Southern Ocean mixing, seasonal sea ice, and glacialinterglacial CO2 variation.
Abstract
Several lines of evidence indicate that a change in the rate of stratification and mixing
in the Southern Ocean, either near the surface or at depth (most probably, both), is a
key factor in explaining glacial-to-interglacial atmospheric CO2 change. Here I
discuss how the Southern Ocean is ventilated today, and propose how this may have
been different in glacial time. Deep convection is confined to regions very close to the
continent. The open ocean south of the Polar front is largely ventilated from below,
with rapid mixing apparently driven by interaction of deep currents with topography
driving high rates of mixing well up into the water column. Between the surface and
~1500m depth, the water column is ventilated from above, stabilized by a halocline
that is due in part to sea ice formation and brine rejection, and probably dominated by
the effects of storms. I propose that in glacial time, more copious sea ice formation
towards the Antarctic continent, together with substantial seasonality and melting
further out, resulted in denser bottom water formation and more fresh water near the
surface. The greater stratification at depth caused lower mixing rates there while
greater winter-time sea ice cover reduced mixing towards the surface. The increased
stratification in the glacial deep ocean led to reduced ventilation of the deep ocean as
a whole, allowing the build up there of biologically-transported carbon. This scenario
is consistent with most proxies, including those on the extent of sea ice, the
productivity and nutrient distribution in the Southern Ocean, and the distribution of
13C. I use a box model, similar to that of Toggweiler (1999) to illustrate that, (in
conjunction with other mechanisms known to have influenced atmospheric CO2), this
scenario can reconcile CO2 variation with current proxies. Specific tests are suggested
that would help distinguish this “Southern Ocean seasonal sea ice” mechanism from
others that have been suggested.
Introduction
The possibility that a change in the rate at which the deep sea is ventilated could
lead to changes in atmospheric CO2 was first raised in 1984, when several papers
were published pointing to the possible role of the high-latitude oceans as controllers
of natural CO2 concentrations (Knox and McElroy 1984; Sarmiento and Toggweiler
1984; Siegenthaler and Wenk 1984). The box models on which these “Harvardton
Bears” papers were based highlighted the dependence of atmospheric CO2 on a
balance between biological productivity and ventilation between the surface and the
deep in the Southern Ocean. Today, water at the surface of the Southern Ocean (and
the North Pacific) contains non-zero mineral macro-nutrients (nitrate and phosphate),
and correspondingly has a higher pCO2 than would be the case if these nutrients were
more fully utilized. Much of the water in the the deep sea is ventilated from this
region, and its “preformed” nutrient and CO2 is set at the surface of the Southern
Ocean. Either increasing biological productivity, or decreasing the exchange of water
between the surface and depth in this region, can cause atmospheric CO2 to be lower
in this kind of model.
1
The results of these papers were initially interpreted in terms of increased
biological export productivity in the polar waters. However, it has become
increasingly clear that proxy evidence does not support the idea of an increased
Southern Ocean productivity in glacial time. There is room for doubt because not all
proxies seem to tell the same story, but a recent “multiproxy” compilation of LGM
export production estimates (Bopp et al. 2003) suggests a coherent picture: an
increase in productivity in the Subantarctic region, particularly in the Atlantic sector,
in glacial time, but less export production from the region south of the (present
position of) the Polar Front.
An alternative possible explanation for lower glacial CO2 is that the rate of
ventilation of the deep ocean was slower in glacial time. Recently, (Toggweiler 1999)
has revived this idea as an important mechanism for causing lower atmospheric CO2.
In fact, given the constraints put on the problem of atmospheric CO2 change by our
present knowledge, we argue in this paper that the deep ocean must have been more
slowly ventilated in glacial time. It is however not trivial to understand why this might
have been the case. In view of this, it is important to fully understand how “Southern
component” deep water ventilates the ocean today. Accordingly, much of this paper is
concerned with reviewing and clarifying this using observations of the modern ocean.
We then suggest why ventilation may have been different in glacial time. Some
calculations using a box-type model are used to show that the mechanisms we
propose, acting in combination with other effects that we know occurred, have the
potential to change atmospheric CO2 by the right amount. This is not the first
suggestion for mechanisms to change ocean ventilation, and we therefore end by
comparing it with other recently proposed mechanisms, suggesting tests that may help
to distinguish between the competing theories.
Our uncertainty about the mechanisms causing glacial-interglacial CO2 change,
even after more than 20 years of research, is sometimes cited as an illustration of our
ignorance of fundamental processes in the Earth system. However, though there are
differences in the detail, it is worth noting that all the viable theories are now
convergent in their most important aspects. Plausible theories for lower glacial
atmospheric CO2 all share a requirement for lower ventilation rates of the deep sea,
obtained by one or more of: more sea ice cover, increased near-surface stratification
of the Southern Ocean or a greater deep stratification. These changes serve to partition
biologically fixed carbon into the deep sea and away from the atmosphere. The higher
carbon concentration in deeper water also leads to an increased alkalinity of the ocean
by “carbonate compensation”, since deep water would otherwise be more corrosive to
carbonate in sediments, leading to an imbalance in the source and sink of ocean
alkalinity. Both the increased efficiency of the carbon pumps and the increased
alkalinity serve to decrease atmospheric pCO2 without requiring large increases in
biological productivity.
The global ocean overturning circulation
Figure 1, adapted from (Toggweiler 1994) and (Sloyan and Rintoul 2001), shows a
schematic of the global ocean overturning circulation. The deepest, densest waters of
the world ocean are formed close to the Antarctic continent, where salty, cold waters
from brine rejection linked to net sea ice formation are found. These transfer into the
deep ocean by down-slope gravity currents and by convection. They mix with fresher
circumpolar deep water as they do, to varying extents, resulting in a variety of distinct
types of Antarctic bottom water (AABW) in the different basins of the polar Southern
2
Ocean (Orsi et al. 1999). North Atlantic Deep Water (NADW), less dense, warmer
and saltier than AABW, is formed by sinking of water in the Northern North Atlantic,
and penetrates southward. Intermediate waters, freshened by net precipitation and sea
ice melt, are formed by the northward transport and subduction of Antarctic surface
waters by Ekman drift under the influence of the westerly winds in the Southern
Ocean.
There are two contrasting views about the mechanics of the overturning
circulation. The more classical view, articulated by (Munk 1966) and elaborated by
(Munk and Wunsch 1998), is that the steady state circulation is governed by a balance
between vertical advection and downward turbulent diffusion of buoyancy. In this
“abyssal recipes” picture, the rate-limiting step is how fast buoyancy can mix
downwards. The vertical mixing rate, represented by the turbulent vertical diffusivity
κz, is highly variable with position in the ocean. Measurements show κz to be very low
in the thermocline (Ledwell et al. 1998) and the interior of the oceans away from
rough topography (Toole et al. 1994). However, internal waves and turbulence
generated by currents interacting with bottom topography generate enhanced mixing
which can extend up through the water column for kilometres (Polzin et al. 1997;
Ledwell et al. 2000). These mixing “hot spots” are now thought to dominate the
overall mixing of the oceans.
The alternative view of the ocean overturning is that it is powered by wind-driven
surface convergence and divergence, especially that associated with the westerlies in
the Southern Ocean (Toggweiler and Samuels 1995, 1998). In this view, the surface
Ekman drift associated with the zonal winds near the latitude of Drake Passage lifts
water out of the deeper ocean. This in turn allows a compensatory sinking of water
elsewhere, i.e in the North Atlantic. This wind-driven overturning is not dependent on
interior diapycnal fluxes, and can proceed even as κz in the interior tends to zero.
(However, as discussed below, it is dependent on surface buoyancy forcing.)
Almost certainly, the real ocean is ventilated by a combination of both of these
modes. Given that rates of diapycnal mixing are agreed to be low in the main
thermocline, advective processes, ultimately wind-driven, will likely dominate
vertical transport through the top kilometre or so of the water column over most of the
world ocean. Some portion the NADW formation is probably wind-driven by Ekman
suction in the Southern Ocean, see for instance (Webb and Suginohara 2001)). On the
other hand, the ventilation of the densest waters in the world ocean must be largely
diapycnally driven. This seems inevitable, because this water has no surface outrcrop
of significant area – it is formed in a few special places, in dense but areally restricted
plumes due to shelf processes or convection. There is no known mechanism to return
it to the surface where it can be destroyed by air-sea flux processes in similarly highly
restricted regions, so its return must involve modification to less dense water in the
interior. On tracing the path of this water away from the Antarctic, it becomes clear
({Mantyla, 1983 #211} also as described further below) that it is so modified, its
properties being altered in the deep sea, well away from the surface, implying large
interior diapycnal fluxes. Therefore, if ventilation of the deepest ocean was slower in
glacial than interglacial time, as Toggweiler (1999) has suggested, this is likely to be
largely a question of changing diapycnal mixing rates. This presents a theoretical
challenge. Why should deep ocean mixing be slower in glacial time?
Recently, (Munk and Wunsch 1998) have pointed out that the energy required to
drive the abyssal circulation is of the same order as the total energy available from
wind and tidally generated currents in the deep sea. Plausibly therefore, the overall
control on the rate at which the deep MOC turns is one of energy limitation. The
3
power required to maintain the diabatic abyssal circulation is readily calculated. From
a scale analysis, the rate at which potential energy must be added to raise bottom
water being formed with a volume flux of Q is W ~ Qgh , where  is the density
difference between the initial and final densities of the water, g the acceleration due to
gravity and h is the scale height of the density change. If the available power supply
is constant, then, other things being equal, Q might be expected to be inversely related
to , with the total buoyancy flux Q remaining constant. Similarly, scale analysis
of the steady-state advective-diffusive balance (e.g. that employed by (Munk 1966))
gives us the relation Q ~ A z / h between the overturning rate, the area A over which
upwelling occurs, the vertical mixing rate and the scale height. Substituting for Q in
the first relation we obtain W ~ A z g , showing that if the power is indeed
constant, then vertical mixing rates should be inversely proportional to density
difference. Recent papers by Nilsson and colleagues (Nilsson et al. 2003, 2004) have
explored in greater detail the implications of an energy-limited overturning for the
modern day circulation with conceptual and numerical models. Here we simply note
that there is some reason to believe that mixing and overturning rates may decrease as
density anomaly increases. Suppose then, that in glacial time the density contrast
between the deepest waters and those overlying them was greater than it is today, this
might help explain why the deep sea was apparently less well ventilated at LGM.
What do we know about stratification of the deep ocean in glacial time? Figure 2,
from the work of (Adkins et al. 2002), shows estimates of the temperature and salinity
in the deep ocean at LGM from measurements on pore waters, compared with present
day conditions. Contours of potential density referenced to 4000m are also shown.
The difference between the stratification of the glacial and modern ocean is striking.
Salinities are greater everywhere, which is to be expected because the greater volume
of fresh water bound up in icecaps in glacial time resulted in a saltier ocean. However,
whereas the modern abyssal ocean is largely stratified by temperature differences, the
density differences in the deep glacial ocean are mostly due to salinity. Temperatures
are close to freezing at all the sites in the glacial ocean, but the deep Southern Ocean
shows as much saltier than the other locations. The density anomaly between the
Southern Ocean site and the others is about three times larger than it is today, and a
crude application of the above (much over-simplified) reasoning might then suggest
that the mixing between this Southern component and the overlying waters was three
times slower at LGM.
Mixing and ventilation of the modern-day Southern Ocean,
Diapycnal mixing and ventilation from below: In recent years a good deal of new
information has become available relating to rates of diapycnal mixing in the
Southern Ocean. In the near-surface there have been several direct measurements
using tracer releases (Law et al. 2003), Goldson, 2004). In the region of the Scotia
Sea, studies making use of new techniques using lowered acoustic doppler current
profilers have suggested very high mixing rates in the region of the Scotia Sea
(Heywood et al. 2002; Garabato et al. 2004). Estimates using CFCs in the broad
plume of bottom water originating from the Weddell Sea and flowing towards the
Indian Ocean have indicated high rates of mixing between this and the overlying
water (Haine et al. 1998). These estimates are drawn together in Table I. There is a
considerable range of values, particularly in the deep ocean, reflecting the huge
influence of topography. However, a consistent picture emerges: in the upper water
column, mixing across and below the summer-time seasonal pycnocline is ~10-5 m2s-1,
4
though with evidence of rates that are several times higher during passage of storms
(Goldson, 2004). In the deep Southern Ocean however rates are ~10-3 m2 s-1, ranging
higher still immediately over rough topography. Results of inverse models
constrained with hydrographic data can also be used to infer diapycnal fluxes, but in
the Southern Ocean, these broad averages are difficult to separate into fluxes due to
air-sea interaction and those due to interior mixing (Ganachaud and Wunsch 2000;
Sloyan and Rintoul 2001).
Table I: Estimates of diapycnal mixing rates in the Southern Ocean
method
Depth
Location
Value (10-4m2s-1)
(Law et al. 2003)
SF6 tracer
release
50-100m
61S,
0.11 0.2
Goldson. (2004)
SF6 tracer
release
50-100m
(Garabato et al.
2004)
LADCP shear
+ hydrography
0-200m
Scotia Sea, Drake
Passage
0.1 – 0.3
(Haine et al. 1998)
CFC budget
> 3500 m
Abyssal S. Ocean
-
52
(Heywood et al.
2002)
LADCP and
basin budget
> 3500m
Scotia Sea
39 
(Garabato et al.
2004)
LADCP shear
+ hydrography
> 2000m
Scotia Sea
5 – 100
West Drake
Passage
10 – 1000
Atlantic and
Indian sector,
-
~10
Source
Upper water
column
?
Deep water
This paper
CFC budget
> 2000m
As shown, for example, by (Haine et al. 1998) chlorofluorocarbon concentrations
can be useful for illuminating the processes of ventilation from the surface to the
interior of the Southern Ocean. Figure 3 shows a compilation of CFC-11
concentrations on five WOCE sections from Drake passage through the Atlantic and
Indian Ocean sectors.. The concentration scale has been adjusted so that
concentrations below 1 pmol kg-1 are highlighted. Deep waters at Drake Passage show
almost undetectable CFCs. Concentrations are also low north of about 50S.
However, the deep outflow of recently ventilated bottom water from the Weddell Sea
is clearly seen as a core at about 60S in both the Atlantic Ocean sections. The section
through the Eastern Scotia Sea shows relatively high concentrations throughout the
5
water column, due perhaps to mixing of this deep outflow up through much of the
water. Likewise, all the other sections show enhanced concentrations in the deepest
waters south of 60S, and easily measurable concentrations right up through the water
column. The more easterly sections also show evidence for extensive “local” deep
water formation close to the Antarctic continent, not associated with the Weddell or
Ross Sea outflows.
Figure 4. shows the means of the CFC concentrations from the Atlantic sections,
averaged south of 50S. The concentrations were averaged along surfaces of constant
4, and are plotted as a function of depth after transforming into depth coordinates,
using the average depth vs 4 relationship. The profiles south of 50S reinforce the
impression of a CFC sourced at the bottom that mixes up through the water column to
depths ~1000m, especially by the very intense mixing in the Scotia Sea. If the bottom
water is the origin of the mid-depth CFC, a simple scale analysis can be used to
estimate the order of magnitude of the vertical mixing that must be responsible, as
follows: CFCs have a time scale for rising in the atmosphere of order 20 years. The
profiles in Fig. 4 have length scales of order 1000m or more. From these scales, a
mixing rate of order z ~ (length scale)2/(2 x time scale) ~ 10-3 m2s-1 follows. This is
broadly compatible with the deep mixing rates obtained by other workers, as
summarized in Table 1.
We can approximately calculate the upwelling
It is probable that not all the mid-water CFC content in Figs 3 and 4 comes from
up-mixing from the bottom source. Newly formed water of less extreme density
probably joins the circumpolar waters at the Weddell-Scotia confluence and
contributes to the mid-depth concentrations. The source of this water is also ice-shelf
interaction, and it too has been subject to intensive mixing. While the above
calculation of diaypcnal mixing would tend to overestimate due to the neglect of this
source, the model of (Haine et al. 1998) for the outflow from the Weddell Sea
considered only data from east of the prime meridian, and still found similarly high
mixing rates.
The effect of this rapid mixing on the properties of newly formed AABW are
substantial. The densest waters are not found north of the ACC. In the process of
transiting away from the continent, they are mixed to lighter densities. For example,
according to the measurements of Haine et al., the Weddell sea water has a transit
time to the Crozet-Kerguelen region of only ~40 years, but this is long enough for its
temperature to rise from < -0.4 C to about 0.5C. Thus the AABW which forms the
bottom water for most of the world’s oceans is substantially less dense than the water
which first sinks to the bottom in the Southern Ocean.
Meridional circulation: Figure 5 shows a schematic of the stream function of
meridonal circulation of the Southern Ocean, averaged zonally. Furthest to the south
is the dense water produced on the continental shelves and sinking to the bottom of
the ocean, penetrating northward and being recirculated by the intense mixing. Above
it lies the circumpolar deep water. The streamfunction here is redrawn from {Karsten,
2002 #87}
The upper water column in the Southern ocean is today largely stratified by a
halocline. The surface water is relatively fresh, and the Ekman flux drives this water
to the north until it is subducted under the warmer and saltier waters north of the
6
Subantarctic front to form Antarctic intermediate water. There are two potential
sources for the freshening that converts salty upwelling circumpolar water into fresh
AAIW. One is meteoric water -- net precipitation minus evaporation, and the other is
fresh water formed from ice melt, for which the associated salt has been removed
from the mixed layer by brine rejection processes.
It is of interest to calculate the order of magnitude of each of these processes
today. Observationally-derived records such as COADS are unreliable in the sparsely
monitored Southern Ocean, so we use the long-term averages of NCEP re-analysis
data to examine the precipitation minus evaporation (P-E) balance. Figure 6 shows
NCEP annual, zonally-summed net freshwater, accumulated northward from 85S.
The net P-E flux between 85S and 60S is 0.32 Sv, while between 85S and 50S it is
0.68 Sv.
To find the amount of fresh water produced by fractionation of sea water in the
formation of sea ice, we estimate the volume of fresh water incorporated into sea ice
annually, as V = Ahr where A is the area covered by seasonal sea ice, h is its draft,
and r is the ratio of the densities of fresh water and ice (about 0.9). The area A is well
documented from satellite observations, and is about 1.5 x 107 km2 (Zwally et al.
2002). Relatively few observations of ice thickness have been made. We need to be
careful also, to ensure that we count only ice formed from freezing of sea water, and
not the additional accumulation of snow on top of this. The observations of (Worby et
al. 1996) in the Bellingshausen and Amundsen seas suggest an average h ~ 0.5m for
this ice thickness. The work of (Harms et al. 2001) suggests total drafts of ~ 0.3m for
sea ice measured at moorings on the Greenwich meridian, with thicker ice (one to two
meters) in the Weddell Sea. If we assume that on average the ice is 0.5 m thick, using
the area above we derive a figure of 0.22 Sv for the annual average rate of fresh
water production due to Southern Ocean seasonal sea ice formation. Thus we
estimate that today, ice formation contributes about 40% of the net fresh water budget
of the surface ocean south of 60S, and about 22% of the budget south of 50S. There
is large uncertainty on these figures, but it is clear that the contribution from sea ice
is important. In glacial time, when the hydrological cycle was weaker (there was less
evaporation because sea surface temperatures were lower) but the area covered
seasonally by sea ice was about twice as great, ice would likely have been the
dominant contributor to the fresh water budget south of the Polar front.
Role of: Temperature, salinity, growth of terrestrial biosphere,
Proxy evidence: salty deep water near the freezing point, low Antarctic productivity
Del13C
The need for a ventilation-related mechanism
The need for slower ventilation of the deep ocean. – a puzzle
Energy-limited mixing of the deep ocean ventilation.
How can deep ocean in glacial
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9
Figure 1: Potential temperature-salinities of deep waters at the LGM derived from porewater
measurements at four ODP sites by Adkins et al., 2002. The modern values at these sites are
also shown. Contours are of 4. The open arrow shows the modern mean ocean salinity, while
the filled arrow is the average salinity adjusted for a 125m drop in eustatic sea level, similar
to that at the LGM. The inset shows the locations of the four sets of measurements, with the
depths in metres written by each symbol.
10
Cumulative, area integrated net freshwater (Sv from 85°S)
1.0
0.8
0.6
0.4
0.2
0.0
-0.2
-0.4
-0.6
-0.8
-1.0
80
70
60
50
40
30
20
10
Latitude (° S)
Figure . Net precipitation – evaporation for the southern hemisphere, integrated zonally and
between 85S and the latitude on the abscissa. Values at 50S and 60S are highlighted by the
dotted lines. The calculation used the time-average of the NCEP re-analyses results from
1968-1996 (available from the NOAA-CIRES Climate Diagnostics Center,
www.cdc.noaa.gov).
11
0
2
1
1000
3
45.60
45.80
45.90
5
2000
46.00
4
Depth (m)
4
4
3000
46.10
5
3
4000
46.15
2
5000
0
0.2
1
0.4
0.6
CFC-11 concentration (pmol kg-1)
0.8
1
12