Main Geographical Factors Shaping the Climate

Undoubtedly, geographical latitude is the main factor determining the weather and climate both in the Arctic and elsewhere. For the purpose of this work, the Arctic has been defined after Atlas Arktiki (1985) (Figure 1.2). From Figure 1.2 it can be seen that the southern boundary of the Arctic thus defined ranges between about 54°N (the Labrador Peninsula) to about 75°N near Spitsbergen. No matter how we define the Arctic, its location in high latitudes limits significantly the magnitude of receiving energy from the sun. In regions lying beyond the Arctic Circle, the most unusual feature is occurrence of seasonal day and night. As we know, the length of both polar night and polar day varies from one day at the Arctic Circle to about six months at the North Pole (Figure 1.3). In addition, because of the atmospheric refraction, the total time when the sun is visible over the horizon during the year is greater at high latitudes than in more temperate latitudes. Also, since the sun crosses the horizon at a shallow angle in the Arctic, dawn and dusk persist for long periods before and after the sun is visible. As a result, winter days arc much longer here than summer nights. The elevation of the sun in noon anywhere cannot be higher than about 47°. This fact is mainly responsible for the lower income of the solar energy (on an annual basis) here than in lower latitudes. However, the total solar radiation in June, which the Arctic receives at the top of the atmosphere, is even higher than in equatorial areas. For example, the solar irradiance flux reaching the upper boundary of the earth-atmosphere system is equal to 129 kJ/cm2 (31 kcal/cm2) and 98.2 kJ/cm2 (23.5 kcal/cm2) at the 80°N and the equator, respectively (Budyko 1971). From a climatological point of view, however, it is only the solar radiation absorbed by the surface which is important. Due to the high albedo of the earth-atmosphere system in the Arctic, this component of the radiation balance is markedly lower than in the rest of the globe.

Looking at the map of the Arctic, one can easily see that the Arctic, in contrast to the Antarctic, consists of an ocean encirclcd by land. The central main part of the ocean is called the Arctic Ocean and is ice-covered year-round, while snow and ice are present on the land for almost all the year. The land encompasses the northern parts of two major land masses - Eurasia and North America - as well as quite a large number of islands, especially on the American side. Of these the largest are Greenland (2,175,600 km2), Baffin Island (476,070 km2), Ellesmere Island (212,690 km2) and Victoria Island (212,200 km2) (The Tunes Atlas of the World 1992). The area of Greenland, including the islands, is 2,186,000 km2(Putnins 1970). A substantial break in the ring of land exists only between Greenland and Norway (Figure 1.2). Other breaks between Asia and America (the Bering Strait) and between the islands of (he Canadian Arctic Archipelago, are of marginal significance. The highest mountains are to be found in south-eastern Greenland, where two sum-

Actual local conditions may vary from values shown

Figure 1.3. Duration of daylight and darkness in the latitude band 60°-90°N (after CIA 1978).

mits rise over 3000 m a.s.!.: Gunnbjorn Fjeld (3700 m, (p = 68°54'N, X = 29°48'W) and Mt. Forel (3360 m, (p = 67°00'N, X = 37°00'W) (The Times Atlas of the World 1992). Much of the Arctic is low lying, except for the Greenland ice sheet, the ice-covered mountains of Ellesmere and Axel Heiberg islands, and the mountains in the northern part of the Beringia region. The differentiated influence of land and sea areas on the climate of the Arctic is significantly lower than in the moderate latitudes. This is true, particularly in winter, when the land and most of the sea areas are covered by snow. The long-term mean depths of snow cover for May, calculated from measurements taken mainly in Russian drifting stations NP3 - NP31 over the period 19541991, vary from 30-40 cm in the centra! part of the Arctic to more than 80 cm in the mountainous regions. The maximum snow-cover depth is most often observed in April or May except in the Canadian Arctic, where it is observed in March. The decay of the snow cover begins in the south of the Arctic in the first ten days of June, and in the vicinity of the Pole in mid-July. The number of days with snow cover is greatest in the central Arctic (more than 350 days). This number decreases towards the south and is equal to about 280-300 days across those Arctic islands which have a continental climate (for more details see sub-section 7.3.).

In general, three physical characteristics of snow - high reflectivity, high infrared emissivity and high insulating property - mean that it plays a very important climatic role. The high albedo of the snow surface significantly reduces the net radiation balance of the surface and low troposphere. The high infrared emissivity of snow is one of the most important factors, which causes near-surface atmospheric temperature inversions, especially in the cold half-year. In addition, it helps in the development and stabilising of the anticyclones. Snow cover, as one of the best insulators of all known natural surfaces, is a veiy important element in the atmosphere-cryosphcre-ocean system, and thus significantly influences heat transport. A snow cover of more than 15 cm in depth may completely stop the heat transport between the atmosphere and land or sea ice.

The Arctic Ocean and its bordering seas occupy an area of 14 million km2 (Barry 1989). In late winter (February-March) almost all this area is covered by sea ice. During the summer (August-September), the sea ice is at its minimum extent (approximately 8 million km2). The role of the sea ice in shaping the climate of the Arctic, and indeed that of the whole globe, is crucial. Generally, four main properties of the sea ice contribute to this. The first property is the significantly higher albedo of sea ice (0.5 to 0.7) in comparison to an open ocean (0.1). As a result, water covered by sea ice absorbs much less radiation than do open waters. A second property is the insulating role of the sea ice, restricting the exchange of heat and moisture between ocean and atmosphere. Maykut (1978) reported that measurements of wintertime sensible heat flux showed that between 10 and 100 times more heat is transferred from a calm open-water ocean to the atmosphere than from an ocean covered by a 2-metre layer of sea ice. A third property is the large latent heat of freezing and melting, which makes sea ice act as a thermal reservoir delaying the seasonal temperature cycle. These processes also alter (he salinity content of the upper layers of the ocean. During freezing, a sea salt is forced out of the sea ice (resulting in an increase of salinity in the water); on the other hand, during melting, the fresh water transferred to the upper layers reduces its salinity. Recently, Wadhams (1995) has drawn our attention to the fact that sca-icc motion (a fourth property), driven mainly by wind stress, is also very important for climate and climate-change studies. Processes connected with this motion, such as divergence and convergence of sea ice, create leads and pressure ridges, respectively. The latter forms contain about half of the total Arctic ice volume (Wadhams 1981). As a result of these processes, atmosphere-ocean heat and moisture fluxes are highly time- and space-dependent.

Sea ice in the Arctic never exists as an unbroken cover or as a floating ice cap. Three categories of sea ice can be distinguished here: Polar Cap Ice, Pack Ice, and Fast Ice (Pickard and Emery 1982). Polar Cap Ice covers about 70% of the Arctic Ocean. It occurs in the vicinity of the Pole near the 1000-m isobath, and consists of ice which is several years old. In winter, the average thickness of undisturbed ice is about 3-4 m, but hummocks can increase the height locally up to 10 m a.s.l. In summer the average thickness decreases to about 2.5 m. The Pack Ice lies outside the polar cap and covers about 25% of the Arctic area. Its areal extent is greatest in May and lowest in September. The Fast Ice grows seawards from the coast to the pack. It is most often anchored to the shore and extends out to about the 20-30-m isobath. The Fast Ice occurs only in wintertime and its thickness reaches 1 to 2 m. Sea ice in the Arctic is continually in motion as a result of the effects of wind, tide, and ocean currents. The same factors create open-water areas known as leads and polynyas. Leads are cracks in the ice which are a few kilometres in width and tens of kilometres long, though which are often short lived. On the other hand, polynyas are large open-water areas in the frozen sea and range in size from a few hundred square meters to thousands of square kilometres. Polynyas appear in winter when the air temperature is well below the freezing point of seawater. The role of open-water areas in the Arctic climate system is sufficiently important to be studied more seriously by climatologists. Through these areas the Arctic surface loses huge amounts of heat because sea-surface temperature in winter can be up to 20°C' higher (as in the case of the so-called North Water polynya in the northern part of Baffin Bay) than that of the surrounding areas and because there is no sea-ice covcr, which significantly reduces the heat exchange between the ocean and the atmosphere, as was mentioned above.

Another type of ice which occurs in the Arctic takes the form of ice-bcrgs and originates as a result of the "calving" of tidewater glaciers. Each year a highly variable number of these navigational hazards (about 1000 across the 55°N latitude) move southward into the Atlantic together with the cold water of East Greenland and Labrador Currents.

ARCTIC WATER

ATLANTIC WATER

BOTTOM WATER

2000

3000

4000

Figure 1.4. Typical temperature and salinity profiles for the Arctic Sea (the Eurasian and Canadian basins) (after Pickard and Emery 1982).

2000

3000

4000

SALINITY °L 30 32 34

ARCTIC WATER

ATLANTIC WATER

BOTTOM WATER

Figure 1.4. Typical temperature and salinity profiles for the Arctic Sea (the Eurasian and Canadian basins) (after Pickard and Emery 1982).

Coachman and Aagaard (1974) distinguished three main water masses in the Arctic Ocean: the surface or Arctic Water from the sea surface to a depth of 200 m, the Atlantic Water from 200 m to 900 m, and the Bottom Water below 900 m. For the study of the Arctic climate knowledge of the Arctic Water is most important and this can be divided into three layers: the Surface Arctic, the Sub-surface Arctic and the Lower Arctic Waters. The physical characteristics of these types of water masses arc provided in Table 1.1 and Figure 1.4. Surface waters extend from the surface to depths of about 25 and 50 m. Both the salinity and temperature of the water is strongly controlled by melting and freezing. As a result, the temperature oscillates near the freezing point of seawater, which varies only from -1.5°C at a salinity of 28%o to -1,8°C at a salinity of 33.5%o. Throughout the year both salinity and temperature show rather small changes, which range up to 2%o and 0.1^).2°C, respectively.

Table 1.1. Arctic Sea water masses (after Pickard and Emery 1982)

Water mass

Properties

Name (circulation direction)

Boundary depth

Temperature (T) and Salinity (S)

Seasonal variation

Surface

ARCTIC SURFACE

25 to 50 m

T: Close to P.P., i.e. -1.5 to -1.9°C S: 28 to 33.5%«

DS: 2

ARCTIC SUB-SURFACE

100 to 150 m

Eurasian Basin -1,6°C to 100 m. then increase S: Both basins 31.5 to 34ifr

(all above masses circulate clockwise)

200 m

Intermediate between Sub-surfaec and Atlantic

ATLANTIC (anticlockwise)

900 m

T: Above 0°C (to 3°C) S: 34.85 to 35%«

Negligible

BOTTOM (uncertain, small)

Bottom

Eurasian Basin -0.8°C -0.6°C S: Both Basins 34.90 to 34.99%»

(rise adiabatic)

The surface water and sea-ice circulation in the Arctic has been largely known from the observed drift of camps on the ice, floe stations, and ships. The earliest information comes from the famous "Frairf drift (1893-1896) and from the icebreaker "Sedov" drift (1937-1940). Observational evidence together with Iheorctical calculations of upper-layer circulation based on water density distribution, give a consistent picture of circulation in the Arctic (Figure 1.5). In the Beaufort Sea the surface waters have a clockwise move-

ment in agreement with the anticyclonic pattern of blowing winds and lead out to the East Greenland Current. From the Eurasian side of the Arctic Ocean the surface waters move towards North Pole and exit the Eurasian Basin as the East Greenland Current. This current is known as the Transpolar Drift Stream. The speeds of these waters are of the order of lem/s to 4 cm/s (300 km/year to 1200 km/year). It is worth to add here that a sea-ice circulation in the Arctic Ocean is similar to the described above circulation of the surface water currents.

Secret Garden Coloring Book

CHUKCHI' SEA^

SHELF

IOMONOSOV RIDGE

ARCTIC

CANADIAN BASIN

EURASIAN ^ BASIN

NOFJTH POLE

SPITSBERGEN

'BAFFIN

GREENLAND / SEA

¡HUDSON

:DAVIS STRj

DEPTH

IRMINGER ff~ CURR

LABRADO«

Figure 1.5, Arctic Sea and North Atlantic adjacent seas: bathymetry and surface currents (after Pickard and Emery 1982).

The circulation of the Atlantic Water is basically counter-clockwise around the Arctic Ocean, i.e. in a direction opposite to that of the Arctic Water above it (Pickard and Emery 1982). The Atlantic Water (West Spitsbergen Current) enters the Eurasian Basin from the Greenland Sea and flows further east along the edge of the Eurasian continental slope. Some waters branch ofT to the north and leave the Arctic as part of the East Greenland Current. The remainder flow across the Lomonosov Ridge into the Canadian Basin. The mixed Arctic (East Greenland Current) and Atlantic (Irminger Current, southwest of Iceland) Waters mass round the southern tip of Greenland and reach the Labrador Sea. Further, they flow as the West Greenland Current to Baffin Bay. This inflow of water is balanced by the southward flow of the Baffin Island Current and Labrador Current. There is also evidence that significant quantities of water of Atlantic origin enter the Arctic Ocean via the Barents and Kara shelves, where they may be considerably modified. Some warm water comes to the Arctic from the Pacific through (he Bering Strait (see Figure 1.5). The principal outflows from the Arctic Ocean arc through the Fram Strait and the Canadian Arctic Archipelago.

According to research conducted by Alekseev et al. (1991 ), the advection of warmth from the lower latitudes supplies more than 50% of the annual heat supply to the Arctic climate system. Most of this warmth (95%) is, however, transported by atmospheric circulation, with the remainder (5%) being transported by oceanic circulation. In winter, during the polar night, only these two fluxes of warmth reach the Arctic and protect it from significant radiation cooling.

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