The changes in mean SLP fields considered above also influence the corresponding changes in the general ice-drift pattern.
In Anon. (F) and Gudkovich and Doronin (2001), the authors constructed the mean multiyear ice-drift pattern in the Arctic Basin, which expresses the most probable ice motion during a prescribed time interval by approximating a field of ice-drift vectors using a two-dimensional polynomial of the form:
where x, y represent the coordinates of the initiation of drift vectors Vxy; m is the degree of polynomial with argument p; t is the degree of polynomial with argument q; and apq is the polynomial coefficient.
When orthogonal components obtained from observations of ice-drift vectors of specified temporal scales (seasonal, semi-annual, and annual) are included in Equation 4.1, they form systems of conventional equations, which can be combined to create expanded matrices of normal equations. Their solution by the least-squares method allows deriving the polynomial coefficients for each component. For m = t = 3, there are 16 coefficients. The derived equations are used to calculate drift vectors in any regular grid within the region observed. An advantage of constructing ice-drift patterns with this method is improved spatial resolution, which is typical of calculations made by hydrodynamic models with the high reliability usually inherent in patterns based on full-scale observations.
For constructing the ice-drift patterns in the studies discussed above, we employed data from various observing platforms (drift of ship-based expeditions, "North Pole (NP)'' drifting stations, ice islands, automatic stations, and radio buoys) through 1976. Most of these observations were made between 1954 and 1975, i.e., they refer to the cooling period that replaced the Arctic warming period of the 1920s-1940s (Zakharov, 1976). A new warming period began in the late 1970s and continued into the beginning of the twenty-first century. Alternation of the cooling and warming periods appears to be controlled by climatic cycles lasting about 60 years.
Using observation data on NP drifting station movement and data from "Argos" buoys, we calculated (Equation 4.1) ice-drift patterns in the Arctic Basin for the 1980-2004 warming period (Gudkovich et al., 2007). Comparing this pattern with a pattern obtained earlier illuminated the main changes that occurred in the ice-drift field at the transition of the climatic system from cooling to warming. Following the practice of Gudkovich and Doronin (2001), we investigated ice-drift vector fields for semiannual periods (October-March and April-September), using the initiating coordinates at 82°30'N, 180°E. The abscissa axis is directed along the 180° meridian to the south, and the ordinate axis is parallel to 270° E.
Figure 4.12a and b show the ice-drift patterns obtained by the method described above for summer and winter of the climate warming epoch that spanned the end of the twentieth and the beginning of the twenty-first century. A comparison of the patterns depicted in this figure with similar patterns published in Gudkovich and Doronin (2001), which mainly characterize the cooling period, shows no significant differences. Similar to the cooling epoch, an increase in ice-drift speed is observed during the last warming period at approaches to Fram Strait in winter, in the displacement of the transarctic flow core from Eurasia to America, and in a decrease in the area of the Beaufort anticyclonic gyre from winter to summer. The average modules of ice-drift speed for monthly and half-year time periods (3.5 and 2.5 cm/s, respectively) practically coincide.
The main differences occur in the patterns (Figure 4.12c, 4.12d) that show the differences in resulting ice-drift vectors during the warming and cooling epochs. Both summer and winter patterns clearly exhibit a cyclonic character in the vector difference fields, indicating increased cyclonicity when the cooling epoch is replaced by the warming epoch. This is expressed more strongly in summer than in winter. A decrease in Arctic atmospheric pressure during warming epochs is confirmed by the pressure charts averaged for the corresponding periods in Figure 4.10. This phenomenon
provides a very good expression of a zonal transfer index in the atmosphere of temperate latitudes (from 40° to 65°N), similar to the Blinova (Blinova, 1943) index (Figure 4.9).
An increase in the recurrence of cyclonic pressure fields over the Arctic Basin at the transition from a cooling to a warming epoch leads to changes in ice cover deformation processes. As shown in Gudkovich and Doronin (2001), Busuyev et al. (1999), Porubayev (2000), Gudkovich and Klyachkin (2001), and Losev et al., (2005), the cyclonic systems of the multiyear ice drift contribute to ice cover divergence. This process is most prevalent in summer, whereas in winter, especially in relatively thin ice zones, ice compacting is usually observed. Anticyclonic SLP fields have the opposite effect. This is confirmed by mathematical modeling of ice cover dynamics (Shoutilin et al., 2005), showing that low-frequency changes in the state of multiyear ice in the Arctic Basin are accompanied by the processes of divergence, compacting, and ridging, which influence the corresponding changes in medium drifting ice. Divergence of multiyear ice in the Arctic Basin during warming is confirmed by experimentally detected gradual displacement of the multiyear ice boundary and its area. This is discussed in section 4.5.
Figure 4.12d shows that the differences in ice-drift vectors in the Fram Strait area in winter are directed from the Greenland Sea to the Arctic Basin. This provides further support for section 4.2 evidence that ice export from the Arctic Basin to the Greenland Sea during warming is weaker than during cooling. Section 4.4 presents additional arguments confirming a direct relationship between the ice cover area and ice export through Fram Strait.
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