above the previous summer surface. From thesemeasurements we derive the annual net surface mass balance at each stake
location. Riming is a serious problem for all installations on Austfonna, so every
year some of the stakes disappear probably because they break down due toheavy riming (Figure 1).
The dominant precipitation direction for Austfonna is from the east. This explains
the general pattern of snow accumulation across the ice cap, which shows apronounced gradient from high values in the SE to lower ones in the NW (Pinglot
and others, 2001; Taurisano and others, 2007). The distribution of snow thickness
is measured using extensive GPR-profiling together with manual snow thickness
probing, and the snow water equivalent is obtained from density measurements insnow pits. The winter snow accumulation varies from year to year by 100 %, but
the pattern of the snow distribution is fairly stable, often with about three times
more snow in the south-east than in the north-west (Taurisano and others, 2007).
The snow thickness is usually between 0.5 m up to 3 m (Figure 2).

Figure 2. Snow distribution and yearly variations across Austfonna.

The net mass balance of Austfonna has been calculated for the period 1986 to
1999 based on shallow core data (Pinglot and others 2001). They drilled altogether
29 shallow cores across the ice cap in 1998 and 1999. The depth down to theChernobyl fallout in 1986 was found in 19 cores in different elevations in theaccumulation area. Thus they could estimate the net mass balance gradients andcalculate the mean overall net surface mass balance to be very close to zero:
93
3
Bn = ∑ bni (z) • Ai [10 m ], Bn ~ 0 km ,Surface elevation measurements both by airborne laser profiling in 1996 and 2002
(Bamber and others 2004) and using our own ground-based GPS profiles show aclear thickening in the interior and a thinning along the lower parts of the ice cap.
Bamber and others (2004) suggest a recent increase in precipitation being
responsible for this. Bevan and others (2007) have suggested that the massbalance over this period was positive, based on the laser data, the shallow coredata and satellite radar interferometry for flux calculations.. Bevan and others(2007) stressed that the result emphasizes the importance of considering the
dynamics of glaciers when attempting to draw any conclusions on climate change
based on snapshot observations of the cryosphere.
54

We calculated the average net mass balance in two large parts of the ice cap,
Etonbreen and South-East area as shown in Figure 1. In each of these areas thenet balance gradient was plotted and assuming that the stakes are representativefor each elevation band, the basin net balances and the overall total net balance22were calculated (Figure 3). The two areas cover ca. 4670 km (of 8120 km ). The
mean total surface net balance for the whole of Austfonna is then given by:
93
3
Bn = ∑ bni (z) • Ai [10 m ] yielding Bn ~ - 2 km ,and the mean specific net balance:
bn = Bn / A [m] = – 0,25 m ± 0,1 m.

Figure 3. Average mass balance gradients for the period 2003 – 2006 plotted as apolynomial curve at Etonbreen basin and as a linear curve on the South-East area.

Schuler and others (2007) used a degree-day approach to model mass balance forthe same period as the measured surface balance and for the year 2004 theycame up with a negative balance: bn 2004 = - 0,32 m w.eq. The model agrees well
to a number of observations including specific balance at stakes, snow
accumulation from GPR, snow distribution at the end of summer as derived fromsatellite imagery and sonic ranger records of ice melt.
New shallow cores drilled in 2004 and 2006 at three of the former locations in theupper accumulation area of the ice cap indicated no recent increase in the
precipitation since 1999. So we stress again the importance of understanding the
dynamics of an ice mass when interpreting observed geometry changes (Hagen
and others 2005, Bevan and others 2007).
Compared to the previous results by Pinglot and others (2001), the surface net
mass balance is clearly negative after 2003, while it was close to zero for theperiod 1986 – 1999. However, we have only three years of data and these also
include the very warm summer of 2004 were melting was observed high up on theice cap.

References
Bamber, J.L., Krabill, W.B., Raper, V. and Dowdeswell, J.A., 2004. Anomalous growth ofpart of a large Arctic ice cap: Austfonna, Svalbard. Geophysical Research Letters, 31,
L12402, doi.10.1029/2004Gl019667.

55

Bevan, S., Luckman, A., Murray, T., Sykes, H. and Kohler, J., 2007. Positive mass balance
during the late 20th century on Austfonna, Svalbard, revealed using satellite radar
interferometry. Annals of Glaciology, 46, 117-122.
Dowdeswell, J.A., 1986. Drainage-basin characteristics of Nordaustlandet ice caps,
Svalbard. Journal of Glaciology, 32, 31-38.
Dowdeswell, J.A., Drewry, D.J., Cooper, A.P.R, Gorman, M.R., Liestøl, O. and Orheim, O.,1986. Digital mapping of the Nordaustlandet ice caps from airborne geophysicalinvestigations. Annals of Glaciology, 8, 51-58.
Hagen, J.O., Liestøl, O., Roland, E., Jørgensen, T., 1993. Glacier Atlas of Svalbard and Jan
Mayen. Norsk Polarinstitutt, 141 pp.
Hagen, J.O., T. Eiken, J. Kohler and K. Melvold, 2005. Geometry changes on Svalbard
glaciers: mass-balance or dynamic response? Annals of Glaciology 42, 255-261.
Pinglot, J.F., Hagen, J.O. Melvold, K. Eiken, T. and Vincent, C., 2001. A mean net
accumulation pattern derived from radioactive layers and radar soundings on Austfonna,
Nordaustlandet, Svalbard. Journal of Glaciology, 47, 555-566.
Schuler, Thomas; Loe, Even; Taurisano, Andrea; Eiken, Trond; Hagen, Jon Ove; Kohler,Jack. 2007. Calibrating a surface mass balance model for the Austfonna ice cap,
Svalbard. Annals of Glaciology, 46, 241-248.
Taurisano, A., Schuler, T.V., Hagen, J.O., Eiken, T., Loe, E., Melvold, K. and Kohler, J.,
2007. The distribution of snow accumulation across the Austfonna ice cap, Svalbard:
direct measurements and modelling. Polar Research, 26, 7-13.

56

THE ROLE OF BASAL FRACTURES IN STORING AND ROUTINGWATERJOEL HARPER1, JOHN BRADFORD2, NEIL HUMPHREY3, TAD PFEFFER4 AND TOBYMEIERBACHTOL1
1

Department of Geosciences, University of Montana
Center for the Geophysical Investigation of the Shallow Subsurface, Boise State University
3
Department of Geology and Geophysics, University of Wyoming
4
Institute of Arctic and Alpine Research, University of Colorado
2

Introduction
Interest has grown in the role of the glacier hydrologic system in ice sliding due torecent observations of increased surface melt of Arctic Glaciers and the margins of
the Greenland Ice Sheet (e.g., Mote, 2007). Fractures in glaciers have received
growing attention as an important mechanism for routing water from the surface tothe bed of glaciers. For example, Boon and Sharp (2003) documented thedrainage of a supraglacial lake to the bed via a newly opened fracture, andobservations at Storglaciären were used to postulate that englacial fractures are
the primary pathway for routing surface water to the bed (Fountain et al., 2005).
Here we investigate the role of englacial fractures in storing basal water.

Methods
Field experiments were conducted at Bench Glacier, Chugach Mountains, Alaska.
Bench Glacier is a temperate valley glacier ~8km long with a straight and simple
geometry. The ice is 180-200 m thick along most of the ablation zone. We
conducted experiments in more than 50 boreholes drilled to the bed of the glacierwith a hot water drill. Borehole video observations were used to quantify theorientation and distribution of englacial fractures. Dye tracing experiments using apair of down-hole fluorometers were used to document pathways and rates ofwater flow through the fractures. Radar surveys using 3D and multi-offset dataacquisition schemes imaged the spatial extent and orientation of cracks within theice mass. Lastly, water level monitoring in a network of boreholes during the drillingof new boreholes quantified englacial water connections and head gradients.

Results and Conclusions
Our multi-faceted observations can be used to draw three main conclusions about
fractures at Bench Glacier.
First, fractures are common and are concentrated in lower two thirds of the ice
mass. Borehole video observations documented open fractures in every boreholewith an average of 2 per hole (Figure 1). The fractures are always steeply dipping,and range in size from a few centimeters up to about 1 m across (Figure 2).
Measurements of fracture orientation show their orientation rotates at depth
following the stress field. Radar profiling shows that scattering events (from water
filled macro scale voids) are concentrated in the lowest two thirds of the ice mass,and distinctly lacking in the 20-50 m of ice below the surface.
57

0
20
40
60
80
100
120
140
160
180
0

1

2
3
4
Number of occurrences

5

Figure 1. Histogram showing open fractures loggedwith a borehole video camera lowered through 16
boreholes drilled in 2003. Logging was terminatedabove the bed due to turbid water; red triangles show
depth where logging was terminated for various holes.
The histogram is biased toward shallow depths since
the lowest 10-20% of the boreholes were not logged.

Second, many fractures are connected to bed. Drilling experiments showed thatboreholes frequently drain or fill during the drilling process, suggesting water drainsto the bed via englacial connections (Figure 3). Video observations show open
fractures coincident with the location of the drill tip at the time of a draining event.
Also, dye tracing experiments revealed water in boreholes was preferentially routed
down fractures intersecting the borehole, rather than past the fracture and down
the borehole.
116 mFigure 2. Photograph from borehole video camera
showing typical open fracture. Fracture depth is
116 m below the ice surface. Vertical bright line isa signal cable which can be seen entering theborehole at the base of the fracture.

Third, some fractures open upward from the bed. We collected two separate time
lapse video sequences that show fractures opening with turbid water and upward
water flow. In 2002, for example, filming of a borehole just after drilling revealed no
open fractures in the lower half of the glacier. Eleven days later we filmed the holeand observed vigorously upwelling turbid water and frazzle ice formation at a depthof 139 m (35 m above the bed). Over a period of several days the turbidity settled
and a new ~15 cm diameter fracture was visible. In fact, the fracture had displaced
the borehole with respect to gravity and the camera could no longer be lowered
into the borehole and would only follow the open fracture when lowered. Water was
clearly flowing up the borehole from a lower depth and then back down thefracture. A similar observation of fracture opening was also made in 2006.
Fracture volume and basal water storage.
We estimated fracture volume from geophysical inversion of radar data (Bradfordand Harper, 2005) and from direct observation of fractures via borehole video.
Based on our sampling of fracture volume, we estimate that the fractures
maintaining an open connection to the bed occupy 0.5% to 3% of the ice mass by58

volume in the lowest 60 m of ice. The storage capabilities of this fracture spaceequates to a layer of water between 3 cm and 18 cm thick covering the bed.
Fractures may therefore be an important element of subglacial drainage systemdevelopment due to their ability to store and release water. Further, they likely play
an important role in the interplay between water and sliding dynamics.

Figure 3. Schematic showing setup of borehole drilling experiments (lower graphic) and
response curves during a draining event (upper plots). Water levels were monitored in up to
six boreholes while a new borehole was being drilled. Dashed line shows approximate
piezometric surface defined by borehole intersections with the bed.Top left plot shows
draining of high head in the drill hole (which is artificially high due to the addition of drilling
water) at 100 m below the surface and 60 m above the bed. Upper right three plots show
simultaneous rise in water levels in boreholes when the drill hole drained. A rise evenoccurred in a borehole which terminated 20 m above bed (far upper right plot).

References
Boon, S., and M. Sharp (2003), The role of hydrologically-driven ice fracture in drainage
system evolution on an Arctic glacier, Geophysical Research Letters, 30(18).
Bradford, J. H., and J. T. Harper (2005), Wavefield Migration as a Tool for Estimating
Spatially Continuous Radar Velocity and Water Content in Glaciers, GeophysicalResearch Letters, 32(L08502), doi:10.1029/2004GL021770.
Fountain, A. G., et al. (2005), Fractures as the main pathways of water flow in temperateglaciers, Nature, 433(7026), 618-621.
Mote, T. L. (2007), Greenland surface melt trends 1973-2007: Evidence of a large increasein 2007, Geophysical Research Letters, 34.

59

MELTWATERRUNOFF AND FIRN DENSIFICATIONPERCOLATION FACIES, GREENLAND ICE SHEET

IN

THE

JOEL HARPER1, NEIL HUMPHREY2, TAD PFEFFER3, JOEL BROWN4, DAVIDSCHULER1 AND DAN STURGIS2
1

Department of Geosciences, University of Montana
Department of Geology and Geophysics, University of Wyoming
3
Institute of Arctic and Alpine Research, University of Colorado
4
Center for the Geophysical Investigation of the Shallow Subsurface, Boise State University
2

We present results from a field campaign focused on meltwater infiltration andhorizontal water transport processes in the percolation zone of the Greenland Ice
Sheet. Field data were collected during a period of heavy melt in June/July 2007
along a ~50 km transect (from ~2000 m to ~1600 m elevation) of the EGIG line ofwest Greenland (Figure 1).

Figure 1. Image showing location of 2007 study transect (Crawford Point to T1) and planned
activities during April/May of 2008 (T1 to Swiss Camp). Study reach follows the EGIG line.

Snow and firn stratigraphies of the upper 10 m were documented with snowpit
measurements, analysis of 21 firn cores drilled to 10+ m, and with over 60 km ofconstant offset radar profiles collected at a variety of frequencies. We also acquired
15 constant midpoint radar profiles to characterize depth-density relationships in
the upper 80 meters of the firn column. Dye tracing experiments were used to
identify meltwater migration pathways and to quantify the relative rates ofhorizontal and vertical water movement. Five thermister strings with 33 channels
and a 30 min time base were installed for long term monitoring of the thermal
signature of meltwater migration and ice layer formation in the upper 10 m of firn
(Figure 2). Two meteorological stations were installed to provide information on
surface boundary conditions.
60

Figure 2. Temperature profiles from six 10 meter holes between CP and T1 (see Figure 1).
Also shown are four temperature profiles collected in 1955 by C. Benson.

Our work shows massive ice layers (up to 0.4 m thick) form at depth underconditions of heavy surface melt. The ice layers, however, lack spatial coherence
over meter length scales and therefore allow vertical meltwater infiltration. Hence,
we found no evidence of significant horizontal water transport along internal ice
layers within this elevation band of the GIS percolation zone. We also found thatdensification rates to 80 m depth show significant elevation dependency, and a
recent increase in surface densification rate at lower elevations.

61

GLACIODYN ACTIVITIES ON AND UNDER ENGABREENMIRIAM JACKSON1 AND TORBORG HAUG2
1Norwegian Water Resources and Energy Directorate, Oslo, Norway.
Institute of Mathematical Sciences and Technology, University of Life Sciences, Ãs,
Norway2Engabreen in northern Norway is an outlet glacier of the Svartisen icecap (Figure1). There is a set of mass balance observations for Engabreen going back nearlyfour decades. These are performed in connection with a hydropower station in thearea that uses water from the glacier that is collected as run-off as well as being
collected directly through subglacial intakes. The mass balance was strongly
positive in the 1990s and this was reflected in a strong advance of the glacier front.
However, mass balance has been mainly negative since 2000, and the glacier front
has retreated 200 m in the last eight years.
Engabreen is a unique glacier in the existence of a subglacial laboratory
underneath the glacier. This laboratory was excavated for the purpose of
glaciological research and has access from the tunnels created as part of thehydropower network.

Helgelandsbukken

7399000
Engabrevatnet

Bautaen (GPS base)

North (m)7397000

7395000

7393000Skjæret7391000

Snøtind

Stake positions

7389000
444000446000448000450000452000454000

East (m)Engabreen

SVARTISEN

0

20000

40000

60000

Figure 1. Location map for
Engabreen, an outlet glacier of
Svartisen.

80000

Several load cells are installed at the glacier bed and measure the total hydrostaticpressure normal to the load cell. The pressure is measured at 15 minute intervals
and shows variations in pressure at the bed due to variations in surface meltwateror changes in the subglacial discharge. Lappegard et al (2006) presented asummary of pressure variations and related them to the hydrological regime at thebase of the glacier.
62

The dynamics of Engabreen have also been studied, including variations in surface
velocity both seasonally (Jackson et al, 2005) and over shorter time scales. As partof the Glaciodyn IPY project, several parameters will be studied simultaneously, inorder to see how different parameters are related to one another, and to better
understand glacier dynamics. In May 2007 a preliminary investigation was madestudying the relation between surface motion and subglacial pressure. Nine stakes
were placed on the glacier tongue in an approximately square pattern at a few
hundred metres spacing (Figure 2). It was not possible to place the stakes in alonger profile across the glacier due to severe crevassin on both sides; the central
part is relatively flat and crevasse-free. Stake positions were measured over a 21day period from 8th May to 29th May 2007, with measurements on days 1 to 4, anddays 16 to 21. Some stakes were not measured on days 1 and 2. Velocities werean average of 40 cm per day over the whole period.

Figure 2. Movement stakes on Engabreen
tongue. Glacier motion is towards the top of
the page, and elevation on the glacierdecreases towards the top of the page.

Stake elevations were also measured and it is interesting to look at changes inthese. The general trend is to lower elevations as the glacier flows downhill.
However, there appears to have been an uplift event that is seen on the westernside of the glacier. It is recorded first in the uppermost stake (SW) on the west sideon 25th May, next in the middle stake (MW) on 26th May and finally in thelowermost stake (NW) on 27th May (Figure 3). It may also be detected in the stake
SC on 27th May.

Figure 3. Elevation of stakes over the 21-day period.

63

Uplift events have been detected on several glaciers (e.g. Iken et al, 1983),
especially in connection with the onset of spring melting. Records at the subglacialpressure cells, that lie about 1.5 km upstream of the measurements stakes, showthat there was a pressure event on about 23rd May (Figure 4). There were also
several stronger events in the preceding week, but the stakes were not measured
in this period. These suggest an increase in the available meltwater at the bed thatwas registered on all the pressure sensors (number 4, 6, 2a, 1e, 97-2 and 97-1 andsituated within 20 m of each other). This then travelled downstream and was
registered predominantly on the side of the glacier that is thickest (due to an
asymmetrical basal topography) and where meltwater presumably drains
predominantly before coming out at the pro-glacial river that is also on the westernside of the glacier (see Figure 1).

Figure 4. Pressure variations over the measurement period.

References
Iken A., H. Röthlisberger, A. Flotron and W. Haeberli. 1983. The uplift of Unteraargletscher
at the beginning of the melt season – a consequence of water storage at the bed?
Journal of Glaciology, 29 (101): 28-47.
Jackson, M., I. Brown and H.Elvehøy, 2005. Velocity measurements on Engabreen. Annals
of Glaciology, 42, 29-34.
Lappegard, G., J. Kohler, M. Jackson and J.O. Hagen. 2006. Characteristics of subglacial
drainage systems deduced from load-cell measurements. Journal of Glaciology 52 (176),
pp. 137-148.

64

NATURAL RADIONUCLIDES IN GLACIER MELTWATER AS USEFULTRACERS
A. KIES1, A. NAWROT2, H. SURBECK3 AND Z. TOSHEVA1
1

Laboratoire Physique des Radiations, University of Luxembourg
Institute of Paleogeography and Geoecology, Adam Mickiewicz University
3
Centre of Hydrogeology, University of Neuchâtel
2

We present an alternative method to investigate the character of the meltwatersupplied by glacierized basins in introducing radioactive ionic measurements
together with classical parameters like electrical conductivity, pH, dissolved gases.
Among the natural radioactive ions the most promising is the nobel gas radon witha short half life of 3.8 days and the possibility of an automated continuous
222
measurement using sensors that can be controlled by data loggers. Radon ( Rn)226
is a radioactive noble gas (T=3.8 d), it originates from the decay of radium ( Ra),238
member of the natural U series. In solids, only the radon produced very close to
the grain surface, within less than a fraction of a micrometer, will be able to reachthe pore space and get dissolved in water.
Glaciers are active erosional agents, continuously crushing and abrading thesubstrate over which they flow resulting in an effective high exchange surface
between substrate and water. Different ionic species are acquired from contact
during water flow through the glacier drainage system and contribute to theincreased electrical conductivity and, in the case of radon, to an uptake of radon
escaping from bed rock and sediments. Such contact is achieved most effectively
at the glacier bed. In meltwater, temporal variations of radon and ionicconcentration may be interpreted in terms of the water flow pathways. Highconductivity, high ionic stable and radioactive element concentrations indicate
contact with reactive sediments and rocks, possibly for extended periods of time.
On the other hand high radon concentrations simply indicate contact with radium
located close to the surface of sediments/rocks, this contact happening less than
20 days before sampling. In both cases the period of contact is important. In thecase of radon a steady equilibrium concentration can be expected whenever
uptake of radon is counterbalanced by decay. The obtained ionic signatures are
generally interpreted in terms of flow through the subglacial drainage system, via
hydraulically inefficient pathways with slow transit speeds and involving muchcontact with large amounts of freshly eroded debris. In contrast, low ionic/radon
concentrations are consistent with no, or only a short period of, flow at the glacierbed, via a hydraulically effective drainage pathway characterized by large water
fluxes and rapid transit speeds.
Radon concentrations depend on the radium concentrations of the rocks, therock/sediment- water contact surface and the contact duration. Meltwatercontaining radon is a secure information that part of this water has been recently incontact with the bedrock, the more radon the higher is the fraction of meltwater incontact with bedrock, the fraction without contact being nearly radon free.
65For our investigation we chose the well studied Werenskioldbreen situated in thesouth-western part of Wedel-Jarlsberg Land, Spitsbergen. Its closeness to thePolish Polar Station in Hornsund has made the glaciated catchment well studied interms of climatology, geomorphology and geology (Bukowska, 2007).
2
Werenskioldbreen is a polythermal glacier with terminus on land, a 27 km area
and a 100-140 m maximum thickness. The terminus is divided into two parts by acentral moraine, one third to the north and two third southern part. Four areas were
sampled between September 2006 and end of July 2007: 2 areas in the northernpart called N-a and N-b, N-a being close to the lateral moraine, W is an importantsummer outflow to the south of the central moraine; N is an area close to thesouthern end of the moraine.
Werenskioldbreen provides evidence for considerable basal routing of water. Highradon concentrations and conductivity of the subglacial upwelling provides
evidence of the passage of icemelt through a subglacial weathering environment
characterized by high rock-sediment:water ratios, prolonged residence times andrestricted access to the atmosphere, preventing radon to degas. At higher
discharges, basal bulk runoff becomes dominated by icemelt from the lower partsof the glacier that is conveyed through a weathering environment, characterized bylow rock: water ratios, short residence times and free contact with atmosphericgases, to the glacier margins. In Werenskioldbreen icemelt is routed via ahydrological system composed of englacial and supraglacial components to theglacier margins by the ice surface slope. Before upwelling, water flows relatively
independently of icemelt to the terminus via a subglacial drainage system, possibly
constituting flow through a sediment layer. Cold basal ice at the terminus forces it
to take a subterranean routing in the latter stages. The existence of spatiallydiscrete flow paths conveying icemelt and subglacial snowmelt to the terminus is
the norm for polythermal-based glaciers on Svalbard (Wadham, 1998).
1098radon (Bq/l)

7
6
5

N-a

4

N-b

3

W

2

S

1
0
N-b N-bW

S

N-a N-b

W

S

N-a

W

N-a

W

S

N-a

W

N-a

W

S

W

S

W

S

13/5 13/5 13/5 13/5 26/5 26/5 26/5 26/5 02/6 02/6 05/7 05/7 05/7 08/7 08/7 17/7 17/7 17/7 18/7 18/7 23/7 23/7

Figure 1. Meltwater radon concentrations measured at 4 different sites in May-July 2007.

66

In September and April air temperatures were below zero, very high radon
concentrations were measured. Due to the lack of surface melting during these
periods one may assume that steady state radon concentrations are approached.
Effectively highest radon concentrations were measured in September (onesample, 22 Bq/l, el. conductivity) at location W, 16 to 28 Bq/L at varying water
outflows in April in N-a, N-b and S (50 samples collected over 3 weeks), (Figure 1).
The sampling dates at Werenskiold glacier covered the different periods of theglacial cycle: end of September 2006; in 2007 April, the start of ablation period, in
May-June and part of the ablation period until end of July. Pulina (1990) refers that,based on the glacial river of the Werenskioldbreen, ablation water in summer is as
much as 80–95% of the total meltwater flow, but, beginning in September, theshare of subglacial sources increases, possibly representing the entire flow by theend of that month.
In September radon and conductivity measurements show highest values in spring
upwelling waters at location W (radon concentration 22 Bq/L) reflecting the
subglacial origin.
Table 1. Natural radioactive isotopic content in meltwater sampled at the Werenskiold
glacier front in April 2007.
+/Date Location
U-238 +/- U-234 +/- Ra-226 +/- Po-210 +/Rn
2007
mBq/l
mBq/l
mBq/l
mBq/l
Bq/l
0.2
0.3
1.1
10.0
2.4
9.04 N-b, 'old' hole,
6.7
5.1
9.9
90
27.4
opened
1.2
0.3
8.1
66.5
1.2
9.04 N-b, new dome
28.4
22.9
50.8
466
11.0
outflow
0.1
0.3
0.8
2.5
1.6
10.04 N-b, outflow
3.6
2.8
5.2
17.8
16.2
small dome
0.1
0.1
0.05
2.5
10.04 N lat. moraine
0.4
0.7
0.3
15.0
0.10 0.02
0.1
0.2
0.6
1.0
1.2
10.04 N-a flowing
4.2
5.5
7.2
8.1
12.0
0.1
0.1
0.2
3.4
1.0
13.04 S outflow small
1.4
1.8
1.1
27.6
10.4
dome
0.1
0.1
0.3
3.2
1.2
21.04 S outflow small
4.1
4.2
1.9
7.1
12.0
dome
0.1
0.1
0.3
0.8
21.04 N-a small dome
5.0
4.2
2.1
2.6
0.86 0.14
running
0.2
6.0
0.2
6.7
1.6
21.04 N-b outflow big
7.2
7.4
1.3
37.6
16.2
dome

From 5th to 21st of April 2007, on six different days, 50 flowing water samples werecollected. Over the 2.5 km glacier front 5 naled ice fields of varying area with ice
domes could be observed, in three of them, N-a, N-b and S, running water outflow
close to the glacier could be accessed and sampled. Flowing water was normallyon or close to ice domes, some of them 2 m high but generally there heights were
below one meter. Often ice had to be broken to access running water, sometimes
under pressure. A particular outflow never lasted more than 3 days before
refreezing. It was not possible to take samples at area W sampled in September.
Table 1 gives an overview of the natural isotopic content of some samples taken to67the home laboratory for radiochemical analyses; radon was measured by liquid
scintillation at the Polish Base; electrical conductivity, temperature and sometimes
pH in situ.
In area N-b highest April radon (12-16 Bq/L) and conductivity values (specific el.
conductivity 400-500 µS/cm) were measured. One outflow, lasting for 3 days,
showed a very high electrical conductivity (2700 µS/cm) and the highest radon
concentrations so far measured at Werenskioldbreen (27 Bq/L). In N-a, radon
concentrations were lower (maximum 11 Bq/L) and very variable, sometimes down
to 0.1 Bq/L, specific electrical conductivities varied around 400 µS/cm. In area S
electrical conductivity was 2.5 times lower than in N-b, whereas with 12 Bq/L radon
concentrations were high.
A general trend was that high radon concentrations are often linked to highelectrical conductivity.
In irregular intervals sampling was done from May to July. Figure 1 resumes themeasured radon concentrations. At location W, at the maximum of upwelling water
outflow, highest radon concentrations (8.2 Bq/L) were measured at the end of July.
If one assumes a simple mixture model of subglacial and radon depleted endosupraglacier water, 37% could be allocated to subglacial water.

Conclusion
At Werenskioldbreen unexpected high radon concentrations were measured.
Combined with traditional investigation tools, radon proves to be a valuable tracer
for the study of artesian glacier meltwater.
Highest radon concentrations were measured in September and April, in periods ofno supraglacier melting, lowest in period of the beginning of thawing, intermediateconcentrations in the periods of maximum thawing with high combined sub-, endoand supra- glacier circulation.
We report here of preliminary investigations with radon and other naturalradionuclides as natural tracers. We have to learn more how to use and read theobtained information

Acknowledgements
Thanks are due to the Polish Polar Station and to friends and colleagues forassistance in the field research.

References
Bukowska-Jania E., 2007. The role of glacier system in migration of calcium carbonate on
Svalbard, Polish Polar Research, vol. 28, no. 2, pp. 137–155.
Pulina M., 1990. Geomorphological effects of the cryochemical processes. Questiones
Geographicae 13/14 (1987/1988), Poznań: 99–112.
Wadham J. L., Tranter M., Dowdeswell J. A., 1998, The hydrochemistry of meltwaters
draining a polythermal-based, high Arctic glacier, south Svalbard: I. The ablation season,
Hydrological Processes, vol.12, no.12, pp 1825-1849.

68

MODIS ALBEDO AND REGIONAL MASS BALANCE OF SVALBARD
GLACIERSJACK KOHLER
Norwegian Polar Institute, Tromsø, Norway

Abstract
MODIS albedo products are compared to glacier mass balance measured on
several glaciers in the high Arctic, in northwestern Svalbard. The glaciers range in2
size from ca. 5-500 km . We use MODIS L3 albedo products (MOD43B3), whichhave a nominal resolution of 1 km, and for which data cover the spring to autumn
months of 2000-2006. We compare the albedo data to four glaciers for which there
are mass balance measurements in each of the years in the study period (Austre
Brøggerbreen, Midtre Lovénbreen, Kongsvegen, and Waldemarbreen), as well as
two glaciers with some years of mass balance measurements (Holtedahlfonna,
Etonbreen). There is still good overall correlation between parameters derived fromannual albedo minima and the equilibrium line altitude (ELA) of the study glaciers.
Simple threshold methods can potentially be used to directly extract an ELA proxyfrom the albedo data.

69

CALVING DYNAMICS OF JAKOBSHAVN ISBRAEMARTIN LÜHTI
Laboratory of Hydraulics, Hydrology and Glaciology, Swiss Federal Institute of Technology(ETH), Zurich, Switzerland

Abstract
Jakobshavn Isbrae, a large outlet glacier of the Greenland ice sheet, has changed
dramatically in the last ten years. Fast surface lowering, calving front retreat anddoubled flow velocity are hallmark changes, similar to those observed at otherGreenland ice streams. To understand the interaction between processes that leadto the fast changes in geometry and dynamics, and to answer the question whether
ice streams react to changes in ocean temperatures or to increased ablation andmelt water production on the ice sheet, we measured flow velocity at high time
resolution along a central flow line from 50 km inland to the terminus. Geodetic
GPS were installed where landing a helicopter was possible, and conventional
survey reflectors tracked by an automatic total station were deployed in the difficult
to access region near the terminus. Tidal forcing has no significant effect on flow
velocities in the terminus area, as was observed twenty years ago. The glacierterminus reacts with a spatially and temporally consistent acceleration to the 1 km
retreat of the terminus during each big calving event, when tens of ice bergs oftypically 500m width, 500 m length and 900 depth break off. The bergs areaccelerated to more than 10 m/s by the energy released by rotation, inducing
seismic events that could be detected at a nearby seismometer.

70

MASS BALANCE OF THE PRINCE OF WALES ICE-FIELD,
ELLESMERE ISLAND, NUNAVUT, CANADADOUGLAS MAIR1, DAVID BURGESS2, MARTIN SHARP3, SHAWN MARSHALL4,FIONA CAWKWELL5, JULIAN DOWDESWELL6 AND TOBY BENHAM6
1

Department of Geography and Environment, School of Geosciences, University ofAberdeen, UK
2Canadian Centre for Remote Sensing, Natural Resources Canada, Ottawa, Canada
3
Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Canada
4
Department of Geography, University of Calgary, Calgary, Canada
5
Department of Geography, University College Cork, Republic of Ireland
6
Scott Polar Research Institute, University of Cambridge, Cambridge, UK

Extended abstract
The mass balance of the Prince of Wales Icefield (POW), Ellesmere Island,
Canada is determined from measurements of surface mass balance and iceberg
calving. Despite being the second largest ice mass of the Canadian High Arctic
there have until recently been only sporadic glaciological measurements carriedout across POW (Koerner, 1977 and 1979) as is not one of the ice caps regularly
measured as part of the Geological Survey of Canada’s long-running massbalance monitoring program (Koerner, 2005). This paper presents the first massbalance measurements from POW in over 25 years.
Alt. 1052 m(core9)
Alt. 1419 m(core5)

137

Cs counts (1200 secs)

300
250
200
150
100
50
0
0

5

10

15

20

25

Depth below surface (m )Figure 1.

137

Cs against depth for two sites showing clear bomb layer peak at depth.

In April-May 2001 and 2002, thirteen shallow cores were drilled across theaccumulation areas of POW, ranging in elevation from 941 m to 1730 m above sea
level. Using a down-borehole gamma spectrometer, gamma activity profiles ofboreholes created by removal of each shallow core was measured. The count137
profiles for
Cs showed a secondary peak at depth (e.g. Figure 1) that is
associated with peak fall-out from atmospheric testing of thermonuclear weapons
in 1962-63. Using the depths of the 1963 radioactive fallout layers andmeasurements of the densities of firn and snow above this layer, the average massbalance, bn, over the period 1963-2001/2002 was calculated at each core site in
the field.
71

From May 2002 to May 2003, measurements were made of changes in the heights
of winter snow and previous summer firn or ice surfaces against a network of massbalance stakes deployed across two transects of the ice-field. In conjunction withmeasurements of snow-pit densities, near-surface firn core densities, andassuming a glacier ice density of 900 kg m-3 across ablation areas, these stakemeasurements were used to calculate bn for the year 2002-2003 at 24 locations
across the ice-field.
Shallow ice core net accumulation measurements and annual mass balance stakemeasurements are used in conjunction with a digital elevation model andknowledge of the location of the dominant moisture source to interpolate andextrapolate the spatial pattern of surface mass balance (SMB) across the Prince ofWales icefield. Ten different approaches yield different estimates of the contributionof SMB to the overall mass balance of the icefield, thereby highlighting andquantifying the methodological uncertainty of such estimates.
The contribution of iceberg calving to the mass balance is calculated fromestimates of (a) the volume of ice flux discharged at the major tidewater glaciertermini per annum and (b) the volume loss or gain due to the observed change oftermini position per annum. The former is calculated from airborne radio echosounding (RES) measurements of ice thickness at the glacier termini and
RADARSAT-1 speckle tracking surface velocity measurements. The latter is basedon comparison of 1959/1960 aerial photography with 1999/2000 satellite imageryof marginal advance or retreat of major glacier basins. The results of this are
summarized in Table 1.
Table 1. Iceberg calving totals, QTotal, and components Qflux and Qv-loss by glacier basin,
where Qflux is the annual volume of ice flux discharged at the tidewater termini and Qv-loss isthe volume loss due to the observed change of terminus position
Glacier basin
Qflux
Qv-loss
QTotal Calving total
3
-1
3
-1
3
-1
(km w.e. a )(km w.e. a )(km w.e. a )Leffert0.01 ± 0.002
0.01 ± 0.002
Cadogan
0.08 ± 0.013
0.02 ± 0.003
0.10 ± 0.016
Eckblaw
0.17 ± 0.043-0.02 ± 0.003
0.15 ± 0.046(terminal advance)
Trinity & Wykeham
1.36 ± 0.221
0.21 ± 0.032
1.57 ± 0.253
SE Margin
0.04 ± 0.006
0.04 ± 0.006
Totals
1.61 ± 0.28
0.26 ± 0.05
1.87 ± 0.32

We conclude that the surface mass balance of the icefield is positive (+1.4 ± 0.7
3-1
km w.e. a ) largely due to high accumulation in south-east catchments facing themain year round moisture source of the Smith Sound portion of the North Open3
Water Polynya. Iceberg calving is a major component of mass loss (1.9 ± 0.3 km-1w.e. a ) and is sufficient to cause the overall mass balance of the icefield to be3-1
most probably negative (-0.5 ± 0.8 km w.e. a ). The Prince of Wales Icefield
therefore has a lower contribution to global eustatic sea level rise thanneighbouring ice masses in the Canadian High Arctic which is consistent withprevious observations showing just a slight reduction in surface area of the icefieldover the period 1959-2000 (-0.7%).
72

73

ABLATION AND VERTICAL GRADIENTS OF AIR TEMPERATURE – A
STUDY FROM HANS GLACIER, SW SPITSBERGEN
K. MIGAŁA1, B. LUKS2, D. PUCZKO2, S. SIKORA1, A. DRZENIECKA–OSIADACZ1,
M. GRABIEC3 AND P. GŁOWACKI2
1

University of Wrocław, Institute of Geography and Regional Development
Institute of Geophysics Polish Academy of Sciences,
3
University of Silesia, Faculty of Earth Sciences
2

Introduction
Glaciers are a derivative of climate and dynamics of glaciological processes
depending on unremitting atmospheric processes in the 3D space. Gradients of: airtemperature, moisture, solar incoming energy, precipitation and chemicalcomposition are the important attributes in the vertical structure of the atmosphere.
Vertical differentiation of atmospheric processes has an impact on stratification of
glacial phenomena. The atmospheric boundary layer (ABL) should be consideredin this case. The depth of the ABL and changes of air temperature in the verticalprofile mark the border in dynamics of the processes concerned with thedevelopment of stratus/stratocumulus clouds and the surface radiation budget(Curry 1985, Bing Lin and Minnis 2003). In consequence, the structure of the ABL
modifies the chemistry of precipitation and dynamics of other processes e.g.
freeze/thaw cycles. Numerous investigations confirm that the ABL in the Arctic isvery shallow, often not exceeding 200 m (Tjernström et al. 2004, Drzeniecka et al.
2007). Knowledge of local climate, orography and glaciers morphology lead to theconclusion that one can expect hypsometric gradients in glaciological processes,
shaped by the Arctic ABL.

o

air temperature [ C]

15

10

5

0
1-07

11-07

21-07

31-07

10-08

20-08

30-08

July - August 2005
PSP 15m

Hs 201m

Hs 320m

Arie 363m

Figure 1. Plot of daily mean air temperature at the Polish Polar Station (PPS), Hans Glacier(Hs) and mountain ridge (Arie), July – August 2005.

74

Table 1. Mean monthly air temperature (MMAT) and sums of positive daily air temperature(PDD) in the period July (VII) - August (VIII) 2005 at the Polish Polar Station (PPS), Hans
Glacier (Hs) and mountain ridge (Arie).
UnglaciatedGlacier
Unglaciated
PPS 15 m
Hs 200 m Hs 320 m Arie 363 m
MMAT (VII)MMAT (VIII)PDD (VII-VIII)

4,3
4,6
275,9

2,4
2,8
162,2

4,2
3,0
221,4

4,6
3,4
247,1

Case study 2005
Experimental measurements were carried out in July–August 2005 at the Polish
Polar Station in Hornsund Fjord, Spitsbergen. Variability of the ABL was describedbased on air temperature measurements, monostatic Doppler sodar and tethered
balloon soundings. The main features of that season were two short-period
episodes of strong temperature inversion evident on the Ariekammen ridge, 363 ma.s.l. (July 6-8 and 23-25) and, long-lasting inversion on Hans Glacier on theelevation of 200 – 320m a.s.l. (Figure 1, Table1). The thickness of the ABL wasmostly around 250-300 m. and the lowest clouds commonly appeared within the
ABL (Table 2). In consequence, the glacier zone of 320 m a.s.l. was warmer thenthat at the height of 200 m a.s.l. It should be stressed that air temperature recordedat the Polish Polar Station (15 m a.s.l.) did not deflected from the long-term mean
conditions.
Table 2. Frequency of the sodar echo types in the classes of height - Hornsund July –August 2005).
Classes of height [m a.g.l.]
≤50
100
150
200
>200
Types of sodar echo
Frequency [%]Ground based layer
16,28 22,21 18,17 14,66
5,22
Elevated layer
5,94
31,96 34,25 27,85
Turbulence
23,47

Figure 2. Measured andcomputed vertical gradientof ablation of the Hans
Glacier in the summer of
2005.

500

Mean vertical gradient of ablation:measured= variability of gradientempirical eq. (1)= 0,36 m w.e./100m
Ohmura eq. (2)= 0,22 m w.e./100m
Khodakov eq. (3a) = 0,26 m w.e./100m
Khodakov eq. (3b) = 0,16 m w.e./100m

altitude [m a.s.l.]

400

300

200

100

0
0,5

1

1,5

2

2,5

ablation w.e. [m]measured

Ohmura eq.(2)Khodakov eq.(3a)empirical eq.(1)Khodakov eq. (3b)avg. 1988-2007

75

3

Long-lasting inversion of air temperature resulted in higher sum of PDD on theheight of 320 m a.s.l (221,4 PDD) then that recorded at the zone of 200 m a.s.l(162,2 PDD). In consequence, values of ablation reversed in the hypsometricprofile. Application of the Khodakov (1965) and Ohmura (1992) formulas tocalculate ablation did not expressed the changes in vertical profile (Figure 2).
Long-term data from Hans Glacier (1988-2007) indicate that the mean ablation has
a vertical gradient of 0,43m w.e./100m and vary in hypsometric profile between
1,91 m and 2,38 m w.e. in the frontal zone (60-120 m a.s.l.) and to 0,80 m w.e. in
the accumulation area, in the altitude of 400-500 m a.s.l.
500
Abla tion
Avg: -1,30m with 0,43/100m
Max (2001); -1,88m with 0,60/100m
Min (1994): -0,61m with 0,40/100m

altitude a.s.l. [m]400boundary of various ABL
effects - (air temperature inversion,
top of clouds and solar energy flux)

300

200

100

0
0,0

0,5

1,0

1,5

2,0

2,5

3,0

3,5

ablation [m w.e.]
2005

mean (1989-2007)

max (1989-2007)

min (1989-2007)

Figure 3. Vertical gradient of ablation effected by ABL structure – case study in 2005, meanand extremes in the period of 1988 – 2007.

Figure 4. Frontal area of Hans Glacier with marked line at the height of 300 m a.s.l.,(left)
and the sea of clouds with top height of 300m a.s.l., Hornsund Fjord. (Photos Marek
Kowalczyk

Plots of the mean gradient as well as the gradients in the seasons with maximum
and minimum ablation indicate a boundary at the height of 200-300m a.s.l. A zone
with decreased rate of ablation is evident in the lower part of the glacier (Figure 3).
It can be concluded that the effect of the arctic ABL and frequent air temperature76

inversions with low level clouds reduce the solar energy income at the surface(Figure 4).
However, the other phenomena, which modify dynamics of ablation, should beconsidered e.g. poorly recognized foehn winds with warming effect and katabatic
winds cooling the lower parts of a glacier.

References
Bing Lin B., Minnis, P., 2003: Cloud liquid water path variations with temperature observed
during the Surface Heat Budget of the Arctic Ocean (SHEBA) experiment, Journal ofGeophysical Research, 108, D4
Curry J.A., 1985: Interactions among Turbulence, Radiation and Microphysics in Arctic
Stratus Clouds, Journal of Atmospheric Sciences, 43,1 :90-106
Khodakov V.G., 1965, O zavisimosti summarnoj abljacii poverchnosti lednikov ot
temperatury vozducha. Meteor. i Gidrol., 7: 48-51.
Drzeniecka-Osiadacz A., Migała K., Sikora S., 2007, Cechy warstwy granicznej atmosfery w
rejonie Hornsundu w lipcu i sierpniu 2005 roku. [w]: Abiotyczne Środowisko
Spitsbergenu w latach 2005-2006 w warunkach globalnego ocieplenia, Wyd.
Uniwersytetu Mikołaja Kopernika: 131-146.
Ohmura A., Kasser P. and Funk M., 1992, Climate at the equilibrium line of glaciers. Journalof Glaciology, 38: 397-411.
Tjernström M, Leck C., Persson P.O.G., Jensen M.L., Oncley S.P., Targino A., 2004, The
Summertime Arctic Atmosphere Meteorological Measurements During The Arctic Ocean
Experiment 2001, Bull. Amer. Meteor. Soc., 85: 1304-1321.

77

INTERNAL STRUCTURE OF ARIEBREEN, SPITSBERGEN, FROM
RADIO-ECHO SOUNDING DATA
F. NAVARRO1, M. GRABIEC2, D. PUCZKO3, U. JONSELL4 AND A. NAWROT5
1

Dept. Matemática Aplicada, ETSI de Telecomunicación, Universidad Politécnica deMadrid, Spain.
2
Dept. of Geomorphology, Faculty of Earth Sciences, University of Silesia, Poland
3
Institute of Geophysics, Polish Academy of Sciences, Poland
4
Dept. of Earth Sciences, Uppsala University, Sweden
5
Institute of Paleogeography and Geoecology, Adam Mickiewicz University, Poland

Introduction

2

Ariebreen (77º 01' N, 15º 29' E) is a small valley glacier (ca. 0.36 km in August2007) located at Hornsund, Spitsbergen, Svalbard, ca. 2.5 km to the west ofHornsund Polish Polar Station. Ariebreen, like many other Svalbard glaciers, hasexperienced a significant recession at least since the 1930s, and most likely since
the end of Little Ice Age (LIA) in the early part of the 20th century. Moreover, thethinning rate of western Svalbard glaciers has shown an acceleration during themost recent decades. Ariebreen follows this general retreat pattern, as is shown inanother contribution to this workshop (Petlicki et al., 2008). Most investigatedglaciers in Hornsund area, in the neighbourhood of Ariebreen, are known to bepolythermal (e.g. Hansbreen and Werenskioldbreen, Pälli et al., 2003). It has beensuggested (Macheret et al., 1992) that the thinning of polythermal glaciers may
result in a switch to cold thermal structure under appropriate conditions. The strong
thinning experienced by Ariebreen during the recent decades makes it an ideal
candidate to undergo such change.
The main aims of this contribution are to understand the internal structure of
Ariebreen, in particular, its hydrothermal regime, and to determine whether theglacier is undergoing or has already experienced a transition from polythermal tocold structure. The main tool to accomplish this will be the analysis of radio-echosounding data.

Ground-penetrating radar measurements and data analysis
During summer 2006 and spring 2007 radar profiles were done providing a fullcoverage of Ariebreen (Figure 1). The total length of useful radar profiles was 2200
m in the summer campaign and 4000 m in the spring campaign, resulting a total
length of profiles of 6200 m. The radar data were acquired using an ice-penetratingradar Ramac/GPR with unshielded antennae, of 200 MHz in summer 2006 and 25
MHz in spring 2007. The choice of the 200 MHz antennae was dictated by theinterest in accurately determining the presence and extent, or the absence, of a firnlayer. For this reason, the sampling frequency within each radar trace was set quite
high, 2012 MHz, resulting a total time window of 509 ns and thus limiting the
maximum depth sampled to about 42 m. Consequently, the bedrock was notreached in the areas of thickest ice. Using the 25 MHz antennae, sampling
frequency was set to 250 MHz, resulting a time window of 2044 ns, so that bedrock78was reached in all cases. In addition to the radar data, navigation information froma stand-alone GPS receiver was also recorded; in the summer campaign, theprofile endpoints were also measured by differential GPS.
No common midpoint measurements were done on this glacier in order todetermine the radio-wave velocity (RWV). However, they are available, both forspring and summer periods, for the neighbouring Hansbreen glacier (Jania et al.
2005). On the basis of such measurements, we took a value of 168 m/µs (typical ofcold ice) for the RWV at Ariebreen. This gives us a vertical resolution of the radar
data of ±1.68 m for the 25 MHz antennae, and ±0.42 m for the 200 MHz antennae,
considering the vertical resolution as one quarter of the wavelength in ice.
The processing of radar data included DC correction, amplitude scaling, band-passfiltering, deconvolution, migration where required and conversion to depth.

Figure 1. Left: Schematics of the radar profiles done in Ariebreen, drawn on an image
showing the glacier extent in 1990. The straight blue lines correspond to profiles surveyed insummer 2006, and the remaining ones correspond to those surveyed in spring 2007. Right:
Ariebreen surface topography based on geodetic measurements done by the authors in
2007.

Results, discussion and conclusions
The ice thickness map for 2006-2007, constructed from radar data, and thesubglacial relief map, constructed by subtracting the ice thickness from the surface
topography map, are shown in Figure 2. The average ice thickness in 2007 was
only 22.29 ± 1.66 m. The thickest ice, with a maximum value of 81.24 ± 1.68 m, isfound in the west-east orientated upper part of the glacier; thickness for the northsouth orientated mid-lower part of the glacier only exceeds 50 m in its uppermost
part. The lower part shows very thin ice. The subglacial relief map shows gentle
slopes in the lower part of the glacier, steadily increasing as we approach theupper part.
79

The analysis of the radar sections shows three main features concerning theinternal structure of Ariebreen:
1. Nearly absence of endoglacial diffractions. This is true for both high (200 MHz)and low (25 MHz) frequency radar data.
2. Absence of any internal reflector, which could be interpreted as an interface
between cold and temperate ice layers. Again true for both high and low
frequencies.
3. Absence of a firn layer or, if any, a very thin one (2-3 m). The vertical resolutionof the radar data is 1.7 m for the 25 MHz radar and 0.4 m for the 200 MHz.
Though the uppermost 2-3 m are obscured in the radar records, neither layering
nor near-surface interface is visible below some 3 m depth.

Figure 2. Left: Ice thickness map of Ariebreen in 2006-2007. Right: Subglacial relief map.

On the basis of the above assertions, we preliminarily conclude that Ariebreen is acold glacier, in contrast with the polythermal structure of the neighbouring
Hansbreen and Werenskioldbreen. This is, however, not surprising, because of the
low ice thickness of Ariebreen, which allows for a relatively fast transmission
through the ice column, and release through the glacier surface, of the geothermalheat flow warming the glacier bed. In the accumulation area, the extremely small
thickness of the firn layer also implies that any possible heat generated byrefreezing of percolating meltwater rapidly reaches the glacier surface andtherefore does not substantially contribute to heating the glacier body. In theablation area, if a polythermal structure were assumed the small ice thickness
would also contribute to a fast transmission to the glacier surface of the latent heatgenerated by the freezing of liquid water in temperate ice at the cold-temperateinterface, and therefore to the downward migration of the freezing front. This,
combined with strong thinning, could justify a slow transition from polythermal tocold structure.
80References
Jania J., Macheret Yu.Ya., Navarro F.J., Glazovskiy A.F., Vasilenko E.V., Lapazaran J.,
Glowacki P., Migala K., Balut A., Piwowar B.A. 2005. Temporal changes in theradiophysical properties of a polythermal glacier in Spitsbergen. Annals of Glaciology 42,
125-134.
Macheret Yu.Ya., Krass M.S., Glazovsky A.F. 1992: Structure, hydrothermal state andregime of subpolar glaciers. In V.M. Kotlyakov (ed.): Regime and evolution of polar ice
sheets. Pp. 48-115. Saint-Petersburg: Gidrometizdat. (In Russian).
Pälli A., Moore J.C., Jania J., Kolondra L., Glowacki P. 2003. The drainage pattern of
Hansbreen and Werenskioldbreen, two polythermal glaciers in Svalbard, Polar Research
22, 355-371.
Petlicki M., Lapazaran J., Navarro F., Glowacki P., Machío F. 2008. Ice volume changes ofAriebreen, Spitsbergen, during 1936-1990-2007. Workshop on the dynamics and mass
budget of Arctic glaciers/ Glaciodyn (IPY) meeting, Obergurgl, Austria, January 2008.

81

POLAR PANORAMAS: IMAGES WORTH A THOUSAND MEGABYTES
MATT NOLAN
University of Alaska Fairbanks

Recent developments in digital camera and photo editing software offer earthscientists new ways to document and explore the landscapes they work in throughuse of spherical and gigapixel panoramas. For polar scientists working in remoteareas with landscapes that are affected on human time-scales by climate change,
panoramic photography offers many new possibilities for long-term change
assessment, après-field landscape analysis, and public outreach. This paper
provides an overview of the basic concepts in photo acquisition, post-processing,and scientific uses.
Panoramic photography has existed since the early days of photography itself, but
the advent of digital cameras and processing have made panoramic photographyeasier, better, and much higher resolution than previously possible. Panoramic
photography takes three basic forms: 1) a single acquisition using a specialized
camera that achieves a wide field of view by sweeping the lens or film, 2) cropping
an elongated image (typically with aspect ratio of 3:1 or higher) from a single
image, or 3) creating a single, seamless mosaic using a series of individual images
taken with a standard camera. This paper describes the latter, called the ‘stitching
technique’, and focuses on pure digital techniques. The digital mosaic can consistof any field of view, from a traditional 90° x 30°, to a single-row cylindrical stripe
covering 360°, to a full spherical view of 360° x 180°. Gigapixel panoramas cancover any field of view, with the distinguishing feature being a total pixel count of at
6least 5 x 10 . The final resolution of the digital mosaic is a function of the focal
length of the lens used and the resolution of the digital sensor in the camera. For
example, one can created a spherical panorama using a 10.5mm fisheye lens or a
50 mm prime lens – with a 12 megapixel (MP) camera, the fisheye lens will require
7 photos for a final resolution of about 11,000 x 5500 pixel (60 MP) and the 50 mm
lens will require 486 photos for a final resolution of over 4 gigapixels (GP).
Obviously there is a trade-off between time spent taking/stitching the photos andthe final resolution desired.
Creating a digital mosaic has several basic steps:
1. acquiring the images
2. performing any pre-stitching photo adjustments (such as conversion from RAW
format, pre-sharpening, exposure adjustments, etc)
3. stitching the images
4. blending the images to obscure the seams, exposure differences, or parallax
errors
5. using a photo editing software such as Photoshop to make any final
adjustments
6. converting to VR format to view at full resolution using pan, tilt and zoom
features82

There are a few principles to keep in mind when acquiring the photos to createhigh-quality panoramas. There are many websites, forums, and a few books thatoffer tutorials on digital stitching of panoramas (of note are http://wiki.panotools.org/
and the Panotools NG forum on Yahoo.com). Here I simply list a few of theprinciples to be aware of:• Panoramic photography is at heart still photography – the better you understandyour camera, its optical physics, the nature of light and its dynamic range, andthe essentials of composition, focus, and exposure, the better your panoramas
will be. The best books I have found on photography are the series by Ansel
Adams: The Camera, The Negative, and The Print.
• Panoramic photography requires you to think and expose in 360 degrees,
meaning that there will always be back-lighting and front-lighting on a sunnyday. Your camera exposure (aperture, shutter speed and white balance) mustremain constant (or nearly so) for all of your photos if you want them to blend tolook like a single photo, and so you need to plan your exposure carefully anduse a camera that has manual settings.
• The camera must be spun about its no parallax point (NPP) which varies withthe lens used, otherwise the photos will not stitch together perfectly, especially if
you have subjects in both near- and far-grounds. The NPP is almost always not
located at the tripod mounting screw in the camera body, and so a special
panoramic rig is required (I use the one from www.reallyrightstuff.com).
• The photos must have enough overlap between them to ensure enough points
in common to stitch them together. Typically 30% is enough. The maindifficulties are with pure blue skies, and here it helps to use a fixed rotational
increment between photos and input this value into the stitching software
directly.
• As with traditional photography, shooting in RAW format is best, as it allows forthe most control of the exposure, white balance, sharpening, and saturation
after the image is taken.
In terms of equipment, there is a lot to choose from, and here I will only describe
the basics along with the equipment that I use and have found reliable. For
spherical panoramas, the easiest method for field scientists is to use a fisheye lensand a monopod (or trekking pole with camera mount) with a tilting head. The Nikkor
10.5 mm fisheye is commonly considered the sharpest lens in the sub-$1000USD
category, though ones from Sigma and Tokina are also used widely. Depending on
the size of the camera sensor and the lens used, you will need between 3 and 7
photos to get full spherical coverage. The issue is that most lens/body
combinations will not result in a true 180° coverage, so there will be a gap at either
the top (zenith) or bottom (nadir) in the photo, and the camera must be tilted up or
down (or both) to fill those holes. My preferred method is to shoot 6 shots aroundwith camera tilted down to eliminate the nadir hole, then tilt the camera up to fill thezenith hole; the monopod must remain vertical during this process (use a bubble
level), otherwise parallax errors will result. For gigapixel panoramas (or any
panoramas using other than the fisheye lens) a tripod is required due to the larger
number of shots. For scenes requiring hundreds of photos, a motordrive is helpful
to help move the camera at a fixed increment and automatically release the
shutter. I have used the motordrive from Seitz and can recommend it for its
reliability and lightweight in remote field work. Nearly any modern D-SLR camera
83can be used for acquiring images; I use the Nikon D2xs and found it rugged,
reliable in any temperature, and with long battery life and recommend its use.
Once the photos are acquired, they must be stitched and blended. There are
numerous applications available for this. Arguably the best software is open-sourcecode called PanoTools, developed by Dr. Helmut Dirsch, and many commercial
packages use this engine. The most popular and among the cheapest commercial
packages is PTgui, which I use exclusively, and it has an active and friendly on-line
user community. Though the packages differ in implementation, the fundamentals
remain the same: find points in common between photos (‘control points’), use
those control points to determine how to shift (or warp) the photos, and then blend
the photos together to obscure any remaining errors. The best approach is tocalibrate your lens (something you can do on your own with test photos) and enter
those calibration coefficients into the stitching software, as this reduces the numberof control points required and prevents the software from warping the photos inways not intended by the user. Stitching and blending a spherical fisheye
panorama at 16 bit and 11,000 x 5,500 resolution can take a fast modern computer
between 5 minutes to 5 hours, depending on the software algorithms used. The
blending algorithms have the most variations, and recently released algorithms
appear to be 2 orders of magnitude faster than those of last year.
The final steps in production are preparing it for distribution and use, typically either
print or on-line viewing. Tools like Photoshop are useful here, especially withcleaning up any visible seams or with overall contrast and exposure adjustments.
Once the flat image is finalized, there are a number of tools that can be used tocreate the on-line display. I use “Pano2VR”, as it is easy to use and inexpensive.
Tools like these will take your flat image (typically a JPG) and convert them to
QuickTime VR or Flash (or any number of other similar formats). The issue here isthat only about 90° field of view can be shown without introducing substantial
distortion, so that user’s field of view must be restricted to this amount. The
Quicktime and Flash applications allow the creator to limit the field of view as
desired, and allow the user to pan, tilt and zoom to see the remaining image. For
gigapixel panoramas, the best software currently available is Microsoft’s HDview,
but Zoomify is also a popular and useful tool. These applications take theenormous gigapixel panorama and split into thousands of smaller images with the‘pyramid resolution layers’, such that the software can carefully manage theavailable bandwidth by only sending the user as much resolution as required given
the zoom level and location of the user’s browser. This is the same way that
Google Earth and many other mapping software works.
To explore the scientific utility of spherical and panoramic photography, I tookapproximately 375 panoramas during an expedition to McCall Glacier Alaska in
July/August of 2007. I have stitched approximately 60 of these at this time, andhave found that it takes me about 2 hours for the complete workflow for each. I
have uploaded all of these to http://arctic.360cities.net and they can also be viewed
in Google Earth from http://arctic.360cities.net/ge.kml.
It remains to be seen what the long-term value of these panoramas will be, but I
can offer several possibilities. I have found that my gigapixel panoramas allow me
84

to explore the landscape from the office nearly as well as or better than standing in
the field does, as the photo resolution is better than my eyes can see alone andthere I can spend hours or days looking at the photo, but typically only a few
minutes in the field. The spherical panoramas, which can be taken in a minute or
so in the field, cover a broad spatial area and so provide nice documentation of thecomplete landscape of our primary glacier field site. I have used these photos on
numerous occasions now to help collaborators who have not been to our study
areas better understand the area, and all have commented on the power of thetechnique. In terms of being a long-term record, based on my own experiencesusing old photos, the panoramic technique is far superior, as it not only gives thewould-be repeat photography a 360 degree view against which to assess change
but also shows exactly the location of the photographer through the nadir image.

Figure 1. Gigapixel panorama of Okpilak Glacier, AK. The upper image is a mosaic
composed of 504 images acquired with a 200 mm lens. The lower image is a cut-out of themosaic demonstrating the detail captured by this technique. This photo-location was firstused 100 years ago by Ernest Leffingwell, when the glacier almost completely filled the fresh
moraine seen here. Gigapixel panoramas like this offer polar scientists new ways to
document climate-sensitive landscapes, detect temporal change, and potentially develop
better insights into the processes involved.

85

RAPID SURFACE ELEVATION LOSSES OF SPITSBERGEN GLACIERS(1936-1990-2005)
CHRISTOPHER NUTH1, GEIR MOHOLDT1, JACK KOHLER2, JON OVE HAGEN1 AND
HARALD FASTE-AAS2
Dept. of Geosciences, Section of Physical Geography, Faculty of Mathematics and Natural
Sciences, University of Oslo, Norway
Norwegian Polar Institute, Norway

Introduction
Elevation changes are a common approach towards extracting glacier changes ofthe largest ice caps Greenland and Antarctica (Krabill et.al., 2000; Wingham et.al.,
1998), as well as regions comprising of smaller glaciers like Alaska (Arendt et.al.,
2002). ICESat contains a satellite laser altimeter that acquires elevation data withremarkable precision and accuracy (Zwally et.al., 2002; Brenner et.al.,2007). In thisstudy, previous long term elevation changes from 1936 to 1990 (Nuth et.al., 2007)
are compared to modern changes derived from comparing the 1990 DEM(Norwegian Polar Institute) to ICESat profiles from 2003 to 2007.

Methodology and Accuracy
Elevations from the 1990 DEM are compared at each ICESat point disregarding
the date of aquisition. Elevation change curves are then averaged altitudinally andby region. These change curves therefore represent annual averages of different
time intervals ranging from 13 years (1990-2003) to 17 years (1990-2007). The
affect of seasonal snow cover, though minimal, results in underestimated thinningand over estimated thickening. The accuracy is derived by comparing nonglacierized stable terrain. The histogram of these changes (Fig. 1) show that 95%
of all the datapoints lie within ±5 meters with a standard deviation of 2.3 meters.

Figure 1. Histogram of elevationdifferences between ICESat andthe 1990 DEM (NP) over nonglacierized terrain.

86Results and Perspectives
Figure 2 shows the elevation change curves for the older time series, 1936-1990(Nuth et.al., 2007) as compared to the newest time interval (1990-2003/07) for six
geographic regions. In all regions, the elevation change curves have become morenegative. Averaged thinning rates at the glacier fronts have doubled between thetwo periods with recent losses of up to -2.5 m/yr. Seasonal snow cover is notaccounted for, though removal only increases thinning rates. The upper altitudes
have generally decreased in elevation. These decreases may not be an indicationof water loss, but rather the result of internal accumulation and firn densification.
Future work should focus on determing how much internal accumulation is
occuring, and how this affects firn densification and thus elevation changes. If thisprocess is occuring and not accounted for, water equivalent volume losses would
be overestimateed. In summary, the glaciers of southern Spitsbergen are recently
thinning more rapidly, even at upper altitudes which are prone to firn densification.
The assumption of constant density for water equivalent conversion therefore
needs to be addressed.

Figure 2. Annual elevation change curves for 6 regions in southern Spitsbergen.

87References
Arendt A.A., K.A. Echelmeyer, W.D. Harrison, C.S. Lingle, V.B. Valentine. 2002. Rapid
wastage of Alaska glaciers and their contribution to rising sea level. Science 297, 382386.
Brenner A., DiMarzio J., and Zwally, J. 2007. Precisionand Accuracy of satellite radar andlaser altimeter data over the continental ice sheets. IEEE Tran.Geo.Rem. 45(2).
Krabill and 9 others. 2000. Greenland ice sheet: high-elevation balance and peripheral
thinning. Science 289, 428 – 430.
Nuth C., J.Kohler, H.F.Aas,O.Brandt.,J.O.Hagen. (2007).Glacier geometry and elevaiton
changes on Svalbard. Ann.Glac 46: pp 106-116.
Wingham D. J., A.J. Ridout, R. Scharroo, R.J. Arthern, C. K. Shum. 1998. Antarctic
elevation change from 1992-1996. Science 282 (5388): pp. 456-458.
Zwally and others(2002). ICESat’s laser measurements of polar ice, atmosphere, ocean,and land. Journal of Geodynamics 34: pp 405-445.

88

ICE VOLUME CHANGES OF ARIEBREEN, SPITSBERGEN, DURING
1936-1990-2007
M. PETLICKI1, J. LAPAZARAN2, F. NAVARRO2, P. GLOWACKI1 AND F. MACHÍO3
1

Institute of Geophysics, Polish Academy of Sciences, Poland,
Dept. Matemática Aplicada, Universidad Politécnica de Madrid, Spain
3
Dept. Electrónica y Comunicaciones, Facultad de Ingeniería Informática, Universidad
Pontificia de Salamanca, Spain
2

Introduction

2

Ariebreen (77º 01' N, 15º 29' E) is a small valley glacier (ca. 0.36 km in August2007) located at Hornsund, Spitsbergen, Svalbard, ca. 2.5 km to the west ofHornsund Polish Polar Station. Many Svalbard glaciers have experienced asignificant recession at least since the 1930s, and most likely since the end of Little
Ice Age in the early 20th century (Werner, 1993). It has manifested as thinning andretreating of ice fronts, though a simultaneous thickening at the uppermost
elevations in many locations has been reported (Bamber et al., 2004; Nuth et al.,2007). Moreover, the thinning rate of western Svalbard glaciers has shown anacceleration during the most recent decades (Kohler et al., 2007).
The main aims of this contribution are to determine whether Ariebreen follows such
retreat pattern and to quantify the retreat it has experienced, in terms of area,
thickness and volume changes, to estimate the average mass balance equivalent
to the ice volume change during the period under investigation, and to estimate thevolume of ice presently stored in Ariebreen. The main tools to accomplish this willbe the analysis of digital terrain models (DTM) of the glacier surface corresponding
to different dates, and the radio-echo sounding of the ice body to determine thepresent ice volume. The latter is described in a separate contribution to thisworkshop (Navarro et al., 2008).

Figure 1. Surface topography of Ariebreen in 1936 (left), 1990 (centre) and 2007 (right).

89

Construction of DTMs and estimation of ice volume changes 19361990-2007 and total volume (2007)
The volume changes 1936-1990 and 1990-2007 were estimated by subtracting theDTMs for the corresponding years, constructed on a common (for each pair) grid
5 m x 5 m. The total volume in 2007 was estimated from the ice thickness map
constructed from the radar data.
The DTM for 1936 (Figure 1, left) was constructed from the sheet B12 (Torellbreen,Spitsbergen) of the 1:100 000 Svalbard topographic map (Norsk Polarinstitutt
1953), which is a photogrammetric compilation from aerial photographs taken in
1936. The errors in the DTM for 1936 are the most prominent ones, because of theuncertainties inherent to the original map and the poor coverage of the source
data, limited to the glacier boundary and a few contour levels. We have taken as
an estimate of the vertical accuracy of the individual data points the value of±12.49 m for Wedel Jarlberg Land (where Ariebreen is located) obtained by Nuth etal. (2007).
The DTM of the glacier surface in 1990 (Figure 1, centre) was extracted from theDTM of Werenskiold area constructed by Kolondra (2002) from infrared false
colour aerial photographs taken in August 1990. The vertical accuracy for theoriginal 1990 DTM was estimated by its author as ±0.5 m. However, for ourreconstruction of the glacier surface on the same common grid 5 m x 5 m as 1936
DTM, a standard deviation of ±2.08 m resulted for the interpolation error of thevertical coordinates by kriging.
The DTM of Ariebreen for 2007 (Figure 1, right) was based on our own geodetic
measurements made in August 2007 with a Leica TCR-1105 total station. A total of
646 points were measured, of which 192 corresponded to the glacier boundary.
The resulting DTM for 2007, with the same regular grid 5 m x 5 m as the 1990
DTM, showed a standard deviation of ±2.42 m for the interpolation error of thevertical coordinates by the kriging method.
The volume changes between any two years were estimated computing, by
Simpson’s method, the volume resulting from the subtraction of the DTMs for thecorresponding years, constructed on a common grid. The total ice volume in 2007
was estimated from the ice thickness map retrieved from the radar data. The ice
thickness data were also subtracted from the 2007 glacier surface DTM,
constructed on the same grid, in order to determine the subglacial bedrock map.
The ice thickness and subglacial relief maps are shown in another contribution tothis workshop (see Figure 2 in Navarro et al., 2008).

Results, discussion and conclusions
The area, average ice thickness and ice volume of Ariebreen in 1936, 1990 and
2007, together with their changes during the periods 1936-1990, 1990-2007 and
1936-2007, are shown in tables 1 and 2, while Figure 2 shows the ice thickness
changes experienced by the glacier during the mentioned periods.

90

Figure 2. Thickness change of Ariebreen 1936-1990 (left), 1990-2007 (centre) and 19362007 (right).

Notice the distinction between the change in average thickness (Table 1) and theaverage thickness change (Table 2), resulting from the computation over different
areas, the latter being that of the start of the period, which is larger and then resultsa smaller average.
Quite remarkable changes are observed for all parameters in Table 1 and 2. The
largest percent change corresponds to the ice volume (that has decreased by
69±29% during the full period analysed), as this figure combines the effects of bothchanges in area –mainly due to ice front retreat– and changes in ice thickness. Thechanges in ice volume during the periods 1936-1990, 1990-2007 and 1936-2007
are equivalent to average annual mass balances of –0.44±0.27, –1.00±0.27 and –
0.57±0.24 m/y w.e., respectively, computed considering an ice density of 900
kg/m3 and an average area for each period.
Table 1. Glacier area, average ice thickness and ice volume of Ariebreen in 1936, 1990 and
2007, and changes for the periods 1936-1990, 1990-2007 and 1936-2007.
1936
1990
2007
ChangeChangeChange% Change
1936-1990 1990-2007 1936-2007 2007/1936
Area
0.689
0.504
0.363
−0.185−0.141−0.325−53 ± 62(km )±0.040±0.038±0.009±0.055±0.039±0.041
Average
50.22
37.70
22.29
−12.52−15.41−27.93−56 ± 26
thickness (m)±12.83±3.61±1.68±12.82±3.98±12.943
Volume (km )0.03460.01900.0108−0.0156−0.0082−0.0238−69 ± 29±0.0098 ±0.0024 ±0.0008±0.0096±0.0022±0.0100
Table 2. Average thickness change of Ariebreen for the periods 1936-1990, 1990-2007 and
1936-2007.
1936-1990
1990-2007
1936-2007
Average−22.65−16.22−34.52
thickness change (m)±12.66±3.19±12.7291

The above data show that Ariebreen has followed the general recession patternobserved in Svalbard, manifested as thinning and retreating of ice fronts, though asimultaneous thickening at the uppermost elevations, an also shows theacceleration in thinning rate observed in western Svalbard glaciers during the mostrecent decades.

References
Bamber J., Krabill W., Raper V., Dowdeswell J. 2004. Anomalous recent growth of part of alarge Arctic ice cap: Austfonna, Svalbard. Geophysical Research Letters 31, L12402,
doi: 10.1029/2004GL019667.
Kohler J., James T.D., Murray T., Nuth C., Brandt O., Barrand N.E., Aas H.F, Luckman A.
2007. Acceleration inthinning rate on western Svalbard glaciers, Geophysical Research
Letters 34, L18502, doi: 10.1029/2007GL030681.
Kolondra L. 2002. Werenskioldbreen and surrounding areas, Spitsbergen, Svalbard,
Norway. Orthophotomap 1:25000. Published by Silesia University, Sosnowiec, andNorsk Polarinstitutt, Tromso.
Navarro F., Grabiec M., Puczco D., Jonsell U., Nawrot A. 2008. Internal Structure of
Ariebreen, Spitsbergen, from radio-echo sounding data. Workshop on the dynamics andmass budget of Arctic glaciers/ Glaciodyn (IPY) meeting, Obergurgl, Austria, January
2008.
Nuth C., Kohler J., Aas H.F., Brandt O., Hagen, J.O. 2007. Glacier geometry and elevation
changes on Svalbard (1936-90): a baseline dataset. Annals of Glaciology 46, 106-116.
Werner A. 1993. Holocene moraine chronology, Spitsbergen, Svalbard: lichenometric
evidence for multiple Neoglacial advances in the Arctic, The Holocene 3, 128-137.

92

ALTIMETRY OF NORDENSKIÖLDBREEN AND VESTFONNA 1991 2007 - PRELIMINARY RESULTSVEIJO POHJOLA1, JIM HEDFORS1, JOHN MOORE2, BJÖRN SJÖGREN1, LESZEKKOLONDRA3, CARLEEN REIJMER4, ELISABETH ISAKSSON5, AND JON OVEHAGEN61

Department of Earth Sciences, Uppsala university, Sweden.
Arctic Centre, University of Lapland, Finland.
3
University of Silesia, Poland.
4
IMAU, University of Utrecht, The Netherlands.
5
Norwegian Polar Institute, Norway.
6
University of Oslo, Norway.
2

Introduction
Regional modelling of the future of the Arctic finds the average warming over theHigh Arctic may likely be in the range of 5-7˚C during the present century (1,2). The-1current rate of sea level increase is 3.2 mm yr (3), to a large part as a product ofmelting ice caps (4). Recent compilations points to that the Antarctic ice sheets are
more or less in balance within the present error bars, and that Greenland delivers
ca 10% of this increase (5,6), suggesting that most of the contribution to the current
sea level increase may be from ice outside these large ice sheets. With the
projected large warming of the Arctic, areas closer to the advection pathways ofheat will initially be more affected than areas more isolated to exchange of heat.
Svalbard is such a sensitive area, due to the proximity to the North Atlantic drift,
and may be expected to show large changes with future warming.

Figure 1. Svalbard and the ice
fields Lomonosovfonna andVestfonna.
©
NorsktPolarinstitutt.

93

Svalbard is ~60% glaciated, and the amount of glaciation increases towards thenorth-eastern part of the archipelago (Figure 1), due to predominant easterly
airflow during precipitation events. The total ice coverage of Svalbard is two times
the ice coverage on the rest of Europe (including Iceland) (7). This suggests thestate and the fate of the large ice fields and ice caps on Svalbard should have ahigh attention in a European perspective, since they encompass an area of highexpected warming and are the largest ice volumes in Europe.

Previous altimetry on Lomonosov- and Vestfonna
Earlier campaigns on the two ice fields by surface DGPS altimetry (8) did not show
any conclusive changes in altimetry on Lomonosovfonna / Nordenskiöldbreen
during the period 1991-1997. This picture was changed for the period 1996-2002
when aerial laser altimetry was flown by NASA (9). The results from the NASA-1
fligths show a subsidence of more than 0.5 m a below 600 m elevation, but ageneral growth above 1000 m elevation (9). The NASA altimetry flown over
Vestfonna the same years show a general subsidence of Vestfonna of the order of-1
0.0 - 0.6 m a along the flight line. A general conclusion, adding the altimetricobservation over Austfonna in the same NASA missions (10), is a growth inelevation on the highest parts of these ice fields, but a general subsidence at least
below 600 m a. sl.

Figure2.
Lomonosovfonna and theoutlet Nordenskiöldbreen.
The altimetry measuredby (8) are marked bypurpledots.
Thegeometrical symbols, andthe DGPS line driven in2007 by the black track.
The background is a
LANDSAT image.

Altimetry campaigns on Lomonosov- and Vestfonna spring 2007
The years after 1997 have been anomalously warm on Svalbard, with longer melt
periods, and also years with high precipitation. This may explain the subsidence ofthe lower elevations, and the growth on higher altitudes found by the NASAmissions. This trend of warmer temperatures and higher precipitation has
continued over the decade. This motivated our mission to re-measure the altimetryof the two ice fields by driving surface borne DGPS over parts of the two ice fields.
94

The mission accomplished in 2007 is only a first campaign of three consecutive
years of planned surface borne profiling. This gives all data is incomplete and only
preliminary at this stage. Figure 2 shows the profiles driven on Lomonosovfonna
2007, and Figure 3 shows the changes over the different measurements periods
from (8,9) and the 2007 campaign. This preliminary data suggest that the trends
found by (9) prevail into the five year period between 2002 and 2007.
-1

1991-1997 dZ = -0m
Δ= 0 ma-1
1997-2002 dZ = -6.2 m Δ = -1.2 m a-1
2002-2007 dZ = +15.6 mΔ = +3.1 m a-1
1991-2007 dZ = +9.8 m Δ = +0.6 m a-1

1991-1996 dZ = -1.3 m Δ = - 0.3 m a-1
1996-1997 dZ = -1.6 m Δ = - 1.6 m a-1
1997-2002 dZ = +0.2 m Δ = 0.0 m a-1
2002-2007 dZ = -14.0 m Δ = - 2.8 m a-1
1991-2007 dZ = -16.7 m Δ = - 1.0 m a

Figure 3. Changes insurface level (dZ) between
the different campaigns on
Lomonosonfonna. Δ is therate of change per year.
The
graynumbers
distributions indicate lesscertain values. Squares andtriangles is measurementsby (8). The lines aremeasurements 2007.

Figure 4. Vestfonna and thealtimetry driven 2007 shown as the
purple track on the ice cap. The
background is a LANDSAT image.
The black rectangle is theapproximate area covered in
Figure 5.

On Vestfonna no ground based campaigns have been accomplished up to date.
The only data is the NASA flight lines (9), and the altimetry flown by Scott Polar
Research Institute in the 1980s (11). Figure 4 shows the DGPS profiling done
2007, and Figure 5 show the trends between 2002 and 2007 over these few points
95the NASA flight lines crosses our ground based profiling. In all three cases ourpreliminary results points at a continued lowering rate throughout the last ten years.
dZ ~ -7.3 m-1→ Δ ~ -1.4 ma
dZ ~ -7.2 m-1→ Δ ~ -1.4 ma

Figure 5. Changes in surface level (dZ)between the different campaigns on
Vestfonna. Δ is the rate of change per
year. The triangles are heights on
position markers for DGPS, and theblue track is kinematic GPS from 2007.
The gray lines are airborne altimetry
flown 2002. All heights are over
WGS84.

dZ ~ -2.5 m-1→ Δ ~ -0.5 ma

Acknowledgements.
The IPY projects GLACIODYN and KINNVIKA, and the logistical support from theNorwegian Polar Institute and the Swedish Polar Research Secreteriat are warmly
acknowledged. We further acknowledge finacial support from the Swedish Science
Council and Svenska Sällskapet för Antropologi och Geografi.

References
1. ACIA scientific report. 2004. http://www.acia.uaf.edu/.
2. IPCC. 2007. Global Change 2007 - the 4th assessment report. http://www.ipcc.ch/
3. Cabanes, C., A.Cazenave and C.LeProvost. 2001. Sea level rise during past 40 years
determined from satellite and in situ observations. Science, 294, 840-842.
4. Alley, R.B., P.U.Clark, P.Huybrechts and I.Joughin. 2005. Ice-sheet and sea-levelchanges. Science, 310, 456-460.
5. Raper, S.C. and R.J.Braithwaite. 2006. Low sea level rise projections from mountainglaciers and icecaps under global warming. Nature, 439(19 jan), 311-313.
6. Luthcke, S.B. and 8 others. 2006. Recent Greenland ice mass loss by drainage systemfrom satellite gravity observations. Science, 314(24 nov), 1286-1289.
7. Meier, M.F., and 7 others. 2007. Glaciers dominate eustatic sea-level rise in the 21st
century. Science, 317(24 Aug), 1064-1066.
8. Hagen, J.O., T.Eiken, J.Kohler and K.Melvold. 2005. Geometry changes on Svalbard
glaciers: mass-balance or dynamic response? Ann.Glaciol., 42, 255-261.
9. Bamber, J.L., W.Krabill, V.Raper, J.A.Dowdeswell and J.Oerlemans. 2005. Elevation
changes measured on Svalbard glaciers and ice caps from airborne laser data.
Ann.Glaciol., 42, 202-208.
10. Bamber, J., W.Krabill, V.Raper and J.Dowdeswell. 2004. Anomalous recent growth ofpart of a large Arctic ice cap: Austfonna, Svalbard. Geophys.Res.Let., 31(L12402,
doi:10.1029/2004GL019667).
11. Dowdeswell, J.A., D.J.Drewry, A.P.R.Cooper, M.R.Gorman, O.Liestøl and O.Orheim.
1986. Digital mapping of the Northaustlandet ice caps from airborne geophysicalinvestigations. Ann.Glaciol., 8, 51-58.

96

MASS BALANCE AND VELOCITYNORDENSKIÖLDBREEN, SVALBARD

OBSERVATIONSONCARLEEN REIJMER1, MARIANNE DEN OUDEN1, VEIJO POHJOLA2 BJÖRNSJÖGREN2 AND HANS OERLEMANS1
1
2

Institute for Marine and Atmospheric Research, Utrecht University, The Netherlands
Department of Earth Sciences, Uppsala University, Sweden.

Introduction
Recent studies suggest that in a warming climate, increasing supply of (melt)water
to the drainage system of a glacier may lead to higher ice velocities, at least
temporarily (Zwally et al., 2002). In order to study the relation between (melt)water
input and ice velocities, and focusing on seasonal and inter-annual variability,
mass balance and continuous ice velocity measurements are carried out on
Nordenskiöldbreen (Svalbard) (Figure 1a). The ice velocity measurements are
performed using relatively low cost stand alone GPSs (Figure 1b). These GPSs
record their position every hour, from which flow velocities are calculated. The
mass balance is determined by traditional stake measurements and by means of asonic height ranger. The locations of the stakes and GPSs are marked in Figure 1.
Here we report first results for the period March 2006 - March 2007.

Figure 1. Left: Nordenskiöldbreen,
Svalbard. Red dots are GPS andmass balance stake locations on theglacier, green dot marks the locationof the GPS reference station, NBB.
Location marked NB5H is the locationof the sonic altimeter. Locations NB1,
2, 4-8 and B were established in
March 2006, NB3 and 9-11 wereplaced in March 2007. Inset: map ofSvalbard with the location of theglacier marked by a red box. Top:
Stand alone GPS mounted on a massbalance stake on Nordenskiöldbreen.

97

Mass balance
Knowledge of the mass balance of Nordenskiöldbreen is limited to accumulation
values for the higher parts of the glacier, the Lomonosovfonna ice field, based on
ground penetrating radar observations and ice cores. Estimates range from 0.41 mw.e. (Pohjola et al. 2002) to 0.52 m w.e. (Pälli et al., 2002). For the lower parts ofthe glacier estimates are thus far based on the average mass balance gradient
presented by Hagen et al. (2003) in combination with the values presented by Pälli
et al. (2002) and an average value of the equilibrium line of 600 m a.s.l. The
measurements presented in figure 2 are the first available for the lower part of
Nordenskiöldbreen. The values in Figure 2a are based on stake readings using an
3
3
ice density of 900 kg/m and snow density of 400 kg/m . They show an increase inmass balance from an ablation of ±2 m w.e. at ±100 m a.s.l. to 0.2 m w.e. ofaccumulation at ±700 m a.s.l. The equilibrium line altitude (ELA) was at ±660 ma.s.l., between sites 7 and 8. Compared to the gradient estimated by Hagen et al.
(2003), the observed gradient is smaller, -0.43 m w.e. and -0.46 m w.e. per 100 m
altitude increase, respectively. Note however that we present only one year of
data.

Figure 2. Left: Annual mass balance (April 2006 - March 2007). Numbers denote the stake
locations in Figure 1. The red dotted line indicates the equilibrium line altitude. The blue
dotted line indicates the estimated mass balance gradient from Hagen et al. (2003). Right:
Time series of distance to the surface as measured with the sonic height ranger (site NB5H
in figure 1).

The temporal variations in the mass balance are measured using a sonic height
ranger (Figure 2b). The data are corrected for the effect of changes in air
temperature on the speed of sound. Several features can be noted in this record. A
snowfall event resulted in a decrease of 0.3 m in distance between sensor andsurface at the end of March. After this event, the distance increased from May till
August, showing about 2 m of snow/ice melt. Furthermore, the autumn of 2006 wasvery dry, no melt took place and accumulation only started in January 2007. Notealso the variability in the record from July onwards, which is caused by the sensor
hanging loose on its pole, moving in the wind. From the record we guess that thesensor came loose on 14 June.
98

Ice velocity
The GPSs record their position every hour. The accuracy of the sampled locations
can be estimated from the time series obtained at the reference site NBB. Figure
3a shows for every sample recorded at NBB the displacement with respect to theaverage location of NBB. Maximum displacement is 5.1 m and the standard
deviation is 1.8 m, which is a measure for the annual accuracy in the displacement.
Time series of velocities are determined by applying a Gaussian weighted runningaverage of 360 hours over the recorded locations, then calculating velocity fromthe resulting hourly displacements and applying a Gaussian weighted runningaverage of 360 hours over the resulting velocities. The results for NB4, 6 and 7 areplotted in Figure 3b. The figure shows large variations in flow velocity throughout
the year with a clear maximum around 10 July 2007, about 3 weeks after the onsetof the melting season in June. There is no clear lead or lag in timing of the
maximum in flow speed. This is partly explained by the limited distance between
the locations (order of km). After this peak in velocities, velocities decrease to on
average lower values in winter. The records also exhibit variability on time scales of
2-3 weeks, for which at present we do not have an explanation. We do not show
NB1, 2, 5 and 8. NB1 and 2 are not on the main flow line of Nordenskiöldbreen(see Figure 1) and NB5 and 8 did not work properly due to technical problems.

Figure 3. Left: Displacement at NBB with respect to its mean location. The red circle
denotes the standard deviation in the average displacement. Right: Flow velocity record at
NB4, 6 and 7.

We use the GPS data also to determine annual flow velocities at all sites. Results
are given in Table 1. The table presents annual mean velocities determined using
three methods, two based on the stand alone GPS records and one based on
traditional DGPS observations. The flow velocities are clearly much lower on NB1
and 2 compared to the sites on the main flow line. Flow velocities at the lower partof the glacier are on the order of 45 m/yr. Compared to the DGPS observations,which have an accuracy of 0.1 m, the stand alone GPSs show similar values within
their uncertainty range of 3.7 m/yr.

99

Table 1. Annual ice velocities. Av refers to values based on velocity averages as presented
in figure 3b, E-S refers to values based on first and last location in the time series. The
DGPS values are based on traditional DGPS measurements. Given percentages refer toamount of available data on which the velocity value is based when smaller than 99%.
StationGPS (m/yr)
GPS (m/yr)
Max (m/yr)
DGPS (m/yr)
(Av)
(E-S)
NB1
9.2
4.0
36.2
NB2

7.9

5.6

18.2

NB4

47.8

47.2

74.3

46.3

NB5H
NB6
NB7
NB8

27.1 (32%)
43.2
46.5
39.6 (90%)

28.6
46.0
45.7
39.0

48.7
81.9
88.5
52.5

40.2
50.2
36.6

OutlookThe stand alone GPS have provided us with mass balance and velocity records at
7 locations on Nordenskiöldbreen. In 2008, data from 4 more locations will become
available. The data will be used for tuning and validation of mass balance anddynamical models, the latter including the effect of (melt)water on the flow
velocities. On the technical site, the GPSs are developed further in order toincrease their accuracy.

References
Pälli, A., J. Kohler, E. Isaksson, J.C. Moore, J.F. Pinlgot, V.A. Pohloja and H. Samuelsson,
2002. Spatial and temporal variability of snow accumulation using ground-penetratingradar and ice cores on a Svalbard glacier. J. Glaciol., 48(162), 417-424
Pohjola, V.A. , T.A. Martma, H.A.J. Meijer, J.C. Moore, E. Isaksson, R. Vaikmäe and R.S.W.
van de Wal, 2002. Reconstruction of three centuries of annual accumulation rates basedon the record of stable isotopes of water from Lomonosovfonna, Svalbard. Ann. Glaciol.,35, 57-62.
Hagen, J.O., J. Kohler, K. Melvold and J.G. Winther, 2003. Glaciers in Svalbard: massbalance, runoff and freshwater flux. Polar Research, 22(2), 145-159.
Zwally, H. J., W. Abdalati, T. Herring, K. Larson, J. Saba and K. Steffen, 2002. Surface MeltInduced Acceleration of Greenland Ice-Sheet Flow. Science, 297, 218-222.

100

SPATIO-TEMPORAL VARIABILITY OF SNOW MELT ON SVALBARDDERIVED FROM SPACEBORNE SCATTEROMETER DATA (QSCAT)GERIT ROTSCHKY, JACK KOHLER, ELISABETH ISAKSSONNorwegian Polar Institute, Tromsø, Norway

Introduction
Evidence for ongoing large changes in the Arctic climate has been accumulating
during the last decade, particularly with regards to sea ice, permafrost and glaciermass balance. Due to the effect of global warming dramatic changes in the Arctic
snow and ice coverage are currently observed, expressed by a reduction of 10%
over the last 30 years associated with an extended and longer lasting melting
season. Previous studies have demonstrated the capacities of active microwave
instruments for the detection of surface melt and freeze-up due to the highsensitivity of radar backscatter to snow wetness (e.g. Wismann, 2000; Nghiem etal., 2001; Smith et al., 2003; Steffen et al., 2004; Wang et al., 2005). Spaceborne
scatterometers provide data at low spatial but high temporal resolution allowing
consistent observations on a daily time scale.
This study focuses on Svalbard characterized by a highly variable climate
throughout the year due to its position within a zone that includes both the polar
ocean and atmospheric fronts between the Arctic Ocean, Nordic Seas and Barents
Sea. Nevertheless, over a long enough period of time we expect to see a generalclimatic trend by monitoring the spatio-temporal variability of snow distribution andmelt all over Svalbard. For this we utilize microwave backscattering measurements
continuously carried out by the Ku-band scatterometer QSCAT since fall 1999.
Meteorological data from weather stations around Svalbard are available toinvestigate links between atmospheric circulation, moisture transport, near surface
air temperature and corresponding deposition and melting of snow, respectively.
The analysis demonstrates that coarse resolution scatterometer data can be
usefully applied to trace consequences of a warming climate in the Arctic.

0

Figure 1. Qscat σ measurements plotted against Ny-Ålesund temperature records.

Methodology
We utilized Qscat enhanced resolution products (Long et al., 1993) to detectspatio-temporal changes in microwave backscattering properties of Svalbard. For agiven frequency the latter are mainly a function of surface wetness, temperature,and density. Due to high absorption of radar waves by liquid water within asnowpack the beginning of the melt season is clearly marked by a sudden drop of101

0

the backscatter coefficient σ linked to air temperatures rising to above-zero values(Figure 1). By taking advantage of this connection we mapped melt onset as wellas number of melt days per year for all of Svalbard since 1999 based on a dynamic0
σ threshold. Melt was thereby assumed when the backscattering within a specific
Qscat resolution cell had lowered to 1.35 times its mean winter value. However, forsome years this simple melt detection approach is easily obscured by single rainfall
events that occur within the winter season visible as spike-like features within theσ0 time series. Therefore, short-term backscatter reductions are removed from theanalysis.

Results
The annual melt intensity, expressed as the number of melt days in each year, isfound to be highest in the south and along the western coast of Svalbard. To
compare spatio-temporal differences we calculate an average backscatter
0
response (σ mean) for 7 Svalbard regions. While there is pronounced regional andinterannual variability in melt onset as well as number of melt days per year, thereis nevertheless a trend of earlier melt onset and an increasing number of melt days
per year apparent for the period of observation (Figure 2), as expected in a
warming climate. Further, in some years striking synchronous melt onsets all overSvalbard can be observed, which are most likely the result of regional-scale rainfall
events.

Figure 2. (a) Summer melt onset and (b) number of melt days per year for Svalbard(regional variability in gray).

References
Long, D.G., P.J. Hardin, P.T. Whiting, 1993. Resolution enhancement of spaceborne
scatterometer data, IEEE Trans. Geosci. Remote Sens., 31(3):700-715.
Nghiem, S.V., K. Steffen, R. Kwok, W.-Y. Tsai, 2001, Detection of snowmelt regions on theGreenland ice sheet using diurnal backscatter change, J. Glaciol., 47(159), 539-547.
Smith L.C., Y. Sheng, R.R. Forster, K. Steffen, K.E. Frey, D.E. Alsdorf, 2003. Melting of
small Arctic ice caps observed from ERS scatterometer time series, Geoph. Res. Lett.,
30(20), 2034.
Steffen, K., S.V. Nghiem, R. Huff, G. Neumann, 2004. The melt anomaly of 2002 on theGreenland Ice Sheet from active and passive microwave satellite observations. Geoph.
Res. Lett., 31, L20402, doi:10.1029/2004GL020444.
Wang L., M. J. Sharp, B. Rivard, S. Marshall, and D. Burgess, 2005. Melt season duration
on Canadian Arctic ice caps, 2000–2004, Geoph. Res. Lett., 32, L19502.
Wismann, V., 2000. Monitoring of seasonal snowmelt on Greenland with ERS scatterometer
data, IEEE Trans. Geosci. Remote Sens., 38(4), 1821-1826.

102

IPY OUTREACH, EDUCATION, AND POLAR SCIENCE DAYSRHIAN SALMONIPY International Programme Office Education and Outreach Coordinator

Abstract
Scientists from a wide range of disciplines and countries are becoming involved inoutreach activities that are unique to their project and reach a wide range ofaudiences. Quarterly IPY Science Days are also now being organised in order todevelop easy ways for the IPY Community to raise awareness of the polar regions
and their research. It is hoped that the Glaciodyn community will become involved
with the March Polar Day that will focus on 'Changing Earth'. This talk anddiscussion hopes to expand on some of these ideas to develop and outreachstrategy that works for the scientists involved.

103

A FIVE-YEAR RECORD OF SUMMER MELT ON EURASIAN ARCTICICE CAPSMARTIN SHARP1 AND LIBO WANG2
1

Department of Earth and Atmospheric Sciences, University of Alberta, Canada
Climate Research Division, Atmospheric Science and Technology Directorate,
Environment Canada, Canada
2

AbstractFive-year climatologies (2000-2004) and annual anomaly patterns of melt season
duration and dates of melt onset and freeze-up were derived for major ice massesin Svalbard, Novaya Zemlya and Severnaya Zemlya using enhanced resolution
backscatter data from QuikScat. Severnaya Zemlya has later melt onset, earlier
freeze-up and shorter melt seasons than the other areas. Latitude (Svalbard) andlongitude (Novaya Zemlya and Severnaya Zemlya) are stronger influences on thefive-year mean melt duration than elevation. Melt duration is generally greatest in
the west of each archipelago, likely due to the influence of warm ocean currents
offshore. 2001 was the longest melt season in all 3 archipelagos, while 2000 was
the shortest. Annual mean melt duration in each region was strongly correlatedwith the mean June + August air temperature at 850 hPa geopotential height(derived from the NCEP/NCAR Reanalysis). Integration of these results withpreviously published results from Greenland and the Canadian Arctic reveals that
2003 was the longest melt season in the 2000-2004 period across the Arctic as awhole, while 2002 was the shortest. Correlation of melt duration anomaly series for
19 discrete areas of the Arctic identifies 7 spatially coherent regions with common
patterns of variability in annual melt duration, and suggests the possibility ofteleconnections between some widely separated regions.

104

REGIONAL DISTRIBUTION OF SNOW ACCUMULATION ON NORTHWESTERN SPITSBERGEN GLACIERS, SVALBARD
IRENEUSZ SOBOTA AND MAREK GRZEŚ
Department of Cryology and Polar Research, Institute of Geography, N. Copernicus
University, Poland

The studies of the snow cover were made on the glaciers of the Kaffiøyra region in
the north-western Spitsbergen. The studies, carried out since 1996, are part of aprogramme which researches the mass balance of the selected glaciers of thisregion (Sobota 2003, 2005, 2007a, 2007b). The main goal of the research was toassess the water resources in snow and the accumulation part of their massbalance. The snow profiles were also used to analyse the selected physical andchemical properties of the snow cover. Spatial diversity of snow accumulation on
the glaciers and its changes resulting from the altitude were analysed. The studiesof the snow cover on the Waldemarbreen began in 1996, on the Irenebreen in
2001, while on the Elisebreen in 2005. In 2007 snow accumulation was alsostudied on Aavatsmarkbreen. All the studies carried out enabled the authors toassess the size and regional diversity of snow accumulation on the glaciers locatedin the north-western part of Spitsbergen.
The glaciers are located in the northern part of the Oscar II Land, Kaffiøyra,north−western Spitsbergen. Kaffiøyra is a coastal lowland situated on the2
Forlandsundet. Waldemarbreen is about 3.5 km long and has an area of 2.6 km .
The ice originates in one cirque and flows from an elevation of more than 500 m tothe present snout at 135 m a.s.l. Irenebreen, a valley glacier located to the south ofWaldemarbreen, flows down towards the Kaffiøyra plain. The area of Irenebreen
22amounts to 4.1 km . Elisebreen area is 10.2 km . Its length is about 7 km, while its
width is up to 1.8 km. To the north the glacier borders Agnorbreen which is often
treated as part of Elisebreen. Aavatsmarkbreen is the largest of all the studied
2
glaciers; its area is about 75 km . Its length is 15 km, while its width is from about 3
km in the mountain zone to 4 km in the ice cliff zone, which ends up in the sea.
The measurements were taken three times at every site, the accuracy of thereadings being ±1 cm. The basis for calculations and graphic presentation of
spatial variation of snow cover thickness resulted from 100 to 150 soundings. Snowpits were dug if a snow layer could not be penetrated through. The measurementsites were located with GPS and marked on the map. Snow density, structure,grain type and hardness values were recorded. The density of snow was measured
in pits and at representative points.
The research has enabled the authors to determine the spatial diversity of snow
accumulation on four glaciers located in the north-western part of Spitsbergen:
Waldemarbreen, Irenebreen, Elisebreen and Aavatsmarkbreen. All those glaciers
had a large spatial diversity of their snow cover thickness, and its distinct
changeability with the changing altitude. Accumulation values observed on the105glaciers were related to their absolute altitude a.s.l. (Fig. 1). The largest
accumulation gradient was recorded for the Elisebreen and the Aavatsmarkbreen.
An important role in forming the spatial diversity of snow accumulation belongs tosnow movement due to wind. The spatial distribution of the winter snow
accumulation on the glaciers of this region shows some rules observed every
winter season (Grabiec et al. 2006; Grześ and Sobota 2000, Sobota and Grześ
2006; Sobota 2003, 2005, 2007a).

Figure 1. Relation between the snow accumulation and altitude in 2007 on selectedglaciers.

The measurements of structure and graining of the snow cover were notundertaken during all of the analysed periods. Snow cover shows some specificphysico-chemical properties. Its vertical profile shows a variety of snow types ofdiverse level of metamorphosis, hardness and wetting. Snow structure reflects
prevailing weather conditions at the time when the snow cover formed. In theindividual years the snow cover of the studied glaciers was dominated by finegrained and medium-grained snow, while the layer above ice contained coarsegrained snow. Numerous ice layers were also found. Ice layers results proveseasonal warming above 0°C, which causes melting followed by freezing of snow
during the formation of the snow cover. Such a phenomenon is an important factor
that stops snow redeposition, as the frozen cover is resistant to wind and it protects
the underlying layers. Moreover, such ice layers deteriorate the conditions for thedeposition of fresh snow. The entire course of the phenomena, in general, isdependant on the weather conditions, and the frequency and the length of thewarming.
During all the years of study, the snow density on the researched glaciers at theindividual depths was similar, and it was directly proportional to the depth. During106most of the accumulation seasons a slight increase of mean snow density wasobserved with the decreasing altitude and the decreasing snow cover thickness.
In 2007 the spatial diversity of snow density of the uppermost layer on all theglaciers was analysed. Although is shows large changeability in individual places ofthe studied glaciers, an inverse proportion was recorded between the snow
density, the altitude and the size of the snow accumulation in a given place. The
snow density distribution is clearly related to the characteristic areas of the glacier:its snout and its accumulative zone (Fig. 2). The interrelation between the meansnow density on a glacier in a given year and the mean snow accumulation was
not recorded.

Figure 2. Maps of surface snow density on glaciers: Waldemarbreen, Irenebreen andElisebreen in 2007.

107

-3

The mean snow density on Waldemarbreen was 375 kgm in the years 1996-3
2007, on Irenebreen it amounted to 360 kgm in the years 2002-2007, while on-3
Elisebreen it was 385 kgm in the years 2005-2007. It was found out that the snow
density on the studied glaciers at the individual depths was similar, and directly
proportional to the depth.
It is characteristic that the regional diversity of snow accumulation on the
researched glaciers is not large and the mean winter accumulation is similar. For
instance, the mean snow accumulation on Waldemarbreen and Irenebreen wasvery similar in the years 2002-2007. However, all the analysed glaciers show alarge spatial diversity of snow accumulation on their surface. It is connected withboth the changing altitude and the local conditions on the glacier.

References:
Grabiec M., Leszkiewicz J., Głowacki P. and Jania J. 2006. Distribution of snow
accumulation on some glaciers of Spitsbergen. Polish Polar Research, 27(4): 309-326.
Grześ M. and Sobota I. 2000. Winter snow accumulation and winter outflow from theWaldemar Glacier (NW Spitsbergen) between 1996 and 1998. Polish Polar Research21(1): 19-32.
Sobota I. 2003. Warunki meteorologiczne i wybrane problemy akumulacji śniegu w regionie
Kaffiøyry (NW Spitsbergen) w okresie od lipca 2001 roku do kwietnia 2002 roku,
Problemy Klimatologii Polarnej 13: 239-249.
Sobota I. 2005. Struktura bilansu masy lodowców Kaffiøyry na tle lodowców Svalbardu.
Kaffiøyra. Zarys środowiska geograficznego Kaffiøyry (NW Spitsbergen), 43-60.
Sobota I. 2007a. Mass balance monitoring of Kaffiøyra glaciers, Svalbard. The Dynamic andMass Budget of Arctic Glaciers. Extended abstracts, Workshop and GLACIODYN (IPY)
Meeting, IASC Working Group on Arctic Glaciology, Utrecht University, 108-111.
Sobota I. 2007b. Selected methods in mass balance estimation of Waldemar Glacier,Spitsbergen. Polish Polar Research, 28(4): 249-268.
Sobota I. and Grześ M. 2006. Charakterystyka pokrywy śnieżnej na lodowcach Kaffiøyry(NW Spitsbergen) w 2005 roku. Problemy Klimatologii Polarnej 16: 147-159.

108

IDENTIFYING CAUSES OF
GREENLAND ICE SHEETPERIPHERALTHINNING

OF

THE

ANDREW SOLE1, TONY PAYNE1, JON BAMBER1, PETER NIENOW2 AND WILLIAMKRABILL3
1

Bristol Glaciology Centre, School of Geographical Sciences, Bristol, UK
School of Geosciences, University of Edinburgh, Edinburgh, UK
3
Cryospheric Sciences Branch, NASA Goddard Space Flight Centre
2

Introduction
Recent observations have shown that the periphery of the Greenland ice sheet isthinning rapidly and that this thinning is greatest around outlet glaciers. Twoprincipal mechanisms for climate induced dynamic thinning have been presented in
the literature.
The first suggests that enhanced ice melt caused by climatic warming is affecting
ice dynamics directly (Zwally et al. 2002). Longer and more widespread surface
melt may increase drainage to the bed through crevasses and moulins and thus
enhance ice flow through basal lubrication. The second mechanism is related tothe collapse of ice shelves and retreat of marine termini. Changes in flow of anoutlet glacier occur in response to changes in its force balance – the sum ofgravitational and resistive forces acting on the ice (Payne et al. 2004). As a floating
ice shelf retreats, changes in longitudinal stresses within the ice may cause its
velocity to increase and dynamic thinning to occur (Vieli et al. 2002). The initial
perturbation could be due to accelerated calving rates (Thomas 2004) caused byincreased basal melting of the ice shelf by warmer ocean water (Payne et al. 2005)
or increased surface melting resulting from higher air temperatures (Alley et al.
2005).

Methodology
By using PARCA laser altimetry data to extract thinning rates (dh/dt) for many
Greenlnad outlet glaciers, the relative importance of these mechanisms can betested using several hypotheses:
1. Thinning is not dynamic, but is related to local mass balance changes. If thiswere the case we would expect thinning to vary only with local mass balancei.e. with altitude and between climatic regions. This is a null hypothesis whichallows us to test whether ‘inland ice’ (slow flowing ice that is not obviously partof an outlet glacier) and outlet glaciers are thinning at different rates. To test thishypothesis we compare ice surface elevation change rates for inland ice (thecontrol flight line) with that of nearby outlet glaciers for two areas in Southern
Greenland.
2. Thinning is principally caused by direct meltwater effects on ice flow. In thiscase, dh/dt of land terminating and tidewater outlet glaciers should bestatistically the same. We test this hypothesis by comparing mean dh/dt for anumber of tidewater and land terminating outlet glaciers in Southern Greenland109

(Figure 1). This region was chosen because there are more tidewater and landterminating outlet glaciers with good flight line coverage. However, tidewateroutlet glaciers are likely to be more sensitive to meltwater input since theirtermini are already ‘floating’. Thus, an identical increase in surface meltwater
input to the bed would be more likely to raise subglacial water pressure tovalues approaching overburden pressure, with correspondingly greater ice
acceleration and dynamic thinning.
3. Thinning is principally caused by calving events of tidewater glaciers. In thiscase land terminating outlet glaciers should either not be thinning at all, or
thinning at a rate related to increases in the number of Positive Degree Days
(PDDs), while thinning rates of tidewater outlet glaciers should be significantly
greater than those expected from changes in air temperature. This hypothesis istested by comparing observed dh/dt for land terminating and tidewater outlet
glaciers with the melt expected from increases in air temperature using a
Positive Degree Day (PDD) Model (Braithwaite 1995).

Tidewater Data

Land Terminating MeanLand Terminating Data

Tidewater Inter -quartile Range

x Tidewater Outliers
x Land Terminating OutliersLand Terminating Inter -quartile Range
Jakobshavn Isbrae

Tidewater Mean

Kangerdlugssuaq Gletscher

Land Terminating Mean

Helheim Gletscher

Tidewater MedianFigure 1. A comparison of surface elevation change rates for a) 11 tidewater and 6 landterminating glaciers for the period 1993 to 1998. Red refers to tidewater outlets, while bluerefers to land terminating outlets. For each elevation band, both the actual data and a box
plot representing the data are shown. Each solid dot represents the mean dh/dt value for therespective elevation band for a particular flight line. The closed boxes represent the interquartile range which includes 50 % of the data, and the whiskers are the 5th and 95th
percentiles. Black, open shapes refer to specific outlet glaciers identified in the text; and b) 9
tidewater and 2 land terminating glaciers for the period 1998 to 2006.

Results and Discussion
Hypothesis 1: Thinning rates are significantly greater on outlet glaciers than onneighbouring areas of slow moving ice and suggest that ice dynamics does play animportant role in the peripheral thinning of the ice sheet. This is in agreement withmany other published results.

110Hypothesis 2: Figure 1 shows that in general, tidewater outlet glaciers have thinned
significantly more than their land terminating counterparts. Mean thinning rates for
tidewater glaciers below ~1000 m are four times larger (significant at the 99 %confidence level) from 1998 to 2006 than from 1993 to 1998. Mean dh/dt values forland terminating outlets, however, remain statistically unchanged over the sametime periods. This suggests that a change in a controlling mechanism specific to
the thinning rates of tidewater outlets occurred in the late 1990s and that thischange did not affect thinning of land terminating outlets. Moreover, the statisticallysignificant differences between mean dh/dt values for tidewater and landterminating outlet glaciers suggest that they can be seen as separate populations.

Figure 2. The percentage of observed thinning which cannot be explained by thinningcalculated by a PDD model for all surveyed glaciers. A zero % value means that for all
glaciers with data in that particular elevation band, the ablation anomaly thinning was
greater than observed thinning. Hollow joined symbols are mean thinning rates, refer toFigure 1 caption for meaning of other symbols.

Hypothesis 3: The percentage difference between dh/dt values predicted by the
PDD model and the observed dh/dt values are shown in Figure 2. For tidewateroutlet glaciers, values increase towards the terminus, reaching ~70 %, while forland terminating outlet glaciers, values are low (mean of ~10 %) and show no
relationship with elevation. It therefore seems likely that thinning of land terminating
outlet glaciers is principally controlled by variations in local mass balance, whilethinning of tidewater outlet glaciers is dependent on ice dynamics. The mechanism
by which this dynamic thinning occurs is still not clear although a good correlation
between calving front retreat and thinning rates (Joughin et al. 2008), as well as
the association of dynamic thinning with tidewater outlet glaciers, suggests that
perturbations at the termini of tidewater outlet glaciers are mainly responsible. It
seems likely that sea surface temperatures, decadal scale surface air temperature
increases and other factors such as glacier geometry, ice thickness and basal
topography must be combined to understand a tidewater glacier’s sensitivity to
climate change (Box and Herrington 2007).
Dynamic thinning related to calving will cease once the ice sheet retreats beyond
the direct influence of the ocean. These findings therefore have important111

implications for the future mass balance of the ice sheet and predictions of sea
level rise over the next few centuries.

References
Alley, R., Clark, P., Huybrechts, P., and Joughin, I., 2005. Ice-sheet and sea-level changes.
Science, 310, 456-460.
Box, J., and Herrington, A. 2007, Was there a 1930's Meltdown of Greenland Glaciers? EosTrans. AGU, 88(52), Fall Meet. Suppl., Abstract C11A-0077.
Braithwaite, R., 1995. Positive degree-day factors for ablation on the Greenland ice sheet
studied by energy-balance modelling. Journal of Glaciology, 41, 153-160.
Joughin, I., Howat, I., Alley, R., Ekstrom, G., Fahenstock, M., Moon, T., Nettles, M., Truffer,
M., and Tsai, V., 2008. Ice-Front Variation and Tidewater Behavior on Helheim andKangerdlugssuaq Glaciers, Greenland. Journal of Geophysical Research,doi:10.1029/2007JF000837, doi:10.1029/2007JF000837.
Payne, A., Holland, P., Feltham, D., and Jenkins, A. 2005, Modelling the coupled evolutionof the ice-shelf/stream flow system and the oceanic circulation in the ice-shelf cavity.
European Geosciences Union, Vienna.
Payne, A., Vieli, A., Shepherd, A., Wingham, D., and Rignot, E., 2004. Recent dramaticthinning of largest West Antarctic ice stream triggered by oceans. Geophysical Research
Letters, 31, doi:10.1029/2004GL021284.
Thomas, R., 2004. Force-perturbation analysis of recent thinning and acceleration of
Jakobshavn Isbrae, Greenland. Journal of Glaciology, 50, 57-66.
Vieli, A., Jania, J., and Kolondra, L., 2002. The retreat of a tidewater glacier: observations
and model calculations on Hansbreen, Spitsbergen. Journal of Glaciology, 48, 593-600.
Zwally, H., Abdalati, W., Herring, T., Larson, K., Saba, J., and Steffen, K., 2002. Surface
melt-induced acceleration of Greenland ice sheet flow. Science, 297, 218-222.

112

RELATIONS BETWEEN HORIZONTAL STRAIN AND ICE ELEVATIONCHANGE
MANFRED STOBER
Stuttgart University of Applied Sciences

Introduction
In a long-term project (1991 until 2006) two research areas, Swiss Camp and ST2,
are surveyed by geodetic methods in order to determine elevation change, ice flow
velocity and deformation of the ice surface. About elevation change and ice flow
were reported in Stober (2007) and Stober and Hepperle (2007). This paper isfocussed on deformation and strain and on the attempt in determination the net
mass balance from geodetic ground measurements.
The research area at Swiss-Camp (ETH/CU-Camp), established in 1991, issituated 80 km east from the West Greenlandic coastal village of Ilulissat, latitude
69°34’N, longitude 49°20’W, elevation 1170 m a.s.l near the equilibrium line
altitude. The network consists of 4 stakes, forming a triangle with a point in its
centre. The side length of the network is about 1.5 km. The 3D-positions of the
stakes are measured by GPS in relation to a fix point on solid rock at the coast. In
2004, the research area was extended by a new deformation network (ST2). ST2 islocated in latitude 69°30’28”N; longitude 49°39’09”W, ellipsoidal height 1000 m.
In case of homogenous strain the distortion of the moving network between two
campaigns can be calculated by a linear affine transformation with 6 unknowns ai,b i:epoch 1 --- epoch 0
X = a0 + a1·x + a2·y
(1)
Y = b0 + b1·x + b2·y
Our network has 4 stakes, so the equation system is over-determined and a leastsquares adjustment is applied, which gives also standard deviations in order tocalculate the significance level of the results.
With (1) the normal and shear strain respectively are calculated fromexx = a1 – 1[ ppm ]
eyy = b2 – 1[ ppm ]
exy = (a2+b1)/2[ ppm ]
Amount and azimuth of the principal strains e1, e2 are derived by
e=
e1 =
e2 =

(e xx − e yy )2 + 4 ⋅ e xy
e xx + e yy + e
2e xx + e yy − e

2

(3)

and

Θ1 =

2 ⋅ e xy1
⋅ arctan,2
e xx − e yy

Θ 2 = Θ1 ± 1002

Now we apply the incompressibility condition of ice:
€

€

(2)

113

gon

(4)

e1 + e2 + e3 = 0
and then the vertical strain component e3 becomes:
e3 = - ( e1 + e2 ).

(5)
(6)

Furthermore we normalize all values on one year and calculate time derivates toget strain rates, dimensions ppm/a. In the following part the same notation forstrain is now used for strain rates, too.
Together with the vertical strain rate e3 we can derive a strain induced elevation
change by:∆He = e3 H,(7)where H = ice thickness .
In reality e3 is not constant within the whole ice column, therefore modelling thevertical strain rate dependent on depth becomes necessary. Strain is caused byvelocity differences, and we assume that e3 is varying with depth in the same way
as the horizontal flow velocity Us. Based on Glen’s flow law the horizontal velocity
in different depths (Uh) becomes (Paterson, 1994):
n n+1
Uh = Us – 2A/(n+1)·(ρ·g·sin α) ·h ,
(8)
with:
Us = horizontal velocity at surface, (Swiss Camp Us = 115 m/a );h = depth, counted from surface h=0 until bedrock h=H;3
ρ = density = 900 kg/m ;2
g = acceleration due to gravity = 9,81 m/s
n = empirically determined constant, for ice sheet usually n = 3;
α = surface slope , Swiss Camp α = 1,14 °, sin α = 0,02
A = flow parameter, depending on ice age and temperature,
-16
-1
-3
A = 4,9·10 [s kPa ], for T = -10°C
-16
-1
-3
A = 2,9·10 [s kPa ], for T = -15°C.
A special investigation has shown that the depth dependent flow velocity is varying
only little from the velocity at the surface, and can be taken into account by ashape factor F, (assuming the ice temperature T = - 15 °C):
Swiss Camp:F = 0,930 ,ST 2:
F = 0,995 .
(9)Applying (9) to (7) the vertical elevation change due to strain finally becomes∆He = F·e3 H.
(10)
According to Patterson (1994), the ice thickness change follows from the massbalance equation and together with (6) and (10) becomes:
ΔHsurface = b - F·H· (e1 + e2)= b + F·H·e3= b + ΔHe
(11)
with b = net mass balance (in ice equivalent) at the research area.
The surface elevation changes were directly measured by GPS at the same stake
positions as in the former campaigns (Stober and Hepperle, 2007), therefore in114

(11) a term due to terrain slope is not included. With ΔHsurface = ΔHGPS thus the net
mass balance (b) follows by
b = ΔHGPS – Δhe .
(12)

Results
Table 1 and figure 1 show the resulting strain rates for Swiss Camp. Except of thecampaigns 1994-96, all strain rates are significant, indicating thinning due to ice
outflow. The net mass balance (in ice equivalent, m/a) varies between campaigns,
but is about zero for the long-term comparison 1994-2002.
Table 1. Strain rates and elevation/thickness changes at Swiss Camp 1994 – 2002.
AzimuthStrain rates
ΔHe=
ΔHGPS
Net mass
e3 * F15*Hbalance (b)EpochsΘ1 for e1
e1
e2
e3
[gon]
[ppm/a] [ppm/a] [ppm/a]
[m/a]
[m/a]
[ m/a ]
1994 - 1995
29,1
927
-922
-5
-0,01
-0,33
- 0,32
1995 - 1996
25,0
924
-846
-78
-0,08
-0,74
- 0,66
1996 - 1999
20,7
1217
-890
-326
-0,33
-0,09
+ 0,25
1999 - 2002
13,5
1222
-904
-318
-0,33
-0,23
+ 0,091994 - 2002
19,3
1118
-894
-224
-0,23
-0,25
- 0,02

Figure 1. Net mass balance at Swiss Camp derived from GPS and strain

The results for the other research area ST2 are shown in table 2 and figure 2. This
deformation network was established in 2004, so here we have only results from
2004, 2005 and 2006. Although the elevation changes by GPS are in the same
magnitude as at Swiss Camp, the vertical strain rates are much larger. The
resulting net mass balance (b) becomes positive with extreme high amount of+2,75 m/a ice equivalent.
Table 2. Strain rates and elevation/thickness changes at ST2 , 2004 - 2006.
AzimuthStrain rates
ΔHe=
ΔHGPS
Net mass
EpochsΘ1 for e1
e1
e2
e3
e3 * F15*Hbalance (b)[gon][ppm/a] [ppm/a] [ppm/a]
[m/a]
[m/a]
[ m/a ]
2004 - 2005
370,3
8548
-1951
-6597
-3,28
-0,39
2,89
2005 - 2006
369,8
8295
-2508
-5787
-2,88
-0,29
2,59
2004 - 2006
370,0
8445
-2243
-6202
-3,08
-0,34
2,75

115

Figure 2. Net mass balance at ST2 derived from GPS and strain.

Summary and acknowledgement
The combined evaluation of elevation changes measured by GPS and derived
from horizontal strain rates seem to be an appropriate tool for the determination of
the net mass balance at a certain area.
Modelling the vertical variability of flow velocity and strain rate show that at Swiss
Camp the reduction in respect of the surface values does not exceed 7%.
The data of ice depth at Swiss Camp (H = 1100 m) and ST2 (H = 500 m) werekindly provided by the Alfred-Wegener-Institut, originally probably measured by the
University of Kansas. The author thanks Dr. Andreas Ahlstrøm (GEUS/
Copenhagen) for intensive discussions into this topic.

References.
Paterson, W.S.B (1994): The Physics of Glaciers, 3. Edition, Pergamon Press, New York,
1994.
Stober, M. (2007): Geodetic investigations on dynamic parameters of the West Greenland
inland ice. The Dynamic and Mass Budget of Arctic Glaciers. Workshop and
GLACIODYN (IPY) meeting, IASC working group on Arctic Glaciology. Pontresina 15 –18 January 2007, Institute for Marine and Atmospheric Research Utrecht, Utrecht
University, 2007.
Stober, M. and Hepperle, J. (2007): Changes in Ice Elevation and Ice Flow-Velocity in the
Swiss Camp Area (West Greenland) between 1991 and 2006. Polarforschung 76 (3), p.
109 – 118, 2006, released 2007.

116

RECENT SURGE ACTIVITY IN SOUTHERN SPITSBERGENMONICA SUND1,2, TROND EIKEN2, JON OVE HAGEN2, ANDY KÄÄB2, THOMAS
VIKHAMAR SCHULER2*, KARSTEN MÜLLER2, GEIR MOHOLDT2 AND CHRIS NUTH2
1

University Centre in Svalbard
Department of Geosciences, University of Oslo*Now at the Norwegian Water Resources and Energy Directorate
2

Introduction
Surges are periods of extraordinarily enhanced glacier dynamics and show cyclic
behaviour. Surges are believed to be related to internal changes in the glaciersystem rather than driven by external causes like changing climate (Meier andPost, 1969). However, the reoccurrence interval of surges is largely controlled byclimate conditions (Harrison and Post, 2003). In Svalbard, surges constitute a
common form of glacier advance (Liestøl, 1969). Glacier surges in this region are
observed to have longer quiescent and active surge phases and relatively low
surge velocities compared to other regions with surging glaciers. Surges in
Svalbard typically have an active phase lasting up to several years and a quiescentphase of 30 to 500 years (Dowdeswell, et al., 1991). The surge mechanism itself isstill object of investigation and several explanations have been proposed (Clarke,
et al., 1984; Kamb, et al., 1985).
In particular, the initiation of surges is poorly studied and understood, due to thedifficulty of recognizing a surge at an early stage when there is little visual evidence
such as surface crevassing. The occurrence of crevasses in previously undisturbed
areas is usually the first indication of a surge initiation (Hagen, et al., 1993).
Currently there are indications that the initiation of surges on Svalbard may be aprocess lasting several years. In this contribution surface changes related to surge
behaviour are presented for several glaciers in southern Spitsbergen. One ofthese, the ca. 9 km long valley glacier Kroppbreen in Kjellströmdalen is subject to amore detailed study.

Methods
Photos obtained on regular flights to and from Longyearbyen over several years
are used in the visual search for changes of glacier surfaces. These photos and
ASTER satellite images were compared to new and old maps and to 1990 aerial
photos by the Norwegian Polar Institute (NPI) in order to detect new crevasses.
Comparison of DTMs of 1990 by NPI, of 2003 from ASTER satellite stereo and of
2006 by Store Norske Spitsbergen Kullkompani (SNSK) is used to reveal large
mass displacements. The ASTER-derived DTM was compiled using
photogrammetric methods within the PCI Geomatica software (Kääb, 2002; Toutin
2002). For Kroppbreen the 2006 DTM is based on airborne LIDAR scanning,
except for a small area near the front which is compiled from aerophotogrammetry
for this purpose. ICESat data from 2003 – 2007 are also used for comparison to
the 1990 DTM.

117

In April 2007 a kinematic Global Navigation Satellite System (GNSS) profile wasmeasured along the centreline. Six stakes were also established along a
longitudinal profile and measured with relative GNSS. These stakes wereremeasured after 25 days and six additional stakes including a cross profile were
drilled and surveyed as well. By middle of August 2007 all stakes wereremeasured.
Ground Penetrating Radar (GPR) profiles were retrieved using a MALÅ 50 MHz
antenna. The GPR data were sampled with a fixed time interval of 1s. The
measurement interval along the transects was dependent on the snowmobile
speed, ~10 km/h, and is in average about 3 m.

Results and discussion
Here we consider changes of the glaciers in the period 1990 – 2007. From the
ASTER images it appears that Ingerbreen, confluencing with Richardbreen in thelower part, was in the initial stage of a surge in 2001 while the whole surface washeavily crevassed in 2005 and the surge was in the terminating phase. On several
glaciers new crevasses are found on the photos taken from the scheduled flights,
when compared with 1990 aerial photos.
• Mendelejevbreen showed a heavily crevassed surface from the ice divide in
2001. A drawdown of 40-50 m is found in the upper basin (Fig. 1).
• Skilfonna has a drawdown somewhat downglacier of the divide and marginal
crevasses along the glacier, also transversal crevasses are observed. The
major changes probably appeared during 2006.
• Zawadzkibreen experienced increased crevassing extending almost to theterminus. There are elevation losses of up to 20 m in the upper basin, withincreases of ~5-10 m lower down (Fig. 1).
• On Polakkbreen there are some crevassed areas. A drawdown of up to 40
meters is experienced in the upper basin while slight thickening is seen at theconfluence with Nathorstbreen (Fig. 1).
• Dobrowolskibreen has crevasses with enhanced marginal crevassing over theentire glacier, but not as heavy as that observed on Skobreen in 2005.

Figure 1. ICESat data (2003-2007)
compared with the 1990 DTM. Both
Zawadzkibreen
andPolakkbreen
experience elevation losses in the upper
part, while an increasing elevation isobserved in the lower part, especially on
Zawadzkibreen. Mendelejevbreen shows
elevation losses in the upper regions.

118

On Kroppbreen, we found a build up in the upper part and surface lowering of thelower part over the period 1936 – 1990. For the period 1990 – 2006 a surface
lowering in upper part and thickening in central part is observed (Fig. 2a). This issimilar to the development of Skobreen in the period 1990 – 2003 (Fig. 2b), which
had proceeded to a fully developed surge (Sund, 2006). Similar changes have also
been mapped on Hessbreen (Liestøl, 1976) prior to its surge (Sund and Eiken,
2004). Kroppbreen has developed large crevasses in the upper part since 1990,
while smaller, and marginal crevasses characterize the surface of the lower part.
Measurements of stake displacement indicate that currently the velocity magnitude
are greatest where the ice has thickened and further indicate block flow across thetransversal profile. GPR measurements on Kroppbreen strongly indicate that thelowermost part of the glacier consists of cold ice.
On Kroppbreen, Zawadzkibreen and Polakkbreen the changes cannot yet beobserved at the terminus. This indicates that during the development of the surge
the mass displacement is limited to a part of the glacier only
On Mendelejevbreen, Ingerbreen, Skobreen and Skilfonna the surge activities
have reached all the way to the terminus. It is also observed that some glaciers,
like Søkkbreen and Knoppbreen, previously only underwent a partial surge thatdeclined downglacier. On Mendelejevbreen, Ingerbreen, Skobreen and Skilfonna
the surge activities have reached all the way to the terminus. It is also observedthat some glaciers, like Søkkbreen and Knoppbreen, previously only underwent a
partial surge that declined downglacier.

Figure 2. a) Surface elevation changes of Kroppbreen over 1990-2006 showing a loweringin the upper part and increase of ice thickness in the central part. b) Surface elevation
changes of Skobreen over 1990-2003 reveal a lowering of the upper part and a mass
displacement to the lower part at the junction to Paulabreen.

119

Concurrent to enhanced crevassing, also a surface lowering in the upper part andthickening in the lower part is found at several of the glaciers. GPR and velocity
data from Kroppbreen indicate that cold ice at the front is damming up thetemperate ice upglacier. There may be also a possible effect on subglacial waterdrainage, such that thermal damming of water at the front eventually leads toincreased subglacial water pressures in the area upglacier. Increased subglacial
water pressure in turn lead to an enhanced sliding motion of the glacier (e.g. Iken
and Bindschadler, 1986)
We have detected changes in the surface texture of several glaciers on Svalbard
during recent years (since ~2000). A comparison of elevation data revealedconsiderable mass displacements downglacier on these glaciers relative to 1990.
Recently enhanced crevassing of previously undisturbed areas indicates changes
in the dynamics of the respective glaciers. The downglacier thickening in
conjunction to thinning of the upglacier areas as observed at Kroppbreen,
Zawadzkibreen and Polakkbreen are typical for geometry changes during a surge.
Based on the experience from previous surges of Hessbreen and Skobreen, weinterpret our observations at a number of glaciers as indicative of an early phase of
a possible surge activity. Some glaciers have shown only a partial surge before theactivity declined. We believe that not all surges have to be large and spectacular,
but mass displacements still can have characteristics of a surge. This implies that
there is a broad range of surge-like behaviour governing the dynamics of Svalbard
glaciers.

Acknowledgement
Store Norske Spitsbergen Kullkompani (SNSK) provided 2006 DTM and aerial
photos. Lufttransport AS helped with recognisance. S. A. K. Stene, K. Høgvard andT. Svenøe contributed to the fieldwork.
The work was also financially supported by the Jan Christensen Endowment.

References
Clarke, G. K. C. , Collins, S. G. and Thompson, D. E. 1984. Flow, thermal structure andsubglacial conditions of a surge-type glacier. Canadian Journal of Earth Science. 21(2),232-240.
Dowdeswell, J. A. Hamilton, G. S. and Hagen J. O. 1991. The duration of the active phaseon surge-type glaciers: contrasts between Svalbard and other regions. Journal of
Glaciology 37 (127), 388-400.
Hagen, J. O., Liestøl, O. Rolan, E. and Jørgensen, T. 1993. Glacier Atlas of Svalbard andJan Mayen. Norsk Polarinstitutt Meddelelser 129.
Harrison, W. D. and Post, A. S. 2003. How much do we really know about glacier surging?
Annals of Glaciology 36, 1-6.
Iken, A. and Bindschadler, R.A. 1986. Combined measurements of subglacial waterpressure and surface velocity of Findelengletscher, Swizerland: conclusions about
drainage system and sliding mechanism. Journal of Glaciology, 32(110): 101-119.
Kamb, B. and 7 others. 1985. Glacier surge mechanism: 1982-1983 surge of VariegatedGlacier, Alaska. Science 227 (4686), 469-479.
Kääb, A. 2002. Monitoring high-mountain terrain deformation from repeated air- andspaceborne optical data: examples using digital aerial imagery and ASTER data. ISPRS
Journal of Photogrammetry & Remote sensing 57(1-2), 39-52.

120

Liestøl, O. 1969. Glacier surges in West-Spitsbergen. Canadian Journal of Earth Science 6,895-897.
Liestøl, O. 1976.Glaciological work in 1974. Norsk Polarinstitutt Årbok 1974. 191-193.
Meier, M. F. and Post, A. 1969. What are glacier surges? Canadian Journal of Earh
Sciences 6, 8907-817.
Sund, M. 2006. A surge of Skobreen. Polar Research 25(2), 115-122.
Sund, M and Eiken, T. 2004. Quiescent –phase dynamics and surge history of a polythermal
glacier: Hessbreen, Svalbard. Journal of Glaciology 50 (171), 547-555.
Toutin, T. 2002. Three-dimensional topographic mapping with ASTER stereo data in rugged
topography. IEEE Transactions on geoscience and remote sensing 40(19), 2241-2247.

121

GLACIOLOGICAL AND METEOROLOGICAL OBSERVATIONS IN
SUNTAR-KHAYATA RANGE, EASTERN SIBERIA, IN 2004-2007
SHUHEI TAKAHASHI1, KONOSUKE SUGIURA2, TAKAO KAMEDA1, HIROYUKIENOMOTO1, YURY KONONOV3 AND MARIA D. ANANICHEVA3
1

Kitami Institute of Technology, Koencho 165, Kitami 165, Japan
JAMSTEC, IORGC, Natsusima-cho, Yokosuka, Japan
3
Geographical Institute, Russian Academy of Sciences, Moscow, Russia
2

IGY and IPY activities in Suntar-Khayata
In the IGY (International Geophysical Years: 1956/57), the Russian Academy has
made extensively glaciological researches in Suntar-Khayata Range in Eastern
Siberia, in which about 180 glaciers were numbered. Especially No. 31 Glacier was
precisely studied; a meteorological station was constructed and lots of glaciological
data were acquired through three whole years. After the IGY, however, almost no
observations were done in this area except getting aerial-photo and satellite
images.
In 2004-2005, meteorological observations were done in this area as an activity of
IPY (International Polar Year). Meteorological instruments were installed at theformer IGY station at the terminus of Glacier No. 31 in Suntar- Khayata Range andat several points around Oimiyakon area. In 2006-2007, also 13 thermometers
were placed along Kolima Street from Magadan to Oimiyakon to study thetemperature distribution around this area (Fig. 1).

No.31 Glacier

Figure 1. Trip routes around SuntarKhayata Range and No.31 Glacier in Suntar
Khayata.

122

2004-2005 Activities around Suntar-Khayata
In July to August, 2004, Meteorological instruments (temperature, wind speed,
wind direction, solar radiation) were installed at Oimiyakon (Fig. 2), and at No. 31
Glacier in Suntar Khayata (Fig. 3). 13 thermometers were set along a road from
Oimiyakon to Tomtor and around No. 31 Glacier to observe temperature
distribution.
The meteorological instruments installed in 2004 were recovered in September
2005. Almost all instruments survived, though several thermometers along the roadbetween Onimiyakon and Tomtor were lost.

Figure 2. Memorial tower of Coldness Polein Oimiyakon The plate explains that -72 C
was recorded in 1924.

Figure 3. Meteorological station at No.31
Glacier. Thermometer, anemometer, solar
radiation meter and snow-depth camera
were installed.

The minimum temperature over a year was -59°C at Oimiyakon (about 680 m
a.s.l.), where it is called “Pole of Cold”, and -45°C at Glacier No. 31 (about 2050 ma.s.l.) (Fig.4), which suggests there was strong temperature-inversion in this areain the period of Siberia high pressure in winter.

Figure 4. Temperature at Oimiyakon (P1-P6) and No.31 Glacier (G0-G3). The minimum
temperature was -59.2°C (P1) in Oimiyakon area and -45.0°C (G1) at No.31 Glacier area.

123

80

Snow depth

Snow depth (cm)70

at No. 31 Glacier,in Suntar Khayata

6050

2004/11/01

2005/02/24

40
30
20
10

2005/05/01

0
8/7

9/18

10/30

2004

12/11

1/22

3/5

4/16

5/28

7/9

8/20

2005

Figure 5. Snow accumulation by interval camera. Snow accumulation variation (left) andphotos of a snow stake (right). Accumulation period was 2004/09/03- 2005/06/27 and the
maximum depth was 72 cm on 2005/05/11.

Snow depth was observed at 3-hour interval digital images of snow stakes, bywhich the variation of snow accumulation was obtained (Fig. 5). By this data, snow
accumulation period was 2004/09/03-2005/06/27 and the maximum snow
accumulation was 72 cm on 2005/05/11.
A feature of the snow accumulation is that the snow depth increased in October, atthe beginning of winter, was almost zero in winter, and increased again in the endof winter. It can be explained as water vapour could not be supplied during theSiberia high-pressure in winter but before/after the winter the atmosphere
disturbance brought water vapour to this area.

2006-2007 Activities along Kolima Street
In September in 2006, thermometers and simple rain-gauges were set along a way
from Magadan to Oimiyakon (Kolima Street) at interval of about 100 km (Fig. 6). To
observe sea ice development at the north end of Okhotsk Sea, a interval camera
was set at Magadan sea shore.

Figure 6. Boarder of Sakha
Republic and Kolima aria on
a way from Magadan to
Oimiyakon

124

Figure 7. Location of thermometers and the minimum temperature of each place [P] means
thermometer was set at plain, [H] at hill, [V] at valley and [Sea] at sea shore.

In Fig. 7 the location of thermometers and the minimum temperature of each placeis shown. The minimum temperature in 2007 was below -50°C in inner region and 59.7°C at Oimiyakon, while higher at coastal region, -23.2°C at Magadan. From therecords of temperature, the temperature at Oimiyakon was below -50°C from
December to February, but higher than that of Magadan in summer.

Results
In Oimiyakon area, strong temperature inversion appears in winter. The minimum
temperature was -59 oC at Oimiyakon (about 680 m a.s.l.) and -45 oC at GlacierNo. 31 (about 2050 m a.s.l.). Along a route from Magadan to Oimiyakon, thecoldest place was Oimiyakon. The minimum temperature was -59.7 oC at
Oimiyakon, while -23.2 oC at Magadan.

125

PRELIMINARY RESULTS OF GIS-BASED SOLAR
MODEL FOR HORNSUND AREA, SW SPITSBERGENRADIATION

M. SZYMANOWSKI1, M. KRYZA1, K. MIGAŁA1, P. SOBOLEWSKI2 AND L.
KOLONDRA3
1

University of Wrocław
Institute of Geophysics Polish Academy of Sciences
3
Silesian University
2

Introduction
GIS-based, solar radiation model is a useful tool to evaluate diversity ofglaciological, geomorphologic and biological processes. Results in elementary form
allow to recognize the local climate diversity and its ranges or to identify cold/warm
sub-areas formed by radiation balance and affected by relief. The model can also
be applied to identify local surfaces with daily thaw/freeze cycles.

Model outlook
The “r.sun” model is applied here to calculate daily sums of direct, diffuse and total“clear-sky” solar radiation (direct + diffuse) (Hofierka 1997, Šuri and Hofierka
2004). The “r.sun” is one of the most developed solar models, which eliminates
numerous drawbacks existed in the well-known commercial modules. The model is
implemented in the GIS-GRASS system, that is Open Source GIS software used in
the project.
The “r.sun” can be used in two modes: (1) to calculate the instantaneous radiation,and (b) to calculate daily/monthly sums of radiation. The first mode is used tovalidate the model against the measurements. The second mode is used here to
calculate the spatial patterns of the solar radiation for the study area.

Input data
The model requires only a few mandatory input parameters:• Digital elevation model (DEM; 10m resolution) based on the source data of theorthophotomap 1:25 000 „Werenskioldbreen and surrounding areas” (© Norsk
Polarinstitutt and Silesian University). DEM, besides elevation, allows to
calculate surface inclination, orientation and shadows;• Astronomical factors (solar constant, declination, latitude, solar hour angle);• Atmospheric attenuation (scattering, absorption) by gases in clear and dry(Rayleigh) atmosphere (relative optical air mass and optical thickness) and bysolid and liquid particles (water vapour and aerosols) described by Linke
turbidity factor (LTF). The LTF expresses a relation between the beam anddiffuse radiation – the higher the LTF, the higher the diffuse component of thetotal radiation flux.
The obtained results (model) are planned to be validated against actinometrical
data form the Polish Polar Station, Hornsund.
126

Results
As an example, few details can be introduced basing on the monthly mean total
solar radiation in June (Figure 1):• Not evident differentiation of solar energy income on the glaciers, ranges only2
between 300 and 350 W/m , e.g. Hans Glacier (H) and Werenskiold Glacier(W).
• Evidently colder slopes, predisposed for long lasting snow cover (NW – NE
2
slopes with monthly mean total radiation < 250 W/ m )• Evidently colder, southern-orientated, small and narrow valleys, e.g. Arie Valley2
with cold-ice glacier, with mean monthly total radiation flux: 200-300 W/m(marked with white circle on the Figure 1).

Figure 1. Mean monthly total solar
2
radiation flux [W/m ] under clear sky
conditions in June (Wedel Jarlsberg
Land, SW Spitsbergen): H – Hans
Glacier, W – Werenskiold Glacier, whitecircle – Arie Valley.

Spatial distribution of solar energy income in the close vicinity of Polish Polar
Station (Hornsund) modelled in hourly intervals for mid-August, leads to theconclusion that “r.sun” model gives a possibility to:• Indicate thaw/freeze cycles zones and zones with more intensive weathering;• Identify glacier’s parts with different dynamics of daily outflow of ablation waterand parts with decreased ablation;• Simulate 24–hour cycle of water outflow intensity.
It should be stressed that multi-year daily mean of air temperature at the Polish
Polar Station varies in the first half of August between 4,4°C and 5,2°C; with mean
daily minimum between 0,3°C and 3,4°C (1988-2007). The analysis of hourly
radiation indicates that freeze/thaw cycles on the southern slopes occur frequently(Figures 2 and 3). In the mid-day under clear sky conditions, southern slopes of theridges (Ariekammen and Fugleberget) receive 50-100% more solar energy (65022750 W/m ) if compared with the coastal areas (350-450 W/m ). Outgoing radiation
starts to dominate on the slopes after 18:00 LSoT, while the coastal plain still2
receives 150-200 W/m (Figure 3). The process extends on all the area after 20:00
LSoT. It means that all processes concerned with heating, water flow from melting
snow patches and from the melted active layer of permafrost etc. are reduced to127minimum. In consequence, a short-time freezing starts, in higher elevations mostly.
Not before 04:00 LSoT solar radiation starts to warm the coastal plain with the2intensity of c.a. 50 W/m .

Figure 2. Orography in the vicinity of Polish Polar Station, Hornsund (left) and hourly mean
2
values of beam solar radiation flux [W/m ] under clear sky conditions at 13:00 LSoT,15 ofAugust (right).

2

Figure 3. Hourly mean of beam solar radiation flux [W/m ] under clear sky conditions at
18:00 LSoT (left) and at 04:00 LSoT (right), 15 of August.

References
Hofierka, J., 1997. Direct solar radiation modelling within an open GIS environment.
Proceedings of JEC-GI'97 conference in Vienna, Austria, IOS Press Amsterdam: 575584.
Šuri M, Hofierka J., 2004. A New GIS-based Solar Radiation. Model and Its pplication to
Photovoltaic Assessments, Transactions in GIS 8 (2): 175-190

128

OBSERVATION

AND MODELLING OF THE SURFACE ENERGYBALANCE IN THE ABLATION ZONE OF THE WEST GREENLAND ICESHEETMICHIEL VAN DEN BROEKE, PAUL SMEETS, JANNEKE ETTEMA, RODERIK VAN DE
WAL, JOHANNES OERLEMANS
Institute for Marine and Atmospheric Research Utrecht, Utrecht University

The surface energy balance (SEB) determines the surface temperature Ts of any
natural surface as well as the amount of melt of snow and ice surfaces, when Ts
has reached the melting point. The SEB is given by:
M = Rnet + SHF + LHF + Gs

(1)

In which M is the energy available for melting, Rnet = SWnet + LWnet is net absorbed
radiation, i.e. the sum of net shortwave radiation and net longwave radiation, SHF
and LHF are the turbulent fluxes of sensible and latent heat, respectively, and Gs is
the subsurface heat flux evaluated at the surface.

Figure 1. Locations of the
AWS used in this study.

129

Four years (August 2003-August 2007) of meteorological observations along the Ktransect, a mass balance stake array along the 67°N latitude circle in west
Greenland, have been collected using three automatic weather stations (AWS).
The AWS are located at 6, 38 and 88 km from the ice sheet margin at elevations of490, 1020 and 1520 m asl (Fig. 1). This part of Greenland is characterized by anapproximately 150 km wide tundra, an ablation zone that extends 100 km onto theice sheet, and an equilibrium line at 1500 m asl. After elaborate data quality checks
and corrections, the hourly AWS data are fed into a surface energy balance (SEB)
model that iteratively solves for surface temperature Ts. If Ts rises above themelting point of ice, it is reset to 0°C and excess energy fluxes are invested in ice
melting (M).

Figure 2. Observed andmodelled
surfacetemperature Ts for a 2005
winter month (top) andsummer month (bottom) at
S6.

To check model performance, modelled Ts is compared to an observed value
derived from outgoing longwave radiation measurements, using Stefan
Boltzmann’s law and assuming the surface to have unit longwave emissivity.
Figure 2a shows the good agreement that is found for a 2005 winter month at S6,with differences between observed and modelled Ts generally smaller than 1 °C, inspite of interdiurnal temperature changes in excess of 25 °C. In summer, nocturnal
longwave radiative losses are almost always too great to sustain nocturnal melting130

at S6, and surface refreezing occurs. Fig. 2b shows that this nocturnal refreezing of
the surface is well captured by the SEB model, as is the daily melt duration at S6.
Now that the SEB has been validated, we can confidently study its various
components. Figure 3 gives an example for a winter and summer month at S5. In
the absence of shortwave radiation and melt, and with moisture concentrations andhence LHF generally small, the main wintertime SEB is a balance between
negative LWnet and positive SHF, i.e. longwave cooling of the surface iscompensated by a downward turbulent transport of sensible heat, heating thesurface and cooling the air. Under clearsky, strong (katabatic) wind conditions, this-2
sensible heat transport easily exceeds 100 W m , remarkably large values under
statically stable conditions that can only be maintained by surface radiative coolingin combination with strong wind shear.

Figure3.
Modelled
components of the surface
energy balance (SEB) at S5
for a 2006 winter month (top)and summer month (bottom).

Under strongly non-stationary conditions, e.g. when cloudiness suddenly increases
and longwave cooling ceases, the subsurface heat flux becomes important totransport heat away from the surface, as happens on 19 February 2006. When, on
the other hand, under clearsky conditions the sensible heat flux ceases, forinstance due to a lack of wind, the subsurface heat flux starts transporting heattowards the surface from the deeper ice/snow layers, to compensate the longwave
cooling. Only when temperatures rise high enough, can sublimation (green line)
become a significant surface heat loss in wintertime.
131A very different SEB is calculated for the summer (Fig. 3b). SHF is still directed
downwards, because the surface cannot further increase its temperature. Because
the melting surface cannot further increase its water vapour pressure either, LHF is
now also positive. Daytime Rnet is dominated by absorption of shortwave radiation,
longwave cooling occurs at night. The sub surface heat flux Gs is small, in reaction
to small temperature gradients in the ice. The result is that the melt flux M attains-2
considerable values in excess of 300-400 W m at daytime, with a very
pronounced daily cycle. For most of the nights, however, summertime melting at
S5 continues. These results will be used to validate regional climate models. Onceagreement is sufficient, these models will be used to calculate melt over the entire
Greenland ice sheet.

Acknowledgements
We thank IMAU technicians Wim Boot, Henk Snellen and Marcel Portanger fortechnical and computer support and Carleen Reijmer for maintaining the IMAUAWS website. This work is funded by Utrecht University and the Netherlands Arctic
Program (NAP) of the Netherlands Organisation of Scientific Research, section
Earth and Life Sciences (NWO/ALW).

132

LARGEAND RAPID VARIABILITY IN THE VELOCITY IN THE
ABLATION ZONE OF THE GREENLAND ICE SHEETR.S.W. V.D. WAL, W. BOOT, M.R. VAN DEN BROEKE, P. SMEETS, C.H.
REIJMER, J.J.A. DONKER AND J. OERLEMANS
Institute for Marine and Atmospheric Research Utrecht, Utrecht University

Abstract
GPS observations show very rapid and large fluctuations along the K-transect in
the ablation zone of the Greenland ice sheet. The velocity of the ice sheet nearly
instantaneously reacts on increased production of melt water, and velocities
increase with a factor four within days, which is much stronger and much fasterthan previously reported. The increase coincides with increased ablation rates
indicating that surface water penetrates through ice layers of 500-1000 m thickwithin days, which is faster than expected from glaciological theory. At the bottom
increased water pressure causes acceleration of the flow. On a longer time scale,
15 years, velocities decrease only slightly, implying that the englacial system
adjusts very rapidly to the variable water input in the system maintaining a more orless constant ice flux. Here we show that despite the rapid fluctuations on shorttime scales, mean velocities are driven by slow changes in slope and thickness.

133