Glacial erratics

What is a glacial erratic?

Glacial erratics, often simply called erratics, or erratic boulders, are rocks that have been transported by ice and deposited elsewhere. The type of rock (lithology) that the glacial erratic is made from is different to the lithology of the bedrock where the erratic is deposited.

For example, an erratic could be a boulder of sandstone is picked up by a glacier, transported, and deposited on top of a limestone bedrock. Some erratics are useful to scientists because they are of a distinctive rock type, which means that their source outcrop can be identified and located. Glacial erratics are therefore useful in reconstructing past glacier flow directions, the timing of glacier retreat, and even the type of glacier flow.

Where do glacial erratics come from?

As a glacier or ice sheet moves, it can erode bedrock. The ice can then pick up, or entrain, the eroded rock. As the ice flows, it transports the bedrock debris in the direction of flow. The ice then deposited the entrained sediment once it begins to retreat.

The process of formation of glacial erratics. A. Glacial erosion entrains a boulder of the bedrock. B. Continued glacier flow transports the boulder. C. As the glacier retreats, the boulder is deposited on a different type of bedrock, forming a glacial erratic.

Erratics can range from large boulders to smaller stones and pebbles. All erratics are of a different rock type. Glacial sediments often contain a range of rocks of different kinds, which can be used to reconstruct the ‘provenance’1, or source, of the sediment and therefore the direction of ice flow.

A granite boulder on sandstone bedrock on Alexander Island, Antarctic Peninsula. There is no granite of this kind on Alexander Island so it was transported by the glacier.

Rocks that are moved by the glacier but are of the same rock type are called ‘glacially-transported’ rocks. All glacially-transported rocks and erratics tend to show evidence of that glacial transport, with scratches (striations), rounded edges and polished faces.

Carboniferous limestone boulder embedded within glacial till at Whitburn Bay, County Durham. Note the faceted shape and scratches (striations). The boulder is shaped by glacial erosion and scratched by contact with other rocks and ice. See: Davies et al., (2009)2

Glacial erratics and glacially-transported rocks can be sourced from rocks falling onto the glacier, rocks being picked up and transported at the base of the glacier, and rocks plucked from valley sides. Rocks transported on the glacier surface are said to be ‘supraglacial’, whilst rocks transported at the base of the ice are ‘subglacially’ transported.

What do glacial erratics tell us about past ice sheets?

The first thing erratics can tell us about past ice sheets is the direction of ice movement. If you find an erratic with a distinctive lithology, you can trace it back to the location where the distinctive bedrock is found.

A good example of this indicator lithology in England is the Shap Granite from Cumbria. Boulders of Shap Granite are found throughout Cumbria, County Durham, North Yorkshire and as far southeast as Bridlington on the Yorkshire Coast2,3. The example shown in the figure below is from Goldsborough Carr in County Durham, which is 40 km east of the Shap Granite.

Examples of erratics from Goldsborough Carr (left) and Assynt (right). Photographs by A. Emery.

Erratics can tell us when the ice sheet retreated, by cosmogenic dating of the boulder. Assuming the boulder was eroded at the base of the glacier, the exposure age given by the cosmogenic dating will tell us when the boulder was deposited by the retreating glacier4. The boulder of Shap Granite in the figure above was deposited by the retreating Eden-Stainmore Ice Stream approximately 19,750 years ago5.

Sampling for cosmogenic nuclide exposure-age dating on an erratic boulder on a moraine in Patagonia.

We can also learn the style of ice-sheet flow from how glacial erratics are grouped. Long lines of glacial erratics are known as dispersal trains. These dispersal trains can show whether flow was focussed into ice streams or as part of a regional, sustained flow6. Boothia-type dispersal trains show that flow over an indicator lithology was focussed into an ice stream, named after the Boothia Peninsula in Arctic Canada. Dubawnt-type dispersal trains have little change in width, which shows that regional, unconstrained flow was active over the indicator lithology.

Styles of glacial erratic boulder dispersal. Left: Boothia-type dispersal. Right: Dubawnt-type dispersal. After Dyke & Morris (1988)6.

Where can you find glacial erratics in the UK?

There are many famous examples of glacial erratics in the UK. These erratics have captured the imagination of amateur and professional geologists for centuries. In 1928, the Yorkshire Geological Society published the work of Frederic Harmer7. This map collated the studies of the Yorkshire Boulder Committee and many similar groups.

As you can see below, the map shows the huge density of glacial erratics in the UK. The Norber erratics in the Yorkshire Dales, near Austwick, Settle, are famous and scenic examples of erratics. More examples of erratics are the Great Stone of Fourstones on the Lancashire/Yorkshire border, and Cloughmore in County Down, Northern Ireland.

The map of glacial erratics and their sources in England and Wales7.
Reproduced from Harmer 1928, Proceedings of the Yorkshire Geological Society, Vol. 21, 79-150, by permission of the Yorkshire Geological Society.

Scotland is full of glacial erratics thanks to its diverse bedrock geology. We use these erratics to reconstruct the dynamics of the British-Irish Ice Sheet. The figure below shows Dubawnt-style dispersal trains in the Assynt region of Scotland. The outcrops of Torridonian Sandstone and the trains of dispersed erratics show that ice flowed towards the west-northwest. The constant width of the dispersal trains shows that the regional flow of ice was a constant velocity over this area8.

Glacial erratic boulder dispersal trains in Assynt. Based on data from Lawson (1995)8. Elevation data: OS Terrain 5


1.          Evans, D. J. A. & Benn, D. I. A practical guide to the study of glacial sediments. (Arnold, 2004).

2.          Davies, B. J. et al. Interlobate ice-sheet dynamics during the last glacial maximum at Whitburn Bay, County Durham, England. Boreas 38, 555–578 (2009).

3.          Clark, C. D. et al. Map and GIS database of glacial landforms and features related to the last British Ice Sheet. Boreas 33, 359–375 (2004).

4.          Raistrick, A. THE GLACIATION OF WENSLEYDALE, SWALEDALE, AND ADJOINING PARTS OF THE PENNINES. Proc. Yorksh. Geol. Soc. 20, 366–410 (1926).

5.          Vincent, P. J., Wilson, P., Lord, T. C., Schnabel, C. & Wilcken, K. M. Cosmogenic isotope (36Cl) surface exposure dating of the Norber erratics, Yorkshire Dales: Further constraints on the timing of the LGM deglaciation in Britain. Proc. Geol. Assoc. 121, 24–31 (2010).

6.          Davies, B. J. et al. Dynamic ice stream retreat in the central sector of the last British-Irish Ice Sheet. Quat. Sci. Rev. 225, 105989 (2019).

7.          Dyke, A. S. & Morris, T. F. DRUMLIN FIELDS, DISPERSAL TRAINS, and ICE STREAMS IN ARCTIC CANADA. Can. Geogr. Géographe Can. 32, 86–90 (1988).

8.          Harmer, F. W. THE DISTRIBUTION OF ERRATICS AND DRIFT. Proc. Yorksh. Geol. Soc. 21, 79–150 (1928).

9.          Lawson, T. J. Boulder Trains as indicators of former ice flow in Assynt, N.W. Scotland. Quat. Newsl. 75, 15–21 (1995).

PATICE interactive map

In this interactive map, you can explore all the glacial landforms and chronologies that were used to generate the new reconstructions of the last Patagonian Ice Sheet from 35,000 years ago to the present day.

Click on the PATICE logo to launch the online interactive map.

The GIF below shows the extent of the ice sheet in 5000 year timeslices. The colours around the margin show where we have high, medium and low confidence in where we have placed the margin.

The image below shows the evidence used to create the reconstruction. You can explore this data yourself using our interactive map that uses ArcGIS Online.


PATICE: The Patagonian Ice Sheet from 35,000 years ago to the present day


This page provides a brand-new reconstruction of the Patagonian Ice Sheet from 35,000 years ago to the present day (called PATICE).

PATICE is a new compilation of published ages and geomorphology, ranked and assessed and recalibrated, which we use to generate new empirical reconstructions of the ice sheet and its ice-dammed palaeolakes.

Citation: Davies et al., 2020. The Evolution of the Patagonian Ice Sheet from 35 ka to the Present Day (PATICE). Earth-Science Reviews. 

You can explore the PATICE data in the ArcGIS online interactive online map (no ArcGIS licence needed).


All the data used in the new PATICE reconstruction has been made available fully open-access, and the paper is available as gold open access.

Below, there is more information about our map and database.

PATICE reconstruction

The Patagonian Ice Sheet formed during the last glaciation along the Andean mountain chain. It blocked the drainage of rivers to the Pacific, so large lakes formed in front of the ice margin as it receded. You can watch the ice sheet separate out into its different parts and the lakes draining and changing through time in the GIF below.

PATICE reconstruction, 35,000 years ago to the present day

The Patagonian Icefields

Patagonia, in southernmost South America, is a region with accelerating glacier recession1,2. Glaciers here are shrinking rapidly, which is enlarging glacial lakes, increasing flood risk3 and causing sea level rise4.

The Southern Andes region lost 1,208 billion tonnes of glacier ice from 1961 to 20165, contributing 0.92 ± 039 mm per year to global sea level rise (27 mm from 1961 to 2016).

Today, there are four main icefields. The Northern Patagonian Icefield (46.4°S to 47.5°S), the Southern Patagonian Icefield (48.3°S to 52°S), the Gran Campo Nevado (52.8°S) and the Cordilleran Darwin Icefield (54.5°S)6.

The glaciers and icefields of Patagonia. Copyright J. Bendle

The region also has numerous small glaciers and icefields, often centred on volcanoes. The lowest latitude Southern Hemisphere glacier that reaches the ocean is found in the Northern Patagonian Icefield (Glaciar San Rafael).

In total, the present-day icefields and glaciers cover a total area of 22,718 km2, equating to a sea level equivalent of 15.1 mm7. These glaciers are all rapidly receding, as you can see in the GIF below.

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

Reconstructing past glacier change in Patagonia

We can use the past record of ice sheet behavior to understand how these glaciers interact with climate, and to better predict how they might behave in the future.

Patagonia has a rich record of glacial geomorphology that can help us to understand how glaciers might behave in the future8. In this study, we created a new GIS database of published glacial geomorphology and published ages to reconstruct the extent of the ice sheet from 35,000 years ago to the present day, in 5000-year time slices.

Patagonian piedmont lobes in the Chilean Lake District. Moraine geomorphology by Glasser and Jansson (2008) and PATICE.

We also reconstructed the evolution of ice-dammed lakes. As the glacier ice shrank back, large and deep lakes formed in front of the ice margin. These lakes existed until the glacier ice shrank back past a col or spillway, allowing the lake to drain. These lakes may have been a key influence on the glaciers, as deep water would have encouraged the calving of icebergs.

Ice dammed lakes form when the normal drainage is blocked by glaciers.

The past behavior of the Patagonian Ice Sheet during different climate states and during rapid climate transitions could shed insights into ways in which the region is sensitive to changes and how it could respond to future change. The southern mid-latitudes are a particularly data-sparse region of the globe, and reconstructions of the Patagonian Ice Sheet provides a unique insight into the past terrestrial glacial and climatic change.

Last Glacial Maximum

The Patagonian Ice Sheet was centred on the central mountain chain of the Andes, and stretched from 38° to 56°S. During glacial maxima, the icefields coalesced to form a single large ice mass9,10.

Its eastern margin comprised fast-flowing lobes of ice that extended out into the Argentinian steppe landscape, the largest fastest of which were ice streams. The western margin reached the continental shelf in the Pacific Ocean.

The Local Last Glacial Maximum occurred at 33,000 to 28,000 years ago from 38°S to 48°S, and earlier, at around 47,000 years ago from 48°S southwards.

PATICE Reconstruction of the Patagonian Ice Sheet at the Local Last Glacial Maximum (35,000 years ago)

At its maximum extent, the Patagonian Ice Sheet covered 492,600 km2, with a sea level equivalent of 1,496 mm (the sea level amount locked up in the ice sheet). It was 350 km long and 2090 km wide.

It was comparable in size to the Antarctic Peninsula Ice Sheet today. For comparison, Sweden is 450,295 km2 and the UK is 242,495 km2.

Comparison of the Patagonian Ice Sheet at 35,000 years ago and the Antarctic Peninsula Ice Sheet today. Both maps are to the same scale.

After the Last Glacial Maximum

The Patagonian Ice Sheet began to retreat and shrink by 25,000 years ago. The ice sheet stabilized and formed large moraines during the period 21,000 to 18,000 years ago, which was then followed by rapid deglaciation, especially between 18,000 and 15,000 years.

By 15,000 years ago, the Patagonian Ice Sheet had separated into several disparate ice masses, draining into large ice-dammed lakes along its eastern margin. These lakes probably encouraged the calving of icebergs, and facilitated rapid melting.

Glacial readvances or stabilisations occurred at least at 14,000 to 13,000 years ago, 11,000 years ago, 6000 to 5000 years ago, 2000 to 1000 years ago, and 500 to 200 years ago.

Glacier lakes

On the eastern side of the Andes, the moraines from these glacier outlet lobes often mark the present-day continental watershed drainage divide. Today, many of the glacial lakes of Patagonia drain westwards, into the Pacific Ocean. The Patagonian Ice Sheet dammed this drainage route, forcing higher lake levels and drainage to the Atlantic Ocean.

During glacier recession, a series of large proglacial lakes formed along the eastern margin, dammed between the ice sheet and higher ground or moraines11–15. As the ice sheet receded, continental scale drainage reversals occurred14,16,17.

In the GIF above, you can see the glacial lakes forming as the ice sheet retreats back towards the high ground of the Andes. By about 10,000 years ago, most lakes have either drained completely or dropped to their current levels.

Evolution of glaciers and lakes from 16,000 years ago to 5,000 years ago around the Northern Patagonian Icefield. The glacier lakes (pink) change as the ice dams recede past particular cols or spillways.

This figure shows how the ice dammed glacial lakes evolved as the glacier ice receded. As the glaciers shrank, new cols and spillways became available, resulting in the lakes dropping to a new level.

Rates of recession

The Patagonian Ice Sheet was relatively stable from 35,000 to 30,000 years ago. Recession from the Local Last Glacial Maximum began by 25,000 years ago, predating the global Last Glacial Maximum. This may be because the Patagonian Ice Sheet was smaller and more dynamic than the larger global ice sheets.

Very rapid recession and widespread deglaciation began after 18,000 years ago, during a period of rapid warming highlighted in the Antarctic ice cores18 and rapid global sea level rise. It may also have been driven by a southwards shift in the Southern Westerly Winds.

The Patagonian Ice Sheet probably contributed ~615 mm to global sea level rise between 20,000 and 15,000 years ago, when it shrank from 359,600 to 121,800 km2.

Glaciers stabilized or re-advanced during the Antarctic Cold Reversal, but the Patagonian Ice Sheet was much smaller at this time (116,700 km2).

Accelerating rates of glacier recession

Rates of recession of the Patagonian Ice Sheet form 35,000 years ago to the present day

Rates of recession were slow through the Holocene until the last few decades (with the caveat that some time periods are highly uncertain). Absolute recession rates (km2 per year; km2 a-1) over recent decades rival those seem between 20,000 and 15,000 years ago for an ice sheet that was more than two orders of magnitude larger.

Relative rates of recession (percentage change per year; % a-1) are higher between 200 years ago and 2011 AD than at any time observed in our reconstruction. It is likely that there were periods of time with especially rapid recession during the last glacial-interglacial transition, when many outlet lobes were calving into large, ice-dammed lakes, but our compilation is unable to capture this.

There are fewer degrees of freedom for ice extent and volume changes during the Holocene (last 10,000 years). Ice margins stabilized not far from present-day positions by the Early Holocene, and dated moraines suggest that readvances were similar to the advance at 200 years ago in size.

Thus we can argue that average rates of ice-marginal recession are currently faster than at any time observed in the Holocene, in line with the recent temperature changes observed in Antarctica and Patagonia, following a sustained period of relative stability, and when glacial lake area remains fairly constant.

Since observations indicate that rates of recession have accelerated in Patagonia over recent decades, from 34.2 km2 a-1 (0.14 % a-1) for 1986 to 2001 AD to 51.2 km2 a-1 (0.22 % a-1) for 2001-2011 AD6, this is especially concerning.

Ice-dammed palaeolakes

Glacial lake area peaked at 13,000 years ago, with an estimated area of 13,999 km2. These large lakes would have accelerated glacial recession at this time by encouraging the outlet glacial lobes to calve icebergs.

Glacial lake area minimum was reached by about 10,000 years ago. Rapid recession from 13,000 to 10,000 years ago led to many cols and spillways opening, and the lakes reached their current spatial extent and volume by about 10,000 years ago.

As the ice dams receded, this cold, fresh water may have been released suddenly into the Pacific Ocean, possibly affecting regional climate14,16.

The remaining lake water today, from lakes that were in the footprint of the palaeolakes, is 6,824 km2. Overall, between the maximum lake extent at 13,000 years ago and today, there has been a reduction in lake area of 7,176 km2.

Area of glacial lakes for each time slice (ka: thousands of years ago)

How was the reconstruction made?

There is a large volume of published ages and glacial geomorphology that help us to reconstruct the past ice sheet extent and dynamics in Patagonia. The geomorphological data provide information on former ice sheet margins, ice-dammed palaeolake evolution, and ice-flow direction. Our new GIS database includes 58,823 landforms and 1,669 published ages.

Published ages and geomorphology used in PATICE

Glacial Geomorphology

Landforms were mapped from satellite images and digital elevation models, most commonly LANDSAT 7 and LANDSAT 8 images, as well as high resolution satellite imagery available in Google Earth.

Our compilation includes moraines, trimlines, glacial lineations (bedrock and sedimentary drumlins or flutes), meltwater palaeochannels, outwash plains, shorelines, deltas and cirques. Our compiled maps also show related topographic landforms such as rivers, lakes and volcanoes.

Numbers of different landforms used in PATICE

In total, we mapped 58,823 landforms, including 25,009 moraines, 2,507 shorelines, 3,926 lineations, 4,309 empty cirques and 4,536 palaeochannels.

Moraines and trimlines give us information about former ice sheet margins, whilst lineations and the pattern of moraines tells us about ice-flow direction. Cirques tell us about regions that were previously glaciated, but now are ice-free. Palaeochannels, outwash plains, shorelines and deltas give us information on ice-dammed palaeolakes.

There is more information on Patagonian Glacial Geomorphology here.  In the view below, you can use Google Earth to explore Patagonian moraines around the North Patagonian Icefield. The arcuate ridges denote the position of the former ice margin.

In general, there are four distinct temperate glacial landsystems in Patagonia.

  1. An upland glacier landsystem, with an assemblage of cirques, lateral and terminal moraines, mountain glaciers and snow patches, flutes, and lakes;
  2. In the lowlands, a land-terminating glacial landsystem, with moraine arcs, outwash plains, meltwater channels, drumlins, and hummocky moraine;
  3. A lowlands glaciolacustrine landsystem, with deltas and shorelines, and ice-contact glaciofluvial landforms;
  4. An offshore glaciomarine landsystem, with fjords, offshore moraine ridges, drumlins, raised fluvial deltas and slope failures, and turbidity current channels.


There are many ways to fix in time particular glacial landforms. Each technique dates something slightly differently, which makes them hard to compare directly. Our GIS database includes 1,669 ages relevant to the timing of deglaciation. Each age is scrupulously checked and recalibrated according to the latest protocols.

Number of ages used in each different dating technique

Our database includes cosmogenic nuclide exposure-age dating of boulders, ideally situated on moraines to give a time of formation for that moraine, radiocarbon dating of organic material, tephrochronology (dating of volcanic ash layers), lichenometry (measuring the size of specific species of lichens to derive an exposure age), dendrochronology (tree-ring dating), historical sources (archival maps and photographs), varve ages (annually laminated lake sediments), and optically stimulated luminescence.

Bethan Davies sampling a boulder for cosmogenic nuclide exposure age dating in Patagonia.

Each age is assessed according to our protocols and given a reliability assessment. The most reliable cosmogenic nuclide ages are used to give an average age for moraine formation.

Reconstructing the ice extent

We used the moraines, dated by various methods, to reconstruct the ice margin. At places where we were confident of the age of the moraines, we could draw short, isolated isochrones.

Secondly, we interpolated between the isochrones to reconstruct overall ice-sheet limits, using moraines and topography.

Thirdly, we provided an assessment of our degree of confidence in each ice margin, from high confidence to medium confidence and low confidence. High confidence ice limits have both well defined glacial geomorphology and a well constrained chronology.

Medium confidence ice limits are defined by geomorphology and are near to published ages, but are less well constrained.

Low confidence limits have no well-defined geomorphology, lie far from published ages, and are first interpretations that require further investigation.

Mapping ice margins and assigning levels of uncertainty. Example from the Southern Patagonian Icefield.


Authors are in alphabetical order.

Bethan J. Davies, Christopher M. Darvill, Harold Lovell, Jacob M. Bendle, Julian A. Dowdeswell, Derek Fabel, Juan-Luis García, Alessa Geiger, Neil F. Glasser, Delia M. Gheorghiu, Stephan Harrison, Andrew S. Hein, Michael R. Kaplan, Julian R.V. Martin, Monika Mendelova, Adrian Palmer, Mauri Pelto, Ángel Rodés, Esteban A. Sagredo, Rachel Smedley, John L. Smellie, Varyl R. Thorndycraft. 


1. Braun, M. H. et al. Constraining glacier elevation and mass changes in South America. Nat. Clim. Chang. 1 (2019).

2. Meier, W. J.-H., Grießinger, J., Hochreuther, P. & Braun, M. H. An updated multi-temporal glacier inventory for the Patagonian Andes with changes between the Little Ice Age and 2016. Front. Earth Sci. 6, 1–21 (2018).

3. Wilson, R. et al. Glacial lakes of the Central and Patagonian Andes. Glob. Planet. Change 162, 275–291 (2018).

4. Malz, P. et al. Elevation and mass changes of the Southern Patagonia Icefield derived from TanDEM-X and SRTM data. Remote Sens. 10, 188 (2018).

5. Zemp, M. et al. Global glacier mass changes and their contributions to sea-level rise from 1961 to 2016. Nature
568, 382–386 (2019).

6.Davies, B. J. & Glasser, N. F. Accelerating shrinkage of Patagonian glaciers from the Little Ice Age (AD 1870) to 2011. J. Glaciol. 58, (2012).

7. Carrivick, J. L., Davies, B. J., James, W. H. M., Quincey, D. J. & Glasser, N. F. Distributed ice thickness and glacier volume in southern South America. Glob. Planet. Change 146, (2016).

8. Coronato, A. & Rabassa, J. Chapter 51 – Pleistocene Glaciations in Southern Patagonia and Tierra del Fuego. in Developments
in Quaternary Sciences
(eds. Jürgen Ehlers, P. L. G. & Philip, D. H.) Volume 15, 715–727 (Elsevier, 2011).

9. Caldenius, C. C. Las glaciaciones cuaternarias en la Patagonia y Tierra del Fuego. Geogr. Ann. 14, 1–164 (1932).

10. Mercer, J. H. Variations of some Patagonian glaciers since the Late-Glacial. Am. J. Sci. 266, 91–109 (1968).

11. García, J.-L., Strelin, J. A., Vega, R. M., Hall, B. L. & Stern, C. R. Deglacial ice-marginal glaciolacustrine environments and structural moraine building in Torres del Paine, Chilean southern Patagonia. Andean Geol. 42, 190–212 (2015).

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

13. Turner, K. J., Fogwill, C. J., McCulloch, R. D. & Sugden, D. E. Deglaciation of the eastern flank of the North Patagonian Icefield and associated continental-scale lake diversions. Geogr. Ann. Ser. A, Phys. Geogr. 87, 363–374 (2005).

14. Thorndycraft, V. R. et al. Glacial lake evolution and Atlantic-Pacific drainage reversals during deglaciation of the Patagonian Ice Sheet. Quat. Sci. Rev. 203, 102–127 (2019).

15. Davies, B. J., Thorndycraft, V. R., Fabel, D. & Martin, J. R. V. Asynchronous glacier dynamics during the Antarctic Cold Reversal in central Patagonia. Quat. Sci. Rev. 200, (2018).

16. Glasser, N. F. et al. Glacial lake drainage in Patagonia (13-8 kyr) and response of the adjacent Pacific Ocean. Sci. Rep. 6, 21064 (2016).

17. García, J.-L. et al. Early deglaciation and paleolake history of Río Cisnes Glacier, Patagonian Ice Sheet (44 S). Quat. Res. 91, 194–217 (2019).

18. Cuffey, K. M. et al. Deglacial temperature history of West Antarctica. Proc. Natl. Acad. Sci. 113, 14249–14254 (2016).

British-Irish Ice Sheet

In this section, we discuss the evidence for the last British-Irish Ice Sheet (BIIS). The BIIS was at its maximum around 27,000 years ago and stretched from the continental shelf on its northern margin to South Wales, north Norfolk and the Vale of York.

Further reading:

The British-Irish Ice Sheet: an introduction

What is the British-Irish Ice Sheet? | The Devensian British-Irish Ice Sheet | How do we know? | What’s in a name? | Activities | References | Comments

What is the British-Irish Ice Sheet?

The British-Irish Ice Sheet is a name given to ice sheets that covered Britain and Ireland at different times during the Quaternary Period. Evidence for at least three major ice sheets is preserved in the sedimentary record on land in Britain and Ireland. These were the Anglian (between 478,000 and 424,000 years ago), the Wolstonian (between 300,000 and 130,000 years ago) and the Devensian (approximately 27,000 years ago, during the Last Glacial Maximum)1. This article deals with the Devensian ice sheet.

The Devensian British-Irish Ice Sheet

The Devensian British-Irish Ice Sheet was a large mass of ice that covered approximately two thirds of Britain and Ireland around 27,000 years ago2. All of Scotland and Ireland, most of Wales, and most of the north of England was underneath the ice sheet during the Last Glacial Maximum. This ice sheet retreated and shrank after 27,000 years ago, and had completely disappeared by 11,300 years ago3.

The maximum extent of the Devensian British-Irish Ice Sheet, compiled from Sejrup et al., 20059, Bradwell et al., 200810 and Sejrup et al., 201611. Bathymetry data are from the GEBCO 2016 dataset ( and land elevation from the European Environment Agency EU-DEM (

The retreat of the British-Irish Ice Sheet was not constant across the entire ice sheet. Different sectors of the ice sheet retreated at different rates, due to different processes affecting the ice sheet margins. Ice margins in contact with the ocean retreated earlier and quicker than ice margins on land2. There was also a period of ice sheet regrowth, known as a readvance, during a period known as the Younger Dryas, also called the Loch Lomond stadial4.

How do we know?

The Devensian ice sheet is the best-understood of the past British-Irish Ice Sheets. Because it is the most recent, evidence for the Devensian ice sheet is well preserved. There is a strong imprint of this ice sheet in the glacial geology of Britain. Many glacial landforms record the story of movement of ice. In the mountains, erosional landforms, such as cirques and roches moutonées, are common. In the lowlands, the ice sheet has left behind sediment deposits, such as tills, and depositional landforms, such as drumlins and moraines.

Drumlins in Ribblesdale
Drumlins formed by the British-Irish Ice Sheet in Ribblesdale, Yorkshire. Their orientation shows ice flowed down the valley. Elevation data: OS Terrain 5 (

The British-Irish Ice Sheet has been studied for nearly 200 years5. This level of study means it is well understood. Knowing how this ice sheet behaved under a warming climate after the Last Glacial Maximum is important for understanding how present-day ice sheets will change in the future6.

What’s in a name?

The British-Irish Ice Sheet has been called many things in the past. Other names you might see are the British Ice Sheet5, the British Isles Ice Sheet7, and the Celtic Ice Sheet8. You might also see the Devensian Stage being referred to as the Late Pleistocene, the Weichselian Stage, or the Late Glacial. In Ireland, the ice sheet is known as the Midlandian stage, because it was historically thought to terminate in the Irish Midlands. Thanks Sam Roberson for that additional information!


Was the land your town is built on under the British-Irish Ice Sheet at the Last Glacial Maximum? Download the .zip file below and import the .kml into Google Earth. Explore which hills and mountain ranges the ice sheet covered. Look at the relationship of the British-Irish Ice Sheet with the water depth in the Atlantic Ocean. Think about areas that were not covered- why not?


1. Ehlers, J. & Gibbard, P. L. Quaternary Glaciations – Extent and Chronology: Part I – Europe. vol. 2 (Elsevier, 2004).

2. Clark, C. D., Hughes, A. L. C., Greenwood, S. L., Jordan, C. & Sejrup, H. P. Pattern and timing of retreat of the last British-Irish Ice Sheet. Quat. Sci. Rev. 44, 112–146 (2012).

3. Small, D. & Fabel, D. Was Scotland deglaciated during the Younger Dryas? Quat. Sci. Rev. 145, 259–263 (2016).

4. Bickerdike, H. L., Evans, D. J. A., Stokes, C. R. & Ó Cofaigh, C. The glacial geomorphology of the Loch Lomond (Younger Dryas) Stadial in Britain: a review. J. Quat. Sci. 33, 1–54 (2018).

5. Clark, C. D. et al. Map and GIS database of glacial landforms and features related to the last British Ice Sheet. Boreas 33, 359–375 (2004).

6. Gandy, N. et al. Marine Ice Sheet Instability and Ice Shelf Buttressing Influenced Deglaciation of the Minch Ice Stream, Northwest Scotland. Cryosph. Discuss. 1–24 (2018) doi:10.5194/tc-2018-116.

7. Boulton, G. & Hagdorn, M. Glaciology of the British Isles Ice Sheet during the last glacial cycle: form, flow, streams and lobes. Quat. Sci. Rev. 25, 3359–3390 (2006).

8. Hughes, A. L. C., Gyllencreutz, R., Lohne, Ø. S., Mangerud, J. & Svendsen, J. I. The last Eurasian ice sheets – a chronological database and time-slice reconstruction, DATED-1. Boreas 45, 1–45 (2016).

9. Sejrup, H. P. et al. Pleistocene glacial history of the NW European continental margin. Mar. Pet. Geol. 22, 1111–1129 (2005).

10. Bradwell, T. et al. The northern sector of the last British Ice Sheet: Maximum extent and demise. Earth-Science Rev. 88, 207–226 (2008).

11. Sejrup, H. P., Clark, C. D. & Hjelstuen, B. O. Rapid ice sheet retreat triggered by ice stream debuttressing: Evidence from the North Sea. Geology 44, 355–358 (2016).

Landsystem of ‘clean’ valley glaciers

Glaciers that carry little to no rock or sediment debris at their surface are known as ‘clean’ or ‘uncovered’ glaciers1.

Truly ‘clean’ valley glaciers represent an ideal end member in the range of valley glacier types, which differ as a result of local basin topography, the amount of surface debris they carry, mass balance, and the amount of meltwater they produce (visit this page for more detail).

Many glaciers may only be partially uncovered, or fall somewhere between being truly clean and completely debris-covered. Therefore, the landsystem of any individual valley glacier may contain only some of the characteristic features of clean glaciers, as well as other landforms not always seen at uncovered glaciers.

The ‘clean’ (or ‘uncovered’) surface of Engasbreen in western Norway. Source: A. Toth.

Dynamics of clean glaciers

With limited debris at the surface, clean valley glaciers tend to respond quickly to shifts in climate, where changes in glacier volume (due to mass loss or gain) cause oscillations (advance and retreat) of the terminus1,2.

The active nature of clean valley glaciers can be explained by the role of debris cover on energy exchange and melting at the glacier surface3. Thick debris cover forms a barrier between the glacier and the atmosphere that insulates the ice surface from melting. Where the debris layer is absent or sparse, on the other hand, a glacier and the atmosphere are able to interact freely, and there is more energy available for melting at the ice surface.

For this reason, clean glaciers react rapidly to atmospheric conditions (e.g. cooling or warming) by gaining or losing mass, and then advancing or retreating1,2. As we will see later, this behaviour has a large impact on their interaction with the landscape.

The surface of Nigardsbreen in western Norway with only sparse debris cover. Source: J. Bendle.

Where are clean valley glaciers found?

Clean valley glaciers are found in most glaciated mountain ranges, although they are more common in low- to moderate relief mountains, and in areas of hard (to erode) bedrock, where the debris supply from valley side mass movements is minimal1.

Excellent examples of clean valley glaciers can be found in the Scandinavian mountains of western Norway and Sweden, in the coastal mountains of British Columbia and the Canadian Rockies, the Andes of Patagonia, and the European Alps.

Saskatchewan Glacier, a clean valley glacier of the Columbia Icefield in the Canadian Rockies of Alberta. Source: B. Menetrier

Landforms of clean valley glaciers

Together with the large-scale features of glaciated valleys2 – which may include U-shaped valley profiles, arêtes, hanging valleys, and ribbon lakes – clean valley glaciers are known to produce a distinctive suite of landforms and sediments1. The most typical are outlined below.

U-shaped valley, with Nigardsbreen in the distance. Source: Sissssou
Glaciated hanging valleys in the Jotunheimen National Park, western Norway. Source: J. Bendle.

Ice-marginal landforms

Latero-frontal moraines are formed at the outer limit of clean valley glaciers1. Largely, they are the result of ice pushing and the squeezing of waterlogged sediments from beneath the ice margin, with few dumps of material from the ice surface4-6. Much of the material that makes up the moraines formed by clean valley glaciers, therefore, derives from subglacial erosion (evidenced by the presence of faceted, striated, and sub-angular rocks in moraine deposits), or is picked up from the foreland during glacier advance4-6.

The active nature of most clean valley glaciers – meaning that they oscillate readily in response to changes in climate – often leads to the formation a large number of moraine ridges on the valley floor, with each ridge representing a period of glacier stability1,2. The size and spacing of these recessional moraine ridges give some indication of the duration and frequency of glacier stillstands during retreat7.

Google Earth image of arcuate recessional moraines that represent periods of temporary stability of Styggedalsbreen, Jotunheimen National Park, Norway.

Typically, the latero-frontal moraines formed by clean valley glaciers are relatively small; often, they are less than 10 metres high1,2. This is result of low supraglacial debris supply that limits the amount of material for moraine formation4-6, oscillating snouts that spread debris across large areas of the valley floor2, and the role of meltwater, which may flush out large volumes of debris from beneath valley glaciers and transport it away from the ice margin in proglacial streams.

Low-relief recessional moraine ridge on the foreland of Nigardsbreen in western Norway. Glacier flow was from right to left. Note the partly rounded cobbles, likely to derive from subglacial erosion, or perhaps reworking of outwash deposits during glacier advance. Source: J. Bendle.

While generally consistent in size, lateral moraines – those formed at the glacier sides – are sometimes larger than those deposited in the valley centre because of the increased supply of debris from valley walls, in the form of rockfalls, landslides, and slumps1,2,4,5,6. Similarly, variations in catchment geology (e.g. areas of weak or strong bedrock) or glacier dynamics may lead to variations in debris supply and, therefore, moraine volume1,2.

Large lateral moraine (left of image) supplied by valley side mass movement at Vadret da Tschierva, Bernina Range, Switzerland. Source: H. Krapf


The former thickness of valley glaciers can often be identified using trimlines on the valley side2,8,9. These trimlines, which mark the upper limit of recent glacial erosion on the valley wall, can be identified by the contrast in vegetation cover on either side of the limit, with bare rock or pioneer vegetation found below the trimline where glacial erosion has occurred, and well-vegetated or forested slopes above it.

The upper extent of valley glaciers may also be inferred from the boundary between frost weathered debris (e.g. talus) above and the ice-moulded bedrock below. This is known as a periglacial trimline8,9.

Google Earth image of an erosional trimline at the lateral margin of the Colonia Glacier from the North Patagonian Icefield in southern South America. Note the sharp distinction between bare, ice-moulded bedrock below the trimline, which marks out the recent thickness of the glacier, and the vegetated slopes above.

Subglacial landforms

Subglacial erosional processes are active at most clean valley glaciers. Zones of ice-moulded bedrock, roches moutonnées, whalebacks, and striations, for example, are often seen emerging from beneath retreating glacier snouts, providing evidence for abrasion and quarrying of the bed by warm-based and sliding ice1,2.

Top: Ice-moulded bedrock emerging from beneath the retreating Nigardsbreen. Bottom left shows a whaleback. Bottom right shows striated and chattermarked bedrock. Source: J. Bendle.

In addition to erosional landforms, the beds of clean valley glaciers may contain flutings (flutes)10, usually on the valley floor between recessional moraines, as well as other streamlined deposits that fill in the hollows behind lumps and bumps in the bedrock and trail away in a downglacier direction. These latter features are known as lee-side cavity fills2.

Google Earth image of the foreland of Maradalsbreen in Jotunheimen National Park, Norway, showing fluted terrain (i.e. streamlined sediment ridges oriented in the direction of former glacier flow) and recessional moraines.

Glaciofluvial landforms

Glaciofluvial landforms and sediments are common in the proglacial zone of clean valley glaciers, particularly those occupying maritime mountain ranges where the amount of glacial meltwater produced each year is high1,2,11. At valley glaciers in dry, arid climates, on the other hand, glaciofluvial features are less abundant.

High melt-season proglacial discharge at Bergsetbreen (top left) in western Norway. Source: J. Bendle.

In valley glacier systems, the movement of meltwater streams is restricted by the valley walls. Therefore, it is common for outwash (sandur) deposits (also known as valley trains) to build up in the valley bottom, or in the low points between recessional moraines1,2,11.

Google Earth image of valley train (sandur) downstream of the Nef Glacier in Patagonia, southern South America. River flow is from left to right.

Where glaciofluvial processes are especially active, meltwater streams may completely rework moraines and other glacial deposits, so that only minor traces of former terminus positions remain9. Over time, river terraces may form in glaciated valleys as a result of fluvial incision into valley fill deposits1,2,11.

Valley train (sandur) downstream of Glacier des Bossons (top of photo) that has largely eroded valley bottom moraines. Source: J. Bendle.

The landsystem of clean valley glaciers

The unique features of clean valley glaciers – i.e. that they carry limited surface debris, that most debris they transport comes from erosion of the bed, and that their snouts fluctuate in response mass loss or gain – leads to the formation of a distinct set of landforms1,2.

In summary, this includes: numerous low-relief moraine ridges crossing the valley floor (occasionally with larger lateral moraines where rock debris falls from the valley walls); areas of ice-moulded bedrock with roches moutonnées, whalebacks, and striations; and the outwash deposits (e.g. valley trains) of proglacial streams.

Spatial pattern of landforms

At some (but not all) clean valley glaciers, the landforms formed by clean valley glaciers are organised into several zones2. This includes an inner erosional zone in the upper valley, where ice-moulded bedrock is extensive; an intermediate zone characterised by both bed erosion and deposition; and an outer depositional zone in the lower reaches of a glaciated valley, where moraines and outwash deposits are most extensive.

This arrangement of landforms, which is not restricted to valley glaciers, reflects the downglacier transport of debris from areas of net erosion higher up in the valley towards the ice margin.

The landsystem of a clean valley glacier at Nigardsbreen, western Norway. Note the presence of an inner erosional zone (containing ice-moulded bedrock) on the steep terrain close to the present-day glacier margin, and an outer depositional zone (containing recessional moraines and outwash deposits) on the low gradient valley bottom. Image: Google Earth.


[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. Hodder-Arnold, London.

[3] Nakawo, M. and Young, G.J. 1981. Field experiments to determine the effect of a debris layer on ablation of glacier ice. Annals of Glaciology 2, 85–91.

[4] Matthews, J.A. and Petch, J.R., 1982. Within‐valley asymmetry and related problems of Neoglacial lateral moraine development at certain Jotunheimen glaciers, southern Norway. Boreas11, 225-247.

[5] Shakesby, R.A., 1989. Variability in Neoglacial moraine morphology and composition, Storbreen, Jotunheimen, Norway: within-moraine patterns and their implications. Geografiska Annaler: Series A, Physical Geography71, 17-29.

[6] Benn, D.I. and Ballantyne, C.K., 1994. Reconstructing the transport history of glacigenic sediments: a new approach based on the co-variance of clast form indices. Sedimentary Geology91, 215-227.

[7] Eyles, N. 1983. The glaciated valley landsystem. In Eyles, N. (ed.) Glacial Geology. Pergamon, Oxford, 91–110.

[8] Ballantyne, C.K., 1997. Periglacial trimlines in the Scottish Highlands. Quaternary International38, 119-136.

[9] Ballantyne, C.K. 2007. Trimlines and palaeonunataks. In Elias, S.A. (ed.), Encyclopedia of Quaternary Science. Elsevier, Oxford, 892–903.

[10] Evans, D.J., Ewertowski, M. and Orton, C., 2017. The glaciated valley landsystem of Morsárjökull, southeast Iceland. Journal of Maps13, 909-920.

[11] Maizels, J.K., 1995. Sediments and landforms of modern proglacial terrestrial environments. In Menzies, J. (ed.), Modern Glacial Environments. Butterworth-Heinemann, Oxford, 365–416.

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


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.


[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.


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 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 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.


[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 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’.


[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.