Calving of freshwater glaciers

This article, while drawing on a variety of literature, is largely based on an excellent review of calving processes by Prof. Doug Benn and colleagues in the journal Earth Science Reviews. If you want to know more about glacier calving, this review paper would make a great starting point.

What is calving?

Calving is the glaciological term for the mechanical loss (or simply, breaking off) of ice from a glacier margin1. Calving is most common when a glacier flows into water (i.e. lakes or the ocean) but can also occur on dry land, where it is known as dry calving2.

The calving margin of Perito Moreno Glacier in Argentinean Patagonia. Photo: Liam Quinn

Why is calving important?

In lake-terminating (or freshwater) glaciers, calving is often a very efficient process of ablation and is therefore an important control on glacier mass balance4-7.

Calving is also important for glacier dynamics and ice retreat rates1. Calving glaciers are often highly dynamic, with patterns of behaviour (e.g. glacier advance and recession) that are at least partly decoupled from climate4-7.

Differences between freshwater and tidewater glaciers

Compared to marine-terminating (or tidewater) glaciers, like those at the margins of the Antarctic and Greenland ice sheets, freshwater glaciers are normally smaller and slower moving. This means that while the processes of calving can be very similar in both settings, freshwater glaciers tend to have lower calving rates1.

Calving basics: the fracture of ice

Before calving occurs, smaller cracks and fractures in glacier ice grow (or propagate) into larger crevasses (see image below). The growth of crevasses effectively divides the ice into blocks that subsequently fall from the snout into an adjacent lake (where they are known as icebergs). Therefore, the fracture of ice is an important control on: where calving will occur, the size of calved icebergs, and how often calving events happen1.

The heavily crevassed snout of Fjallsjökull, Iceland, with ice blocks ready to be released in claving events. Photo: Wojciech Strzelecki

Growth of fractures

Cracks and fractures in glacier ice will grow when the stress acting on a fracture is larger than the fracture strength of ice8. When this condition is met, the ice will fracture in a brittle fashion, causing existing cracks to become deeper and wider. Large stresses occur in many situations in glaciers. Good examples include: where glacier ice is extended (‘pulled apart’) or compressed (‘squeezed together’) as a result of flow.

Example of large fractures in the snout of Goldbergkees Gletscher in the Austrian Alps. Photo: Ewald Gabardi

Water-filled crevasses

Water plays a key role in the depth of crevasses and the likelihood of calving (see diagram below). In a water-free crevasse, stress at the crevasse tip is offset by the weight of overlying ice. This causes a fracture to close up. However, in a water-filled crevasse, the pressure of water offsets the weight of ice. This allows a crevasse to extend deeper into the ice, and often to the glacier bed8.

Crevasses often open up due to extending glacier flow that causes the ice to ‘stretch’ or be pulled apart. In a water free crevasse, the weight of ice forces the fracture shut. In a water-filled crevasse, by contrast, the added water pressure counteracts the weight of ice, allowing a fracture to deepen. (diagram modified from ref. 8)

Calving processes

There are several main calving mechanisms at freshwater glaciers, all of which are related to stress at the glacier terminus1.

Stretching and crevassing of ice

At a grounded lake-terminating glacier, ice flow commonly becomes faster (due to basal sliding) near the snout. This happens because the snout is close to floating in lake water, which reduces frictional drag at the bed1,9. The faster flow near the terminus causes the ice to ‘stretch out’, and crevasses to propagate through the glacier (see diagram below). This process, known as longitudinal stretching, creates heavily crevassed glacier snouts (see image below). Calving occurs along the lines of weakness formed by crevassing1,9,10.

Faster ice flow near the glacier margin, due to reduced basal drag, causes the ice to stretch out and crevasses to open. When stresses are high, crevasses propagate through the glacier and calving occurs.

The heavily crevassed terminus of Grey Glacier, Chilean Patagonia, formed in part due to longitudinal stretching of ice. Photo: NASA Earth Observatory

Crevasses may also form in areas further upglacier, such as in icefalls, where ice flows rapidly across steep terrain8. The crevasses formed in icefalls provide likely zones for calving when they move downglacier to the terminus (see diagram below)1.

Crevasses formed in areas of steep terrain, such as icefalls, provide likely zones of iceberg calving when they reach the glacier terminus. T1 = crevasses form in icefall. T2 = crevasses move downglacier and promote calving.

Force imbalances at glacier terminus

At a floating glacier terminus, the outward-directed cryostatic pressure (i.e. the pressure exerted by ice) and inward-directed hydrostatic pressure (i.e. the pressure exerted by water) are out of balance (see diagram below)11. Below the lake waterline, hydrostatic pressure in part balances cryostatic pressure. However, above the waterline, there is very little inward-directed force (from the atmosphere) to counter cryostatic pressure11. This imbalance creates a zone of high stress at the ice surface, opening crevasses and promoting calving1.

Diagram to illustrate differences between outward-directed cryostatic pressure and inward-directed hydrostatic pressure. Above the lake waterline, the atmosphere provides little offset to cryostatic pressure, which results in high stresses in the terminal ice cliff and calving.

Undercutting of a terminal ice cliff

Glacier ice at or below a lake waterline often melts at a faster rate than the ice above a lake waterline. Waterline melting will often erode a notch that undercuts the calving ice cliff (see image below)6,12,13. Once undercut, calving may occur by forward toppling of overhanging ice blocks, or where the roof of a waterline notch collapses1.

Waterline notches often develop during summer but cease forming in winter when glacial lake temperatures are cooler and/or when the lake surface freezes over. Calving by notch erosion therefore tends to follow a seasonal pattern6,12,13.

Example of a thermo-erosional notch cut into the terminal ice cliff. Photo: Michael Clarke

Melting at or below a lake waterline can erode a notch into a terminal ice cliff (T1). As the notch grows over time, the ice cliff becomes unstable and blocks topple outwards (T2).

Calving event at Perito Moreno Glacier in Argentinean Patagonia, due to collapse of an englacial tunnel roof. Photo: Rafael Bernstein

Buoyant forces at a glacier terminus

Where a glacier surface thins to below the level needed for ice flotation, the margin will become buoyant and lift off the bed7. If the surface continues to thin, buoyancy increases, causing large bending forces at the grounding line, the growth of large crevasses, and eventually calving7. This processes often produces large icebergs.

Buoyancy can cause large bending forces at the glacier grounding line when the glacier surface falls to flotation level. Large tabular icebergs are released by this process.

Buoyant forces may also cause calving below the lake surface. Subaqueous calving often occurs where an ‘ice foot’ has developed due to calving losses above the waterline (e.g. due to notch erosion and toppling ice cliffs). The loss of ice above the waterline decreases the ice overburden pressure pressing down on the ‘ice foot’, allowing upward buoyant forces to fracture the ice and cause calving1. In such events, icebergs can rapidly shoot up to the lake surface, and sometimes emerge 100s of metres away from the ice front.

Buoyant forces cause calving of a subaqueous ‘ice foot’ due to a loss of ice above the waterline and reduced ice overburden pressure.


[1] Benn, D.I., Warren, C.R. and Mottram, R.H., 2007. Calving processes and the dynamics of calving glaciers. Earth-Science Reviews, 82, 143-179.

[2] Diolaiuti, G., Smiraglia, C., Vassena, G. and Motta, M., 2004. Dry calving processes at the ice cliff of Strandline Glacier northern Victoria Land, Antarctica. Annals of Glaciology, 39, 201-208.

[4] Warren, C.R., 1994. Freshwater calving and anomalous glacier oscillations: recent behaviour of Moreno and Ameghino Glaciers, Patagonia. The Holocene, 4, 422-429.

[5] Naruse, R. and Skvarca, P., 2000. Dynamic features of thinning and retreating Glaciar Upsala, a lacustrine calving glacier in southern Patagonia. Arctic, Antarctic, and Alpine Research, 32, 485-491.

[6] Warren, C.R. and Kirkbride, M.P., 2003. Calving speed and climatic sensitivity of New Zealand lake-calving glaciers. Annals of Glaciology, 36, 173-178.

[7] Boyce, E.S., Motyka, R.J. and Truffer, M., 2007. Flotation and retreat of a lake-calving terminus, Mendenhall Glacier, southeast Alaska, USA. Journal of Glaciology, 53, 211-224.

[8] Benn, D.I., and Evans, D.J.A., 2010. Glaciers and Glaciation. Routledge. pp. 802.

[9] O’Neel, S., Pfeffer, W.T., Krimmel, R. and Meier, M., 2005. Evolving force balance at Columbia Glacier, Alaska, during its rapid retreat. Journal of Geophysical Research: Earth Surface, 110(F3).

[10] Warren, C. and Aniya, M., 1999. The calving glaciers of southern South America. Global and Planetary Change, 22, 59-77.

[11] Reeh, N., 1968. On the calving of ice from floating glaciers and ice shelves. Journal of Glaciology, 7, 215-232.

[12] Kirkbride, M.P. and Warren, C.R., 1997. Calving processes at a grounded ice cliff. Annals of Glaciology, 24, 116-121.

[13] Haresign, E. and Warren, C.R., 2005. Melt rates at calving termini: a study at Glaciar León, Chilean Patagonia. Geological Society, London, Special Publications, 242, 99-109.

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