Greenland Ice Sheet mass balance

By Dr Tom Slater, University of Leeds, UK

How does mass balance vary over Greenland?

The mass balance of the Greenland Ice Sheet is the net difference between ice gains through snowfall, and ice losses through melting at its surface or underneath its floating ice tongues, and through the calving of icebergs from glaciers flowing into the ocean.

How is Greenland Ice Sheet mass balance changing?

Contemporary changes in Greenland Ice Sheet mass balance

Between 1992 and 2018, the Greenland Ice Sheet lost more ice through ablation than it gained through accumulation, losing 3.9 trillion tonnes of ice in total at an average rate of 150 billion tonnes per year5. During this period the rate of ice loss from Greenland increased seven-fold, rising from 34 billion tonnes per year in the 1990s to 234 billion tonnes per year in the 2010s. Approximately 360 billion tonnes of ice loss will raise global sea levels by 1 mm.

Unlike Antarctica, which is losing almost all of its mass through ice dynamics, recent losses in Greenland have been almost equally split between dynamic losses and decreasing surface mass balance5. While ice discharge has increased6, the acceleration in ice losses has been mainly driven by increased runoff as the atmosphere above Greenland has warmed, melting more of the surface during the summer7.

How is Greenland Ice Sheet mass balance expected to change in the future?

Climate models predict that Greenland will continue to lose ice this century8. As observed over the past three decades, surface melting and runoff will continue to increase in Greenland as the climate warms and dominate its mass balance in future. The largest losses are expected from southwest Greenland9 which typically experiences the most melting due to how warm air is transported over the ice sheet.

As more of Greenland’s surface melts it becomes darker, reduces the surface albedo  and absorbs more solar radiation, creating a positive-feedback loop which exacerbates melt. Although Greenland’s surface mass balance will be the main source of ice loss in future, ice flow through its marine terminating glaciers is also expected to speed-up, increasing rates of iceberg calving and ice retreat9.

Based on an ensemble of 256 ice sheet models, Greenland is expected to raise global sea levels by between 2 and 10 cm by 21008.

What are the different components of mass balance in Greenland?

In order to understand how Greenland is changing now and how it might change in the future, we must account for each individual component of its mass balance.

Surface mass balance

Greenland Ice Sheet surface mass balance is the net difference between ice gains (accumulation) and losses (ablation) at the ice sheet surface1. Greenland gains ice through precipitation, mainly as snowfall.

Ice is lost through melting of snow and ice at the surface when temperatures exceed 0 °C, most of which flows downhill into the ocean (runoff), and to a lesser extent through sublimation (the change of solid ice to vapour), evaporation (the change of liquid water on the surface to vapour), and the erosion of snow in windy weather.

Unlike Antarctica, runoff is a much larger component of Greenland’s surface mass balance, which is further away from the poles and warms above 0 °C during the summer at low altitudes2.

surface meltwater on the Greenland Ice Sheet
Surface meltwater flows off the Greenland Ice Sheet and into the ocean through channels, which form over several summers. Credit: Ian Joughin

Ice dynamical changes

Ice dynamical changes relate to changes in ice discharge into the ocean and in the speed of ice flow. The Greenland Ice Sheet is drained by a series of narrow marine terminating glaciers dotted along the coast, where ice is lost to the ocean through the calving of icebergs and ocean driven melting underneath their floating ice tongues3. If calving increases, then there is more mass lost to the oceans. Calving may increase because ice velocity increases, because ocean currents change, or because the glaciers are thinning and increasingly floating.

Ice velocity, Greenland Ice Sheet
Map showing the locations of Greenland’s ice streams and marine terminating glaciers from ice speed measured through satellite optical imagery. Data taken from Gardner et al., 20194

How do we measure Greenland Ice Sheet mass balance?

The Greenland Ice Sheet is massive (about 9 times bigger than the UK, for example) and remote – how do we measure it’s mass balance? This is only possible through satellites orbiting the Earth, which provide repeat observations and comprehensive coverage over the polar ice sheets. Satellites launched by the European Space Agency and NASA have continuously monitored the ice sheets since the 1990s and allow scientists to measure Greenland Ice Sheet mass balance in 3 ways:

  • Altimetry reveals changes in ice sheet volume through measuring the height of the surface, which can be related to mass through the density of the ice lost or gained10.
  • Optical/Radar imagery provide measurements of ice speed which can be used to measure ice discharge through Greenland’s marine terminating glaciers11.
  • Gravimetry measures changes in ice mass through changes in Earth’s gravity field12.
Ice height change, Greenland Ice Sheet
Maps of elevation change from satellite altimetry reveal where the Greenland Ice Sheet is changing mass. Map created using data acquired by the CryoSat-2 satellite radar altimeter. Credit: CPOM.

Validating satellite measurements in the field

To increase confidence in satellite measurements, we can validate them by travelling to the Greenland Ice Sheet and conducting field campaigns. On the ground, we can collect ice cores to measure snow properties such as density and stratigraphy, which improve our understanding of the surface conditions observed by the satellite13. We can also quantify the accuracy of satellite measurements directly, by acquiring overlapping measurements from similar sensors mounted on aircraft14.

Validating glacier mass changes on the Greenland Ice Sheet. Scientist with helicopter.
Scientists acquire ice cores on the Greenland Ice Sheet to help validate satellite measurements. Credit: Anna Hogg.

About the author

Dr Tom Slater is a research fellow at the NERC Centre for Polar Observation at the University of Leeds. His research focusses on using satellite radar altimetry to study the Antarctic and Greenland ice sheets, and measure their contribution to global sea levels.

Webpage: Tom Slater

Twitter: @_tslater


1.         Lenaerts, J. T. M., Medley, B., Broeke, M. R. van den & Wouters, B. Observing and Modeling Ice Sheet Surface Mass Balance. Reviews of Geophysics 57, 376–420 (2019).

2.         Noël, B., Berg, W. J. van de, Lhermitte, S. & Broeke, M. R. van den. Rapid ablation zone expansion amplifies north Greenland mass loss. Science Advances 5, eaaw0123 (2019).

3.         King, M. D. et al. Dynamic ice loss from the Greenland Ice Sheet driven by sustained glacier retreat. Communications Earth & Environment 1, 1–7 (2020).

4.         Gardner, A. S., Fahnestock, M. A. & Scambos, T. A. ITS_LIVE Regional Glacier and Ice Sheet Surface Velocities. (2019).

5.         The IMBIE Team. Mass balance of the Greenland Ice Sheet from 1992 to 2018. Nature 579, 233–239 (2020).

6.         Enderlin, E. M. et al. An improved mass budget for the Greenland ice sheet. Geophysical Research Letters 41, 866–872 (2014).

7.         Hanna, E., Mernild, S. H., Cappelen, J. & Steffen, K. Recent warming in Greenland in a long-term instrumental (1881–2012) climatic context: I. Evaluation of surface air temperature records. Environ. Res. Lett. 7, 045404 (2012).

8.         Edwards, T. L. et al. Projected land ice contributions to twenty-first-century sea level rise. Nature 593, 74–82 (2021).

9.         Goelzer, H. et al. The future sea-level contribution of the Greenland ice sheet: a multi-model ensemble study of ISMIP6. The Cryosphere 14, 3071–3096 (2020).

10.       Simonsen, S. B., Barletta, V. R., Colgan, W. T. & Sørensen, L. S. Greenland Ice Sheet Mass Balance (1992–2020) From Calibrated Radar Altimetry. Geophysical Research Letters 48, e2020GL091216 (2021).

11.       Mouginot, J. et al. Forty-six years of Greenland Ice Sheet mass balance from 1972 to 2018. PNAS 116, 9239–9244 (2019).

12.       Velicogna, I. et al. Continuity of Ice Sheet Mass Loss in Greenland and Antarctica From the GRACE and GRACE Follow-On Missions. Geophysical Research Letters 47, e2020GL087291 (2020).

13.       Otosaka, I. N. et al. Surface Melting Drives Fluctuations in Airborne Radar Penetration in West Central Greenland. Geophysical Research Letters 47, e2020GL088293 (2020).

14.       MacGregor, J. A. et al. The Scientific Legacy of NASA’s Operation IceBridge. Reviews of Geophysics 59, e2020RG000712 (2021).


Videos about glacier mass balance

Watch this brief introductory video, made by Time for Geography with Bethan Davies and Simon Cook, about glacier mass balance. This is suitable for GCSE and A-level students.

Video on glacier mass balance by Time for Geography

This more in-depth lecture (29 minutes) introduces the concept of glacier mass balance, and then goes on to discuss how glaciers worldwide are shrinking, and what this means for global water resources and sea level rise. This lecture is aimed at A-Level and Post-16 students.

This lecture by Bethan Davies, from the Reading Climate Festival in November 2020, goes into more detail about global glacier change (1 hr).

From Snow to Firn to Glacier ice


How do we build a glacier? We start with a snowflake. Snow, over time, is compressed into firn, and then into glacier ice.

Snow falls in cold regions, such as mountain tops or in polar regions. In glaciology, snow refers to material that has not changed since it fell1.

Snow is very light and fluffy, and has a very low density. If the snow is wetter, it will have an increased density. Snowflakes have a hexagonal structure, and fallen snow has a significant amount of air in it.

Snow flakes by Wilson Bentley. Bentley was a bachelor farmer whose hobby was photographing snow flakes. ; Image ID: wea02087, Historic NWS Collection ; Location: Jericho, Vermont ; Photo Date: 1902 Winter. From Wikimedia Commons


Firn is usually defined as snow that is at least one year old and has therefore survived one melt season, without being transformed to glacier ice.

Firn is transformed gradually to glacier ice as density increases with depth, as older snow is buried by newer snow falling on top of it. Year after year, successive accumulation layers are built up. In the accumulation zone of a glacier, density therefore increases with depth; the rate depends on the local climate and rate of accumulation1. Firn transforms to glacier ice at a density of 830 kg m-3.

New snow (immediately after falling, calm conditions50-70
Damp new snow100-200
Settled snow200-300
Wind-packed snow350-400
Very wet snow and firn700-800
Glacier ice830-923
Typical densities (kg m-3). From Cuffey and Paterson, 2010.
A scientist collecting snow and ice samples from the wall of a snow pit. Fresh snow can be seen at the surface and en:glacier ice at the bottom of the pit wall. The snow layers are composed of progressively denser en:firn. Taku Glacier, Juneau Icefield, en:Tongass National Forest, en:Alaska. From Wikimedia Commons

Firn transforms to glacier ice in 3-5 years in the temperate Upper Seward Glacier in the St Elias Mountains near the Alaska-Yukon border. Firn becomes ice at a depth of about 13 m1. At sites like this with rapid snow accumulation, the depth of a firn layer, and the load on it, increases rapidly with depth.

However, in cold, dry East Antarctica, firn becomes ice at a depth of 64 m at Byrd and 95 m at Vostok. 280 years are needed at Byrd, and 2500 at Vostok. Low temperatures slow the transformation. Temperatures at Vostok, the coldest region of Earth, are 30°C lower than Byrd, which explains the slower increase in density. In addition, slow accumulation gives slow burial, and a small load each year; the increase in density takes much longer.

Typically, the transformation of firn to ice takes 100-300 years, and a depth of 50 – 80 m1.

Glacier ice

Firn becomes glacier ice when the interconnecting air or water-filled passageways between the grains are sealed off (“pore closure”)1. Air is isolated in separate bubbles. This occurs at a density of 830 kg m-3. The air space between particles is reduced, bonds form between them, and the particles grow larger. This is a process known as sintering. Increasing pressure compresses the bubbles, placing the enclosed air under pressure and increasing the density of the ice2.

Fresh snowflakes, which have a complex shape, have a large surface area. Over time and under pressure, the surface area is reduced, the surface is smoothed, and the total surface area is reduced. Fresh, complex snowflakes are transformed into rounded particles.

Formation of glacier ice. Luis Maria Benitez, Wikimedia Commons

The transformation of firn to ice is much faster where there is melting and refreezing2.  Meltwater can percolate downwards, infilling porespaces, and the displaced air escapes upwards. If the snow is under 0°C, the water will freeze, producing areas of compact ice.  This will produce high density ice much more rapidly than in colder regions without melting.

The density of pure glacier ice is usually taken as 917 kg m-3. This strictly is only true at 0°C and in the upper layers of ice sheets and mountain glaciers; the density may be greater at the mid-depth ranges in polar ice sheets, where there are no bubbles and temperatures are -20°C to -40°C1.

Below 4 km of ice, such as at the centre of the East Antarctic Ice Sheet, the pressure would increase the density to 921 kg m-3.


Bubbles are common in glacier ice. Bubbles can contain liquid water or atmospheric gases, making them very useful for ice core research. The air in the bubble largely reflects the atmospheric concentrations when the ice formed1. In polar environments, bubbles in the ice occupy about 10% of the volume when firn turns to ice.

Glacier ice with many bubbles exposed on the ice shelf. It is melting and thinning rapidly.
Close up of white bubble-rich ice. Note the sharp junction between the coarse-clear ice and bubble-rich ice.

With greater depth in polar ice sheets, bubbles shrink as the overlying ice increases. The gas pressure within the bubbles therefore increases, and at certain depths, the gas attains a dissociation pressure. The bubbles begin to disappear as the gas molecules form clathrate hydrates1.  This process takes thousands of years.


Glacier ice contains various impurities in tiny amounts. By most scales, glacier ice is a very pure solid-earth material, because the processes leading to snowfall on a glacier – evaporation, condensation, precipitation – act as a natural distillation system1.

However, glaciers can contain impurities. The dirtiest glaciers are mountain glaciers, where debris can fall directly onto the ice surface. On ice sheets and glaciers, dust and other debris may blow onto the ice surface.

Iceberg laden with debris from a glacier, Antarctic Peninsula

Debris on the ice surface can affect the absorption of energy at the ice surface, and lead to increased or decreased melting.

Supraglacial debris on Unnamed Glacier, James Ross Island, Antarctic Peninsula

Layers in the ice

Glaciers are composed of sedimentary layers in their accumulation zones, formed of annual layers of snowfall. These layers are initially parallel to the glacier surface. This is the primary stratification in structural glaciology.

In temperate and subpolar settings, the annual sedimentary layers consist of alternating thick layers of bubble-rich ice, which originated as winter snow, and thin layers of clear ice, which are the refrozen meltwater from the summer melt season.

Glacier ice exposed in an ice-cored moraine. Note the foliation with coarse clear ice and white bubble-rich ice.
Primary stratification on a glacier on James Ross Island, Antarctic Peninsula.

Debris horizons may form, when summer melting concentrates debris (such as rockfall and wind-blown dust) on the ice surface.

In cold polar regions, annual layering forms instead by seasonal variation of snow metamorphism and wind deposition1.

This 19 cm long of GISP2 ice core from 1855 m depth shows annual layers in the ice. This section contains 11 annual layers with summer layers (arrowed) sandwiched between darker winter layers. From the US National Oceanic and Atmospheric Administration, Wikimedia Commons.

Blue glacier ice

Glacier ice is blue because the longer visible wavelengths are absorbed. The more energetic, blue, wavelengths are scattered back2.  The effect is greatest with deep, basal ice, which is bubble free and has large crystals. The blue colour tends therefore to be most intense in the calls of calved icebergs or fresh fractures.

Rough, weathered ice and fresh snow will appear white because preferential absorption does not occur.

This iceberg is formed from basal glacier ice. It is blue an has basal dirt. Differential melting forms holes all over its surface.

Further reading


1.           Cuffey, K. M. & Paterson, W. S. B. The Physics of Glaciers, 4th edition. (Academic Press, 2010).

2.           Benn, D. I. & Evans, D. J. A. Glaciers & Glaciation. (Hodder Education, 2010).

Mass Balance teaching resources

This page highlights some of the excellent teaching resources available for exploring glacier mass balance.

For more ideas, see the Resources for Teachers page.

Case study: USGS Benchmark Glaciers

The USGS has an excellent resource on the mass balance of Lemon Creek Glacier, a World Reference Glacier, and the other USGS Benchmark Glaciers.

Lemon Creek Glacier, Alaska

This has resulted in a publication showing a reanalysis of the USGS Benchmark Glaciers (O’Neel et al., 2019). Point data were collected at each glacier over many years. These point datasets allow glacier-wide mass balance to be calculated. The full datasets are rather complicated, probably too much so for post-16 education, but the “Glacier-wide solutions” spreadsheets could be used to calculate glacier mass balance from annual winter and summer balances.

The data are available for download from the Alaska Science Centre.

I have produced some summary sheets and teaching resources that can be used with this dataset to understand glacier mass balance.

First, watch the video on Changing Glaciers:

Video aimed at post-16 students studying glacier mass balance (29 minutes)

The Summary Sheet gives the key points to remember about glacier mass balance.

The activity below guides students to explore the 5 USGS Benchmark Glaciers and their changing mass balance over time.

Students could also use the dataset from the WGMS to plot cumulative glacier mass balance over time.

World Glacier Monitoring Service (WGMS)

The World Glacier Monitoring Service (WGMS) provides a host of resources for the reference glaciers monitored for mass balance.

World Glacier Monitoring Service

There are annual mass balance reports, and these are presented with some clear graphics showing cumulative glacier mass balance.

Annual mass balance of reference glaciers with more than 30 years of ongoing glaciological measurements. From the WGMS
Cumulative mass change of reference glaciers. Cumulative values relative to 1976 are given. From the WGMS.

WGMS Glacier Browser

The WGMS have produced a browser where you can view the mass balance records of different reference glaciers. The map is based on ArcGIS Online and allows students to explore reference glaciers worldwide.

WGMS Fluctuations of Glaciers Browser

The numbers in the circules highlight the number of types of measurement and glaciers in each area. As you zoom in, you can click through the individual glaciers and see a graph of mass balance observations, surges, and front variations over time.

This resource allows you to explore observations on glaciers worldwide, and examine the dataset easily to see if glaciers really are receding.

Raw mass balance data

These data are all available in the following publication:

WGMS (2020, updated, and earlier reports). Global Glacier Change Bulletin No. 3 (2016-2017). Zemp, M., Gärtner-Roer, I., Nussbaumer, S. U., Bannwart, J., Rastner, P., Paul, F., and Hoelzle, M. (eds.), ISC(WDS)/IUGG(IACS)/UNEP/UNESCO/WMO, World Glacier Monitoring Service, Zurich, Switzerland, 274 pp., publication based on database version: doi:10.5904/wgms-fog-2019-12.

The WGMS provides mass balance data, which could be used by students to plot as graphs or as data: You can also download and explore the full database.

Case study: Bahia del Diablo, Vega Island, Antarctic Peninsula

An example of a student exercise could be to look at the Mass Balance Point data for a single year for a single glacier from the WGMS dataset, and plot elevation against balance for each point. Students could then plot a graph of elevation against mass balance and get the mass balance gradient through time. These are plotted in the WGMS Bulletin for comparison.

An example could be Glaciar Bahia del Diablo on Vega Island on the Antarctic Peninsula. This is a reference glacier that has been monitored since 2009.

James Ross Island and Vega Island, northern Antarctic Peninsula.

Glaciar del Diablo is on the northern side of the island, and is a land-terminating glacier.

Using the mass balance point data from the WGMS, students could attempt to plot the net balance for each point over certain years.

Net balance versus altitude from Glaciar Bahia del Diablo. WGMS 2020

OGGM Glacier Simulator

You can explore more about glacier mass balance using the OGGM Glacier Simulator.

This is an interactive web application that allows you to learn about how glaciers flow, shrink and grow, and what parameters influence their size.

The webpage has information and guided tutorials.


O’Neel, S., McNeil, C., Sass, L., Florentine, C., Baker, E., Peitzsch, E., . . . Fagre, D. (2019). Reanalysis of the US Geological Survey Benchmark Glaciers: Long-term insight into climate forcing of glacier mass balance. Journal of Glaciology,65(253), 850-866. doi:10.1017/jog.2019.66

WGMS (2019): Fluctuations of Glaciers Database. World Glacier Monitoring Service, Zurich, Switzerland. DOI:10.5904/wgms-fog-2019-12. Online access:

WGMS 2020. Global Glacier Change Bulletin No. 3 (2016-2017). Zemp, M., Gärtner-Roer, I., Nussbaumer, S. U., Bannwart, J., Rastner, P., Paul, F., and Hoelzle, M. (eds.), ISC(WDS)/IUGG(IACS)/UNEP/UNESCO/WMO, World Glacier Monitoring Service, Zurich, Switzerland, 274 pp., publication based on database version: doi:10.5904/wgms-fog-2019-12. PDF (20 MB)

Mass Balance

What is glacier mass balance?

The Mass balance of a glacier can be thought of as the health of a glacier. Mass balance is the total sum of all the accumulation (snow, ice, freezing rain) and melt or ice loss (from calving icebergs, melting, sublimation) across the entire glacier.

If glaciers have a mass balance that is in equilibrium with climate, then the inputs are equal to the outputs. The glacier remains the same size, and does not grow or shrink. Ice continues to be transferred from the top of the glacier (the accumulation zone) to the bottom of the glacier (ablation zone).

If the amount of melting across the glacier increases, then the glacier will have a negative mass balance, and the glacier will shrink.

If the amount of snow or ice that the glacier receives increases but the amount of melt stays the same, then the glacier will grow. The glacier will have a positive mass balance.

Further reading

Other pages that are relevant:

The surface energy balance

This page draws from the excellent review of glacier melt by Hock (2005) and the equally informative chapter on ‘Snow, Ice and Climate’ (Chapter 2) in Glaciers and Glaciation by Benn and Evans (2010). I highly recommend these resources if you want to dive deeper into the processes of glacier melting and the surface energy balance.

What is the surface energy balance?

The balance of energy at the glacier surface control melting.

The term surface energy balance is specifically used to describe the balance between all surface energy inputs and outputs over a given interval of time. By calculating the surface energy balance, glaciologists can work out the change in temperature at a glacier surface and rate of glacier melt1-4.

The surface energy balance is the sum of all energy fluxes at the glacier surface, and is represented by the simple equation:

QM = SW + LW + QH + QE + QR

where QM is the energy available for melting SW is shortwave radiation, LW is longwave radiation, QH is sensible heat transfer, QE is latent heat transfer, QR is the energy supplied by rain.

Meltwater at the surface of the Greenland Ice Sheet. Glacier surface melt is the product of all positive and negative energy fluxes. Source: NASA.

This page will explain each component of the energy balance equation in order. But first, we need to understand how snow/ice is melted.

The melting of snow and ice

For snow or ice to melt, its temperature must first be increased from sub-freezing (i.e. temperatures lower than 0°C) to the melting point, which is 0°C. Once at 0°C, further energy surplus will cause melting3.

For this reason, an air temperature of 0°C or warmer does not automatically result in melting. It may heat the glacier surface, but unless it raises the temperature of snow/ice to the melting point, no change in state will arise3.

The opposite transpires when there is an energy deficit at the glacier surface. This deficit either cools the ice further or forms new ice by freezing surface water or through the condensation of water vapour3,4.

For snow or ice to melt, energy is needed to first warm it to 0°C. Once at 0°C any further energy surplus will cause melting. Image shows the surface of Root Glacier, Alaska. Source: NPS / J.W. Frank.

The energy needed to change state

The relationship between energy input (in the form of heat) and the change in temperature of a material is known as the specific heat capacity. The specific heat capacity is different for different materials. To warm 1 gram of ice at –10°C by 1°C requires an energy input of 2 joules3,4.

The melting of snow and ice requires an energy input of 334 joules per gram3,4. Conversely, 334 joules are released when water is frozen. The energy released during this change of state (as well as the consumption of energy in melting) is known as latent heat.

The energy required for evaporation, and that released by condensation, is much higher than for melting/freezing, being 2500 joules per gram.

The energy (in joules per gram) required for state changes between water, snow/ice, and water vapour.

The surface energy balance

Now that we have the basics of the heating and melting of snow and ice, let’s break down the main components of the surface energy balance1-3.

Shortwave radiation

QM = SW + LW + QH + QE + QR

We’ll start with the radiation emitted by the sun, otherwise called ‘shortwave’ radiation as it occurs at short wavelengths (0.2 – 4 micrometres). The shortwave radiation receipt varies from glacier to glacier depending on several factors.


First, is latitude. The sun’s energy is focused over a much smaller area at the equator than at the poles, owing to the higher angle of the sun’s rays in the low latitudes. Therefore, more shortwave radiation is available per unit area for heating the Earth’s surface at the equator than at the poles3,4.

Some solar radiation is scattered by air/water molecules and other particles in the atmosphere as it moves towards Earth’s surface3,4. Because the distance between the sun and the poles is greater than between the sun and the equator, more energy is lost by the time it reaches the high latitudes, leaving less energy available for melting.

Shortwave radiation receipt tends to be highest at low latitude glaciers in high altitude mountain ranges (e.g. the Andes and Himalayas) where the sun angle is high and the thin, relatively cloudless air at high altitude limits the amount of solar energy lost by scattering5-7. In contrast, shortwave radiation receipts are often lower for mid- or high latitude glaciers in persistently cloudy (e.g. coastal) areas, such as Patagonia or Alaska.

Low latitude and high altitude Llaca Glacier in the Cordillera Blanca of Peru. Such glaciers receive high energy inputs from solar radiation. Source: Edubucher.


Second, the local slope gradient and aspect impact the angle at which solar rays meet the land surface and, therefore, the amount of shortwave radiation received2,8. The terrain around a glacier may also play an important role, by shading the ice from direct solar radiation8,9.

Time of day

Shortwave radiation receipt also varies as the sun angle changes throughout the day, with a peak at midday when the sun angle is highest.

Topographic shading of a Baffin Island glacier, Arctic Canada, by the adjacent valley walls. Source: M. Beauregard.

The albedo effect

Only some of the shortwave radiation that reaches Earth’s surface is absorbed. A portion is also reflected. This is known as albedo. The energy available for melting, therefore, is the difference between incoming and outgoing radiation3,4.

The amount of energy reflected depends on the surface material10. Clean ice or snow reflect more solar radiation than dirty or debris-covered ice, leaving less energy for melting.

Clean ice or snow has a higher albedo and reflects more solar radiation than dirty or debris-covered ice. Source: Papphase

Longwave radiation

QM = SW + LW + QH + QE + QR

Longwave radiation, so-called because it occurs at the ‘long’ wavelengths of 4 to 120 micrometres, is emitted from both the land surface and atmosphere. Glaciers also emit longwave radiation, so the net energy available for melting is the difference between that received and that emitted by ice3,4.

The main sources of incoming longwave radiation are the valley walls that surround glaciers and water vapour (as well as CO2 and ozone) in the atmosphere11,12. Heat emitted from warm rockwalls increase the energy available for melting, particularly around the glacier sides. As water vapour absorbs and emits longwave radiation, it is most important where the air is humid, and weather cloudy3,4.

For most glaciers, the most important components of the surface energy are the combined flux of shortwave (solar) and longwave radiation, which may supply over 75% of the energy for melting1,3,12.

Longwave radiation is emitted from terrestrial surfaces and the atmosphere. The longwave radiation from valley walls provides energy for melting snow/ice, especially at the glacier edges. Source: Ibex73.

Sensible heat transfer

QM = SW + LW + QH + QE + QR

Sensible heat is the thermal energy passed directly from one material to another, in this case, from the atmosphere to a glacier (or vice versa). It is called sensible heat as it we can sense or feel it (a cold breeze, for example).

The amount of sensible heat transferred from the atmosphere to a glacier depends on the temperature gradient near the glacier surface (i.e. the difference in temperature between the ice and air above it) and wind speed3,12.

The higher the temperature gradient and the faster the wind speed, the more sensible heat is transferred to a glacier.

There are several examples of such weather conditions. The first are Föhn winds, which are dry, strong winds that blow down the leeside of mountains, warming the air, and glacier surface, as they descend. The second are valley winds, which are warm, low-level winds that are drawn up alpine valleys as the air over mountain ranges heats up during the day.

Common wind types in alpine mountains (e.g. the European Alps) that transfer sensible heat to glaciers.

Latent heat transfer

QM = SW + LW + QH + QE + QR

Latent heat is the energy consumed or released during a change of state at the ice surface; condensation (vapour to liquid), evaporation (liquid to vapour), deposition (vapour to solid) and sublimation (solid to vapour).

Like sensible heat above, these processes depend on wind speed over the glacier and the humidity of the air at and above the ice. For this reason, sensible and latent heat transfers are often grouped together and termed the turbulent fluxes3,13.

Where the air is above the glacier is more humid than at the surface, evaporation or sublimation will occur. In contrast, where the air is above the glacier is less humid (drier) than at the surface, condensation or deposition will occur. This change in state, releases energy for warming or melting ice.

Winds are important in latent heat transfer as they can stir up the air at the ice surface (which is often humid) and mix it with the air above.

Latent heat fluxes are important at high-altitude mountain glaciers that experience cold and dry weather conditions, where sublimation occurs14. However, they can also be important in the energy balance of maritime glaciers15.

Strong, dry winds redistribute snow at the calving front of Pine Island Glacier. These conditions can cause sublimation (i.e. the change from a solid [ice] to a gas [water vapour] in the atmosphere) and latent heat release. Source: NASA.

The energy supplied by rain

QM = SW + LW + QH + QE + QR

Rainfall that falls on the surface of a glacier, although generally a minor component of the overall energy balance (except for short periods, e.g. during warm fronts or storms16), can supply energy for melting.

Rain will cool to the temperature of the surface snow or ice and ultimately, freeze. This releases latent heat that contributes to heating or melting ice. If the glacier surface is already at the melting point, then rain will add to melting directly.

Rain cloud moving over a small cirque glacier in the Jotunheimen mountains of western Norway. Source: J. Bendle.

How is glacier melt modelled?

So, we have covered the main components of glacier surface energy balance, but how is glacier melt actually modelled? Well, there are two main types of model for this purpose3.

First, are point models, where the energy balance is estimated at a single point on a glacier surface. Usually, this is at the site of a weather station.

Second, are distributed models, where the energy balance is estimated across an area1,2,3. This method has become more common in recent years as satellite datasets (such as digital elevation models) and computer power have improved.

Schematic representation of point and distributed melt models. Point models estimate glacier melt at a single location on a glacier, usually at the site of a weather station that automatically records temperature, humidity, wind speed, radiation etc. Distributed models estimate the energy balance over a larger area (often the entire glacier) using gridded datasets, such as digital elevation models. The energy balance is calculated for each square in the grid, so can take account of variations in melt across the glacier. Image from Google Earth.

Distributed models are used to investigate how the individual energy balance components influence ice melting over different parts of a glacier. Detailed information of this nature is needed to better understand how glacier ablation trends will react to climate and weather changes, and to predict how glacier mass balance will change in the future.

Why study glacier energy balance?

To sum up – why should we care about the surface energy balance of glaciers and ice sheets?

Well, mainly because it allows glaciologists to understand current – and, therefore, predict future – trends in glacier melting.

At the global scale, this information can help us estimate the glacier contribution to sea-level rise with climate change.

At the regional scale, this information helps to predict river discharge and geomorphic activity downstream of mountain glaciers, where huge numbers of people rely on glacial freshwater for drinking, the irrigation of crops, and hydroelectric power. It is also important for the forecasting of floods and the safety of downstream communities.

Nearly two billion people rely on meltwater from snow and glaciers for drinking, farming and electricity generation. The image above shows fields irrigated by Himalayan meltwater. Chapursan Valley, Pakistan. Source: Imran Shah.


[1] Arnold, N.S., Willis, I.C., Sharp, M.J., Richards, K.S. and Lawson, W.J., 1996. A distributed surface energy-balance model for a small valley glacier. I. Development and testing for Haut Glacier d’Arolla, Valais, Switzerland. Journal of Glaciology42, 77-89.

[2] Hock, R. and Holmgren, B., 2005. A distributed surface energy-balance model for complex topography and its application to Storglaciären, Sweden. Journal of Glaciology51, 25-36.

[3] Hock, R., 2005. Glacier melt: a review of processes and their modelling. Progress in Physical Geography29, 362-391.

[4] Benn, D.I., and Evans, D.J.A., 2010. Glaciers and Glaciation. Hodder-Arnold, London.

[5] Benn, D.I., Wiseman, S. and Hands, K.A., 2001. Growth and drainage of supraglacial lakes on debris-mantled Ngozumpa Glacier, Khumbu Himal, Nepal. Journal of Glaciology47, 626-638.

[6] Mölg, T., Hardy, D.R. and Kaser, G., 2003. Solar‐radiation‐maintained glacier recession on Kilimanjaro drawn from combined ice‐radiation geometry modeling. Journal of Geophysical Research: Atmospheres108 (D23).

[7] Pellicciotti, F., Helbing, J., Rivera, A., Favier, V., Corripio, J., Araos, J., Sicart, J.E. and Carenzo, M., 2008. A study of the energy balance and melt regime on Juncal Norte Glacier, semi‐arid Andes of central Chile, using melt models of different complexity. Hydrological Processes22, 3980-3997.

[8] Arnold, N.S., Rees, W.G., Hodson, A.J. and Kohler, J., 2006. Topographic controls on the surface energy balance of a high Arctic valley glacier. Journal of Geophysical Research: Earth Surface111 (F2).

[9] Olson, M. and Rupper, S., 2019. Impacts of topographic shading on direct solar radiation for valley glaciers in complex topography. The Cryosphere13, 29-40.

[10] Paterson, W.S.B., 1994. Physics of glaciers. Butterworth-Heinemann.

[11] Brock, B.W., Willis, I.C., Sharp, M.J. and Arnold, N.S., 2000. Modelling seasonal and spatial variations in the surface energy balance of Haut Glacier d’Arolla, Switzerland. Annals of Glaciology31, 53-62.

[12] Oerlemans, J. and Klok, E.J., 2002. Energy balance of a glacier surface: analysis of automatic weather station data from the Morteratschgletscher, Switzerland. Arctic, Antarctic, and Alpine Research34, 477-485.

[13] Morris, E.M., 1989. Turbulent transfer over snow and ice. Journal of Hydrology105, 205-223.

[14] Cullen, N.J., Mölg, T., Kaser, G., Steffen, K. and Hardy, D.R., 2007. Energy-balance model validation on the top of Kilimanjaro, Tanzania, using eddy covariance data. Annals of Glaciology46, 227-233.

[15] Conway, J.P. and Cullen, N.J., 2013. Constraining turbulent heat flux parameterization over a temperate maritime glacier in New Zealand. Annals of Glaciology54, 41-51.

[16] Hay, J.E. and Fitzharris, B.B., 1988. A comparison of the energy-balance and bulk-aerodynamic approaches for estimating glacier melt. Journal of Glaciology34, 145-153.

The role of debris cover on glacier ablation

Rock and sediment debris often cover part or all of a glacier’s surface, where it plays an important role in surface energy balance and the rate of glacier ablation.

Debris partly covers the surface of Hopper Glacier in the Hunza Valley of northern Pakistan. Source: J. Moshin.

The relationship between debris thickness and glacier melting

Measurements taken at the surface of glaciers show there to be a strong relationship between the thickness of debris and the rate of ice melting1,2.

Relationship between surficial debris thickness and glacier melt. The rate of melting increases to a maximum at ~2 cm debris thickness. Further debris thickness decreases the rate of ice melt. Based on refs. 1 and 6.

Where debris covering the ice surface is thin, the rate of melting rises. The rate of melting continues to rise until the debris layer reaches around 2 cm thick. Where debris cover is thicker than 2 cm, the rate of melting falls exponentially1.

This relationship is explained by both the albedo effect and the insulation effect.


Albedo (usually denoted by the Greek letter ‘α’) is the term used for the proportion of incoming shortwave (solar) radiation that is reflected by a surface3. The albedo of a surface can be determined using the following simple equation:

α = SWout / SWin

where α is albedo, SWout is outgoing shortwave radiation (i.e. the amount reflected by a surface), and SWin is incoming shortwave radiation (i.e. the amount received by a surface).

Surface Albedo
Dry snow 0.80–0.97
Melting snow 0.66–0.88
Firn 0.43–0.69
Clean ice 0.34–0.51
Slightly dirty ice 0.26–0.33
Dirty ice 0.15–0.25
Debris-covered ice 0.10–0.15
from: Paterson (1994)

The albedo of dirty and debris-covered ice (i.e. the parts of a glacier’s surface littered with bare rock and sediment) is lower than that of clean ice or snow (see table above) and, as a consequence, they absorb more incoming shortwave radiation4. This increases the amount of energy that is available for melting.

The partly debris-covered snout of Franz Josef glacier in the Southern Alps of New Zealand. Contrast the low albedo surfaces at the terminus with the highly reflective (high albedo) surfaces further upglacier. Source: Papphase


The second effect of debris on ice surface melting is insulation. Surface debris forms a barrier between the glacier and the atmosphere, reducing the amount of energy that reaches the ice surface and, therefore, insulating the ice from melting3.

Thick debris cover insulates the snout of Exploradores Glacier from melting in Patagonia, southern South America. Source: J. Bendle.

Which effect is most important?

That depends – as you may have guessed – on the thickness of debris at the glacier surface. The albedo effect has a greater influence on ablation rates where the debris cover is sparse or absent, whereas the insulation effect is more important where the debris cover is thick3.

Thin, patchy debris cover on the Biafo Glacier in the Karakoram Mountains of Gilgit Baltistan, Pakistan. In this part of the glacier, the albedo effect has a strong influence on melt rate. Source: Yousaf Feroz Gill.

Calculating the impact of surface debris on ice melt

The influence of surface debris on ice melt can be assessed by calculating how much heat is transferred vertically through a debris layer to the top of the glacier. This heat transfer is known as the conductive heat flux (Qc) and can be estimated by a simple equation5:

Qc = k (TsTi) / hd

where k is the thermal conductivity of the debris layer (i.e. the ability of the debris to conduct heat, which varies depending on rock type), Ts and Ti are the temperature at the top and base of the debris layer, and hd is the debris layer thickness.

The above equation is handy as it provides a simple method of calculating the conductive heat flux, which can then be used to calculate ice melt rates (using the equation below) beneath debris cover. However, it makes several assumptions.

Most importantly, the equation assumes that the change in temperature between the top and base of the debris layer is linear (i.e. that the temperature changes at an even rate).

In nature, however, this temperature gradient is rarely stable, but is always changing in response to fluctuations in the receipt of energy at the surface6. In short, it is non-linear. This means that the above equation may not always give a reliable estimate of the passage of heat through a debris layer and, thus, melt rate.

However, as the difference in temperature between the top and base of a debris layer is linear when averaged out over the course of a day, it is possible to get around this problem by using daily mean surface temperature data (which can be accessed from local weather stations) to calculate ice melt rates6.

Using daily mean surface temperatures, the equation above gives the daily average heat flux through a debris layer, which can be used to calculate ice melt rate (M) using the following simple equation:

M = M/Lf(M > 0)

where M is the energy flux available for melting (i.e. the daily average heat flux from above) and Lf is latent heat given off by melting.

Why should we care about the effect of debris on glacier ablation?

Many of the world’s alpine glaciers are covered by debris to some extent7, and this debris (as explained above) affects the rate of ice melting1,2,5. This, in turn, impacts the overall mass balance of glaciers, as well as the landforms produced at ice margins7.

Debris-covered terminus of the Tasman Glacier in the Southern Alps of New Zealand. Source: T. Hisgett.

Understanding the relationships between surface debris and glacier melting is also important for accurately predicting how debris-covered glaciers in regions such the Himalayas, Andes, and Southern Alps of New Zealand, will react to climate change, and whether changes in the patterns of ice melting will threaten communities living downstream (e.g. flooding)8.

Debris-covered surface in the lower part of the Khumbu Glacier, Everest region of Nepal. Source: Bokeyby.


[1] Østrem, G. 1959. Ice melting under a thin layer of moraine, and the existence of ice cores in moraine ridges. Geografiska Annaler Series A, 41, 228–230.

[2] Kayastha, R.B., Takeuchi, Y., Nakawo, M. and Ageta, Y., 2000. Practical prediction of ice melting beneath various thickness of debris cover on Khumbu Glacier, Nepal, using a positive degree-day factor, IAHS-AISH P264, 71-81.

[3] Benn, D.I., and Evans, D.J.A., 2010. Glaciers and Glaciation. Hodder-Arnold, London.

[4] Paterson, W.S.B., 1994. Physics of glaciers. Butterworth-Heinemann.

[5] Nakawo, M. and Young, G.J., 1981. Field experiments to determine the effect of a debris layer on ablation of glacier ice. Annals of Glaciology2, 85-91.

[6] Nicholson, L. and Benn, D.I., 2006. Calculating ice melt beneath a debris layer using meteorological data. Journal of Glaciology52, 463-470.

[7] Benn, D.I., Kirkbride, M.P., Owen, L.A. and Brazier, V., 2003. Glaciated valley landsystems. In Evans, D.J.A. (ed.) Glacial Landsystems, pp. 372-406.

[8] Benn, D.I., Bolch, T., Hands, K., Gulley, J., Luckman, A., Nicholson, L.I., Quincey, D., Thompson, S., Toumi, R. and Wiseman, S., 2012. Response of debris-covered glaciers in the Mount Everest region to recent warming, and implications for outburst flood hazards. Earth-Science Reviews114, 156-174.

What is the global volume of land ice and how is it changing?

How much land ice is there in the World?

Most (99.5%) of the permanent ice volume in the world is locked up in ice sheets and glaciers. The Antarctic Ice Sheet is the largest store of frozen freshwater; it would raise sea levels by 57.9 m (its “sea level equivalent”, or SLE) on full melting (BedMachine). The Antarctic Ice Sheet covers 8.3% of the Earth’s land surface.

The Greenland Ice Sheet has a sea level equivalent ice volume of 7.42 m, and covers 1.2% of the global land surface (BedMachine).

Finally, glaciers and ice caps have a sea level equivalent ice volume of 0.32 m, covering just 0.5% of the global land surface (Figure 1). There is a nice illustration of this here.

Global glaciers (in yellow) and ice sheets (white). From IPCC AR5

Figure 1. Global land ice. Glaciers are highlighted in yellow, ice shelves in green, ice sheets in white.

Other sources of global ice

There are also small amounts of ice stored in the ground in permafrost regions, frozen lakes and rivers, seasonal snow cover, and so on.

Sea ice (frozen sea water) and ice shelves (frozen floating extensions of land ice; green on Figure 1 above) do not have a “sea level equivalent” of ice volume as they are already floating, so would not raise sea levels on full melting.

Measuring changes in global ice volume

Changes in global ice volume are often expressed in gigatonnes per year (yr-1). A gigatonne is 1,000,000,000 tonnes. 1 kmwater = 1 Gt water; 361.8 Gt of ice will raise global sea levels by 1 mm.

Greenland Ice Sheet

Mass balance of the Greenland Ice Sheet

The Greenland Ice Sheet has been losing mass for over 20 years. The most recent estimates suggest that the Greenland Ice Sheet from 2012 to 2016 had a negative mass balance, losing 247 ± 15 Gigatonnes (Gt) per year of ice volume, contributing 0.69 ± 0.04 mm per year to sea level rise[2]. The mass balance of Greenland has been increasingly negative since 1995, and it is now equivalent to the global contribution to sea level rise from glaciers and ice caps (Figure 2).

Figure 2. Cumulative ice mass loss from Greenland ice sheet 1992–2012[1] (from IPCC AR5).

Driven by changes in surface mass balance

These changes have largely been driven by changes in surface mass balance. While in Greenland 60% of mass loss is through ice discharge across the grounding line to the ocean (as icebergs or melting in the ocean), 40% of mass loss is from surface melt. Increases in surface melt (ablation) are largely responsible for the increasing melting of Greenland [3].

On June 15, 2016, the Advanced Land Imager (ALI) on NASA’s Earth Observing-1 satellite acquired a natural-color image of an area just inland from the coast of southwestern Greenland (120 kilometers southeast of Ilulisat and 500 kilometers north-northeast of Nuuk). From Wikimedia Commons

Figure 3. Surface meltwater on the Greenland Ice Sheet.

The estimates of Greenland Ice Sheet mass balance above include the peripheral glaciers surrounding the larger ice sheet. These peripheral glaciers account for around 15-20% of the total mass imbalance of the ice sheet[2, 4].

These increases in surface melt and mass losses from Greenland are due to recent increases in winter and summer air temperatures, with increases in the size of the ice sheet ablation area (the area with net melting over one year). This is associated with changes in the surface albedo, as ice has a lower albedo than white snow, exacerbating melt. Overall, this is leading to a lowering of the Greenland Ice Sheet surface elevation (Figure 4), and a decrease in ice volume.

Acceleration in outlet glaciers

Ice discharge from the major outlet glaciers of the Greenland Ice Sheet has also increased, with glaciers accelerating in western Greenland (e.g. Jakobshavn Isbrae, JI) (Figure 4). This faster ice flow leads to these outlet glaciers discharging more ice volume to the ocean as icebergs than is replaced by snow, so the outlet glaciers are also thinning, as can be seen by the red on the figure below.

Figure 4. Average rates of surface elevation change (dh/dt) through time (2010-2017) for the Greenland and Antarctic Ice Sheets[2].

Antarctic Ice Sheet

Antarctic Ice Sheet ice volume

The best estimates of Antarctic volume come from BEDMAP2 [5]. BEDMAP2 provides us with a detailed map of the base of the ice sheet, derived mostly from radar data. There are three ice sheets in Antarctica, each with their own unique characteristics. They are the larger East Antarctic Ice Sheet (EAIS), with an SLE of 53.3 m, the West Antarctic Ice Sheet (WAIS), with an SLE of 4.3 m, and the Antarctic Peninsula Ice Sheet (APIS) with an SLE of 0.2 m.

Surface elevation of the Greenland and Antarctic ice sheets (IPCC AR5)

Figure 5. BEDMAP2 (Fretwell et al., 2013; IPCC AR5).

Antarctica surface mass balance

It is very cold in Antarctica, with very limited surface melt [6]. There is abundant accumulation in the coastal parts of Antarctica, especially western West Antarctica and on the APIS.  The figure below shows where surface mass balance is highest; reds and yellows indicate far more snowfall than is lost through surface melting. It is cold and dry in the centre of the East Antarctic Ice Sheet, with very little snowfall or surface melt.

The average ice-sheet integrated surface mass balance of Antarctica is +2418 ± 181 Gt yr-1 [6].

Figure 6. Mean (1979–2010) surface mass balance [mm w.e. y−1]. [6]

Changes in Antarctic mass balance

Most mass loss in Antarctica is driven through ocean melting and iceberg calving[7, 8]. This ice discharge to the ocean through the grounding line is increasing as outlet ice streams are accelerating and grounding lines are retreating (see here). Thus increased ice flow in Antarctica accounts for almost all recent increases in mass losses.

The sea level rise contribution from Antarctica was 0.49 – 0.73 mm yr-1 from 2012-2017, mostly from the APIS and WAIS and due to acceleration of outlet glaciers in Amundsen Sea Embayment (e.g. Pine Island Glacier/Thwaites Glacier) (Figure 4; 7)[2].

Ice streams of Antarctica with Pine Island Glacier and Thwaites glacier highlighted.

Figure 7. Location of Pine Island and Thwaites Glacier in Antarctica, with ice velocity from Rignot et al. 2011

Including ice gained and lost through all mechanisms, the current mass balance of Antarctica from 1992 to 2017 was:

  • EAIS: +5 ± 46 Gt yr-1
  • WAIS: –94 ± 27 Gt yr-1
  • APIS: –20 ± 15 Gt yr-1
  • Total Antarctic Ice Sheet: -109 ± 56 Gt yr-1

Antarctic Ice Sheet mass balance changed from 2012 to 2017 to -219 ± 43 Gt yr-1 [8] . Mass losses from West Antarctica are driving most of the total mass losses from Antarctica, with the mass balance of East Antarctica showing negligible changes [8].

Shepherd et al. 2018

Figure 8. Mass changes in Antarctica (Shepherd et al. 2018).

Glaciers and Ice caps

Glacier extent

The amount of ice contained in global glaciers and ice caps is mapped by the Randolph Glacier Inventory[9, 10]. This inventory uses satellite imagery and a formalised methodology to organise researchers working on mapping glaciers and glacier change. The Randolph Glacier inventory estimates that there are 198,000 glaciers worldwide (Figure 9); however, this is an arbitrary number as it depends on:

  • Subdivision of glaciers and mapping of ice divides
  • Accuracy of the digital elevation model used
  • Minimum area threshold; it is hard to map glaciers smaller than 0.2 km2 and so this is usually set as a minimum area threshold. There could be up to 400,000 glaciers if small glacierettes are included (but they only account for 1.4% of glacierised area).

Bamber et al. 2018

Figure 9. Global glaciers (yellow) and their area (pie charts) [2, 10].

The RGI estimates a total glacierised area of: 726,000 km2

  • Subantarctic and Antarctic: 132,900 km2
  • Arctic Canada North: 104,900 km2
  • Asia: 62,606 km2
  • Low latitudes: 2346km2
  • 44 % is in Arctic regions, 18% in Antarctic & Subantarctic.

Global glacier ice volume

An estimate of global ice volume in glaciers and ice caps remains a “grand challenge” in glaciology; there are few glaciers with direct measurement by radar [11]. Bed topography and thus ice thickness is usually then estimated, either by volume-area scaling [12, 13], inversions of ice surface slope and velocity [14, 15], or from numerical modelling of ice flow [16].

Our best current estimate of global glacier ice volume is[16]:

  • 170 x 103 ± 21 x 103 km3 (moutain glaciers & ice caps outside Greenland & Antarctica)
  • = 0.43 ± 0.06 m SLE.

Glacier recession

Glaciers worldwide are receding. The key methods for mapping glacier change include:

  • Satellite images (1970s-present)[17]
  • Topographic maps (~1900 to present)
  • Geomorphological evidence of glacier extent (LIA/sig. advances)
  • Automated and manual mapping from satellite imagery
  • Limit realistically of mapping glaciers min. 0.2 km2

Mass loss can also be quantified from analysis of glacier surface elevation change (dh/dt)[18, 19] using digital elevation model differencing, satellite gravimetry or altimetry, and in-situ surface mass balance measurements [20].

The figure below shows the current best estimates of ice volumes lost from Antarctica and Greenland from 2012-2016 (taken from Bamber et al. 2018) and from glaciers around the world. Bamber et al. 2018 do not provide an individual assessment of ice volume lost from each area, so here I have plotted ice volumes lost from 2003-2009 from Gardner et al. 2013. Each region corresponds to those mapped out in Figure 9 and glacier outlines are from GLIMS and the Randalph Glacier Inventory.

Note that peripheral glaciers around Greenland and Antarctica are included in the assessment for the ice sheets (cf. Bamber et al. 2018). These glaciers are however changing rapidly, and indeed account for a large portion of the overall change.

World glaciers and ice sheets mass balance

Figure 10. Global glacier mass budgets from 2012-2016 by Bamber et al. 2018 (ice sheets) and 2003-2009 (glaciers; Garder et al. 2013).

These data, recently compiled by Bamber et al. 2018, give a global estimate of mass loss from glaciers of -227 ± 31 Gt yr-1 (2012-2016). This does not include losses from peripheral glaciers around Greenland and Antarctica, which are included in the ice sheet mass balance assessments.


Figure 11. Global glacier melt (IPCC AR5)[1]

This has led the World Glacier Monitoring Service (WGMS) to state: “rates of early 21st-century mass loss are without precedent on a global scale, at least for the time period observed and probably also for recorded history” [21].

This global melt is a challenge for society. While the sea level rise from glaciers is ultimately constrained by their small ice volume globally, they remain important as sources of freshwater [22]; their melting poses new hazards to mountain communities[23-25], and they remain important for local economies [26].


Global changes in land ice volume were recently summarised by Bamber et al. (2018):

Ice mass Total ice volume % Global land surface Volume change 2012-2016 Sea level contribution 2012-2016
EAIS 53.3 m SLE 8.3% -19 ± 20 Gt yr-1 0.05 ± 0.06 mm yr-1
WAIS & APIS 4.5 m SLE -172 ± 27 Gt yr-1 0.48 ± 0.08 mm yr-1
Greenland 7.36 m SLE 1.2% -247 ± 15 Gt yr-1 0.69 ± 0.04 mm yr-1
Global glaciers and ice caps* 0.43 m SLE

(113,915 to 191,879 Gt)

0.5% -227 ± 31 Gt yr-1 0.63 ± 0.08 mm yr-1
Total 12.5% -665 ± 48 Gt yr-1 1.85 ± 0.13 mm yr-1

*excl. glaciers peripheral to ice sheets

Accelerating mass loss from land ice

Mass loss is accelerating (Figure 12), with changes in ocean melt driving recession in Antarctica, increased ice discharge and surface melt driving changes in Greenland, and negative surface mass balances largely driving glacier recession worldwide. Losses from Greenland are now the most significant contributor to global sea level rise (this includes the peripheral glaciers around the ice sheet), recently overtaking glaciers as the largest contributor.

Bamber et al. 2018

Figure 12. Mass losses from glaciers and ice sheets, annually (Bamber et al. 2018)

Below is a nice summary of the key changes and processes from the IPCC AR4:

Figure 13. Summary of global changes in land ice, IPCC AR5 (2013).

Further reading


  1. Vaughan, D.G., et al., Observations: Cryosphere, in Climate Change 2013: The Physical Science Basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change, T.F. Stocker, et al., Editors. 2013, Cambridge University Press: Cambridge, UK. p. 317-382.
  2. Bamber, J.L., et al., 2018, Environmental Research Letters.
  3. van den Broeke, M., et al., 2017, Current Climate Change Reports. 3, 345-356.
  4. Bolch T, et al., 2013, Geophys. Res. Lett. . 40, 875-881.
  5. Fretwell, L.O., et al., 2013, The Cryosphere. 7, 375-393.
  6. Lenaerts, J.T.M., et al., 2012, Geophysical Research Letters. 39, L04501.
  7. Pattyn, F., et al., 2018, Nature Climate Change.
  8. Shepherd, A., et al., 2018, Nature. 558, 219-222.
  9. Arendt, A., et al., Randolph Glacier Inventory [v2.0]: A Dataset of Global Glacier Outlines. 2012, Global Land Ice Measurements from Space: Boulder Colorado, USA.
  10. Pfeffer, W.T., et al., 2014, Journal of Glaciology. 60, 537.
  11. Gärtner-Roer, I., et al., 2014, Global and Planetary Change. 122, 330-344.
  12. Bahr, D.B., Estimation of glacier volume and volume change by scaling methods, in Encyclopedia of Snow, Ice and Glaciers. 2014, Springer. p. 278-280.
  13. Bahr, D.B., W.T. Pfeffer, and G. Kaser, 2014, Reviews of Geophysics.
  14. Carrivick, J.L., et al., 2018, Geografiska Annaler: Series A, Physical Geography, 1-23.
  15. Carrivick, J.L., et al., 2016, Global and Planetary Change. 146, 122-132.
  16. Huss, M. and D. Farinotti, 2012, Journal of Geophysical Research: Earth Surface. 117, F04010.
  17. Davies, B.J. and N.F. Glasser, 2012, Journal of Glaciology. 58, 1063-1084.
  18. Willis, M.J., et al., 2011, Remote Sensing of Environment. 117, 184-198.
  19. Willis, M.J., et al., 2012, Geophys. Res. Lett. 39, L17501.
  20. Gardner, A.S., et al., 2013, Science. 340, 852-857.
  21. Zemp, M., et al., 2015, Journal of Glaciology. 61, 745-762.
  22. Immerzeel WW, van Beek L P H, and B.M.F. P, 2010, Science. 328, 1382–85.
  23. Emmer, A., 2017, Quaternary Science Reviews. 177, 220-234.
  24. Emmer, A., Glacier Retreat and Glacial Lake Outburst Floods (GLOFs), in Oxford research Encyclopedias–Natural Hazard Science. 2017, Oxford University Press. p. 1-38.
  25. Harrison, S., et al., 2017.
  26. al, H.M.e., 2017 Earth’s Future 5 418-35.

Glacier accumulation and ablation

Glacier accumulation | Glacier ablation | Equilibrium line altitude | Glaciers as a system | Further reading | References | Comments |

Glacier accumulation

A glacier is a pile of snow and ice. In cold regions (either towards the poles or at high altitudes), more snow falls (accumulates) than melts (ablates) in the summer season. If the snowpack starts to remain over the summer months, it will gradually build up into a glacier over a period of years.

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

The key input to a glacier is precipitation. This can be “solid precipitation” (snow, hail, freezing rain) and rain1. Further sources of accumulation can include wind-blown snow, avalanching and hoar frost. These inputs together make up the surface accumulation on a glacier.

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. After Cogley et al., 2011.

In general, glaciers receive more mass in their upper reaches and lose more mass in their lower reaches. The part of the glacier that receives more mass by accumulation than it loses by ablation is the accumulation zone.

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

Formation of glacial ice

Over time, the snowfall (by far the most important input to a glacier) is gradually compressed and compacted by the weight of further snowfall on top it. The beautiful pointy edges of the snowflake gradually lose their tips and shape, becoming first granular ice, then firn, and finally glacial ice.

Layers of ice on Davies Dome Glacier, James Ross Island, Antarctic Peninsula.

The processes of transformation from snow to ice include partial melting, refreezing and fusing. The rate of transformation varies according to climate (temperature and precipitation regimes). The image below is from an ice core. Note the summer and winter layers in the ice. You can also no longer see the individual crystals that make up the glacier ice at this depth.

This 19 cm long of GISP2 ice core from 1855 m depth shows annual layers in the ice. This section contains 11 annual layers with summer layers (arrowed) sandwiched between darker winter layers. From the US National Oceanic and Atmospheric Administration, Wikimedia Commons.

Glacier ice is a crystalline material, and the crystal size and depth varies with the history of the ice.

Glacier ablation

As ice flows downhill, it either reaches warmer climates, or it reaches the ocean.  This causes various processes of melt, or ablation, to occur. In a land-terminating glacier (a glacier that ends on dry land), the main processes of ablation will be surface melt, because air temperatures generally increase as you lose altitude. This meltwater runs off the glacier and forms a number of rivers that typically drain the glacier.

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

This surface meltwater may runoff as surface runoff (as shown above; this is a supraglacial meltwater stream on the surface of the glacier), or it may make its way to the bed of the glacier through cracks in the ice (see the figure below). The water at the glacier bed eventually makes it way to the margin of the glacier, where it exits as a meltwater stream.

Meltwater propagates to the glacier bed through crevasses and moulins

Glaciers that reach the sea or terminate in a lake (Marine-terminating and lacustrine-terminating respectively) additionally will calve icebergs and melt underwater.   In large parts of Antarctica, melting underneath the base of floating ice shelves and calving from the margin of the glaciers dominate over surface melt.

Upsala Glacier, from the Southern Patagonian Ice Field, terminates in a large lake. Note the calved icebergs drifting out across the lake. Credit: NASA

The lower part of the glacier generally loses more mass from ablation than it receives from accumulation. This part of the glacier is the ablation zone.


Small tidewater (marine-terminating) glaciers calving into Croft Bay, Antarctic Peninsula

Equilibrium line altitude

Most glaciers receive more inputs and accumulation in their upper reaches, and lose more mass by ablation in their lower reaches. The Equilibrium Line Altitude (ELA) marks the area of the glacier separating the accumulation zone from the ablation zone, and were annual accumulation and ablation are equal2.

Equilibrium line altitudes in a hypothetical glacier

Glaciers as a system

Glacier ice is actually a viscous fluid, which flows and deforms under its own weight. Glaciers can therefore be thought of as systems, which receive snow and ice, flow downslope, and melt. Snow and ice are stored in the glacier until they melt as the glacier reaches lower elevations. This concept is explored in more detail in the Introduction to Glacier Mass Balance page and the pages on Glacier Flow.

In the European Alps and North America, most glaciers receive snowfall throughout the winter, and the main glacier ablation occurs in the summer. The Mass Balance, the balance of accumulation and ablation, is usually therefore positive in the winter and negative in the summer3. These glaciers, which receive more snow in winter and less in summer, are known as Winter Accumulation Type Glaciers. These glaciers form the majority of the world’s glaciers4.

In contrast, in places like the Himalaya, the monsoon brings more precipitation in the summer and less in the relatively cold, dry winter. These glaciers therefore receive more accumulation in the summer, and are known as Summer Accumulation Type Glaciers.

Further reading


1              Cogley, J. G. et al. Glossary of Glacier Mass Balance and related terms.  (IHP-VII Technical Documents in Hydrology No. 86, IACS Contribution No. 2, UNESCO-IHP, 2011).

2              Bakke, J. & Nesje, A. in Encyclopedia of Snow, Ice and Glaciers   (eds Vijay P. Singh, Pratap Singh, & Umesh K. Haritashya)  268-277 (Springer Netherlands, 2011).

3              Naito, N. in Encyclopedia of Snow, Ice and Glaciers   (eds Vijay P. Singh, Pratap Singh, & Umesh K. Haritashya)  1107-1108 (Springer Netherlands, 2011).

4              Kumar, A. in Encyclopedia of Snow, Ice and Glaciers   (eds Vijay P. Singh, Pratap Singh, & Umesh K. Haritashya)  1227-1227 (Springer Netherlands, 2011).


Mass balance of the Antarctic ice sheet from 1992 to 2017

A new paper with a whole host of authors has just been published in Nature (IMBIE Team, 2018). It provides a new estimate of mass balance of the entire Antarctic Ice Sheet over the last 25 years, the longest and most thorough estimate of this to date.

This article argues that the Antarctic Peninsula, the smallest ice sheet in Antarctica, has lost an average of 20 Gigatonnes (Gt) of ice per year over the 25 year study. This increased during the study and especially since the year 2000.  The West Antarctic Ice Sheet lost 53±29 Gt yr-1 from 1992-1997, but this accelerated to 159±26 Gt yr-1 from 2012-2017. The East Antarctic Ice Sheet is more stable, with small gains (with large errors) over the study period. Continue reading