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.

Latitude

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.

Topography

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.

References

[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

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

Insulation

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.

References

[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 58.3 m (its “sea level equivalent”, or SLE) on full melting. 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 7.36 m, and covers 1.2% of the global land surface.

Finally, glaciers and ice caps have a sea level equivalent ice volume of 0.43 m, covering just 0.5% of the global land surface[1] (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.

IPCC AR5

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].

Summary

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

References

  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.

croft-bay2

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

References

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

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.

The total mass balance of Antarctica was recently updated here.

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

References


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:

References


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).

Glacier response time

What is glacier response time? | The role of glacier size | The role of mass balance gradients | The role of glacier slope and hypsometry | Calculating glacier response times | Summary | Further reading | References | Comments |

What is glacier response time?

Glacier response time is the length of time taken for a glacier to adjust its geometry to a new steady state after a change in glacier mass balance1, caused by a changing climate2. Mountain glaciers worldwide are currently thinning and receding, but their behaviour in the future is highly variable, as their recession is controlled by many parameters. Understanding glacier response time is important for understanding how quickly particular glaciers will change in length and volume under a given climatic scenario; if we are to be able to estimate future sea level rise from shrinking glaciers, we need to know how quickly they can change2. Will glaciers shrink in response to short term climate fluctuations, or are longer and more sustained climatic changes required before a glacier changes its geometry?

Mountain glacier mass balance since 1970, excluding the Greenland and Antarctic ice sheets. From the Global Warming Art Project.

Mountain glacier mass balance since 1970, excluding the Greenland and Antarctic ice sheets. Mountain glaciers worldwide currently show significant thinning and recession. From the Global Warming Art Project.

The response time of a glacier is largely a function of its mean thickness and terminus ablation3 (Table 1), and of its hypsometry and mass balance gradient4,5 (the change in accumulation and ablation with elevation; a glacier with a steeper mass balance gradient receives more accumulation and has more ablation than a dry glacier with little accumulation and ablation). High mass balance gradients are indicative of high flux through the ELA. High gradients are usually found in mid-latitude maritime regions, where the maritime environment represents a major heat and moisture source6. Glacier size also affects glacier response time7.

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

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

Table 1. Estimated glacier response times as a function of thickness and terminus ablation. From Cuffey and Patterson, 2010.

  Thickness (m) Terminus ablation (m per year) Response time (years)
Glaciers in temperate maritime climate 150-300 5-10 15-60
Glacier in high-polar climate 150-300 0.5-1 150-600
Ice caps in Arctic Canada 500-1000 1-2 250-1000
Greenland Ice Sheet 3000 1-2 1500-3000

The role of glacier size

Small glaciers contribute significantly to present observed sea level rise (Gardner et al., 2013)

Small glaciers contribute significantly to present observed sea level rise (Gardner et al., 2013)

Smaller glaciers have shorter response times, and this is one of the reasons why small glaciers currently contribute significantly to sea level rise10,11. The high sensitivity of small ice masses to climate change is a function of their small system scale, and their proximity (compared with larger ice sheets) to the melting point9. However, it is difficult to apply a uniform response time widely across multiple glaciers. There is no straightforward relationship between glacier size and change in  ice volume under any given climate scenario12.

Short, steep glaciers in maritime environments (with a correspondingly steep mass balance gradient) reach equilibrium following a change in mass balance forcing after only a few years, larger valley glaciers in around a century, and continental ice caps with gentle slopes much longer4. The smallest cirque glaciers will reflect annual changes in mass balance, almost without delay8. Franz Josef Glacier, in New Zealand, is 11 km long and 35 km2 in area, has a steep surface gradient and a steep mass balance gradient, and has a response time of 21-24 years9. The Greenland Ice sheet, on the other hand, is estimated to have a response time of 1500-3000 years5.

The implication of this is that, given the same climatic forcing, glaciers of different lengths and thicknesses will respond in different ways, with variable numbers of advances and retreats. These should generally correlate regionally. Moraines occurring at only a few glaciers at a specific time are interpreted as reflecting shorter term climatic oscillations and a glacier with a short response time4. Larger, flatter glaciers tend to smooth the climatic signal, with a delay of several decades.

The role of mass balance gradients

Mass balance gradients of hypothetical glaciers

Mass balance gradients of hypothetical glaciers

A key factor in controlling glacier response to climatic perturbations is the mass balance gradient, the change in net balance with altitude, which is largely governed by the temperature lapse rate2,6. There is a linear relationship between mass balance gradient and mass balance sensitivity. Glaciers with a steeper mass balance gradient flow faster. The mass balance gradient of a glacier is affected by glacier size, hypsometry (the variation of glacier area with altitude13) and glacier slope.

Wet, maritime glaciers have a steeper mass balance gradient than dry continental glaciers. Warm, wet glaciers with a steep mass balance gradient and high sensitivity have a small response time, and cold and dry glaciers with low mass balance gradients and sensitivity have typically a longer response time. Temperature-sensitive glaciers may show the most rapid response to climate change at present, but may not be the most important contributors to global sea level rise over longer time periods2.

The role of glacier slope and hypsometry

Glacier hypsometry controls the mass balance elevation distribution over a glacier13. Glacier hypsometry is determined by valley shape, topographic relief and ice volume distribution. The altitudinal distribution of a glacier controls its sensitivity to a rise in the Equilibrium Line Altitude (ELA); glaciers with a large, relatively flat accumulation area will be more sensitive to a small increase in ELA than glaciers with a steeper accumulation area.

Response time of glaciers. From Haeberli, 1995

Response time of glaciers. From Haeberli, 1995

Steeper glaciers therefore typically have shorter response times4,5,8. In the New Zealand Southern Alps, smaller, steeper mountain glaciers have recently had minor readvances following short term climatic oscillations, while larger, low-gradient valley glaciers have continued to recede, as they have done for the last century14.  This is because, in general, if the glacier surface gradient is small, changes in mass balance occur slowly with distance15, whilst with steep glaciers, changes in mass balance occur rapidly with distance.

Calculating glacier response times

Mathematical estimates

In reality, response time actually refers to the time taken for a glacier to complete most of its adjustment to a change in mass balance5. This is because glaciers continue to adjust at an ever decreasing rate for a very long time after the change. Response time therefore typically refers to the time taken for a glacier to complete all but a factor 1/e (or 37%) of its net change1,5,9. Glacier response time can be estimated for a particular glacier from the simple equation,

 equation1

Where H is the thickness of the glacier and bt is the scale of ablation at the terminus of the glacier15,16. This calculates response time taken for the volume of the glacier to reach steady state following a change in mass balance. It will only give order-of-magnitude estimates2. This equation predicts that response time will increase linearly with glacier thickness, and that larger glaciers will have longer response times15. However, thicker glaciers are also longer, and longer glaciers push their snouts further into the ablation zone. The mass balance therefore gets increasingly negative at the snout as the glacier gets longer.

Numerical models

Glacier response times are usually calculated using glacier models9,17,18. Numerical flowline models take into account glacier geometry and mass balance when calculating glacier response time and climate sensitivity. Oerlemans (1997) defines glacier response time as the time taken for the glacier volume to go from one steady state (V1) to another (V2) under a given climatic forcing (C1 to C2)18. Response time for glacier volume can therefore be written as:

equation2

Oerlemans (1997) used a step change of 0.5K and -0.5K to force a change in Franz Josef Glacier18, and observed how long it took to reach equilibrium for both length and volume. More recently, Anderson et al. (2008) used a numerical model to impose a step-change in mass balance on Franz Josef Glacier, and observed how long it took for the glacier to complete two-thirds of its response9. This approach has also been used to compare the response time of glaciers of different sizes using different models1.

Summary

Glacier response time is an important factor to take into account when analysing glacier response to climate change. Not all glaciers will respond in a uniform way to a change in environmental conditions, as their response times are governed by ice thickness, ablation at the terminus, mass balance gradients, hypsometry and ice surface slope.

Further reading

References


1.            Leysinger Vieli, G.J.M.C. & Gudmundsson, G.H. On estimating length fluctuations of glaciers caused by changes in climatic forcing. Journal of Geophysical Research: Earth Surface 109, F01007 (2004).

2.            Raper, S.C.B. & Braithwaite, R.J. Glacier volume response time and its links to climate and topography based on a conceptual model of glacier hypsometry. The Cryosphere 3, 183-194 (2009).

3.            Jóhannesson, T., Raymond, C. & Waddington, E. Time-scale for adjustment of glaciers to changes in mass balance. Journal of Glaciology 35, 355-369 (1989).

4.            Kirkbride, M.P. & Winkler, S. Correlation of Late Quaternary moraines: impact of climate variability, glacier response, and chronological resolution. Quaternary Science Reviews 46, 1-29 (2012).

5.            Cuffey, K.M. & Paterson, W.S.B. The Physics of Glaciers, 4th edition, 704 (Academic Press, 2010).

6.            Rea, B.R., Evans, D.J.A., 2007. Quantifying climate and glacier mass balance in north Norway during the Younger Dryas. Palaeogeography, Palaeoclimatology, Palaeoecology 246, 307-330.

7.            Harrison, W.D., Elsberg, D.H., Echelmeyer, K.A. & Krimmel, R.M. On the characterization of glacier response by a single time-scale. Journal of Glaciology 47, 659-664 (2001).

8.            Haeberli, W. Glacier fluctuations and climate change detection. Geogr. Fis. Dinam. Quat 18, 191-199 (1995).

9.            Anderson, B., Lawson, W. & Owens, I. Response of Franz Josef Glacier Ka Roimata o Hine Hukatere to climate change. Global and Planetary Change 63, 23-30 (2008).

10.            Gardner, A.S., Moholdt, G., Cogley, J.G., Wouters, B., Arendt, A.A., Wahr, J., Berthier, E., Hock, R., Pfeffer, W.T., Kaser, G., Ligtenberg, S.R.M., Bolch, T., Sharp, M.J., Hagen, J.O., van den Broeke, M.R. & Paul, F. A Reconciled Estimate of Glacier Contributions to Sea Level Rise: 2003 to 2009. Science 340, 852-857 (2013).

11.          Hock, R., de Woul, M., Radic, V. & Dyurgerov, M. Mountain glaciers and ice caps around Antarctica make a large sea-level rise contribution. Geophysical Research Letters 36, L07501 (2009).

12.          Oerlemans, J., Anderson, B., Hubbard, A., Huybrechts, P., Jóhannesson, T., Knap, W.H., Schmeits, M., Stroeven, A.P., van de Wal, R.S.W., Wallinga, J. & Zuo, Z. Modelling the response of glaciers to climate warming. Climate Dynamics 14, 267-274 (1998).

13.          Jiskoot, H., Curran, C.J., Tessler, D.L. & Shenton, L.R. Changes in Clemenceau Icefield and Chaba Group glaciers, Canada, related to hypsometry, tributary detachment, length-slope and area-aspect relations. Annals of Glaciology 50, 133-143 (2009).

14.          Winkler, S., Chinn, T., Gärtner-Roer, I., Nussbaumer, S.U., Zemp, M. & Zumbühl, H.J. An introduction to mountain glaciers as climate indicators with spatial and temporal diversity. Erdkunde, 97-118 (2010).

15.          Bahr, D.B., Pfeffer, W.T., Sassolas, C. & Meier, M.F. Response time of glaciers as a function of size and mass balance: 1. Theory. Journal of Geophysical Research: Solid Earth 103, 9777-9782 (1998).

16.          Jóhannesson, T., Raymond, C.F. & Waddington, E.D. A Simple Method for Determining the Response Time of Glaciers. in Glacier Fluctuations and Climatic Change, Vol. 6 (ed. Oerlemans, J.) 343-352 (Springer Netherlands, 1989).

17.          Klok, E.J. & Oerlemans, J. Deriving historical equilibrium-line altitudes from a glacier length record by linear inverse modelling. The Holocene 13, 343-351 (2003).

18.          Oerlemans, J. Climate Sensitivity of Franz Josef Glacier, New Zealand, as Revealed by Numerical Modeling. Arctic and Alpine Research 29, 233-239 (1997).

Deformation and sliding

Glacier mass balance | Glacier flow | Internal deformation | Basal sliding | Subglacial deformation | Different types of glacier flow | References | Comments |

Glacier mass balance

Components of mass balance of a glacier. From the USGS

Components of mass balance of a glacier. From the USGS

How do glaciers move? A glacier is a pile of ice, and as such, deforms under the force of gravity. Glaciers flow downslope because they accumulate mass (ice) in their upper portions (from precipitation and from wind-blown snow) and ablate (melt, sublimate and calve ice bergs) in their lower portions.

This means that a glacier in a steady state (equilibrium) will not change in steepness or size, because accumulation = ablation. The altitude with zero net accumulation or ablation on the glacier is the equilibrium line altitude. Changes in rates of accumulation or ablation will lead to glacier advance or recession; if the accumulation area of a glacier shrinks, for example, and the equilibrium line altitude rises, then the glacier will recede.

Glacier mass balance is the difference between accumulation and ablation. It is therefore controlled by both temperature and precipitation. If accumulation is greater than ablation, then the glacier has positive mass balance and will advance. If ablation is greater than accumulation, then the glacier has negative mass balance and will recede.

Note: see Common Misconceptions in Glaciology. Glaciers always flow downslope under the weight of their own gravity. A receding or shrinking glacier still flows (although it might flow very slowly!); it’s just that it’s melting faster than it’s acquiring snow in its upper reaches. As a result, the glacier will thin and the snout position will recede backwards.

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

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

In theory, glaciers discharge ice from the accumulation area to the ablation area and maintain a steady-state profile. The velocity, the balance velocity, is controlled by glacier mass balance and glacier geometry (Jiskoot et al., 2011). Some glaciers have dynamic flow driven by other factors, for example, surging glaciers, tidewater glaciers, ice streams or ice-shelf tributary glaciers.

Glaciers flow through ice deformation and sliding

Glaciers always flow downslope, through the processes of deformation and sliding. Glacier flow, velocity and motion is controlled several factors (Jiskoot et al., 2011), including those listed below:

  • Ice geometry (thickness, steepness),
  • Ice properties (temperature, density),
  • Valley geometry,
  • Bedrock conditions (hard, soft, frozen or thawed bed),
  • Subglacial hydrology,
  • Terminal environment (land, sea, ice shelf, sea ice), and
  • Mass balance (rate of accumulation and ablation).

When glaciers flow downslope, gravitational driving stresses are resisted (resistive stress). The driving stress is controlled by gravitational acceleration, ice density and temperature, ice thickness and ice surface slope. Resistive stresses mainly operate at the glacier bed, and comprise basal drag or lateral drag against valley walls.

This driving stress means that glaciers move in one of three ways:

  1. Internal deformation (creep)
  2. Basal sliding
  3. Soft bed subglacial deformation.

All glaciers flow by creep, but only glaciers with water at their base (temperate or polythermal – see Glacial Processes) have basal sliding, and only glaciers that lie on soft deformable beds have soft sediment deformation.  If all three factors are present, you have the ingredients to contribute to fast ice flow (see Ice Streams).

Internal deformation

If the glacier flows just by internal deformation, then it is likely that rates of creep decrease with depth, with fastest ice movement at the surface and slowest (or no) ice movement at the base and at the valley sides, where resistive stresses are highest (Jiskoot et al., 2011). Ice deforms because it is plastic. If large stresses are applied it can crack in a brittle manner (forming crevasses or calving ice bergs).

The video below shows a huge calving event at Helheim Glacier, Greenland, in July 2010. It was made by the Swansea Glaciology Group.

Basal sliding

Glaciers can slide because ice melts under pressure, resulting in a film of water at the ice-bed interface. This can facilitate decoupling and enhance fast ice flow. If the glacier bed is rough, with many bumps and obstacles, this increases melting and ice flow. This process is known as regelation. If water pressures become high enough, cavities can form at the ice-bed interface, causing sliding with bed separation. This reduces basal friction and allows faster ice flow. Sliding velocity is controlled by basal shear stress and effective pressure, which is the difference between ice overburden pressure and water pressure (Jiskoot et al., 2011).

This video shows the glacier flowing above a cavity beneath the ice on Mont Blanc glacier.

Subglacial deformation

Subglacial till (see Glacial Processes) comprises unconsolidated, unsorted or poorly sorted sediments ranging from boulders to clay. In Norfolk, till sequences are over 20 metres thick. Fine sediments, such as clay and sand, are not cohesive and therefore deform readily when shear stress is applied to them if they have a high pore-water pressure (so, like basal sliding, subglacial deformation depends on high basal water pressures). If basal shear stress (the gravitational driving dress) is greater than the yield strength of the till, deformation occurs, resulting in some fantastic glaciotectonic sequences (see picture gallery below and papers by Davies et al., 2009,  2012a, 2012b).

Different types of glacier flow

Glaciers do not just flow in a steady state, however. We have cold-based glaciers, which have little flow velocity; polythermal glaciers, which are partly frozen to their bed; wet-based glaciers which have sheet flow (as is described in the above sections); ice streams, which have very rapid flow velocities; and surging glaciers, which have periods of rapid flow separated by quiescent periods of slow flow. The key differences in temperate glacier flow is summarised in the table below. Each of these different flow regimes results in a set of different and diagnostic glacial landforms.

Further Reading

Other sources:

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