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.
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.
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.
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.
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.
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.
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.
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.
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. Boreas38, 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. Boreas33, 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,
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,
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.
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.
Below, there is more information about our map and database.
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.
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 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.
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.
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.
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.
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.
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.
After the Last Glacial Maximum
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.
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
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.
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.
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
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.
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
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.
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
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.
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
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.
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.
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.
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
How was the reconstruction
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.
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.
we mapped 58,823 landforms, including 25,009 moraines, 2,507 shorelines, 3,926
lineations, 4,309 empty cirques and 4,536 palaeochannels.
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
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.
An offshore glaciomarine landsystem,
with fjords, offshore moraine ridges, drumlins, raised fluvial deltas and slope
failures, and turbidity current channels.
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.
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
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
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.
The PATICE Team
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. Change162, 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. Change146, (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. Geomorphology204, 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).
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.
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 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.
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).
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
 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.
 Benn, D.I., and Evans, D.J.A., 2010. Glaciers and Glaciation. Hodder-Arnold, London.
 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.
 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. Boreas, 11, 225-247.
 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 Geography, 71, 17-29.
 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 Geology, 91, 215-227.
 Eyles, N. 1983. The glaciated valley landsystem. In Eyles, N. (ed.) Glacial Geology. Pergamon, Oxford, 91–110.
 Ballantyne, C.K., 1997. Periglacial trimlines in the Scottish Highlands. Quaternary International, 38, 119-136.
 Ballantyne, C.K. 2007. Trimlines and palaeonunataks. In Elias, S.A. (ed.), Encyclopedia of Quaternary Science. Elsevier, Oxford, 892–903.
 Evans, D.J., Ewertowski, M. and Orton, C., 2017. The glaciated valley landsystem of Morsárjökull, southeast Iceland. Journal of Maps, 13, 909-920.
 Maizels, J.K., 1995. Sediments and landforms of modern proglacial terrestrial environments. In Menzies, J. (ed.), Modern Glacial Environments. Butterworth-Heinemann, Oxford, 365–416.
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.
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.
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.
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).
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.
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.
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.
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.
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.
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.
 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 Reviews, 106, 14–28.
 Golledge, N.R. (2007) An ice
cap landsystem for palaeoglaciological reconstructions: characterizing the
Younger Dryas in western Scotland. Quaternary Science Reviews26,
 Golledge, N.R. (2010)
Glaciation of Scotland during the Younger Dryas stadial: a review. Journal
of Quaternary Science, 25, 550–566.
 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,
 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.
 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,
 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.
 Evans, I.S. 1977. World-wide
variations in the direction and concentration of cirque and glacier
aspects. Geografiska Annaler: Series A, Physical Geography, 59,
 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.
 Ballantyne, C.K. (2002) The
Loch Lomond Readvance on the Isle of Mull, Scotland: glacier reconstruction and
palaeoclimatic implications. Journal of Quaternary Science, 17,
 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.
 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,
 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.
 Carr, S.J. & Coleman, C.G.
(2007) An improved technique for the reconstruction of former glacier
mass-balance and dynamics. Geomorphology, 92, 76–90.
 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
 Ballantyne, C.K. (2002) A
general model of paraglacial landscape response. The Holocene, 12,
 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 lakes are a common feature of glaciated mountain ranges. They form wherever glacial ice blocks the drainage of rivers or meltwater. This includes:
a glacier blocks a trunk or tributary valley; and
a glacier fills an overdeepened valley created by glacial erosion
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
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.
 Teller, J.T., 1995. History and
drainage of large ice-dammed lakes along the Laurentide Ice Sheet. Quaternary
International, 28, 83-92.
 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. Boreas, 26,
 Teller, J.T., 2003. Subaquatic
landsystems: large proglacial lakes. In Evans, D.J.A. Glacial
Landsystems (pp. 348-371). Arnold London.
 Breckenridge, A., 2013. An
analysis of the late glacial lake levels within the western Lake Superior basin
based on digital elevation models. Quaternary Research, 80,
 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 Reviews, 203, 102-127.
 Broecker, W.S., 1966. Glacial
rebound and the deformation of the shorelines of proglacial lakes. Journal
of Geophysical Research, 71, 4777-4783.
 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
Bulletin, 106, 19-31.
 Østrem, G., Haakensen, N. and
Olsen, H.C., 2005. Sediment transport, delta growth and sedimentation in Lake
Nigardsvatn, Norway. Geografiska Annaler: Series A, Physical Geography, 87,
 Benn, D.I. and Evans, D.J.A.,
2010. Glaciers and Glaciation (pp. 570-573)Routledge, London.
 Nemec, W., 1990. Aspects of sediment
movement on steep delta slopes. In Coarse-grained deltas (Vol.
10, pp. 29-73).
 Smith, D.G. and Jol, H.M.,
1997. Radar structure of a Gilbert-type delta, Peyto Lake, Banff National Park,
Canada. Sedimentary Geology, 113, 195-209.
 Bell, C.M., 2009. Quaternary
lacustrine braid deltas on Lake General Carrera in southern Chile. Andean
Geology, 36, 51-66.
 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 Bulletin, 117,
 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 Research, 80,
 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 Publications, 53,
 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 Research, 63, 171-181.
 Ashley, G.M., 2002.
Glaciolacustrine environments. In Modern and past glacial environments (pp.
 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
Reviews, 226, 105964.
 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,
 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
Change, 162, 275-291.
 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 Reviews, 29, 1212-1227.
 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 Geography, 87(2), pp.363-374.
 Bell, C.M., 2008. Punctuated
drainage of an ice‐dammed
quaternary lake in southern south america. Geografiska Annaler: Series
A, Physical Geography, 90(1), pp.1-17.
 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 Reports, 6, p.21064.
 Benito, G. and Thorndycraft,
V.R., 2019. Catastrophic glacial-lake outburst flooding of the Patagonian Ice
Sheet. Earth-Science Reviews, p.102996.
 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 Reviews, 177, 314-339.
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.
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).
What controls valley
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.
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.
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’
Debris supply to
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.
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’.
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.
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.
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
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.
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
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’.
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Kirkbride, M.P., Owen, L.A. and Brazier, V., 2003. Glaciated valley
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Science Reviews, 27(25-26), 2361-2389.
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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 Holocene, 15(1),
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insights into glacier dynamics and climatic controls. Boreas, 41(3),
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