Although icebergs and sea ice often appear together, they are very different phenomena. Icebergs are floating masses of freshwater ice that have broken off from the seaward end of either a glacier or an ice shelf. They are found in the oceans surrounding Antarctica, in the seas of the Arctic and subarctic, in Arctic fjords, and in lakes fed by glaciers. In contrast, sea ice is essentially frozen seawater. It also occurs within the seas surrounding Antarctica and the Arctic Ocean, as well as its adjacent seas as far south as China and Japan.
Icebergs are magnificent structures composed of a freeboard section (height above the waterline) and an underwater section. Their size is greatest when they initially calve, and Arctic bergs tend to be much smaller than the largest Antarctic ones. After they calve from their parent glacier or ice shelf, icebergs may be released into the open ocean, where they slowly erode from wave action or melt from rising environmental temperatures.
Icebergs of the Antarctic calve from floating ice shelves and are a magnificent sight, forming huge, flat “tabular” structures. A typical newly calved iceberg of this type has a diameter that ranges from several kilometres to tens of kilometres, a thickness of 200–400 metres (660–1,320 feet), and a freeboard, or the height of the “berg” above the waterline, of 30–50 metres (100–160 feet). The mass of a tabular iceberg is typically several billion tons. Floating ice shelves are a continuation of the flowing mass of ice that makes up the continental ice sheet. Floating ice shelves fringe about 30 percent of Antarctica’s coastline. The transition area where floating ice meets ice that sits directly on bedrock is known as the grounding line.
Under the pressure of the ice flowing outward from the centre of the continent, the ice in these shelves moves seaward at 0.3–2.6 km (0.2–1.6 miles) per year. The exposed seaward front of the ice shelf experiences stresses from subshelf currents, tides, and ocean swell in the summer and moving pack ice during the winter. Since the shelf normally possesses cracks and crevasses, it will eventually fracture to yield freely floating icebergs. Some minor ice shelves generate large iceberg volumes because of their rapid velocity; the small Amery Ice Shelf, for instance, produces 31 cubic km (about 7 cubic miles) of icebergs per year as it drains about 12 percent of the east Antarctic Ice Sheet.
Iceberg calving may be caused by ocean wave action, contact with other icebergs, or the behaviour of melting water on the upper surface of the berg. With the use of tiltmeters (tools that can detect a change in the angle of the slope of an object), scientists monitoring iceberg-calving events have been able to link the breaking stress occurring near the ice front to long storm-generated swells originating tens of thousands of kilometres away. This bending stress is enhanced in the case of glacier tongues (long narrow floating ice shelves produced by fast-flowing glaciers that protrude far into the ocean). The swell causes the tongue to oscillate until it fractures. In addition, on a number of occasions, iceberg calving has been observed immediately after the collision of another iceberg with the ice front. Furthermore, the mass breakout of icebergs from the Larsen Ice Shelf between 1995 and 2002, though generally ascribed to global warming, is thought to have occurred because summer meltwater on the surface of the shelf filled nearby crevasses. As the liquid water refroze, it expanded and produced fractures at the bases of the crevasses. This phenomenon, known as frost wedging, caused the shelf to splinter in several places and brought about the disintegration of the shelf.
Map showing the extent of collapse of the Larsen Ice Shelf. The Larsen A Ice Shelf disintegrated in 1995, whereas the Larsen B Ice Shelf broke apart in 2002. Both events were caused by water from surface melting that ran down into crevasses, refroze, and wedged each shelf into pieces.
Most Arctic icebergs originate from the fast-flowing glaciers that descend from the Greenland Ice Sheet. Many glaciers are funneled through gaps in the chain of coastal mountains. The irregularity of the bedrock and valley wall topography both slows and accelerates the progress of glaciers. These stresses cause crevasses to form, which are then incorporated into the structure of the icebergs. Arctic bergs tend to be smaller and more randomly shaped than Antarctic bergs and also contain inherent planes of weakness, which can easily lead to further fracturing. If their draft exceeds the water depth of the submerged sill at the mouth of the fjord, newly calved bergs may stay trapped for long periods in their fjords of origin. Such an iceberg will change shape, especially in summer as the water in the fjord warms, through the action of differential melt rates occurring at different depths. Such variations in melting can affect iceberg stability and cause the berg to capsize. Examining the profiles of capsized bergs can help researchers detect the variation of summer temperature occurring at different depths within the fjord. In addition, the upper surfaces of capsized bergs may be covered by small scalloped indentations that are by-products of small convection cells that form when ice melts at the ice-water interface.
The Arctic Ocean’s equivalent of the classic tabular iceberg of Antarctic waters is the ice island. Ice islands can be up to 30 km (19 miles) long but are only some 60 metres (200 feet) thick. The main source of ice islands used to be the Ward Hunt Ice Shelf on Canada’s Ellesmere Island near northwestern Greenland, but the ice shelf has been retreating as ice islands and bergs continue to calve from it. (The ice shelf is breaking into pieces faster than new ice can be formed.) Since the beginning of observations in the 1950s, the Ward Hunt Ice Shelf has virtually disappeared. The most famous of its ice islands was T-3, which was so named because it was the third in a series of three radar targets detected north of Alaska. This ice island carried a manned scientific station from 1952 to 1974. Ice islands produced by Ellesmere Island calve into the Beaufort Gyre (the clockwise-rotating current system in the Arctic Ocean) and may make several circuits of the Canada Basin before exiting the Arctic Ocean via Fram Strait (an ocean passage between Svalbard and Greenland).
Mountain peaks project through the ice cap on northern Ellesmere Island, Canada. Fred Bruemmer
A third source of ice islands, one that has become more active, is northeastern Greenland. The Flade Isblink, a small ice cap on Nordostrundingen in the northeastern corner of Greenland, calves thin tabular ice islands with clearly defined layering into Fram Strait. Observations in 1984 showed 60 grounded bergs with freeboards of 12–15 metres (40–50 feet) off Nordostrundingen in 37–53 metres (120–175 feet) of water. Similar bergs acted as pinning points for pressure ridges, which produced a blockage of the western part of Fram Strait for several years during the 1970s. In 2003 the multiyear cover of fast ice along the northeastern Greenland coast broke out. This allowed a huge number of tabular icebergs to emerge from the fast-flowing Nioghalvfjerdsfjorden Glacier and Zachariae Isstrøm in northeastern Greenland. Some of these reached the Labrador Sea two to three years later, while others remained grounded in 80–110 metres (260–360 feet) of water on the Greenland shelf.
A newly calved Antarctic tabular iceberg retains the physical properties of the outer part of the parent ice shelf. The shelf has the same layered structure as the continental ice sheet from which it flowed. All three features are topped with recently fallen snow that is underlain by older annual layers of increasing density. Annual layers are often clearly visible on the vertical side of a new tabular berg, which implies that the freeboard of the iceberg is mainly composed of compressed snow rather than ice. Density profiles through newly calved bergs show that at the surface of the berg the density might be only 400 kg per cubic metre (25 pounds per cubic foot)—pure ice has a density of 920 kg per cubic metre (57 pounds per cubic foot)—and both air and water may pass through the spaces between the crystal grains. Only when the density reaches 800 kg per cubic metre (50 pounds per cubic foot) deep within the berg do the air channels collapse to form air bubbles. At this point, the material can be properly classified as “ice,” whereas the lower- density material above the ice is more properly called “firn.”
Corresponding to a layer some 150–200 years old and coinciding approximately with the waterline, the firn-ice transition occurs about 40–60 metres (130–200 feet) below the surface of the iceberg. Deeper still, as density and pressure increase, the air bubbles become compressed. Within the Greenland Ice Sheet, pressures of 10–15 atmospheres (10,100–15,200 millibars) have been measured; the resulting air bubbles tend to be elongated, possessing lengths up to 4 mm (0.2 inch) and diameters of 0.02–0.18 mm (0.0008–0.007 inch). In Antarctic ice shelves and icebergs, the air bubbles are more often spherical or ellipsoidal and possess a diameter of 0.33–0.49 mm (0.01–0.02 inch). The size of the air bubbles decreases with increasing depth within the ice.
As soon as an iceberg calves, it starts to warm relative to its parent ice shelf. This warming accelerates as the berg drifts into more temperate regions, especially when it drifts free of the surrounding pack ice. Once the upper surface of the berg begins to melt, the section above the waterline warms relatively quickly to temperatures that approach the melting point of ice. Meltwater at the surface can percolate through the permeable uppermost 40–60 metres (130–200 feet) and refreeze at depth. This freezing releases the berg’s latent heat, and the visible part of the berg becomes a warm mass that has little mechanical strength; it is composed of firn and thus can be easily eroded. The remaining mechanical strength of the iceberg is contained in the “cold core” below sea level, where temperatures remain at −15 to −20°C (5 to −4°F). In the cold core, heat transfer is inhibited owing to the lack of percolation and refreezing.
As ice depth increases to 12 metres (40 feet) and beyond, the temperature difference between icebergs and ice shelves is negligible. Encyclopædia Britannica, Inc.
For many years, the largest reliably measured Antarctic iceberg was the one first observed off Clarence Island (one of the South Shetland Islands) by the whale catcher Odd I in 1927. It was 180 km (110 miles) long, was approximately square, and possessed a freeboard of 30–40 metres (100–130 feet). In 1956 an iceberg was sighted by USS Glacier off Scott Island (a small island about 500 km [300 miles] northeast of Victoria Land in the Ross Sea) with unconfirmed length of 335 km (210 miles) and width of 100 km (60 miles).
However, there have been many calvings of giant icebergs in the Ross and Weddell seas with dimensions that have been measured accurately by satellite. In 2000 iceberg B-15 broke off the Ross Ice Shelf with an initial length of 295 km (about 185 miles). Although B-15 broke into two fragments after a few days, B-15A—the larger portion, measuring 120 km (75 miles) long by 20 km (12 miles) wide—obstructed the entrance to McMurdo Sound and prevented the pack ice in the sound from clearing out in the summer. In October 2005 B-15A broke up into several large pieces off Cape Adare in Victoria Land because of the impact of distant swell. Iceberg C-19 was an even larger but narrow iceberg that broke off the Ross Ice Shelf in May 2002. It fragmented before it could drift far.
The Antarctic Peninsula has been warming significantly in recent decades (by 2.5 °C [4.5 °F] since the 1950s). Three ice shelves on the peninsula, the Wordie and Wilkins ice shelves on the west side of the peninsula and the Larsen Ice Shelf on the east side, have been disintegrating. This has caused the release of tremendous numbers of icebergs. The Larsen Ice Shelf has retreated twice since 2000; each event involved the fracture and release of a vast area of shelf ice in the form of multiple gigantic icebergs and innumerable smaller ones. The breakout of 3,250 square km (1,250 square miles) of shelf over 35 days in early 2002 effectively ended the existence of the Larsen B portion of the shelf.
Although these events received much attention and were thought to be symptomatic of global warming, the Ross Sea sector does not seem to be warming at present. It is likely that the emission of giant icebergs in this zone was an isolated event. Intense iceberg outbreaks, such as the one described above, may not necessarily be occurring with a greater frequency than in the past. Rather, they are more easily detected with the aid of satellites.
In the typically ice-free Southern Ocean, surveys of iceberg diameters show that most bergs have a typical diameter of 300–500 metres (1,000–1,600 feet), although a few exceed 1 km (0.6 mile). It is possible to calculate the flexural (bending) response of a tabular iceberg to long Southern Ocean swells, and it has been found that a serious storm is capable of breaking down most bergs larger than 1 km into fragments.
Arctic bergs are generally smaller than Antarctic bergs, especially when newly calved. The largest recorded Arctic iceberg (excluding ice islands) was observed off Baffin Island in 1882; it was 13 km (8 miles) long by 6 km (4 miles) wide and possessed a freeboard, or the height of the berg above the waterline, of 20 metres (65 feet). Most Arctic bergs are much smaller and have a typical diameter of 100–300 metres (330–1,000 feet). Owing to their origin in narrow, fast-flowing glaciers, many Arctic bergs calve into random shapes that often develop further as they fracture and capsize. Antarctic bergs also evolve by the erosion of the weak freeboard or via further calving into tilted shapes. Depending on the local shape of the ice shelf at calving, the surfaces of icebergs, even while still predominantly tabular, may be domed or concave.
Most of the erosion taking place on Antarctic icebergs occurs after the bergs have emerged into the open Southern Ocean. Melt and percolation through the weak firn layer bring most of the freeboard volume to the melting point. This allows ocean wave action around the edges to penetrate the freeboard portion of the berg. Erosion occurs both mechanically and through the enhanced transport of heat from ocean turbulence. The result is a wave cut that can penetrate for several metres into the berg. The snow and firn above it may collapse to create a growler (a floating block about the size of a grand piano) or a bergy bit (a larger block about the size of a small house). At the same time, the turbulence level is enhanced around existing irregularities such as cracks and crevasses. Waves eat their way into these features, causing cracks to grow into caves whose unsupported roofs may also collapse. Through these processes, the iceberg can evolve into a drydock or a pinnacled berg. (Both types are composed of apparently independent freeboard elements that are linked below the waterline.) Such a berg may look like a megalithic stone circle with shallow water in the centre.
Icebergs are typically divided into six types. Encyclopædia Britannica, Inc.
In the case of Arctic icebergs, which often suffer from repeated capsizes, there is no special layer of weak material. Instead, the whole berg gradually melts at a rate dependent on the salinity (the salt concentration present in a volume of water) and temperature at various depths in the water column and on the velocity of the berg relative to the water near the surface.
On the basis of their observations of iceberg deterioration, American researchers W.F. Weeks and Malcolm Mellor have proposed a rough formula for predicting melt loss:
–Z = KD
where Z = loss in metres per day from the walls and bottom of the iceberg, K = a constant of order 0.12, and D = mean water temperature in °C averaged over the draft of the iceberg. This yields a loss of 120 metres (400 feet) from iceberg sides and bottom during 100 days of drift in water at 10 °C (50 °F)—a rate that corresponds quite well to survival times of icebergs in waters off the coast of Newfoundland as measured by the International Ice Patrol. It has been suggested that if the melt rate could be reduced by interposing a layer of fabric between the ice and water, an iceberg could theoretically survive long enough to be towed across the Atlantic Ocean from Newfoundland to Spain for use as a water and power source.
In Arctic icebergs, erosion often leads to a loss of stability and capsizing. For an Antarctic tabular berg, complete capsize is uncommon, though tiltmeter measurements have shown that some long, narrow bergs may roll completely over a very long period. More common is a shift to a new position of stability, which creates a new waterline for wave erosion. When tabular icebergs finally fragment into smaller pieces, these smaller individual bergs melt faster, because a larger proportion of their surface relative to volume is exposed to the water.
In the Antarctic, a freshly calved iceberg usually begins by moving westward in the Antarctic Coastal Current, with the coastline on its left. Since its trajectory is also turned to the left by the Coriolis force owing to Earth’s rotation, it may run aground and remain stationary for years before moving on. For instance, a large iceberg called Trolltunga calved from the Fimbul Ice Shelf near the Greenwich meridian in 1967, and it became grounded in the southern Weddell Sea for five years before continuing its drift. If a berg can break away from the coastal current (as Trolltunga had done by late 1977), it enters the Antarctic Circumpolar Current, or West Wind Drift. This eastward-flowing system circles the globe at latitudes of 40°–60° S. Icebergs tend to enter this current system at four well-defined longitudes or “retroflection zones”: the Weddell Sea, east of the Kerguelen Plateau at longitude 90° E, west of the Balleny Islands at longitude 150° E, and in the northeastern Ross Sea. These zones reflect the partial separation of the surface water south of the Antarctic Circumpolar Current into independently circulating gyres, and they imply that icebergs found at low latitudes may originate from specific sectors of the Antarctic coast.
Once in the Antarctic Circumpolar Current, the iceberg’s track is generally eastward, driven by both the current and the wind. Also, the Coriolis force pushes the berg slightly northward. The berg will then move crabwise in a northeasterly direction so that it can end up at relatively low latitudes and in relatively warm waters before disintegrating. In November 2006, for instance, a chain of four icebergs was observed just off Dunedin (at latitude 46° S) on New Zealand’s South Island. Under extreme conditions, such as its capture by a cold eddy, an iceberg may succeed in reaching extremely low latitudes. For example, clusters of bergs with about 30 metres (100 feet) of freeboard were sighted in the South Atlantic at 35°50´ S, 18°05´E in 1828. In addition, icebergs have been responsible for the disappearance of innumerable ships off Cape Horn.
In the Arctic Ocean, the highest latitude sources of icebergs are Svalbard archipelago north of Norway and the islands of the Russian Arctic. The iceberg production from these sources is not large—an estimated 6.28 cubic km (1.5 cubic miles) per year in a total of 250–470 cubic km (60–110 cubic miles) for the entire Arctic region. An estimated 26 percent originates in Svalbard, 36 percent stems from Franz Josef Land, 32 percent is added by Novaya Zemlya, about 6 percent begins in Severnaya Zemlya, and 0.3 percent comes from Ushakov Island. Many icebergs from these sources move directly into the shallow Barents or Kara seas, where they run aground. Looping trails of broken pack ice are left as the bergs move past the obstacles. Other bergs pass through Fram Strait and into the East Greenland Current. As these icebergs pass down the eastern coast of Greenland, their numbers are augmented by others produced by tidewater glaciers, especially those from Scoresby Sund. Scoresby Sund is an inlet that is large enough to have an internal gyral circulation. Water driven by the East Greenland Current enters on the north side of the inlet and flows outward on the south side. This pattern encourages the flushing of icebergs from the fjord. In contrast, narrower fjords offer more opportunities for icebergs to run aground; they also experience an estuarine circulation pattern where outward flow at the surface is nearly balanced by an inward flow at depth. An iceberg feels both currents because of its draft and thus does not move seaward as readily as sea ice generated in the fjord.
Satellite image of Scoresby Sund, Greenland. Jacques Descloitres, MODIS Rapid Response Team, NASA/GSFC
As the increased flux of icebergs reaches Cape Farewell, most bergs turn into Baffin Bay, although a few “rogue” icebergs continue directly into the Labrador Sea, especially if influenced by prolonged storm activity. Icebergs entering Baffin Bay first move northward in the West Greenland Current and are strongly reinforced by icebergs from the prolific West Greenland glaciers. About 10,000 icebergs are produced in this region every year. Bergs then cross to the west side of the bay, where they move south in the Baffin Island Current toward Labrador. At the northern end of Baffin Bay, in Melville Bay, lies an especially fertile iceberg-producing glacier front produced by the Humboldt Glacier, the largest glacier in the Northern Hemisphere.
Some icebergs take only 8–15 months to move from Lancaster Sound to Davis Strait, but the total passage around Baffin Bay can take three years or more, owing to groundings and inhibited motion when icebergs are embedded in winter sea ice. The flux of bergs that emerges from Davis Strait into the Labrador Current, where the final part of the bergs’ life cycle occurs, is extremely variable. The number of bergs decreases linearly with latitude. This reduction is primarily due to melting and breakup or grounding followed by breakup. On average, 473 icebergs per year manage to cross the 48° N parallel and enter the zone where they are a danger to shipping—though numbers vary greatly from year to year. Surviving bergs will have lost at least 85 percent of their original mass. They are fated to melt on the Grand Banks or when they reach the “cold wall,” or surface front, that separates the Labrador Current from the warm Gulf Stream between latitudes 40° and 44° N.
Graph of the change in iceberg number with decreasing latitude in the Northern Hemisphere. Encyclopædia Britannica, Inc.
Much work has gone into modeling the patterns of iceberg drift, especially because of the need to divert icebergs away from oil rigs. It is often difficult to predict an iceberg’s drift speed and direction, given the wind and current velocities. An iceberg is affected by the frictional drag of the wind on its smooth surfaces (skin friction drag) and upon its protuberances (form drag). Likewise, the drag of the current acts upon its immersed surfaces; however, the current changes direction with increasing depth, by means of an effect known as the Ekman spiral. Another important factor governing an iceberg’s speed and direction is the Coriolis force, which diverts icebergs toward the right of their track in the Northern Hemisphere and toward the left in the Southern Hemisphere. This force is typically stronger on icebergs than on sea ice, because icebergs have a larger mass per unit of sea-surface area. As a result, it is unusual for icebergs to move in the same direction as sea ice. Typically, their direction of motion relative to the surface wind is some 40°–50° to the right (Northern Hemisphere) or left (Southern Hemisphere). Icebergs progress at about 3 percent of the wind speed.
When an iceberg runs aground, it can plow a furrow several metres deep in the seabed that may extend for tens of kilometres. Iceberg scour marks have been known from the Labrador Sea and Grand Banks since the early 1970s. In the Arctic, many marks are found at depths of more than 400 metres (1,300 feet), whereas the deepest known sill, or submerged ridge, within Greenland fjords is 220 metres (about 725 feet) deep. This unsolved anomaly suggests that icebergs were much deeper in the past or that sedimentation rates within the fjords are so slow that marks dating from periods of reduced sea level have not yet been filled in. It is also possible that an irregular berg can increase its draft by capsizing, though model studies suggest that the maximum gain is only a few percent. Since not all iceberg-producing fjords have been adequately surveyed, another possibility is that Greenland fjords exist with entrances of greater depth. In the Antarctic, the first scours were found in 1976 at latitude 16° W off the coast of Queen Maud Land in the eastern Weddell Sea, and further discoveries were made off Wilkes Land and Cape Hallett at the eastern entrance to the Ross Sea.
In addition, iceberg scour marks have been found on land. On King William Island in the Canadian Arctic, scour marks have been identified in locations where the island rose out of the sea—the result of a postglacial rebound after the weight of the Laurentide Ice Sheet was removed. Furthermore, Canadian geologist Christopher Woodworth-Lynas has found evidence of iceberg scour marks in the satellite imagery of Mars. Scour marks are strong indicators of past water flow.
Observations indicate that long furrows like plow marks are made when an iceberg is driven by sea ice, whereas a freely floating berg makes only a short scour mark or a single depression. Apart from simple furrows, “washboard patterns” have been seen. It is thought that these patterns are created when a tabular berg runs aground on a wide front and is then carried forward by tilting and plowing on successive tides. Circular depressions, thought to be made when an irregular iceberg touches bottom with a small “foot” and then swings to and fro in the current, have also been observed. Grounded bergs have a deleterious effect on the ecosystem of the seabed, often scraping it clear of all life.
Both icebergs and pack ice transport sediment in the form of pebbles, cobbles, boulders, finer material, and even plant and animal life thousands of kilometres from their source area. Arctic icebergs often carry a top burden of dirt from the eroded sides of the valley down which the parent glacier ran, whereas both Arctic and Antarctic bergs carry stones and dirt on their underside. Stones are lifted from the glacier bed and later deposited out at sea as the berg melts. The presence of ice-rafted debris (IRD) in seabed-sediment cores is an indicator that icebergs, sea ice, or both have occurred at that location during a known time interval. (The age of the deposit is indicated by the depth in the sediment at which the debris is found.)
Noting the locations of ice-rafted debris is a very useful method of mapping the distribution of icebergs and thus the cold surface water occurring during glacial periods and at other times in the geologic past. IRD mapping surveys have been completed for the North Atlantic, North Pacific, and Southern oceans. The type of rock in the debris can also be used to identify the source region of the transporting iceberg. Caution must be used in such interpretation because, even in the modern era, icebergs can spread far beyond their normal limits under exceptional conditions. For instance, reports of icebergs off the coast of Norway in spring 1881 coincided with the most extreme advance ever recorded of East Greenland sea ice. It is likely that the bergs were carried eastward along with the massive production and outflow of Arctic sea ice.
It is ice-rafted plant life that gives the occasional exotic colour to an iceberg. Bergs are usually white (the colour of snow or bubbly ice) or blue (the colour of glacial ice that is relatively bubble-free). A few deep green icebergs are seen in the Antarctic; it is believed that these are formed when seawater rich in organic matter freezes onto the bottoms of the ice shelves.
Most scientists maintain that adding large numbers of melting icebergs to the oceans may cause an increase in global sea level if the polar regions are not replenished by an equal amount of snowfall. Also, the addition of large amounts of freshwater to the oceans may lower the salinity of the upper layers of the ocean and possibly alter the present convection-current regime.
Apart from local weather effects, such as fog production, icebergs have two main impacts on climate. Iceberg production affects the mass balance of the parent ice sheets, and melting icebergs influence both ocean structure and global sea level.
The Antarctic Ice Sheet has a volume of 28 million cubic km (about 6.7 million cubic miles), which represents 70 percent of the total fresh water (including groundwater) in the world. The mass of the ice sheet is kept in balance by a process of gain and loss—gain from snowfall over the whole ice sheet and ice loss from the melting of ice at the bottom of the ice shelf and from the calving of icebergs from the edges of the ice shelf. The effect of summer runoff and from sublimation off the ice surface is negligible.
Annual snowfall estimates for the Antarctic continent start at 1,000 cubic km (240 cubic miles). If the Antarctic Ice Sheet is in neutral mass balance, the annual rate of loss from melting and iceberg calving must be close to this value; indeed, estimates of iceberg flux do start at this value, though some run much higher. Such apparently large fluxes are still less than the mean flow rate of the Amazon River, which is 5,700 cubic km (about 1,370 cubic miles) per year. In Antarctica the annual loss amounts to only one ten-thousandth of its mass, so the ice sheet is an enormous passive reservoir. However, if losses from iceberg calving and ice-shelf melting are greater than gains from snowfall, global sea levels will rise.
At present, the size, and even the sign, of the contribution from Antarctica is uncertain. Consequently, Antarctic ice flux has not been included as a term in the sea-level predictions of Climate Change 2007, the fourth assessment report of the Intergovernmental Panel on Climate Change (IPCC). What is more certain is that the retreat of glaciers in the Arctic and mountain regions has contributed about 50 percent to current rates of sea-level rise. (The rest is due to the thermal expansion of water as the ocean warms.) An increasing contribution is coming from a retreat of the Greenland Ice Sheet, and part of this contribution is occurring as an iceberg flux.
In considering the effect of iceberg melt upon ocean structure, it is found that the total Antarctic melt is equivalent to the addition of 0.1 metre (0.3 foot) of fresh water per year at the surface. This is like adding 0.1 metre of extra annual rainfall. The dilution that occurs, if averaged over a mixed layer 100–200 metres (330–660 feet) deep, amounts to a decrease of 0.015–0.03 part per thousand (ppt) of salt. Melting icebergs thus make a small but measurable contribution to maintaining the Southern Ocean pycnocline (the density boundary separating low-salinity surface water from higher-salinity deeper water) and to keeping surface salinity in the Southern Ocean to its observed low value of 34 ppt or below.
It is interesting to note that the annual production of Antarctic iceberg ice is about one-tenth of the annual production of Antarctic sea ice. Sea ice has a neutral effect on overall ocean salinity, because it returns to liquid during the summer months. Nevertheless, when sea ice forms, it has an important differential effect in that it increases ocean salinity where it forms. This is often near the Antarctic coast. Increased salinity encourages the development of convection currents and the formation of bottom water (masses of cold and dense water). Icebergs, on the other hand, always exert a stabilizing influence on the salinity of the water column. This stabilizing influence manifests itself only when the icebergs melt, and this occurs at lower latitudes.
Individual Arctic icebergs, although similar in numbers to Antarctic bergs (10,000–15,000 emitted per year), are smaller on average, so the ice flux is less. This, however, was not necessarily the case during the last glacial period. It has been postulated that, during the first stage of the retreat of the Laurentide Ice Sheet of North America, a large ice-dammed glacial lake (Lake Agassiz) formed in Canada over much of present-day Manitoba. When the ice dam broke, an armada of icebergs was suddenly released into the North Atlantic. As the icebergs melted, they added so much fresh water at the surface that the normal winter convection processes were turned off in the North Atlantic Ocean. As a result, the Gulf Stream was weakened, and northern Europe was returned to ice-age conditions for another millennium—the so-called Younger Dryas event.
An iceberg is a very large object that can be detected in the open sea both visually and by radar. In principle an iceberg can also be detected by sonar. In the open sea, an iceberg produces squealing, popping, and creaking sounds caused by mechanical stresses and cracking, and these sounds can be detected underwater up to 2 km (more than a mile) away. In summer, bergs can also produce a high-pitched hissing sound called “bergy seltzer,” which is due to the release of high-pressure air bubbles from the ice as it melts in the warmer water.
The discovery of an iceberg depends on the alertness of a ship’s watchkeepers. A decaying iceberg poses additional hazards because of its trail of growlers and bergy bits. Although small in size, they have masses (up to 120 tons for growlers, up to 5,400 tons for bergy bits) that are capable of damaging or sinking ships. As they drop into the sea, icebergs often roll over and lose their snow layers. In a heavy sea, the bergs’ smooth wetted ice surfaces produce a low radar cross section. This makes them difficult to discriminate by eye against foam and whitecaps. Because a ship may steer to avoid a large parent berg, it may be in greater danger from undetected growlers or bergy bits drifting nearby.
The problem of protecting shipping from icebergs is most critical in two regions, the high-latitude Southern Ocean and the northwestern section of the North Atlantic. The Southern Ocean threat is increasing because large container ships—those unable or unwilling to use the Panama Canal—can reach high southern latitudes on transit from Australia or New Zealand to Cape Horn. No special measures are currently in place to protect such vessels. In the North Atlantic, the International Ice Patrol was established in 1914 following the loss of the RMS Titanic to an iceberg in April 1912. Its task is to track icebergs as they enter shipping lanes via the Labrador Current and to keep a continuous computer plot of the known or estimated whereabouts of every berg. Reports are transmitted twice a day to ships. In the past, iceberg positions were sited by ships or aircraft. However, it is becoming more common that icebergs are sited by the interpretation of satellite imagery.
The most useful type of sensor is synthetic aperture radar (SAR), which combines high resolution with day-and-night weather-independent capability. Tools with a pixel size of about 20 metres (65 feet) are capable of resolving most bergs. The new generation of SAR in the early 21st century, such as the Canadian RADARSAT and the European ENVISAT, also surveys wide swaths (up to 400 km [250 miles] wide) in every orbit and thus is capable of surveying the entire danger zone once per day.
During the 1950s and 1960s, attempts were made by the U.S. Coast Guard to find ways of fragmenting icebergs that posed a threat to shipping. All were unsuccessful. Explosive techniques were particularly so, since ice and snow are so effective at absorbing mechanical shock. Often the yield of fragmented ice was no greater than the mass of explosive used. Because of the need to defend offshore drilling and production platforms from icebergs, the viability of explosive techniques has been readdressed more recently. It was found that very cold ice, such as the type found in the lower part of an iceberg, can be fragmented successfully by the use of slow-burning explosives such as Thermit. Thermit can be implanted by drilling. However, implantation is a dangerous process because of the possibility of capsize.
Until these techniques are perfected, icebergs cannot be destroyed. Current protocols call for the location and tracking of threatening icebergs. Iceberg trajectories are then predicted by increasingly sophisticated computer models. If necessary, icebergs are captured and towed out of the way of their targets.
Most sea ice occurs as pack ice, which is very mobile, drifting across the ocean surface under the influence of the wind and ocean currents and moving vertically under the influence of tides, waves, and swells. There is also landfast ice, or fast ice, which is immobile, since it is either attached directly to the coast or seafloor or locked in place between grounded icebergs. Fast ice grows in place by freezing of seawater or by pack ice becoming attached to the shore, seafloor, or icebergs. Fast ice moves up and down in response to tides, waves, and swells, and pieces may break off and become part of the pack ice. A third type of sea ice, known as marine ice, forms far below the ocean surface at the bottom of ice shelves in Antarctica. Occasionally seen in icebergs that calve from the ice shelves, marine ice can appear green due to organic matter in the ice.
Sea ice undergoes large seasonal changes in extent as the ocean freezes and the ice cover expands in the autumn and winter, followed by a period of melting and retreat in the spring and summer. Northern Hemisphere sea ice extent typically ranges from approximately 8 million square km in September to approximately 15 million square km in March. (One square km equals approximately 0.4 square mile.) Southern Hemisphere sea ice extent ranges from approximately 4 million square km in February to approximately 20 million square km in September. In September 2007 the sea ice extent in the Northern Hemisphere declined to roughly 4.1 million square km, a figure some 50 percent below mean sea ice coverage for that time of year. Globally, the minimum and maximum sea ice extents are about 20 million square km and 30 million square km, respectively. Measured routinely using data obtained from orbiting satellite instruments, the minimum and maximum sea ice extent figures vary annually and by decade. These figures are important factors for understanding polar and global climatic variation and change.
As seawater freezes and ice forms, liquid brine and air are trapped within a matrix of pure ice crystals. Solid salt crystals subsequently precipitate in pockets of brine within the ice. The brine volume and chemical composition of the solid salts are temperature-dependent.
Liquid ocean water has an average salinity of 35 parts per thousand. New ice such as nilas has the highest average salinity (12–15 parts per thousand); as ice grows thicker during the course of the winter, the average salinity of the entire ice thickness decreases as brine is lost from the ice. Brine loss occurs by temperature-dependent brine pocket migration, brine expulsion, and, most importantly, by gravity drainage via a network of cells and channels. At the end of winter, Arctic first-year ice has an average salinity of 4–6 parts per thousand. Antarctic first-year ice is more saline, perhaps because ice growth rates are more rapid than in the Arctic, and granular ice traps more brine.
In summer, gravity drainage of brine increases as the ice temperature and permeability increase. In the Arctic, summer gravity drainage is enhanced by flushing, as snow and ice meltwater percolate into the ice. Consequently, after a few summers the ice at the surface is completely desalinated and the average salinity of Arctic multiyear ice drops to 3–4 parts per thousand. Antarctic multiyear ice is more saline because the snow rarely melts completely at the ice surface, and brine flushing is uncommon. Instead of percolating into the ice, snow meltwater refreezes onto the ice surface, forming a layer of hard, glassy ice. In contrast, even though it forms from platelets in seawater, marine ice contains little or no salt. The reasons for this remain unclear, but possible explanations include the densification of the ice crystals or their desalination by convection within the “mushy” crystal layer.
Because sea ice is porous and permeable and the brine held within it contains nutrients, sea ice often harbours rich and complex ecosystems. Viruses, bacteria, algae, fungi, and protozoans inhabit sea ice, taking advantage of the differences in salinity, temperature, and light levels. Algae are perhaps the most obvious manifestation of the sea ice ecosystem because they are pigmented and darken the ice. Algae are found at the top, bottom, and interior of Antarctic sea ice; however, they are found primarily at the bottom of Arctic sea ice, where they can occur as strands many metres in length. Sea ice algae are important as a concentrated food source for krill and other zooplankton. Melting sea ice rich in algae may also be important for seeding phytoplankton blooms in the previously ice-covered ocean.
Sea ice that is not more than one winter old is known as first-year ice. Sea ice that survives one or more summers is known as multiyear ice. Most Antarctic sea ice is first-year pack ice. Multiyear ice is common in the Arctic, where most of it occurs as pack ice in the Arctic Ocean.
Pack ice is made up of many individual pieces of ice known as cakes, if they are less than 20 metres (about 66 feet) across, and floes, which vary from small (20–100 metres [about 66–330 feet] across) to giant (greater than 10 km [about 6 miles] across). As the ice drifts, it often breaks apart, and open water appears within leads and fractures. Leads are typically linear features that are widespread in the pack ice at any time of year, extend for hundreds of kilometres, and vary from a few metres to hundreds of metres in width. In winter, leads freeze quickly. Both new and young ice are often thickened mechanically by rafting and ridging, when they are compressed between thicker floes. A pressure ridge is composed of a sail above the waterline and a keel below. In the Arctic most keels are 10–25 metres (about 33–80 feet) deep and typically four times the sail height. Keel widths are typically 2–3 times the sail width. Antarctic pressure ridges are less massive than Arctic pressure ridges. Though they only make up about 25 percent of the total ice area in both polar regions, approximately 40–60 percent of the total ice mass is contained within pressure ridges.
Ice crystals growing on the ocean surface typically break down quickly into smaller pieces that form a soupy suspension known as frazil or grease ice. Under calm conditions the crystals freeze together to form a continuous sheet of new ice called nilas. It is up to 10 cm (about 4 inches) thick and looks dark gray. As the sheet ice thickens by freezing at the bottom, it becomes young ice that is gray to grayish white and up to 30 cm (about 1 foot) thick. If new and young ice are not deformed into rafts or ridges, they will continue to grow by a bottom-freezing process known as congelation. Congelation ice, with its distinctive columnar crystal texture due to the downward growth of the ice crystals into the water, is very common in Arctic pack ice and fast ice.
Pressure ridge in multiyear sea ice thrust up against the northernmost coast of Ellesmere Island, Queen Elizabeth Islands, Canada. M.O. Jeffries, University of Alaska Fairbanks
Under more turbulent conditions, when the water is disturbed by wind and waves, frazil crystals agglomerate into discs known as pancakes. As they grow from a few centimetres to a few metres across, they solidify and thicken mechanically by rafting on top of each other. Pancakes freeze together to form cakes and floes, which contain a large amount of ice with a granular texture. The “frazil-pancake cycle,” though it occurs in both hemispheres, is particularly important in Antarctica, where it accounts for the rapid expansion of ice cover during the autumn and winter. Consequently, Antarctic ice floes generally contain a larger amount of granular ice and a smaller amount of columnar ice than Arctic ice floes.
Frazil, grease, and pancake ice formation also occur in polynyas, which are recurrent features that remain partially or totally ice-free in areas normally expected to be covered with sea ice. They are particularly common in Antarctica, where katabatic winds blowing off the continent force the ice at the coast away from shore, leaving the ocean surface ice-free and open to further ice growth. Ice formation and removal can be almost continuous in coastal polynyas. Consequently, they are sometimes referred to as “ice factories.”
Antarctic ice floes also contain a significant amount of granular ice because the weight of snow is often sufficient to depress the ice surface below sea level, soaking the base of the snow with seawater and producing a slush. When the slush freezes, a layer of granular snow ice is added near the top of the floes.
Platelet ice is perhaps the most exotic form of sea ice besides marine ice. In Antarctica, where cold, relatively low-salinity seawater flows out from beneath ice shelves, platelet ice grows both in the water column and at the bottom of the sea ice on the ocean surface. Whereas platelet ice has been found frozen into pack ice floes, it is most common in fast ice such as the type found in McMurdo Sound. In the Arctic, platelet ice grows primarily in pools of low-salinity water. These pools form at the base of ice floes during the summer months from meltwater runoff.
The large-scale drift of sea ice in the Arctic Ocean is dominated by the Beaufort Gyre (a roughly circular current flowing clockwise within the surface waters of the Beaufort Sea in the western or North American Arctic) and the Transpolar Drift (the major current flowing into the Atlantic Ocean from the eastern or Eurasian Arctic). The clockwise rotation of the Beaufort Gyre and the movement of the Transpolar Drift, the result of large-scale atmospheric circulation, are dominated by a high-pressure centre over the western Arctic Ocean. The pattern is not constant but varies in both strength and position about every decade or so, as the high-pressure centre weakens and moves closer to both Alaska and the Canadian Arctic. This decadal shift in the high-pressure centre is known as the Arctic Oscillation.
The Transpolar Drift exports large volumes of ice from the Arctic Ocean south through Fram Strait and along the east coast of Greenland into the North Atlantic Ocean. Ice drift speeds, determined from buoys placed on the ice, average 10–15 km (about 6–9 miles) per day in the Fram Strait. Ice can drift in the Beaufort Gyre for as much as seven years at rates that vary between zero at the centre to an average of 4–5 km (about 2.5–3 miles) per day at the edge. Together, the Beaufort Gyre and Transpolar Drift strongly influence the Arctic Ocean ice thickness distribution, which has been determined largely from submarine sonar measurements of the ice draft. Ice draft is a measurement of the ice thickness below the waterline and often serves as a close proxy for total ice thickness. The average draft increases from about one metre (about three feet) near the Eurasian coast to 6–8 metres (about 20–26 feet) along the coasts of north Greenland and the Canadian Arctic islands, where the ice is heavily ridged.
In Antarctica the large-scale sea circulation is dominated by westward motion along the coast and eastward motion farther offshore in the West Wind Drift (also known as the Antarctic Circumpolar Current). The average drift speed is 20 km (about 12 miles) per day in the westward flow and 15 km (about 9 miles) per day in the eastward flow. Where katabatic winds force the ice away from the coast and create polynyas, local sea ice motion is roughly perpendicular to the shore. There are gyres in the Ross Sea and Weddell Sea where the westward-moving ice is deflected to the north and meets the eastward-moving ice further offshore. Unlike the Beaufort Gyre in the Arctic Ocean, these gyres do not appear to recirculate ice. Ice thickness data from drilling on floes, visual estimates by observers on ships, and a few moored sonars indicate that Antarctic sea ice is thinner than Arctic sea ice. Typically, Antarctic first-year ice is less than one metre (about three feet) thick, while multiyear ice is less than two metres (about 6.5 feet) thick.
The growth and decay of sea ice influences local, regional, and global climate through interactions with the atmosphere and ocean. Whereas snow-covered sea ice is an effective insulator that restricts heat loss from the relatively warm ocean to the colder atmosphere, there is significant turbulent heat and mass transfer from leads and polynyas to the ocean and atmosphere during the winter months. These losses are manifested as frost smoke from evaporation and condensation at the water surface, and they affect atmospheric processes hundreds of metres above and hundreds of kilometres downstream from leads and polynyas. Brine rejected from ice growing within leads and polynyas drives the deep mixing of the ocean. Rejected brine also affects global ocean circulation and ventilation processes by increasing the salt concentration of the water it is released into. The conversion of both new and young ice into pressure ridges creates rough top and bottom surfaces that enhance the transfer of momentum from the atmosphere to the ocean. Ridges at the ice surface act as sails and catch the wind. The subsequent movement of the ice floes transfers energy to the underlying water via the keels on the underside of the ice.
Frost smoke, open water, and new and young sea ice at a small lead surrounded by pack ice and icebergs in the Bellingshausen Sea, Antarctica. M.O. Jeffries, University of Alaska Fairbanks
Snow and ice reduce the amount of solar radiation available for organisms residing in the ice and water. This decrease in the amount of available energy affects and often reduces the productivity of plants, animals, and micro-organisms. Snow has a high albedo (it reflects a significant proportion of solar shortwave radiation back to the atmosphere), and thus the temperature at the surface remains cool. In the Arctic the surface albedo decreases in summer as the snow melts completely, ponds of meltwater form on the ice surface that absorb a greater share of incoming shortwave radiation, and the overall ice concentration (the ratio of ice area to open water area) decreases. The increase in shortwave radiation absorption by meltwater ponds and the open ocean accelerates the melting process and further reduces surface albedo. This ice-albedo positive feedback plays a key role in the interaction of sea ice with climate.
Submarine sonar data obtained since 1958 have revealed that the average ice draft in the Arctic Ocean in the 1990s decreased by over 1 metre (about 3 feet) and that ice volume was 40 percent lower than during the period 1958–76. The greatest ice draft reduction occurred in the central and eastern Arctic. Remote sensing also revealed a reduction of 3 percent per decade in Arctic sea ice extent from 1978, with particularly rapid losses occurring from the late 1980s. This included the eastern Arctic, where both the ice concentration and the duration of the ice-covered season also decreased. Computer simulations suggest that sea ice changes in this region were due to changes in atmospheric circulation, and thus ice dynamics, rather than higher air temperatures. Yet it is not clear whether these changes are due to natural variability—i.e., the Arctic Oscillation—or whether they represent a regime shift that will persist and perhaps become even more severe in the future.
Since computer models of climate change predict that the consequences of global warming will occur earlier and be most pronounced in the polar regions, particularly the Arctic, monitoring and understanding the behaviour of sea ice are important. Continued reductions of Arctic sea ice extent could have potentially severe ecological impacts. One such event may have arisen in western Hudson Bay, Canada, where a significant decline in the physical condition and reproductive success of polar bears occurred as the duration and extent of sea ice cover decreased during the 1980s and 1990s. On the other hand, a reduction in sea ice could be advantageous for oil and mineral exploration, production, and transport, and for navigation through the Northern Sea Route (Northeast Passage), a water route connecting the Atlantic and Pacific Oceans along the northern coast of Europe and Russia, and the Northwest Passage, a similar route along the northern coast of North America.
Whaling records suggest that Antarctic sea ice extent decreased by approximately 25 percent between the mid-1950s and early 1970s, whereas ice core samples suggest a 20 percent decrease in sea ice extent since 1950. Since then, remote sensing data have indicated an increase in Antarctic sea ice extent parallel to the decrease in Arctic sea ice extent through the 1980s and 1990s. Yet the increase in Antarctic sea ice extent has not been uniformly distributed. A reduction in sea ice extent west of the Antarctic Peninsula has been correlated with slight declines in Adélie penguin numbers and a significant rise in the Chinstrap penguin population. There is speculation that if ice extent continues to decrease in this region, krill numbers will diminish significantly as they lose their under-ice habitat and face growing competition from salps.
