Thermohaline Circulation Of The Oceans

Circulation of water at the ocean surface is largely a response to the overlying atmospheric circulation which exerts drag at the surface. However, circulation of deeper waters in the oceans of the world is a consequence of density variations, which result from differences in temperature and salinity brought about by sensible and latent heat fluxes, precipitation, and runoff at the ocean surface; this is termed the thermohaline circulation. In areas where surface waters become relatively dense, due to cooling and/or evaporation (thereby increasing salinity) they will sink to the level at which they reach equilibrium (neutral buoyancy) with surrounding water masses.21 Also, areas of sea ice formation, which result in salt expulsion from the ice, produce brine

21 A water mass, like an air mass, is recognized by its distinctive physical properties (principally temperature and salinity), which enable it to be distinguished from adjacent water masses. As a water mass moves from its source region it will slowly mix with other waters and gradually lose its original identity.

which increases the water density, causing water to sink and form a cold dense water mass. Dense water masses flow away from their source regions as either Bottom Waters or Intermediate Waters, depending on their relative density. Much of the deepest sections of the world's oceans are filled by dense Antarctic Bottom Water (AABW), which originates from areas of sea-ice formation adjacent to the Antarctic continent. Antarctic Bottom Water (AABW) is characterized by a temperature of —0.4 °C and a salinity of ~34.7%o. In the Atlantic Ocean, much of the deepest sections (>2 km) are occupied by North Atlantic Deep Water (NADW), which forms mainly in the Norwegian Sea (-60° N, east of Iceland) and in the Greenland Sea (north and west of Iceland) (Kellogg, 1987; Hay, 1993). Deepwater does not form at present in the North Pacific, where waters are less saline than in the North Atlantic.

Overlying these dense water masses (at ~1 km depth) are Intermediate Waters, which generally have slightly lower salinities and/or higher temperatures. Much of the world ocean is occupied by Antarctic Intermediate Water (AAIW), which has a temperature of 2-4 °C and a salinity of ~34.2%o and originates in the circum-Antarctic polar frontal zone. At high latitudes of the North Atlantic, intermediate water from the Labrador Sea (3-4 °C, 34.92%o) is found (sometimes referred to as Upper North Atlantic Deep Water, or Northwestern Atlantic Deep Water) and farther south saline water flowing from the Mediterranean Sea can also be traced at intermediate levels.

Changes in deepwater circulation are important for paleoclimatology because, as a result of deepwater formation and compensating fluxes of water in the upper mixed layer, large quantities of heat are carried around the globe. Of particular significance is the thermohaline circulation associated with the formation of North Atlantic Deep Water (Dickson and Brown, 1994). As noted earlier, NADW forms north of -60° N as surface waters cool (by evaporation and sensible heat loss) and their salinity increases, thereby creating a dense water mass that moves southward at depth (Fig. 6.43) and carries saline water to the South Atlantic and other ocean basins (Fig. 6.44). The loss of water in this manner is compensated for by the poleward movement of warm, saline surface waters to the North Atlantic, in the Gulf Stream and associated North Atlantic Drift. These water masses are responsible for the relatively mild temperatures western Europe experiences, even in winter.

The movement of warm salty water to high latitudes of the North Atlantic, the formation of dense NADW, and the displacement of water as this water mass exits from the North Atlantic can be considered as a linked system or conveyor belt; disturbances to the system may cause it to change its speed (rate of water exchange) or even to cease operation altogether (Broecker, 1991). In fact, models indicate that the system is quite sensitive to disturbance, particularly by freshwater inflows into the North Atlantic (Manabe and Stauffer, 1988; Rahmstorf, 1994; Weaver and Hughes, 1994). Currently, the North Atlantic Basin loses slightly more freshwater via evaporation than it gains from either precipitation or river runoff (-1.21 m a1, v. +0.87 m and 0.21 m, respectively). It is this fact, together with a flux of saline Gulf Stream water and strong cooling (especially in winter) that leads to NADW formation. However, if the freshwater flux were to increase (as it may have when the major continental ice sheets melted) it would create an upper, low salinity water layer that

Nadw And Aabw And Miw





Subtropical Convergence

Antarctic Polar Frontal Zone

Antarctic Divergence i_

Subtropical Convergence

Thermohaline Circulation

FIGURE 6.44 Schematic diagram showing the thermohaline ("conveyor belt") circulation. Near surface waters are shown by the darker shaded arrows, deepwater by the lighter colored arrows. Sinking of surface waters occurs in the North Atlantic, and to a much lesser extent in the North Pacific and Indian Ocean. Estimates of the volume of water transported in each section are shown in the circles (in Sverdrups; I sv = I06 m3 s"'). (Schmitz 1995).

FIGURE 6.43 Meridional cross section of the Atlantic Ocean showing the principal water masses and their distribution today. NADW = North Atlantic Deep Water; AABW = Antarctic Bottom Water; Al W = Atlantic Intermediate Water; AAIW = Antarctic Intermediate Water; M = Mediterranean Intermediate Water. During the last glacial maximum, NADW was more limited in extent, whereas AABW penetrated farther north into the deep basins of the North Atlantic (see Fig. 6.45) (adapted from Brown et ai, 1989).

FIGURE 6.44 Schematic diagram showing the thermohaline ("conveyor belt") circulation. Near surface waters are shown by the darker shaded arrows, deepwater by the lighter colored arrows. Sinking of surface waters occurs in the North Atlantic, and to a much lesser extent in the North Pacific and Indian Ocean. Estimates of the volume of water transported in each section are shown in the circles (in Sverdrups; I sv = I06 m3 s"'). (Schmitz 1995).

would disrupt this process, and shut-down NADW formation. This in turn would "turn-off" the conveyor belt of NADW, eventually leading to a reduction of Gulf Stream water to replace that lost through sinking in the North Atlantic. This would mean less heat transported into the North Atlantic and generally colder climatic conditions leading to less freshwater runoff. Eventually the process of NADW formation would be restored, turning the conveyor back on and allowing the overall circulation to revert back to its former condition. Thus, the North Atlantic thermohaline circulation can be thought of as having two distinct modes ("conveyor on" or "conveyor off") controlled by the relative balance of freshwater flux to the surface waters of the North Atlantic Basin (Broecker et al., 1985b; Broecker and Denton, 1989; Broecker et al., 1990a, b; Broecker, 1994). When the conveyor is off, the rate of salt export from the North Atlantic (in NADW) is less than the rate of salt buildup resulting from evaporation and water vapor export to adjacent regions. Salinity gradually increases until some critical density threshold is reached, at which point the conveyor switches on, which brings more saline water to the North Atlantic via the Gulf Stream. Providing that meltwater flux to the North Atlantic is less than the salt build-up, NADW will continue to form. However if meltwater and/or salt export exceeds that threshold, NADW formation will be greatly reduced or eliminated altogether (Broecker et al., 1990a). In this way the ocean-atmosphere-cryosphere systems are in dynamic equilibrium, in which disturbance of one part of any system may lead to a nonlinear response in another system (Broecker and Denton, 1989).

In a reassessment of this model, Boyle and Rosener (1990) question whether the coupled system really has only two modes or whether in fact there have been multiple stable circulation patterns as suggested by model simulations (Rahmstorf, 1994, 1995). They suggest that rather than being controlled by an "on-off switch" the system may be considered as responding more to a "valve" whereby there are many possible quasi-stable circulation states. Unfortunately, the resolution of deep-sea sediments is frequently insufficient (because of low sedimentation rates and biotur-bation) to resolve which of these models is correct, though there is evidence that when NADW was not formed (e.g., in the Last Glacial Maximum) a distinct North Atlantic Intermediate Water was being produced (Boyle and Keigwin, 1987). This suggests an additional scenario, perhaps the result of some balance of factors neither at one extreme nor the other. Lehman and Keigwin (1992a, b) make the argument that the formation of deepwater in the Norwegian Sea ("Lower NADW") was disrupted quite often in Late Glacial times (see Section 6.10.2) but deepwater from the Labrador Sea ("Upper NADW") continued to form. Another scenario is proposed by Veum et al. (1992), who suggest that deepwater, formed by brine rejection during sea-ice formation, formed in the marginal ice zone of the Greenland-Iceland-Norwegian Seas throughout the Last Glacial Maximum, ventilating the deep basins of this region. By contrast, no deepwater formed to the south in the open North Atlantic Ocean, the deep basins of which were occupied by AABW until -12,600 yr B.P. There may thus be many different states of deepwater circulation that have developed over time, with complete shutdown of both Lower and Upper NADW being the extreme end-member state of a whole range of possible conditions.

Before discussing further the evidence for such changes, it is first necessary to consider the means by which changes in deepwater circulation can be identified. Each water mass has certain geochemical characteristics that can be identified by analysis of benthic forams living in those waters. The geochemistry of benthic forams in marine sediments thus serves as a "tracer" of deepwater conditions at the time the forams were deposited. Of particular importance is the 13C/12C ratio (813C) in the carbonate tests, derived from dissolved C02 in the water column. The 813C value for the atmosphere is -7.2%0; because of fractionation effects, seawater in equilibrium with the atmosphere has a 813C value of ~+3.5%o (at 2 °C) (Mook et al., 1974). By contrast, organic matter has a 813C of -20 to -25%o. Oxidation of organic matter falling through the water column therefore causes the 813C of the water to decline. If surface waters are nutrient-rich and productive, the large input of organic material to the deep ocean will result in low 813C and reduced oxygen levels. The low 813C is balanced to some extent by dissolution of the carbonate tests of planktic forams as they fall through the water column, because these have a 813C value close to that of total dissolved C02 in the upper water column. Hence the overall 813C of a water mass reflects a balance between the amount of organic matter oxidized, and dissolution processes. Nevertheless, the global distribution of 813C strongly reflects the nutrient content and organic productivity of the water mass (Kroopnick, 1985). For example, Antarctic waters are nutrient-rich and productive, resulting in deep-water (AABW), which is depleted in 13C; North Atlantic deepwater, on the other hand, has lower nutrient levels, is less productive, and has a higher S13C value (Duplessy and Shackleton, 1985). These characteristics, preserved in the tests of benthic forams, can be used to trace the presence and distribution of NADW and AABW over glacial-interglacial cycles (Curry et al., 1988; Raymo et al., 1990). Furthermore, in areas of deepwater formation, the vertical distribution of 813C is fairly homogeneous (due to convective mixing) so that the 813C signal in the calcareous tests of both benthic and planktic foraminifera are similar. With increasing distance from areas of deepwater formation, the difference between surface and deepwater 813C increases, and this is reflected in the tests of the deep-dwelling and surface forams (Duplessy et al., 1988). Hence 813C can be used to identify changes in areas of deepwater formation, and to track their movement over time.

There are two slight complications to this neat approach. First, not all benthic foram species record the same values of 813C, apparently due to a species-dependent habitat effect; this can be resolved by selecting only benthic forams (such as Cibici-doides wuellerstorfi) that do not show such an effect, or by using those whose effect is known (Zahn et al., 1986). Second, the global mean 813C level decreased by 0.3-0.4%o during glacial times due to a reduction in terrestrial biomass, and a remobi-lization of 813C-depleted organic material on the exposed continental shelves (when sea level was up to 120 m lower) (Boyle and Keigwin, 1985; Duplessy et al., 1988; Curry et al., 1988; Keigwin et al., 1994). This affects all records equally so it is easily accommodated; 813C thus provides a valuable proxy indicator of water masses in the past (Fig. 6.45).

Another useful tracer of deepwater is the Cd/Ca ratio in benthic foram tests (Boyle and Keigwin, 1982; Boyle, 1988). Cadmium is a proxy for oceanic nutrient


Thermohaline Circulation Shutdown





FIGURE 6.45 Cross section through the Atlantic Ocean showing the distribution of 8I3C today (GEOSECS data) and in benthic foraminifera from the last glacial maximum (LGM).The 8I3C is used to characterize particular water masses; thus the lowest values are indicative of deep water produced in the sub-Antarctic (AABW). At the LGM, NADW (which has higher 8I3C values) penetrated only to intermediate depths, whereas today it sinks to greater depths and occupies most of the deep Atlantic basins, as far south as the Equator (see Fig. 6.43). Changes in this and other tracers of deepwater circulation reflect important differences in the thermohaline circulation over time (Duplessy and Maier-Reimer, 1993).

levels; Antarctic Bottom Waters have relatively high Cd levels compared to NADW so Cd/Ca provides a useful index of these water masses (Boyle, 1992). Although Cd levels in the world ocean were higher during glacial time, sediments from the Bermuda Rise show that during the last glacial maximum (isotope stage 2) and during the Younger Dryas interval, Cd levels in the deep Atlantic Ocean increased even more, indicating that NADW flux was reduced and replaced by AABW at those times (Boyle and Keigwin, 1987). This supports the 813C data, which also points to a shift in deepwater circulation towards increased AABW flux (see Fig. 6.45) not only in the last glacial but also in isotope stage 6 (135 ka B.P.) and earlier glacial periods (Duplessy and Shackleton, 1985; Boyle and Keigwin, 1985; Oppo and Fairbanks, 1987; Curry et al., 1988; Raymo et al., 1990). The fairly rapid changes in deepwater recorded during the late glacial/Younger Dryas interval clearly link the well-documented changes in surface oceanic conditions and climate around the North Atlantic (mainly in western Europe) with deepwater variations.

Another important aspect of deepwater formation relates to the transport of oxygen into the deep ocean basins of the world. Water at the surface is generally well oxygenated but as deepwaters form and sink, oxygen levels decline as oxidation of organic matter falling through the water column proceeds. In effect, deep-water formation ventilates the deep ocean by carrying oxygenated water to great depths. Ventilation rates can be estimated by measuring the radiocarbon content of the water; once the water is isolated from the atmosphere, radiocarbon is no longer in equilibrium with the atmospheric reservoir and 14C levels will decline. The radiocarbon age of deepwater therefore reflects the time since isolation from the surface; this obviously varies from one area to another (see Fig. 3.7) but typically deepwater in the Atlantic has a 14C age of -400 yr, in the Indian Ocean -1200 yr, and in the Pacific Ocean, ~1600yr.22 Broecker et al. (1988a) compared planktic and benthic forams from the last glacial maximum (LGM) in Atlantic and Pacific Ocean sediments. On average, the discrepancy between ages of forams in the upper and lower water columns increased during the LGM, from 400 to 600 yr in the Atlantic and from 1600 to 2100 yr in the Pacific, indicating that ventilation rates were significantly lower in glacial times compared to the present. A similar conclusion was reached by Bard et al. (1994), who found evidence for reduced North Atlantic ventilation during the cold Younger Dryas oscillation.

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