Formally defining thermohaline circulation is difficult and numerous definitions exist within the literature. Historically, thermohaline circulation has been defined as the component of oceanic circulation driven by fluxes of heat and freshwater (sometimes combined into a buoyancy flux) through the ocean's surface. This particular definition of thermohaline circulation is prevalent among ocean modelers, wherein ocean models driven exclusively by boundary conditions on heat and freshwater, with wind forcing set to zero, lead to a global-scale meridional overturning.
However, in many cases, anomalies of the surface buoyancy flux, and particularly its thermal component, strongly depend on the thermohaline circulation itself. For example, an intensified meridional overturning cell would lead to stronger poleward heat transport in the ocean. At steady state, this heat transport anomaly must be balanced by more heat loss to the atmosphere over the regions towards which the heat is being transported, and by more heat gain by the ocean over regions from which the heat is being transported (Figure T1).
A further problem with the above definition arises when one tries to separate the thermohaline ocean circulation from the wind-driven circulation. For example, turbulent fluxes (latent and sensible heat fluxes) represent a sizable part of the net heat flux and are highly dependent on the near surface wind speed. Furthermore, evaporation, which represents a negative part of the total freshwater flux into the ocean is essentially given by the latent heat loss, and hence is also proportional to the strength of surface winds. In addition, ocean models and the ocean component of coupled models typically specify a fixed value of diapycnal mixing (mixing of heat and salt across surfaces of constant density in the ocean). This representation of diapycnal mixing is a parametrization of internal wave breaking which in turn originates from wind (and tidal) forcing.
Thermohaline circulation is also defined as the densitydriven global-scale oceanic circulation. This too has serious shortcomings, as density gradients in the ocean are largely set up by the winds. In particular, the density structure in subtropical and subpolar gyres is determined by Ekman dynamics. The wind stress anomalies themselves are functions of sea surface temperature. Finally, in the ocean, which is typically in geostrophic balance over the spatial and temporal scales of interest to climate and paleoclimate researchers (100s of km and years, respectively) one cannot separate the existence of the density gradients from the existence of the currents.
Wunsch (2002) defined the thermohaline circulation as the circulation of temperature and salt. This definition, however, is not one that is readily accessible to paleoclimatologists. A more practical approach is to move away from the use of the term thermohaline circulation and to focus on the three dimensional (3D) global overturning circulation (GOC). For many paleoclimatological applications, a sufficient representation of the GOC is in terms of the meridional overturning circulation (MOC) and its manifestation in the Atlantic as the Atlantic meridional overturning (AMO), which represent mass transport in the 2D meridional/vertical plane (Figure T1). Inherent in this approach is the acknowledgement that wind and buoyancy forcing are inseparable, and that wind and tidal forcing play a fundamental role in diapycnal mixing within the ocean.
Here, we focus our discussion on the MOC as a fundamental diagnostic in understanding the role of the ocean in past, present and future climate change. Furthermore, sudden changes in the strength of the MOC, and in particular the AMO, are associated with the abrupt climate change prevalent in high resolution proxy records over the last glacial cycle. The importance of the AMO to climate lies in its association with much of the total oceanic poleward heat transport in the present-day Atlantic, peaking at about 1.2 ± 0.3 PW (1 PW equals 1015 Watts) at 24° N. Since the MOC represents the ocean circulation in the meridional-vertical plane, its existence and structure is fundamentally connected with the locations of deep water formation in the ocean.
Two distinct forms of deep water formation occur: nearcontinent and open ocean. The former involves one of two simple processes: evaporation, or more typically brine rejection, above a continental shelf produces dense heavy water which sinks down and along the slope under the combined forces of gravity, friction and the Coriolis force (Killworth, 1983); alternatively supercooled water may be formed at the base of a thick ice shelf during freezing or melting and this dense water may in turn flow down-slope. By contrast, open ocean convection is observed in areas remote from land and is characterized by a large-scale cyclonic mean circulation which causes a doming of the isopycnals and weakens the static stability over an area tens of kilometers wide. The convection itself is short-lived, is restricted to a narrow circumference and is typically driven through intense surface cooling. These narrow convective chimneys are almost purely vertical and usually do not entrain large volumes of surrounding water as in near-continent convection.
The two main constituent water masses of the deep North Atlantic Ocean - North Atlantic Deep Water at the bottom and Labrador Sea Water at an intermediate level - are currently formed in the Greenland-Iceland-Norwegian Seas and the Labrador Sea, respectively. Deep convection also occurs at a number of locations around Antarctica (Adéle Coast, Amery Ice Shelf, Ross Sea, and Weddell Sea), but the dense bottom water is susceptible to being trapped by topographic sills (as in the Bransfield Strait), or by local circulation patterns (not excluding the Antarctic Circumpolar Current - ACC). In the Southern Ocean, around the southern tip of South America, an enhanced formation of low salinity Antarctic Intermediate Water (AAIW) also occurs.
In today's climate, there are no sources of deep water in the North Pacific Ocean, although North Pacific Intermediate Water (NPIW), characterized by a salinity minimum in the subtropical gyre at depths of 300-800 m within a narrow density range, is locally produced. This water mass is mainly constrained to the subtropical gyre, unlike the Labrador Sea intermediate water mass which is largely confined to the subpolar gyre of the North Atlantic. The original source water for NPIW is thought to be sinking in the Sea of Okhotsk.
Deep convection also occurs in the Mediterranean Sea, but the dense salty outflow that exits into the Atlantic gains buoyancy by mixing with lighter environmental water and spreads out at mid-depth rather than sinking to the bottom. Similar, but weaker, saline outflows have been detected issuing at mid-depth from the Red Sea and Persian Gulf.
In summary, two main bottom water masses exist whose trajectories can be traced by their temperature and salinity characteristics throughout the rest of the world ocean. These are (see Figure T1): (a) the component of Antarctic Bottom Water (AABW) which is largely produced in the Weddell Sea before mixing with Circumpolar Deep Water in the ACC and then flowing into the major ocean basins; (b) North Atlantic Deep Water (NADW), which lies above the AABW at all latitudes except north of 40° N in the North Atlantic. The North Pacific Ocean, although a source of North Pacific Intermediate Water is not a source of bottom water.
In addition, as discussed below, AAIW plays a critical role in linking the Pacific and Atlantic Oceans and in particular on the stability of the AMO. In the present climate, about 13 Sv (1Sv = 106 m3 s-1) of cold, fresh Antarctic Intermediate Water (AAIW) flows into the South Atlantic, about 8 Sv of which is converted into thermocline water through mixing by 32° S (Rintoul, 1991). Further mixing converts the remaining 5 Sv of AAIW into thermocline water as it approaches the equator (Schmitz and McCartney, 1993; Table T1). At 24° N there is about 13 Sv of thermocline water which flows northward into the high northern North Atlantic where it is converted into NADW. AABW flows into the South and North Atlantic where it is modified into and exported as NADW into the Southern Ocean.
|Water type||32° S||24° N|
|Intermediate water (AAIW)||5||-|
|Deep water (NADW)||-17||-18|
|Bottom water (AABW)||4||5|
Sea water density is a function of not only temperature, but also salinity. Furthermore, at low temperatures, such as those that exist in the deep water formation regions, the density of sea water is more sensitive to changes in salinity than to changes in temperature. In addition, throughout the Quaternary, large quantities of freshwater have been periodically stored on land at middle to high latitudes in the form of continental ice sheets. During the growth or melt of these continental ice sheets, tens of meters of sea level equivalent freshwater have been taken from or released back to the ocean. As a consequence, numerous studies have been conducted in an attempt to understand the role of freshwater perturbations on the stability of the MOC.
The dependence of sea water density on salinity is a key player in what is now known as a hysteresis behavior of the AMO. Recently it has been discovered that the AMO behaves like a “see-saw,” with relatively warm surface conditions in the North Atlantic and relatively cool surface conditions in the South Atlantic occurring when NADW formation is active, and conversely, with relatively cool surface conditions in the North Atlantic and relatively warm surface conditions in the South Atlantic occurring when NADW formation is inactive (Figure T2). Such a finding is supported by both modeling studies, proxy records, and from combined model/proxy record studies. More recently still, it has been realized that the “see-saw” nature of the AMO is fundamentally linked to a coupling between the strength of AAIW formation and the strength of NADW formation (Figure T3).
Early ice core records for the last glaciation have revealed large amplitude variability on the millennial timescale characterized by abrupt warming events (interstadials) lasting from several hundred to several thousand years. These oscillations (Figure T4 and T5), known as Dansgaard-Oeschger (D-O) oscillations, also appear in North Atlantic sediment records suggesting a role or response of the ocean. Evidence from the Santa Barbara basin and the Northeast Pacific suggests that a signature of these D-O oscillations is also present in the Pacific Ocean, while further recent sediment analyses suggest they may be an inherent part of late and early Pleistocene climate.
In an attempt to provide a mechanism for the observed D-O variability, it was initially proposed that a stable AMO was not possible during glacial times, when the northern end of the Atlantic Ocean was surrounded by ice sheets. Furthermore, when the AMO was weakened or shut down and ice sheets were growing, there was little oceanic salt export from the Atlantic to the other ocean basins. Assuming a net evaporation over the North Atlantic, its salinity continued to increase as moisture was deposited on land as snow, thereby expanding the ice sheets. Upon reaching a critical salinity, deep convection, and subsequently the AMO, turned on, transporting and releasing heat to the North Atlantic and thereby melting back the ice sheets. The fresh water flux into the North Atlantic from the melting ice sheets (or enhanced iceberg calving) eventually reduced or shut off the AMO, causing the process to begin anew. This hypothesis fundamentally assumed that most of the evaporation from the Atlantic basin was balanced by a net freshwater transport into the Atlantic associated with an active AMO. While this in itself has not yet been established, it could be true for a climate with an active AMO. However, it is clear that the weakening of the AMO would trigger a reorganization of the circulation in the South Atlantic, but it is not clear whether this reorganized circulation would be able to balance the required Atlantic evaporation.
Recent models, including an interactive continental ice sheet, showed that the earlier mechanism for D-O oscillations was in fact opposite to what actually occurred within the coupled system. During the cold, glacial climate, when the AMO was still active, the mass balance over continental ice sheets was positive. That is, rather than melting the ice sheets when the AMO was active, they grew, since the warmer atmosphere allowed greater precipitation (in the form of snow). The models also found a mechanism for D-O oscillations that involved an interaction between the AMO and the adjacent continental ice sheets, by assuming that the rate of iceberg calving into the North Atlantic was not constant in time, which is consistent with observations of ice-rafted debris. It was shown that after a delay of several hundred years following a major surging event and a reduction of the iceberg calving rate, the AMO intensified and the North Atlantic experienced an abrupt shift to a much warmer climate. This mechanism appears to be more intuitively appealing than a competing hypothesis involving stochastic resonance of the AMO, which relies upon the existence of an unknown 1,500 year external periodic forcing.
Heinrich (1988), in analyzing marine sediments in three cores from the North Atlantic, noted the presence of six anomalous concentrations of lithic fragments over the last glaciation. Since the source for these fragments was the land (and in particular Canada), he argued that this provided evidence for six anomalous surges of icebergs into the North Atlantic (Figure T5). The so-called Heinrich events (H) were even more striking when expressed in terms of the ratio of lithic fragments to the sum of lithic fragments and foraminiferal shells, due to the low foraminiferal counts in the Heinrich sediment layers. A simple model illustrates the mechanism for Laurentide Ice Sheet instability, which ultimately gives rise to an ice-sheet surge and a Heinrich event. It has been argued that the Heinrich 4 glacial cycles recorded in the Vostok ice core cycle consists of two phases. In the growth phase, the Laurentide Ice Sheet grew through snow accumulation while remaining attached to the bedrock. In the purge phase, the high pressures caused by the deep ice sheet caused thawing near the base of the ice sheet thereby allowing the ice sheet to surge seaward over a lubricated base through Hudson Strait. The resulting freshwater discharge into the North Atlantic would be of the order of 0.1-0.2 Sv over a period as short as 250-500 years.
Heinrich events, appearing about every 10,000 years, occur at the end of a sequence of D-O cycles during a prolonged cold period. Bond et al. (1993) noted that the sequences of D-O oscillations tend to follow a saw-tooth cycle (now termed a Bond Cycle) with successive D-O oscillations involving progressively cooler interstadials (Figure T5). They argued that this Bond Cycle was terminated by a Heinrich event, after which a rapid warming occurred and the process began anew.
The last cold event, known as the Younger Dryas (YD) (also known as Heinrich event zero, H0) took place between 12,700 and 11,650 years bp and terminated abruptly within a few decades. Proxy data suggest that during the YD, even the AMO was weakened. Further analysis of the proxy record indicates that the YD was preceded by a warm period in the North Atlantic climate, known as the Bølling-Allerød (B-A) warm interval (Figure T5). The B-A interval was accompanied by a strong global sea-level rise, apparently due to the melting of continental ice sheets. Until recently, it was not clear how to reconcile the warm North Atlantic climate during the B-A with melting of adjacent ice sheets. The latter would have caused a reduction of AMO, leading to a cooling, rather than warming of the North Atlantic. A possible explanation lies in the hypothesis that a significant portion of the observed rapid sea level rise during the B-A could have originated from the melting of Antarctic ice sheets, rather than solely from melting of the ice sheets surrounding the North Atlantic. A meltwater pulse into the Southern Ocean, originating from the partial collapse of an Antarctic ice sheet, could trigger an enhanced formation of NADW and an onset of warming in the North Atlantic (Weaver et al., 2003).
Despite these fundamental advances over the last few years, many challenges remain. In particular, a complete explanation including the essential role of the MOC, for the existence of millennial timescale (D-O) variability and its packaging into Bond Cycles in cold climates, its association with Heinrich events, and its dependence on the mean climatic state remains elusive to the paleoclimate community. Another major climate mystery is the cause of atmospheric CO2 and CH4 changes during the glacial cycles (Figure T4) and their possible connection to long-term changes of structure and strength of the GOC.
Antarctic Bottom Water and Climate Change
Binge-Purge Cycles of Ice Sheet Dynamics
Millennial Climate Variability
North Atlantic Deep Water and Climate Change
The Atlantic Ocean is the body of water bordered by the American continents, Greenland, Europe, Africa, and the Antarctic region. It is...
The vast body of water covering 70.8% of the surface of the globe, or any one of its principal divisions: the Antarctic, atlantic , Arctic,...
Abstract: The Atlantic Ocean, in particular its Meridional Overturning Circulation (MOC), is sensitive to the patterns of atmospheric forcing, in p