Cirque glaciation landsystem of upland Britain

The Loch Lomond Stadial in Britain

Between around 13 and 11 thousand years ago, the climate in Britain, as well as across much of Northern Europe, cooled abruptly1. This short-lived cold period temporarily reversed the general pattern of warming that drove the retreat of ice sheets after the Last Glacial Maximum, causing glaciers to readvance in many mountain regions.

In Britain, this cold snap is known as the Loch Lomond Stadial. In the Loch Lomond Stadial, an ice cap grew over the western Highlands of Scotland2,3, along with other smaller icefields, valley glaciers, and cirque glaciers that formed in the mountains and uplands of Scotland, England (e.g. Lake District4) and Wales (e.g. Snowdonia5).

Loch Lomond Stadial cirque glaciers

In the Loch Lomond Stadial, cirque glaciers formed in areas that were close to the threshold for glaciation6, such as around the margins of larger icefields, or in areas where the climate was not suited (e.g. warmer melt season temperatures7) to forming larger glaciers, typically further away from the main centre of glaciation in the Scottish Highlands.

Cross-section through an idealised cirque glacier occupying a mountainside hollow.

Cirque glaciers occupied bedrock hollows (cirques) in mountain sides or the lee (downwind) side slopes of escarpments. Cirques with a north or northeasterly aspect were particularly favourable sites for glaciation5 as they protected the ice from direct solar radiation for much of the day, resulting in less ice-melt across the year8.

Map of Loch Lomond Stadial cirque glaciers in Snowdonia, North Wales, showing a strong preference for northeasterly facing ice masses. Image from Bendle & Glasser (2012; ref. 5)

In addition, southwesterly prevailing winds blew snow from mountain summits and plateaus into the cirques below that, along with avalanches from steep cirque sides, added to glacier mass5,9.

In summary, topography played an important role in Loch Lomond Stadial cirque glaciation in Britain and, in turn, the landsystem these cirque glaciers created6.

Landforms created by Loch Lomond Stadial cirque glaciers

The cirque glacier landsystem of upland Britain6 contains landforms created directly by glacier ice, and landforms related to periglacial and paraglacial activity outside the limits of glacier cover.

Inside the limits of glaciation

Moraines

The maximum extent of Loch Lomond Stadial cirque glaciers is typically marked by a terminal moraine6. This may occur as a single, arcuate terminal ridge, or as a small belt of moraines around the maximum ice extent. Sometimes, although not always, recessional moraines extend some way back into the cirque floor, recording active glacier retreat5.

Arcuate, boulder-covered moraine ridges marking the limit of a former cirque glacier in Cwm Cau, mid-Wales. Note the hummocky drift enclosed inside the outer terminal moraine. Note also, talus slopes formed outside the limit of glaciation. Photo: J. Bendle.

Sometimes, terminal moraines are large in comparison to the cirque glacier that formed them. Such large moraines form in two main scenarios. First, where a glacier snout remained stable at a given location for a prolonged period of time6, allowing a large amount of debris to build up around its margin. Second, where glacier advance entrained a large amount of debris from the cirque floor and sides (possibly left behind by earlier glacial and paraglacial activity).

Lateral moraine of a Loch Lomond cirque glacier that formerly filled the Snowdon Horseshoe valley, North Wales. Photo: J. Bendle.

Drift limits

The limit of cirque glaciers is not always marked by moraine ridges5. Sometimes, glacial extent is recorded by ‘drift’ – a fairly homogenous blanket of glacial diamict (‘till’). The drift covered floor of some cirques contrasts greatly with the drift-free slopes above, allowing the vertical thickness of ice to be estimated10.

Glacial drift (the light-coloured ground dotted with boulders) marking the limit of cirque glaciers in Snowdonia, North Wales. Image from Google Earth.

Hummocky drift

Inside terminal moraines, it is common to observe hummocky drift mounds, which display no obvious alignment to a former ice margin5,6 and may reflect the wastage of ice and/or the chaotic dumping of debris during deglaciation.

Hummocky drift mounds (foreground) on the floor of Cwm Idwal, Snowdonia (North Wales) and a chain of lateral moraine ridges (background). Photo: J. Bendle.

Erosional landforms

Along with the depositional landforms described above, the floor and sides of cirques were often eroded by Loch Lomond Stadial glaciers, forming of ice-moulded bedrock, roches moutonnées, and striations5, which show that cirque glaciers were (at least at times) warm based.

Striations (thin grooves running from right to left) scored into bedrock on the floor of a cirque in Snowdonia, North Wales. Photo: J. Bendle.

Outside the limits of glaciation

Summit blockfields and frost weathered debris

Blockfields and frost-weathered debris are commonly found on the mountain summits above cirque basins, and talus slopes often blanket cirque sides above the limit of glaciation. These periglacial features, formed by frost-weathering in extremely cold conditions11, are therefore a useful indicator of the vertical thickness of ice12.

Frost-weathered debris on mountain summits above the limit of Loch Lomond Stadial cirque glaciation in Snowdonia, North Wales. Photo: J. Bendle.
Valley side talus slopes marking the upper (vertical) limit of cirque glaciation in Cwm Idwal, Snowdonia. Photo: J. Bendle.

Protalus ramparts

Protalus ramparts have the appearance of moraine ridges but were not formed by glacier ice13,14. Instead, they formed around perennial snowbeds, where debris weathered from the cirque backwall or sides fell on to the snowbed and slid or rolled downslope to accumulate as ridges around the snowbed margin.

Protalus ramparts – while looking like moraines – form where rock debris falls, rolls, and slides across a perennial snowbed to build up a ridge (or rampart) at the snowbed edge.

Rockslope failures

Rockslopes failures often create moraine-like ridges and/or hummocky deposits that may be mistaken for glacier limits, especially when they occur in cirques15. Rockslope failures are, however, paraglacial features (i.e. features formed by unstable conditions following the retreat of glacial ice from an area16), mostly formed because of high seismic activity caused by postglacial rebound following the last ice sheet glaciation of Britain17.

Boulder-capped rockslope failure from the backwall of Cwm Bochlwyd in Snowdonia, North Wales. Photo: J. Bendle.

The cirque glacier landsystem of the Loch Lomond Stadial

In summary, the cirque glaciation landsystem6 created throughout upland Britain during the Loch Lomond Stadial contains: (1) outer limits marked by moraines and drift, with recessional moraines on some cirque floors indicating active retreat in deglaciation; (2) erosional landforms, such as roches moutonnées and striations that provide evidence of warm based ice; and (3) periglacial (e.g. talus slopes) and paraglacial (rock slope failures) landforms created outside glacier limits.

An example of the Loch Lomond Stadial cirque glacier landsystem at Cwm Idwal, Snowdonia (North Wales). Lateral moraines – and a bouldery ridge beyond Llyn (lake) Idwal that approximates the former glacier terminus – enclose hummocky drift and areas of ice-moulded bedrock. Outside the glacier limits, talus slopes blanket the valley side. Photo: J. Bendle.

References

[1] Rasmussen, S.O., Bigler, M., Blockley, S.P., Blunier, T., Buchardt, S.L., Clausen, H.B., Cvijanovic, I., Dahl-Jensen, D., Johnsen, S.J., Fischer, H. & Gkinis, V. (2014) A stratigraphic framework for abrupt climatic changes during the Last Glacial period based on three synchronized Greenland ice-core records: refining and extending the INTIMATE event stratigraphy. Quaternary Science Reviews106, 14–28.

[2] Golledge, N.R. (2007) An ice cap landsystem for palaeoglaciological reconstructions: characterizing the Younger Dryas in western Scotland. Quaternary Science Reviews 26, 213–229.

[3] Golledge, N.R. (2010) Glaciation of Scotland during the Younger Dryas stadial: a review. Journal of Quaternary Science, 25, 550–566.

[4] McDougall, D.A. (2013) Glaciation style and the geomorphological record: evidence for Younger Dryas glaciers in the eastern Lake District, northwest England. Quaternary Science Reviews, 73, 48–58.

[5] Bendle, J.M. & Glasser, N.F. (2012) Palaeoclimatic reconstruction from Lateglacial (Younger Dryas Chronozone) cirque glaciers in Snowdonia, North Wales. Proceedings of the Geologists’ Association, 123, 130–145.

[6] Bickerdike, H.L., Ó Cofaigh, C., Evans, D.J.A. & Stokes, C.R. (2018) Glacial landsystems, retreat dynamics and controls on Loch Lomond Stadial (Younger Dryas) glaciation in Britain. Boreas, 47, 202–224.

[7] Ballantyne, C.K. (2007) Loch Lomond Stadial glaciers in North Harris, Outer Hebrides, North-West Scotland: glacier reconstruction and palaeoclimatic implications. Quaternary Science Reviews, 26, 3134–3149.

[8] 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.

[9] Mitchell, W.A. (1996) Significance of snowblow in the generation of Loch Lomond Stadial (Younger Dryas) glaciers in the western Pennines, northern England. Journal of Quaternary Science, 11, 233– 248.

[10] Ballantyne, C.K. (2002) The Loch Lomond Readvance on the Isle of Mull, Scotland: glacier reconstruction and palaeoclimatic implications. Journal of Quaternary Science, 17, 759–771.

[11] Curry, A., Jennings, S., Scaife, R. & Walden, J. (2007) Talus accumulation and sediment reworking at Mynydd Du. In Carr, S.J., Coleman, C.G., Humpage, A.J. & Shakesby, R.A. (eds.): The Quaternary of the Brecon Beacons: Field Guide, 120–127. Quaternary Research Association, London.

[12] Benn, D.I. & Ballantyne, C.K. (2005) Palaeoclimatic reconstruction from Loch Lomond Readvance glaciers in the West Drumochter Hills, Scotland. Journal of Quaternary Science, 20, 577–592.

[13] Shakesby, R.A. & Matthews, J.A. (1993) Loch Lomond Stadial glacier at Fan Hir, Mynydd Du (Brecon Beacons), South Wales: critical evidence and palaeoclimatic implications. Geological Journal, 28, 69– 79.

[14] Carr, S.J. & Coleman, C.G. (2007) An improved technique for the reconstruction of former glacier mass-balance and dynamics. Geomorphology, 92, 76–90.

[15] Carr, S.J., Coleman, C.G., Evans, D.J.A., Porter, E.M. & Rea, B.R. (2007) An alternative interpretation of Craig y Fro based on mass balance and radiation modelling. In Carr, S.J., Coleman, C.G., Humpage, A.J. & Shakesby, R.A. (eds.): The Quaternary of the Brecon Beacons: Field Guide, 120–127. Quaternary Research Association, London.

[16] Ballantyne, C.K. (2002) A general model of paraglacial landscape response. The Holocene, 12, 371–376.

[17] Ballantyne, C.K., Sandeman, G.F., Stone, J.O. & Wilson, P. (2014) Rock-slope failure following Late Pleistocene deglaciation on tectonically stable mountainous terrain. Quaternary Science Reviews, 86, 144–157.

Ice-dammed lake landsystems

Ice-dammed lakes are a common feature of glaciated mountain ranges. They form wherever glacial ice blocks the drainage of rivers or meltwater. This includes:

  • where a glacier blocks a trunk or tributary valley; and
  • where a glacier fills an overdeepened valley created by glacial erosion
Ice-dammed lakes form where glaciers block the flow of water in either a trunk or tributary valley (left), and where a glacier terminates in an overdeepened basin (right) that lake water cannot escape from.

Today, ice-dammed lakes exist at the margins of many mountain valley or icefield glaciers. During the last Ice Age, when glaciers were expanded globally, huge ice-dammed lakes formed when continental ice sheets advanced and blocked the flow of river systems, causing water to pond up against their margins1,2.

Glacial lake dammed by the Perito Moreno glacier of the North Patagonian Icefield in southern South America. Source: L. Galuzzi.

Ice-dammed lakes create a unique landsystem that reflects the action of both glacial ice and water on the landscape3. The main landform and sediment assemblages related to ice-dammed lake activity are described below.

Landforms of ice-dammed lakes

The most characteristic landforms of ice-dammed lakes are features created at lake margins, which result from the erosional and depositional action of waves.

Shorelines

Some of the most common landforms related to ice-dammed lakes are wave-cut shorelines4,5. Shorelines are seen as distinct benches or terraces in glaciated landscapes that dip towards a current or former glacial lake and run unbroken for hundreds of metres up to tens of kilometres where large glacial lakes once existed. Shorelines are useful as they mark out the extent and elevation of ice-dammed lakes that no longer exist4,5.

Ice-dammed lake at the margin of the Viedma glacier in Patagonia. Notice the lake shorelines raised above the current lake level, which records a time when the lake level was at a higher elevation. Image: Planet Labs

At the very largest glacial lakes that formed in the last Ice Age, shorelines are seen to tilt upwards when moving upvalley from a former glacier terminus4-7. This is caused by the rebound of Earth’s crust after ice has retreated and no longer weighs down on the land surface6,7. Glacial lake shorelines can, therefore, be used to work out the rate of Earth surface rebound (known as postglacial rebound) caused by the weight of former ice sheets.

Shorelines of a former ice-dammed lake at Lago Buenos Aires, in central Patagonia, eroded into the flank of a terminal moraine. Photo: J. Bendle.

Deltas

Deltas are another common landform related to ice-dammed lakes. Deltas are masses of sediment that build out into lakes at the point where a river meets standing water. In glaciated areas, rivers often carry large sediment loads that allow deltas to grow rapidly in size8.

Delta at the point a river flows into Tuttilik Lake in eastern Greenland. Nigertuluk Glacier in the background. Photo: Qeqertaq

There are many types of delta, but the most common at ice-dammed lakes are known as Gilbert-type deltas (after the American geologist Grove Gilbert)9. Gilbert-type deltas have three main parts10,11: topsets, fluvial sediments deposited at the delta surface, foresets, sediments deposited underwater on the steep delta front that dips downward into the lake, and bottomsets, sediments deposited in deeper water at the base of the delta.

Relict lake delta (formed when a glacial lake higher than at present) raised above an actively-forming lake delta. Note the distinction between horizontally bedded topsets and steeply dipping foresets in the relict delta. Photo: J. Bendle.

Similar to shorelines, the surface of a delta (the topsets) marks the water level of a former ice-dammed lake. Often, a ‘staircase’ of deltas will form as the level of a lake (and the river that flows into it) drops over time (see photo above)5,12. Ice-dammed lakes can also partially or completely refill after being drained, and this may lead to new shorelines being cut into the front of older deltas by wave action5,12.

Shorelines cut into the front of a raised delta in Patagonia (south America) formed by changes in ice-dammed lake level. Image: Google Earth.

Beaches

Beaches are commonly found in close proximity to raised deltas and lake shorelines5,12, and form in shallow water near the lake edge3. Like coastal beaches, those formed at the edges of ice-dammed lakes are the product of wave action and longshore drift that deposits sand, gravel and cobbles around the lake margin3.

Beach sediments collecting at the shore of the Perito Moreno ice-dammed lake in Patagonia (south America). Photo: J. Lascar.

It is also common to observe beach ridges that closely mirror the shape of the lake shoreline and reflect short periods of time when waves moved sediment up the beach to a specific elevation13,14.

Beach ridges along the shore of Lago Buenos Aires in Patagonia (south America). Photo: J. Bendle.

Iceberg features

Unique to glacial lakes are features created by icebergs15,16. Icebergs broken off from the glacier drift across the lake pushed by wind and lake currents. As they drift, their keels may scour long grooves and plough marks into the sediment at the lake floor. When an iceberg becomes grounded on the lake bottom (usually in shallower water near the lake edges), it sinks down into soft lake sediments, creating craters and hollows that remain in the landscape after the icebergs have melted and lake drained3.

Iceberg scours on the former lake bed of Glacial Lake Agassiz, Manitoba, Canada. Linear scours are between around 100 and 150 m wide and reach up to 10 km in length. Image: Google Earth.

Grounding line fans

While moraines can form at the margins of glaciers that terminate in lakes, more common are subaqueous grounding line fans3,9. These are fan-shaped deposits that build up around meltwater channels at the base of a glacier as when meltwater drops the sediment load it is carrying as it enters deep lake water. When a glacier remains stable for some time, it is common for fans to link up along the base of the glacier margin, forming a chain of connected fans3,9 that – much like a moraine – record a former glacier position.

Simplified diagram of and subaqueous fan forming at the grounding line of a lake-terminating glacier. Glacial meltwater transfers subglacial sediment from beneath the ice and into the lake, where it accumulates in a fan that spreads out from the ice-margin. Source: J. Bendle.

Sediments of ice-dammed lakes

Ice-dammed lakes are sinks for the sediments transported by glacial meltwater or rivers.

Close to the glacier margin

As we have already seen with deltas, the largest particles are dropped from rivers at the lake edges, where they enter relatively still water3,9. In a similar way, the largest grains (sand, gravel and cobbles) entering the lake in meltwater plumes directly from a glacier are deposited close to the ice margin, often forming fans along the ice front (see above)3,9. Along with fans, debris (such as boulders) may fall from the glacier surface into the water and accumulate at the base of the terminus9.

At the lake bottom

Further away from the ice margin, fine silt and clays settle out of the water column to the lake bottom3. This material is moved to deeper parts of the lake in meltwater currents that flow from the ice margin known as underflows (which travel along the lake bed) interflows (which travel through the lake at intermediate levels) or overflows (which travel across the lake surface)9,17.

Underflows, interflows, and overflows entering an ice-dammed lake from glacier meltwater and river systems, and carrying fine-grained sediments (silt and clay) out into the lake. Source: J. Bendle.

It is common for this material to settle on the lake bottom as coarse (silt) and fine (clay) couplets known as varves18. This happens as only the heaviest material (silt) can fall to the lake floor during the summer period, when glacial meltwater disturbs the water column. The lightest material (clay) falls from suspension in winter, when meltwater stops entering the lake, and when the lake surface freezes over preventing disturbance of the water column by winds18.

Light-dark (summer-winter) couplets, known as varves. These varve sediments were formed in an Ice Age lake in Patagonia (south America). Four varves (or four years’ worth of sediment) are present in the photo. Source: J. Bendle.

Ice-rafted debris

A final unique feature of glacial lake sediments is ice-rafted debris, material that is contained in or on icebergs and which falls to the lake bottom when icebergs roll, tip, break up, or melt16,19.

Debris contained in and on the surface of icebergs in Bering Glacier lake in Alaska. as icebergs melt, roll and tip, debris is released into the lake and falls to the lake bottom. Source: Sam Beebe
Ice-rafted debris in Patagonian varve sediments. Source: J. Bendle

Case study: Ice Age glacial lakes in Patagonia

Patagonia is an area known for its numerous ice-dammed lakes, both in the present day20 around North and South Patagonian Icefields, and in the past when glaciers were larger than today5.

During the cool climate of the last Ice Age, glaciers of the North and South Patagonian Icefields expanded and joined together to form a large mountain ice sheet21. This barrier of ice blocked the flow of rivers to the ocean, and huge volumes of water ponded at the ice sheet edge. The best known of these lakes are the Lago Buenos Aires and Lago Pueyrredón ice-dammed lakes that formed around the expanded North Patagonian Icefield5.

The extent of the Patagonian Ice Sheet at the Last Glacial Maximum (LGM) and the location of the Lago Buenos Aires and Lago Pueyrredón ice-dammed lakes. Source: J. Bendle.

As glaciers retreated at the end of the last Ice Age, these lakes expanded greatly, forming shorelines, deltas and beaches that extend over one hundred kilometres upvalley of the maximum ice extent5,22-24.

Satellite image and mapped landforms formed at the margins of the former Lago Buenos Aires ice-dammed lake. Source: J. Bendle.

The gradual retreat of ice opened up new valleys over time, causing water to drain away and lower the lake surface5. This left great staircases of shorelines and raised deltas in the landscape, which record several large (about 100 m) drops in the lake level and the escape of meltwater along river valleys5,12.

The history of glacial lakes in central Patagonia during the end of the last Ice Age (see ref. 5) Source: J. Bendle.

Eventually, as glaciers broke up and the North and South Patagonian Icefield split apart, a huge flood of meltwater was released5,25. This sped along the Río Baker river and out to the Pacific Ocean, eroding deep gorges into bedrock and depositing huge bars topped with house-sized boulders as it went.

Landforms created by the drainage of ice-dammed lakes, including giant flood bars and eroded bedrock gorges (see ref. 25). Source: J. Bendle.

Today, glacial geologists use the landforms and sediments of these vast ice-dammed lakes to work out when and how glaciers changed during the demise of the last Ice Age5,26, how outburst floods changed the landscape25, and how meltwater released to the ocean may have altered regional climate24.

References

[1] Teller, J.T., 1995. History and drainage of large ice-dammed lakes along the Laurentide Ice Sheet. Quaternary International28, 83-92.

[2] Jensen, J.B., Bennike, O.L.E., Witkowsi, A., Lemke, W. and Kuijpers, A., 1997. The Baltic Ice Lake in the southwestern Baltic: sequence‐, chrono‐and biostratigraphy. Boreas26, 217-236.

[3] Teller, J.T., 2003. Subaquatic landsystems: large proglacial lakes. In Evans, D.J.A. Glacial Landsystems (pp. 348-371). Arnold London.

[4] Breckenridge, A., 2013. An analysis of the late glacial lake levels within the western Lake Superior basin based on digital elevation models. Quaternary Research80, 383-395.

[5] Thorndycraft, V.R., Bendle, J.M., Benito, G., Davies, B.J., Sancho, C., Palmer, A.P., Fabel, D., Medialdea, A. and Martin, J.R., 2019. Glacial lake evolution and Atlantic-Pacific drainage reversals during deglaciation of the Patagonian Ice Sheet. Quaternary Science Reviews203, 102-127.

[6] Broecker, W.S., 1966. Glacial rebound and the deformation of the shorelines of proglacial lakes. Journal of Geophysical Research71, 4777-4783.

[7] Clark, J.A., Hendriks, M., Timmermans, T.J., Struck, C. and Hilverda, K.J., 1994. Glacial isostatic deformation of the Great Lakes region. Geological Society of America Bulletin106, 19-31.

[8] Østrem, G., Haakensen, N. and Olsen, H.C., 2005. Sediment transport, delta growth and sedimentation in Lake Nigardsvatn, Norway. Geografiska Annaler: Series A, Physical Geography87, 243-258.

[9] Benn, D.I. and Evans, D.J.A., 2010. Glaciers and Glaciation (pp. 570-573)Routledge, London.

[10] Nemec, W., 1990. Aspects of sediment movement on steep delta slopes. In Coarse-grained deltas (Vol. 10, pp. 29-73).

[11] Smith, D.G. and Jol, H.M., 1997. Radar structure of a Gilbert-type delta, Peyto Lake, Banff National Park, Canada. Sedimentary Geology113, 195-209.

[12] Bell, C.M., 2009. Quaternary lacustrine braid deltas on Lake General Carrera in southern Chile. Andean Geology36, 51-66.

[13] Fisher, T.G., 2005. Strandline analysis in the southern basin of glacial Lake Agassiz, Minnesota and North and South Dakota, USA. Geological Society of America Bulletin117, 1481-1496.

[14] Lepper, K., Buell, A.W., Fisher, T.G. and Lowell, T.V., 2013. A chronology for glacial Lake Agassiz shorelines along Upham’s namesake transect. Quaternary Research80, 88-98.

[15] Woodworth-Lynas, C.M.T. and Guigné, J.Y., 1990. Iceberg scours in the geological record: examples from glacial Lake Agassiz. Geological Society, London, Special Publications53, 217-233.

[16] Eyles, N., Eyles, C.H., Woodworth-Lynas, C. and Randall, T.A., 2005. The sedimentary record of drifting ice (early Wisconsin Sunnybrook deposit) in an ancestral ice-dammed Lake Ontario, Canada. Quaternary Research63, 171-181.

[17] Ashley, G.M., 2002. Glaciolacustrine environments. In Modern and past glacial environments (pp. 335-359). Butterworth-Heinemann.

[18] Palmer, A.P., Bendle, J.M., MacLeod, A., Rose, J. and Thorndycraft, V.R., 2019. The micromorphology of glaciolacustrine varve sediments and their use for reconstructing palaeoglaciological and palaeoenvironmental change. Quaternary Science Reviews226, 105964.

[19] Ovenshine, A.T., 1970. Observations of iceberg rafting in Glacier Bay, Alaska, and the identification of ancient ice-rafted deposits. Geological Society of America Bulletin, 81, 891–894.

[20] Wilson, R., Glasser, N.F., Reynolds, J.M., Harrison, S., Anacona, P.I., Schaefer, M. and Shannon, S., 2018. Glacial lakes of the Central and Patagonian Andes. Global and Planetary Change162, 275-291.

[21] Hein, A.S., Hulton, N.R., Dunai, T.J., Sugden, D.E., Kaplan, M.R. and Xu, S., 2010. The chronology of the Last Glacial Maximum and deglacial events in central Argentine Patagonia. Quaternary Science Reviews29, 1212-1227.

[22] Turner, K.J., Fogwill, C.J., McCulloch, R.D. and Sugden, D.E., 2005. Deglaciation of the eastern flank of the North Patagonian Icefield and associated continental‐scale lake diversions. Geografiska Annaler: Series A, Physical Geography87(2), pp.363-374.

[23] Bell, C.M., 2008. Punctuated drainage of an ice‐dammed quaternary lake in southern south america. Geografiska Annaler: Series A, Physical Geography90(1), pp.1-17.

[24] Glasser, N.F., Jansson, K.N., Duller, G.A., Singarayer, J., Holloway, M. and Harrison, S., 2016. Glacial lake drainage in Patagonia (13-8 kyr) and response of the adjacent Pacific Ocean. Scientific Reports6, p.21064.

[25] Benito, G. and Thorndycraft, V.R., 2019. Catastrophic glacial-lake outburst flooding of the Patagonian Ice Sheet. Earth-Science Reviews, p.102996.

[26] Bendle, J.M., Palmer, A.P., Thorndycraft, V.R. and Matthews, I.P., 2017. High-resolution chronology for deglaciation of the Patagonian Ice Sheet at Lago Buenos Aires (46.5°S) revealed through varve chronology and Bayesian age modelling. Quaternary Science Reviews177, 314-339.

Introduction to glaciated valley landsystems

Glaciated valley landsystems refer to the landforms and sediments produced by valley glaciers in upland and mountainous environments1. As valley glaciers currently exist under a broad range of topographic and climatic settings across the globe2,3, the landsystems they create are equally varied.

The glaciated valley landsystems section of ‘AntarcticGlaciers’ will give examples of the range of different landscapes formed by valley glaciers. But before diving into specific examples, we suggest reading this page, which outlines the broad controls on the ‘style’ of valley glacier and the landforms and sediments they create.

Valley glaciers exist in many mountain ranges across the globe. The valley glacier pictured above, the Alestchgletscher, flows from the Jungfrau mountain area in the Swiss Alps. Photo: D. Beyer

What valley glaciers have in common

Let’s first look at what nearly all valley glaciers have in common. Most important, valley glacier behaviour and the landforms they create is largely related to two main factors1:

  • Topography, which strongly controls glacier size and shape (known as its morphology), as well as the transfer of mass (ice) and debris. As all valley glaciers are, by definition, confined by valley walls, their flow and interaction with the land surface is closely related to topography.
  • The amount of rock and sediment debris received from adjacent valley sides and carried at the ice surface (which, as we’ll see below, varies from glacier–to–glacier).
Fox Glacier in New Zealand (2013). Note that debris from the valley side has partly covered the ice surface. Photo: M. Basler

What controls valley glacier style?

Topography

Topography is important at several scales.

At the largest scale, the tectonic history of a region defines the size, number and altitude of mountains where glaciers can exist3. Valley glaciers occupying the highest mountain ranges, such as the Himalayas, for example, exist under a different set of climatic conditions than glaciers in lower altitude mountains, such as in Norway or Sweden. For this reason, valley glaciers can have a range of thermal regimes, which control glacier flow, debris erosion and transport, and the creation of landforms.

The debris-covered Khumbu glacier in the ‘high-relief’ Everest region of Nepal. Notice the steep valley sides that rise far above the glacier and supply its surface with rock and sediment debris. Photo: Vlunyak

At a more local scale, topography (and especially relief) to a large extent determines how much debris is supplied to the glacier surface1-3. For example, a valley with very steep sides is more likely to undergo regular mass movement (e.g. rock falls, landslides, slumps) that supply the glacier surface with rock and sediment debris than a valley with shallower sides.

Rockfalls and slumps from steep valley walls above Morteratschgletscher in the Swiss Alps. Photo: Samedan

Similarly, where there are large areas of rock exposed above the glacier, the chance of debris falling on to the ice surface is much greater than where there are very few exposed rocks on the valley walls that surround the glacier. Valleys with steep, high sides (that often rise >1000 m above the valley floor) are known as ‘high-relief’ areas, whereas valleys with less steep and lower sides are known as ‘low-relief’ areas.

Debris supply to glacier surfaces

As touched on above, the amount of debris covering a valley glacier surface can vary. Glaciers can be ‘clean’, meaning they have very little to no debris at the surface, or they can be ‘debris-covered’, where large areas (typically in the ablation zone) are completely mantled with rock and sediment debris.

Whether a glacier is ‘clean’ or debris-covered depends largely on how much and how often debris is supplied to the ice surface1. As we have seen above, the glaciers of high-relief areas, such as the Himalayas, Andes, or Southern Alps of New Zealand, are surrounded by large, high, and very steep valley sides that release huge volumes of debris to glacier surfaces through rock falls, slumps and landslides4. Some mountain areas are also tectonically active. In these cases, earthquakes can trigger extremely large rock avalanches that run out on to glaciers in the valley bottom, significantly increasing the amount of debris at the ice surface1,4.

The debris-covered tongue of the Tasman glacier in the Southern Alps of New Zealand. Photo

In other mountain areas – for example, where there is less exposed rock directly above a glacier’s surface, where the valley sides are less steep (and less prone to mass movement), or where the local geology is more resistant to failure and rockfall, the supply of debris to the glacier surface will be lower and the ice comparatively ‘clean’.

The largely ‘clean’ glacier surface of Nigardsbreen, western Norway. Photo: J. Bendle

How does debris cover influence glacier behaviour?

The amount of debris on the surface of a valley glacier can change its behaviour in several ways. First, it alters the glacier response to climate. Debris-covered glaciers have a muted response to climate (e.g. warming air temperature) as the debris that covers the ice surface (where thicker than several centimetres) insulates it against melting1-3. For this reason, the terminus position of debris-covered valley glaciers is generally stable for long periods of time. ‘Clean’ glaciers, on the other hand, respond rapidly to climate with shifts in terminus position, as the insulating effect of debris cover is far less important.

The debris-covered snout of the Exploradores glacier in Patagonia (South America). Thick debris cover can slow the rate of ice-melt by insulating it from solar radiation. Photo: J. Bendle

Second, it alters the type of landforms that valley glaciers create. At debris-covered glaciers, huge volumes of debris build-up at the relatively stable ice margins, often leading to the deposition of large latero-frontal moraines5,6. These moraines, in turn, influence the glacier response to climate, by providing a barrier to snout advance3.

Large latero-frontal moraines enclose the Mueller glacier, New Zealand, and its proglacial lake. Image: Google Earth (see below for a photo of the moraines).
Latero-frontal moraine of the Mueller glacier (see Google Earth image above) that rises around 80–100 m above the ice-front and proglacial lake. Large latero-frontal moraines like this form where valley sides release large volumes of debris to the glacier surface. Photo: K. Golik

At ‘clean’ glaciers, by contrast, there is less debris at the ice margin, and snout fluctuations mean that this debris may be ‘spread out’ across a larger area so that, in general, landforms such as moraines are smaller but more numerous (e.g. recessional moraines7-9).

The amount of meltwater

The amount of meltwater flowing through a valley glacier is controlled by annual temperature and precipitation (and is therefore related to climate) and water storage in the catchment (e.g. does water move quickly through a glacier, or does it get stored in glacial lakes?)

Where sediment and rock debris are transported quickly through a glacier by large volumes of meltwater, a greater amount of glaciofluvial (e.g. outwash) landforms are formed1,10 and the debris available to deposit moraines is reduced (leading to smaller moraines). These type of valley glaciers exist in humid mountain ranges that receive a lot of precipitation in a year. Examples include southern Chile, New Zealand, and Alaska.

Braided proglacial river network transporting meltwater and sediment away from the Tasman proglacial lake (New Zealand) and to form an outwash (sandur) plain. Photo: F. Rindler

By contrast, in colder, drier mountain areas, less meltwater is produced in a year and less sediment is washed away in proglacial streams. Therefore, debris transported to the glacier margins forms moraines, which can grow to be extremely large in size over time1. This type of glacier tends to exist in high-altitude and arid mountain ranges, such as parts of the Andes and Himalayas.

The main types of valley glacier

As we have seen, there are many (interrelated) factors that influence valley glacier style and, in turn, the landsystems they create. To summarise, they can be divided into types1 based on the amount of surface debris cover, with ‘clean’ and ‘debris-covered’ types, and based on the amount of meltwater they produce, where it is possible to have glaciers with efficient meltwater systems that wash large volumes of sediment from within the glacier and from around its margin, and glaciers with less efficient meltwater systems, where large volumes of debris can build up around their margins.

It is important to bear in mind that these four glacier types are ‘idealised’ examples. In reality, valley glaciers are extremely variable, as are the landforms and sediments they create. We will explore the various types of valley glacier and their landsystems further in this section of ‘AntarcticGlaciers’.

References

[1] 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.

[2] Benn, D.I. and Evans, D.J.A., 2010. Glaciers and glaciation. Routledge.

[3] Bennett, M.M. and Glasser, N.F.G., 2011. Glacial geology: ice sheets and landforms. John Wiley & Sons.

[4] Hambrey, M.J., Quincey, D.J., Glasser, N.F., Reynolds, J.M., Richardson, S.J. and Clemmens, S., 2008. Sedimentological, geomorphological and dynamic context of debris-mantled glaciers, Mount Everest (Sagarmatha) region, Nepal. Quaternary Science Reviews27(25-26), 2361-2389.

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

[6] 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.

[7] Matthews, J.A., 2005. ‘Little Ice Age’ glacier variations in Jotunheimen, southern Norway: a study in regionally controlled lichenometric dating of recessional moraines with implications for climate and lichen growth rates. The Holocene15(1), 1-19.

[8] Beedle, M.J., Menounos, B., Luckman, B.H. and Wheate, R., 2009. Annual push moraines as climate proxy. Geophysical Research Letters36(20).

[9] Lukas, S., 2012. Processes of annual moraine formation at a temperate alpine valley glacier: insights into glacier dynamics and climatic controls. Boreas41(3), 463-480.

[10] Kirkbride, M.P., 2000. Ice-marginal geomorphology and Holocene expansion of debris-covered Tasman Glacier. New Zealand, IAHS-AISH P264, pp. 211-217.

Glaciated valley landsystems

This section of ‘AntarcticGlaciers’ deals with valley glaciers and the landsystems they create.

Before diving into examples from specific glacier systems, we suggest first reading the Introduction to glaciated valley landsystems page, which covers valley glacier behaviour and how this influences the landforms and sediments they leave in the landscape.

Active temperate glacier landsystem

Temperate glaciers reach the pressure-melting point throughout, for at least for part of the year. Today, temperate glaciers are found in mild maritime climates such as southern Iceland, western Norway, New Zealand, and southern Chile, where both winter snowfall and summer melt rates are high.

Temperate glaciers are often very sensitive to changes in climate and will periodically advance (e.g. during the winter) even when in overall recession. This type of glacier is defined as active.

The lake-terminating ice margin of the active temperate Fjallsjökull glacier, Iceland. Photo: Wojciech Strzelecki

The active temperate glacier landsystem reflects the wet-based thermal regime of temperate glaciers, and their tendency to oscillate in response to seasonal temperature. The landforms created by lowland temperate glaciers (such as those in southern Iceland) fall into three groups: ice marginal landforms, subglacial landforms, and glaciofluvial and glaciolacustrine landforms.

Landform assemblages of active temperate glaciers

Ice-marginal landforms

One of the most characteristic features of the active temperate glacier landsystem are the sharp, low relief moraine ridges found on their forelands1,2. These moraines are typically <10 m high and mimic the shape of the glacier snout when deposited, often taking on a saw-tooth pattern that reflects the pattern of ice-margin crevasses3.

Sawtooth push moraines on the foreland of Skaftafellsjökull, southern Iceland. Photo: Chensiyuan

These moraines are formed by a combination of ice pushing and the dumping of sediment from the snout. Some sediment may also be squeezed out from beneath an advancing snout (either during winter advance, or during the summer when sediment beneath the snout can become saturated with water and more mobile).

Because active temperate glaciers often advance during winter and retreat during summer, a series of annual push moraines can form during deglaciation4-6.

Annual push moraines formed at the retreating terminus of Skálafellsjökull, Iceland (see ref. #3). The moraines display a sawtooth planform that closely mimics the shape of the ice margin. Image: Google Earth.

Subglacial landforms

The beds of former active temperate glaciers are characterised by landforms of both erosion and deposition1,2.

Where exposed at the land surface, bedrock is polished, moulded and striated. The bedrock may also be shaped into roches moutonnées, indicating that both abrasion and quarrying occur at active temperate glacier beds.

Streamlined subglacial landforms, such as flutes and drumlins, are also common on temperate glacier forelands1,7. These features form in large groups at right angles to push moraines (i.e. in the direction of former glacier flow) by some combination of subglacial deformation8,9 and the ploughing (erosion) of soft sediments by the overriding glacier10.

The foreland of Svínafellsjökull, Iceland, showing flutes and debris stripes aligned at right angles to push moraines in the direction of former ice flow (Evans et al., 2019). Image: Google Earth.

Temperate glacier forelands sometimes contain overridden moraines1,2, which are more subdued than push moraines and have flutes across their surfaces. These serve as evidence for ice-overriding during a glacier advance.

Glaciofluvial and glaciolacustrine landforms

While not unique to the active temperate glacier landsystem, glaciofluvial and glaciolacustrine landforms are common owing to the high volumes of meltwater released by temperate glaciers during the spring and summer months1,2.

Proglacial streams that flow away from the snout produce outwash (also referred to as sandur) fans11, whereas meltwater draining around the sides of the glacier form kame terraces and narrow outwash corridors1,2.

As sandur fans form in contact with the snout, they often develop ‘pitted’ surfaces where glacial ice is buried and later melts out, leaving small lakes at the outwash surface1,2.

Outwash (sandur) deposits around the margin of Skeiðarárjökull, southern Iceland. The many pits and pockmarks that break up the outwash surface form by the melting of buried ice over time. Image: Google Earth.

Some temperate glacier forelands also contain eskers, which are narrow, often sinuous ridges of glaciofluvial sand and gravel that form in subglacial, englacial and supraglacial (all ice-walled) channels, which give some indication of the patterns of meltwater drainage in former glaciers12.

Sinuous esker ridges on the foreland of the Breiðamerkurjökull glacier, southern Iceland. (see ref. #12). Image: Google Earth.

Ice-dammed or proglacial lakes are commonly found around the margins of receding temperate glaciers. These lakes interrupt the path of sediment-containing meltwater streams, allowing thick sequences of sediment to accumulate at the lake bottom1,2. Lake shorelines and deltas also form around glacial lake margins and often remain clear in the landscape after a lake has drained1.2.

Proglacial lake developing around the Skaftafellsjökull ice front (Iceland). Photo: Óðinn

The active temperate glacier landsystem

The active temperate glacier landsystem1,2 serves as a clear and detailed signature of past glacial activity, particularly of ice-front oscillations and meltwater drainage patterns.

Research has shown that active temperate glaciers existed in a wide range of formerly glaciated regions. For example, some Ice Age (around 20,000 years ago) glaciers of the Laurentide13 and Patagonian ice sheets14 and the New Zealand mountain ice cap15 produced landform assemblages typical of active temperate glacier activity.

Because the behaviour (e.g. seasonal advance and retreat patterns) of active temperate glaciers is closely tied to climate, identifying their landsystem in formerly glaciated areas can serve as a record of past climate.

In summary, the active temperate glacier landsystem1,2 usually contains: large areas of low amplitude push, dump and squeeze moraines (that mark out former glacier positions), which often record active annual recession; flutes, drumlins, and ice-moulded bedrock between moraine ridges; and extensive glaciofluvial (outwash, eskers, kame terraces) and glaciolacustrine (shorelines) features that provide evidence of abundant meltwater around the glacier snout.

References

[1] Evans, D.J.A. and Twigg, D.R., 2002. The active temperate glacial landsystem: a model based on Breiðamerkurjökull and Fjallsjökull, Iceland. Quaternary Science Seviews21, 2143-2177.

[2] Evans, D.J.A., 2003. Ice-marginal terrestrial landsystems: active temperate glacier margins. In Evans, D.J.A. (Ed.) Glacial Landsystems. Hodder–Arnold, London.

[3] Evans, D.J.A., Ewertowski, M. and Orton, C., 2016. Fláajökull (north lobe), Iceland: active temperate piedmont lobe glacial landsystem. Journal of Maps12, 777-789.

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

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

[6] Chandler, B.M., Evans, D.J.A. 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.

[7] Evans, D.J.A., Nelson, C.D. and Webb, C., 2010. An assessment of fluting and “till esker” formation on the foreland of Sandfellsjökull, Iceland. Geomorphology114, 453-465.

[8] Boulton, G.S., 1976. The origin of glacially fluted surfaces-observations and theory. Journal of Glaciology17, 287-309.

[9] Benn, D.I., 1994. Fluted moraine formation and till genesis below a temperate valley glacier: Slettmarkbreen, Jotunheimen, southern Norway. Sedimentology41, 279-292.

[10] Tulaczyk, S.M., Scherer, R.P. and Clark, C.D., 2001. A ploughing model for the origin of weak tills beneath ice streams: a qualitative treatment. Quaternary International86, 59-70.

[11] Evans, D.J.A. and Orton, C., 2015. Heinabergsjökull and Skalafellsjökull, Iceland: active temperate piedmont lobe and outwash head glacial landsystem. Journal of Maps11, 415-431.

[12] Storrar, R.D., Evans, D.J.A., Stokes, C.R. and Ewertowski, M., 2015. Controls on the location, morphology and evolution of complex esker systems at decadal timescales, Breiðamerkurjökull, southeast Iceland. Earth Surface Processes and Landforms, 40, 1421-1438.

[13] Evans, D.J., Lemmen, D.S. and Rea, B.R., 1999. Glacial landsystems of the southwest Laurentide Ice Sheet: modern Icelandic analogues. Journal of Quaternary Science14, 673-691.

[14] Darvill, C.M., Stokes, C.R., Bentley, M.J., Evans, D.J.A. 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.

[15] Sutherland, J.L., Carrivick, J.L., Evans, D.J.A., Shulmeister, J. and Quincey, D.J., 2019. The Tekapo Glacier, New Zealand, during the Last Glacial Maximum: An active temperate glacier influenced by intermittent surge activity. Geomorphology343, 183-210.

Surging glacier landsystem

Glaciers gain mass in their upper reaches (accumulation zone) and lose mass at their snout (ablation zone). The majority of glaciers flow (and transfer mass) at a steady rate. However, some glaciers switch between periods of slow and fast flow.

Surging glaciers have relatively long periods of ice build-up and slow ice flow before a sudden release of mass and a short-lived period of much faster (sometimes up to 1000 times faster) ice flow1,2. These surge cycles are largely driven by internal processes and are unrelated to climate3,4 (note, however, that surge glaciers are generally found within an optimal climate envelope5).

The surge of Variegated Glacier, Alaska, in 1982–83 recorded by timelapse photography. (my advice: turn sound off!)

The two phases of glacier surging are known as the active phase and quiescent phase6. In the active phase, ice is moved rapidly from a reservoir zone (most commonly high up on the glacier) to the snout. In this phase, ice may flow at a rate of 10s of metres per day. This fast transfer of mass also tends to cause an advance of the glacier snout. In the quiescent phase between surges, glacier flow slows down, the snout stagnates, and ice once again builds up in the reservoir zone. The active phase of a surge can last from 1 year to 3-10 years, whereas the quiescent phase can be 10s or even 100s years long7,8,9.

As glacier fluctuations are commonly used to reconstruct past climate changes, the ability to distinguish between climate-driven advances and those not related to climate (i.e. glacier surges) is important.

There is no one landform on its own that is indicative of glacier surging. However, glacier surges leave behind a distinctive assemblage (or group) of landforms in the landscape.

Landform assemblages of surging glaciers

Push moraines and thrust-block moraines

The maximum limit of a surge is marked by moraines10,11. These may be single ridge push moraines, or zones of multiple closely spaced ridges that are pushed up in a single ‘block’ as a rapidly advancing snout deforms, compresses, and ‘thrusts up’ sediment on the glacier foreland. For this reason, these are known as thrust block moraines.

Push moraine ridges on the foreland of the Brúarjökull surge-type glacier in southern Iceland. Notice how sets of flutes grade to the moraine ridge. (see ref. 11 for further detail). Image: Google Earth.
Thrust-block moraines on the foreland of the Eyabakkajökull surge-type glacier in southern Iceland. Notice how each thrust block contains multiple ridges. (see ref. 11 for further detail and diagram below for explanation). Image: Google Earth.
Thrust-block moraines form as proglacial sediments are deformed (‘thrust up’) by the rapidly advancing surge snout. This creates a thrust moraine zone with numerous distinct ridges.

Hummocky moraine

During a surge, rapid flow causes the ice to stretch, bend, fold and fracture7. This deformation of ice (particularly a process known as ‘thrusting’) can move large volumes of sediment from the bed (or from within the glacier) up on to the glacier surface12. In the quiescent phase, after the surge has taken place, the melting of the stagnant snout gradually lowers this sediment to the land surface, where it forms a zone of hummocky mounds and hollows.

A simplified model of hummocky moraine formation owing to the thrusting of debris to the ice surface during a surge. Sediment is thrust to the glacier surface in the active phase. In the quiescent phase, the snout stagnates and melting gradually lowers debris to the land surface. Over time, moraine ice cores melt away leaving behind hummocky topography.

In the surging glacier landsystem10,11, belts of hummocky moraine can be found on the ice-contact slopes of push moraines or thrust-block moraines, as this is the area where stagnating ice develops most extensively during the quiescent phase.

Ice-cored outwash

Another common feature of surge glacier forelands are areas of ice-cored outwash10,11. Toward the end of surge, some glaciers release large volumes of meltwater that transport and deposit sediment across the stagnant glacier snout. As this sediment-covered ice melts over time, small lakes (kettles) may form at the outwash surface.

Ice-cored outwash on the foreland of the Eyabakkajökull glacier in southern Iceland. The numerous small lakes that give the outwash a ‘pitted’ appearance form as buried ice melts out over time (see ref. 11 for further details). Image: Google Earth.

Flutes

Flutes are streamlined ridges of sediment formed at the glacier bed. In themselves, flutes are not diagnostic of glacier surging, as they are found at many temperate glacier types. However, flutes formed by a glacier surge are often particularly long (over 1 km long in some cases) and unbroken, as they were created by a single, rapid glacier advance13.

Long (100s of metres) flutes and crevasse-squeeze ridges preserved on the foreland of the Brúarjökull surge-type glacier in southern Iceland. Notice how the crevasse-squeeze ridges do not parallel ice flow, but cut obliquely across streamlined flutes (see ref. 11 for further detail). Image: Google Earth.

In contrast to non-surging glacier types, flutes are often also found is close association with crevasse-squeeze ridges (see below) when formed during surges.

Crevasse-squeeze ridges

During a surge, glacier ice stretches and fractures7, which creates many crevasses that pass all the way through the glacier (video below), from the ice surface to the bed.

A helicopter flight over a surging glacier in Yukon’s Saint Elias Mountains shows extensive crevassing during the active phase of a surge. (note: Gwenn Flowers of Simon Fraser University also provides a great explanation of glacier surging and why surge-type glaciers are important). Filming: CBC News: The National.

Once the surge is over (when there is a drop in basal water pressure) sediment is squeezed upwards into open basal crevasses. As the snout stagnates and ice melts away during the quiescent phase following a surge, a cross-cutting network of crevasse-squeeze ridges is left behind in the landscape12,14.

A simplified explanation of crevasse-squeeze ridge formation by glacier surging. 1) Crevasses form through the snout during the active surge phase. 2) A drop in basal water pressure (pw) following the surge allows sediment to fill basal crevasses. 3) The glacier snout stagnates and thins in the quiescent phase. 4) Melting of glacial ice leaves behind crevasse-squeeze ridges in the landscape.

The surging glacier landsystem

While there are slight differences from glacier to glacier, the surging glacier landsystem10,11 tends to be spatially arranged in three main zones: an outer zone of push and thrust-block moraines (this represents the maximum extent of a surge), an intermediate zone of hummocky moraine (where snout stagnation occurs post-surge), and an inner zone of flutes, crevasse-squeeze ridges (where ice has overridden the foreland), and areas of ice-cored and pitted outwash.

The surging glacier landsystem of Eyabakkajökull glacier in southern Iceland (see ref. 11 for further details). Note the outer zone of push and thrust-block moraines, the intermediate zone of hummocky moraine, and the inner zone of flutes, crevasses-squeeze ridges and overridden moraines. Image: Google Earth.

Because glacier surging is cyclical in nature, the surging glacier landsystem may also contain evidence of several surge events, such as overridden moraines, which are smoothed by the overriding ice and often have flutes across their surface.

In summary, surging glaciers leave a distinct imprint on the land surface. No one landform is concrete evidence of a past glacier surge, but where a full assemblage of landforms (the surging glacier landsystem) occurs, the activity of past surging glaciers (in areas that are no longer glaciated) can be studied and reconstructed15,16.

References

[1] Clarke, G.K., Collins, S.G. and Thompson, D.E., 1984. Flow, thermal structure, and subglacial conditions of a surge-type glacier. Canadian Journal of Earth Sciences21, 232-240.

[2] Raymond, C.F., 1987. How do glaciers surge? A review. Journal of Geophysical Research: Solid Earth92), 9121-9134.

[3] Meier, M.F. and Post, A., 1969. What are glacier surges? Canadian Journal of Earth Sciences6, 807-817.

[4] Sharp, M., 1988. Surging glaciers: behaviour and mechanisms. Progress in Physical Geography12, 349-370.

[5] Sevestre, H. and Benn, D.I., 2015. Climatic and geometric controls on the global distribution of surge-type glaciers: implications for a unifying model of surging. Journal of Glaciology61, 646-662.

[6] Cuffey, K.M. and Paterson, W.S.B., 2010. The Physics of Glaciers, 4th Edn, Oxford: Butterworth-Heinemann.

[7] Kamb, B., Raymond, C.F., Harrison, W.D., Engelhardt, H., Echelmeyer, K.A., Humphrey, N., Brugman, M.M. and Pfeffer, T., 1985. Glacier surge mechanism: 1982-1983 surge of Variegated Glacier, Alaska. Science227, 469-479.

[8] Dowdeswell, J.A., Hamilton, G.S. and Hagen, J.O., 1991. The duration of the active phase on surge-type glaciers: contrasts between Svalbard and other regions. Journal of Glaciology37, 388-400.

[9] Eisen, O., Harrison, W.D. and Raymond, C.F., 2001. The surges of Variegated Glacier, Alaska, USA, and their connection to climate and mass balance. Journal of Glaciology47, 351-358.

[10] Evans, D.J.A. and Rea, B.R., 1999. Geomorphology and sedimentology of surging glaciers: a landsystems approach. Annals of Glaciology28, 75-82.

[11] Evans, D.J.A. and Rea, B.R., 2003. Surging glacier landsystem. In Evans, D.J.A. (Ed.) Glacial Landsystems. Hodder–Arnold. London.

[12] Sharp, M., 1985. “Crevasse-fill” ridges—a landform type characteristic of surging glaciers? Geografiska Annaler: Series A, Physical Geography67, 213-220.

[13] Evans, D.J.A., Twigg, D.R., Rea, B.R. and Shand, M., 2007. Surficial geology and geomorphology of the Brúarjökull surging glacier landsystem. Journal of Maps3, 349-367.

[14] Rea, B.R. and Evans, D.J.A., 2011. An assessment of surge‐induced crevassing and the formation of crevasse squeeze ridges. Journal of Geophysical Research: Earth Surface116 (F4).

[15] Evans, D.J.A., Lemmen, D.S. and Rea, B.R., 1999. Glacial landsystems of the southwest Laurentide Ice Sheet: modern Icelandic analogues. Journal of Quaternary Science14, 673-691.

[16] Sutherland, J.L., Carrivick, J.L., Evans, D.J.A., Shulmeister, J. and Quincey, D.J., 2019. The Tekapo Glacier, New Zealand, during the Last Glacial Maximum: An active temperate glacier influenced by intermittent surge activity. Geomorphology343, 183-210.

Palaeo-ice stream landsystem

Ice streams are corridors of fast-flowing ice within ice sheets that are flanked on either side by slowly moving ice1. Palaeo-ice streams are ice streams that existed in former ice sheets2,3, such as the continental ice sheets that grew during the last Ice Age. Glaciologists know that these palaeo-ice streams existed as they left a clear imprint on the landscape over large parts of North America4, Scandinavia5, and Britain6.

Landsat 7 ETM+ satellite image of Byrd Glacier, an ice stream in West Antarctica. Ice flow is towards the top of the image. Note how flow converges into the main ice stream trunk. Also, note the sharp boundary between fast- and slow-flowing ice. Image: NASA.

Why are ice streams important?

Ice streams in Greenland and Antarctica are the main control on ice sheet mass balance and discharge to the world’s oceans7. Understanding how ice streams behave and change over time, therefore, is important for predicting and managing the impacts of future climate change.

But this is easier said than done…

Firstly, records of modern-day ice stream activity only cover the most recent ~50 years (the length of the satellite record), which is not enough to confidently predict how they may change in the future.

Secondly, it is almost impossible for glaciologists to study the processes that occur at ice stream beds – which control fast ice flow and, ultimately, ice stream discharge to the oceans1 – owing to the great thickness (up to ~3 kilometres) of ice sheets.

Fast-flowing ice streams (blue and white areas) drain the interior of the East and West Antarctic ice sheets, controlling ice sheet mass balance and discharge to the oceans. Image: Jonathan Bamber

Why study palaeo-ice streams?

Therefore, the landforms and sediments left behind by palaeo–ice streams in areas like North America, Scandinavia, and Britain, are very important.

Firstly, they allow glaciologists to study how ice streams have evolved over thousands to tens of thousands of years, through important stages, such as ice sheet build-up, at a glacial maximum, and during deglaciation2,3,8,9.

Secondly, the landform record offers a window into the processes that occurred at former ice stream beds, allowing researchers study how they flowed, shifted, turned ‘on’ and ‘off’, and interacted with the landscape2.

The palaeo–ice stream record, therefore, can be used to better understand how ice streams change over long timescales and under different climate conditions, in order to improve predictions of future ice sheet change.

The palaeo–ice stream landsystem

Ice streams have three important characteristics that are reflected in the landforms they create10,11,12. First, they flow very rapidly – orders of magnitude faster than a typical valley glacier13 – by a combination of internal deformation, sliding, and subglacial deformation1,10. Second, they have convergent onset zones1,10 (onset zones are areas where ice flow changes from slow- [sheet flow] to fast-moving [stream flow] at the head of an ice stream). Third, their lateral margins are very sharp1,10.

Characteristics of an ice stream (fast-flowing ice, a convergent onset, and sharp lateral margins) displayed at Byrd Glacier, West Antarctica. Image: NASA.

Fast ice flow

Mega-scale glacial lineations are the most striking landforms created by fast ice flow in palaeo–ice streams14,15. They are streamlined sediment ridges formed at the bed in the main ice stream trunk zone16. You can think of these landforms as ‘stretched’ out flutes or drumlins, as they are similar in shape, but much larger and more elongate14,15.

In size, mega-scale glacial lineations are between 10–100 kilometres long and 200–1300 metres wide11, making it difficult to identify them on the ground. Instead, they are most easily mapped from satellite images (see below). When viewed from space, it is also obvious that mega-scale glacial lineations are not isolated features, but occur together in large groups. Within these groups, they run parallel to one another over great distances11,14,15.

Mega-scale glacial lineations formed at the bed of the Duawnt Lake palaeo-ice stream in Canada (see ref. 23). Note how individual lineations are highly elongate and closely parallel each other. In this example, the palaeo-ice stream flowed from right to left. Image: Google Earth.

Convergent onset zones

Shorter subglacial bedforms, such as flutes and drumlins, form in palaeo–ice stream onset zones, where ice velocity is lower than in the ice stream trunk zone11,17. These landforms are arranged in a fan-like pattern that flows in toward (or converges on) a narrower corridor of fast-flow landforms that include mega-scale glacial lineations.

Convergent flow in the onset zone of the Transition Bay palaeo-ice stream, Arctic Canada (see ref. 17). Also, note how this ice stream flow path (white) crosscuts an older ice stream flow path (black). Image: Google Earth.

Sharp ice stream margins

In modern ice streams, shear zones – areas of intense deformation several kilometres wide, marked by crevassing at the ice-surface18 – develop at the margins of ice streams, where fast- and slow-moving ice meet19.

Surface crevasses in a shear zone at Recovery Glacier ice stream in East Antarctica Image: NASA.

Ice stream shear margin moraines are sediment ridges deposited subglacially in the shear zone20. At first glance, they look similar to mega-scale glacial lineations, but they are generally wider and longer20. Shear margin moraines can be used to identify the edges (and thus lateral extent) of palaeo–ice streams11,12.

Shear margin moraine (arrowed) with a fast-flow assemblage (e.g. drumlins, mega-scale glacial lineations) ‘inside’ the palaeo-ice stream flow path (right of shear moraine) and ice-stagnation landforms ‘outside’ the ice stream flow path (left of shear moraine). Example from the M’Clintock Channel palaeo-ice stream in Arctic Canada (see ref. 20). Image: Google Earth.

Flow-direction changes

Ice streams do not always follow the same flow pathway; they are capable of switching flow-direction over time owing to glaciological (e.g. ice thickness) or topographic (e.g. basin infilling) changes9,21.

In the palaeo–ice stream landsystem, flow-direction changes can be mapped where one group of flow assemblages (e.g. drumlins) crosscuts another11,12,14. It is usually possible to work out the relative order of flow changes by studying the pattern of crosscutting (see the Transition Bay palaeo-ice stream diagram above).

Ice stream shutdown

While the palaeo–ice stream landsystem is dominated by features relating to fast ice-flow (e.g. mega-scale glacial lineations), these may be overprinted by other landform assemblages. For example, during deglaciation, moraine ridges and ice-stagnation landforms may be deposited over the top of fast-flow landforms as the active ice-front moves back2,11,12.

Similarly, ribbed moraines (transverse sediment ridges) may form over the top of glacial lineations22. Ribbed moraines are thought to form where ice-flow changes from an extensional (ice streaming) to a compressional regime. Where they lie on top of glacial lineations, therefore, they may record the slowing or shutdown of palaeo-ice streams during ice sheet deglaciation22.

Ribbed moraines lying on top of glacial lineations at the bed of the former Dubawnt Lake palaeo-ice stream. This ordering of landform assemblages records ice stream shutdown during deglaciation (see ref. 22).

Summary

Ice streams shape the land surface they flow over, leaving behind a distinctive landsystem11 that includes mega-scale glacial lineations, which record the passage of fast-moving ice14, convergent bedforms in onset zones, and shear margin moraines that mark their sharp lateral margins20. In addition, the palaeo–ice stream landsystem often displays evidence of dynamic ice sheet changes5,6, such as switches in flow-direction9,21 (crosscutting landforms) and velocity.

Related content

Professor Chris Clark’s Sheffield University webpages also host a wealth of information on mega-scale glacial lineations, drumlins, and ribbed moraines!

References

1. Bennett, M.R., 2003. Ice streams as the arteries of an ice sheet: their mechanics, stability and significance. Earth-Science Reviews61, 309-339.

2. Stokes, C.R. and Clark, C.D., 2001. Palaeo-ice streams. Quaternary Science Reviews20, 1437-1457.

3. Livingstone, S.J., Cofaigh, C.Ó., Stokes, C.R., Hillenbrand, C.D., Vieli, A. and Jamieson, S.S., 2012. Antarctic palaeo-ice streams. Earth-Science Reviews111, 90-128.

4. Margold, M., Stokes, C.R., Clark, C.D. and Kleman, J., 2015. Ice streams in the Laurentide Ice Sheet: a new mapping inventory. Journal of Maps11, 380-395.

5. Kleman, J., Hättestrand, C., Borgström, I. and Stroeven, A., 1997. Fennoscandian palaeoglaciology reconstructed using a glacial geological inversion model. Journal of glaciology43, 283-299.

6. Hughes, A.L., Clark, C.D. and Jordan, C.J., 2014. Flow-pattern evolution of the last British Ice Sheet. Quaternary Science Reviews89, 148-168.

7. Rignot, E., Velicogna, I., van den Broeke, M.R., Monaghan, A. and Lenaerts, J.T., 2011. Acceleration of the contribution of the Greenland and Antarctic ice sheets to sea level rise. Geophysical Research Letters38 (5).

8. Stokes, C.R., Margold, M., Clark, C.D. and Tarasov, L., 2016. Ice stream activity scaled to ice sheet volume during Laurentide Ice Sheet deglaciation. Nature530, 322-326.

9. Ó Cofaigh, C., Evans, D.J. and Smith, I.R., 2010. Large-scale reorganization and sedimentation of terrestrial ice streams during late Wisconsinan Laurentide Ice Sheet deglaciation. GSA Bulletin122, 743-756.

10. Clark, C.D., 1999. Glaciodynamic context of subglacial bedform generation and preservation. Annals of Glaciology28, 23-32.

11. Clark, C.D and Stokes, C.R. 2003. Palaeo-ice stream landsystem. In Evans, D.J.A. (Ed.) Glacial Landsystems. Hodder–Arnold, UK.

12. Stokes, C.R. and Clark, C.D., 1999. Geomorphological criteria for identifying Pleistocene ice streams. Annals of Glaciology28, 67-74.

13. Rignot, E., Mouginot, J. and Scheuchl, B., 2011. Ice flow of the Antarctic ice sheet. Science333, 1427-1430.

14. Clark, C.D., 1993. Mega‐scale glacial lineations and cross‐cutting ice‐flow landforms. Earth Surface Processes and Landforms18, 1-29.

15. Stokes, C.R. and Clark, C.D., 2002. Are long subglacial bedforms indicative of fast ice flow? Boreas31, 239-249.

16. King, E.C., Hindmarsh, R.C. and Stokes, C.R., 2009. Formation of mega-scale glacial lineations observed beneath a West Antarctic ice stream. Nature Geoscience2, 585-588.

17. Angelis, H.D. and Kleman, J., 2008. Palaeo‐ice‐stream onsets: examples from the north‐eastern Laurentide Ice Sheet. Earth Surface Processes and Landforms, 33, 560-572.

18. Raymond, C., 1996. Shear margins in glaciers and ice sheets. Journal of Glaciology42, 90-102.

19. Schoof, C. 2004. On the mechanics of ice-stream shear margins. Journal of Glaciology50, 208-218.

20. Stokes, C.R. and Clark, C.D., 2002. Ice stream shear margin moraines. Earth Surface Processes and Landforms27, 547-558.

21. Winsborrow, M.C., Stokes, C.R. and Andreassen, K., 2012. Ice-stream flow switching during deglaciation of the southwestern Barents Sea. GSA Bulletin124, 275-290.

22. Stokes, C.R., Lian, O.B., Tulaczyk, S. and Clark, C.D., 2008. Superimposition of ribbed moraines on a palaeo‐ice‐stream bed: implications for ice stream dynamics and shutdown. Earth Surface Processes and Landforms33, 593-609.

23. Stokes, C.R. and Clark, C.D., 2003. The Dubawnt Lake palaeo‐ice stream: evidence for dynamic ice sheet behaviour on the Canadian Shield and insights regarding the controls on ice‐stream location and vigour. Boreas32, 263-279.

Introduction to glacial landsystems

What are glacial landsystems | The landsystems approach | Studying landform–sediment assemblages | Why are glacial landsystems useful? | Summary | Key terms | References

What are glacial landsystems?

Research in glacial geology has increasingly concentrated on glacial landsystems1,2. In broad terms, the landsystems concept attempts to understand how a landscape was created through investigation of the complete collection of landforms and sediments within it.

Formally, a landsystem can be defined as1:

An area with a common set of features that is different to that of neighbouring areas. This includes the surface topography (of which landforms are a part), but also its underlying sediments and soils, and overlying vegetation.

There are two key principals to the landsystem approach:

The first is that landforms (such as an esker, moraine, or roche moutonnée) and sediments are considered not in isolation, but in combination as landform–sediment assemblages that make up a landscape (see “Key terms” at bottom of page).

The second is that, in landsystems research, the emphasis is on linking landform–sediment assemblages to the processes that create them – to produce process–form models – which, when applied to a particular landscape, can even be linked to local environmental (e.g. climate) and geological (e.g. topography) controls.

The glacial tongue and foreland of Skaftafellsjökull glacier in Iceland, with assemblages of closely-spaced sawtooth push moraines. Iceland has been an important testing ground for the study of process–form models at active glacier margins (see Evans, 2003; ref. 1) Photo: Chensiyuan

The landsystems approach

For the landsystems approach to be effective, glacial geologists must carefully study the glacial landform–sediment assemblages of two settings:

One – they investigate the landforms and sediments being actively created in currently glaciated areas (such as Iceland, the European Alps, or the Southern Alps of New Zealand) to establish clear links with the glaciological processes that produce them, and;

The partially debris-covered terminus of the Fox Glacier, South Island, New Zealand, in 2013. Here, a glacial geologist could study the process–form relationships of a temperate valley glacier with a high supraglacial debris load. Photo: M. Basler

Two – they study the landforms and sediments of areas where glaciers are no longer present (such as the British Isles) to reconstruct past glacial systems and the processes that operated within them.

The cirque floor of Cwm Idwal, North Wales, with a chain of moraine ridges flanking Llyn (lake) Idwal and a staircase sequence of lateral moraines descending the opposite valley side (right of image). This is a good example of landform–sediment assemblages produced by a small cirque glacier during the Loch Lomond Stadial in upland Britain. Photo: J. Bendle.

From this, you should see that the accuracy of the landsystem approach – and the reconstruction of former glacier systems – relies on a clear understanding of the processes that create specific landform–sediment assemblages at active glaciers – i.e. on process–form relationships.

Studying landform–sediment assemblages

In practice, studying process–form relationships at active glaciers, and the application of the landsystem approach to a formerly glaciated landscape, involve broadly the same method: a detailed physical inspection of the landscape to identify landform–sediment assemblages and patterns in their spatial distribution.

This is typically achieved by mapping the type, size, and shape of landforms from aerial and satellite imagery, or through fieldwork (this can even be achieved using Google Earth imagery, making it an interesting and accessible topic for A-level and undergraduate research projects).

Google Earth image of the Skálafellsjökull glacier margin and immediate foreland (Iceland). Glacial geologists use satellite images such as this to map the type and distribution of landform–sediment assemblages.

In addition, the sedimentological characteristics of landform-assemblages – for example, the size, shape, and roundness of particles in a moraine – are recorded in the field or analysed in a laboratory. This extra information tells us much about how a feature was formed, such as whether the sediment was laid down directly by ice or by glacial meltwater2,3.

Exposure in the ‘Little Ice Age’ moraine of the Exploradores Glacier, Patagonia (South America). Glacial geologists carefully record the constituents (e.g. the size, roundness and shape of particles) of landforms such as moraines to better understand how they were created. Photo: J. Bendle

Why are glacial landsystems useful?

One of the main advantages of the landsystems approach is the ability to reconstruct – not just the size and shape of former glaciers – but the distinct characteristics of these glaciers and the processes that once operated there, more accurately and in much greater detail than is possible by studying individual landforms in isolation1.

One example would be the ability to determine the thermal regime of a former glacier based on the suite of landform-sediment assemblages it left behind. Active temperate (warm-based) glaciers, for example, are often associated with low-amplitude push, dump and squeeze moraines, flutes, drumlins, and glaciofluvial features, such as eskers and outwash plains4,5. These assemblages develop under wet-based conditions that encourage basal sliding and subglacial deformation. Cold-based glaciers, by contrast, which are frozen to their beds, produce different landform–sediment assemblages typified by ice-contact fans, thrust-block moraines, and periglacial screes, with no subglacial features like flutes and drumlins6,7.

Google Earth image of the Svínafellsjökull glacier (Iceland) showing the landform–sediment assemblages typical of an active temperate glacier margin, such as flutes and debris stripes, sawtooth push moraines, and outwash deposits (see ref. 8)

Another advantage of the landsystem approach occurs where distinctive landform-sediment assemblages overprint or overlap one another, as they contain evidence of temporal changes in glacier processes or characteristics. One example relates to assemblages of cross-cutting drumlins (and/or other subglacially moulded landforms) or bedrock striations, which indicate changes in the direction of ice flow over time.

Summary

The landsystems approach is a holistic method of studying glacier and landscape history that (i) makes inferences using the full suite of landform–sediment assemblages that constitute a landscape, and (ii) is supported by process–form models established at active glaciers.

Other pages in this section of the website give examples of the main glacial landsystems to be identified in both actively and formerly glaciated areas.

Key terms

Landform-sediment assemblage: a distinctive group of landforms and sediments that together reflect a common process or age. A glaciated landscape is typically made up of lots of distinctive landform–sediment assemblages related to different (e.g. subglacial, supraglacial, ice-marginal) processes.

Process–form model: a theoretical model (based on detailed observation) that links physical processes to the landforms and sediments they create. In glaciated systems, processes of glacier erosion, ice and debris transfer, and deposition create landform–sediment assemblages.

References

[1] Evans, D.J.A., 2003. Glacial landsystems. Hodder–Arnold.

[2] Benn, D.I., and Evans, D.J.A., 2014. Glaciers and glaciation. Routledge.

[3] Evans, D.J.A., and Benn, D.I., 2004. The practical guide to the study of glacial sediments. Hodder–Arnold

[4] Evans, D.J.A., and Twigg, D.R., 2002. The active temperate glacial landsystem: a model based on Breiðamerkurjökull and Fjallsjökull, Iceland. Quaternary Science Reviews21 (20-22), 2143-2177.

[5] Evans, D.J.A., 2003. Ice-marginal terrestrial landsystems: active temperate glacier margins (ed.) in Evans, D.J.A., Glacial Landsystems. Hodder–Arnold, pp. 89-110.

[6] Fitzsimons, S.J. 2003. Ice-marginal terrestrial landsystems: polar continental glacier margins (ed.) in Evans, D.J.A., Glacial Landsystems. Hodder–Arnold, pp. 89-110.

[7] Hambrey, M.J., and Fitzsimons, S.J., 2010. Development of sediment–landform associations at cold glacier margins, Dry Valleys, Antarctica. Sedimentology57 (3), 857-882.

[8] Evans, D.J.A., Ewertowski, M.W., and Orton, C., 2019. The glacial landsystem of Hoffellsjökull, SE Iceland: contrasting geomorphological signatures of active temperate glacier recession driven by ice lobe and bed morphology. Geografiska Annaler: Series A, Physical Geography101 (3), 249-276.

Types of glaciers

Earth’s glaciers are incredibly varied in their size and shape, ranging from small ice masses that cling precariously to steep mountain sides, to vast ice sheets that submerge entire continents below kilometres thick ice1,2.

The form, shape and structure – known as the morphology – of these two extreme examples, as well as all glacier types in between, is a function of two key variables: climate and topography.

Climate

Climate controls the annual temperature cycle of a region as well as the amount of precipitation that falls as snow. Because of this, climate governs the annual mass balance of glaciers and hence their size (a key part of glacier morphology).

Where climatic conditions lead to mass inputs (e.g. snowfall) that are larger than mass outputs (e.g. melting) a glacier will grow. Conversely, where mass outputs exceed mass inputs a glacier will shrink.

Regular and heavy snowfall over Monte San Valentín (4058 m) and the glacier accumulation zone(s) of the North Patagonian Icefield contribute to regional mass balance. Photo: M. Foubister

All other factors being equal, therefore, it follows that the coldest places on Earth’s surface, the polar regions, will contain the largest and most extensive glaciers. However, climate is only one part of the story.

Topography

Topography is also a major control on glacier morphology. Topography not only provides the land surface (e.g. high altitude mountains) on which glacial ice can develop, but it also controls the physical dimensions of glaciers and how they flow.

The Alps mountains above Chamonix, France, not only rise to a high enough altitude that glaciers can exist there due to the cold conditions, but the very steep slopes running away from Mont Blanc (top left) dictate the form and flow of Glacier des Bossons (front) and Glacier de Tacconaz (behind). Photo: S. Räsänen

Consider this example. A deep valley that is several kilometres in length will contain a thicker and longer ice mass than a small mountain cirque. Because of its greater thickness, this hypothetical valley glacier will, in turn, flow more rapidly because thicker ice increases driving stresses at the glacier bed and raises basal temperatures3, which increase the rate of ice deformation and basal sliding.

Types of glacier

Bearing in mind the combined influences of climate and topography in shaping glacier morphology, the broad range of glacier types at Earth’s surface fit into two main groups, known as unconstrained glaciers and constrained glaciers1,2, which are defined as follows:

  • Unconstrained glaciers have a morphology and flow pattern that is in the most part independent of underlying topography, whereas;
  • Constrained glaciers have a morphology and flow pattern that is strongly dependent on underlying topography.

Unconstrained glaciers

Ice sheets and ice caps

Ice sheets and ice caps take the same basic form, having a broad, upstanding, and slowly moving ice dome at their centre, with channels of faster moving ice that transfer mass to their margins.

Surface elevation maps of the Greenland and Antarctic ice sheets, showing their dome-like structure (IPCC, AR5)

However, they differ in terms of scale2. Ice sheets are larger, being more than 50,000 km2 in size, with ice domes that can be more than 3000 m thick. In contrast, ice caps only reach thicknesses of several hundred metres.

GoogleEarth image of the Vatnajökull ice cap, Iceland, with a central ice dome drained by valley glaciers along its southern margin. The snowline is marked by the boundary between bare ice (grey and black) and snow (white).

There are two more key features of ice sheet and ice cap morphology. Firstly, they almost completely submerge the landscape, with only the tips of mountain peaks (known as nunataks) piercing the ice surface.

Starr nunatak rising above the ice surface in the Victoria Land region of the Antarctic Ice Sheet. Photo: S. Bannister

Secondly, their flow patterns are (at least in the most part) unaffected by underlying topography. The exception to this general rule are the fast-flowing ice streams and outlet glaciers that often reside within glacial troughs closer to the periphery of ice sheets and ice caps4,5.

Ice streams

Ice streams are corridors of rapidly moving ice in an ice sheet4. A feature unique to ice streams is that they are bordered on either side not by bedrock, but by slowly moving ice.

The crevassed surface of the Recovery ice stream that drains part of the East Antarctic Ice Sheet. Photo: NASA

Ice streams are extremely large (>20 km wide and >150 km long) and when viewed from space, we observe that they are fed by numerous tributaries that are connected to a central ice dome6,7. Ice streams are critically important to the overall dynamics and mass balance of ice sheets as they control the vast majority (~90% in Antarctica) of ice and sediment discharge to the oceans4,6,7.

Ice streams of Antarctica draining the ice sheet interior. From: Rignot et al. (2011)

Constrained glaciers

Ice fields

Unlike ice caps, ice fields do not have a simple dome-like structure. Instead, their morphology and flow are controlled by topography. Ice fields (such as the Patagonian ice fields) develop in mountainous terrain where the land surface reaches an altitude that enables snow and ice to accumulate. They are drained by large valley glaciers.

The North Patagonian Icefield of southern South America with its numerous radiating valley glaciers. Image: NASA

Valley glaciers

Valley glaciers (as their name suggests) exist within bedrock valleys and are overlooked by ice-free slopes. They are found in many alpine and high mountain environments, including the European Alps, Southern Alps of New Zealand, the Andes, and the Himalayas (to name just a few).

The Aletsch Glacier (or ‘Aletschgletscher’ in German) in the Swiss Alps is a classic example of a valley glacier. Note in the upper reaches that the glacier has several cirque basin tributaries that feed the main glacier trunk. Photo: D. Beyer

Valley glaciers are fed in their upper parts by ice and snow discharged from surrounding ice fields or cirques (see the Aletsch Glacier above) in addition to snow and ice avalanches from overlooking slopes. In terms of morphology, valley glaciers can be single features or made up of a branching network of tributaries (see image below), and range in length from several kilometres to over 100 kilometres.

Large valley glaciers in Alaska (USA) seen from space by the Sentinel-2 satellite. Notice the
tributary glaciers feeding the main trunk of Columbia Glacier (centre).

Transection glaciers

Transection glaciers are, in essence, a system of interconnected valley glaciers that flow in several different directions, often in a radiating (or web-like) pattern. Transection glacier networks develop where bedrock valleys are deeply dissected, allowing ice to overflow the cols between adjacent valleys.

GoogleEarth image of the Spitsbergen island of the Svalbard archipelago, showing transection glaciers.

Examples of active transection glaciers can be found in Greenland, Svalbard (see above), and Alaska. Such systems also developed during the last glacial period in the European Alps8, and parts of the Loch Lomond Stadial ice cap in Scotland are also thought to have formed transection glacier networks9.

Piedmont glaciers

Piedmont glaciers have a distinctive form characterised by large terminal ice lobes that splay outwards onto lowland terrain after exiting a confining bedrock valley. Topography, therefore, exerts varying degrees of control on piedmont glacier morphology and flow at different points along the glacier length.

The Agassiz (left) and Malaspina (right) piedmont glacier lobes spilling out from the St. Elias mountains, Alaska, on to flat coastal plains. Image: NASA

Another common feature is that large areas of a piedmont glacier are situated below the equilibrium line altitude in the ablation zone. The Malaspina Glacier in Alaska (see image above) is the most famous example of a piedmont glacier. This glacier, which drains the Mt. St. Elias ice field, has a terminal lobe that is around 40 km long and almost 65 km across at its widest.

Cirque glaciers

Cirque glaciers are among the most common types of glacier on Earth, being found in nearly all alpine landscapes that support ice accumulation. Cirque glaciers are either localised to armchair-shaped bedrock hollows on a mountain side (see image below), or to the uppermost parts of a glacial trough, where they flow into larger valley glaciers.

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

The morphology of a cirque glacier largely depends on the topography in which it sits. The cirque basin itself dictates the size and shape of the cirque glacier and directs its flow, while the terrain surrounding a cirque basin is an important source of wind-blown snow and therefore glacier mass balance10.

Niche glaciers

Smaller in size than cirque glaciers, niche glaciers form where ice accumulates in a mountain side recess (or niche), such as a rock bench, couloir, or depression. Niche glaciers represent the early stages of glacier development and are commonly found in climatically favourable settings, such as in shaded north-facing slopes of mountains in the Northern Hemisphere11. Similar to niche glaciers, but adhering to steep mountain sides, are ice aprons.

Niche glacier occupying a bedrock recess at the summit of Blick vom Gatschkopf (2945 m) in the Austrian Alps. Photo: Kogo

Activities

Using GoogleEarth (or similar) explore Earth’s mountain regions and, using the definitions and images in this article, try to identify examples of unconstrained and constrained glacier types. While doing this, think about the possible climatic and topographic factors that control the size, shape, and flow of glaciers.

You may also like to compare the size of different glacier types, as well as other physical metrics such as ice surface gradient. You can do this by experimenting with the “Measure Tool” in GoogleEarth, which enables you to measure distance and area.

References

[1] Sugden, D.E., John, B.S., 1976. Glaciers and Landscape: Arnold.

[2] Benn, D.I., Evans, D.J.A., 2010. Glaciers and Glaciation, 2nd edition: Routledge.

[3] Cuffey, K.M. and W.S.B. Paterson, 2010. The Physics of Glaciers, 4th edition: Academic Press.

[4] Bennett, M.R., 2003. Ice streams as the arteries of an ice sheet: their mechanics, stability and significance. Earth-Science Reviews, 61, 309-339.

[5] Winsborrow, M.C.M., Clark, C.D., Stokes., C.R. 2010. What controls the location of ice streams? Earth-Science Reviews, 103, 45-59.

[6] Joughin, I., Smith, B.E., Howat, I.M., Scambos, T., Moon, T., 2010. Greenland flow variability from ice-sheet-wide velocity mapping. Journal of Glaciology, 56, 415-430.

[7] Rignot, E., Mouginot, J., Scheuchl, B., 2011. Ice flow of the Antarctic ice sheet. Science, 333, 1427-1430.

[8] Wirsig, C., Zasadni, J., Christl, M., Akçar, N., Ivy-Ochs, S., 2016. Dating the onset of LGM ice surface lowering in the High Alps. Quaternary Science Reviews, 143, 37-50.

[9] Golledge, N.R., Hubbard, A., 2005. Evaluating Younger Dryas glacier reconstructions in part of the western Scottish Highlands: a combined empirical and theoretical approach. Boreas, 34, 274-286.

[10] Lie, Ø., Dahl, S.O., Nesje, A., 2003. A theoretical approach to glacier equilibrium-line altitudes using meteorological data and glacier mass-balance records from southern Norway. The Holocene, 13, 365-372.

[11] Harrison, S., Knight, J., Rowan, A.V., 2015. The southernmost Quaternary niche glacier system in Great Britain. Journal of Quaternary Science, 30, 325-334.