188.8.131.52 North Atlantic Subpolar Gyre, Labrador Sea and Nordic Seas
In the North Atlantic subpolar gyre, Labrador Sea and Nordic Seas, large salinity changes have been observed that have been associated with changed inputs of fresh water (ice melt, ocean circulation and river runoff) and with the NAO. Advection of these surface and deep salinity anomalies has been traced around the whole subpolar gyre including the Labrador and Nordic Seas. These anomalies are often called ‘Great Salinity Anomalies’ (GSAs; e.g., Dickson et al., 1988; Belkin, 2004). During a positive phase of the NAO, the subpolar gyre strengthens and expands towards the east, resulting in lower surface salinity in the central subpolar region (Levitus, 1989; Reverdin et al., 1997; Bersch, 2002). Three GSAs have been thoroughly documented: one from 1968 to 1978, one in the 1980s and one in the 1990s. Observational and modelling studies show that the relative influence of local events and advection differ between different GSA events and regions (Houghton and Visbeck, 2002; Josey and Marsh, 2005).
These surface salinity anomalies have affected the Labrador Sea and the production of Labrador Sea Water (LSW), a major component of the North Atlantic Deep Water (NADW) and contributor to the lower limb of the MOC. The LSW appears to alternate between dense, cold types and less dense, warm types (Yashayaev et al., 2003; Kieke et al., 2006) possibly with more production of dense LSW during years of positive-phase NAO (Dickson et al., 1996). Since 1965 to 1970, the LSW has had a significant freshening trend with a superimposed variability consisting of three saltier periods, coinciding with warmer water, and two freshening and cooling periods in the 1970s and 1990s (Figure 5.7). During the period 1988 to 1994, an exceptionally large volume of cold, fresh and dense LSW was produced (Sy et al., 1997; Lazier et al., 2002), unprecedented in the sparse time series that extends back to the 1930s (Talley and McCartney, 1982). The Labrador Sea has now returned to a warmer, more saline state; most of the excess volume of the dense LSW has disappeared, the mid-layers became warmer and saltier, and the production of LSW shifted to the warmer type (e.g., Lazier et al., 2002; Yashayaev et al., 2003; Stramma et al., 2004). This warming and increased salinity and reduction in LSW was associated with the weakening of the North Atlantic subpolar gyre, seen also in satellite altimetry data (Häkkinen and Rhines, 2004).
Figure 5.7. The longest available time series of salinity (psu; upper panel) and potential temperature (°C, lower panel) in the central Labrador Sea from 1949 to 2005 (updated from Yashayaev et al., 2003). The dashed lines are contours of potential density (kg m-3, difference from 1,000 kg m-3) and are the same on both panels.
The eastern half of the subpolar North Atlantic also freshened through the 1980s and into the 1990s, but the upper ocean has been increasing in salinity or remaining steady since then, depending on location. About two-thirds of the freshening in this region has been attributed to an increase in precipitation associated with a climate pattern known as the East Atlantic Pattern (Josey and Marsh, 2005), with the NAO playing a secondary role. From 1965 to 1995, the subpolar freshening amounted to an equivalent freshwater layer of approximately 3 m spread evenly over its total area (Belkin, 2004; Curry and Mauritzen, 2005).
Subsurface salinity in the Nordic Seas has also decreased markedly since the 1970s (Dickson et al., 2003), directly affecting the salinity of the Nordic Sea overflow waters that contribute to NADW. This decrease in subsurface salinity was associated with lower salinity of the Atlantic waters entering the Nordic seas and related to the high NAO index and intensification of the subpolar gyre. Since 1994, the salinity of the inflow from the North Atlantic has been increasing, reaching the highest values since 1948, largely due to a weakening of the subpolar gyre circulation that allowed more warm water into the Nordic Seas, associated with a decreasing NAO index (Hátún et al., 2005).
The densest waters contributing to NADW and to the deep limb of the MOC arise as overflows from the upper 1,500 m of the Nordic Seas through the Denmark Strait and Faroe Channel. The marked freshening of the overflow water masses exiting the Arctic was associated with growing sea ice export from the Arctic and precipitation in the Nordic Seas (Dickson et al., 2002, 2003). The transports of the overflow waters, of which the largest component is through Denmark Strait, have varied by about 30% (Macrander et al., 2005), but there has been no clear trend in this location. Overall, the overflows that contribute to NADW from the Nordic Seas have remained constant to within the known variability.
The overall pattern of change in the North Atlantic subpolar gyre is one of a trend towards fresher values over most of the water column from the mid-1960s until the mid-1990s. Since then, there has been a return to warmer and more saline waters (Figure 5.7), which coincides with the change in NAO and East Atlantic Pattern. However, this return to saltier waters has not been sustained for a long enough period to change the sign of the long-term trends (Figure 5.5 Atlantic).
Box 5.1: Has the Meridional Overturning Circulation in the Atlantic Changed?
The global Meridional Overturning Circulation consists primarily of dense waters that sink to the abyssal ocean at high latitudes in the North Atlantic Ocean and near Antarctica. These dense waters then spread across the equator with comparable flows of approximately 17 and 14 Sv (106 m3 s–1), respectively (Orsi et al., 2002; Talley et al., 2003a). The North Atlantic overturning circulation (henceforth ‘MOC’) is characterised by an inflow of warm, saline upper-ocean waters from the south that gradually increase in density from cooling as they move northward through the subtropical and subpolar gyres. They also freshen, which reduces the density increase. The inflows reach the Nordic Seas (Greenland, Iceland and Norwegian Seas) and the Labrador Sea, where they are subject to deep convection, sill overflows and vigorous mixing. Through these processes NADW is formed, constituting the southward-flowing lower limb of the MOC.
Climate models show that the Earth’s climate system responds to changes in the MOC (e.g., Vellinga and Wood, 2002), and also suggest that the MOC might gradually decrease in transport in the 21st century as a consequence of anthropogenic warming and additional freshening in the North Atlantic (Bi et al., 2001; Gregory et al., 2005; see also Chapter 10). However, observations of changes in the MOC strength and variability are fragmentary; the best evidence for observational change comes from the North Atlantic.
There is evidence for a link between the MOC and abrupt changes in surface climate during the past 120 kyr, although the exact mechanism is not clear (Clark et al., 2002). At the end of the last glacial period, as the climate warmed and ice sheets melted, there were a number of abrupt oscillations, for example, the Younger Dryas and the 8.2 ka cold event (see Section 6.4), which may have been caused by changes in ocean circulation. The variability of the MOC during the Holocene after the 8.2 ka cooling event is clearly much smaller than during glacial times (Keigwin et al., 1994; see Section 6.4).
Observed changes in MOC transport, water properties and water mass formation are inconclusive about changes in the MOC strength (see Section 184.108.40.206). This is partially due to decadal variability and partially due to inadequate long-term observations. From repeated hydrographic sections in the subtropics, Bryden et al. (2005) concluded that the MOC transport at 25°N had decreased by 30% between 1957 and 2004, but the presence of significant unsampled variability in time and the lack of supporting direct current measurements reduces confidence in this estimate. Direct measurements of the two major sill overflows have shown considerable variability in the dominant Denmark Strait Overflow without enough years of coverage to discern long-term trends (Macrander et al., 2005). The observed freshening of the overflows and the associated reduction in density from 1965 to 2000 (see Section 5.3.2) has so far not led to a significant weakening of the MOC (Dickson et al., 2003; Curry and Mauritzen, 2005). Moreover, large decadal variability observed since 1960 in salinity and temperature of the surface waters, including the recent increase in salinity of the surface waters feeding the MOC, obscures the long-term trend (Hátún et al., 2005; ICES 2005) and hence conclusions about potential MOC changes.
Changes in the MOC can also be caused by changes in Labrador Sea convection, with strong convection corresponding to higher MOC. Convection was strong from the 1970s to 1995, but thereafter the Labrador Sea warmed and re-stratified (Lazier et al., 2002; Yashayaev et al., 2003) and convection has been weaker. Based on observed SST patterns, it was concluded that the MOC transport has increased by about 10% from 1970 to the 1990s (Knight et al., 2005; Latif et al., 2006). From direct current meter observations at the exit of the subpolar North Atlantic, Schott et al. (2004) concluded that the Deep Water outflow, while varying at shorter time scales, had no significant trend during the 1993 to 2001 period.
In summary, it is very likely that up to the end of the 20th century the MOC was changing significantly at interannual to decadal time scales. Given the above evidence from components of the MOC as well as uncertainties in the observational records, over the modern instrumental record no coherent evidence for a trend in the mean strength of the MOC has been found.