Moraine formation

Ridges, mounds and hummocks formed at the margin of glaciers are generally termed moraines. The study of moraines is particularly useful as it can shed light on the physical processes occurring at both active and former ice margins1,2 and because moraines are markers of former glacier extent, so can be used to track glacier change (e.g. size) over time3.

Moraine ridge forming at the terminus of Easton Glacier, Washington, USA. Photo: W. Siegmund

How do moraines form?

Moraines form through several main processes, which may vary from glacier to glacier, on a temporal (e.g. seasonal basis), and with changes in climate. The key moraine-forming processes are shown in the diagram below and explained through this page.

Summary of the three main moraine-forming processes. Push moraines (top) form during periods of ice-front stillstand or advance that bulldoze proglacial sediments. Dump moraines (middle) consist of rock and sediment that fall, flow and slump from the ice margin by gravity. Ablation moraines (bottom) form due to the varying rates of ice melt across the snout. Where debris cover is sparse (i.e. where the ice is ‘clean’) melting is relatively rapid and the glacier surface lowers quickly. Where surface debris cover is thick, the ice is insulated from melting and ice-cored moraines can exist. Created by J. Bendle.

Push moraines

Push moraines form at the snout of active glaciers. Rock and sediment debris at the ice margin is moulded into ridges by the bulldozing of material (ice pushing) by an advancing glacier4,5.

Due to the nature of their formation, push moraines tend to take on the shape of the ice margin during the time at which they formed4,5 (see image below). They are often found at the margin of active temperate glaciers (such as those found in southern Norway and Iceland) that experience brief periods ice-front stability or advance despite a general pattern of recession4,5. In some cases, a series of annual push moraines may form, where low-relief ridges are formed during winter advances of the glacier snout, leaving behind a detailed record of glacier extent over time6-10.

Push moraine ridges formed at the retreating terminus of Skálafellsjökull, Iceland. The moraines display a ‘sawtooth’ planform that closely mimics the ice margin geometry. These push moraines have been shown to form annually, driven by local climate conditions (Chandler et al., 2016; ref. 10). Image from GoogleEarth.

Debris squeezing

As well as the bulldozing of debris, sediment may also be squeezed out from beneath the glacier margin, either as a glacier advances in winter, or in the ablation season when till becomes water-soaked and easily displaced by the weight of overlying ice4,11. This process also contributes to the formation and growth of push moraines.

Dump moraines

Dump moraines form where debris flows or falls from a glacier surface due to gravity and accumulates at the ice front or side as a ridge. They form where the ice front is stationary and there is a regular supply of debris to the snout, normally due to the melt-out of rock debris stored in the ice4.

Dump moraine size is related to the amount of debris accumulating at the snout and the length of time the glacier margin is stationary. The volume of debris on the glacier surface is high where (i) the debris content within the ice is high; and (ii) where ice velocity is high, as faster flowing ice can transfer more debris to the margin11.

Loch Lomond Stadial moraines

Debris dumped from the ice front may be bulldozed into push moraines by advance(s) of the glacier margin2,12. Moraines formed by a combination of both the dumping and pushing of debris include those constructed by certain Scottish cirque and valley glaciers during the Loch Lomond Stadial2,13 (see image below). These moraines are similar in their genesis and morphology to those created by Icelandic glaciers today, which suggests that Loch Lomond Stadial glaciers in Britain were likely temperate and active during deglaciation2,13.

Loch Lomond Stadial moraine ridges formed by a combination of the dumping and bulldozing of rock and sediment debris, Coire Ardair. Photo: B. Davies.

Ablation moraines

Ablation moraines form where rock and sediment debris accumulate on the glacier surface near the margin and subsequently undergo melt-out4,11. The accumulation of dark-coloured material on the glacier surface lowers the ice albedo (i.e. its reflectiveness) and increases the amount of solar radiation absorbed at the glacier surface, which causes ice melt to speed up. However, where the debris layer is more than a few centimetres thick it insulates the ice surface from heating, slowing the rate of ice melt. Where the debris cover is extensive across a large part of the snout, the ice margin may detach completely from the main body of the glacier and become stagnant (see image below).

The debris-covered and stagnant ice margin of Exploradores Glacier in central Patagonia, Chile. Photo: J. Bendle.

When a debris-covered snout melts over time material is gradually let down from the ice surface to produce an area of ‘hummocky moraine’. This melt-out process can produce a variety of moraine types, from a chaotic assortment of sediment mounds and hollows (see image below)1 to more regular transverse ridges (often termed controlled moraines) that reflect the former pattern of debris in a parent glacier14.

Example of chaotic mounds and hollows in southern Patagonia, South America, which are interpreted to have formed by ice stagnation (see Darvill et al., 2017; ref. 15).


[1] Kjær, K.H. and Krüger, J., 2001. The final phase of dead‐ice moraine development: processes and sediment architecture, Kötlujökull, Iceland. Sedimentology48, 935-952.

[2] Lukas, S., 2005. A test of the englacial thrusting hypothesis of ‘hummocky’ moraine formation: case studies from the northwest Highlands, Scotland. Boreas34, 287-307.

[3] Schomacker, A. 2011. Moraine (Eds.) Singh, V.P., Singh, P. and Haritashya, U.K. Encyclopedia of Snow, Ice and Glaciers. Springer.

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

[5] Boulton, G.S., 1986. Push‐moraines and glacier‐contact fans in marine and terrestrial environments. Sedimentology33, 677-698.

[6] Sharp, M., 1984. Annual moraine ridges at Skálafellsjökull, south-east Iceland. Journal of Glaciology30, 82-93.

[7] Bradwell, T., 2004. Annual moraines and summer temperatures at Lambatungnajökull, Iceland. Arctic, Antarctic, and Alpine Research36, 502-508.

[8] Beedle, M.J., Menounos, B., Luckman, B.H. and Wheate, R., 2009. Annual push moraines as climate proxy. Geophysical Research Letters36.

[9] Lukas, S., 2012. Processes of annual moraine formation at a temperate alpine valley glacier: insights into glacier dynamics and climatic controls. Boreas, 41, 463-480.

[10] Chandler, B.M., Evans, D.J. and Roberts, D.H., 2016. Characteristics of recessional moraines at a temperate glacier in SE Iceland: Insights into patterns, rates and drivers of glacier retreat. Quaternary Science Reviews135, 171-205.

[11] Bennett, M.M. and Glasser, N.F. 2009. Glacial Geology: Ice Sheets and Landforms. John Wiley & Sons.

[12] Boulton, G.S. and Eyles, N., 1979. Sedimentation by valley glaciers: a model and genetic classification. Moraines and varves33, pp. 11-23.

[13] Jones, R.S., Lowe, J.J., Palmer, A.P., Eaves, S.R. and Golledge, N.R., 2017. Dynamics and palaeoclimatic significance of a Loch Lomond Stadial glacier: Coire Ardair, Creag Meagaidh, Western Highlands, Scotland. Proceedings of the Geologists’ Association128, 54-66.

[14] Evans, D.J.A., 2009. Controlled moraines: origins, characteristics and palaeoglaciological implications. Quaternary Science Reviews28, 183-208.

[15] Darvill, C.M., Stokes, C.R., Bentley, M.J., Evans, D.J. and Lovell, H., 2017. Dynamics of former ice lobes of the southernmost Patagonian Ice Sheet based on a glacial landsystems approach. Journal of Quaternary Science32, 857-876

Moraine types

Moraines are distinct ridges or mounds of debris that are laid down directly by a glacier or pushed up by it1. The term moraine is used to describe a wide variety of landforms created by the dumping, pushing, and squeezing of loose rock material, as well as the melting of glacial ice.

Moraine ridges on the forefield of the Matanuska Glacier, Alaska. Photo: Frank K.

In terms of size and shape, moraines are extremely varied. They range from low-relief ridges of ~1 m high and ~1 m wide formed at the snout of actively retreating valley glaciers2, to vast ‘till plains’ left behind by former continental ice sheets3.

Low-relief moraine ridges on the forefield of the actively retreating Skaftafellsjökull Glacier in Iceland. The moraines mark former ice extent and mirror the shape of the glacier terminus at the time of formation. Photo: TommyBee

Moraines consist of loose sediment and rock debris deposited by glacier ice, known as till. They may also contain slope, fluvial, lake and marine sediments if such material is present at the glacier margin, where it may be incorporated into glacial ice during a glacier advance, or deformed by glacier movement4,5.

Moraine composed of loose rock and sediment forming at the lateral margin of the Boulder Glacier, Washington, USA. Photo: W. Siegmund.

Moraines are important features for understanding past environments. Terminal moraines, for example, mark the maximum extent of a glacier advance (see diagram below) and are used by glaciologists to reconstruct the former size of glaciers and ice sheets that have now shrunk or disappeared entirely6.

Summary of the main moraine types and their spatial patterns. The top diagram is a cross-section through a cirque glacier. The bottom diagram is drawn in plan view, looking down on the surface of a valley glacier made up of several tributaries. Image created by J. Bendle.


The most common moraine types are defined below:

A terminal moraine is a moraine ridge that marks the maximum limit of a glacier advance. They form at the glacier terminus and mirror the shape of the ice margin at the time of deposition. The largest terminal moraines are formed by major continental ice sheets and can be over 100 m in height and 10s of kilometres long7,8.

Terminal moraine marking the limit of the former Patagonian Ice Sheet at the Last Glacial Maximum (~25 to 18 thousand years ago). Photo: J. Bendle.

Recessional moraines are found behind a terminal moraine limit and form during short-lived phases of glacier advance or stillstand that interrupt a general pattern of glacier retreat. In some cases, recessional moraines form on a yearly basis (normally as a result of winter glacier advances) and are known as annual moraines9,10,11.

Recessional moraines (arrowed) marking the shrinkage of a South American valley glacier. The glacier (not shown) retreated towards the south-west, leaving behind a moraine-dammed glacial lake. Imagery from GoogleEarth, diagram created by J. Bendle.

Lateral moraines form along the glacier side and consist of debris that falls or slumps from the valley wall or flows directly from the glacier surface12 (see image below). Where the rate of debris supply is high, lateral moraines can reach heights of more than 100 metres12–15.

Lateral moraine of the Callequeo Glacier of the San Lorenzo Icefield in central Patagonia, South America. Photo: J. Martin.

The term latero-frontal moraine is used where debris builds up around the entire glacier tongue14. These moraine types are common in mountain settings such as the European Alps, the Southern Alps of New Zeland (see the Mueller Glacier moraines below) and the Himalayas, where the high supply of rock debris from unstable valley sides, rapidly build up at the glacier margins.

Latero-frontal moraine complex of the Mueller Glacier, South Island, New Zealand. The debris-covered and downwasting Mueller Glacier is flanked by lateral moraines of ~100 m in height, which continue down valley and merge into terminal moraines. Imagery from GoogleEarth, diagram created by J. Bendle.

Medial moraines are debris ridges at the glacier surface running parallel to the direction of ice flow4,5. They are the surface (or supraglacial) expression of debris contained within the ice. Medial moraines form where lateral moraines meet at the confluence of two valley glaciers, or where debris contained in the ice is exposed at the surface due to melting in the ablation zone16.

Medial moraines on the surface of an Alaskan valley glacier. In this example, surface debris is concentrated at the point where two glaciers merge. Imagery from GoogleEarth, diagram created by J. Bendle.

Ground moraine is a term used to describe the uneven blanket of till deposited in the low-relief areas between more prominent moraine ridges6. This type of moraine, which is also commonly referred to as a till plain, form at the glacier sole as due to the deformation and eventual deposition of the substratum.


1. Hambrey, M. J. 1994. Glacial Environments. UCL Press.

2. Krüger, J., Schomacker, A. and Benediktsson, Í.Ö., 2010. 6 Ice-Marginal Environments: Geomorphic and Structural Genesis of Marginal Moraines at Mýrdalsjökull. Developments in Quaternary Sciences13, 79-104.

3. Dyke, A.S. and Prest, V.K. 1987. Late Wisconsinan and Holocene history of the Laurentide Ice Sheet. Geographie Physique et Quaternaire XLI, 237–63.

4. Benn, D.I. and Evans, D.J.A., 2010. Glaciers and Glaciation. Hodder Education. 

5. Bennett, M.M. and Glasser, N.F. 2011. Glacial Geology: Ice Sheets and Landforms. John Wiley & Sons.

6. Schomacker, A. 2011. Moraine (Eds.) Singh, V.P., Singh, P. and Haritashya, U.K. Encyclopedia of Snow, Ice and Glaciers. Springer.

7. Dyke, A.S., Andrews, J.T., Clark, P.U., England, J.H., Miller, G.H., Shaw, J. and Veillette, J.J., 2002. The Laurentide and Innuitian ice sheets during the last glacial maximum. Quaternary Science Reviews21, 9-31.

8. Glasser, N.F., Jansson, K.N., Harrison, S. and Kleman, J., 2008. The glacial geomorphology and Pleistocene history of South America between 38°S and 56°S. Quaternary Science Reviews27, 365-390.

9. Sharp, M., 1984. Annual moraine ridges at Skálafellsjökull, south-east Iceland. Journal of Glaciology30, 82-93.

10. Bradwell, T., 2004. Annual moraines and summer temperatures at Lambatungnajökull, Iceland. Arctic, Antarctic, and Alpine Research36, 502-508.

11. Beedle, M.J., Menounos, B., Luckman, B.H. and Wheate, R., 2009. Annual push moraines as climate proxy. Geophysical Research Letters36.

12. Lukas, S., Graf, A., Coray, S. and Schlüchter, C., 2012. Genesis, stability and preservation potential of large lateral moraines of Alpine valley glaciers–towards a unifying theory based on Findelengletscher, Switzerland. Quaternary Science Reviews38, 27-48.

13. Benn, D.I. and Owen, L.A., 2002. Himalayan glacial sedimentary environments: a framework for reconstructing and dating the former extent of glaciers in high mountains. Quaternary International97, 3-25.

14. Benn, D.I., Kirkbride, M.P., Owen L.A. and Brazier, V. 2003. Glaciated Valley Landsystems (Ed.) Glacial Landsystems, Arnold, London.

15. Evans, D.J., Shulmeister, J. and Hyatt, O., 2010. Sedimentology of latero-frontal moraines and fans on the west coast of South Island, New Zealand. Quaternary Science Reviews29, 3790-3811.

16. Eyles, N. and Rogerson, R.J., 1978. A framework for the investigation of medial moraine formation: Austerdalsbreen, Norway, and Berendon Glacier, British Columbia, Canada. Journal of Glaciology20, 99-113.


Glacial cirques, known locally as corries or coires (Scotland) and cwms (Wales), are large-scale erosional features common to many mountainous regions1,2. Classic cirques take the form of armchair-shaped hollows (see image below), with a steep headwall (which often culminates in a sharp ridge, or arête) and a gently-sloping or overdeepened valley floor (see diagram below).

Classic glacial cirque basin. Cwm Clyd in the Glyderau mountains of Snowdonia. Image from GoogleEarth.
Cross-section of a classic glacial cirque with an overdeepened (and lake filled) valley floor and a steep headwall mantled with slope deposits, such as scree. Image created by J. Bendle based on Barr & Spagnolo (2015; ref. 2)

In actively glacierized terrain, cirques are important basins for the accumulation of snow. They may host small cirque glaciers (see image below) that are confined to their bedrock hollows, or act as the source area for larger valley glaciers.

Cirque glacier (Styggebrean) in Jotunheimen National Park, Norway. Photo: J. Bendle.

In other mountainous areas, such as the British uplands, the occurrence of ice-free cirques (see image below) serve as a reminder of past glacier activity by recording former sites of glacier build-up3,4,5.

Cwm Cau, a formerly glacierized cirque basin in Snowdonia, Wales. Photo: J.Bendle.

Types of cirques

Far from being the same in all mountain areas, a wide range of cirque types occur. The most common are1,6:

  • Simple cirques, which are distinct and independent features
  • Compound cirques, where the upper part of a cirque basin contains two similarly sized simple cirques
  • Cirque complexes, where the upper part of a cirque basins contains more than two similarly sized simple cirques
  • Staircase cirques, where one cirque occurs above another
  • Cirque troughs, where a cirque basin occurs at the upper end of a glacial trough
Different types of glacial cirques. The top three examples are drawn in plan view, whereas the bottom two are drawn in cross-section. Image created by J. Bendle based on Barr & Spagnolo (2015; ref. 2)

The formation and growth of cirques

Cirques form through the gradual expansion of mountainside hollows associated with earlier fluvial, volcanic, or mass movement (e.g. landsliding) activity7. When these hollows become filled with snow8 they start to enlarge by nivation (a group of processes that includes freeze-thaw activity, chemical weathering, and seasonal snow melt)9.

True cirque growth only occurs once the thickness of snow patches increases to a point at which glacier ice can form by compaction. Once formed, glaciers widen and deepen cirques by subglacial abrasion and quarrying of the hollow floor and lower headwall3 (see diagram below). Cirques can also grow by backwards headwall erosion (wear back) due to frost-action, free-thaw, and mass movement3,10.

Cirque glaciers erode their hollows by subglacial plucking and abrasion, which are most effective under a warm-based, sliding glacier. Meltwater that drains to the bed through the randkluft (the gap between the glacier and headwall), bergshrund (a large crevasse near, but not touching, the headwall) or other crevasses, promotes subglacial erosion. Periglacial erosion (e.g. freeze-thaw) occurs on the headwall and in the randkluft. Image created by J. Bendle.

Case study: glacial cirques of Snowdonia

The glacial cirques of Snowdonia formed over several glaciations, and have a long history of investigation, first being visited by Charles Darwin over 150 years ago11. The most recent period of glacier activity in Snowdonia was during the mountain glaciation of upland Britain in the Loch Lomond Stadial (between ~12 and 10 thousand years ago)5,12,13.

Loch Lomond Stadial (~12 to 10 thousand years ago) cirque glaciers in Snowdonia, North Wales. Image from Bendle & Glasser (2012; ref. 5)

Why are cirques important?

Because cirques are areas of snow accumulation, the direction in which they point (their aspect) can tell us something about the links between climate and glacier growth in the past2,14.

If looking from above (see image above), an interesting observation is that most cirques in Snowdonia face to the north or east14 and these also held most (as well as the largest) Loch Lomond Stadial glaciers5,12.

Controls on cirque aspect

This is due to two factors. Firstly, north-facing cirques receive less solar radiation than south-facing cirques (in the Northern Hemisphere), resulting in lower air temperatures and less ice-melt across the year15.

Secondly, where prevailing winds blow mainly from the west, the snow on high ground will be blown down into east-facing cirques, adding to glacier mass5,15.


Using GoogleMaps or GoogleEarth, enter “Snowdon” in the navigation search bar and explore the cirques of Snowdonia.

Try to identify different cirque types (e.g. ‘simple’, ‘compound’, ‘complex’), and compare their sizes, shapes, and aspects.


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

[2] Barr, I.D. and Spagnolo, M., 2015. Glacial cirques as palaeoenvironmental indicators: their potential and limitations. Earth-Science Reviews151, 48-78.

[3] Evans, I.S., 2006. Allometric development of glacial cirque form: geological, relief and regional effects on the cirques of Wales. Geomorphology80, 245-266.

[4] Ballantyne, C.K., 2007. Loch Lomond Stadial glaciers in North Harris, Outer Hebrides, North-West Scotland: glacier reconstruction and palaeoclimatic implications. Quaternary Science Reviews26, 3134-3149.

[5] Bendle, J.M. and Glasser, N.F., 2012. Palaeoclimatic reconstruction from Lateglacial (Younger Dryas Chronozone) cirque glaciers in Snowdonia, North Wales. Proceedings of the Geologists’ Association123, 130-145.

[6] Gordon, J.E., 1977. Morphometry of cirques in the Kintail-Affric-Cannich area of northwest Scotland. Geografiska Annaler: Series A, Physical Geography59, 177-194.

[7] Turnbull, J.M. and Davies, T.R., 2006. A mass movement origin for cirques. Earth Surface Processes and Landforms 31, 1129-1148.

[8] Sanders, J.W., Cuffey, K.M., MacGregor, K.R. and Collins, B.D., 2013. The sediment budget of an alpine cirque. Geological Society of America Bulletin125, 229-248.

[9] Thorn, C.E., 1976. Quantitative evaluation of nivation in the Colorado Front Range. Geological Society of America Bulletin87, 1169-1178.

[10] Sanders, J.W., Cuffey, K.M., Moore, J.R., MacGregor, K.R. and Kavanaugh, J.L., 2012. Periglacial weathering and headwall erosion in cirque glacier bergschrunds. Geology40, 779-782.

[11] Darwin, C.R., 1842. Notes on the effects produced by the ancient glaciers of Caernarvonshire, and on the boulders transported by floating ice Lond. Edinb. Dublin Philos. Mag. J. Sci. 21, 180-188.

[12] Gray, J.M., 1982. The last glaciers (Loch Lomond Advance) in Snowdonia, N. Wales. Geological Journal17, 111-133.

[13] Hughes, P.D., 2009. Loch Lomond Stadial (Younger Dryas) glaciers and climate in Wales. Geological Journal44, 375-391.

[14] Evans, I.S., 2006. Local aspect asymmetry of mountain glaciation: a global survey of consistency of favoured directions for glacier numbers and altitudes. Geomorphology73, 166-184.

[15] Evans, I.S., 1977. World-wide variations in the direction and concentration of cirque and glacier aspects. Geografiska Annaler: Series A, Physical Geography59, 151-175.

Glaciers as a water resource

Mountains as Water Towers of the World

In many mountainous parts of the world with a seasonal rainfall, glaciers are a reliable water resource in the dry season. Mountains could be called the “Water Towers of the World”1, providing water from glacier melt and orographic rainfall to lowland regions. 

Glacierised drainage basins cover 26% of the global land surface outside of Greenland and Antarctica, and are populated by almost one-third of the World’s population2. Upland areas (above 2000 m above sea level) in southeast Asia supply the five basins of the Indus, Ganges, Yellow, Brahmaputra and Yangtze rivers, providing water to 1.4 billion people (over 20 % of the global population).

The Himalayan river basins and the number of people living in each one.

(Source: Redrawing the map of the world’s international river basins)

High Mountain Asia river basins


Glacier meltwater and runoff

Glacier meltwater and runoff contribute to and module downstream water flow, affecting freshwater availability for irrigation, hydropower, and ecosystems3.

Glacier runoff is typically seasonal, with a minimum in the snow-accumulation season, and a maximum in the melt season. This meltwater can compensate for seasons or years with low streamflow or droughts in downstream regions4.

Mountain glacier and lake in Peru

Global glacier recession

Mountain glaciers around the World are currently shrinking5-7, and this is expected to continue throughout the next century. Globally, glaciers are shrinking by 227 ± 32 gigatonnes per year8, enough to raise global sea levels by 0.63 ± 0.08 mm per year.

The areas shrinking fastest are in north America (-50 gigatonnes per year), northern Arctic Canada (60 gigatonnes per year), the Himalaya region (26 gigatonnes per year), and South America (29 gigatonnes per year)8.

World glaciers and ice sheets mass balance. Glaciers are shown in black. Green circles show glacier area, red circles are how much ice is lost annually.

Glacier “Peak meltwater”

As glaciers shrink, meltwater is released from storage within the glacier. Annual meltwater therefore increases, until a maximum is reached3,9. This maximum has been called ‘Peak Meltwater’.

After Peak Meltwater, runoff decreases as smaller glacier volumes can no longer support rising meltwater volumes. As the glacier retreats and disappears, annual runoff from direct precipitation may return to something like the original value, as water is no longer stored as snow. However melt-season runoff may decline substantially, as the glacier no longer acts as a reservoir. Seasonality of water availability may therefore increase, leading to droughts in dry years or dry seasons.

Essentially, as the glaciers shrink, they provide less and less melt water from long-term storage, which impacts seasonal freshwater availability3.  

Peak Meltwater and glacier recession under a warming climate.

Adapted from Huss and Hock (2018) and Rowan et al. (2018).

The degree to which glacier runoff contributes to downstream meltwater varies according to the basin, with glacier contributions being as much as 25% of the annual water budget. In many of these basins, peak meltwater is expected to have passed (e.g., Ref. 10), or will be passed in the next 20-30 years (e.g. Ref. 11). Ultimately, some projections suggest that up to half of the world’s population could be living in water scarcity by 2100 AD12.

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

Global scale peak meltwater?

A recent study by Huss and Hock (2018, Nature Climate Change) computed glacier runoff changes for the Earth’s 56 large-scale (>5000 km2) glacierised drainage basins with at least 30 km2 of ice to 2100 AD, and analysed the effect of glacial recession on streamflow.

In half of the basins, peak meltwater has already been reached. In the remaining basins, the modelled annual glacier runoff continues to rise until the maximum is reached, and then runoff declines. Peak water tends to occur later in basins with larger glaciers and higher ice-cover fractions3.

The researchers used a glacier model and climate model outputs forced by three different emissions scenarios, with peak emissions occurring at 2020 AD (RCP 2.6), 2050 AD (RCP 4.5) and after 2100 AD (RCP 8.5)3.  RCP 2.6 is the closest scenario to the targets of the Paris 2015 climate agreement. Projected temperature increases between 1990-2010 and 2080-2100, range from 1.6 ± 1.1°C (RCP 2.6) to 5.4 ± 2.2°C.

Between 2010 and 2100 AD, glacier volume in the 56 investigated basins was projected to decrease by 43±14% (RCP 2.6), 58±13% (RCP 4.5) and 74±11% (RCP 8.5). For the mid-range RCP 4.5, glacier volume reductions in the individual basins ranged from 37 to 99%1.

Reaching Peak meltwater

Peak meltwater has already been reached in 45% of the basins (year 2017 AD), but annual runoff is expected to continue to rise beyond 2050 AD in 22% of the basins. Basins with larger glaciers and high glacier cover (e.g. Susitna, Jökulsá) tend to reach peak meltwater towards the end of the twenty-first century.

In basins dominated by small glaciers (e.g. western Canada, central Europe, South America), peak meltwater has already passed and meltwater will decline over coming decades.

In most basins fed by High Mountain Asia (Aral Sea, Indus, Tarim, Brahmaputra), annual glacier runoff is projected to rise until the middle of the century, followed by steadily declining glacier meltwater runoff thereafter3.

By the end of the twenty-first century, the seasonal glacier runoff maximum is reduced in 93% of the basins compared with the 1990-2010 average, and runoff is less concentrated during the melt season.

Colours show the modelled year of peak water computed from 11-year moving averages of annual glacier runoff from all the glaciers located in the 56 investigated drainage basins, aggregated in 0.5 × 0.5° grid cells. Peak water is also shown with grey scales for all the macroscale basins, classified in 30-year intervals. The results refer to runoff from the initially glacierized area, and are based on the multimodel mean of 14 GCMs and the RCP4.5 emission scenario. The numbers in brackets below the basin names refer to basin glacierization in per cent. The insets show the modelled annual glacier runoff normalized with the average runoff in 1990–2010 for three selected basins. Triangles depict peak water (± standard deviation), thin lines show results for individual GCMs and G denotes the percentage ice cover. From Huss and Hock, 2018

In 19 of the 56 basins, the glacier runoff change between 2000 and 2090 AD accounts for at least one melt-season month with a reduction in runoff of at least 10% (i.e. glacier runoff reduction exceeds 10% of the basin runoff). This is sufficient to cause water scarcity in these basins.

The most significantly affected basins are in High Mountain Asia (Aral Sea, Indus, Tarim, Balkhash), Peru (Santa), South America (Colorado, Baker, Santa Cruz), and North America (Fraser, Skeena, Taku, Nass)3.

The ratio of glacier runoff change to basin runoff is evaluated for the period July to October (January to April for the southern hemisphere, and throughout the year in the tropics). For basins with substantial glacier runoff decreases in at least one month, the ratio refers to the month (given in brackets below the basin names) with the largest reduction in glacier runoff. Basins with negligible glacier impact (|ΔQ′g/Qbasin|< 5%) are shown in grey, and the remaining basins, which show glacier runoff increases that exceed 5% in at least one month, in dark blue. The results refer to multi-GCM means and RCP4.5. Small dots refer to population density > 100 km−2 on a 0.5 × 0.5° grid as an indicator for potential downstream socio-environmental impacts.

Case study: Glaciers and water resources in the Himalaya

In the Himalaya, Karakorum and Hindu Kush mountains, millions of people rely on the 90,000 glaciers as a water resource9. These glaciers form the headwaters of the Indus, Ganges and Brahmaputra rivers. Glaciers here are highly sensitive to climate change, and are rapidly shrinking7,13. The developing countries in these catchments use this water for agriculture and hydropower, and are vulnerable to changes in their water supply14.

The contribution of glaciers to runoff varies in each basin, ranging from 18.8% in the Dudh Koshi catchment (a major tributary to the Ganges), to 80.6% in the Hunza catchment, which drains into the Indus basin9.

In High Mountain Asia, the glacial ice acts to protect against extreme water shortages on seasonal and longer timescales, because the glacial melt is sustained through droughts while all other stores of water in the basin decline14. Hydrological modelling predicts a decline in glacial meltwater contribution to the overall catchment hydrology by 2065 AD of -8% in the Indus, -18% in the Ganges and -20% in the Brahmaputra1.

In southern China, just north of the border with Nepal, one unnamed Himalayan glacier flows from southwest to northeast, creeping down a valley to terminate in a glacial lake. At the end of the glacier’s deeply crevassed snout sits a glacial lake, coated with ice in this wintertime picture. Just as nearby mountains cast shadows, the crevassed glacier casts small shadows onto the lake’s icy surface. This glacial lake is bound by the glacier snout on one end, and a moraine—a mound formed by the accumulation of sediments and rocks moved by the glacier—on the other. Source:

Glacier water resources in the Indus catchment

In the westerly Indus catchment, meltwater dominates water inputs during drought summers, and predicted glacier loss will add considerably to drought-related water stress14. The Indus and Aral basins are dominated by wet winters, dry summers, and have extensive glaciation14. The summer monsoon in these more westerly basins is also less dominant than that further east.

Map of the Indus River basin with tributaries labeled. Yellow regions are non-contributing parts of the watershed (e.g. the Thar Desert). From Wikimedia Commons (Keenan Pepper,

In these basins, the highest proportion of glacial melt to overall basin hydrology occurs in the upper basins, closer to the glaciers. In the Indus basin, two thirds of the population (>120 million people) lives in the middle altitudes, where glacial meltwater is more significant. The use of water for hydropower and irrigation is concentrated at dams and barrages with average altitudes of 936-1484 m above sea level, in these middle altitudes14. In the Indus, 121 of 143 existing or planned dams are glacier-fed. In the upper Indus, without glaciers, summer monthly water flows would be reduced by 38%, and up to 58% in drought years. Water stress is likely to peak in the relatively dry summers in drought years as the glacier melt declines.

Further reading


1              Immerzeel WW, van Beek L P H & P, B. M. F. Climate change will affect the Asian water towers. Science 328, 1382–1385 (2010).

2              Beniston, M. Climatic Change in Mountain Regions: A Review of Possible Impacts. Climatic Change 59, 5-31, doi:10.1023/a:1024458411589 (2003).

3              Huss, M. & Hock, R. Global-scale hydrological response to future glacier mass loss. Nature Climate Change 8, 135-140, doi:10.1038/s41558-017-0049-x (2018).

4              Barnett, T. P., Adam, J. C. & Lettenmaier, D. P. Potential impacts of a warming climate on water availability in snow-dominated regions. Nature 438, 303, doi:10.1038/nature04141 (2005).

5              Bamber, J. L., Westaway, R. M., Marzeion, B. & Wouters, B. The land ice contribution to sea level during the satellite era. Environmental Research Letters (2018).

6              Zemp, M. et al. Historically unprecedented global glacier decline in the early 21st century. Journal of Glaciology 61, 745-762, doi:10.3189/2015JoG15J017 (2015).

7              Zemp, M. et al. Global glacier mass changes and their contributions to sea-level rise from 1961 to 2016. Nature, 1 (2019).

8              Gardner, A. S. et al. A Reconciled Estimate of Glacier Contributions to Sea Level Rise: 2003 to 2009. Science 340, 852-857, doi:10.1126/science.1234532 (2013).

9              Rowan, A. V. et al. The sustainability of water resources in High Mountain Asia in the context of recent and future glacier change. Geological Society, London, Special Publications 462, 189-204 (2018).

10           Frans, C. et al. Implications of decadal to century scale glacio‐hydrological change for water resources of the Hood River basin, OR, USA. Hydrological processes 30, 4314-4329 (2016).

11           Immerzeel, W., Pellicciotti, F. & Bierkens, M. Rising river flows throughout the twenty-first century in two Himalayan glacierized watersheds. Nature geoscience 6, 742 (2013).

12           Hejazi, M. I. et al. Integrated assessment of global water scarcity over the 21st century under multiple climate change mitigation policies. Hydrology and Earth System Sciences 18, 2859-2883 (2014).

13           Bolch, T. et al. The state and fate of Himalayan Glaciers. Science 336, 310-314 (2012).

14           Pritchard, H. D. Asia’s glaciers are a regionally important buffer against drought. Nature 545, 169-174, doi:10.1038/nature22062 (2017).

Roches moutonnées

Roches moutonnées are asymmetric bedrock bumps or hills with a gently sloping and abraded upglacier (stoss) face and a quarried (or plucked) downglacier (lee) face that is typically blunter1,2. A good example of a roche moutonnée is shown in the image below.

Roche moutonnée from near Castle Loch, southwest Scotland, with a gently sloping (abraded) stoss face and a blunt (quarried) lee face. Ice flow was from left to right. Photo: David Baird

Roches moutonnées range in size from several metres to several hundreds of metres across, and often occur in clusters1 (see image below). They may be found emerging from beneath actively deglaciating ice masses (see image below), or on the sides and bottom of deglaciated valleys where they were once overridden by glacial ice3,4. Their distinctive form, which is partly linked with the orientation of glacier flow, make roches moutonnées useful to glaciologists aiming to reconstruct the flow direction of former glaciers.

Cluster of roches moutonnées (white arrows) in Porsangerfjorden, northern Norway. Ice flow was from right to left. Photo: Arnstein Rønning

Roches moutonnées emerging from beneath Goldbergkees glacier (Austria) as the ice thins and retreats. Photo: Ewald Gabardi

How do roches mountonnées form?

Roches mountonnées develop their distinctive morphology due to the pattern of stress on a bedrock surface beneath a sliding glacier, as shown in the diagram below. On the stoss side of bedrock bumps, normal stresses are relatively high and particles embedded in the ice are moved across the underlying surface where they carry out abrasion5,6. The evidence of such abrasion is the common occurrence of striations (i.e. scores and scratches on bedrock) on the sloping upper surface and flanks of roches moutonnées (see image below).

Formation of a roche moutonnée as a result of stress differences over the bedrock surface. High normal stress (pressure) on the stoss face results in bedrock abrasion, whereas lower normal stresses (pressure) on the lee face often allow a cavity to form, which promotes quarrying of bedrock along lines of existing weakness (e.g. bedrock joints). Diagram: Jacob M. Bendle

Striations on the flank of a roche moutonnée in Mount Rainier National Park, USA, giving evidence of glacial abrasion. Photo: Walter Siegmund

On the lee side of bedrock bumps, normal stresses are lower, which allows a cavity to form between the ice and bed (see diagram above) and prevents abrasion. In its place, bed cavities increase stress build up in the bedrock immediately upstream of the cavity, causing rock fracture and erosion by quarrying (or plucking). This process is particularly efficient where water pressure at the bed regularly changes3,7,8,9 (see diagram below).

The importance of bed cavities in roche moutonnée formation. In T1, the water pressure (pw) present in the bed cavity in the lee of a bedrock bump offsets the downward directed ice overburden pressure (pi), preventing bedrock fracture. However, in T2, the water has drained, and a high stress zone (red) develops in the bedrock around the cavity edge, which causes rock fracture and quarrying (plucking) to occur. Diagram: Jacob M. Bendle

The quarrying of rock at the lee end of roches mountonnées is also strongly influenced by the joint distribution in the parent rock3, and determines the size and shape of quarried rock fragments (see diagram below).

The importance of bedrock joint structure in the evolution of quarrying of a roche moutonnée lee face. The orange lines depict the progressive upglacier migration of the lee face as bedrock fragments are progressively plucked along lines of weakness (joints). Diagram: Jacob M. Bendle

What do roches mountonnées tell us about former glaciers?

Through an understanding of how roches mountonnées are formed, glaciologists are able to make inferences about the nature of past glacier systems where such landforms are found.

As roches mountonnées are most likely to form where cavities exist at the glacier bed, it is common for them to develop where the ice overburden pressure is low (i.e. where ice is relatively thin). Such conditions occur beneath thin cirque or valley glaciers, or near the margins of ice sheets3,4,10. This also means that roches moutonnées may be more likely to develop during deglaciation, when a glacier or ice sheet thins, ice overburden pressure decreases, and gaps between the ice and bed open up11 (see diagram below).

During full glacial conditions, when ice is at its thickest, ice overburden pressure (pi) is high and the glacier presses down into bumps in the bed. As the ice thins during deglaciation, the ice overburden pressure (pi) decreases and cavities open up at the bed, promoting favourable conditions for roche moutonnée formation. Diagram: Jacob M. Bendle (based on Roberts and Long, 2005)

Because roche mountonnée formation is also aided by fluctuations in basal water pressure, they are most likely to occur beneath warm-based (temperate) glaciers with hydrological systems that direct meltwater the bed10. The fact that they contain abraded (i.e. polished and striated) surfaces (see image above) also informs glaciologists that the ice responsible for their formation was (at least at times) warm based and moving by basal sliding, as well as carrying a basal debris load.


[1] Bennett, M.R., and Glasser, N.F. (2009) Glacial Geology: Ice Sheets and Landforms. Wiley-Blackwell.

[2] Benn, D.I., and Evans, D.J.A. (2010) Glaciers and Glaciation. Routledge.

[3] Sugden, D.E., Glasser, N.F., and Clapperton, C.M. (1992) Evolution of large roches moutonnées. Geografiska Annaler, 74A, 253-264.

[4] Glasser, N.F. (2002) Scottish Landform Example 28: The large roches moutonnées of upper Deeside. Scottish Geographical Journal, 118, 129-38.

[5] Boulton, G.S. (1974) Processes and patterns of glacial erosion. In Glacial Geomorphology (ed. D.R. Coates) Springer, Dordrecht, pp. 41-87.

[6] Hallet, B. (1979) A theoretical model of glacial abrasion. Journal of Glaciology, 23, 39-50.

[7] Iverson, N.R. (1991) Potential effects of subglacial water pressure fluctuations on quarrying. Journal of Glaciology, 37, 27-36.

[8] Hallet, B. (1996) Glacial quarrying: a simple theoretical model. Annals of Glaciology, 22, 1-8.

[9] Cohen, D., Hooyer, T.S., Iverson, N.R., Thomason, J.F. and Jackson, M. (2006). Role of transient water pressure in quarrying: A subglacial experiment using acoustic emissions. Journal of Geophysical Research: Earth Surface, 111(F3).

[10] Hall, A.M., and Glasser, N.F. (2003) Reconstructing the basal thermal regime of an ice stream in a landscape of selective linear erosion: Glen Avon, Cairngorm Mountains, Scotland. Boreas, 32, 191-208.

[11] Roberts, D.H., and Long, A.J. (2005) Streamlined bedrock terrain and fast flow, Jakobshavns Isbrae, West Greenland: implications for ice stream and ice sheet dynamics. Boreas, 34, 25-42.


Subglacial erosion

What is subglacial erosion?

Subglacial erosion refers to processes that act at a glacier or ice sheet bed that cause the Earth’s surface to be worn down, broken up, and transported by ice. These processes leave behind some of the classic signs of glacial activity, in the form of erosional landforms and landscapes.

Subglacial erosion is one of the key components of the glacial system, yet it remains poorly understood despite decades of research. This is largely due to the inaccessible nature of glacier beds, which limit the opportunity for direct observations or measurements1,2.

Because of this, processes of subglacial erosion have been based on theoretical models3-6 or inferred through investigations of landforms left behind in deglaciated areas. Nonetheless, direct access to glacier beds has been achieved in rare instances7,8, by accessing natural or artificial tunnels, which allow the installation of monitoring equipment and direct measurements to be made.

These approaches have identified two main mechanisms of subglacial erosion:

  • Glacial abrasion, the wearing down of bedrock surfaces
  • Glacial plucking or quarrying, the removal of rock fragments and blocks from the bed

Glacial abrasion

Glacial abrasion is the wear of a bedrock surface by rock fragments transported at the glacier base. This can happen by (i) the scoring (striation) of bedrock by rock particles (usually > 1 cm) embedded in the glacier sole, due to ice flow across a rock surface (see image below); and (ii) the polishing of bedrock surfaces by smaller, silt-sized particles that are dragged across the bedrock1.2.

Fine-grained debris frozen to the basal ice of Nigardsbreen glacier, west Norway, with debris coming into contact with underlying bedrock. Photo: Jacob M. Bendle

Scoring results in the formation of thin, linear grooves across a bedrock surface (see image below). These are known as striations (or striae). While striations may appear smooth, close inspection of striae beds show they form by a series of small rock fractures due to the build-up of stress below a mobile rock particle9.

Crossing-cutting glacial striations in bedrock, Maine, USA. Photo: Neil P. Thompson

Polishing, on the other hand, results in the overall smoothing down of rough areas of the bed (see image below). This process can be likened to the effect of sandpaper on wood.

Glacially smoothed bedrock recently uncovered due to retreat of Nigardsbreen glacier, west Norway. Photo Jacob M. Bendle

Controls on glacial abrasion

Rate of basal sliding

As glacial abrasion is caused by the movement of rock particles across bedrock, it is closely associated with basal sliding1,2. In warm-based (temperate) glaciers, where ice exists at the pressure melting point throughout, basal sliding occurs and a high flux of debris is dragged across a bedrock surface. By contrast, cold-based glaciers are frozen to their beds, so sliding rates are very low and the ability to abrade the bed limited1,2.

Debris concentration

Along with the rate of basal sliding, the amount of debris embedded in the glacier base also influences the rate of abrasion5. However, it is not as simple as a higher debris load resulting in faster rates of abrasion. In fact, a glacier with a high basal debris concentration results in friction between the ice and its bed, slowing the rate of basal sliding (see diagram below). Instead, glacial abrasion is most effective where basal debris is relatively sparse, as the reduced friction promotes faster sliding1.

The effect of basal debris concentration and glacier sliding on abrasion rate. Graph redrawn after Bennett and Glasser (2009)

Glacial plucking or quarrying

Plucking or quarrying is the fracture and removal of larger rock fragments (>1 cm) from the bed. In much the same way as the striation process, plucking occurs where stress build-up beneath an overriding rock particle results in the expansion of pre-existing cracks in the bedrock and the detachment of rock fragments1,2 (see image below). The fractured bedrock can then entrained by the overriding glacier and transported downglacier.

Vertical  joints and fractures observed in the bedrock of the formerly glaciated Ibañez valley, central Patagonia. Ice flow was from left to right. In the foreground, a former zone of plucking is illustrated by the ‘missing’ blocks of bedrock at the downglacier end of a roches mountonnée. Photo: Jacob M. Bendle.

There are several main ways by which plucked rock fragments can be entrained into the base of a glacier:

  • Debris can be frozen-on to the glacier sole as meltwater refreezes in the low-pressure zone in the lee (i.e. the downglacier end) of bedrock obstacles (see diagram below); this process is therefore strongly associated with the regelation mechanism of basal sliding
  • Debris can be frozen-on to the glacier sole at the boundary between warm-based (temperate) and cold-based ice, for example, approaching the glacier snout where ice thickness decreases and pressure melting of basal ice is inhibited (see diagram below)
  • Debris can be simply dragged from the bedrock and enveloped into basal ice, particularly where it is very loose

Freeze-on of plucked (quarried) debris in a low pressure zone at the downglacier end of a bedrock bump, due to refreezing of meltwater associated with regelation. Source: Jacob M. Bendle.

Meltwater and debris frozen in to basal ice layers at the transition from warm based to cold based ice (in this instance, at the glacier margin). Source: Jacob M. Bendle

Controls on plucking

Bedrock lithology

As described above, plucking tends to be focused along pre-existing cracks in bedrock8. The lithology (or rock type) of the bed will therefore influence its resistance to erosion. For example, in well-jointed rocks with deep, near-continuous cracks (e.g. shale), plucking rates will be higher than in rocks with fewer or more widely-spaced joints and cracks (e.g. granite).


Both theoretical models10 and direct observations beneath modern glaciers8 show that the presence of cavities at the bed is an important control on plucking. When a cavity is filled with water, water pressure offsets the overburden pressure resulting from the weight of overlying ice (‘T1’ in diagram below). In this situation, stresses in the bed are highest adjacent to the cavity. If water leaves a cavity and the water pressure drops, stresses in the bedrock increase considerably and lead to plucking of rock fragments (‘T2’ in diagram below).

In T1, the water pressure (pw) associated with a water-filled cavity in the lee of a bedrock bump offsets the downward directed ice overburden pressure (pi), preventing bedrock fracture. In T2, the water has drained, and a high stress zone (red) develops in the bedrock around the cavity edge, which may result in bed fracture and plucking. Source: Jacob M. Bendle

Therefore, plucking rates will be highest where the bed surface is undulating (i.e. there are abundant sites for cavities to form) and when the supply of meltwater causes fluctuation in water pressure, such as during the ablation season, where diurnal (day-night) melting patterns often develop1,2.


Subglacial erosion occurs at all ice masses, from small cirque glaciers to large continental ice sheets. It is also fundamentally linked to ice motion (e.g. sliding) and, in turn, mass balance regime and glacier thermal regime. Subglacial erosion processes therefore offer an excellent example of the connections between various components of the glacial system.

Other pages in this section of the site explore the effect of glacial erosion on Earth’s surface morphology.


[1] Bennett, M.R., and Glasser, N.F. (2009) Glacial Geology: Ice Sheets and Landforms. Wiley-Blackwell.

[2] Benn, D.I., and Evans, D.J.A. (2010) Glaciers and Glaciation. Routledge.

[3] Boulton, G.S. (1974) Processes and patterns of glacial erosion. In Glacial Geomorphology (ed. D.R. Coates) Springer, Dordrecht, pp. 41-87.

[4] Hallet, B. (1979) A theoretical model of glacial abrasion. Journal of Glaciology23, 39-50.

[5] Hallet, B. (1981) Glacial abrasion and sliding: their dependence on the debris concentration in basal ice. Annals of Glaciology2, 23-28.

[6] Iverson, N.R. (2012) A theory of glacial quarrying for landscape evolution models. Geology40, 679-682.

[7] Cohen, D., Iverson, N.R., Hooyer, T.S., Fischer, U.H., Jackson, M. and Moore, P.L. (2005). Debris‐bed friction of hard‐bedded glaciers. Journal of Geophysical Research: Earth Surface110(F2).

[8] Cohen, D., Hooyer, T.S., Iverson, N.R., Thomason, J.F. and Jackson, M. (2006). Role of transient water pressure in quarrying: A subglacial experiment using acoustic emissions. Journal of Geophysical Research: Earth Surface111(F3).

[9] Drewry, D.J. (1986) Glacial Geologic Processes. Edward Arnold, London.

[10] Morland, L.W., and Morris, E.M. (1977) Stress in an elastic bedrock hump due to glacier flow. Journal of Glaciology, 18, 67-75.


Glacial depositional landforms

This section of the website includes many examples of landforms created underneath and around the margins of glaciers. These depositional landforms typically form in two domains: subglacial landforms and ice-marginal landforms.

Subglacial landforms include:

  • A continuum of lineated bedforms, ranging from small scale (flutes), through to intermediate scale (10s of metres; Drumlins), through to large scale (kilometres; Megascale glacial lineations).
  • Sediments and landforms associated with meltwater, such as eskers.

Ice-marginal landforms include:

  • Piles of debris formed at the ice margin, such as moraines;
  • Till plains formed underneath the ice sheet;
  • Fluvioglacial landforms such as kames, outwash plains, meltwater channels.

There are lots of examples of these types of landforms across the Patagonian Ice Sheet.

Improvements to

Here at we have been busy making many updates to the website. We are particularly keen to update the website to bring it in to line with the reformed A-Level syllabus, and also to update and rewrite some of the older content, and improve the website as a resource to promote public understanding of glaciers and climate change.

Since was founded 6.5 years ago, we have undergone substantial improvements and learned a lot over the years. This outreach endeavour, motivated by a desire to publicly communicate the risks that climate change and rising sea levels pose to our world’s glaciers and ice sheets, has evolved into one of the premier sites on this subject. This website aims to inspire both interested adults and also young people and school children with geology and geomorphology, and specifically targets teachers to supply them with engaging, original content that they can use in lesson planning.

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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).