Thermohaline Circulation and Climate
Introduction
These notes discuss the thermohaline circulation of the ocean and its effect on
climate. These notes cover some basic physical properties of the ocean that are relevant
to the understanding of the system, discuss the global pattern of the thermohaline
circulation, and states some of the important dynamical effects in the system. The effect
this circulation has on global warming, and some discussion of observed model
variability in the system are then shown.
1. Ocean Basics
Compared with the atmosphere above it, the ocean is very dense. Water masses
approximately one tonne per cubic meter. In contrast, the entire mass of an atmospheric
column one square meter in area is ten tonnes, meaning that the mass contained in a ten
meter water column is equal to the mass of the entire atmosphere above it. Furthermore,
since the heat capacity of water is four times the heat capacity of air (per unit mass) the
upper 2.5 meters of the ocean has the same heat capacity as the entire atmospheric
column above it. This large heat capacity is the major source of the ocean’s ability to
affect climate.
Ocean transports are often measured in Sverdrups, defined as 106 m3/s of water—
equivalent to a 1 m/s current moving through a surface a square kilometer in area.
Average rainfall on earth is about 3mm/day, or 16 Sverdrups. Since rivers draw their
transport from rain falling on a basin, river flow is generally a small fraction of a
Sverdrup. By contrast, the Gulf Stream’s transport varies from about 30 Sverdrups near
Florida to about 150 Sverdups near Cape Cod and the transport of the Antarctic
Circumpolar Current is about 150 Sverdrups..
Figure 1 illustrates the variation of seawater density with Salinity and
Temperature. Density is given in sigma units, where  = ( - 1000 kg/m3) --since density
typically varies from 1020 kg/m3 to 1028 kg/m3 in the ocean, only the final two digits are
of significance. There are two forms of sigma in use, sigma-t and sigma-theta, also
written t and . Sigma-t is the sigma calculated using in-situ ocean temperature, and
sigma-theta is a ‘pressure-corrected’ density, defined as the sigma a water parcel would
have if it were adiabatically raised to the ocean surface from depth.
Seawater gets lighter as temperature increases and as salinity decreases. This is in
contrast to fresh water, which has a density minimum in temperature; above a salinity of
~23 ppt, density is monotonic with temperature. At low temperature, the density surfaces
on the graph become almost vertical, indicating that at low temperature density hardly
depends on temperature and is therefore controlled by salinity.
Figures 2 and 3 show surface density maps for Northern and Southern hemisphere
winters, respectively. In the northern winter, the densest surface waters in the world exist
in the Labrador and Greenland Seas, while in the southern winter, the Weddell Sea near
Antarctica has the densest surface waters with the Ross Sea a close second. These are the
only sites in the world where surface densities become large enough to drive surface
water to the bottom of the abyss.
Surface salinity is largely controlled by the difference between evaporation and
precipitation on the waters, E-P. Because the atmosphere tends to diverge moisture from
the Atlantic and deposit it in the Pacific, the Atlantic generally has more E than P and the
opposite is the case for the Pacific. Thus, Atlantic surface waters are generally denser
than Pacific surface waters. The Atlantic is also generally warmer than the Pacific; since
the Gulf Stream transports a large amount of equatorial water to northern latitudes. As
the salty Gulf Stream flows north it loses heat to evaporation and sensible heating and,
once it reaches the Greenland Sea, is cold and salty. This allows the Greenland Sea
surface waters to become dense enough to convect to the bottom.
It takes thousands of years to make large alterations to the deep water mass
properties by overturning, since the deep sinking in the Greenland Sea and around
Antarctica drives a circulation of about 15 Sverdrups (the time scale is set by the time it
takes to renew the entire volume of the ocean by 15 Sverdrup flows). Thus, this deep
convection serves to set the deep ocean water properties and guarantees it is cold even
when the surface is very warm (e.g. in the tropics). Over the entire As heat is constantly
diffusing downwards from the equatorial surface, the deep waters would eventually warm
up if this diffusion was not countered by deep sinking injecting cold water to the bottom.
Thus the world’s thermohaline circulation is also responsible for maintaining the stable
stratification of the world’s oceans.
2. Atlantic THC
The thermohaline convection in the Greenland and Labrador Seas serves as the
forcing for a global thermohaline circulation, or THC, shown in figure 4. Water in the
Greenland Sea sinks to the bottom of the Atlantic, where it moves southwards along the
western boundary. It flows through the Atlantic until it reaches the Antarctic
Circumpolar Current (ACC), where it meets deep water formed in the Weddell Sea
during southern winter. From here the ACC sweeps the water into the Indian and Pacific
basins, where it slowly upwells towards the surface, and begins returning to the Atlantic.
Once in the Atlantic, it joins the Gulf Stream and is returned to the Greenland Sea. The
length of time for a to complete this circulation is on the order of 3000 years.
This circulation pattern has been inferred largely through oceanic tracers. Figure
5 shows water ‘age’ (defined as the time elapsed since the water was last at the surface)
at 3 kilometers depth, based on Carbon-14 tracer data. It shows that the youngest deep
water occurs around the Greenland, Labrador and Weddell Seas, where the deep water
forms. The water becomes progressively older as it travels along the western boundary
of the Atlantic, to the Indian and, finally, into the Pacific ocean, where the oldest deep
water masses in the world ocean are found.. This pattern also shows up in other tracers,
such as in oxygen utilization. Studies of anthropogenic freon, which has only existed in
the atmosphere over the last century (bsically since the 1940s), show freon appearing in
the North Atlantic’s deep western boundary current and only slowly working its way into
the interior ocean.
The flow of the THC through the Atlantic enhances the Gulf Stream since the
total vertically integrated transport along the western boundary is set by the magnitude of
the wind stress curl on the surface of the Atlantic. By contrast, the southward flowing
THC weakens the southward Brazil Current. The rapidly flowing Gulf Stream brings
more heat northward than the Brazil current brings south: the heat transport is northward
everywhere in the Atlantic. This makes the North Atlantic warmer than the South
Atlantic creating a larger evaporative flux from the northern oceans, removing fresh
water from the northern hemisphere ocean. Thus, the THC creates a net transport of heat
from the southern hemisphere to the northern hemisphere, and a net flux of fresh water in
the opposite direction.
3. Modelling
Since the THC operates over millennial timescales, the main channel to studying
its dynamics is computer modeling. Most models used to study the long term THC
dynamics have been coarse-resolution, in order to achieve timesteps long enough to
integrate the many thousands of years needed to simulate the equilibrium circulations and
major features of thermohaline variability. Such models have been able to reproduce the
general pattern of the THC. They also appear to show that while Antarctic Bottom Water
formation appears relatively stable, North Atlantic Deepwater formation can be turned on
and off by changes in the North Atlantic surface salinity (Manabe and Stouffer, 1988).
The timescale for a complete trip around the overturning circulation is about 3000
years. However, disturbances in the system may propagate much faster than this
timescale would indicate. Models of frictional boundry layers have shown that a tracer
released in the deep Atlantic can propagate into the Pacific basin after only 500 years, as
shown in figure 6. (Goodman, 1996) Model runs indicate that turning on a deep water
source in the North Atlantic can propagate Kelvin waves with will affect the depth of the
Pacific thermocline within 12 years, which in turn can modulate ENSO strength.
Therefore the estimate of 3000 years for water renewal by the THC should be treated
with caution: faster time scales exist as a result of water in the frictional boundary layer
and as a result of wave processes.
4. THC and Global Warming
The Thermohaline Circulation acts as a negative feedback mechanism to global
warming in the northern hemisphere. As temperatures rise, freshwater increases at high
latitudes, both by additional precipitation and by icecap melting. This will have the effect
of lowering surface densities at the deep water formation sites, slowing or even shutting
off the THC. Without the THC, net global northward heat transport would be lessened,
creating a negative feedback; in the southern hemisphere, conversely, less heat would be
exported, creating a positive global warming feedback.
An intriguing possibility for the THC not slowing down was recently proposed by
Latif and collaborators. The idea is that the Atlantic THC depends in an essential way on
the divergence of moisture from the Atlantic to the Pacific. If global warming increases
the regularity and intensity of El Niño in the tropical Pacific, then the region of persistent
precipitation in the western tropical Pacific will move eastward and draw moisture from
the subtropical Atlantic. The Gulf Stream will therefore deliver saltier water to the
sinking regions in the GIN seas and this can counteract the effects of freshening due to
precipitation and ice melt. Since most coupled models of the climate used for global
warming calculations don’t do El Niño correctly, this hypothesis remains to be verified.
5. Variablity
Model integrations show the THC to be sensitive to surface salinity. F. Bryan
(1986) ran a coarse resolution model of a 60o sector of the earth from pole to pole,
representing an idealized ocean basin. He found that perturbing a symmetrically
circulating model THC by salting or freshening the south could result in a reversal of the
THC on a 50-100 year timescale, as shown in figure 7. These quick timescale changes to
the circulation are possible due to frictional flows along the model boundry.
Weaver et al. (1993) ran a similar model and found that the equilibria for the THC
altered as the forcing was increased. If the northern latitudes were salty, one stable state
for the circulation existed. If the freshening were increased, two stable states were
observed in the model. Increasing the freshening further, however, put the model into a
regime which had one stable state, and one unstable, as shown in figure 8. The unstable
state experiences chaotic fluctuations in the strength of the THC, until eventually settling
down to the stable state. If the freshening is increased even further, the model oscillations
eventually go unstable, and the THC stops entirely. This final state is characterized by
long quiescent periods followed by short ‘flushing’ events, when the deep ocean is
warmed sufficiently by diffusion to become lighter than the surface waters. This wide
variability, controlled by freshwater content in the northern surface waters, may
contribute to the observed Dansgaard-Oeschger oscillations in the paleoclimate record.
Figure Captions
Figure 1. Potential density as a function of temperature and salinity for values
representative of the world ocean. Note the nearly vertical line slopes near 0 oC,
indicating that at low temperature, density is largely a function of salinity. (From
Figure 2. Northern hemisphere winter sea surface density, in sigma-t. Highest densities
occur in the Labrador and Greenland Seas. (Levitus, 1982)
Figure 3. Southern hemisphere winter sea sufrace density, in sigma-t. Higest densities
occur near Antartica. (Levitus, 1982)
Figure 4. Schematic diagram of the thermohaline circulation. (CLIVAR Website)
Figure 5. Radiocarbon ages for deep-sea water. Dots indicate station locations and ages
of water taken from 3 kilometers depth. Shaded areas indicate sites of deep water
formation. (Broecker, 1985)
Figure 6. Fraction of model water column tagged with tracer at 100 years and 500 years
after tracer release in the North Atlantic. The diamond in the Greenland Sea indicates
tracer release point. The tracer moves quickly along the western boundary, allowing it to
reach the Pacific after only 500 years. (Goodman, 1996)
Figure 7. Streamfunction for zonally integrated thermohaline circulation in an idealized
model ocean after 1300 years integration. Center top figure shows control run, which
was forced with salinity perturbations which were symmetric around the equator. Left
column shows streamfunction after 15, 44 and 68 years when the southern hemisphere
salinity was decreased 1‰. Right column shows streamfunction after 342, 411 and 534
years when the southern hemisphere salinity was increased 2‰. Positive values indicate
clockwise circulation; contour interval is 2.5 Sverdrups. (Bryan, 1986)
Figure 8. Upper panels show streamfunction for zonally integrated thermohaline
circulation for two equilibria of an idealized model ocean. Upper right panel shows
unstable equilibrium circulation, upper left panel shows stable circulation. Lower right
panel shows eventual evolution of model from the unstable to the stable circulation mode.
Lower left panel shows evolution of model under ideantacle forcing, but spun up from
rest. Contour interval is 1 Sverdrup. (Weaver et al., 1993)
References
Broecker, W, 1985: “Building a Habitable Planet”, Eldigio Press.
Bryan, Frank, 1986: High Latitude Salinity Effects and Interhemispheric Thermohaline
Circulations. Nature, 323, 301-304.
CLIVAR Website:
http://www.clivar.org/publications/other_pubs/clivar_transp/pdf_files/av_d3_992.pdf
Goodman, Paul, 1996: The Role of North Atlantic Deep Water Formation in the Global
Ocean Circulation. University of Washington MSc. Thesis.
Levitus, S., 1982: Climatological Atlas of the World Ocean, NOAA Professional Paper
13. US Dept of Commerce: National Oceanic and Atmospheric Administration.
Manabe, S, and RJ Stouffer, 1988: Two Stable Equilibria of a Coupled OceanAtmosphere Model. Journal of Climate: Vol. 1, No. 9, pp. 841–866.
Weaver, Andrew, Jochem Marotzke, Patrick Cummins and E Sarachik, 1993: Stability
and Variability of the Thermohaline Circulation. Journal of Physical Oceanography: Vol.
23, No. 1, pp. 39–60.