Thus, while perhaps not as dramatic as cyclones, anticyclones have their own unique and interesting features and impacts.
See also
Cold Air Damming.Dynamic Meteorology:Potential Vorticity.Quasi-geostrophic Theory.Static Stability.
Further Reading
Bluestein HB (1992) Synoptic-Dynamic Meteorology in Midlatitudes. New York: Oxford University Press.
Holton JR (1992)An Introduction to Dynamic Meteorol-ogy. New York: Academic Press.
Palmen E and Newton CW (1969)Atmospheric Circulation Systems: Their Structure and Physical Interpretation.
New York: Academic Press.
feedbacks and couplings between the Arctic and global ocean, the Arctic has gained a prominent role in the climate change debate.
Key Physical Features
Most of the area north of 701N is occupied by the Arctic Ocean. Except for the sector between about 201E and 201W, the ocean is surrounded by land
(Figure 1). The Arctic Ocean is hence often referred to as a Mediterranean-type sea. The dominant feature of the ocean surface is its sea ice cover, which ranges in areal extent from about 14.8!106 km2in March to about 7.8!106 km2 in September, but with large seasonal and interannual variability. The ice is typically 1–5 m thick but also highly variable (Barry et al. 1993). Most of the land surface is snow-covered from October through May, with the
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duration of snow cover increasing with latitude. Over the central Arctic Ocean, snow cover is generally present for 10 months of the year. However, precip-itation is generally scant in the Arctic. Much of the land is classified as polar desert, often with less than 5% plant cover. In the low Arctic, the tundra commonly includes shrub vegetation of birch and willow. Permanent land ice is restricted primarily to Greenland (the Greenland ice sheet) and the ice caps and glaciers of north-eastern Canada. However, most of the land is underlain by perennially frozen ground (permafrost), overlain by an active layer exhibiting seasonal thaw. Permafrost acts to keep moisture near the surface. Many areas are covered by shallow thaw lakes.
Atmospheric Circulation
Large-Scale Features
The primary feature of the northern high-latitude, mid-tropospheric circulation is the polar vortex.
The vortex is strongly asymmetric during winter (Figure 2A) with major troughs over eastern North America and eastern Asia and a weaker trough over western Eurasia (the Urals trough). A strong ridge is located over western North America.
The lowest winter pressure heights are located over northern Canada. These features are related to orography, land–ocean distribution and radiative forcing. The polar vortex weakens during summer and is more symmetric than its winter counterpart (Figure 2B).
The dominant sea level features of the mean winter circulation (Figure 3A) are the Icelandic Low off the southeast coast of Greenland, the Aleutian Low in the north Pacific basin, and the Siberian high over central Eurasia. The Icelandic and Aleutian lows are main-tained by low-level thermal effects of the compara-tively warm underlying ocean, position downstream of the major mid-tropospheric stationary troughs where eddy activity is favored and regional cyclone development processes. The Siberian High is a cold, shallow feature, driven largely by long-wave radiative cooling. The Icelandic and Aleutian lows are much weaker during summer (Figure 3B). Summer also sees replacement of the Siberian High by mean low pressure. A weak high pressure cell is found in the Beaufort Sea. An area of mean low pressure is also found near the pole.
Extratropical Cyclone Activity and Polar Lows In accord with the mid-tropospheric steering currents (Figure 2A), winter cyclone activity is most prominent over the Atlantic side of the Arctic (Figure 4A).
Cyclones in this area take a northerly to easterly track, and collectively represent part of the North Atlantic cyclone track. Activity peaks in the vicinity of the Icelandic Low. This is a region of frequent cyclogenesis and strong cyclone deepening rates. Development is enhanced because of contrasts between the warm, northward-flowing North Atlantic drift current and the cold, southward-flowing East Greenland current, proximity to the sea ice margin, and vorticity produc-tion on the lee of the Greenland ice sheet. The North Atlantic track is weaker in summer, but cyclone acti-vity increases over land (Figure 4B). Summer cyclo-genesis occurs in preferred regions over east central Eurasia and over Alaska and extending south-east.
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A summer cyclone maximum is also found over the central Arctic Ocean. This feature arises largely from the migration and subsequent occlusion of systems generated over Eurasia and along the weakened North Atlantic track. Serreze (1995) provides further reading.
Polar lows are mesoscale systems that form within or at the leading edge of polar airstreams. They are commonly found in the Arctic peripheral seas during the winter season. Polar lows are typically less than 500 km in diameter. They may intensify rapidly and surface winds speeds can reach hurricane force (Businger and Reed 1989).
Frontal Activity
Early Canadian analysis schemes adopted a three-front model of the westerlies, with the northernmost representing Arctic fronts, hence separating Arctic from polar air masses. More recent studies based on aircraft data collected during the winter season present clear evidence of separate Arctic jet streams with well-defined tropopause folds (as diagnosed from potential vorticity) between the lower (approximately 5 km) Arctic tropopause to the north and the higher
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(7 km) polar tropopause to the south (Shapiroet al.
1987).
The concept of preferred geographical regions of frontal activity in northern high latitudes emerging as distinct from frontal activity in middle latitudes (termed the ‘Arctic Frontal Zone’) has a long history.
A maximum in frontal frequencies is found during summer along northern Eurasia from about 601 to 701N, best expressed over the eastern half of the continent. A similar relative maximum is found over Alaska, which although best expressed in summer is present year-round. These features are clearly separate from the polar frontal zone in the middle latitudes of the Pacific basin. While some separation between high- and middle-latitude frontal activity is observed in every season, the summer is distinguished by the development of a mean baroclinic zone aligned along the Arctic Ocean coastline and associated wind maxima in the upper troposphere. While it has been postulated that the frontal zone arises from contrasts in energy balance between the tundra and boreal forest, it appears that coastal baroclinicity and focus-ing of the baroclinicity by orography play stronger roles. Regions of maximum summer frontal frequency correspond to preferred areas of summer cyclogenesis over Eurasia and Alaska (Serrezeet al. 2001).
Surface Energy Budget
Figure 5shows typical monthly values of radiative flux components for the central Arctic Ocean. The outgo-ing long-wave flux from the surface decreases from about 320 W m"2in summer (when the sea ice surface is melting) to about 200 W m"2in winter. The incom-ing long-wave flux varies between 160 W m"2 in winter to 300 W m"2in July. For all months, the net long-wave flux is directed away from the surface. The downwelling short-wave flux is zero during winter, rising to about 300 W m"2in June. Because of the high surface albedo (exceeding 0.80 when covered with fresh snow), comparatively little of the solar flux is absorbed by the surface. A fraction of the incoming solar radiation (typically 15%) penetrates into the snow and ice. Net radiation is directed away from the surface from October through March and peaks in June at about 80 W m"2. During winter there is a conductive heat flux through the ice to the surface. On an annual basis, the sensible and latent heat fluxes together account for 20–50% of the net radiation.
During summer, most of the net radiation is used to melt snow and ice. Locally, over areas of open water or thin ice where strong temperature gradients are formed in the boundary layer, winter sensible heat fluxes may reach 600 W m"2. Condensate plumes
emanating from wide (410 km) open water areas (leads and polynyas) that extend to 4 km in the atmosphere have been observed (Barryet al. 1993).
The fundamental difference between the surface energy budgets of the Arctic Ocean, glaciers and tundra is the portion energy used to melt snow and ice.
Once the snow is melted from the tundra, energy can be used in sensible heating and to evaporate water. The consumption of heat through melt on the ocean and glaciers is about 4–6 times larger than on the tundra.
Consequently, sensible heat is transferred from the atmosphere to the surface of the oceans and glaciers, while it is carried from the surface to the atmosphere in the tundra. Evaporation is, on average, the most significant heat sink on tundra, and is considerably larger than on the ocean and glaciers (Ohmura 1984).
A key control on Arctic surface energy budgets is cloud cover. Winter cloud fractions range from 40 to 70%, greatest over the Atlantic side. Total cloud fractions rise to 70–90% in summer. There is a rapid
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increase between April and May, characterized by the development of extensive low-level stratus over the ocean. For most of the year, the cloud radiative forcing is positive, meaning that clouds have a warming effect at the surface. The sign and magnitude of the cloud radiative forcing depends on the solar flux above the clouds, cloud albedo, optical thickness and tempera-ture, surface albedo, and multiple reflections between the surface and cloud base. Curryet al.(1996) reviews Arctic cloud characteristics, radiative forcing, and feedback processes.
Air Temperature and Boundary Layer
The winter surface air temperature field is strongly controlled by downwelling long-wave radiation.
However, as seen in Figure 6A, winter surface air temperatures decrease sharply from the northern North Atlantic to the central Arctic Ocean. The higher temperatures over the Atlantic sector arise from ocean heat fluxes (which keep the region free of sea ice), extensive cloud cover, and horizontal atmospheric heat transports associated with the North Atlantic cyclone track. The lowest winter air temperatures are found over east-central Eurasia in association with the Siberian high. Heat fluxes through areas of open water and thin ice result in comparatively higher tempera-tures over the central Arctic Ocean. Low temperatempera-tures over Greenland reflect elevation. Summer air temper-atures exhibit a much more zonal distribution (Figure 6B). Because of the melting sea ice cover, summer temperatures over the Arctic Ocean are close to zero.
Higher temperatures over land reflect lower latitude and sensible heating of the snow-free surface. Note the strong temperature gradients along the coastline.
Winter radiation deficits give rise to strong surface-based temperature inversions (Kahl 1990). Away from the moderating effects of the Atlantic sector, winter inversions are typically 1000 m deep, with a temper-ature difference across the inversion layer of 10–121C.
Inversion depth and strength vary widely, however, in response to local topographic conditions, winds, and cloud cover. Inversions over the central Arctic ocean tend to be weaker than over land owing to heat fluxes through areas of open water and thin ice. Inversions are also common in summer, although they are weaker than their winter counterparts and are typically separated from the surface by a mixed layer.
Hydrologic Budget
Precipitation and Precipitation Minus Evaporation
(P "E )
Precipitation in the Arctic is difficult to measure be-cause of gauge undercatch of blowing snow, changes
in instrument types and reporting practices, and the sparse precipitation monitoring network. Figure 7 shows the distribution of annual precipitation based on a gridded climatology that compiles data from several sources. The highest totals are found off the south-east coast of Greenland (locally 42400 mm), with amounts decreasing north-east to about 400 mm in the Kara Sea. This pattern manifests the pattern of cyclone activity shown inFigure 2A. High totals are also found over southern Alaska. The lowest annual totals (o200 mm) are found over the Beaufort Sea and northern Canada. The winter pattern is qualitatively similar to that seen in the annual mean. For example, January precipitation ranges from over 200 mm in the northern North Atlantic to less than 10 mm over northern Canada and east-central Eurasia. The Atlan-tic side maximum largely disappears in summer.
Precipitation is more uniform across the Arctic, with markedly higher totals as compared with winter over land areas. This is consistent with seasonal changes in synoptic activity (Figure 2B). Convective precipitation is not unknown over Arctic land areas during summer.
Winter precipitation is largely stored in the snow-pack. Maximum spring snow depths are highly variable due to differences in precipitation, tempera-ture, topographic setting and redistribution by wind.
Values of 20–50 cm over the Arctic Ocean and 40–70 cm over the subarctic can be considered typical.
Mean hydrographs for Arctic rivers exhibit a late spring to early summer peak in discharge due to melt of the snow pack.
Direct estimates of evaporation are very scanty.
However, large-scale estimates of precipitation minus evaporation ðP"EÞ (net precipitation) can be ob-tained through evaluation of the atmospheric vapor flux convergence (Cullather et al. 2000). Estimated mean annualP"E(Figure 8) is typically 150–300 mm over land, 200 mm over the central Arctic Ocean and over 1000 mm in the vicinity of the Icelandic Low.
Although precipitation over much of the land area peaks in summer,P"E for this season (not shown) tends to be small or even negative, pointing to significant regional recycling of water vapor.
Freshwater Budget and Circulation of the Arctic Ocean
The Arctic Ocean is unique in receiving discharge from four of the world’s major rivers (the Ob, Yenisei, Lena, and Mackenzie). River discharge contributes about 360 mm of fresh water to the Arctic Ocean annually.
Along with the inflow of ocean water through the Bering Strait andP"Eover the Arctic Ocean itself, river discharge helps to maintain a relatively fresh surface layer. This layer extends down to about 200 m, and is often well mixed down to about 50 m.
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Figure 6 Mean surface air temperature (1C) for (A) January and (B) July, based on the University of Washington International Arctic Buoy Programme/Polar Exchange at the Sea Surface (IABP/POLES) data set. (Reproduced with permission from Rigor IG, Colony RL and Martin S (2000) Variations in surface air temperature observations in the Arctic, 1979–1997.Journal of Climate13: 896–914.) 152 ARCTIC CLIMATE
Relatively warm waters of Atlantic origin are found at depths between 200 and 900 m, which if brought to the surface would quickly melt the sea ice cover.
However, at low water temperatures of the Arctic Ocean, the density structure is determined by salinity.
Hence the fresh surface layer suppresses vertical
mixing with the Atlantic layer, allowing sea ice to form readily in winter and inhibiting melt during summer.
The large-scale mean annual drift of the sea ice cover is characterized by the clockwise Beaufort Gyre, centered in the Canada Basin, and a mean drift of ice from the Siberian coast, across the pole and through Fram Strait, known as the Transpolar Drift Stream.
This pattern reflects roughly equal contributions by winds and surface currents, the latter ultimately wind driven to a large extent. The mean annual sea ice circulation hence bears a strong resemblance to the mean annual sea level circulation of the atmosphere (Figure 9). Fresh water exported out of the Arctic Ocean, largely via Fram Strait in the form of low-salinity sea ice and liquid water, is believed to impact on the overturning cell of the global ocean through influencing convection in the subarctic gyres which in turn feed the North Atlantic (Lewis 2000).
Variability and Change in Arctic Climate
Arctic climate exhibits pronounced variability on interannual to decadal scales. A major source of variability is associated with the North Atlantic Oscillation (NAO), which describes mutual strength-ening and weakstrength-ening of the Azores High and the Icelandic Low. The NAO has concentrations of power at 24, 8, and 2.1 years and also has a multidecadal signal. Under the positive mode of the NAO (a deep Icelandic Low), positive temperature anomalies are found over the western Eurasian Arctic with negative anomalies over north-eastern Canada and the northern North Atlantic. In turn, the North Atlantic cyclone track extends deeper into the Arctic Ocean. Roughly opposing anomalies are associated with negative NAO states. More recently, attention has been paid to the ‘Arctic Oscillation’ or AO, which represents the leading empirical orthogonal function of monthly sea-level pressure anomalies poleward of 201N (Thompson and Wallace 1998). Pressure vari-ability associated with the AO is characterized by a primary center of action over the Arctic Ocean and opposing anomalies in midlatitudes of the Pacific and Atlantic basins. The AO signal represents a strength-ening and weakstrength-ening of the polar vortex. Time series of the AO and NAO are highly correlated, and the exact relationship between these teleconnections is being debated. Arctic climate variability is also linked to the El Nin˜o Southern Oscillation (ENSO), particularly as it relates to variability in the strength and location of the Aleutian Low, as well as other teleconnections, such as the North Pacific Oscillation (NPO).
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There has been pronounced change in the northern high-latitude environment from the 1970s through the 1990s, in part linked to generally positive phases of the AO and NAO. This includes winter and spring warming over both continents (partly compensated by cooling over the northern North Atlantic). Warm-ing is also evident over the central Arctic Ocean. There has been a downward tendency in sea ice extent and thickness and increased areal extent of the Arctic Ocean’s Atlantic layer. Negative snow cover anoma-lies have dominated over both continents since the late 1980s and terrestrial precipitation has increased in some areas. Small Arctic glaciers have exhibited generally negative mass balances. While permafrost has warmed in Alaska and Russia, it has cooled in
eastern Canada (Serreze et al. 2000). Paleoclimate evidence suggests that Arctic temperatures of the late twentieth century are the highest of the past 400 years (Overpecket al.1997).
General circulation models predict that the effects of anthropogenic greenhouse warming will be ampli-fied in the Arctic owing to feedbacks in which variations in snow and sea ice extent, the stability of the lower troposphere, and thawing of permafrost play key roles. However, regional patterns of Arctic warming differ greatly among simulations. Projected warming is greatest for late autumn and winter, largely because of the delayed onset of sea ice and snow cover.
Retreat of snow cover and sea ice is accompanied by increased winter precipitation.
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Figure 9 Mean annual sea ice drift (cm s"1), based on data from drifting buoys, manned and unmanned camps and mean annual sea-level pressure (hPa). (Reproduced with permission from Barry RG, Serreze MC, Maslanik JA and Preller RH (1993) The Arctic sea-ice climate system: observations and modelling.Reviews of Geophysics31: 397–422.)
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Further Reading
Barry RG (1983) Arctic Ocean ice and climate: perspectives on a century of polar research. Annals, Association of American Geographers73: 485–501.
Barry RG, Serreze MC, Maslanik JA and Preller RH (1993) The Arctic sea-ice climate system: observat-ions and modeling. Reviews of Geophysics 31:
397–422.
Businger S and Reed RJ (1989) Cyclogenesis in cold air masses.Weather and Forecasting4: 133–156.
Cullather RI, Bromwich DH and Serreze MC (2000) The atmospheric hydrologic cycle over the Arctic basin from reanalyses. Part I: Comparison with observa-tions and previous studies. Journal of Climate 13:
923–937.
Curry JA, Rossow WB, Randall D and Schramm JL (1996) Overview of Arctic cloud and radiation characteristics.
Journal of Climate9: 1731–1764.
Kahl JD (1990) Characteristics of the low-level temperature inversion along the Alaskan Arctic coastline. Inter-national Journal of Climatology10: 537–548.
Lewis EL (ed.) (2000)The Freshwater Budget of the Arctic Ocean. NATO Science Series. Series 2. Environmental Security; vol. 70. Dordrecht: Kluwer.
Ohmura A (1984) Comparative energy balance study for Arctic tundra, sea surface, glaciers and boreal forests.
Geojournal8: 221–228.
Overpeck J, Hughen K, Hardy D, et al. (1997) Arctic environmental change of the last four centuries.Science 278: 1251–1256.
Rigor IG, Colony RL and Martin S (2000) Variations in surface air temperature observations in the Arctic, 1979–
1997.Journal of Climate13: 896–914.
Serreze MC (1995) Climatological aspects of cyclone devel-opment and decay in the Arctic.Atmosphere–Ocean33:
1–23.
Serreze MC, Lynch AH and Clark MP (2001) The Arctic frontal zone as seen in the NCEP/NCAR reanalysis.
Journal of Climate(in press).
Serreze MC, Walsh JE, Chapin FS III,et al. (2000) Obser-vational evidence of recent change in the northern high latitude environment.Climatic Change46: 159–207.
Shapiro MA, Hampel T and Krueger AJ (1987) The Arctic tropopause fold. Monthly Weather Review 115(2):
444–454.
Thompson DWJ and Wallace JM (1998) The Arctic Oscil-lation signature in the wintertime geopotential height and temperature fields. Geophysical Research Letters 25:
1297–1300.