Different types of trimlines exist, dependent on the length of time since the last glacial advance . In recently glaciated valleys, they are marked by a distinct change in vegetation. Above the trimline, dense vegetation exists with species characteristic of the region[2,3]. And below the trimline, valley sides are characteristic of bare, ice-scoured rock, or with early stages of vegetational development [2,3]. These are termed ‘Vegetational Trimlines‘, and are commonly associated with glacial activity since the Little Ice Age, within the historical era [2,3].
For older glaciations, occurring prior to the Little Ice Age, the vegetation change may be less distinctive as vegetation succession has occurred [2,4]. Therefore, these trimlines mark the boundary between the smooth, ice-scoured bedrock below the trimline, and the frost-shattered regolith from periglacial weathering above . These can be termed ‘Peri-glacial Trimlines’ [2,3,5].
What are Trimlines Used For?
Unlike other glacial landforms which show the lateral extent, or behaviour of a past glaciations, trimlines mark the maximum vertical extent of the ice surface [3,4]. This enables the production of 3-D model reconstructions of ice sheets and valley glaciers[4,6].
In currently occupied glacial regions, trimlines can be compared to the modern ice surface elevation to assess the role of ice surface thinning, and responses to climate change .
 McCarroll D. (2014) Trimline. In: Encyclopedia of Planetary Landforms. Springer, New York, NY. doi.org/10.1007/978-1-4614-9213-9_383-1
 Benn, D.I., and Evans, D.J.A., 2010. Glaciers and Glaciation. Hodder-Arnold, London
 Rootes, C.M., and Clark, C.D. (2020) Glacial trimlines to identify former ice margins and subglacial thermal boundaries: A review and classification scheme for trimline expression. Earth-Science Reviews. 210. 103355.
 McCarroll, D. (2016) Trimline Trauma: The wider implications of a paradigm shift in recognising and interpreting glacial limits. Scottish Geographical Journal. 132(2).
 Ballantyne, C. K. (1997). Periglacial Trimline in the Scottish Highlands. Quaternary
International, 38, 119–136.
 Ballantyne, C. K. (2010). Extent and Deglacial Chronology of the Last British‐Irish Ice Sheet: Implications of Exposure Dating Using Cosmogenic Isotopes. Journal of Quaternary Science, 25(4), 515–534.
 Kohler, J., James, T., Murray, T., Nuth, C., Brandt, O., Barrand, N., Aas, H. & Luckman, A.
(2007). Acceleration in Thinning Rate on Western Svalbard Glaciers. Geophysical Research
Glacial flutes are elongated, low-relief ridges formed subglacially and orientated in the direction of glacier flow [1,2,3]. Their size can range between several centimetres to a few metres both wide and high, and occur in groups of streamlined ridges known as ‘swarms’ .
Flutes are formed subglacially and are found in glacial foregrounds. They are more likely to be found in modern glacier foreground as they can be subjected to erosion because of their till-like composition .
Flutes are found in a variety of glaciated regions including Iceland, Sweden, Norway, New Zealand, and Alaska [1,2,3]. Because of their relatively small size, they are often hard to identify whilst at ground level. Therefore high-resolution satellite data or LiDAR methods are used to map them .
Flutes can be formed subglacially beneath both polythermal, and warm-based glaciers. There have been several models proposed about the formation of flutes, but the most widely accepted model is the Cavity Infill Model [1,2,4].
This model proposes that a boulder causes an obstruction beneath the actively flowing glacier. The glacier then forces highly saturated sediment into a cavity on the leeside of the boulder obstruction [1,2]. The pressure on the leeside of the boulder drops, allowing the saturated sediment to freeze, and is carried forward by the ice, forming the elongated flute shape [1,3].
 Benn, D.I., and Evans, D.J.A., 2010. Glaciers and Glaciation. Hodder-Arnold, London
 Ely, J.C., Graham, C., Barr, J.D., Rea, B.R., Spagnolo, M., and Evans, J. (2016) Using UAV acquired photography and structure from motion techniques for studying glacier landforms: application to the glacial flutes at Isfallsglaciären. Earth Surface Processes and Landforms. DOI: 10.1002/esp.4044.
 Gordon, J.E., Whalley, W.B., Gellatly, A.F., Vere, D.M. (1992) The formation of glacial flutes: Assessment of models with evidence from Lyngsdalen, North Norway. QSR. 13(7). PP. 709-731.
 Boulton, G.S. (1976) The origin of glacially fluted surfaces – observations and theory. Journal of Glaciology. 17. PP. 287-309.
Tidewater glaciers are glaciers which extend out, and terminate into the sea . They are part of a group of glaciers known as calving glaciers, as their main method of ice loss is through iceberg calving, instead of surface melt [1,2]. Calving icebergs currently accounts for up to 70% of the worlds annual mass transfer from glacial regions to the ocean .
Tidewater glaciers are found at latitudes of 45ᵒ and above, and are present in different glacial regions including Antarctica, Alaska, Greenland, Svalbard, and Patagonia [1,2].
Types of Tidewater Glacier
Mountain glaciers terminating into the ocean are called ‘tidewater glaciers’. ‘Tidewater outlet glaciers’ are glaciers which reach the ocean through fjords, branching off from ice caps, ice sheets or icefields .
Tidewater glaciers can either be grounded – where the glacier is in constant contact with the bed. Or they can be floating – when the terminus is floating on the sea water, or flowing into an ice shelf [1,3]. Grounded glaciers tend to be located in temperate regions such as Alaska, or Canada. And floating tidewater glaciers are commonly found in polar regions, namely Greenland, Svalbard, and Antarctica.
Tidewater Glaciers and Iceberg Calving
Calving icebergs are the most efficient method of losing mass from a glacier . It is the dominant cause of mass loss from the Antarctic Ice Sheet , therefore, it is important to understand the process behind these calving events [1,4]
Iceberg calving occurs when there are faults in the glacier, known as crevasses. Crevasses can form when there stress and strain thresholds are reached on the glacier. The trigger for the iceberg calving events vary for both grounded and floating tidewater glaciers.
Floating tidewater glacier
For a floating tidewater glacier, submarine melting of the underside of the glacier causes a direct loss of ice, as well as undercutting the floating glacier terminus or ice shelf . This causes instability, resulting in complete collapse [1,6].
When the floating section of the tidewater is removed, the ice on the land is no longer supported from the buttressing ice shelf. It is then able to rapidly, and continuously calve icebergs. For example, the Larsen B ice shelf collapse in 2002 on the Antarctic Peninsula .
Grounded tidewater glaciers
Grounded tidewater glaciers calve when there is either a rapid thinning of the glacier surface, or a localised change in sea level. This change forces the glacier terminus to be out of equilibrium with the ocean, resulting in the terminus to be lifted and detached from the bed, causing the terminus to become buoyant . During this process, the crevasses are able to isolate large blocks of ice which are then calved into icebergs.
 Vieli, A., 2011. Tidewater glaciers, in: Singh, V.P., Singh, P., Haritashya, U.K. (Eds.), Encyclopedia of Snow, Ice and Glaciers. Springer, pp. 1175–1179.
 Benn, D.I., Hulton, N.R.J., Mottram, R.H., 2007. “Calving laws”, “sliding laws” and the stability of tidewater glaciers, in: Sharp, M. (Ed.), Annals of Glaciology, Vol 46, 2007, Annals of Glaciology. Int Glaciological Soc, Univ Ctr Svalbard UNIS, NO-9171 Longyearbyen, Norway. Benn, DI, Univ Ctr Svalbard UNIS, Box 156, NO-9171 Longyearbyen, Norway., pp. 123–130.
 van der Veen, C.J., 2002. Calving glaciers. Prog. Phys. Geogr. 26, 96–122. https://doi.org/10.1191/0309133302pp327ra
 Benn, D.I., and Evans, D.J.A., 2010. Glaciers and Glaciation. Hodder-Arnold, London
 Shepherd, A., Ivins, E., Rignot, E., Smith, B., Van Den Broeke, M., Velicogna, I., Whitehouse, P., Briggs, K., Joughin, I., Krinner, G., IMBIE Team, 2018. Mass balance of the Antarctic Ice Sheet from 1992 to 2017. Nature 558, 219–222. https://doi.org/10.1038/s41586-018-0179-y
 Benn, D.I., Astrom, J., Zwinger, T., Todd, J., Nick, F.M., Cook, S., Hulton, N.R., and Luckman, A. (2017) Melt under-cutting and buoyancy-driven calving from tidewater glaciers: new insights from discrete element and continuum model simulations. Journal of Glaciology. 63(240).
During the Younger Dryas, the Lake District was covered by plateau icefields and cirque glaciers. The image below shows the larger plateau icefields (green) and the smaller cirque glaciers (red) in the Lake District and Snowdonia.
You can explore all of the locations in this page using the Younger Dryas Glacial Map. This version of the map is focused on the Lake District.
These ice masses left behind numerous moraines as they retreated. The image below (credit Dr Richard Waller) shows some 360º imagery of a fabulous set of hummocky moraines in Greenup Gill, near Borrowdale, in the Central Lake District. These glacial landforms show the retreat of a small plateau icefield outlet glacier during the Loch Lomond Stadial[1,3].
The moraines show the retreat of the glacier all the way up onto the plateau, showing the active retreat of this plateau icefield. The continuous moraine sequences shows that the outlet glacier retreated towards its plateau source area without becoming disconnected from the plateau icefield accumulation area.
This 360º image shows a lateral moraine (a moraine formed at the sides of the glacier) in Bannerdale (credit Dr Richard Waller). Bannerdale held a small cirque glacier during the Younger Dryas, with lateral moraines demarcating the glacier limits.
In the image below, the lateral moraine is visible as a linear mound of sediment against the valley side walls, in the immediate foreground. Some glacially transported boulders are visible on the ridge of the moraine.
Haweswater also held a plateau icefield during the Younger Dryas [1,3].
This is a set of moraines above Haweswater (credit Dr Richard Waller). The moraines track the recession of the plateau icefield outlet glaciers onto the upland areas .
You can explore these moraines yourself in Google Earth or in the Younger Dryas Glacial Map. The moraines are visible in the satellite imagery as rounded hummocks with scattered boulders.
The Younger Dryas Glacial Map shows the locations of these moraines, just at the head of Haweswater Reservoir.
Gillercomb, the valley just to the west of Seathwaite (Cumbria), preserves a number of glacier moraines deposited during the Younger Dryas.
The location can be explored in Google Maps. The moraines are visible as the smoothed, elongated mounds in the valley floor.
These moraines were formed during the recession of the plateau icefield that covered this part of Cumbria during the Younger Dryas .
A number of small elongate moraines exist in the bottom of the Derwent river valley, just south of Seathwaite. These moraines have a number of glacially transported boulders on their summits.
Between 12,900 and 11,700 years ago, gradual warming of Britain’s climate was interrupted by a sudden period of renewed cooling. During this period, known as the Loch Lomond or Younger Dryas Stadial, glaciers regrew in many areas of upland Britain.
Evidence of these glaciers is preserved in a range of different glacial landsystems in Britain. Even though these glaciers have long since disappeared, by studying the Younger Dryas glacial landsystems they left behind, we can understand what processes operated in these glacial environments.
The alpine icefield landsystem
The most widespread landsystem of the Loch Lomond Stadial is the alpine icefield, evidence of which is found throughout the mountainous areas of the Western Grampian Highlands of Scotland and on several of the Western Isles, including Skye and Mull.
This landsystem is a type of glaciated valley landsystem, usually consisting of a series of steep-sided, glacial valleys, separated by arêtes and spurs. The size and shape of these glaciers was strongly controlled by the topography, with ice confined to within the valleys.
In some places, ice from two or more separate valleys would join together over lower sections of the mountain ridges, called cols. This created networks of connected valley glaciers called icefields.
Landforms of the alpine icefield landsystem
The diagram below shows the types of landform usually found in the Loch Lomond Stadial alpine icefield landsystem in Britain. The numbered features are discussed below.
#1. Recessional moraines
The most widespread feature of this landsystem is sequences of recessional moraines, which are arranged in concentric ridges on the valley floors and lower slopes. Moraines are piles of debris, usually mud, sand, and boulders, all deposited in piles at the ice terminus. They are typically unsorted and chaotic.
These moraines formed during short phases of glacier advance and retreat that interrupted the general pattern of glacier retreat and are typical of active temperate glaciers.
#2. Moraine mounds
Sometimes there are small patches of more chaotically arranged moraine mounds (2) within these sequences. Areas of extensive moraines indicate that the Loch Lomond Stadial glaciers transported large volumes of debris.
Some of this debris likely fell onto the glacier surfaces from the surrounding valley slopes but it is also thought that the glaciers reworked large volumes of debris that was already present in the landscape.
In some places, eskers (3) are present on the valley floors, but these are less common. Eskers are ridges of sand and gravel, deposited by glacial meltwater flowing through tunnels within and underneath glaciers. After the glacier disappears, these sediments are left behind as a ridge in the landscape.
#4. Medial moraines
Similarly, medial moraines (4) may mark locations at the confluence of two valley glaciers, but evidence of these within this landsystem is rare. Medial moraines form where two glaciers met.
#5. Terminal moraines
In some valleys, particularly those with cirques at their heads, recessional moraines are only found in the area around the former glacier terminus (5).
In these valleys with cirques at their head, the upper valley might be covered within a thin blanket of till or show evidence of flutes (6). Flutes are streamlined ridges of sediment, sometimes with a boulder or obstacle at their head, that formed subglacially underneath temperate ice.
#7, #8. Erosional landforms
At the heads of these valleys, erosional glacial landforms can often be found. These can include roches moutonées (7), formed by abrasion and quarrying of the bedrock under the sliding glacier, and ice-smoothed bedrock (8).
In many areas, the height of the former glacier surface is marked by trimlines (9). These features show the height of the former glacier surface on the valley slopes.
Trimlines can be identified by the contrast between glacial landforms below the trimline (in the area covered by the former glacier), and evidence of frost-shattering and periglacial processes above the trimline (in areas that remained above the glacier surface).
Cast study: Isle of Mull Alpine Icefield
The Isle of Mull on the West coast of Scotland shows glacial geomorphology typical of the alpine icefield landsystem. The numbers on the map match with the features described above.
The Isle of Mull had an independent ice domes that deflected mainland ice around it during the Last Glacial Maximum. During the Younger Dryas, it was glaciated with an independent mountain icefield.
Ice drained from the broad uplands of Sgurr Dearg and the Beinn Talaidh-Corra-bheinn ridge to form the Ba and Forsa outlet glaciers to the northwest and north, respectively.
The lower slopes of these valleys are covered with nested lateral moraines, chains of recessional moraines and thick drift of glacial sediments. The terminus of the glaciers is obscured by glaciofluvial outwash sands and gravels.
There were six cirque glaciers around the margins of the icefield. They were not connected to the main icefield.
In summary, the Loch Lomond Stadial alpine icefield landsystem is found in upland areas of Britain with interconnected steep-sided glacial valleys.
The landsystem contains: sequences of recessional moraines on the valley floors and lower slopes (typical of active temperate glaciers); flutings or glacial erosional landforms in the upper valleys; and trimlines marking the former glacier surface.
Use the map to zoom to the Isle of Mull. Zoom in and out and explore the landforms. Turn the basemap to satellite imagery and investigate the geomorphological evidence for yourself. Can you see the features in the satellite imagery?
Hannah Bickerdike completed her BSc in Geography at the University of St Andrews. She subsequently undertook a PhD at Durham University, studying the geomorphology of the Loch Lomond/Younger Dryas Stadial glaciers of Britain. A key element of this work was compiling geomorphological evidence of these glaciers, mapped in previous research, into a GIS database of over 95,000 features, a version of which can be found on this site.
How do we build a glacier? We start with a snowflake. Snow, over time, is compressed into firn, and then into glacier ice.
Snow falls in cold regions, such as mountain tops or in polar regions. In glaciology, snow refers to material that has not changed since it fell1.
Snow is very light and fluffy, and has a very low density. If the snow is wetter, it will have an increased density. Snowflakes have a hexagonal structure, and fallen snow has a significant amount of air in it.
Firn is usually defined as snow that is at least one year old and has therefore survived one melt season, without being transformed to glacier ice.
Firn is transformed gradually to glacier ice as density increases with depth, as older snow is buried by newer snow falling on top of it. Year after year, successive accumulation layers are built up. In the accumulation zone of a glacier, density therefore increases with depth; the rate depends on the local climate and rate of accumulation1. Firn transforms to glacier ice at a density of 830 kg m-3.
New snow (immediately after falling, calm conditions
Damp new snow
Very wet snow and firn
Typical densities (kg m-3). From Cuffey and Paterson, 2010.
Firn transforms to glacier ice in 3-5 years in the temperate Upper Seward Glacier in the St Elias Mountains near the Alaska-Yukon border. Firn becomes ice at a depth of about 13 m1. At sites like this with rapid snow accumulation, the depth of a firn layer, and the load on it, increases rapidly with depth.
However, in cold, dry East Antarctica, firn becomes ice at a depth of 64 m at Byrd and 95 m at Vostok. 280 years are needed at Byrd, and 2500 at Vostok. Low temperatures slow the transformation. Temperatures at Vostok, the coldest region of Earth, are 30°C lower than Byrd, which explains the slower increase in density. In addition, slow accumulation gives slow burial, and a small load each year; the increase in density takes much longer.
Typically, the transformation of firn to ice takes 100-300 years, and a depth of 50 – 80 m1.
Firn becomes glacier ice when the interconnecting air or water-filled passageways between the grains are sealed off (“pore closure”)1. Air is isolated in separate bubbles. This occurs at a density of 830 kg m-3. The air space between particles is reduced, bonds form between them, and the particles grow larger. This is a process known as sintering. Increasing pressure compresses the bubbles, placing the enclosed air under pressure and increasing the density of the ice2.
Fresh snowflakes, which have a complex shape, have a large surface area. Over time and under pressure, the surface area is reduced, the surface is smoothed, and the total surface area is reduced. Fresh, complex snowflakes are transformed into rounded particles.
The transformation of firn to ice is much faster where there is melting and refreezing2. Meltwater can percolate downwards, infilling porespaces, and the displaced air escapes upwards. If the snow is under 0°C, the water will freeze, producing areas of compact ice. This will produce high density ice much more rapidly than in colder regions without melting.
The density of pure glacier ice is usually taken as 917 kg m-3. This strictly is only true at 0°C and in the upper layers of ice sheets and mountain glaciers; the density may be greater at the mid-depth ranges in polar ice sheets, where there are no bubbles and temperatures are -20°C to -40°C1.
Below 4 km of ice, such as at the centre of the East Antarctic Ice Sheet, the pressure would increase the density to 921 kg m-3.
Bubbles are common in glacier ice. Bubbles can contain liquid water or atmospheric gases, making them very useful for ice core research. The air in the bubble largely reflects the atmospheric concentrations when the ice formed1. In polar environments, bubbles in the ice occupy about 10% of the volume when firn turns to ice.
With greater depth in polar ice sheets, bubbles shrink as the overlying ice increases. The gas pressure within the bubbles therefore increases, and at certain depths, the gas attains a dissociation pressure. The bubbles begin to disappear as the gas molecules form clathrate hydrates1. This process takes thousands of years.
Glacier ice contains various impurities in tiny amounts. By most scales, glacier ice is a very pure solid-earth material, because the processes leading to snowfall on a glacier – evaporation, condensation, precipitation – act as a natural distillation system1.
However, glaciers can contain impurities. The dirtiest glaciers are mountain glaciers, where debris can fall directly onto the ice surface. On ice sheets and glaciers, dust and other debris may blow onto the ice surface.
Debris on the ice surface can affect the absorption of energy at the ice surface, and lead to increased or decreased melting.
Layers in the ice
Glaciers are composed of sedimentary layers in their accumulation zones, formed of annual layers of snowfall. These layers are initially parallel to the glacier surface. This is the primary stratification in structural glaciology.
In temperate and subpolar settings, the annual sedimentary layers consist of alternating thick layers of bubble-rich ice, which originated as winter snow, and thin layers of clear ice, which are the refrozen meltwater from the summer melt season.
Debris horizons may form, when summer melting concentrates debris (such as rockfall and wind-blown dust) on the ice surface.
In cold polar regions, annual layering forms instead by seasonal variation of snow metamorphism and wind deposition1.
Blue glacier ice
Glacier ice is blue because the longer visible wavelengths are absorbed. The more energetic, blue, wavelengths are scattered back2. The effect is greatest with deep, basal ice, which is bubble free and has large crystals. The blue colour tends therefore to be most intense in the calls of calved icebergs or fresh fractures.
Rough, weathered ice and fresh snow will appear white because preferential absorption does not occur.
There are many ArcGIS Story Maps around. Some are better than others; some take too long to load or are not well thought through. But some are excellent.
Who owns Antarctica? Antarctic geopolitics
This excellent and well presented, professionally built StoryMap illustrates who owns Antarctica and introduces the Antarctic Treaty.
Glaciation: past, present and future
This great ESRI Storymap introduces glaciers in the present day and at the Last Glacial Maximum. It uses shapefiles from resources like GLIMS or the Randolph Glacier Inventory, and the Global LGM shapefiles from Ehlers and Gibbard. It’s well made and up to date. It also uses the BRITICE map to introduce glacial landforms across the British Isles and the LGM and Younger Dryas in Britain.
Glacier lake hazards in Alaska
This ArcGIS Story by the Alaska Climate Science Centre is better than most. It’s all about glacier hydrology and glacier lakes in Alaska. The videos and pictures are well made, and interspersed with explanatory figures.
A nice, well illustrated, introduction to glacier recession and mapping glacier change over recent decades.
The recession of Glacier National Park Glaciers
This is a great introduction to using imagery to track and map glacier recession from 1966 to the present day.
Glacial Landforms Story Map
This ESRI Story Map introduces a host of glacial landforms. It was a little slow to run for me, though.
An Introduction to Sea Ice
A lovely storymap that introduces and illustrates sea ice, with illustrations of how it changes with the seasons at both poles.
Mapping Mount Everest
This storymap, by Alex Tait from the National Geographic Society, tells us all about Mount Everest and how we map it, with some beautiful graphics.
Glacial Landforms of Snowdonia
A straightforward storymap that highlights the glacial landforms in Snowdonia.
Connecticut’s landscape is the story of glaciers
Learn about the glacial landforms of Connecticut.
The River Tees from Source to Mouth
More fluvial than glacial, but this is a very nice storymap that covers the River Tees. It is suitable for pre-16 as well as post-16 education.
Eskers are ridges made of sands and gravels, deposited by glacial meltwater flowing through tunnels within and underneath glaciers, or through meltwater channels on top of glaciers. Over time, the channel or tunnel gets filled up with sediments. As the ice retreats, the sediments are left behind as a ridge in the landscape.
Eskers are important, because they can tell us about how ice sheets and glaciers behaved. They can tell us about meltwater, and help us reconstruct the former ice surface, and the orientation of the glacier’s snout.
What do eskers look like?
Eskers are usually metres to tens of metres high, and tens to hundreds of metres wide e.g., 2,3. In cross-section, their shape can be sharp-crested (triangular), round-crested (semi-circular), flat-topped (trapezoid), or multi-crested (having two or more crests).
Eskers can range in length from hundreds of metres to hundreds of kilometres. The individual esker ridges that formed beneath the huge, continental-scale ice sheet that covered North America, for example, can extend up to ~100 km in length. Groups of aligned ridges can form fragmented esker chains up to ~300 km long4. Similarly long eskers in Scandinavia were formed by the Eurasian Ice Sheet.
Why are eskers important?
Eskers that formed in subglacial tunnels are valuable tools for understanding the nature and evolution glaciers and ice sheets. They record the paths of basal meltwater drainage near to the ice margin.
The weight of the overlying ice means that the subglacial meltwater is under high pressure. It can therefore flow uphill! This means that, on a local scale, eskers commonly go uphill and climb up local topography.
The path taken by the pressurised meltwater in subglacial channels is controlled mostly by the slope of the ice surface, rather than the slope of the bed. Eskers therefore tend to be oriented parallel to ice flow, and transverse to the ice terminus. As a result, the path of an esker section can be used to reconstruct the slope of the ice surface, and the orientation of the ice terminus at the time of its formation.
Eskers produced by the last North American and Eurasian Ice Sheets probably record the final retreat of those ice sheets as climate warming increased the rate of meltwater production towards the end of the Pleistocene. Therefore, by studying eskers, we can better understand how glaciers and ice sheets respond to climate warming.
These palaeoglaciological insights are essential for predicting the responses of the contemporary Antarctic and Greenland Ice Sheets to human-induced climate change, and their potential contributions to sea level rise.
Where do eskers form?
Eskers are abundant across the land that was once covered by the former North American (Laurentide) Ice Sheet6, the Eurasian Ice Sheet6, and the British-Irish7 Ice Sheet. You can explore British Ice Sheet eskers using Britice map.
Subglacial eskers that formed in subglacial meltwater channels (termed R-channels, which are incised upwards into the basal ice) are the most common among those preserved on palaeo-ice-sheet beds. Good examples of more recently formed eskers are seen, for example, at Breiðamerkurjökull in Iceland2, and Høybyebreen in Svalbard8.
In the embedded Google Map below, look for the raised ridges of the eskers that formed in front of Breiðamerkurjökull. The zig-zagging eskers are largely in the direction of flow, whereas the moraines are parallel to the ice margin.
Eskers on paleo-ice-sheet beds are more abundant in areas of crystalline bedrock with thin coverings of surficial sediment than in areas of thick deformable sediment e.g., 9,4. This is because meltwater flowing at the bed is more likely to incise upwards into the ice to form an R-channel where the bed is hard; where the bed is deformable, meltwater is more likely to incise downwards10.
How long does it take to form an esker?
The timescales over which eskers form is a key topic of ongoing debate. Long eskers extending hundreds of kilometres over paleo-ice-sheet beds are not thought to have formed ‘synchronously’ i.e. at a single moment in time in continuous conduits extending deep into ice sheet interiors. Rather, their formation is thought to have been ‘time-transgressive’, with eskers ‘growing’ at their headward ends as their parent glaciers and associated meltwater conduits retreat across the landscapee.g., 11,12.
Under this mechanism, meltwater conduits need not extend more than a few tens of kilometres into the ice interior, beyond which the weight of the overlying ice would make it hard to form stable drainage conduits. Long esker systems may therefore take hundreds to thousands of years to form12.
Shorter eskers (hundreds of metres to tens of kilometres in length) could form synchronously, possibly over periods of days-to-weeks, during high-magnitude drainage events such as glacial outburst floods13,14.
About the Author
Dr Frances Butcher is a planetary scientist researching glaciers on Earth and Mars. She completed her PhD entitled ‘Wet-Based Glaciation on Mars’ at the Open University (UK) in 2019. She is currently a member of the European Research Council (ERC) funded PALGLAC project at The University of Sheffield (UK), using glacial landforms on Earth to reconstruct the dynamics of the former Scandinavian Ice Sheet. Frances has been involved in preparations for the ESA-Roscosmos ExoMars (‘Rosalind Franklin’) Rover mission, which launches to Mars in 2022.
Welcome to the Younger Dryas Glacial Map! Here, you can explore the glaciation of the UK during the Younger Dryas glaciation. In the UK, this period is also called the “Loch Lomond Stadial”.
At this time (12,900 to 11,700 years ago), there was a period of abrupt cooling. Glaciers began to grow again in much of upland Britain. There was a large ice field, running the length of the Western Highlands in Scotland. This icefield was surrounded by numerous smaller icefields, ice caps, valley glaciers and cirque or niche glaciers. Glaciers also grew in Snowdonia, the Brecon Beacons, the Lake District, and on the Hebridean Islands.
Younger Dryas Glacial Map
This glacial readvance left behind a very distinctive geomorphological imprint on the UK. You can explore these data using our Younger Dryas Glacial Map! This is an ArcGIS Online Map that shows the geomorphological evidence for glaciation and the reconstructed glaciers and ice caps.
The Younger Dryas Glacial Map includes all the geomorphological information relating to the Younger Dryas glaciation of the UK. The data were compiled in a series of papers by Bickerdike et al. (2016, 2018a, 2018b). Here, we have hosted these shapefiles on ArcGIS Online and made them publically available. Click the image below to launch the Younger Dryas Glacial Map.
How to use the map
Launch the Younger Dryas Glacial Map by clicking here.
Read the information on the Spash Screen and then say OK. You will see a map of the UK (as the figure above).
Zoom in and out
You can zoom in on a selected area by pressing Shift and drawing a box with your mouse (SHIFT and DRAG). Click the Home icon on the left hand tool bar to return to the default extent. You can also zoom by scrolling with your mouse, or by using the + and – buttons.
Tools and toolbars
There are a number of icons in the left hand corner. You can hover over them in the webmap for an explanation.
On the vertical bar, the + and – buttons allow you to zoom in and out, and the House returns you to the full map extent. The circle shows you where you are. The square box makes the map full-screen; press escape to return. The left and right arrows take you to the past or next extent.
The five horizontal buttons are, in turn, Measure, Basemap, Legend, Layer, and Information (same as the Splash screen). You can turn layers on and off using Layer List, and fiddle with their transparency and other settings.
You can view the Metadata by clicking on the three little dots to the right of each layer in the Layer List.
Change the Basemap
If you wish, you can change the Basemap to a digital terrain model, satellite imagery, or some other kind of basemap. Change it to “Light Grey Canvas” to speed up drawing time.
Try zooming in, and then changing the Basemap to “Satellite Imagery” to see high-resolution satellite imagery of your field of view.
Choose “Terrain with Labels” to see a digital terrain model.
Try zooming in over some features on an area of the map (press shift and draw a square with the mouse). Clicking a feature will select it, and bring up a popup. The Popup has the attribute information (including the reference of the original authors who first mapped the feature), and a description and photograph of a typical example.
Zoom to a new area and once the map has loaded, click the little black tab at the base of the screen. This will bring up the Attribute Table for all the layers. By default, they are filtered to the map extent. You can therefore view the attribute data held for each layer in the current map view.
Try zooming over a small area and viewing the attribute data for each type of landform.
The Select Tool is in the top right corner. You can use this tool to select different landforms, and then view them in the attribute table. You may wish to turn Basemap to “Light Grey Canvas” if drawing time is slow.
Tick the check box in the Select Tool drop-down menu for the layer that you want to select. Then draw a box with your mouse to select some features.
You can view the selected features in the Attribute Table, by clicking on the three little dots on the right of the Select drop-down menu.
Geomorphological Data in the Younger Dryas Map
The glacier outlines were reconstructed in Bickerdike et al (2018a) using a number of key landforms: moraines, meltwater channels, drift limits, and trimlines. From these data, Bickerdike et al. (2018a) reconstructed the ice limits and the extent of ice-dammed lakes. The geomorphological evidence has been organised into a series of shapefiles, each containing a different landform type.
The database contains over 95,000 individual features, which are organised into thematic layers and each attributed to its original citation. The evidence includes moraines, drift and boulder limits, drift benches, periglacial trimlines, meltwater channels, eskers, striations and roches moutonneés, protalus ramparts and ice-dammed lakes.
There are three moraine shapefiles in the Younger Dryas Glacial Map. Moraine ridges are the most abundant; these are the moraine ridges that were deposited at the terminus of glaciers during the Younger Dryas glaciation of Britain.
Some researchers represent the shape and distribution of individual moraine mounds as polygons, whilst others record only ridge crests as line features or just general areas of ‘hummocky moraine’. These different styles of mapping cannot be reconciled within a single layer in ArcMap, and so are split into separate layers accordingly (‘Moraines (detail)’, ‘Moraine_Ridges’ and ‘Moraine_Hummocky_Area’).
Generally, in areas where mapping of individual features overlapped areas of general ‘hummocky moraine’, only the detailed features were digitised to prevent the database becoming too cluttered.
Drift limits denote the extent of glacier till or sediment that was directly deposited by the glacier. Drift limits can be used alongside moraines to demarcate the maximum extent of the ice limits.
Trimlines form at the ice surface in valleys; the area below the trimline was covered in glacier ice and wasa subjected to glacial erosion. The area above the trimline was exposed above the ice surface, and subjected to frost-shattering and periglacial processes. This leads to a distinct difference in appearance in bedrock above and below the trimline. This can help to reconstruct the ice surface.
Meltwater channels were formed by the passage of water and can form under the glacier (subglacial), in front of the glacier (proglacial) or on the valley side, between the glacier and any valley sides (ice-marginal).
Ice-marginal meltwater channels can extend up onto plateaus, where they form extensive networks. Ice-marginal meltwater channels often form where the ice is thin and cold-based, and frozen to the substrate. Subglacial drainage was inhibited, for at least some of the melt season.
Meltwater channels can therefore help to determine the location of the ice margin, and also determine the nature and style of glaciation.
Using the instructions above, allow the students to explore, interact with, and get used to the map. Encourage them to use all the functions and investigate the glacial landforms themselves.
Does anyone live near to any glacial landforms? Has anyone been to anywhere on holiday that might have been glaciated during the Younger Dryas glaciation?
Encourage the students to use the popups to work out which was the biggest icefield, and which were second and third largest. How big were they? Use the Measure tool to measure the length of the icefields.
Zoom in to an area with detailed moraine ridges, like the terminal moraines in the Lake District. Turn off the icefields by unchecking “Younger_Dryas_Extent” in the Layer List. Turn on the satellite imagery and see if the students can view the landforms for themselves.
“Fluvioglacial” means erosion or deposition caused by flowing meltwter, from melting glaciers, ice sheets and ice caps. Glacial meltwater is usually very rich in sediment, which increases its erosive power.
Fluvioglacial landforms include sandar (also known as outwash plains; they are braided, sediment-rich streams that drain away downslope away from a glacier), kames and kettles, meltwater channels, and eskers.
Glaciofluvial systems are characterised by strong changes in flow magnitude and frequency. Flow magnitudes can fluctuate strongly on a daily basis, as melt increases and decreases over day and night. It also fluctuates seasonally, in the summer (ablation) and winter (accumulation) seasons.
Glacier meltwater can flow supraglacially (on top of the glacier ice), englacially (within the glacier ice), subglacially (below the glacier ice) and proglacially (in front of, and away from, glacier ice). Surface meltwater can reach the bed by draining through the bases of crevasses and moulins.
These pages outline some of the key glaciofluvial landforms associated with the passage of glacial meltwater. For more information, see Glacier Hydrology.