17.1: How Ice Is Melting

Surface melt and runoff

In Greenland, ice and snow that melts at the surface of the ice sheet is transported by supraglacial rivers, which end up in large holes in the ice, called moulins, that go straight into the ice for hundreds of meters before migrating in a complex set of horizontal and vertical galleries that ultimately connect the water from the surface to the bed (Figure 17.1.1). Water pressure slowly builds up at the bed in early summer and eventually becomes high enough to overcome the overburden pressure of ice, lifting the ice off the bed so it can slide faster downhill. Subglacial water flows beneath the ice along a network of subglacial channels that are initially disconnected. Eventually, the subglacial channels become connected and reach the ocean, subglacial water pressure is released, and ice again comes into contact with the bed, which stops its enhanced sliding. As more meltwater pours down from the surface to reach the bed, the speedup occurs sooner and is larger in magnitude, but it extends over shorter time periods because the subglacial channels become connected sooner. Overall, in Greenland the summer speedup averages about 10% for a period of 2–3 months; that is, enhanced sliding has only a small impact on the annual mass loss from the ice sheet. This enhanced sliding is often presented in news media as a main process of acceleration of ice toward the ocean, but in reality, we have learned that it is not a major process of ice loss.

Figure 17.1.1 Ice and snowmelt at the surface of Greenland are transported in supraglacial rivers like this one, which ends up in a big hole called a moulin. The moulin goes 100 meters vertically, then connects horizontally and eventually brings water to the bed. The bed is wet. Water pressure at the bed can be high enough to lift the ice. When that happens, the ice moves faster, at least temporarily. Photograph reproduced with permission from The New York Times from Davenport et al. 2015.

Other aspects of snow/ice melt are important. A higher production of meltwater at the surface results in the formation of supraglacial lakes that may break up and cause flooding, a natural hazard. In addition, as snow melts, it is replaced by the underlying ice, and as ice melts, it is replaced by standing liquid water. The albedo of fresh snow is 90%; that is, it reflects 90% of the incoming sun energy and absorbs only 10% as heat that melts snow. In comparison, the albedo of ice is 35%, so 65% of the incoming energy is absorbed. The albedo of liquid water is only 3%, meaning that 97% of the incoming energy is absorbed as heat. As the ice sheet melts away, its surface changes from strongly reflective to strongly absorptive by nearly a factor of 10. This positive feedback keeps pushing the ice sheet out of mass equilibrium.

Another positive feedback is the elevation feedback: As ice melts, the snow/ice surface migrates to a lower elevation, becomes exposed to warmer air temperatures, and melts faster. Similarly, as the ice margins get exposed to the atmosphere when ice melts away, more ice melts, and rocks, fine glacial debris, and dust with no vegetation to hold it are exposed. Dust gets blown away onto the glaciers by the prevailing winds, which makes ice and snow dirty, lowering their albedo, and increasing melt.

The largest positive feedback affecting climate warming, however, is a change in ice flow dynamics, that is, the speed of glaciers, as discussed next.

Ice discharge

Figure 17.1.2 To a photographer in space, Antarctica looks like a vast white continent with no life. Using radar imagery and tracking the scintillation of the surface, scientists can measure the rate of deformation of the ice with amazing precision—a few millimeters per day. This visualization shows ice flow in Antarctica. Flow speed is highest in purple to red areas and low in brown to green. The rivers of ice, or glaciers, reach far into the continent and drain ice into narrow corridors at a few locations along the coast. The glaciers control the flow of ice into the sea. If they flow slowly, the ice sheet grows; if they flow quickly, the ice sheet shrinks. From Rignot et al. 2011.

The second major process of “melting” or ice removal from the ice sheets is the rate at which ice is transported by glaciers and ice streams into the ocean (Figure 17.1.2). When ice reaches the glacier front, several scenarios may take place. Ice may melt in contact with ocean waters, or it may break up into blocks called icebergs. In Greenland, it is not uncommon for icebergs to “explode” into a myriad of small pieces during their detachment from an ice face, because the blocks of ice are too small to remain cohesive under their own weight. Iceberg debris generated in a matter of minutes quickly dissipate into the ocean. Ocean water is a very efficient “solvent” of land ice because the melting point of the seawater-ice mixture is to −2°C, reduced from 0°C for pure water.

As the climate warms up, glaciers produce smaller and smaller icebergs. In Antarctica, where the climate is colder, icebergs are larger and more cohesive and do not explode into a myriad of pieces. They are called tabular icebergs because they form large “tables of ice” several kilometers to tens of kilometers in size afloat in the ocean. They do not flip over as they detach, as the icebergs in Greenland do because they are taller than wide and hence flip to the side when they detach, which helps them break up and melt. In contrast, an Antarctic iceberg will survive for years or decades in the ocean.

Traditional books of glaciology state that iceberg production is the main process of mass loss in Antarctica. We have learned in the last 20 years, though, that a significant part of mass loss proceeds directly from below due to ice melt by the ocean (Figure 17.1.3). This means that ocean warming, or changes in the advection of ocean heat toward the glaciers, or both, is a climate forcing that plays a major role in the evolution of glaciers around Greenland and Antarctica. As frontal ice breaks up and melts away, the inland ice flows faster, effectively unplugging the land ice to spill out into the ocean.

Figure 17.1.3 When ice reaches the ocean, it melts in contact with warm, salty ocean water, and it also breaks into icebergs like this one on Helheim Glacier that is 800 meters tall and 3 kilometers wide. Everyone has heard about icebergs. For a long time, scientists thought that most of the meltwater reaching the ocean was due to the breakup of icebergs. In the 1990s, we learned that a lot of ice is melted from below, by the ocean, orders of magnitude faster than at the surface. Photograph reproduced with permission from Denise Holland.

Ice that is already afloat in the ocean does not change the mass of the ocean or sea level when it melts; in fact, it lowers sea level (by only 2.6%) due to dilution. Ice that rests on land above sea level does raise sea level when it melts into the ocean, at a rate of 1 millimeter for every 361 Gt of ice.

Iceberg calving

Icebergs detach from glaciers or ice shelf fronts when the tensile stress of ice, that is, the rate at which ice is stretched longitudinally by the speedup of ice flow toward the ocean margin, exceeds a threshold. That threshold depends on ice fabrics, ice temperature, and preconditioning of the ice to break up, for example, the presence of cracks or bottom crevasses. Icebergs may also detach when ice blocks are sufficiently close to flotation that they freely rotate off the ice face and fall into the ocean. In the case of an ice shelf, where ice is partly attached to land and partly floating on water, we have also witnessed a domino-like effect: as unstable blocks of ice detach and rotate off the ice shelf front, they may bang back into the ice shelf and generate more calving events. If cracks preexist in the ice shelf, as in the case of the Larsen B ice shelf in the Antarctic Peninsula, the detaching of an iceberg triggers a chain reaction that breaks up the entire ice shelf in a matter of weeks. Conversely, it would take centuries to re-form these ice shelves if we were to let the land ice expand freely into the ocean again, at the same original speed, until it reaches the former position of the ice shelf front.

Scientists have examined the impact of surface meltwater and its role in hydrofracturing ice. Another process is “ice cliff failure” whereby ice cliffs above a certain height can no longer support the ice pressure, fail, and break off. These two calving processes may explain episodes of rapid sea level change that took place in the past. When applied to the present-day evolution of the Antarctic ice sheet, these calving processes yield a sea level rise greater than 1 meter by the end of the century.

Another form of calving that is important in Greenland but which has been highlighted only recently is undercutting. In that process, ice melting by the ocean is most effective at depth; that is, the ocean waters undercut the glacier front. The ice above the cut is then not supported by its base, so it breaks up, independent of the tensile stress or height above flotation. This form of calving is driven by the temperature of the ocean: the warmer the water, the higher the melt rate, and the faster the ice is undercut. It is also affected by the production of subglacial meltwater, discharged at the base of the glacier front, which is buoyantly driven up the water column and entrains warm ocean water along the ice face to melt it. In Greenland, the process of undercutting is comparable in magnitude to the mass loss from “dry” calving by tensile stress and block rotation, but the partitioning between the two varies considerably from one glacier to the next.

Understanding how the calving of icebergs changes as a result of climate warming is an important topic of ongoing research. We need to quantify the roles of hydrofracture, the mixture of ice and sea ice that glues large pieces of ice shelf together before they break away from an ice front, fabrics, temperature, and the stress regime at calving margins. We expect the calving rate of glaciers to increase in a warmer climate, but quantifying the increase has remained challenging because of a lack of observations.

The marine ice sheet instability

A most important process of evolution of ice sheets and glaciers is called the marine ice sheet instability (MISI). This concept was proposed in the 1960s by a number of glaciologists, including Weertman (1974), Thomas and Bentley (1978), and Hughes (1981), and observed in the case of marine-terminating glaciers in Alaska, where it was referred to as the “tidewater glacier cycle,” by Meier and Post (1988) (Figure 17.1.4). We had to wait until the 1990s and early 2000s to verify the concept of marine instability in ice sheets, but scientists studying tidewater glaciers in Alaska knew it decades earlier.

Figure 17.1.4 The most important concept for ice sheets in a warmer world is the concept of marine ice sheet instability (MISI). If a marine-terminating glacier ends in the ocean on a normal (prograde) slope, that is, with bed elevation increasing inland, any glacier retreat or advance will be slow, as there are many stable positions. If the glacier is on a reverse (retrograde) slope, that is, with bed elevation dropping in the inland direction, there are only two stable states: either it reaches the edge of the continental shelf, or it becomes entirely afloat in the ocean. Once initiated, the MISI retreat is in principle (considering one-dimensional geometry) unstoppable. From Thomas 1979.

If a glacier stands on a retrograde slope, that is, where the bedrock slopes downward in the inland direction, there are only two stable states for the glacier: either it reaches the outer edge of that slope and remains stable at that location, or it retreats inland until either the bedrock slopes upward or the entire glacier is afloat in water between the glacier and the bedrock. In other words, glaciers resting on retrograde slopes are inherently unstable and prone to rapid retreat. In Alaska, the edge of the slope is a moraine created by the glacier during a prior advance. If the glacier retreats from the moraine because of climate forcing, for example, warmer ocean waters melt the ice faster than in the past, the retreat will proceed rapidly, in a nonclimatic fashion, until the ice front reaches a new bed position where the bedrock elevation rises again in the inland direction. This could be many kilometers inland. For Columbia Glacier in Alaska, the retreat that started in the 1980s will proceed for another 50 kilometers along a retrograde bed. Conversely, if the bedrock slope is prograde (slopes downward in the ocean direction), the glacier retreat will be slow and may even stop.

The effect of a retrograde slope is large because ice deformation exhibits a nonlinear dependence on ice thickness. The strain of deformation of ice varies as the third power of the thickness, and the speed of the ice varies as the fourth power of ice thickness. Hence a drop in bed elevation as the ice front moves inland translates into an accelerating response of the ice tensile stress, with sliding and ice breakup into the ocean. This is the marine ice sheet instability.

In practice, the MISI is complicated by two dimensional effects. The glacier and its floating extension in the ocean—the ice shelf—experience friction along valley walls (lateral shear), islands (longitudinal backstress), and bumps in bedrock topography (basal friction), so a glacier on a retrograde slope may not be systematically unstable and hence capable of unstoppable and rapid retreat. To address that possibility, precise observations of bed topography, detailed understanding of the ice flow mechanics, proper scenarios of climate forcing, and the usage of a coupled ocean-ice numerical model are essential.

We have already witnessed examples of MISI in present-day ice sheets: (1) the Jakobshavn Isbrae Glacier, the largest discharger of ice in Greenland; and (2) the Pine Island, Thwaites, and Smith Glaciers draining into the Amundsen Sea Embayment of West Antarctica, the largest dischargers of ice in Antarctica. Scientists think that the retreat of ice in these areas is ongoing and unstoppable. There are, however, other sectors at risk of MISI around North Greenland and East Antarctica.

Ocean heat and its impact on ice sheets

How could more ocean heat be brought into contact with ice in a warming climate? The physical processes are the same in the Arctic and Antarctic—ocean heat is driven by wind—but the details differ. In both polar regions, we find cold, fresh water at the top of the water column and warm, salty water at the bottom, about 200–300 meters below the surface in Greenland and 400–500 meters in Antarctica (Figure 17.1.5). This configuration is opposite to that in the tropics, where warm, fresh water sits atop cold, salty water. In polar regions, colder air temperatures cool the surface, and the freezing of seawater forms sea ice, which loses its salt. This produces salty water that sinks to the ocean bottom and participates in the global thermohaline circulation of the ocean.

Figure 17.1.5 In polar seas, ocean heat is found several hundred meters below the surface in the form of warm, salty water. Surface waters are comparatively cold and fresh. Winds transport ocean heat toward the glacier, and a thermohaline circulation ventilates the ice shelf cavities as in this figure. Melt rates of 100 meters per year in Antarctica and several meters per day in Greenland are generated in this manner. In the Antarctic, heat comes from the Antarctic Circumpolar Current. In Greenland, it comes from the subtropics. Reproduced from NASA.

Depending on prevailing winds and the depth of the seafloor around the ice sheets, ocean heat may or may not reach the glaciers. Prevailing winds are affected by climate change. Seafloor depth differs where deep channels have been carved into the seafloor by prior advances of the glaciers. Wind characteristics, depth, and ocean temperature all must be known if they are to be used in numerical ocean and ice models.

When warm seawater reaches the glaciers or ice shelves, it fuels rapid rates of ice melt because ocean water has a melting point of −2°C (versus 0°C for freshwater ice) and the melting point decreases with pressure by 0.75°C per kilometer of water, hence it is −3.5°C at 2 kilometers depth. If warm seawater does not reach the glaciers, that is, if the glaciers stand in cold water, ice melts slowly. In Antarctica, the largest observed melt rates are of the order of 100–200 meters per year. In Greenland, the melt rates reach 2–3 meters per day, that is, one order of magnitude larger than in Antarctica. Conversely, in cold parts of Antarctica, melt rates may drop to values as low as 10 centimeters per year. In some places, ice may even cease to melt completely. Instead, typically up to 100 kilometers from the grounding line, seawater may freeze onto the ice shelf bottom, creating a layer of “marine ice.” Marine ice may accumulate by about 100 meters over time (decades to centuries).

Figure 17.1.6 In Antarctica, ocean heat is contained in the Antarctic Circumpolar Current (ACC). Warm water, about +2°C, is present at 400−700 meters below the surface. At the surface, the water is cold and fresh. Prevailing westerly winds push the ACC in a clockwise direction. The Coriolis force displaces the surface water to the north (left) and displaces the subsurface warm, salty water to the south (right). With climate warming in most of the world and cooling in Antarctica under the ozone hole, the winds are getting stronger and displacing more warm water toward Antarctica. From Gille et al. 2016.

In the Antarctic, ocean heat originates from the Antarctic Circumpolar Current (ACC), which is a broad area of subsurface warm, salty water that encircles the continent, clockwise, pushed by the westerly winds (Figure 17.1.6). In some parts of the Southern Ocean, the ACC is close to the continent, for example, in the Amundsen Sea Embayment of West Antarctica and in the western Antarctic Peninsula. In other parts it is far from the coast, for example, in the Weddell Sea and in the Ross Sea. Since the 1980s, the westerlies have increased in strength and started to contract southward. This is due to an increase in the temperature difference between Antarctica and the rest of the world. Antarctica is not warming as fast as the rest of the world, because a decrease in the ozone concentration in the stratosphere above cools Antarctica, and it experiences a slower rate of warming from greenhouse gas emissions than the rest of the world because of a lack of albedo feedback. As a result of the Coriolis effect, the winds tend to push surface waters to the north (to the left of the wind), which contributes to the slight extension of sea ice cover with time, and to push subsurface waters to the south (to the right), which brings more subsurface ocean heat toward Antarctica’s glaciers. As more heat reaches Antarctica’s coast, glaciers and ice shelves melt faster from below, which reduces the buttressing force in front of them, leads the glaciers to speed up, and increases sea level.

In Greenland, warm water is transported north from the subtropics by the Gulf Stream, which is deviated by Iceland to form the Irminger Current, which runs along southeast Greenland, rounds the tip of Greenland, and reaches the Labrador Sea. Some of that warm water makes it into Baffin Bay through Davis Strait, circles counterclockwise inside Baffin Bay, and returns south along the coast of Canada. To the east, a branch of the North Atlantic Current returns cold water from the High Arctic along the east coast of Greenland and meets the Irminger Current southeast of Greenland. These currents contribute to the North Atlantic gyre, which allows warm, salty Atlantic Water (AW) to intrude onto the continental shelf and glacial valleys by following troughs on the seafloor that have been carved by former glacier advances during ice ages. The strength of the North Atlantic gyre is affected by fluctuations of the jet stream, itself affected by climate change.

At present, the Arctic is warming up 2–3 times faster than the rest of the world. As the temperature differential between the Arctic and the rest of the world decreases, the strength of the jet stream is reduced, allowing it to wobble, that is, undergo large excursions and incursions north and south. In some of the lobes of the Rossby waves—global air pressure waves high in the atmosphere—cold Arctic air flows unusually far south, which creates cold snaps along the east coast of the United States, for instance. In other lobes, warm air from the subtropics intrudes far north, which creates unusually warm winters in Greenland. Models suggest that as the jet stream wobbles, the lobes tend to become stationary, hence the unusual flow of cold, and air masses may persist for long periods of time. This simple explanation of Arctic changes is in debate but offers an explanation for changes taking place in the north. If the models are correct, the wobbling will send more warm air and ocean masses toward Greenland than in the past, which will melt the glaciers from above and below faster than in the past.

In both Greenland and Antarctica, the amount of warm, salty subsurface ocean water pushed by the prevailing winds toward the glaciers has been changing in response to climate change, which is caused by human activities.