The Patagonian Icefields

Geographic setting

Patagonia, between ~40°S to 56°S, is the most southerly part of the South American continent. The landscape of this region is one of contrasts. Dense temperate rainforests cover the western coast, whereas the eastern plains are flat, vast, and arid. Perhaps most striking, however, are the high, Patagonian Andes, which rise steeply (up to around 4000 m asl) above deeply carved glacial fjords and valleys, and are home to the Patagonian Icefields.

The rugged, and densely vegetated west coast of Patagonia (Pacific fjords in distance). The high, ice-covered Andes are shown in the foreground. Photo: Miguel Vieira Wikimedia Commons.

Climate of Patagonia

Patagonia is one of the windiest and wettest places on Earth. The region has a temperate maritime climate, with a strong west-east precipitation gradient as a result of the year-round passage of westerly winds over the Patagonian Andes(1,2). On the west coast, annual precipitation reaches up to 7,500 mm year, whereas less than 1,500 mm a year falls east of the icefields(3,4). The snowfall brought by the westerly winds are the main source of accumulation for the glaciers of this region.

Windswept Nothofagus Antarctica tree, Ushuaia, Tierra del Fuego, Patagonia, Argentina. Photo: Leonardo Pallotta, Wikimedia Commons.

The Patagonian Icefields

While the Patagonian Andes are home to hundreds of small caps and valley glaciers(5,6,7,8), most ice is locked up in three large icefields, called the North, South, and Cordillera Darwin Icefields (see map below). Together, these icefields contain around 5,500 gigatons of ice, enough to raise global sea level by around 15 mm if completely melted(9).

The three major Patagonian Icefields, which occur south of ~46°S, are the largest expanse of ice in the Southern Hemisphere outside Antarctica(10). While large, today’s icefields are the remnants of a much larger Patagonian Ice Sheet, which formed during the global LGM around 21,000 years ago.

The icefields and glaciers of Patagonia. Major outlet glaciers are labelled. Glacier outlines from Pfeffer et al. (2014). Copyright J. Bendle

Dynamic outlet glaciers

Outlet glaciers of the Patagonian Icefields are fed by heavy snowfall (up to 10,000 mm per year) in the accumulation area, and have high melt rates at lower elevations. As a result, they have high mass balance gradients and are very dynamic(10), especially on the western coast snowfall is highest(11).

Heavy cloud cover and snowfall over Monte San Valentín (4058 m asl) in the accumulation zone of the North Patagonian Icefield. Photo: Murray Foubister Wikimedia Commons.

North Patagonian Icefield

Size, elevation, and structure

The North Patagonian Icefield (NPI) stretches over 100 km from 46°30’S to 47°30’S (see map above). Its current size is estimated at ~3,700 km2(8,12).

The NPI is drained by 29 main outlet glaciers (>10 km2), with glacial tongues that terminate in different settings(7). 11 (38%) outlet glaciers terminate on land, 17 (59%) terminate in lakes (e.g. Glacier Exploradores), and only one (3%), Glacier San Rafael, terminates in the sea.

The average elevation of NPI glaciers is around 400 m lower on the western side of the icefield (1240 m asl) than on the east (1640 m asl), reflecting the higher rates of snowfall in the west(8).

The North Patagonian Icefield seen from space. Image: NASA.

Glacier Flow speeds

The fastest-flowing glacier of the NPI is Glacier San Rafael (NASA Fig). This glacier is flowing at 7.6 km per year (around 20 metres per day11). In total, three NPI glaciers flow faster than 1 km per year (San Rafael, San Quintín, and Colonia11).

Outlet glacier flow speeds calculated from satellite imagery. The fastest-moving glaciers (shown by green and yellow colours) terminate in the Pacific Ocean fjords and embayments, or in large glacial lakes (e.g. Glacier Upsala). Image: NASA (based on data from Mouginot and Rignot, 2015)

South Patagonian Icefield

Size, elevation, and structure

The South Patagonian Icefield (SPI) is the largest of the South American icefields. It stretches over 350 km from 48°20’S to 51°30’S, and its current size is estimated at ~12,200 km2 (around three times larger than the NPI8).

The SPI is drained by 53 main outlet glaciers (>10 km2) that are generally larger than those of the NPI. The SPI has a similar amount of lake-terminating glaciers (59%), but more (17 or 32%) marine-terminating glaciers(7).

The South Patagonian Icefield seen from space. Image: NASA

Glacier Flow speeds

The SPI also has some of the fastest flowing glaciers in the world (see glacier flow speed diagram above). Of the ten fastest SPI outlet glaciers, which all move faster than 2.5 km per year, eight flow out into Pacific Ocean fjords (e.g. the Jorge Montt glacier shown below) and two into large glacial lakes (e.g. Glacier Upsala). The fastest, Glacier Penguin, flows at 10.3 km per year (around 28 metres per day11). These data suggest that ocean heat plays an important role in melting at glacier fronts, and rapidly drawing down ice from the SPI interior.

The Jorge Montt outlet glacier of the Southern Patagonian Icefield, flowing out into a Pacific Ocean fjord that is choked with icebergs. Image: NASA.

The Upsala outlet glacier of the Southern Patagonian Icefield, flowing into a iceberg filled proglacial lake (Lago Argentino). Image: NASA

Cordillera Darwin Icefield

Size, elevation, and structure

The Cordillera Darwin Icefield (CDI) is the smallest (~2600 km2), and southernmost of Patagonia’s icefields(13), existing between 54°40’S to 55°00’S. It is also the least studied. Most CDI glaciers descend to sea level, and many flow into the Pacific Ocean(8).

Glacier Flow speeds

Most CDI outlet glaciers have flow speeds of between 1 and 3 metres per day(13). However, the marine-terminating Marinelli Glacier (the largest and fastest moving of the CDI glaciers) and Darwin Glacier, flow much faster, at around 8-10 metres per day(13).

The Cordillera Darwin Icefield seen from space. Image: NASA

Glacier change

Most glaciers of the Patagonian Icefields are experiencing negative mass balance, as a result of glacier thinning and the widespread retreat of ice-fronts(5,11,14). Only Glacier Pío XI (the largest glacier in South America) from the South Patagonian Icefield has grown in recent years(15,16). The overall pattern of ice loss make the Patagonian Icefields one of the largest current sources of global sea level rise (they contribute ~10% of that from all glaciers and ice caps worldwide17,18).

Since the Little Ice Age

Since the end of the Little Ice Age at around 1870 AD, over 90% of Patagonian outlet glaciers have shrunk(5,8,19) (see GIF below). Reconstructions of Little Ice Age glacier extents(19) show that the North Patagonian Icefield has lost around 103 ± 20.7 km3 of ice, and the South Patagonian Icefield around 503 ± 101.1 km3. This gradual melting of the icefields has contributed around 0.0018 mm to global seal level rise per year (or 0.27 mm in total) since 1870 AD(19).

Retreat of the North Patagonian Icefield between 1870 AD (the end of the Little Ice Age) to 2011.

Over the 21st century

During the last 40-50 years, however, the rate of Patagonian ice loss has sped up substantially. Between 1975 and 2000, the North and South Patagonian Icefields have lost a combined 15.0 ± 0.7 gigatons of ice per year(17). Between 2000 and 2011, the rate of ice loss increased further to 24.4 ± 1.4 gigatons per year(20,21), and has since remained similar (between 2011 and 2017 the icefields lost 21.29 ± 1.98 gigatons per year16).

Retreat of the HPS-12 outlet glacier of the Southern Patagonian Icefield between 1985 and 2017. The glacier terminus has retreated approximately 13 km in the last 30 years. Images from NASA. Compiled by: J. Bendle

Contribution to rising sea level

The current contribution of the Patagonian Icefields to global sea level rise is around 0.067 ± 0.004 mm per year(21), over one order of magnitude greater than the long-term contribution since the Little Ice Age(19). The rate of glacier melting is expected to continue, and maybe even increase, in coming decades(22).

Causes of icefield retreat: a warming atmosphere

The rapid melting of the Patagonian Icefields during the 21st century has several possible causes. For example, a warming atmosphere has resulted in a greater number of warm ‘summer’ days each year, and more melting at the glacier surface(23). At 50°S, for example, a warming of 0.5°C between 1960 to 1999 has resulted in a 0.5 m water equivalent increase in annual glacier melt in the ablation area(24).

Less snowfall, more rainfall

The amount of snowfall falling over the Patagonian Icefields decreased by around 5% between 1960 and 1999, whereas the amount of rainfall increased, both as a result of warmer air temperatures(24). Alongside faster surface melting, therefore, the amount of annual accumulation is getting smaller, preventing glacier growth.

Continued warming in Patagonia, therefore, will have a major impact on glacier accumulation and ablation trends(25), and will ultimately dictate the fate of the Patagonian Icefields.


[1] Garreaud, R.D., Vuille, M., Compagnucci, R. and Marengo, J., 2009. Present-day South American climate. Palaeogeography, Palaeoclimatology, Palaeoecology281, 180-195.

[2] Garreaud, R., Lopez, P., Minvielle, M. and Rojas, M., 2013. Large-scale control on the Patagonian climate. Journal of Climate26, 215-230.

[3] Carrasco, J.F., Casassa, G. and Rivera, A., 2002. Meteorological and climatological aspects of the Southern Patagonia Icefield. In The Patagonian Icefields (pp. 29-41). Springer, Boston, MA.

[4] Schneider, C., Glaser, M., Kilian, R., Santana, A., Butorovic, N. and Casassa, G., 2003. Weather observations across the southern Andes at 53°S. Physical Geography2, 97-119.

[5] Davies, B.J. and Glasser, N.F., 2012. Accelerating shrinkage of Patagonian glaciers from the Little Ice Age (~AD 1870) to 2011. Journal of Glaciology58, 1063-1084.

[6] Falaschi, D., Bravo, C., Masiokas, M., Villalba, R. and Rivera, A., 2013. First glacier inventory and recent changes in glacier area in the Monte San Lorenzo Region (47 S), Southern Patagonian Andes, South America. Arctic, Antarctic, and Alpine Research45, 19-28.

[7] Pfeffer, W.T., Arendt, A.A., Bliss, A., Bolch, T., Cogley, J.G., Gardner, A.S., Hagen, J.O., Hock, R., Kaser, G., Kienholz, C. and Miles, E.S., 2014. The Randolph Glacier Inventory: a globally complete inventory of glaciers. Journal of Glaciology60, 537-552.

[8] Meier, W.J-H., Grießinger, J., Hochreuther, P. and Braun, M.H., 2018. An updated multi-temporal glacier inventory for the Patagonian Andes with changes between the Little Ice Age and 2016. Frontiers in Earth Science, 6, 62.

[9] Carrivick, J.L., Davies, B.J., James, W.H., Quincey, D.J. and Glasser, N.F., 2016. Distributed ice thickness and glacier volume in southern South America. Global and Planetary Change146, 122-132.

[10] Warren, C.R. and Sugden, D.E., 1993. The Patagonian icefields: a glaciological review. Arctic and Alpine Research25, 316-331.

[11] Mouginot, J. and Rignot, E., 2015. Ice motion of the Patagonian icefields of South America: 1984–2014. Geophysical Research Letters42, 441-1449.

[12] Dussaillant, I., Berthier, E. and Brun, F., 2018. Geodetic Mass Balance of the Northern Patagonian Icefield from 2000 to 2012 using two independent methods. Frontiers in Earth Science6, 8.

[13] Melkonian, A.K., Willis, M.J., Pritchard, M.E., Rivera, A., Bown, F. and Bernstein, S.A., 2013. Satellite-derived volume loss rates and glacier speeds for the Cordillera Darwin Icefield, Chile. The Cryosphere7, 823-839.

[14] Rivera, A., Benham, T., Casassa, G., Bamber, J. and Dowdeswell, J.A., 2007. Ice elevation and areal changes of glaciers from the Northern Patagonia Icefield, Chile. Global and Planetary Change59, 126-137.

[15] Schaefer, M., Machguth, H., Falvey, M., Casassa, G. and Rignot, E., 2015. Quantifying mass balance processes on the Southern Patagonia Icefield. The Cryosphere, 9, 25-35.

[16] Foresta, L., Gourmelen, N., Weissgerber, F., Nienow, P., Williams, J.J., Shepherd, A., Drinkwater, M.R. and Plummer, S., 2018. Heterogeneous and rapid ice loss over the Patagonian Ice Fields revealed by CryoSat-2 swath radar altimetry. Remote Sensing of Environment211, 441-455.

[17] Rignot, E., Rivera, A. and Casassa, G., 2003. Contribution of the Patagonia Icefields of South America to sea level rise. Science302, 434-437.

[18] Gardner, A.S., Moholdt, G., Cogley, J.G., Wouters, B., Arendt, A.A., Wahr, J., Berthier, E., Hock, R., Pfeffer, W.T., Kaser, G. and Ligtenberg, S.R., 2013. A reconciled estimate of glacier contributions to sea level rise: 2003 to 2009. Science340, 852-857.

[19] Glasser, N.F., Harrison, S., Jansson, K.N., Anderson, K. and Cowley, A., 2011. Global sea-level contribution from the Patagonian Icefields since the Little Ice Age maximum. Nature Geoscience4, 303-307.

[20] Willis, M.J., Melkonian, A.K., Pritchard, M.E. and Ramage, J.M., 2012. Ice loss rates at the Northern Patagonian Icefield derived using a decade of satellite remote sensing. Remote Sensing of Environment117, 184-198.

[21] Willis, M.J., Melkonian, A.K., Pritchard, M.E. and Rivera, A., 2012. Ice loss from the Southern Patagonian ice field, South America, between 2000 and 2012. Geophysical Research Letters39. L17501.

[22] IPCC, 2013. The Physical Science Basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, UK.

[23] Monahan, P.A. and Ramage, J., 2010. AMSR-E melt patterns on the Southern Patagonia Icefield. Journal of Glaciology56, 699-708.

[24] Rasmussen, L.A., Conway, H. and Raymond, C.F., 2007. Influence of upper air conditions on the Patagonia icefields. Global and Planetary Change59, 203-216.

[25] Schaefer, M., Machguth, H., Falvey, M. and Casassa, G., 2013. Modelling past and future surface mass balance of the Northern Patagonia Icefield. Journal of Geophysical Research: Earth Surface118, 571-588.

Patagonian Ice Sheet at the LGM

What was the former Patagonian Ice Sheet?

The Patagonian Ice Sheet was a large, elongated mountain ice mass that developed over the Andes mountains of southern South America during cold periods[1]. The Patagonian Ice Sheet has advanced and retreated at least 5 times in the last million years[2] in response to changes in global climate (i.e. cooling and warming).

What evidence is there for this former mountain ice sheet?

The evidence for past glaciations of the Patagonian Ice Sheet is preserved in the landscape in the form of landforms (such as moraines) and sediments (such as fine-grained lake sediments and coarse, poorly sorted glacial sediments). The great quantity and variety of glacial landforms in Patagonia[3] make the record of Patagonian Ice Sheet activity one of the longest and most complete anywhere in the world[4].

Map of the Patagonian Ice Sheet at the Last Glacial Maximum around 21,000 years ago. The modern North (NPI), South (SPI), and Cordillera Darwin (CDI) Icefields, and other smaller mountain glaciers, are shown in light blue for comparison.

Last Glacial Maximum (21,000 years ago)

During the global Last Glacial Maximum (LGM) around 21,000 years ago, the Patagonian Ice Sheet almost completely submerged the Patagonian Andes between around 38 to 56°S[1]. In length, the distance between the most northern and southern tips of the ice sheet was around 2000 km (see map above).

How big was the Patagonian Ice Sheet?

Computer simulations estimate that, when at its largest, the Patagonian Ice Sheet had a volume of ~525,000 km3[5]. However, as climate warmed after the LGM, the Patagonian Ice Sheet rapidly thinned and retreated[6,7,8,9]. In all, the ice sheet shrank by ~500,000 km3 in ~8-9000 years, and contributed ~1.2 m to global sea level[5].

Ice sheets at the Last Glacial Maximum worldwide, around 27,000 to 21,000 years ago

What is left of the Patagonian Ice Sheet?

Today, small remnants of the Patagonian Ice Sheet exist in the form of three main mountain icefields, these are: the Northern Patagonian Icefield (NPI; shown in the GIF below), Southern Patagonian Icefield (SPI), and the Cordillera Darwin Icefield (CDI). These ice masses are currently rapidly retreating under the influence of global warming.

Recession of the North Patagonian Icefield, AD 1870 (Little Ice Age) to 2011.

The structure of the former Patagonian Ice Sheet

The Patagonian Ice Sheet was divided into two main parts: a western part and an eastern part that spread out from a central ice-divide along the Andean mountains (see map at top of page). The Patagonian Ice Sheet was drained by at least 66 major outlet glaciers[1]. These outlet glaciers transported ice from the interior parts of the ice sheet to the margins and, in doing so, they controlled the overall form of the ice sheet[1].

But rather than simply flowing east or west from the main ice-divide, these outlet glaciers were strongly influenced by topography (see the arrows showing former ice flow directions in the diagram below), being funnelled through a complicated network of bedrock valleys[10,11].

Outlet glacier flow pathways around the NPI at the Last Glacial Maximum (red line). Glaciers flowed along bedrock valleys (dashed lines) and fed into large, fast-flowing outlet glaciers (solid lines) that filled the widest and deepest troughs.

Pacific Ocean fjords

On the west side of the ice sheet, most outlet glaciers flowed into Pacific Ocean fjords (see the present-day example in the satellite image below). We cannot currently be sure how far these outlet glaciers advanced, because the seafloor has not yet been explored for glacial landforms.

However, computer simulations suggest that most outlet glaciers would have reached the continental shelf edge at the LGM[5]. These simulations also show that, on the west side of the ice sheet, glaciers were fast-flowing, with ice velocities of up to 400 m per year due to plentiful snowfall over the Andean mountains.


A glacier of the modern South Patagonian Icefield (top right) flowing into a deep valley filled with sea water, known as a fjord (bottom left). Image from NASA.

Piedmont lobes

In the northern parts of the former Patagonian Ice Sheet, such as the Chilean Lake District (see map below), west-flowing glaciers did not extend to the Pacific Ocean, but instead formed large piedmont lobes that remained on land[8,12].

The geomorphological map below shows moraines (red lines) that delimit the maximum extent of former outlet glaciers in the Chilean Lake District. Note the piedmont lobe glaciers that spilled out on to flat coastal plains, with their source areas high in the mountains.

Patagonian piedmont lobes in the Chilean Lake District. Moraines mapped by Glasser and Jansson (2008) (ref. 3)

Below is an example of a present-day piedmont lobe glacier in Alaska.

The Agassiz (left) and Malaspina (right) piedmont glaciers spilling out from the Alaskan mountains on to flat coastal plains. Former outlet glaciers in the Chilean Lake District would have looked something like this. Image from NASA.

Moraines and glacial lakes

On the eastern side of the ice sheet, outlet glaciers flowed along large valleys that emerged on the flat Argentinian plains. At the LGM, the largest outlet glaciers advanced more than 150 km east of the modern icefield limits[13]. When they moved on to the flat plains they stabilised, and constructed arcuate terminal moraines[14,15,16,17,18], such as those shown in the photograph below.

During periods of Patagonian Ice Sheet retreat (such as after the LGM) many valleys were flooded with glacial lakes (see the shorelines photographed below, which provide evidence for these former lakes) as meltwater was trapped between terminal moraines and the receding glacier margins[15,19,20,21]. These lakes, which in some valleys were more than 500 m deep[21], had an important role on ice dynamics, likely increasing the rate of glacier retreat through the calving of icebergs[22].

Arcuate terminal moraine (crestline marked by arrows) formed by a major outlet glacier in central Patagonia. The moraine ridge is made up of unconsolidated glacial sediment (that either fell from the glacier surface, or was pushed out from beneath the ice margin), and marks the terminal (or end) point reached by the glacier. Image: J. Bendle.

Top: glacial lake shorelines (marked by white arrows) cut into a terminal moraine. Bottom: a raised lake delta (a landform created when a river enters a lake and deposits sediment) and beach. These landforms can be used to work out the depth of former glacial lakes. Image: J. Bendle.

Why is it important to study the Patagonian Ice Sheet?

Ice sheets are sensitive to changes in the temperature and circulation patterns of the atmosphere and oceans. This means that, firstly, the reconstruction and dating of former ice sheet activity can be used to better understand ice sheet-climate interactions[23]. Such information may be critical in understanding how modern ice sheets will respond to continued global warming[24].

Secondly, in the Southern Hemisphere, which is dominated by oceans, Patagonia is one of only a few areas of land from which scientists can develop records of past environmental change. Such records, which include records of long-term glacial change, allow us to more fully understand how the Southern Hemisphere climate system works, and how it may interact with climate changes happening at the global scale[25].


[1] Glasser, N.F., Jansson, K.N., Harrison, S. & Kleman, J. 2008. The glacial geomorphology and Pleistocene history of South America between 38°S and 56°S. Quaternary Science Reviews, 27, 365–390.

[2] Coronato, A. & Rabassa, J. 2011. Pleistocene glaciations in Southern Patagonia and Tierra del Fuego. In Ehlers, L., Gibbard, P.L., Hughes, P.D. (Eds.) Developments in Quaternary Sciences, 15, Elsevier. pp. 715–727.

[3] Glasser, N.F. & Jansson, K. 2008. The glacial map of southern South America. Journal of Maps, 4, 175–196.

[4] Rabassa, J. & Clapperton, C.M., 1990. Quaternary glaciations of the southern Andes Quaternary Science Reviews, 9, 153–174.

[5] Hulton, N.R., Purves, R.S., McCulloch, R.D., Sugden, D.E. & Bentley, M.J. 2002. The last glacial maximum and deglaciation in southern South America. Quaternary Science Reviews, 21, 233–241.

[6] Hein, A.S., Hulton, N.R., Dunai, T.J., Sugden, D.E., Kaplan, M.R. & Xu, S. 2010. The chronology of the Last Glacial Maximum and deglacial events in central Argentine Patagonia. Quaternary Science Reviews, 29, 1212–1227.

[7] Boex, J., Fogwill, C., Harrison, S., Glasser, N., Hein, A., Schnabel, C. & Xu, S. 2013. Rapid thinning of the Late Pleistocene Patagonian Ice Sheet followed migration of the Southern Westerlies. Scientific Reports 3, 2118.

[8] Moreno, P.I., Denton, G.H., Moreno, H., Lowell, T.V., Putnam, A.E. & Kaplan, M.R. 2015. Radiocarbon chronology of the last glacial maximum and its termination in northwestern Patagonia. Quaternary Science Reviews, 122, 233–249.

[9] Hall, B.L., Porter, C.T., Denton, G.H., Lowell, T.V. & Bromley, G.R. 2013. Extensive recession of Cordillera Darwin glaciers in southernmost South America during Heinrich stadial 1. Quaternary Science Reviews, 62, 49–55.

[10] Glasser, N.F. & Jansson, K.N. 2005. Fast-flowing outlet glaciers of the last glacial maximum Patagonian Icefield. Quaternary Research, 63, 206–211.

[11] Glasser, N.F. & Ghiglione, M.C. 2009. Structural, tectonic and glaciological controls on the evolution of fjord landscapes. Geomorphology, 105, 291–302.

[12] Denton, G.H., Heusser, C.J., Lowel, T.V., Moreno, P.I., Andersen, B.G., Heusser, L.E., Schlühter, C. & Marchant, D.R. 1999. Interhemispheric linkage of paleoclimate during the last glaciation. Geografiska Annaler: Series A Physical Geography, 81, 107–153.

[13] Caldenius, C.C. 1932. Las glaciaciones cuaternarios en la Patagonia y Tierra del Fuego. Geografiska Annaler, 14, 1–164.

[14] Kaplan, M.R., Ackert, R.P., Singer, B.S., Douglass, D.C. & Kurz, M.D. 2004. Cosmogenic nuclide chronology of millennial-scale glacial advances during O-isotope stage 2 in Patagonia. Geological Society of America Bulletin, 116, 308–321.

[15] Sagredo, E.A., Moreno, P.I., Villa-Martínez, R., Kaplan, M.R., Kubik, P.W. & Stern, C.R. 2011. Fluctuations of the Última Esperanza ice lobe (52°S), Chilean Patagonia, during the last glacial maximum and termination 1. Geomorphology, 125, 92–108.

[16] Darvill, C.M., Stokes, C.R., Bentley, M.J. & Lovell, H. 2014. A glacial geomorphological map of the southernmost ice lobes of Patagonia: the Bahía Inútil–San Sebastián, Magellan, Otway, Skyring and Río Gallegos lobes. Journal of Maps, 10, 500–520.

[17] García, J.L., Hall, B.L., Kaplan, M.R., Vega, R.M. & Strelin, J.A. 2014. Glacial geomorphology of the Torres del Paine region (southern Patagonia): Implications for glaciation, deglaciation and paleolake history. Geomorphology, 204, 599–616.

[18] Bendle, J.M., Thorndycraft, V.T. & Palmer, A.P., 2017. The glacial geomorphology of the Lago Buenos Aires and Lago Pueyrredón ice lobes of central Patagonia. Journal of Maps, 13, 654–673.

[19] McCulloch, R.D., Bentley, M.J., Tipping, R.M. & Clapperton, C.M., 2005. Evidence for late glacial ice dammed lakes in the central Strait of Magellan and Bahía Inútil, southernmost South America. Geografiska Annaler: Series A Physical Geography 87, 335–362.

[20] Lovell, H., Stokes, C.R., Bentley, M.J. & Benn, D.I. 2012. Evidence for rapid ice flow and proglacial lake evolution around the central Strait of Magellan region, southernmost Patagonia. Journal of Quaternary Science, 27, 625–638.

[21] Glasser, N.F., Jansson, K.N., Duller, G.A., Singarayer, J., Holloway, M. & Harrison, S. 2016. Glacial lake drainage in Patagonia (13-8 kyr) and response of the adjacent Pacific Ocean. Scientific Reports, 6. 21064.

[22] Carrivick, J.L. & Tweed, F.S. 2013. Proglacial lakes: character, behaviour and geological importance. Quaternary Science Reviews, 78, 34–52.

[23] Kaplan, M.R., Fogwill, C.J., Sugden, D.E., Hulton, N.R.J., Kubik, P.W. & Freeman, S.P.H.T. 2008. Southern Patagonian glacial chronology for the Last Glacial period and implications for Southern Ocean climate. Quaternary Science Reviews, 27, 284–294.

[24] IPCC, 2013. The Physical Science Basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, UK. doi:10.1017/CBO9781107415324.

[25] Killan, R. & Lamy, F. 2012. A review of glacial and Holocene paleoclimate records from southernmost Patagonia (49-55°S). Quaternary Science Reviews, 1–23.


Drumlins around Lago Viedma

Although the Patagonian Icefields aren’t generally associated with drumlins (Glasser et al., 2008), there are some around Lago Viedma in the South Patagonian Icefield. They have been described in detail (Ponce et al., 2013) but they show up beautifully in the Landsat map below. The mapping below is by me (Bethan Davies) and Glasser and Jansson (2008). Drumlins are probably rare in Patagonia as the temperate ice masses release large amounts of meltwater, which may destroy any bedforms.

Drumlins around Lago Viedma, South Patagonian Icefield. The background image is Landsat 7 ETM+ from 2001. Panels C and D are to the same scale.

Drumlins around the world

Drumlins have been observed at the beds of former palaeo ice-sheets across the world. They are found across the Pennines of Britain, in Anglesey, and in the Lake District (Livingstone et al., 2008; Hughes et al., 2010, 2014). In the photo below there is a drumlin on Anglesey. The farmer has helpfully put a stone wall along the long-axis of the drumlin.

Drumlin in Anglesey. Photograph: Bethan Davies

The most ubiquitous subglacial landforms

Drumlins are therefore one of the most ubiquitous landforms formed underneath ice sheets (Clark et al., 2009).  They are typically oval-shaped hills, with a long-axis parallel to ice flow. The up-ice (stoss) face is typically steeper than the down-ice (lee) face (Stokes et al., 2011). They are typically between 250 – 1000 m long, 120-300 m wide, and 1.7 to 4.1 times as long as they are wide (Clark et al., 2009). They also generally occur in clusters, or swarms, as can be seen in the images from around Lago Viedma above.

Drumlins. Top: cross-profile, Second: view from above. Third: a swarm of drumlins. Adapted from work on Wikimedia Commons.

Drumlins in palaeo-ice sheet reconstruction

Glacial geologists frequently use these swarms of drumlins in palaeo-ice sheet reconstruction, because they can be directly related to the direction of former ice flow. They can therefore be used to reconstruct the dynamic behaviour of former ice sheets (Livingstone et al., 2010; Livingstone et al., 2012).  Their length may be related to ice velocities, with a tendency to become more elongated under fast ice-flow conditions. At their longest, they grade into Mega-Scale Glacial Lineations, typically found under former ice-streams.

The mystery of how drumlins are made

Although drumlins have a typical morphology, they can be made up of lots of different kinds of internal sediments, ranging from mainly bedrock, to mainly till, to mainly sorted sediments (Stokes et al., 2011). A diagnostic process for their formation is therefore challenging to deduce.

Cross-section through a drumlin on Anglesey. This one is mostly made of glacial till. Photo: Bethan Davies

One theory growing in popularity is the ‘instability theory’, which states that small perturbations in the bed of the ice sheet grow under a positive-feedback mechanism into the large landforms we see on land (e.g., in Britain and in Patagonia) today (Stokes et al., 2013). Such instabilities tend to grow exponentially, with a rate dependent on wavelength.  Essentially, the process amplifies the relief at the ice-bed interface and results in the formation of bedforms like drumlins in recognisable patterns.

Further reading


Clark, C.D., Hughes, A.L.C., Greenwood, S.L., Spagnolo, M., Ng, F.S.L., 2009. Size and shape characteristics of drumlins, derived from a large sample, and associated scaling laws. Quaternary Science Reviews 28, 677-692.

Glasser, N., Jansson, K., 2008. The Glacial Map of southern South America. Journal of Maps 4, 175-196.

Glasser, N.F., Jansson, K.N., Harrison, S., Kleman, J., 2008. The glacial geomorphology and Pleistocene history of South America between 38°S and 56°S. Quaternary Science Reviews 27, 365-390.

Hughes, A.L.C., Clark, C.D., Jordan, C.J., 2010. Subglacial bedforms of the last British Ice Sheet. Journal of Maps 6, 543-563.

Hughes, A.L.C., Clark, C.D., Jordan, C.J., 2014. Flow-pattern evolution of the last British Ice Sheet. Quaternary Science Reviews 89, 148-168.

Livingstone, S.J., Evans, D.J.A., Ó Cofaigh, C., Davies, B.J., Merritt, J.W., Huddart, D., Mitchell, W.A., Roberts, D.H., Yorke, L., 2012. Glaciodynamics of the central sector of the last British–Irish Ice Sheet in Northern England. Earth-Science Reviews 111, 25-55.

Livingstone, S.J., Ó Cofaigh, C., Evans, D.J.A., 2008. Glacial geomorphology of the central sector of the last British-Irish Ice Sheet. Journal of Maps 2008, 358-377.

Livingstone, S.J., Ó Cofaigh, C., Evans, D.J.A., 2010. A major ice drainage outlet of the last British-Irish Ice Sheet: the Tyne Gap, northern England. Journal of Quaternary Science 25, 354-370.

Ponce, J.F., Rabassa, J., Serrat, D., Martínez, O.A., 2013. El campo de drumlins, flutes y megaflutes de lago Viedma, Pleistoceno Tardío, provincia de Santa Cruz. Revista de la Asociación Geológica Argentina 70, 115-127.

Stokes, C.R., Fowler, A.C., Clark, C.D., Hindmarsh, R.C.A., Spagnolo, M., 2013. The instability theory of drumlin formation and its explanation of their varied composition and internal structure. Quaternary Science Reviews 62, 77-96.

Stokes, C.R., Spagnolo, M., Clark, C.D., 2011. The composition and internal structure of drumlins: Complexity, commonality, and implications for a unifying theory of their formation. Earth-Science Reviews 107, 398-422.

Observing glacier change from space

Satellite observation of glaciers | Mapping structural glaciology | Mapping glacier change | Mapping glacier thinning | Mapping glacier velocity | Other glaciological applications | Different satellites with different capabilities | Summary and conclusions | Further reading | References | Comments |

Satellite observation of glaciers

The Terra (EOS AM-1) satellite, operated by NASA, with the MODIS and ASTER sensors, is regularly used by glaciologists due to its affordability, wide swath (60km) and good resolution (15 m). Image credit: Wikipedia.

The Terra (EOS AM-1) satellite, operated by NASA, with the MODIS and ASTER sensors, is regularly used by glaciologists due to its affordability, wide swath (60km) and good resolution (15 m). Image credit: Wikipedia.

Ice cover on Earth is changing. We have records going back decades, measuring the mass balance of glaciers in the Alps. We can measure the distance between the modern glacier terminus and its maximum position during the Little Ice Age, around 1850 AD in the northern hemisphere. We can visit the same glacier each year for ten years and visually see the recession of its snout. However, empirical field measurements of glacier mass balance and recession are arduous, time consuming and expensive. Point measurements are collected from just a small number of reference glaciers. But the advent of remote sensing over the last 40 years has opened up unparalleled opportunities for glacier observation.

Cameras in space

ASTER image of James Ross Island, taken 03 March 2009. Swath size is 60km.

ASTER image of James Ross Island, taken 03 March 2009. Swath size is 60km.

“Remote Sensing” just means the acquisition of information about an object without touching it. We have been “remote sensing” with our eyes for millennia. But put a camera up into space, and suddenly, we can observe glacier behaviour at a much larger scale. With repeated images over time, we can easily see how glaciers have changed.

“Remote sensing” includes obtaining information from aerial photographs from drones or aeroplanes, multibeam swath bathymetry data from ships, ground penetrating radar data and more. In this article, we focus only on remote sensing from satellites.

Earth Observation satellites orbit the Earth, taking pictures in various bands of the electromagnetic spectrum, from visible to thermal infrared. The large swaths (footprint of the image on the ground) make regional assessments simple. Each satellite takes pictures at a different resolution, and so each is used for different applications. Optical satellite remote sensing allows for the regular monitoring of glacier surface elevation, velocity, area, length, equilibrium line altitude, terminus position and more1,2.

Hundreds of applications for remote sensing

Section of an ASTER image of Ulu Peninsula, James Ross Island.

Section of an ASTER image of Ulu Peninsula, James Ross Island.

Remote sensing is a mature technique with many applications. The following gives some examples of how satellite imagery and remote sensing is used in the world of glaciology. Feel free to add a Comment and tell us how you use remote sensing in your glaciological research!

Scroll to the Different satellites with different capabilities section to find out all about each of the different satellites that glaciologists use for remote sensing of cryospheric change.

Mapping structural glaciology

Image swath and image resolution.

Image swath and image resolution.

Satellites providing visual images of Earth are great for mapping regional glacier and ice-shelf structures. ASTER and Landsat images are frequently used for this kind of work, because their large swaths (footprint) means that a regional view of the ice is provided, but their relatively fine resolution means that even small glacier structures, such as crevasses and melt ponds, can be imaged and mapped. The long time series of images from these satellites (ASTER since 1999, Landsat since the 1970s), is useful for mapping changes over time.

Larsen B Ice Shelf

Antarctic Peninsula Ice Shelves

Antarctic Peninsula Ice Shelves

As an example, Glasser and Scambos used optical satellite imagery (ASTER and Landsat 7) to map detailed changes in ice structures in the build-up to the collapse of Larsen B ice shelf in 2002. Their maps showed rifts becoming more pronounced in the years before the ice-shelf collapse, and that ice-shelf collapse was preconditioned by partial rupturing of the sutures between flow units3. This shows one of the great advantages of satellite remote sensing – after a catastrophic event, such as ice-shelf collapse, you can go back in time and analyse how the ice behaved prior to disintegration.

Prince Gustav Ice Shelf

Prince Gustav Ice Shelf in 1988. It collapsed in 1995, and the glaciers which flowed into it subsequently accelerated and thinned, transmitting lots of ice into the ocean and resulting in measureable sea level rise.

Landsat image of Prince Gustav Ice Shelf in 1988. It collapsed in 1995, and the glaciers which flowed into it subsequently accelerated and thinned, transmitting lots of ice into the ocean and resulting in measureable sea level rise.

The situation is similar for Prince Gustav Ice Shelf, where the ice shelf suture zones formed zones of weakness, abetting ice-shelf collapse4, which occurred in 1995. The Landsat image opposite shows clear sutures forming between flow units. Meltwater ponding on the surface eventually aided iceberg formation and ice-shelf disintegration by melting rapidly downwards through the ice shelf.

Ice shelves on the SW Antarctic Peninsula

Holt et al. also used Landsat images to map the surface features of ice shelves on the south-western Antarctic Peninsula. Tom Holt was able to map the ice front, grounding zone, longitudinal surface structures (flow stripes), pressure ridges, crevasses, fractures and rifts, surface meltwater, ice dolines, ice rises and ice rumples5. Together, these data help us to understand the structure of more ‘stable’ ice shelves, and helps us to assess their likely future response and vulnerabilities to climate change6.

Glaciological structures on George VI Ice Shelt, Antarctic Peninsula, illustrating the dominance of meltwater channels, longitudinal structures and rifts towards the south ice front. From Holt et al., 2013.

Glaciological structures on George VI Ice Shelt, Antarctic Peninsula, illustrating the dominance of meltwater channels, longitudinal structures and rifts towards the south ice front. From Holt et al., 2013.

Flow stripes on Antarctic ice streams

Glasser and Gudmundsson (2012) have mapped surface structures on a number of Antarctic ice streams7. Because the pattern of velocity varies across the ice streams, crevasses and flow lines form.

Mapping glacier change

Glacier recession in Patagonia

Mapping the glaciers of the Northern Patagonian Icefield.

Mapping the glaciers of the Northern Patagonian Icefield.

Because of the long history of optical satellite imagery of the Earth, satellite remote sensing offers fabulous opportunities for mapping glacier recession. It is simple to derive 40+ year histories of regional glacier recession8-11. Landsat and ASTER images are frequently used for this application. For example, remote sensing of glacier ice and Little Ice Age moraines in Patagonia (formed circa 1870) and glacier extent in 2011, 2001, 1986 and 1975 provides a detailed analysis of rates of recession over the last 140 years12.


Recession of the North Patagonian Icefield, AD 1870 (Little Ice Age) to 2011.

Conducting glacier inventories

Glacier inventories are also an essential part of glaciological monitoring, and many parameters, such as glacier length, altitudinal range, snow cover and so on can be calculated from satellite images. These data are essential for wider studies and modelling of glaciers. The GLIMS Geospatial Database13 provides detailed guidelines on the correct methods for conducting glacier inventories14-19, and a free online database where any user can download outlines of all inventoried glaciers.

Additional data often mapped by satellite and added to glacier inventories include changing ice debris cover20, rock glaciers21, equilibrium line altitude22, grounding zones23 and glacial lake extent24,25.  Volume-area scaling laws mean that you can calculate changing ice volume26. The applications are as limitless as your imagination.

Mapping glacier thinning

While it’s important to map glacier recession, it’s not the whole story. Glaciers may be thinning and downwasting in situ, with little change at their terminus. Mapping glacier thinning is therefore an important application of glacier remote sensing. One way of doing this is to generate two digital elevation models from two different time periods and to subtract one from the other (e.g., ref. 27). However, this isn’t easy as it assumes that the change in surface elevation is greater than the errors and uncertainties in the digital elevation models.

Dynamic thinning around Pine Island Glacier and the Antarctic Peninsula

One analysis by Kunz et al. of images taken from aeroplanes (1948-2005) and ASTER stereo-pairs (2001-2010) showed that glaciers in the western Antarctic Peninsula have been thinning over the last few decades28. All the glaciers show surface lowering in their ablation zones since the 1960s, with the  highest rates of thinning occurring in the north-western Antarctic Peninsula region28.

Another method of calculating elevation change is to use ICESat data. This satellite measures ice surface elevation, and repeated tracks show elevation change through time. Analysis of ICESat data across George VI Ice Shelf on the south-western Antarctic Peninsula agreed with the DEM differencing work conducted by Kunz et al. Holt et al. analysed surface elevation change of George VI Ice Shelf using ICESat’s GLAS repeat measurements from 2003 to 20086. Significant thinning was observed in the southern section of the ice shelf, which is coupled with widespread recession of the ice front terminus.

Thinning of George VI Ice Shelf. Non-significant elevation change (less than the uncertainties) was observed in the northern part of the ice shelf. Widespread and significant elevation change was recorded in the southern section, coupled with retreat of the grounding zone. From Holt et al., 2013.

Thinning of George VI Ice Shelf. Non-significant elevation change (less than the uncertainties) was observed in the northern part of the ice shelf. Widespread and significant elevation change was recorded in the southern section, coupled with retreat of the grounding zone. From Holt et al., 2013.

Rate of surface elevation change from Antarctica and Greenland from ICEsat data (laser altimetry). From: Pritchard et al. 2009, Nature.

Rate of surface elevation change from Antarctica and Greenland from ICEsat data (laser altimetry). From: Pritchard et al. 2009, Nature. Reprinted by permission from Macmillan Publishers Ltd [Nature] (Pritchard et a., 2009), copyright 2009.

This highly accurate data has also been applied to the wider Antarctic continent, where it shows widespread thinning as a result of accelerated ice flow around Pine Island Glacier and the Antarctic Peninsula29. The high-resolution ICESat dataset is able to resolve even small glaciers on the complex topography of the Antarctic Peninsula. This “dynamic thinning”, a result of fast ice flow, has now intensified on key Antarctic grounding lines, endures for decades after ice-shelf collapse, penetrates far into the interior of the ice sheet and is spreading as ice shelves melt and thin (due to warming from below by ocean currents).

Thinning ice shelves

Antarctic ice shelf thickness changes. Note the rapid thinning of Pine Island Glacier ice shelf in West Antarctica. From Pritchard et al., 2012, Nature. Reprinted by permission from Macmillan Publishers Ltd: Nature (Pritchard et al. 2012), copyright (2012).

Antarctic ice shelf thickness changes calculated from ICEsat data. Note the rapid thinning of Pine Island Glacier ice shelf in West Antarctica. From Pritchard et al., 2012, Nature. Reprinted by permission from Macmillan Publishers Ltd: Nature
(Pritchard et al. 2012), copyright (2012).

ICESat has also been used to calculate ice-shelf thinning and basal melt in ice shelves around Antarctica30. Pritchard et al. used a combination of satellite laser altimetry and modelling of the surface firn layer to show ice-shelf thinning around Antarctica as a result of increased basal melt. This melt is the primary control on Antarctic ice-sheet loss, as the thinner ice shelves are less able to buttress ice in the interior, leading to faster ice flow. The strongest thermal forcing and highest melt rates were found near Pine Island Glacer, West Antarctica.



Mapping glacier velocity

Measuring regional glacier and ice stream velocity, and its change through time, is a critical application of glacier remote sensing. There are several methods; the first relies on repeated optical satellite imagery of one region. An algorithm applied to the images calculates the distance that features on the ice surface have moved (feature tracking) (e.g., 27). Cosi-Corr is frequently used for feature tracking in this way31,32. A second method uses repeat radar images (Synthetic Aperture Radar interferometry, or InSAR) to calculate glacier velocity. Two pairs of images are used to calculate ascending-pass and descending-pass interferograms. Velocity fields can then be calculated6.

Ice velocity on George VI Ice Shelf

For example, Tom Holt used a combination of radar and optical feature tracking to examine the response of George VI Ice Shelf to environmental change6. This analysis showed ice-shelf acceleration towards the north and south ice fronts, combined with thinning throughout and an increase in fracture distribution towards the southern ice front.

Surface speed calculations at the north ice front of George VI Ice Shelf for ca. 1989 (A), 1995 (B), 2002 (C) and 2007 (D). From Holt et al., 2013.

Surface speed calculations at the north ice front of George VI Ice Shelf for ca. 1989 (A), 1995 (B), 2002 (C) and 2007 (D). From Holt et al., 2013.

Mapping Antarctic ice stream velocity

Map showing location of modern ice streams around Antarctica, made using velocity data from Rignot et al. 2011

ICEsat data is used to map the ice streams around Antarctica. Image made using velocity data from Rignot et al. 2011

Eric Rignot compiled masses of InSAR velocity data to present a comprehensive, high resolution map of Antarctic ice velocity. The data revealed widespread enhanced flow, with tributary glaciers reaching deep into the ice-sheet’s interior.

Other glaciological applications

There are many hundreds of applications of remote sensing data for glaciers, so I will just pick out a few here.

Mapping Equilibrium Line Altitudes

Glacier inventory data (information on glacier length, altitudinal range, thickness, snow cover etc) can be used to calculate regional Equilibrium Line Altitudes (ELAs)22. These data provide important information on glacier mass balance, an important glaciological parameter.

Mapping snow cover

Related to above is mapping of snow cover. This has important applications for understanding glacier mass balance as well as the behaviour of perennial snow patches, water resources and more.  Scientists may set up algorithms to automatically classify terrain with different reflectance values (white snow, dark water, green land etc). This makes it simple to quickly map changing snow cover in a wider region.

Mapping glacier landforms

Glacial geology has been revolutionised by satellite remote sensing. Large swaths mean that images that are subtle and difficult to map on the ground can be quickly and easily mapped from space. A typical example is mega-scale glacial lineations; these features  have now been recognised in glaciated terrains across the northern hemisphere, and are easily visible in a digital elevation model created from satellite images (e.g., see 33-35).

Measuring the ice-sheet bed

Information on bed topography is one of the key parameters needed to drive ice-sheet and glacier numerical models. Unfortunately, the bed of a glacier is rather difficult to get at. Fortunately, radar mounted on satellites can penetrate the ice, helping to map the geometry of the ground beneath the ice. Synthetic aperture radar (SAR) can be used to generate digital elevation models of the ice bed, even in the most remote of regions36.

Different satellites with different capabilities

There are various satellites used for observation of ice sheets. They have different resolutions (pixel size in the image) and swath sizes (the footprint of the image on the ground). They have different return rates (returning to take the same image). As a result, different applications are best suited to different satellites.  ASTER and Landsat are some of the most widely used sensors in glaciology, due to their long history of taking images, broad swaths allowing for regional analysis, and high image resolution. ASTER also takes images as stereo-pairs, meaning that digital elevation models can be made from its images (ASTER GDEM).

Some of the most common satellites used in glacier and ice sheet observation are:

  Satellite name Ground resolution (pixel size) Swath size Operation dates Sensors Applications

Visualising the Earth

ASTER 15 m 60 km Launched December 1999 Multispectral (14 bands) sensor. Three separate sensors:
VNIR (Visible Near Infrared), (15 m)SWIR (ShortWave Infrared), (30 m)TIR (Thermal Infrared) (90 m)
ASTER data is used to create detailed maps of land surface temperature, reflectance, and elevation.ASTER captures high spatial resolution data in 14 bands, from the visible to the thermal infrared wavelengths, and provides stereo viewing capability for digital elevation model creation.
Landsat 7 ETM+ 30 m 185 km Launched April 1999 Multispectral (7 bands) Series of satellites (currently Landsat 8), with the earliest images dating from the 1970s.
Spot-5 5 m and 2.5 m 60 x 60 km and 60 x 120 km Launched May 2002. Twin sensors.Multispectral and panchromatic.1 shortwave infrared. The coverage offered by SPOT-5 is a key asset for applications such as medium-scale mapping (at 1:25 000 and 1:10 000 locally). Revisit time: 2-3 days
ENVI Sat 30 m 58 km March 2002 to May 2012. Multispectral with synthetic aperture radar. Repeat cycle of 35 days. After losing contact with the satellite on 8 April 2012, ESA formally announced the end of Envisat’s mission on 9 May 2012
MODIS 250 m (bands 1-2)
500 m (bands 3-7)
1000 m (bands 8-36)
2330 km Launched May 2002. Multispectral (36 spectral bands) MODIS (or Moderate Resolution Imaging Spectroradiometer) is a key instrument aboard the Terra (EOS AM) and Aqua (EOS PM) satellites.Terra MODIS and Aqua MODIS are viewing the entire Earth’s surface every 1 to 2 days, acquiring data in 36 spectral bands, or groups of wavelengths
Quickbird 0.61 m 16.5 Km x 16.5 Km at nadir Launched October 2001 Multispectral QuickBird is a high resolution satellite owned and operated by DigitalGlobe. This satellite is an excellent source of environmental data useful for analyses of changes in land usage, agricultural and forest climates.
Geo-Eye 0.41 m panchromatic; 1.65 m multispectral 15.2 km Launched September 2008 Pan-sharpened multispectral imagery. Revisit time of 3 days.
Worldview-2 0.46 m panchromatic; 1.85 m multispectral 16.4 km Launched October 2009 Multispectral (8 bands) Revisit time of every 1.1 days.

Measuring elevation change

Cryosat 250 m wide strips. Launched April 2010 Synthetic Aperture Interferometric Radar Altimeter (SIRAL). The altimeter makes a measurement of the distance between the satellite and the surface. Cryosat is designed to calculate the elevation of ice and sea-ice ‘freeboard’, which is the height protruding from the water. Its aim is to determine changes in land-ice and sea-ice thickness.It reaches 88°N and S on every orbit.
ICEsat Can detect changes in elevation of up to 1 cm per year. 60 m January 2003-2009. Decommissioned August 2010. ICEsat-2 is planned in 2016. Lasers for measuring elevation change Ice, Cloud, and land Elevation Satellite. The ICESat mission provided multi-year elevation data needed to determine ice sheet mass balance as well as cloud property information. ICESat ended its science mission in February 2010 with the failure of the last of its three lasers.
SRTM 90 m for the DEM 225 km February 2000 (single mission) Radar The Shuttle Radar Topography Mission (SRTM) obtained elevation data on a near-global scale to generate the most complete high-resolution digital topographic database of Earth. SRTM consisted of a specially modified radar system that flew on-board the Space Shuttle Endeavour during an 11-day mission in February of 2000.

Measuring gravity

GRACE March 2002 The mission uses a microwave ranging system to accurately measure changes in the speed and distance between two identical spacecraft flying in a polar orbit about 220 kilometres (140 mi) apart, 500 kilometres (310 mi) above Earth. The Gravity Recovery And Climate Experiment (GRACE).  Twin satellites making detailed measurements of the Earth’s gravity field. Changes in the gravity are interpreted in terms of changes to the Earth’s ice sheets.

Summary and conclusions for remote sensing of glaciers

Satellite remote sensing has revolutionised glaciology and glacial geology by enabling quick, cheap and logistically simple ways to map large landscapes. It has enabled glacier inventories of entire countries, underpinned our understanding of glacier recession and advance, helps us to map glacier snow cover and mass balance, track changes in ice sheet thickness and ice flow velocities, and allowed detailed changes in remote locations to be observed.

Do you have a favourite glacier remote sensing technique? How do you use satellite images in your work? Add in an example in the Comments box and tell us!

Further reading


1.            Racoviteanu, A.E., Williams, M.W. & Barry, R.G. Optical remote sensing of glacier characteristics: a review with focus on the Himalaya. Sensors 8, 3355-3383 (2008).

2.            Paul, F. et al. The glaciers climate change initiative: Methods for creating glacier area, elevation change and velocity products. Remote Sensing of Environment.

3.            Glasser, N.F. & Scambos, T.A. A structural glaciological analysis of the 2002 Larsen B ice shelf collapse. Journal of Glaciology 54, 3-16 (2008).

4.            Glasser, N.F. et al. From ice-shelf tributary to tidewater glacier: continued rapid glacier recession, acceleration and thinning of Röhss Glacier following the 1995 collapse of the Prince Gustav Ice Shelf on the Antarctic Peninsula. Journal of Glaciology 57, 397-406 (2011).

5.            Holt, T., Glasser, N. & Quincey, D. The structural glaciology of southwest Antarctic Peninsula Ice Shelves (ca. 2010). Journal of Maps 9, 523-531 (2013).

6.            Holt, T.O., Glasser, N.F., Quincey, D. & Siegfried, M.R. Speedup and fracturing of George VI Ice Shelf, Antarctic Peninsula. The Cryosphere 7, 797-816 (2013).

7.            Glasser, N.F. & Gudmundsson, G.H. Longitudinal surface structures (flowstripes) on Antarctic glaciers. The Cryosphere 6, 383-391 (2012).

8.            Abermann, J., Lambrecht, A., Fischer, A. & Kuhn, M. Quantifying changes and trends in glacier area and volume in the Austrian Ötztal Alps (1969-1997-2006). The Cryosphere 3, 205-215 (2009).

9.            Arendt, A. et al. Randolph Glacier Inventory [v2.0]: A Dataset of Global Glacier Outlines. (Global Land Ice Measurements from Space, Boulder Colorado, USA., 2012).

10.          Berthier, E., Schiefer, E., Clarke, G.K.C., Menounos, B. & Remy, F. Contribution of Alaskan glaciers to sea-level rise derived from satellite imagery. Nature Geoscience 3, 92-95 (2010).

11.          Bolch, T., Menounos, B. & Wheate, R. Landsat-based inventory of glaciers in western Canada, 1985-2005. Remote Sensing of Environment 114, 127-137 (2010).

12.          Davies, B.J. & Glasser, N.F. Accelerating recession in Patagonian glaciers from the “Little Ice Age” (c. AD 1870) to 2011. Journal of Glaciology 58, 1063-1084 (2012).

13.          Raup, B. et al. The GLIMS geospatial glacier database: A new tool for studying glacier change. Global and Planetary Change 56, 101-110 (2007).

14.          Raup, B. & Khalsa, S.J.S. GLIMS Analysis Tutorial, 15 (GLIMS, Global Land Ice Measurements from Space, NSIDC,, 2010).

15.          Paul, F. et al. Guidelines for the compliation of glacier inventory data from digital sources, 23 (GLIMS, Global Land Ice Measurement from Space, NSIDC, University of Colorado, Boulder, 2010).

16.          Racoviteanu, A.E., Paul, F., Raup, B., Khalsa, S.J.S. & Armstrong, R. Challenges and recommendations in mapping of glacier parameters from space: results of the 2008 Global Land Ice Measurements from Space (GLIMS) workshop, Boulder, Colorado, USA. Annals of Glaciology 50, 53-69 (2009).

17.          Paul, F. et al. Recommendations for the compilation of glacier inventory data from digital sources. Annals of Glaciology 50, 119-126 (2009).

18.          Raup, B. et al. Remote sensing and GIS technology in the Global Land Ice Measurements from Space (GLIMS) Project. Computers & Geosciences 33, 104-125 (2007).

19.          Rau, F., Mauz, F., Vogt, S., Khalsa, S.J.S. & Raup, B. Illustrated GLIMS Glacier Classification Manual, Version 1.0, 36 (GLIMS (Global Land Ice Measurement from Space), NSIDC, GLIMS Regional Centre, ‘Antarctic Peninsula’, 2005).

20.          Nagai, H., Fujita, K., Nuimura, T. & Sakai, A. Southwest-facing slopes control the formation of debris-covered glaciers in the Bhutan Himalaya. The Cryosphere 7, 1303-1314 (2013).

21.          Esper Angillieri, M.Y. A preliminary inventory of rock glaciers at 30°S latitude, Cordillera Frontal of San Juan, Argentina. Quaternary International 195, 151-157 (2009).

22.          Braithwaite, R.J. & Raper, S.C.B. Estimating equilibrium-line altitude (ELA) from glacier inventory data. Annals of Glaciology 50, 127-132 (2009).

23.          Rignot, E., Mouginot, J. & Scheuchl, B. Antarctic grounding line mapping from differential satellite radar interferometry. Geophys. Res. Lett. 38, L10504 (2011).

24.          Frey, H., Huggel, C., Paul, F. & Haeberli, W. Automated detection of glacier lakes based on remote sensing in view of assessing associated hazard potentials. Grazer Schriften der Geographie und Raumforschung Band 44, 23-30 (2010).

25.          Loriaux, T. & Casassa, G. Evolution of glacial lakes from the Northern Patagonia Icefield and terrestrial water storage in a sea-level rise context. Global and Planetary Change 102, 33-40 (2013).

26.          Radic, V., Hock, R. & Oerlemans, J. Analysis of scaling methods in deriving future volume evolutions of valley glaciers. Journal of Glaciology 54, 601-612 (2008).

27.          Willis, M.J., Melkonian, A.K., Pritchard, M.E. & Ramage, J.M. Ice loss rates at the Northern Patagonian Icefield derived using a decade of satellite remote sensing. Remote Sensing of Environment 117, 184-198 (2011).

28.          Kunz, M. et al. Multi-decadal glacier surface lowering in the Antarctic Peninsula. Geophys. Res. Lett. 39, L19502 (2012).

29.          Pritchard, H.D., Arthern, R.J., Vaughan, D.G. & Edwards, L.A. Extensive dynamic thinning on the margins of the Greenland and Antarctic ice sheets. Nature 461, 971-975 (2009).

30.          Pritchard, H.D. et al. Antarctic ice-sheet loss driven by basal melting of ice shelves. Nature 484, 502-505 (2012).

31.          Herman, F., Anderson, B. & Leprince, S. Mountain glacier velocity variation during a retreat/advance cycle quantified using sub-pixel analysis of ASTER images. Journal of Glaciology 57, 197-207 (2011).

32.          Leprince, S., Barbot, S., Ayoub, F. & Avouac, J.-P. Automatic and precise orthorectification, coregistration, and subpixel correlation of satellite images, application to ground deformation measurements. IEEE Transaction on geoscience and remote sensing 45, 1529-1558 (2007).

33.          Stokes, C.R. et al. Formation of mega-scale glacial lineations on the Dubawnt Lake Ice Stream bed: 1. size, shape and spacing from a large remote sensing dataset. Quaternary Science Reviews in press(2013).

34.          Stokes, C.R. & Clark, C.D. The Dubawnt Lake palaeo-ice stream: evidence for dynamic ice sheet behaviour on the Canadian Shield and insights regarding the controls on ice-stream location and vigour. Boreas 32, 263-279 (2003).

35.          Clark, C.D. & Stokes, C.R. Extent and basal characteristics of the M’Clintock Channel Ice Stream. Quaternary International 86, 81-101 (2001).

36.          Paden, J., Akins, T., Dunson, D., Allen, C. & Gogineni, P. Ice-sheet bed 3-D tomography. Journal of Glaciology 56, 3-11 (2010).

Antarctic Ice Sheet mass balance

How does mass balance vary over Antarctica? | Surface mass balance in the past | Surface mass balance in the future | References | Comments |

How does mass balance vary over Antarctica?

Is Antarctica currently losing or gaining mass? Will this massive ice sheet grow or shrink in the future? And what effect will increased snowfall have over coming centuries? In order to answer these questions, we must analyse the surface mass balance of the Antarctic Ice Sheet.

First, let’s introduce some definitions.

  • Mass balance is the sum of all processes of accumulation and ablation, including those at the ice surface and at the bed, but does not include mass changes due to ice flow1. See this page (Introduction to Glacier Mass Balance) for more information.
  • Surface mass balance is the net balance between the processes of accumulation and ablation on a glacier’s surface (it does not include dynamic mass loss and basal melting)1.
  • Climatic mass balance includes surface mass balance and internal accumulation1.
  • Ice dynamical changes may include changes to ice discharge and acceleration or deceleration of flow, which can lead to dynamic thinning or thickening, ice-shelf collapse, marine ice sheet instability, and other factors resulting in changes in the glacier beyond surface mass balance.

Surface mass balance

Surface mass balance varies extensively over Antarctica. The Antarctic Peninsula has the highest accumulation rates (up to 1500 mm per year), followed by coastal West Antarctica, which has around 1000 mm accumulation per year2. Compare this with the interior of the Antarctic Ice Sheet, where it is dry and cold; here accumulation can be less than 25 mm per year.

Surface mass balance of the Antarctic and Greenland ice sheets. From Van den Broeke et al., 2011.

Surface mass balance of the Antarctic and Greenland ice sheets. From Van den Broeke et al., 2011.

Surface mass balance estimates are constantly improving as scientists gain better understandings of glacio-isostatic adjustment, improve glacier modelling techniques and gain access to higher resolution satellite datasets over longer timescales3. Surface mass balance estimates therefore tend to improve over time, but are subject to large uncertainties4. For this reason, there tends to be differences between the results of different techniques used to measure surface mass balance. Surface mass balance of the grounded Antarctic Ice Sheet is currently estimated at ~2000 gigatonnes per year2, 5, 6, and it is subject to large variability across the ice sheet and through time.

Total mass balance

The figure below shows some recent estimates for total mass balance (including basal processes) over Antarctica7. Each box is bounded by the time interval studied and the uncertainties identified.

Summary of estimates of rates of ice mass change for Antarctica and Greenland. Reprinted by permission from Macmillan Publishers Ltd: [Nature] (Hanna et al., 2013) copyright (2013)

Summary of estimates of rates of ice mass change for Antarctica and Greenland. Reprinted by permission from Macmillan Publishers Ltd: [Nature] (Hanna et al., 2013) copyright (2013)

Overall, a recent estimate puts Antarctic net mass balance at -71 ± 53 gigatonnes per year8, so just negative over the 19 year survey. Mass losses are increasing in West Antarctica and the Antarctic Peninsula. The mass balance of West Antarctica is dominated by dynamic losses from the Amundsen Sea sector, and dynamic gains from the Kamb Ice Stream8. From the period 2005-2010, Shepherd et al. (2012) estimate the mass balance of the entire Antarctic Ice Sheet to be -81 ± 37 gigatonnes per year8.

An unweighted average of recent estimates suggests that Antarctica moved from a weakly negative mass balance in the 1990s to a faster rate of mass loss at a rate of between -45 and -120 gigatonnes per year7. Larger dynamic losses in West Antarctica are being partially offset by increases in accumulation over East Antarctica.

Accelerating total mass losses from Antarctica

The GRACE (Gravity Recovery and Climate Experiment) satellite gravity mission shows that total mass loss in Antarctica is accelerating over time. They found that total mass loss increased by 26 ± 14 gigatonnes per year from 2002 to 20099. Rignot et al. (2011) found a smaller acceleration of 14.5±2 gigatonnes per year from 1993-20115, but this change is still three times larger than that found for mountain glaciers and ice caps.

Surface mass balance of Antarctica in the past

How has the surface mass balance of Antarctica changed in the past? Firn and ice-core records can hold the key to providing a longer perspective on surface mass balance than is currently available from satellite records. Frezzotti et al. used 67  of these cores to reconstruct surface mass balance over the last 800 years. They found that current surface mass balance is not exceptionally high compared with the last 800 years10. Periods of high accumulation occurred in the past, in the 1370s and 1610s AD, but there has been an increase of 10% in snow accumulation in some coastal regions since 1850 – a fact that agrees with independent work on the Antarctic Peninsula11.

Surface mass balance of Antarctica in the future

Climate models predict that, for a generally warmer climate, snowfall will increase over Antarctica7. Surface melt will increase around the more northerly Antarctic Peninsula, and dynamic changes such as increased ice discharge12, ice-shelf collapse and grounding line recession13, and marine ice-sheet instability are likely to offset any increases in precipitation7. However, if no dynamical ice response is assumed, then increases in snowfall over the entire continent of 6-16% to 2100 AD and 8-25% to 2200 AD are likely to result in a drop in sea level of 20-43 mm in 2100 and 73-163 in 2200, compared with today14. However, it is more likely that the Greenland and Antarctic ice sheets will lose mass over the next century, with rapid coastal changes, increases in ice flow and ice-shelf collapse all likely4. As a result of these complex expected changes, there are a number of uncertainties in past, present and future ice sheet mass balance.

Further reading


1.            Cogley, J.G., Hock, R., Rasmussen, B., Arendt, A., Bauder, A., Braithwaite, R.J., Jansson, P., Kaser, G., Moller, M., Nicholson, L., & Zemp, M. Glossary of Glacier Mass Balance and related terms. Paris: IHP-VII Technical Documents in Hydrology No. 86, IACS Contribution No. 2, UNESCO-IHP. 124 (2011).

2.            Lenaerts, J.T.M., van den Broeke, M.R., van de Berg, W.J., van Meijgaard, E., & Kuipers Munneke, P. A new, high-resolution surface mass balance map of Antarctica (1979–2010) based on regional atmospheric climate modeling. Geophysical Research Letters. 39, L04501 (2012).

3.            Van den Broeke, M., Bamber, J., Lenaerts, J., & Rignot, E. Ice Sheets and Sea Level: Thinking Outside the Box. Surveys in Geophysics. 32, 495-505 (2011).

4.            Alley, R.B., Spencer, M.K., & Anandakrishnan, S. Ice-sheet mass balance: assessment, attribution and prognosis. Annals of Glaciology. 46, 1-7 (2007).

5.            Rignot, E., Velicogna, I., Van den Broeke, M., Monaghan, A., & Lenaerts, J. Acceleration of the contribution of the Greenland and Antarctic ice sheets to sea level rise. Geophysical Research Letters. 38, (2011).

6.    Agosta, C., Favier, V., Krinner, G., Gallée, H., Fettweis, X., & Genthon, C. High-resolution modelling of the Antarctic surface mass balance, application for the twentieth, twenty first and twenty second centuries. Climate Dynamics. 41, 3247-3260 (2013).

7.            Hanna, E., Navarro, F.J., Pattyn, F., Domingues, C.M., Fettweis, X., Ivins, E.R., Nicholls, R.J., Ritz, C., Smith, B., Tulaczyk, S., Whitehouse, P.L., & Zwally, H.J. Ice-sheet mass balance and climate change. Nature. 498, 51-59 (2013).

8.            Shepherd, A., Ivins, E.R., A, G., Barletta, V.R., Bentley, M.J., Bettadpur, S., Briggs, K.H., Bromwich, D.H., Forsberg, R., Galin, N., Horwath, M., Jacobs, S., Joughin, I., King, M.A., Lenaerts, J.T.M., Li, J., Ligtenberg, S.R.M., Luckman, A., Luthcke, S.B., McMillan, M., Meister, R., Milne, G., Mouginot, J., Muir, A., Nicolas, J.P., Paden, J., Payne, A.J., Pritchard, H., Rignot, E., Rott, H., Sørensen, L.S., Scambos, T.A., Scheuchl, B., Schrama, E.J.O., Smith, B., Sundal, A.V., van Angelen, J.H., van de Berg, W.J., van den Broeke, M.R., Vaughan, D.G., Velicogna, I., Wahr, J., Whitehouse, P.L., Wingham, D.J., Yi, D., Young, D., & Zwally, H.J. A Reconciled Estimate of Ice-Sheet Mass Balance. Science. 338, 1183-1189 (2012).

9.            Velicogna, I. Increasing rates of ice mass loss from the Greenland and Antarctic ice sheets revealed by GRACE. Geophysical Research Letters. 36, (2009).

10.            Frezzotti, M., Scarchilli, C., Becagli, S., Proposito, M., & Urbini, S. A synthesis of the Antarctic surface mass balance during the last 800 yr. The Cryosphere. 7, 303-319 (2013).

11.            Thomas, E.R., Marshall, G.J., & McConnell, J.R. A doubling in snow accumulation in the western Antarctic Peninsula since 1850. Geophysical Research Letters. 35, L01706 (2008).

12.          Winkelmann, R., Levermann, A., Martin, M.A., & Frieler, K. Increased future ice discharge from Antarctica owing to higher snowfall. Nature. 492, 239-243 (2012).

13.          Barrand, N.E., Hindmarsh, R.C.A., Arthern, R., Williams, C.R., Mouginot, J., Scheuchl, B., Rignot, E., Ligtenberg, S.R.M., van den Broeke, M.R., Edwards, T.L., Cook, A.J., & Simonsen, S.B. Computing the volume response of the Antarctic Peninsula Ice Sheet to warming scenarios to 2200. Journal of Glaciology. 59, 397-409 (2013).

14.          Ligtenberg, S.R.M., Berg, W.J., Broeke, M.R., Rae, J.G.L., & Meijgaard, E. Future surface mass balance of the Antarctic ice sheet and its influence on sea level change, simulated by a regional atmospheric climate model. Climate Dynamics. 41, 867-884 (2013).

An introduction to Glacier Mass Balance

Glacier mass balance | Measuring mass balance | Mass balance gradients | Mass balance through time | Further readingReferences | Comments |

Glacier mass balance

Glacier mass balance and atmospheric circulation. By NASA. From Wikimedia Commons.

Glacier mass balance and atmospheric circulation. By NASA. From Wikimedia Commons.

The mass balance of a glacier is a concept critical to all theories of glacier flow and behaviour. It is simple enough, really: mass balance is simply the gain and loss of ice from the glacier system1. A glacier is the product of how much mass it receives and how much it loses by melting.

Mass balance can be thought of as the ‘health of a glacier’; glaciers losing more mass than they receive will be in negative mass balance and so will recede. Glaciers gaining more mass than they lose will be in positive mass balance and will advance. Glaciers gaining and losing approximately the same amount of snow and ice are thought of as ‘in equilibrium’, and will neither advance nor recede.

For clarification: when we talk about glaciers advancing, receding or being in equilibrium, we are talking about the position of their snout. Glaciers will continually flow under the force of gravity; ice is continually being moved from the upper reaches to the lower reaches, where it melts.

Accumulation zone

Unnamed Glacier, Ulu Peninsula, James Ross Island. Small valley glacier.

Unnamed Glacier, Ulu Peninsula, James Ross Island. The accumulation zone for this glacier extends from the plateau downwards.

The glacier system receives snow and ice through processes of accumulation. Surface accumulation processes include snow and ice from direct precipitation, avalanches and windblown snow. There may be minor inputs from hoar frost. The snow and ice is then transferred downslope as the glacier flows.

Precipitation falling as rain is usually considered to be lost to the system. Internal accumulation may include rain and meltwater percolating through the snowpack and then refreezing. Basal accumulation may include freezing on of liquid water at the base of the glacier or ice sheet2.

The figure below summarises the inputs and outputs from a glacier system; the inputs are the processes of accumulation (including precipitation (snow, hail and rain) and other sources of accumulation such as wind-blown snow and avalanching.

The Glacier as a System. Inputs are largely from precipitation, and also from wind-blown snow and avalanches. The glacier loses mass (ablates) mainly by the processes of calving and surface and subaqueous melt. In this simplified figure, processes of internal and basal accumulation are ommitted. See Cogley et al. 2011 for more information.

Ablation zone

Meltwater stream on Mendenhall Glacier, Alaska. From: Gillfoto, Wikimedia Commons

Meltwater stream on Mendenhall Glacier, Alaska. From: Gillfoto, Wikimedia Commons

The Glaciers as a System figure above summarises the key processes of ablation for a glacier.

Glaciers lose mass through processes of ablation. Surface ablation processes include surface melt, surface meltwater runoff, sublimation, avalanching and windblown snow. Glaciers on steep slopes may also dry calve, dropping large chunks of ice onto unwary tourists below. Glaciers terminating in the sea or a lake will calve photogenic icebergs.

Other processes of ablation include subaqueous melting, and melting within the ice and at the ice bed, which flows towards the terminus2.


Equilibrium line altitude

Accumulation usually occurs over the entire glacier, but may change with altitude. Warmer air temperatures at lower elevations may also result in more precipitation falling as rain. The zone where there is net accumulation (where there is more mass gained than lost) is the accumulation zone. The part of the glacier that has more ablation than accumulation is the ablation zone. Where ablation is equal to accumulation is the Equilibrium line altitude.

Equilibrium line altitudes in a hypothetical glacier

Equilibrium line altitudes in a hypothetical glacier

The snowline at the end of the summer season is often used to demarcate the equilibrium line on satellite images of glaciers. Above the snowline, where there is more accumulation than ablation, snow remains all year around and the glacier is a bright white colour. Below the snowline, there is more ablation than accumulation, so there is no snow left at the end of the summer, and the duller, grey-blue coloured glacier ice is visible.

The figure below shows an outlet glacier of the North Patagonian Icefield. The bright white parts in the upper reaches of the glacier are in the accumulation zone; the darker more blue areas on the glacier tongues are in the ablation zone. The Equilibrium Line Altitude here is approximately equal to the snow line.

Medial moraines on the North Patagonian Icefield (Landsat image). Each medial moraine separates out an individual flow unit.

So what is Glacier Mass Balance?

So, glacier mass balance is the quantitative expression of a glacier’s volumetric changes through time.In the figure below, Panel A shows how temperature varies with altitude. It is colder at the top than it is at the bottom of the glacier. This is crucial, as surface air temperature strongly controls melting and accumulation (as in, how much precipitation falls as snow or ice).

Mass balance (b) is the product of accumulation (c) plus ablation (a). Mass balance (b) = c + a Mass balance is usually given in metres water equivalent (m w.e.). It varies over time and space; accumulation is greater in the higher reaches of the glacier, and ablation is greater in the lower, warmer reaches of the glacier (Panel B in the figure).

Mass balance also varies throughout the year; glaciers typically get more accumulation in the winter and more ablation in the summer (Panel C in the figure). Glacier mass balance therefore usually can therefore be expressed as a mass balance gradient curve, showing how c + a varies attitudinally across the glacier (Panel D in the figure). The balance gradient is the rate of change of net balance with altitude3. A glacier’s net mass balance is a single figure that describes volumetric change across the entire glacier across the full balance year.

Principles of glacier mass balance

Principles of glacier mass balance

Measuring Mass Balance

Jonathan Carrivick prepares to stake out Glacier IJR45 on James Ross Island.

Jonathan Carrivick prepares to stake out Glacier IJR45 on James Ross Island.

Glacier mass balance is normally measured by staking out a glacier. A grid of ‘ablation stakes’ are laid out across a glacier and are accurately measured. They can be made of wood, plastic, or even bamboo like you’d use in your garden. These stakes provide point measurements at the glacier surface, providing rates of accumulation and ablation.

These methods are time consuming, logistically challenging and arduous; the stakes will need to be visited several times through the balance year. Accumulations and ablation are generally measured by reference to stakes inserted to a known depth into the glacier, and fixed by freezing and packing in3. The location is fixed with GPS.

Automatic weather stations on the glacier surface are key to understanding energy fluxes on the glacier. Probing, snowpits and crevasse stratigraphy are also used to measure mass balance on glaciers, ideally supplemented with stakes.

Remote sensing of glacier mass balance is obviously a good alternative, as it allows many glaciers to be assessed using desk-based studies. It is a cheap and simple alternative to arduous fieldwork, but ground truthing of mass balance measurements will always be necessary. Researchers from Aberystwyth University use satellite measurements to track changes in the mass balance of the Greenland Ice Sheet.

Mass balance gradients

Mass balance gradients of some typical glaciers.

The mass balance gradient of a glacier is a key control in factors such as the glacier’s response time. A glacier’s mass balance gradient is critically determined by the climatic regime in which it sits; temperate glaciers at relatively low latitudes, such as Fox Glacier in New Zealand, may be sustained by very high precipitation. They will therefore have a greater mass balance gradient (more accumulation, more ablation). These wet, maritime glaciers may have a shorter response time and higher climate sensitivity than cold, polar glaciers that receive little accumulation but also have correspondingly low ablation. These cold, dry glaciers may respond more slowly to climate change.

In the figure on the left, temperate glaciers with greater mass balance gradients are represented by the shallower lines; more mass is transferred from the top to the bottom of the glacier. Cold, polar-type glaciers with smaller mass balance gradients are represented by the steeper lines.

Mass balance through time

The Cumulative mass balance is the mass of the glacier at a stated time, relative to its mass at some earlier time. Some glaciers have mass balance measurements going back decades, which means that scientists can analyse how mass balance is changing over time.

These measurements give us detailed information about climate change, as glaciers are sensitive ‘barometers’ to our changing world. Usually, the net mass balance over the balance year is plotted on a graph. There are several projects monitoring glaciers all over the world, and these analyses show that glacier mass balance is generally decreasing (becoming more negative) over time.

30 year glacier mass balance for 30 reference glaciers in the Alps.

30 year glacier mass balance for 30 reference glaciers in the Alps. From the World Glacier Monitoring Service and Alpine Glacier Mass Balance.

In Europe, European Environment Agency has records of many glaciers, and makes their cumulative mass balance measurements publically available. The Glaciers (CLIM 007) analysis shows that the vast majority of European glaciers are receding, with the rate of recession accelerating since the 1980s.

Cumulative specific net mass balance of European glaciers (mm water equivalent) from 1946 to 2010

Cumulative specific net mass balance of European glaciers (mm water equivalent) from 1946 to 2010. From the Glaciers (CLIM 007) assessment.

The North American region shows a similar trend, with a generally declining mass balance each year.

North American glacier mass balance. Image courtesy of Mauri Pelto

North American glacier mass balance. Image courtesy of Mauri Pelto

Further afield, the IPCC AR4 shows cumulative specific net mass balance of glacierised regions worldwide. The differing behaviours of different regions shows the variable strength of climate change.

Cumulative mean specific mass balances (a) and cumulative total mass balances (b) of glaciers and ice caps, calculated for large regions (IPCC AR4)

Cumulative mean specific mass balances (a) and cumulative total mass balances (b) of glaciers and ice caps, calculated for large regions (IPCC AR4)

Further reading

More information on glacier accumulation and ablation

How glaciers flow:

Also of interest:

Wider reading:


1.            Benn, D.I. &Evans, D.J.A. Glaciers & Glaciation. London: Hodder Education. 802 (2010).

2.            Cogley, J.G., Hock, R., Rasmussen, B., Arendt, A., Bauder, A., Braithwaite, R.J., Jansson, P., Kaser, G., Moller, M., Nicholson, L., & Zemp, M. Glossary of Glacier Mass Balance and related terms. Paris: IHP-VII Technical Documents in Hydrology No. 86, IACS Contribution No. 2, UNESCO-IHP. 124 (2011).

3.            Hubbard, B. &Glasser, N.F. Field Techniques in Glaciology and Geomorphology. Wiley. 412 (2005).

Palaeo ice sheet reconstruction

Why should we reconstruct past ice sheets? | The past is the key to the future | How do we reconstruct past ice-sheet change? | How do we relate ice-sheet change to climate? | Further Reading | Comments |

Why should we reconstruct past ice sheets?

Glacial geologists love to go out into the field, collecting rock samples and bags of gravel. They spend hours mapping a single moraine in detail. They hammer tubes into beds of sand and cover themselves, and the sand, in black bin liners. But what is the focus of all this effort? Why do we care how big or how thick past ice sheets were?

Instrumental record. This image shows the instrumental record of global average w:temperatures as compiled by the w:NASA's w:Goddard Institute for Space Studies. (2006) "Global temperature change". Proc. Natl. Acad. Sci. 103: 14288-14293. Following the common practice of the w:IPCC, the zero on this figure is the mean temperature from 1961-1990. This figure was originally prepared by Robert A. Rohde from publicly available data and is incorporated into the Global Warming Art project. Wikimedia Commons.

Instrumental record. This image shows the instrumental record of global average w:temperatures as compiled by the w:NASA’s w:Goddard Institute for Space Studies. (2006) “Global temperature change”. Proc. Natl. Acad. Sci. 103: 14288-14293. Following the common practice of the w:IPCC, the zero on this figure is the mean temperature from 1961-1990. This figure was originally prepared by Robert A. Rohde from publicly available data and is incorporated into the Global Warming Art project. Wikimedia Commons.

Scientists reconstruction past ice sheets because they want to know how glaciers and ice sheets interact with climate and with the ocean. We can observe modern glaciers melting; we can look at the surface of ice sheets with satellite images and calculate changes in mass balance and we can map mountain glacier recession. It’s pretty obvious that glaciers are shrinking and melting worldwide. It’s also clear that air temperatures are increasing, ocean currents are penetrating deeper onto the Antarctic continental shelf, and precipitation patters are changing. What does this mean for our mountain glaciers and ice sheets? We know that they are already contributing to sea level rise, but by how much? And by how much will they change in the future?

It is the role of palaeo ice sheet reconstruction, or palaeoglaciology, to answer these questions. We investigate past ice-sheet and glacier response to climate change to understand:

The past is the key to the future

Sea level rise to 2100. Modified from the IPCC sea level rise estimates (from Wikimedia Commons) and using estimates from Bamber and Aspinall 2013, assuming a uniform rate of sea level rise.

Sea level rise to 2100. Modified from the IPCC sea level rise estimates (from Wikimedia Commons) and using estimates from Bamber and Aspinall 2013, assuming a uniform rate of sea level rise.

By examining how ice sheets responded to change in the past, glaciologists hope to uncover details that will help them understand how they are likely to change in the future. By looking at how ice sheets changed over long timescales, they can extend the short observation period over the Antarctic Ice Sheet. For example, ice streams in Antarctica have been observed to change, accelerate, switch off and recede. Is this normal behaviour? By looking at past ice stream change, for example in the last British Ice Sheet, we can gain a far better understanding of how ice streams evolve over far longer timescales.

Thresholds and tipping points are crucial. We worry that the West Antarctic Ice Sheet is unstable and could catastrophically collapse. Has this happened before? Will it pass a threshold and then collapse quickly, or slowly melt away? What is this threshold? We can only know by looking at the palaeo record.

Scientists also hope to better understand processes of change. The bed of the Greenland and Antarctic ice sheet is difficult to access; far better to examine the exposed bed of ancient ice sheets, like the last British-Irish Ice Sheet!

The ultimate goal of palaeo ice sheet reconstruction is gain a better understanding of how ice sheets and glaciers responded to change in the past. This will enable us to predict how they will respond to change in the future, and this will mean that we can give more precise and accurate estimates of future sea level rise.

How do we reconstruct past ice-sheet change?

Fortunately, we have many tools at our disposal to reconstruct past ice sheets. Glacial geologists usually have three main objectives:

Past ice extent and thickness

Measuring moraines in Antarctica. You can't go wrong with a notebook, tape measure, camera and pencil!

Measuring moraines in Antarctica. You can’t go wrong with a notebook, tape measure, camera and pencil!

Reconstructing past ice extent means finding past moraines and glaciated terrain. Glacial geologists compile databases and maps of these moraines, interpolating between them and creating isochrones of past ice-sheet extent.

Glacial geologists can reconstruct past ice-sheet thickness in a number of ways. They can use trimlines to demarcate the height of the former ice surface. They can use equations that predict ice thickness by assuming a certain surface slope to reach a certain extent. And they can use measurements of the Earth’s isostatic rebound to calculate the past volume of ice.

Examining past sea level rise in far-field locations, such as the Bahamas, gives us a record of changing global ice volume over long timescales. Examining local sea level rise in near-field locations (where near-field means close to a present or former ice sheet) helps us understand rates of isostatic adjustment, and hence local ice volume.

Past ice flow

Recent research shows that past ice flow was not isochronous; in fact, the last British Ice Sheet had numerous complex phases of ice flow. Ice streams flowed in different places at different times as ice divides changed in response to changing ice sheet thickness. Mapping landforms such as mega scale glacial lineations, drumlins, roche moutonnées and streamlined, striated bedrock helps glacial geologists to reconstruct these changing ice flow pathways and dynamic ice stream behaviour.

Massive thrusts in tectonised sediments in Norfolk yield insights as to palaeo ice sheet dynamics

Massive thrusts in tectonised sediments in Norfolk yield insights as to palaeo ice sheet dynamics

Looking in detail at the sediments laid down at the ice-bed interface allows glacial geologists to understand in detail the processes of past ice flow. How quickly did the ice flow? Was it cold-based or wet-based? Did the ice flow due to deformation of its bed, slipping, or deformation of the ice? By looking at sections of sediment under the microscope, we can infer detailed information about the processes of sediment deposition.

Looking at sediments and landforms together as a whole allows glacial geologists to reconstruct a Glacial Landsystem that they can use to understand the style of glaciation and take a broader view of processes and external influences.

Dating past ice fluctuations

Taking cosmogenic nuclide samples in Erratic Valley. Photo credit: Ian Hey

Taking cosmogenic nuclide samples in Erratic Valley, Alexander Island, Antarctic Peninsula. Photo credit: Ian Hey

Of course, it is all very well understand how big an ice sheet was and what direction the ice flowed in. But this is not useful for understanding rates and magnitudes of change if we cannot put it into some kind of time-scale context.  Fortunately, we have many techniques at our disposal.

Cosmogenic nuclide dating is extensively used to date the formation of moraines and the speed of glacier recession. This technique gives an exposure age, that is, the time since the boulder was exposed at the Earth’s surface. Cosmogenic nuclide dating in this context gives a maximum age for moraine formation; the moraine must be younger than the exposure age (unless your boulder has rolled, moved, has an inheritance, or any of the other multitude of factors that may result in an inaccurate age). Cosmogenic nuclide dating can also be used to date trimlines and constrain past ice thickness.

Radiocarbon dating is an essential part of the glacial geologists’ toolkit. This technique relies on organic remains. One example of an application may be for a lake dammed by a moraine. A radiocarbon age from the base of the sediment core gives a minimum age for moraine formation; the moraine must be older than the radiocarbon age. Radiocarbon ages are also extensively used from marine sediment cores around the margins of ice sheets, such as the Antarctic Peninsula Ice Sheet. Transitional Glaciomarine Sediments, those glaciomarine sediments laid down immediately after ice sheet recession and that overlie subglacial tills, provide a minimum age for ice-sheet extent at the core’s position; the ice margin was at this position before the radiocarbon age.

Optically stimulated luminescence is used to date beds of sand that accumulated in front of the ice margin by proglacial rivers. Like any other technique, dating glaciofluvial sands has its own challenges, but it can be used to provide limiting ages on ice-sheet extent.

How do we relate ice-sheet change to climate?

With all the data that has been collected mapping out former ice sheet extent and thickness, changes in its flow regime and rates of thinning and recession, we have a good idea of how the ice sheet evolved through time. The next stage is to examine the external forcings that drove this change. There are several ways of doing this:

  • By using ice cores and other proxies to reconstruct past climate
  • By relating cryospheric change to climate through a numerical model

Reconstructing past climate

Ice sheets and glaciers are an integral part of the atmosphere-ocean-cryosphere system. Change in one inevitably leads to change in all others. So if we want to understand what may have driven past ice sheet change, we need to understand past climate. Fortunately, there are many ways in which we can do this.

This photograph shows an ice core sample being taken from a drill. Photo by Lonnie Thompson, Byrd Polar Research Centre, Ohio State University. From Wikimedia Commons.

This photograph shows an ice core sample being taken from a drill. Photo by Lonnie Thompson, Byrd Polar Research Centre, Ohio State University. From Wikimedia Commons.

Ice cores preserve the composition of the ancient atmosphere. We have detailed accounts of past air temperature over the last 800,000 years. Cores from the ocean also tell us how air temperatures and global ice volume has changed over millions of years. We have detailed records of past environmental change from proxy records from pollen, beetles, insects, algae and in fact any organism that lives within limited environmental conditions. Together, all these data build up a picture of global to local environmental change over short (decadal to centennial) to longer (millennial to epoch) timescales.

Numerical ice sheet modelling

The equations that govern ice flow and mass balance (the balance between snow gained and snow melted) are complex. They depend on velocity, thickness, thermal gradients, ice temperature, air temperature lapse rates, and many other factors.

Fortunately, these physical relationships are reasonably well understood and can be fed into a computer model. These computer models can take in input data (such as past climate from an ice core) and drive a numerical ice sheet model, allowing scientists to investigate detailed processes and these key questions of past rates and magnitudes of change. Models that provide a good fit to the geological record have a high confidence, and allow scientists to investigate atmosphere-ocean-cryosphere relationships. Ultimately, all these data together help scientists to answer the critical question of how the cryosphere will change over coming centuries, and how much global sea levels will rise.

As an example, I finish with this simulation of the glaciers in Yosemite National Park responding to climate change during the Last Glacial Maximum, 21,000 years ago.

Further reading