Greenland

East Greenland Current

^ICELAND

Irminger Sea <

C taille Gibbs Fracture Zone

Figure 6.21 (a) Temperature section across the Denmark Strait between Greenland and Iceland, at about 65° N. The Denmark Strait is the deepest overflow channel, and so this is the largest outflow, (b) Sections across the Greenland continental slope at about 63° N, showing the overflow downstream of the Denmark Strait, as indicated by (i) isopycnals (for contour values see p. 230) and (ii) long-term mean current speed (m s~1). The flow in (b) consists of the Denmark Strait overflow, plus the overflows from east of Iceland (cf. Figure 6.20). Note the different vertical and horizontal scales in (a) and (b).

The paths followed by the densest, least-mixed overflow water are shown on Figure 6.20 in olive green. The overflow water from east of Iceland follows the topography around to the west, passes through the Charlie Gibbs Fracture Zone at -55° N and combines with the largest overflow from the Denmark Strait (see Figure 6.21(a)). Because of the effect of the Coriolis force, the overflow water 'hugs the topography', with the isopycnals sloping up steeply to the right of the flow (Figure 6.21(b)(i)).

Deep water may be added in the Irminger Sea off southern Greenland, where winter cooling of homogeneous Sub-Arctic Upper Water (Figure 6.15) leads to deep convection from surface to bottom. However, much more significant is the addition of dense water in the Labrador Sea, to form North-West Atlantic Deep Water, or western North Atlantic Deep Water (the lighter of the broad green arrows in Figure 6.20), which flows south in the western basin of the Atlantic and also spreads eastwards to overlie the denser North-East Atlantic Deep Water in the eastern basin.

The influence of Labrador Sea Water on North Atlantic Deep Water ma\ he observed in northern and western parts of the North Atlantic. What other intermediate water mass has a significant influence on the characteristics of North Atlantic Deep Water, particularly on the eastern side ol the ocean?

Mediterranean Water. As described earlier, the least-mixed layer of Mediterranean Water is generally found at a depth of about 1000 m (Figure 6.14). but its influence extends down to more than 2000 m.

Figure 6.21 (a) Temperature section across the Denmark Strait between Greenland and Iceland, at about 65° N. The Denmark Strait is the deepest overflow channel, and so this is the largest outflow, (b) Sections across the Greenland continental slope at about 63° N, showing the overflow downstream of the Denmark Strait, as indicated by (i) isopycnals (for contour values see p. 230) and (ii) long-term mean current speed (m s~1). The flow in (b) consists of the Denmark Strait overflow, plus the overflows from east of Iceland (cf. Figure 6.20). Note the different vertical and horizontal scales in (a) and (b).

40 km

J ote ran n nal variability in formation of the components of NADW

As discussed in Section 6.3.1. there is great variability from year to year in the depth of mixing in the Labrador Sea. and in the amount of Labrador Sea Water formed. Occasionally. Labrador Sea Water forms to the west of the gyre, and can enter directly irito the deep western boundary current (sec DWJJC on Figure 6.20) without being caught up in the gyral circulation. However, it usually circulates at depth around the Labrador Sea a number of times, so thai marked interannual variations in the signature of Labrador Sea Water in North-West Atlantic Deep Water are somewhat "smoothed out',

Nevertheless, over the last few decades, océanographers have been noticing significant fluctuations in both the contribution of Labrador Sea Water to North-West Atlantic Deep Water and (he rale of production of deep water in the Greenland Sea. Indeed, i I seems that there may be a 'see-saw' relationship between convection in the two locations, iti that as one increases the other decreases, and vie? versa. These fluctuations may he pan of changes related to the North Atlantic Oscillation (Section 4.5). Around the early 1990s, the NAO indes. was positive (Figures 4.40 and 4.411. and the Labrador Sea was subjected to unusually cold winds and winter storms, with the result that mixing there became deeper, and large amounts of Labrador Sea Water were formed (cf Figure ft. 17): at the same time, the Norwegian and Greenland Seas became wanner and less stormy, and convection became less intense and less deep, and the rale of formation of deep water there declined. Not surprisingly, there was a decline in the rate Of formation of Labrador Sea Water m the early 1970s, when the Great Salinity Anomaly passed through the north-vvesi Atlantic (Figure 5,28).

As will be discussed in Section 6A. North Atlantic Deep Water plays a crucial role in the global climate system, so it is important to know how it is affected by changes in the rates of production of deep water in the Greenland and Labrador Seas. Ocea nog ra pliers have been investigating the variation in the volume of dense water overflowing the Greenland-Scotland Ridge, and although (as mentioned earlier) individual overflow events seem to he triggered by meteorological changes over the Arctic region, (here is evidence that during the course of the I%0s to 1990s, the rale at which ilense water was spilling over into the deep Atlantic approximately doubled, and then returned in its previous value.

Antarctic Bottom Water

Antarctic Botlom Water (A ABW) is the most widespread water mass in the world and is found in ail three ocean basins, particularly m their southern parts (Figure 6.22), It seems to have two separate source areas -the continental shelf around the continent of Antarctica and deep levels in the Antarctic Circumpolar Current (cf. Figure 5.30).

Before we discuss how Antarctic Bottom Water forms, we should emphasize an important aspect of the stratification of polar waters, in the Arctic as well as the Antarctic, Except where ice is actually forming, surface w ater is relatively fresh, because of the combined effects of excess precipitation and the production of melt water: ihis surface water is aiso very cold. Below the surface layer there is generally a layer of relatively warm water which is nevertheless denser because of its relatively high salinity. The stability of this layering may he destroyed through turbulent mixing caused by wind, combined w ith an increase in density of the surface waters. In northern polar latiludes. the density of surface waters is increased source regions

Arctic Deep Water

AA8W PIOCW Pacific and Indian Ocean Common Water

A 000 m depth contour

Figure 6.22 The global distribution of deep and bottom water masses (between a depth of about 1500 m and the sea-floor). At high southern latitudes, the distribution of Antarctic Bottom Water is shown schematically; for a more realistic picture, see Figure 6.25. The source regions are shown by dark blue tone. The fine dashed line is the 4000 m isobath. (The unlabelled regions are to a large extent occupied by Pacific and Indian Ocean Common Water - see text.)

through winter cooling by cold winds (cf. Figure 6.7(b)) and. particularly around the Arctic basin, through ice-formation. In the Southern Ocean, where the seasonal production of ice is more extensive (Figure 5.27). the interaction between ice and surface waters plays an even greater role in the formation of dense water.

Ice-water interaction is perhaps best illustrated through discussion of polvnyas. extensive areas of ice-free water within the winter ice cover. In the Antarctic region, there are two types of polynya: coastal polynyas and open ocean polynyas. Coastal polynyas develop when strong winds blowing off the Antarctic continent drive newly formed ice away from the shoreline, exposing a zone of open sea that might be 50-100 km wide. Open-ocean polynyas develop far from the coast within the pack ice (both over the continental shelf and in deeper water), and include the largest and most long-lived polynyas. The Weddell Polynya was an enormous area of open ocean that appeared in the Weddell Sea during three consecutive winters (1974-76): at its largest it measured about 1000 km by 350 km. The Weddell Polynya reappeared in approximately the same position each year (Figure 6.23) above a sea-bed topographic high, known as the Maud Rise. This implies that its formation was related to the upward deflection of relatively warm subsurface water, and in fact this is a possible explanation for a number of other open-ocean polynyas.

Recurrent polynyas are also found in the Arctic ice cover, but do not seem to be associated with sinking of dense surface water. In the Antarctic, both coastal and open-ocean polynyas are sites where surface water may be sufficiently cooled for deep convection to occur.

We stated earlier that water in contact with ice is cooled by conduction, until a new heat balance is attained, \Vh>. then, are polynyas - regions with no ice cover - characterized by significant heat loss ?

Figure 6.23 Nimbus satellite images showing the position of the Weddell Polynya during the southern winters of (a) 1974, (b) 1975 and (c) 1976. The westward movement of the polynya has been attributed to the generally westward current flow in the region (cf. Figure 5.30). The light blue tone corresponds to regions that are ice-free, while the purplish tone corresponds to regions that are more or less completely covered with ice. The Antarctic continent has been superimposed in white.

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Figure 6.23 Nimbus satellite images showing the position of the Weddell Polynya during the southern winters of (a) 1974, (b) 1975 and (c) 1976. The westward movement of the polynya has been attributed to the generally westward current flow in the region (cf. Figure 5.30). The light blue tone corresponds to regions that are ice-free, while the purplish tone corresponds to regions that are more or less completely covered with ice. The Antarctic continent has been superimposed in white.

A layer of ice 'insulates' the ocean from the atmosphere and significantly reduces heat loss through conduction/convection, Qh (as well as preventing heat loss through evaporation, Qc). When there is no ice cover, heat losses to the atmosphere are greatly increased, especially when cold winds blow over the sea-surface.

In fact, the main mechanisms for heat loss to the atmosphere are different for the two types of polynya. Coastal polynyas have been described as 'sea-ice factories': the wind drives sea-ice away from the continent as soon as it freezes, re-exposing the sea-surface to the atmosphere so that more ice can form (Figure 6.24). The continued production of ice removes large amounts of heat from surface water, mainly in the form of latent heat of freezing. This is the heat lost to the atmosphere by water while it remains at freezing point (-1.9 °C for seawater) as ice crystals are in the process of forming, and is analogous to the latent heat of condensation released when water

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