The role of Southern Ocean winds for the global meridional

Diploma Thesis at the Department of Physics
Freie Universität Berlin
The role of Southern Ocean winds for the global meridional
overturning circulation in the Earth System Model of
Intermediate Complexity CLIMBER-3α
Jacob Schewe
Berlin, December 10, 2007
Abstract
The global meridional overturning circulation (MOC) is a system of large-scale ocean currents that
spans the entire world ocean and has various important implications for global climate. The driving
mechanism of this circulation is subject of ongoing discussion. There are two candidates: Turbulent
mixing of heat in low latitudes, and upwelling driven by the wind stress in the Southern Ocean (SO).
Climate projections conducted for the latest IPCC report suggest a significant strengthening of SO winds
for the near future.
This study aims to investigate the role of Southern Ocean winds for the meridional overturning
circulation in a coupled climate model of intermediate complexity. We have carried out a number of
simulations with the wind stress forcing in the Southern Ocean varying around the observed field. In
order to test the robustness of the results, simulations with anomalous freshwater flux into the North
Atlantic were carried out.
The results of these experiments are used to examine the influence of the so-called Drake Passage effect
and associated Southern Ocean upwelling on the Atlantic, and global, meridional overturning circulation,
and to outline an extended concept of the mechanisms responsible for setting the magnitude and structure
of the circulation. We find the circulation, at present-day boundary conditions, being entirely driven by
Southern Ocean winds. The distribution of Southern Ocean upwelling onto the ocean basins changes,
however, with varying winds, and is determined by meridional density gradients. The concept integrates,
to some extent, two opposing general theories and has the potential to contribute to a more coherent
view of the global oceanic circulation (Schewe and Levermann, 2008).
Furthermore, we discuss the effects of varying SO wind stress on the interhemispheric distribution
of surface temperatures, which shows an unexpected response. In simulations where the Atlantic MOC
is reduced by means of anomalous freshwater input, the surface cools in the northern hemisphere, and
warms in the southern hemisphere - the so-called ’bipolar see-saw’. However, when the AMOC weakens
due to reduced Southern Ocean wind stress, we observe a relative cooling in both hemispheres. This can
be explained by reduced sea ice export from the Antarctic. Increased sea ice extent and albedo result in
a cooling which overcompensates the warming caused by the reduced northward oceanic heat transport
(Levermann et al., 2007).
iii
Die Bedeutung der Winde über dem Südlichen Ozean für die
globale meridionale Umwälzzirkulation im Erdsystemmodell
mittlerer Komplexität CLIMBER-3α
Kurzfassung Die globale meridionale Umwälzzirkulation (MOC) ist ein System großskaliger
Ozeanströmungen, das alle Weltmeere umspannt. Sie ist in vielerlei Hinsicht von großer Bedeutung für
das Klima der Erde. Ihr Antriebsmechanismus ist noch immer Gegenstand einer seit langem geführten
wissenschaftlichen Auseinandersetzung. Zur Diskussion stehen zwei Kandidaten: Turbulente Durchmischungsprozesse, durch die in niedrigen Breiten Wärme in den tiefen Ozean gelangt; und windgetriebene
Aufwärtsströmungen im Südlichen Ozean (SO). Für den aktuellen IPCC-Bericht durchgeführte Klimaprojektionen deuten auf eine signifikante Verstärkung der Winde über dem SO in der nahen Zukunft hin.
Die vorliegende Arbeit untersucht die Bedeutung der südlichen Winde für die globale meridionale
Umwälzzirkulation in einem gekoppelten Klimamodell mittlerer Komplexität. Wir haben in einer Reihe
von Simulationen die Windstärke über dem SO um das beobachtete Feld herum variiert. Zusätzlich wurden Simulationen mit verstärktem Süßwassereintrag in den Nordatlantik durchgeführt, um die Robustheit
der Ergebnisse zu überprüfen.
Anhand der Resultate analysieren wir den Einfluss des sogenannten Drake Passage-Effekts und der
damit verbundenen Aufwärtsströmungen auf die atlantische und globale MOC. Wir skizzieren ein erweitertes Konzept der Mechanismen, die Stärke und Struktur der Zirkulation bestimmen. Wir stellen fest,
dass die MOC unter heutigen Randbedingungen vollständig von südlichen Winden angetrieben wird. Die
Verteilung des aufströmenden Wassers auf die einzelnen Ozeane ändert sich jedoch, wenn die Windstärke
variiert wird, und ist durch meridionale Dichtegradienten bestimmt. Das Konzept verbindet zu einem
gewissen Grad zwei bisher disjunkte Theorien und könnte zu einer einheitlicheren Beschreibung der globalen Ozeanzirkulation beitragen (Schewe and Levermann, 2008).
Des weiteren betrachten wir die Auswirkungen von Veränderungen der südlichen Winde auf die
Oberflächentemperaturen von Nord- und Südhemisphäre. In Simulationen, in denen die atlantische MOC
durch verstärkten Süßwassereintrag abgeschwächt wird, beobachtet man eine Abkühlung der Oberfläche
auf der Nord-, und eine Erwärmung auf der Südhalbkugel. In unseren Experimenten, in denen die
AMOC durch reduzierte südliche Winde abgeschwächt wird, beobachten wir jedoch überraschenderweise
eine Abkühlung beider Hemisphären relativ zum heutigen Zustand. Wir können dies damit erklären,
dass die geringere Windstärke den Abtransport von Meereis aus antarktischen Breiten behindert. Ausgedehntere Eisbedeckung und höhere Albedo führen zu einer Abkühlung, die die aufgrund des reduzierten
nordwärtigen Wärmetransports zu erwartende Erwärmung überkompensiert (Levermann et al., 2007).
v
Contents
1 Introduction
1
2 Model description
4
3 The meridional overturning circulation
7
3.1
The schematic of an oceanic conveyor belt . . . . . . . . . . . . . . . . . . . . . . . . . . .
7
3.2
Upwelling waters and AMOC driving mechanisms . . . . . . . . . . . . . . . . . . . . . . .
10
4 Geostrophy and the Drake Passage Effect
13
4.1
Geostrophic balance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
13
4.2
The Drake Passage effect . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
14
5 Control mechanisms of the MOC
17
5.1
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
17
5.2
Experiments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
18
5.3
Wind-driven Southern Ocean upwelling . . . . . . . . . . . . . . . . . . . . . . . . . . . .
19
5.4
Volume transport and meridional density gradients . . . . . . . . . . . . . . . . . . . . . .
21
5.5
Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
28
6 Lack of bipolar see-saw
30
6.1
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
30
6.2
Experiments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
31
6.3
Temperature response to SO wind changes . . . . . . . . . . . . . . . . . . . . . . . . . . .
33
6.4
Sea ice mechanism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
33
6.5
Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
35
7 Overall discussion and conclusions
38
vii
List of Abbreviations
AABW . . . . . . . . . . . Antarctic Bottom Water
ACC . . . . . . . . . . . . . Antarctic Circumpolar current
AMOC . . . . . . . . . . . Atlantic meridional overturning circulation
AOGCM . . . . . . . . . atmosphere-ocean general circulation model
DP . . . . . . . . . . . . . . . Drake Passage
EMIC . . . . . . . . . . . . Earth System Model of Intermediate Complexity
IPCC . . . . . . . . . . . . . Intergovernmental Panel on Climate Change
MOC . . . . . . . . . . . . . Meridional overturning circulation
NADW . . . . . . . . . . . North Atlantic Deep Water
SO . . . . . . . . . . . . . . . Southern Ocean
SRES . . . . . . . . . . . . . Special Report on Emissions Scenarios
THC . . . . . . . . . . . . . Thermohaline circulation
WBC . . . . . . . . . . . . . Western Boundary current
viii
1
1
Introduction
The global meridional overturning circulation (MOC), a system of large-scale deep currents encompassing
the entire world ocean, is probably the oceanic feature most vital to our present global climate (figure 1).
Particularly its Atlantic branch has been a research focus for many years because of its various functions in
the climate system with respect to, for example, heat and volume transport and the hydrological (Vellinga
and Wood , 2002, 2007) and biogeochemical cycles (Zickfeld et al., 2007b). Recent studies have suggested
that abrupt reorganizations of the Atlantic MOC (AMOC) have been involved in major climatic changes
in the geological past (Bond et al., 1992; McManus et al., 2004), and raised concerns about a possible
slow-down, or even abrupt collapse, of the AMOC due to future anthropogenic greenhouse gas emissions
(Manabe and Stouffer , 1994; Rahmstorf and Ganopolski, 1999; Wood et al., 1999; Schaeffer et al., 2002;
Bryden et al., 2005; Zickfeld et al., 2007c). Most likely, this would have major implications for global and
regional climatic aspects, including a possible suppression of the El Niño-Southern Oscillation (ENSO)
phenomenon (Timmermann et al., 2005); a southward shift of the Intertropical Convergence Zone over
the Atlantic sector (Vellinga and Wood , 2002); a decline of the marine ecosystem (Schmittner , 2005);
and a rise of sea level in the North Atlantic (Levermann et al., 2005). The MOC is described in more
detail in section 3.
Despite of the importance of the AMOC, the question is still unresolved of which physical mechanisms
ultimately drive it, in the sense of providing the energy for sustaining a steady-state circulation (see
Kuhlbrodt et al. (2007) and references therein). There extists a controversy over two fundamental theories:
One of a circulation driven mainly by vertical downward diffusion of heat in low latitudes (Sandström,
1916; Jeffreys, 1925; Munk and Wunsch, 1998); and a second one involving wind-driven upwelling in the
Southern Ocean due to the so-called Drake Passage effect (Toggweiler and Samuels, 1993a, 1995, 1998).
The two processes could imply different sensitivities to external forcing and thus different evolutions
of the AMOC under future global climate change (Schmittner and Weaver , 2001; Prange et al., 2003).
Knowledge of their relative importance is therefore crucial for a comprehensive assessment of future risks
of a collapse of the AMOC, and investigation into their interactions is of great interest even beyond
scientific curiosity.
There are further processes, such as surface fluxes of heat and salinity, which influence the magnitude
and structure of the MOC without adding net energy to the system. Therefore, in addition to the question
of the energy source, the question has to be addressed which parameters set the actual flow pattern of
the circulation. On a global scale, this concerns first of all the relative strength of the overturning in the
individual ocean basins.
The motivation for the present study is to put the theory of a wind-driven AMOC to the test in the
global coupled climate model of intermediate complexity CLIMBER-3α. To this end, we have analysed the
2
1
INTRODUCTION
equilibrium response of the model to variations in the surface wind stress upon the Southern Ocean. Our
results not only suggest that wind-driven Southern Ocean upwelling is the principal driver of the AMOC
in our model; they also point towards a coherent picture of the conditions determining the distribution
of the upwelling onto the ocean basins. Moreover, they reveal a novel effect of reduced Southern Ocean
winds on surface temperatures.
This study is organized as follows: The coupled climate model is briefly described in section 2. Section
3 introduces some basic oceanographic concepts concerning the meridional overturning circulation, as well
as the two main theories with respect to its driving mechanism. In section 4, the Drake Passage effect is
explained and its functionality is demonstrated using results of our model simulations.
The main results are presented in sections 5 and 6. Section 5 deals with the mechanisms that determine
the strength and structure of the global MOC in our model. Southern Ocean upwelling is found to be
proportional to the wind stress gradient. We observe that, for present-day boundary conditions, the net
upwelling equals the outflow from the Atlantic basin, and conclude that the AMOC is completely winddriven. The distribution of Southern Ocean upwelling among the ocean basins is controlled by meridional
density gradients. We outline a comprehensive explanation on the grounds of classical boundary layer
theory. The results of this section are currently being prepared for separate publication (Schewe and
Levermann, 2008).
In section 6, we describe a previously unrecognized effect of reduced Southern Ocean winds on the
climate system: Instead of the so-called bipolar see-saw pattern of surface temperature change that is
usually associated with an AMOC reduction, we observe the declining AMOC to be accompanied by
surface cooling in both hemispheres. We give an explanation related to changes in the export of sea
ice from the Antarctic ocean. Possible implications regard paleoclimatic events as well as present-day
climate sensitivity (i.e., the equilibrium change in global mean surface temperature following a doubling
of the atmospheric equivalent CO2 concentration). Recent model simulations, performed for the Fourth
Assessment Report of the Intergovernmental Panel on Climate Change (IPCC) (Meehl et al., 2007a),
suggest a strengthening of surface wind stress in response to atmospheric CO2 increase under certain
climate change scenarios (Fyfe and Saenko, 2006). We have published the results of this section in
Geophysical Research Letters (Levermann et al., 2007).
3
Figure 1: Strongly simplified sketch of the global meridional overturning circulation
(MOC), from Kuhlbrodt et al. (2007), after Rahmstorf (2002). In the Atlantic, warm,
saline surface waters from the tropics are exported northwards via the Gulf stream and
the North Atlantic current (red). Due to surface cooling, they sink in the Nordic Seas
and the Labrador Sea to form North Atlantic Deep Water (NADW), which returns
southward at depth (light blue). In the Southern Ocean, the Antarctic Circumpolar
current (ACC) spans the entire globe and connects all other basins. Sinking regions
also exist south of the ACC; the deep water masses formed here are denser than
the NADW and thus circulate at greater depths (dark blue). The upwelling which
must balance the sinking at these localised sites could in principle be provided by two
processes, low-latitude vertical mixing and wind-driven Southern Ocean upwelling,
both of which are indicated here; their relative importance is a main subject of this
study.
4
2
2
MODEL DESCRIPTION
Model description
A large variety of numerical models is presently used for the study of the climate system, ranging from
simple zero-dimensional energy balance models, to zonally (i.e., in east-west direction) averaged box
models, through to three-dimensional coupled general circulation models of various spatial and temporal
resolutions (Meehl et al., 2007b). The computational cost of a given model is, of course, largely determined
by its resolution and by the complexity of the physical processes included. Applicability for studies of
large-scale, long-term climate phenomena, on the other hand, requires a certain computational efficiency.
Simple reduced models, which often consist only of a few general equations and may represent whole
ocean basins as a single box, are used to address questions of rather conceptual nature. They allow
comprehensive analyses of the parameter space and integrations over large time spans (e.g. several glacial
cycles). On the other end of the model spectrum, high-resolution, three-dimensional atmosphere-ocean
general circulation models (AOGCMs) are capable of realistically reproducing oceanic or atmospheric
flow patterns and individual weather systems, and the complex interactions between different climate
subsystems. They are thus suitable for, e.g., projections of near-future climate evolution under climate
change scenarios. However, their high complexity and resolution make it difficult, if not impossible,
to integrate them sufficiently long to reach a stable equlibrium - let alone, to conduct simulations on
paleoclimatic timescales. They have to be initialized with observationally based data sets. This can be
problematic because erroneous data can bias the results if the spin-up integration is too short. Also, the
time evolution of the results may be represented incorrectly if the inertia of the system is not reproduced
well. Only if models are thoroughly run into equilibrium, they can provide results which are truly
independent of initial conditions and possibly unknown inertial feedbacks, and thus allow inferences to
more fundamental relationships.
The global coupled climate model CLIMBER-3α (Montoya et al., 2005) belongs to the class of Earth
System Models of Intermediate Complexity (EMICs). EMICs are designed to bridge the gap between
conceptual and comprehensive models. They describe most processes covered by the more complex
models, albeit in a reduced, more parameterized form. They combine several components of the climate
system, e.g., ocean, atmosphere, cryosphere, vegetation and biogeochemical cycles, and explicitly simulate
their interactions. On the other hand, their relative simplicity and coarse resolution permits integrations
over several 10000 years or even (e.g. in the case of CLIMBER-2, the predecessor of CLIMBER-3α) full
glacial cycles. Thus, stable equilibrium states can be simulated, as was done for the present study.
CLIMBER-3α combines a three-dimensional ocean general circulation model based on the Princeton
MOM-3 code (Pacanowski and Griffies, 1999) with the 2.5-dimensional statistical-dynamical atmosphere
model POTSDAM-2 (Petoukhov et al., 2000) and a dynamic and thermodynamic sea-ice model (Fichefet
and Maqueda, 1997). The oceanic horizontal resolution is 3.75◦ × 3.75◦ (see figure 2) with 24 variably
5
Figure 2: Coastlines and bottom topography of the oceanic model grid of CLIMBER3α. Note that the 24 depth levels are variably spaced, see table 1. For a detailed
model description see Montoya et al. (2005).
spaced vertical levels, ranging from 25m in the upper four layers, to ∼ 500m in the lowest layer (table 1).
In order to reduce numerical diffusion a second order moment tracer advection scheme is applied (Prather ,
1986; Hofmann and Maqueda, 2006). The atmospheric grid has a resolution of 22.5◦ in longitude and
7.5◦ in latitude and is congruent with the oceanic grid. The time step of all model components is 1/2
day.
Slightly different model versions are used in sections 5 and 6.
The experiments and analyses
presented in section 6 were conducted earlier in the course of my diploma thesis than those presented
in section 5. In the meantime, the model was improved compared to the previous version, which is
described extensively in Montoya et al. (2005) together with a comparison to present-day observations.
The improvements include a deeper Indonesian throughflow, wind stress prescribed according to the
reanalysis data of Trenberth et al. (1989), and a background vertical diffusivity of 0.3 cm2 /s instead of
0.1 cm2 /s. A main difference in the output of the two model versions is that the rate of the meridional
overturning circulation is generally higher in the improved version used in section 5. However, the
qualitative results presented in this study are robust against these changes.
6
2
MODEL DESCRIPTION
Table 1: Depths of grid points in the oceanic component of CLIMBER-3α, and corresponding resolutions.
Level
Depth (m)
∆z (m)
Level
Depth (m)
∆z (m)
1
2
3
4
5
6
7
8
9
10
11
12
12.50
37.50
62.50
87.50
112.50
140.54
177.50
229.21
301.00
397.76
523.75
682.52
25
25
25
25
26.50
32.50
44.34
61.74
84.27
111.37
142.38
176.53
13
14
15
16
17
18
19
20
21
22
23
24
876.81
1,108.48
1,378.48
1,686.81
2,032.52
2,413.75
2,827.76
3,271.00
3,739.21
4,227.50
4,730.51
5,242.50
212.98
250.83
289.16
327.01
336.46
397.46
428.62
455.72
478.25
478.25
507.49
513.49
7
3
The meridional overturning circulation
3.1
The schematic of an oceanic conveyor belt
The global meridional overturning circulation (MOC) is a system of large scale oceanic currents that
encompasses all basins of the world ocean and whose principal flow is in meridional (north-south) direction. It is portrayed in a simplified sketch in figure 1. Its most prominent part is the Atlantic meridional
overturning circulation (AMOC), often termed the oceanic conveyor belt (Broecker , 1991). The most
important water masses and processes associated with the Atlantic part of the MOC are illustrated in
figure 3. The AMOC consists of a northward surface branch, comprising the Gulf Stream and the North
Atlantic current, and a deep southward return flow along the western boundary of the Atlantic ocean.
It is of great importance to the global climate, particularly due to its role as a giant heat pump: Warm,
salty surface waters from the Southern Ocean and the tropical Atlantic are exported to the north at a
rate of 15 ± 3 Sv (Sverdrups, 1 Sv = 106 m3 /s) (Talley, 2003; Ganachaud and Wunsch, 2000)), where they
cool and subsequently sink to depth to form North Atlantic Deep Water (NADW) (Dickson and Brown,
1994; Brix and Gerdes, 2003). Thus, about 1 P W (= 1015 W ) of heat is transported to the North Atlantic (Hall and Bryden, 1982; Ganachaud and Wunsch, 2000; Trenberth and Caron, 2001), significantly
raising water and air temperatures in northern high latitudes and thereby contributing to a relatively
mild regional climate, particularly in winter. This redistribution of heat is a major global climatic factor.
Moreover, the AMOC provides for the ventilation of the deep ocean. The deep water formation
associated with the AMOC transfers surface properties to depths of up to several thousand meters,
while the compensating upwelling lifts nutrient rich deep waters to the surface. The timescale of this
ventilation is of the order of millennia and depends on the strength of the AMOC. It is crucial for the
deep ocean’s function as a storage of heat, nutrients, and chemical components such as carbon dioxide
(CO2 ). Presently, for example, the Atlantic ocean acts as a net sink of anthropogenic CO2 and thus
slows down man-made global warming (Sabine et al., 2004; Zickfeld et al., 2007a). At the same time, the
circulation sets the living conditions for numerous marine ecosystems, which fisheries around the world
are highly dependent on (Schmittner , 2005).
The AMOC can be visualized in numerical climate models by displaying the so-called meridional
overturning streamfunction Ψ (Gill , 1982). It is defined by ∂y Ψ = ρ0 w̄ and ∂z Ψ = −ρ0 v̄, where v̄, w̄ are
the meridional and vertical velocity, respectively, integrated zonally across the entire basin. The upper
panel in figure 4 shows such a streamfunction1 of the Atlantic overturning for the standard configuration of
CLIMBER-3α. The contours can be interpreted as flowlines, with 1 Sv of mass transport between any two
1 Because of continuity, the streamfunction is only defined within closed zonal boundaries; in particular, there is no
meaningful streamfunction representation of the circulation in any open sector of the Southern Ocean. This is why figure 4
only displays the Atlantic and Indopacific basins north of 30◦ S. For a global streamfunction including the Southern Ocean,
see figure 6b.
8
3 THE MERIDIONAL OVERTURNING CIRCULATION
Figure 3: Idealized sketch representing a zonally averaged picture of the Atlantic
and the Southern Ocean, from Kuhlbrodt et al. (2007). The color shading depicts
a zonally averaged density profile derived from observational data (Levitus, 1982),
where blue is light and yellow is dense. Straight arrows indicate the main volume
transport associated with the AMOC. Two principal water masses circulate in the
deep ocean: North Atlantic Deep Water (NADW), which forms in the Nordic Seas
due to atmospheric cooling of salty surface waters and, after flowing over a shallow sill
between Greenland and Scotland, sinks to greater depths in the Atlantic basin; and
Antarctic Bottom Water (AABW), which forms in the Weddell and Ross Seas in the
Southern Ocean. These deep water masses are separated from the lighter and warmer
surface waters by a region of large temperature gradient, the thermocline. Breaking
of internal waves, generated at the surface as well as over bottom topography, leads
to upwelling through the thermocline via turbulent mixing along the density gradient
(diapycnal mixing). In the Southern Ocean, deep waters are upwelled by surface wind
stress divergence due to the Drake Passage effect (section 4). Part of these waters
recirculate within the so-called Deacon cell. Note that this sketch is only meant to
illustrate some terms and concepts mentioned in the text; the processes depicted are
highly simplified. Note also that the typical ratio of meridional to vertical extent of
the real ocean is of the order of 5000.
3.1 The schematic of an oceanic conveyor belt
9
Figure 4: Meridional overturning streamfunction of the Atlantic (top) and Indopacific
(bottom) basins, in Sv, for the standard configuration of CLIMBER-3α. Positive
(negative) values and red (blue) shading indicate clockwise (anticlockwise) circulation.
Spacing between contours is 1 Sv.
adjacent contours, and positive (negative) values indicating clockwise (anticlockwise) circulation. Simply
counting the contour lines crossing, e.g., the equator, it shows that the interhemispheric circulation in the
upper and intermediate Atlantic amounts to some ∼ 10 Sv, which flow northwards above, and southwards
below, about 800m. The depth level separating the northward branch from the southward branch of the
circulation is called the level of no mean motion.
Below this clockwise overturning cell (red in figure 4), there is a cell of Antarctic Bottom Water
(AABW) (Orsi et al., 1999; Brix and Gerdes, 2003) rotating in the opposite direction (blue), with about
4 Sv flowing into the basin at the bottom, and leaving it between 2600m and 3500m.
The Pacific and Indian branches of the MOC are much less pronounced in our present-day model
configuration than the Atlantic branch, which is supported by observations (Ganachaud and Wunsch,
2000; Talley, 2003). While an intermediate overturning cell exists in the North Pacific, the net interhemispheric volume transport in the upper and intermediate Indopacific is virtually zero; only well below
1000m, there is some circulation of Antarctic Bottom Water (figure 4, lower panel). However, the Pacific
10
3 THE MERIDIONAL OVERTURNING CIRCULATION
and Indian oceans do come into play for different surface boundary conditions (see section 5). The Atlantic, Pacific and Indian branches of the MOC are connected via the Southern Ocean, a continuous band
of water encompassing the globe between Antarctica and the Southern tips of the American, African,
and Australian continents. It allows for the existence of a strong wind-driven zonal current, the so-called
Antarctic Circumpolar current (ACC), encircling the Antarctic continent at a rate of about 120 − 160 Sv
(Zlotnicki et al., 2007). Once the outflow from one of the adjacent basins enters the Southern Ocean, it
is rapidly advected by the ACC and spread across the entire Southern Ocean.
The sinking associated with the MOC is rather localised. There are two deep water formation regions
in the North Atlantic, one in the Labrador Sea southwest of Greenland, the other in the Nordic Seas
between Greenland, Iceland and Norway (compare figure 1). Salty surface waters that enter the South
Atlantic Ocean are advected northward by the Gulf Stream and the North Atlantic current, and give
off heat to the atmosphere as they reach high northern latitudes. This makes them denser than the
surrounding, fresher water masses and ultimately causes them to sink. The positive feedback implied by
the fact that the AMOC itself provides for the high salinity necessary to enhance sinking is known as the
Stommel-feedback, and is the main reason for the existence of a stable off-state of the AMOC (Stommel ,
1961). The deep water formation in the Weddell and Ross Seas in the Southern Ocean is largely due to
brine rejection associated with the annual build-up of sea ice; the cooling that occurs during the opening
of so-called polynyas, large ice-free patches in the sea-ice cover, also plays an important role (Maqueda
et al., 2004).
3.2
Upwelling waters and AMOC driving mechanisms
Unlike the sites of deep water formation, much less is known about the spatial distribution of the
upwelling that is needed to balance it and close the circulation in the vertical. Vertical velocities in
the ocean generally exhibit large variability on all time scales, and it is virtually impossible to directly
measure their net long-term components for any larger region. Deep water formation rates can be
estimated from inventories of chlorofluorocarbons, anthropogenic substances which have entered the
surface ocean within the last decades (Smethie and Fine, 2000). Obviously, this method is not applicable
to the upwelling of deep, old water masses. Moreover, upwelling takes place much less localised than
sinking, and is thus much more difficult to observe. Analyses of prebomb radiocarbon distributions
(Toggweiler and Samuels, 1993b), biogenic silica production rates (Gnanadesikan and Toggweiler , 1999),
and other tracers (Wunsch et al., 1983; Robbins and Toole, 1997) have provided evidence for substantial
upwelling of abyssal water masses in the Southern Ocean. High-resolution model simulations support
these observations (Döös and Coward , 1997). However, observational data is sparse, which makes
comparison with model estimates or theoretical considerations difficult.
3.2 Upwelling waters and AMOC driving mechanisms
11
The question of where the upwelling takes place is closely connected to the question about the mechanism providing for the upwelling. Two distinct mechanisms are presently being discussed. One was
proposed by Sandström (1916) and Jeffreys (1925) and described in detail by Munk and Wunsch (1998).
In their view, the formation of deep water masses is balanced by mixing-induced diapycnal upwelling in
the low latitudes of the Atlantic. The term diapycnal refers to mixing across surfaces of equal density
(isopycnals) and is used in distinction to isopycnal mixing, which requires much less energy. The breaking
of internal waves generated by winds and tides leads to turbulent mixing in the upper ocean layers (Munk
and Wunsch (1998); Huang (1999);Wunsch and Ferrari (2004)). The subsequent downward diffusion of
heat in low latitudes, where the atmosphere is generally warmer than the ocean, lightens deeper layers
of water and causes them to rise. The resulting surface and intermediate waters are then again advected
northward, where they are transformed into denser water masses in the NADW formation regions, and
finally sink and slowly flow southward to close the loop (compare figure 3). A meridional density gradient between low and high latitudes is thus established. This is the traditional thermohaline driving
mechanism, referring to a circulation controlled by fluxes of heat and salinity. The terms thermohaline
circulation (THC) and meridional overturning circulation (MOC) are sometimes used interchangeably.
However, they are not synonymous, because THC refers to a specific driving mechanism, while MOC
describes the flow pattern geometrically (Wunsch, 2000; Rahmstorf , 2003). We will consequently use the
neutral terms MOC or AMOC.
The second mechanism was put forward by Toggweiler and Samuels (1993a, 1995, 1998) upon the
observation that the actual amount of diapycnal mixing in the central Atlantic was not sufficient to
sustain an estimated AMOC rate of about 15 to 20 Sv. Munk (1966) and Munk and Wunsch (1998)
estimated that an average diapycnal diffusivity of 1cm2 s−1 was necessary to balance the total sinking
in the North Atlantic and the Southern Ocean. On the other hand, observations suggest a coefficient of
only 0.1cm2 s−1 in the ocean interior (Moum and Osborn, 1986; Ledwell et al., 1993; Oakey et al., 1994),
which would translate into 3 Sv of global upwelling (Webb and Suginohara, 2001). Although much higher
mixing rates are found near highly variable bottom topography or at continental slopes (Polzin et al.,
1997; Ledwell et al., 2000; Garabato et al., 2004), it has been doubted that this could raise the global
average diffusivity by an order of magnitude (Webb and Suginohara, 2001). Toggweiler and Samuels
(1995) suggested that most of the upwelling instead occurs in the Southern Ocean, driven by surface
wind stress via the so-called Drake Passage effect. Drake Passage (DP) is the name of the opening
between the tip of South America and the Antarctic Peninsula. The unique dynamic constraint imposed
by the existence of an uninterrupted zonal band of water in the latitudes of DP sets the conditions for
wind-driven upwelling from depths of several thousand meters. The DP effect is explained in the following
section.
12
3 THE MERIDIONAL OVERTURNING CIRCULATION
When reflecting on driving mechanisms, it is important to distinguish between the energy source
required to maintain a steady-state circulation, and other factors that exert control on the magnitude and
structure of the circulation without actually powering it. Although the fluxes of salt and heat (subsumed
under the term buoyancy fluxes) associated with deep water formation certainly play an important role
in setting the overall pattern of the present-day oceanic circulation, they are not able to sustain a largescale, steady-state deep circulation (Sandström, 1916; Kuhlbrodt et al., 2007). The required input of
mechanical energy is therefore thought to be associated with the complementary processes of upwelling,
be it turbulent mixing in low latitudes or wind forcing in the south. In both cases, the energy ultimately
originates from the influence of the sun and the moon, and is transferred by the action of winds and
tides. The difference is that in the case of wind-driven Southern Ocean upwelling, the wind stress itself
drives large-scale motions of water masses from the abyssal ocean towards the surface, while in the case
of low-latitude mixing, the perturbations caused by winds and tides dissipate into small-scale turbulent
motions and thus facilitate the downward mixing of heat.
13
4
Geostrophy and the Drake Passage Effect
4.1
Geostrophic balance
Much of the complexity of atmospheric and oceanic dynamics ultimately arises from the fact that the
~ ≃ 2πday −1 (neglecting the small effect of
Earth is rotating. This rotation with an angular velocity Ω
the Earth’s rotation around the sun) requires the addition of a Coriolis term to the momentum equation
of any moving fluid, be it water or air. The main part of the governing equations of ocean dynamics,
namely, the continuity equation and the three-dimensional momentum equation, thus read2
ρ−1
0
Dρ
+ ∇ · ~u = 0
Dt
D~u
~ × ~u = −ρ−1 ∇p − ~g + ν ∇2 ~u
+ 2Ω
0
Dt
(1)
(2)
with the velocity field ~u ≡ (u, v, w), pressure p, gravitational acceleration ~g ≡ g~ez , and average density
ρ0 , where the Boussinesq approximation has been made. Turbulent effects are approximated here as a
diffusion process with the kinematic viscosity ν, which is a good approximation on scales where viscosity
changes and compressibility can be neglected (Gill , 1982).
Given the large ratio of horizontal to vertical length scales in the ocean, the horizontal equations for
a steady-state solution reduce to
2
−f v = −ρ−1
u
0 ∂x p + ν ∇ ~
(3)
2
f u = −ρ−1
u
0 ∂y p + ν ∇ ~
(4)
with3 f = 2Ω sin φ. A further simplification is possible where viscous effects are negligible. In general,
this is the case away from surfaces and boundaries. The last term on the r.h.s. of (3) and (4) can then
be set to zero, and one obtains what is called the geostrophic balance, where horizontal flows are entirely
balanced by horizontal pressure gradients. The observation that the flow is at right angle to the pressure
gradient balancing it is well known, for example, from the atmosphere with its cyclonic and anticyclonic
circulations around pressure extrema.
In the ocean, pressure gradients can be established, for example, by density gradients or by differences
in sea surface elevation, or by the existence of continents, which balance any pressure imposed on their
boundaries. The North Atlantic current, for instance, which crosses the Atlantic ocean from the North
American coast towards the Nordic Seas, is balanced by a sea level difference across the current of the
order of 1m (Levermann et al., 2005).
2 It is common in earth sciences, unless processes close to the poles are examined, to use a rectangular coordinate system
with the x and y axes pointing eastwards and towards the North Pole, respectively, and z pointing towards the center of
the Earth - the so-called β-plane approximation (Gill , 1982).
3 The relation between y and the spherical latitude coordinate φ is linear: φ = y/r
earth .
14
4
4.2
GEOSTROPHY AND THE DRAKE PASSAGE EFFECT
The Drake Passage effect
The latitude band of Drake Passage is unique on earth in that it contains no continental boundaries in
the zonal direction; the highest topographic ridges lie at a depth of about 2500m (about 1380m in our
model). This allows for a closed zonal current, the strong Antarctic Circumpolar current (ACC). The
lack of continental boundaries also means that no net zonal pressure gradient can exist in these latitudes,
because both density and sea level differences amount to zero when integrated around the Southern
Ocean. Thus, according to (3), no net, geostrophically balanced meridional flow can establish in this
region above 2500m:
I
f v = ρ0−1
I
∂x p = 0
(5)
where the line integral represents any closed contour in the latitudes of the DP and between the surface
and the highest topographic point at the respective latitude.
At the same time, strong westerly winds cause waters in the surface layer of the Southern Ocean to
move northwards. This phenomenon, called Ekman transport (Ekman, 1905; Price et al., 1987), is again
due to the action of the Coriolis effect. The transport rate is Sy = τx /(f ρ), where τx is the eastward
surface wind stress. Because the wind stress in the latitudes of DP has a strong meridional divergence
(compare figure 7a), the Ekman transport is also divergent, i.e. gets larger towards the equator. To
balance this divergence, water from below the surface layer has to be pumped up at a rate which is
proportional to the wind stress gradient:
SEk = ∂y Sy = (βρ0 )−1 ∂y τx
(6)
where β = ∂y f .
Strong, divergent winds occur at other latitudes, too, for example as trade winds north and south of
the equator. There, the upwelling caused by Ekman pumping is fed by a return flow just below the surface
layer (figure 5a). That way, shallow overturning cells - the so-called Ekman cells - are formed, that reach
no deeper than a few hundred meters and do not involve any significant transformation of density. In
the latitudes of DP, however, no net geostrohically balanced return flow can develop in the upper ocean.
The Ekman pumping thus reaches down to greater depths, where water can flow southward along the
flanks of topographic ridges (figure 5b). This is how the Drake Passage effect leads to the upwelling of
deep water masses. These water masses are substantially denser than those transported northward at
the surface, which creates the need for a compensating transformation of light to dense water somewhere
else in the ocean. Toggweiler and Samuels (1995) suggested on the grounds of this consideration that
15
4.2 The Drake Passage effect
there must be a connection between Southern Ocean upwelling and NADW formation.
depth
a
DP
b
S
Surface Ekman transport
Return flow
N
Figure 5: Simplified sketch of a south-north section through an ocean basin, illustrating the occurance of (a) shallow Ekman cells and (b) deep upwelling due to divergent
northward Ekman transport at the surface. The opening labelled ’DP’ represents the
zonal gap of Drake Passage.
In the following, because of (6), we will use the quantity ∂y τx (more precisely, the maximum of
h∂y τx ix in DP latitudes) as reference. However, the results of our study do not change much qualitatively
if one substitutes τx for ∂y τx , and we will sometimes use the terms wind stress or winds instead of wind
stress gradient.
The effect can be observed rather clearly in our coupled climate model. It is illustrative to plot the
R NP
H
zonally integrated upwelling budget through a given depth level h, defined as y dy ′ dx w(z = h),
where NP denotes the North Pole. Figure 6a shows such a budget for the level of no mean motion,
h = 780m, and thus represents the sinking and upwelling associated with the global MOC. The most
rapid upwelling is observed in DP latitudes. Two rather narrow, closed overturning cells are visible as
negative excursions: One between 60◦ N and 30◦ N , representing the recirculation of a part of the Deep
Water formed in the northern North Atlantic, which we will discuss later on; and a stronger one between
40◦ S and 60◦ , which is called the Deacon cell (Döös and Webb, 1994; Speer et al., 2000). Both these cells
can also be identified in the corresponding global meridional overturning streamfunction (6b). Neither
of them contribute to the interhemispheric volume transport. When they are subtracted, the global
upwelling budget basically consists of about 15 Sv sinking in high northern latitudes, about half of which
are slowly upwelled in low and mid latitudes; the other half is then compensated by upwelling due to the
DP effect. South of the DP, there is some circulation of AABW. The rates of both the upwelling and
the interhemispheric circulation scale with the Southern Ocean wind stress, while the general pattern
is preserved (compare dashed and dot-and-dash lines in figure 6a; see also section 5.3). Note that the
16
4
GEOSTROPHY AND THE DRAKE PASSAGE EFFECT
discussion above refers to the global MOC, while the role of the DP effect for the Atlantic MOC will be
discussed later.
10
0
Sv
−10
−20
−30
−40
a
−50
0
−1000
m
−2000
−3000
−4000
−5000
b
−80
−60
−40
−20
0
20
Latitude (°N)
40
60
80
Figure 6: (a) Global upwelling budget through the 780m depth level for the standard
configuration (solid line) and for 0.5 (dot and dash) and 1.5 (dashed) times the observed wind stress field. The main upwelling region is located in the Southern Ocean
around 60◦ S. The location of Drake Passage is indicated by blue lines. (b) Global
meridional streamfunction for the standard configuration. The DP opening (blue)
and the 780m depth level (dashed blue) are indicated.
17
5
Control mechanisms of the MOC
In this section, we examine experiments with Southern Ocean wind stress ranging from 0.2 to 1.5 times
presently observed fields, combined with varying surface freshwater fluxes to the North Atlantic and
North Pacific. We show that deep upwelling in the Southern Ocean is proportional to the surface wind
stress in the latitudinal band of Drake Passage. At the same time, large scale meridional overturning
circulation in each basin is determined by the respective meridional density gradient. The distribution
of the Southern Ocean upwelling onto the oceanic basins is thus controlled by buoyancy distribution; the
inflow into each basin being proportional to the respective meridional density gradient. The constant
of proportionality is the same for all basins. For strongly reduced wind stress in the Southern Ocean,
the circulation enters a regime where Atlantic overturning is maintained through Pacific upwelling, in
order to satisfy the transports set by the density gradients. We therefore propose that both Southern
Ocean upwelling and meridional density gradients set up a system of conditions that determine the global
meridional overturning circulation.
5.1
Introduction
There has been continued discussion about what controls the strength of the Atlantic meridional overturning circulation (AMOC). A common view is that the formation of North Atlantic Deep Water (NADW)
is balanced by diapycnal upwelling in the low latitudes of the Atlantic, driven by downward diffusion of
heat due to turbulent mixing in the upper ocean layers (Munk and Wunsch (1998); Huang (1999);Wunsch
and Ferrari (2004)). Various conceptual models have been designed which reflect this general picture.
Especially in zonally averaged models, meridional density (or pressure) differences have often been assumed to be key in setting the overturning rate of the AMOC (e.g. by Stommel (1961); see also Wright
et al. (1998) and references therein). Rahmstorf (1996) was the first to show a linear relation between
meridional density gradients and the overturning rate in an ocean general circulation model (OGCM) in
response to freshwater perturbations. The reasoning behind this is that meridional pressure gradients
in the North Atlantic create a primarily westward flow which, once encountering the flank of the North
American continent, continues southward as a Western Boundary current, with friction balancing the
Coriolis pseudo force (Marotzke, 1997).
An alternative theory (Toggweiler and Samuels (1993a, 1995, 1998)) involves wind-driven deep upwelling in the Southern Ocean due to the so-called Drake Passage effect (see section 4): Strong westerlies
induce a divergent northward Ekman transport in the latitudes of Drake Passage. Because of the lack
of continental boundaries in these latitudes, the net zonal pressure gradient above the sill of the Passage
must be zero, and no net geostrophically balanced north-south flow can develop to compensate for the
northward surface transport. Therefore, waters from below the sill are drawn up to the surface. The
18
5
CONTROL MECHANISMS OF THE MOC
AMOC can be seen as the closing branch of this wind-driven overturning. Nof (2003), upon integrating
the momentum balance of a water column along a closed contour around the Southern Ocean, came to
the conclusion that the Atlantic inflow and outflow is controlled exclusively by Southern Ocean wind
stress.
Griesel and Morales-Maqueda (2006) have recently conducted various experiments with an OGCM
coupled to a simple moisture balance model, and pointed out that both Southern Ocean winds and
meridional density/pressure gradients play a role in determining the strength of the Atlantic MOC. In
the following, we attempt to extend and generalize this idea, and give a picture of the global MOC in
which both wind-driven Southern Ocean upwelling and meridional density gradients combine to set the
magnitude and structure of the circulation. To this end, we analyse experiments in a coupled coarseresolution climate model with both attenuated and amplified Southern Ocean wind stress, as well as
simulations with anomalous freshwater flux applied to the North Atlantic. Since our model has a very
low background vertical diffusivity and thus operates close to the limit of no diapycnal mixing in the
ocean interior, the AMOC is, under present-day boundary conditions, entirely driven by Southern Ocean
winds. However, different regimes are associated with different boundary conditions. As SO winds are
reduced, the other ocean basins come into play and partly overtake the Southern Ocean’s role of balancing
North Atlantic sinking. Adversely, for enhanced winds, the Pacific and Indian basins contribute to the
enhanced upwelling with a net outflow.
We show that, independent of the regime, the meridional density gradient along the Western Boundary
current (WBC) in each of the basins is a measure for the magnitude of the respective branch of the global
MOC. A universal linear relation exists between the overturning rate and the density gradients. An
explanation is given on the basis of classical boundary layer theory, whose application is justified by
the fact that the main meridional volume transport in our model is associated with WBCs. This also
means that the concept presented here is independent of the mechanism of water mass transformation
that closes the circulation, because only the density distribution within the WBC is considered. The
meridional density gradients are altered by changing SO winds through the upwelling of deep, dense
water masses; but also by other forcings that may be applied to modulate the MOC, such as anomalous
freshwater fluxes in the North Atlantic. We conclude that there is no contradiction between a wind-driven
and a buoyancy flux-driven circulation.
5.2
Experiments
The global coupled climate model CLIMBER-3α is described in section 2. Here, we use an improved
version of the model, comprising a deeper Indonesian throughflow and wind stress prescribed according
to the reanalysis of Trenberth et al. (1989), and apply a very low background value of vertical diffusivity
5.3 Wind-driven Southern Ocean upwelling
19
(0.3 · 10−4m2 /s). Thus the mixing induced upwelling in both the Atlantic and Pacific ocean is very small
(Mignot et al., 2006).
For the present study, the zonal oceanic wind stress component τx between 71.25◦S and 30◦ S is
modified following τx (x, y) = α · τx0 (x, y), where τx0 (x, y) is the standard wind stress according to Trenberth
et al. (1989). Experiments were conducted for α = 0.2, 0.5, 1.0 and 1.5, and thus span a range of
both reduced and enhanced forcing conditions (figure 7a). Since the meridional distribution of τx , in
the aforementioned latitudinal interval, has a quasi-sinusoidal shape (with close-to-zero wind stress at
71.25◦S and 30◦ S, and maximum eastward wind stress at about 51◦ S), this implies a modulation of both
the wind stress magnitude, τx , and the wind stress gradient, Ty ≡ ∂y τx . In particular, in the latitude
band of Drake Passage, between 62◦ S and 56◦ S, the maximum of the wind stress gradient Ty varies from
4.1 · 10−8 N/m3 in the α = 0.2 experiment, to 29.9 · 10−8 N/m3 in the α = 1.5 experiment. For each
experiment, the model was run for several thousand years until an equilibrium state was reached.
For comparison, three additional runs were carried out with freshwater flux (FWF) into the North
Atlantic enhanced by 0.35 Sv, and with a compensating negative FWF into the North Pacific. In these
runs, the observed SO wind stress field was multiplied by α = 0.5, 1.0 and 1.5, as described above.
5.3
Wind-driven Southern Ocean upwelling
The total integrated upwelling in the Southern Ocean (figure 8, diamonds) is proportional to the wind
stress, varying from about 24 Sv for α = 1.5, to about 2 Sv for α = 0.2, and approximately crossing the
origin when extrapolated to α = 0. This upwelling is computed from the global meridional overturning
streamfunction (compare, e.g., figure 7b) at 30◦ S at a depth of 780 m, which corresponds to the approximate level of no mean motion between the upper northward flow into the Atlantic ocean and the outflow
of North Atlantic Deep Water in our model. This depth is fairly constant throughout our experiments.
Part of this total upwelling is returned to depth within the SO as Antarctic Bottom Water (AABW).
Thus, the rate of water exported northwards from the SO is equivalent only to a net SO upwelling,
defined as the north-south streamfunction difference across the Southern Ocean south of 30◦ S (figure
8, triangles), whereby the AABW cell is cancelled out. This net upwelling can then be decomposed by
computing the Southern Ocean outflow (i.e., the flow entering the SO from the north) at 30◦ S below
780m from the Atlantic (figure 8, circles) and Indopacific (squares) ocean basins. Both scale linearly with
the wind stress gradient. In the standard configuration, all the water upwelled in the Southern Ocean
is drawn out of the Atlantic, while the net outflow from the Indopacific basin is virtually zero. As wind
stress is increased, Southern Ocean upwelling rises faster than does the outflow from the Atlantic. The
difference is supplied from the Pacific and Indian oceans. For reduced wind stress conditions, on the
other hand, there is net inflow into the Indopacific basin, which balances the remaining outflow from the
20
5
3
CONTROL MECHANISMS OF THE MOC
1.5
a
dyn/cm
2
2
1.0
1
0.5
0.2
0
−1
0
−1000
m
−2000
−3000
−4000
−5000
b
−70
−60
−50
Latitude (°N)
−40
−30
Figure 7: (a) Mean annual zonal mean wind stress in Southern Ocean latitudes, for
the experiments with winds varying from 0.2 to 1.5 times the observed field. (b)
Global meridional overturning streamfunction in corresponding latitudes, for wind
stress forcing according to observations. Spacing between contours is 2 Sv, dashed
contours mean anticlockwise circulation. The latitudinal extent and depth of the
Drake Passage gap are indicated (blue). At the very surface, the gap is somewhat
wider (compare fig. 2).
Atlantic. This Indopacific inflow is divided rather evenly among the Pacific and the Indian basin (figure
9).
In summary, in CLIMBER-3α the present-day AMOC is entirely maintained by wind-driven upwelling.
Given the small vertical diffusivity applied in the model, low-latitude upwelling in the Atlantic is very
small (Mignot et al., 2006, compare also the upper panel of figure 4). Note that the maximum of the
overturning is somewhat larger than the Southern Ocean outflow, i.e. there is some recirculation of
NADW within the North Atlantic. This recirculation occurs mainly along the Western Boundary of the
basin and is likely to be, at least partly, caused by spurious upwelling due to the poorly resolved Western
Boundary layer (Griesel and Morales-Maqueda, 2006; Yang, 2003). In any case, it does not contribute
to the interhemispheric transport of volume and heat. Griesel and Morales-Maqueda (2006) emphasized
the importance of clearly distinguishing between the maximum of the Atlantic MOC and the Southern
Ocean outflow, the latter being the appropriate measure of the interhemispheric circulation.
The magnitude of the South Atlantic outflow, however, is not solely controlled by SO wind stress.
21
5.4 Volume transport and meridional density gradients
25
SO upwelling including AABW
SO upwelling
Atlantic outflow
Indopacific outflow
20
15
Sv
10
5
0
−5
−10
0
5
10
15
20
SO wind stress gradient (10−8 N/m3)
25
30
Figure 8: The upwelling south of 30◦ S is a linear function of the Southern Ocean wind
stress gradient. The total upwelling (diamonds) is proportional to the SO wind stress
gradient; it extrapolates approximately to the origin (dot-and-dash line). Because
part of this water sinks within the SO to form Antarctic Bottom Water (AABW),
the net upwelling across the SO (triangles) is offset from the total upwelling by the
sinking rate of AABW. In the standard configuration (indicated by the vertical dashed
line), all of the net SO upwelling is drawn out of the Atlantic (circles). For high wind
stress, the Indopacific basin (squares) supplies additional outflow into the SO. For
low winds, on the other hand, there is a net upwelling in the Indopacific, which is
balanced by comparatively high outflow from the Atlantic. Solid and dashed lines are
linear interpolations. The vertical dashed line indicates the standard configuration.
Reducing wind stress induces a shift of the circulation regime towards an AMOC sustained by Indopacific
low-latitude upwelling, instead of Southern Ocean upwelling. As will be discussed in the following section,
buoyancy distribution determines the overturning rates in the different basins.
5.4
Volume transport and meridional density gradients
In search for the mechanism responsible for setting the AMOC rate, we compute average meridional
density gradients for the Atlantic, Pacific, and Indian ocean, as the density difference at 1100m depth
between a selected region in the north of each basin and a corresponding Southern Ocean region (figure
10). All regions are located along the eastern flanks of the continents, comprising the Western Boundary
currents responsible for the main meridional volume transport (compare figure 11). At the same time,
22
5
CONTROL MECHANISMS OF THE MOC
20
Global
Pacific
Indian
15
Sv
10
5
0
−5
−10
0
5
10
15
20
SO wind stress gradient (10−8 N/m3)
25
30
Figure 9: Southern ocean outflow from the Pacific (filled squares) and the Indian
basin (empty squares). Net Southern Ocean upwelling is plotted again as in fig. 8 for
comparison (triangles). The vertical dashed line indicates the standard configuration.
the Southern Ocean regions are chosen to capture the density difference induced by the change of wind
stress north of the outcropping region. As Southern Ocean wind stress is changed, the Southern Ocean
experiences the largest density variations (figure 12, solid lines) because the wind-driven upwelling of
deep, dense water masses and the associated outcropping of isopycnals are directly affected (figure 13):
Comparing α = 0.2 and α = 1.5, corresponding isopycnals have a much larger slope for high wind stress
than for low wind stress and, north of about 50◦ S, reach deeper. Therefore, Southern Ocean waters in
the region relevant to the meridional density gradients (green lines in figure 13) become lighter as winds
are enhanced. Density is also altered in the Northern Hemisphere (figure 12, dashed lines). However, the
strongest response is seen in the SO. The steady state ocean circulation obviously adapts to altered SO
wind stress with a global redistribution of buoyancy.
Additional simulations with enhanced freshwater flux into the North Atlantic were conducted. This is
a standard experimental setup to investigate the robustness and sensitivity of AMOC dynamics (Manabe
and Stouffer , 1995; Rahmstorf , 1996; Rahmstorf et al., 2005). The anomalous freshwater flux reduces
deep water formation and, by this means, weakens the Atlantic MOC (figure 14, circles). At the same
time, sinking increases in the Indopacific, where a compensating negative freshwater flux is applied
5.4 Volume transport and meridional density gradients
Figure 10: Ocean regions used to compute meridional density gradients. Densities
are taken at 1100m depth, i.e. within the lower branches of the MOC. The regions
are chosen to match the Western Boundary currents (WBCs) responsible for the
main meridional volume transport. Horizontal velocities of the α = 1.5 experiment,
averaged between 1000m and 2000m depth, are overlaid as vectors (velocities south
of 30◦ S are omitted for the sake of clarity).
Figure 11: Zonal section of meridional velocity at 10◦ S in the α = 1.0 experiment,
in cm/s. Principal deep currents flow along the western boundaries of the African,
Australian, and American continents.
23
24
5
CONTROL MECHANISMS OF THE MOC
0.5
Atlantic
Atlantic FWF
Pacific
Pacific FWF
−0.5
0.5
0
Normalised density σ − σ (kg/m3)
0
0
−0.5
0.5
0
Indian
−0.5
0
Indian FWF
10
20
30
10
−8
SO wind stress gradient (10
20
30
3
N/m )
Figure 12: Densities σsouth (solid lines) and σnorth (dashed lines), normalised to σsouth
at standard wind conditions (σ0 ), for the Atlantic (top), Pacific (center) and Indian
ocean (bottom). Panels on the right show the FWF experiments. Densities in the
south are more strongly affected by changing SO winds. Vertical dashed lines denote
the standard wind stress configuration.
(squares). Note that the combined outflow from both basins remains constant to balance Southern
Ocean upwelling (triangles). The respective functions for the FWF experiments can thus be obtained
by a parallel translation from the experiments without FWF. FWF experiments are denoted by empty
symbols in figures 15 and 12.
Southern Ocean outflow of all experiments depends linearly on the respective density gradient ∆σ
(figure 15). The slope of the linear regression is
s=
Sv
∆MSO
.
≈ 44
∆(∆σ)
kg/m3
(7)
It can be understood as follows: Within the Western boundary layer, the zonal component of the steady
state geostrophic flow equation is
f u = −ρ−1
0 ∂y p + ∇(ν ∇v) ≃ 0
(8)
5.4 Volume transport and meridional density gradients
Figure 13: Average density difference between α = 0.2 and α = 1.5 (colour shading,
1.5 minus 0.2, in kg/m3 ) in the southern Atlantic (top), Pacific (center), and Indian
ocean (bottom). The green lines indicate the latitudes and depth of the three Southern
Ocean regions depicted in fig. 10. Five selected, corresponding isopycnals are overlaid
each for α = 1.5 (black) and α = 0.2 (red) to illustrate the change in outcropping and
stratification due to changed wind stress.
25
26
5
CONTROL MECHANISMS OF THE MOC
20
SO upwelling (no FWF)
SO upwelling FWF
Atlantic outflow FWF
Indopacific outflow FWF
15
Sv
10
5
0
−5
−10
0
5
10
15
20
SO wind stress gradient (10−8 N/m3)
25
30
Figure 14: As figure 8, but with anomalous freshwater fluxes (FWF) applied in the
North Atlantic, and compensated for in the North Pacific. SO upwelling (filled triangles) does not change significally compared to the runs without FWF (empty triangles), but the distribution of the upwelled water among the ocean basins changes:
The overturning circulation in the Atlantic is reduced (circles), while the outflow from
the Indopacific increases (squares). However, slopes with respect to Southern Ocean
wind stress are approximately the same as in figure 8. Lines are linear fits to the
data. The vertical dashed line indicates the standard configuration.
i.e., a meridional pressure gradient produces a meridional flow along the boundary, with the Coriolis force
balanced by friction. In contrast to (4), the viscosity appears here as a tensor ν. Small-scale turbulent
flows (eddies), which significantly contribute to the dissipation of momentum, cannot be resolved by global
climate models. They are commonly parameterized by making the viscosity artificially large, in which
case it is also called the eddy viscosity (Gill , 1982). Because flows across isopycnals require much more
energy than flows within layers of equal density, turbulence is much higher along isopycnals than across
them. Since isopycnals are approximately horizontal in most of the ocean, the horizontal components
νhor of the eddy viscosity tensor ν are set to a much larger value (by eight orders of magnitude) than
the vertical component (Montoya et al., 2005). On the other hand, the divergence of the velocity in
y-direction can be neglected. Therefore, the frictional term in (8) is dominated by νhor ∇2 v. Assuming a
quadratic velocity profile v(x) = −4 vmax
x(x−Lx )
L2x
for the lower branch of the Western Boundary current
(figure 16), where Lx is the width of the boundary layer and vmax is the maximum southward velocity
27
5.4 Volume transport and meridional density gradients
15
12
Southern outflow (Sv)
9
Atlantic
Pacific
Indian
Atlantic FWF
Pacific FWF
Indian FWF
6
3
0
−3
−6
0
0.15
σ
north
3
−σ
south
0.3
0.45
(kg/m )
Figure 15: Throughout all experiments, southern outflow from the Atlantic, Indian,
and Pacific oceans scales linearly with the respective meridional density gradient at
1100m depth. The slope of this relation is universal for all basins.
within the layer,
∇2 v ≃ ∂x2 v(x) = −8
vmax
L2x
(9)
Integration across the boundary layer gives the associated transport
MW BC ≃ δ
Z
Lx
dx v(x) =
0
2
vmax Lx δ
3
(10)
where δ is the thickness of the boundary layer. We substitute ∂y p in (8) by
∂y p = g ∂y
Z
η
D+δ
dz ρ(z) = g
Z
D+δ
dz ∂y ρ(z) = g δ ∂y ρ(ξ)
(11)
D
where we have used the mean value theorem in the third step, and assumed that δ does not change
much along the Western Boundary current. In the second step, we have used that all horizontal pressure
gradients vanish at the level of no motion. D is the position of the level of no motion, which is also the
upper limit of the current. η is the sea surface elevation. We will later identify ξ ∈ (D, D + δ) with the
depth at which the meridional density gradients are diagnosed. However, the large grid spacing does not
28
5
CONTROL MECHANISMS OF THE MOC
allow to capture this depth too exactly in the model.
v
v max
x
0
Lx
Figure 16: Sketch of an idealized zonal profile of the southward velocity within a
Western Boundary current.
Combining the above equations, and assuming a mean meridional density gradient h∂y ρiy = ∆ρ/Ly ,
we obtain (with the minus sign indicating southward transport)
MW BC = −
g
δ 2 L3x ∆ρ
12νhor Ly ρ0
(12)
The largest uncertainties, besides the idealized velocity profile, are included in the thickness of the
boundary layer, δ, and its lateral extent Lx . Both are of the order of the model resolution and thus
difficult to diagnose. Nevertheless, the observed slope s can be obtained, for example, with a choice of
δ = 2.5 km and Lx = 300 km, which is in good agreement with the ranges implied by our model, and
also seems reasonable with respect to the real world.
5.5
Discussion
We have presented results from experiments with a coupled climate model designed to investigate the role
of Southern Ocean wind stress and buoyancy distribution with respect to the global oceanic circulation.
The experiments consist of simulations with Southern Ocean wind stress varying from 0.2 to 1.5 times
the observed field, and additional simulations with enhanced freshwater flux into the North Atlantic. For
each experiment, the model was run into an equilibrium state.
Southern Ocean upwelling is proportional to the wind stress gradient. A small part of the upwelled
water recirculates within the Southern Ocean. The rest is drawn out of the ocean basins to the north. The
distribution of the upwelling among the Atlantic, Pacific, and Indian basins is determined by buoyancy
differences. For the present-day wind stress field, all the net upwelling in the SO is drawn out of the
Atlantic, and thus balances the net sinking in the North Atlantic. The meridional flows generated by
5.5 Discussion
29
this dipole of northern sinking and southern upwelling constitute the conveyor belt of the Atlantic MOC.
For higher wind stress in the SO, part of the additional upwelling is drawn out of the Pacific and Indian
basins. For low wind stress, on the other hand, there is a net inflow into the Pacific and Indian oceans
that partly balances the remaining outflow from the Atlantic.
We have shown that the strength - and direction - of the meridional overturning circulation in the
different basins is set by meridional density gradients. The Southern Ocean outflow rates of all basins obey
a universal linear relation with respect to the meridional density gradient along the Western Boundary of
the respective basin. An explanation was offered on the grounds of classical boundary layer theory, and
the slope of the relation could be confirmed to a degree of accuracy imposed by the model resolution. We
conclude that both wind-driven Southern Ocean upwelling and meridional density gradients combine to
set up a coherent system of conditions that determines the magnitude and structure of the global MOC.
The concept presented here is independent of the mechanism of density transformation, because only
gradients along the Western Boundary currents are considered. For example, convection in the North
Atlantic is not crucial by itself for the functioning of the MOC, but only in setting the density of the
deep water mass. With regard to the Indopacific, both directions of circulation, i.e. southern out flow and
inflow, obey the linear relation with respect to meridional density gradients, although the mechanisms
at work are certainly different for sinking than for upwelling. Nonetheless, detailed description and
explanation of these mechanisms are absolutely necessary for a comprehensive understanding of the
circulation, and remain subject to further investigations. Also, we have neglected to a large degree the
dynamics governing the circulation within the Southern Ocean; for instance, the important question of
what determines the magnitude of the recirculation and formation of AABW are beyond the scope of
the present thesis and will be investigated in a future study.
30
6
6
LACK OF BIPOLAR SEE-SAW
Lack of bipolar see-saw
A cessation of the Atlantic meridional overturning circulation (AMOC) significantly reduces northward
oceanic heat transport. In modelling studies where the AMOC cessation is achieved by the application
of an anomalous freshwater flux, this leads to the classic ’bipolar see-saw’ pattern of northern cooling
and southern warming in surface air and ocean temperatures. By contrast, as shown here in our coupled
climate model, both northern and southern cooling are observed for an AMOC reduction in response to
reduced wind stress in the Southern Ocean (SO). For very weak SO wind stress, not only the overturning
circulation collapses, but sea ice export from the SO is strongly reduced. Consequently, sea ice extent
and albedo increase in this region. The resulting cooling overcompensates the warming by the reduced
northward heat transport. The effect depends continuously on changes in wind stress and is reversed for
increased winds. It may have consequences for abrupt climate change, the last deglaciation and climate
sensitivity to increasing atmospheric CO2 concentration.
6.1
Introduction
The northward heat transport by the Atlantic meridional overturning circulation (AMOC) (Ganachaud
and Wunsch, 2000; Talley, 2003) has important consequences for global climate (Zickfeld et al., 2007c;
Vellinga and Wood , 2007). Presently, it favors mild northern European surface air temperature (SAT)
and a northward excursion of the Intertropical Convergence Zone in the Atlantic sector (Vellinga and
Wood , 2002). An AMOC collapse significantly reduces the northward heat transport and yields a net
heat transport to the Southern Hemisphere. Without any other processes involved this induces a timedelayed positive temperature anomaly in the south and therewith the so-called bipolar see-saw pattern
of northern cooling and southern warming (Crowley, 1992).
For the last glacial period abrupt climatic events (Dansgaard-Oeschger events) can be observed in
Greenland ice cores (NGRIP , 2004). These are accompanied by temperature variations recorded in
Antarctic cores with Antarctica generally warming during a cold (stadial) phase in Greenland (EPICAProject , 2006). A possible dynamical explanation is associated with an abrupt reorganization of the
AMOC in response to fluctuations in surface freshwater fluxes to the North Atlantic (Ganopolski and
Rahmstorf , 2001). In this theory, the time-delayed bipolar see-saw effect explains the characteristic
temperature responses in the Northern and Southern Hemispheres. The temporal lead of the Northern
Hemisphere can however not be inferred from paleodata (Steig and Alley, 2002). An alternative view
is that perturbations are induced from the SO and reach the North Atlantic with some time delay [e.g.
Blunier et al. (1998); Weaver et al. (2003)]. In order to reproduce the temporal sequence recorded in
Greenland and Antarctic ice cores that would require a warming in the SO to be associated with a
warming in the North Atlantic. The mechanism presented here yields such a lack of bipolar see-saw
6.2 Experiments
31
through perturbations in SO wind stress.
The importance of SO winds for the deep AMOC is still a matter of intense investigation (Kuhlbrodt
et al., 2007). Toggweiler and Samuels (1995) proposed that surface wind stress in the SO drives the AMOC
and therefore influences its strength, while other studies emphasize the role of the thermohaline component
of the circulation (e.g. Rahmstorf and England (1997)). For the last glacial period, reconstructions of
wind stress are presently not well constrained (Crowley and North, 1991; Wunsch, 2003, and ref. therein).
Toggweiler et al. (2006) argue that mid-latitude westerly winds are generally shifted equatorward during
glacial periods. In combination with associated changes in SO overturning circulation and the carbon
cycle, they suggest a positive feedback as an explanation for the observed Antarctic warming during the
last deglaciation. The effect presented here in the global coupled climate model CLIMBER-3α yields
such a warming for increased wind stress in the SO and could therefore augment their mechanism.
The main novelty of our simulations is based on the fact that the AMOC in CLIMBER-3α is driven
by SO winds with almost no upwelling in the Atlantic and Pacific basin (Mignot et al., 2006) which
is supported by observations but not reproduced by most state-of-the-art climate models. Potential
implications for future climate change become apparent as recent model simulations, performed for the
Fourth Assessment Report of the IPCC (Meehl et al., 2007a), suggest a strengthening of surface wind
stress in response to atmospheric CO2 increase under the A2 SRES4 scenario (Fyfe and Saenko, 2006).
In combination with our results this would lead to a global warming of about 0.15K and hence contribute
to climate sensitivity.
6.2
Experiments
The global coupled climate model CLIMBER-3α is described in section 2. The model version used here
applies a very low vertical diffusivity coefficient (10−5 m2 /s). Thus the mixing induced upwelling both in
the Atlantic and Pacific are small (Mignot et al., 2006) and the AMOC is almost entirely driven by SO
wind stress divergence through the so-called Drake Passage Effect (Toggweiler and Samuels, 1995).
Oceanic wind stress is prescribed to the NCEP-NCAR reanalysis τ 0 (x, y) (Kistler et al., 2001). We
performed a series of experiments, in which the zonal oceanic wind stress component τx is multiplied
with a factor α between 71.25◦S and 30◦ S with α = 0.01, 0.1, 0.2, 0.5, 1.0, 1.5 and 2.0. The simulations
thus span a range of both reduced and enhanced surface forcing. Since the meridional distribution of τx ,
in the aforementioned latitudinal interval, has a quasi-sinusoidal shape (being close-to-zero at 71.25◦ S
and 30◦ S, and maximum at about 51◦ S), this implies a modulation of both the wind stress magnitude,
τx , and the wind stress gradient. In particular, in the latitude band of Drake Passage, between 62◦ S and
4 Special Report on Emissions Scenarios. The SRES provides a standard set of scenarios of future greenhouse gas
emissions for the IPCC climate projections. The A2 scenario familiy refers to a differentiated evolution of global economy
and relatively slow technological change. The SRES is available at www.ipcc.ch/ipccreports/sres/emission/.
32
6
LACK OF BIPOLAR SEE-SAW
Figure 17: Annual SAT difference (in K, on- minus off-state): AMOC cessation
is obtained by (a) application of 0.35 Sv of anomalous freshwater flux (FWF) to
the northern North Atlantic (b) reduction in SO wind stress (α = 0.01). Note the
asymmetric color scale. The thick contours in the Southern Ocean in (b) mark the
level of 70% average winter sea ice cover for normal wind conditions (α = 1, black),
for collapsed AMOC (α = 0.01, red), and for pre-industrial climate reconstructions
(Rayner et al., 2002) (year 1870, green).
6.3 Temperature response to SO wind changes
33
56◦ S, the maximum of the zonally averaged wind stress gradient, Ty ≡ ∂y τx , varies from 0.2 · 10−8 N/m3
in the α = 0.01 experiment, to 20.3 · 10−8N/m3 in the α = 2 experiment. For each experiment, the model
was run for several thousand years until an equilibrium state was reached. For comparison a simulation
with 0.35 Sv = 0.35 · 106m3 /s of enhanced freshwater flux to the North Atlantic (from 52◦ N to 80◦ N and
from 48◦ W to 15◦ E) was carried out, leading to a collapse of the AMOC (experiment FWF, following
Levermann et al. (2005)). This experiment was run for more than 5000 years with anomalous freshwater
flux and prescribed NCEP/NCAR wind stress forcing.
6.3
Temperature response to SO wind changes
In experiments with enhanced freshwater flux into the North Atlantic such as the FWF experiment
(Manabe and Stouffer (1995); Stouffer et al. (2006)), surface air and ocean temperatures are found to
cool in the Northern Hemisphere, and warm in the Southern Hemisphere as the AMOC is reduced (figure
17a). This aforementioned bipolar see-saw pattern of opposite surface temperature changes originates
from a reduction in northward oceanic heat transport associated with the AMOC.
In CLIMBER-3α a similar AMOC reduction is obtained for near-zero SO wind stress (fig. 18a).
However, these experiments do not exhibit the bipolar see-saw pattern. Instead, SAT cools in both
hemispheres when the AMOC is reduced (figure 17b). The effect does not show abrupt changes but
depends continuously on SO wind stress gradient (figure 18c). All experiments with α < 1 exhibit global
cooling relative to the standard simulation, while a slight global warming is observed for increased Ty .
The same behavior is found for sea surface temperature while subsurface temperatures in the SO increase
for weakened Ty due to reduced AMOC and thereby northward oceanic heat transport (not shown).
Note that the negative surface temperature anomaly does not arise from reduced SO upwelling, since
temperatures decrease with depth in this region. Due to reduced wind-driven outcropping of isopycnals
and reduced brine rejection, convection is strongly reduced in the SO for weak wind stress. This could,
in principle, lead to a local cooling as observed in response to anomalous freshwater forcing in the North
Atlantic. However, the global cooling seen in fig. 17b and the corresponding sea surface temperature
pattern can not be caused by a redistribution of heat within the ocean. A net heat loss of the lower
troposphere is necessary and thus surface processes have to be responsible.
6.4
Sea ice mechanism
The absence of a bipolar see-saw in our experiments is due to changes in sea ice cover and thereby
altered albedo and atmosphere-ocean heat exchange in the SO. The southern subpolar westerlies induce
a northeastward Ekman sea ice transport. A reduction in SO surface wind stress consequently reduces
northward sea ice export from the Antarctic region (fig. 19). As a result, sea ice cover increases in the SO.
34
6
LACK OF BIPOLAR SEE-SAW
SO outflow (Sv)
15
10
FWF on
5
FWF off
a
0
0.5
FWF off
0.45
0.3
0.2
Heat flux anomaly (W/m2)
FWF off
b
Surface albedo
0.7
0
0
FWF off
−20
−4
−40
−8
c
0
10
20
−12
30
−8
SO wind stress gradient (10
SAT (deg C)
Sea ice fraction
0.7
3
N/m )
Figure 18: Changes with varying SO winds: (a) SO outflow (at 33.75◦S). The line
gives a linear approximation. For comparison, the AMOC on- and off-states of the
anomalous freshwater flux (FWF) experiments are indicated by open circles. (b) Sea
ice fraction (red dots) and surface albedo (black squares), averaged south of 40◦ S.
Sea ice fraction increases as wind stress is reduced, enhancing surface albedo. The
lines give exponential approximations with saturation at 0.24 and 0.33 for the sea
ice fraction and surface albedo, respectively, as Ty → ∞. (c) Ocean-atmosphere heat
flux anomalies (red dots) and SAT (black squares), averaged south of 40◦ S. For
decreasing wind stress, reduced heat flux into the atmospheric surface layer reduces
surface temperature. The lines gives exponential approximations with saturation at
4W/m2 and −2.6◦ C for heat flux anomaly and for SAT, respectively, as Ty → ∞. In
(b) and (c), for sea ice fraction, albedo, and SAT, the AMOC off-states-values of the
FWF experiment are given for comparison. FWF-on-state-values are equivalent to
the α = 1 experiment. See figure 20 for a decomposition of the ocean-atmosphere heat
flux into its radiative and turbulent components. The x-axis shows the maximum of
the zonally averaged wind stress gradient Ty between 63.75◦ S und 52.5◦S.
35
6.5 Discussion
Northward transport (Sv)
0.4
2.0
1.0
0.3
0.2
0.5
0.1
0
−75
−70
−65
−60
−55
Latitude
−50
−45
Figure 19: Zonally integrated annual mean northward sea ice transport (Sv) for
α = 0.5, 1.0, and 2.0.
The thick contours in figure 17b indicate the 70%-winter-sea-ice-cover. The standard simulation (α = 1,
black) overestimates sea ice extent slightly compared to pre-industrial reconstructions (green) (Rayner
et al., 2002). For smaller Ty , the 70%-sea-ice-margin is shifted northward by up to ∼ 15 degrees for
α = 0.01 (red) compared to standard conditions. The sea ice area south of 40◦ S (fig. 18b, red) increases
from ∼30% with standard wind stress up to 70 % for α = 0.01. For enhanced wind stress sea ice cover
declines and saturates around 24% (obtained from the red exponential curve in fig. 18b).
Increased sea ice cover enhances surface albedo (fig. 18c) and insulates the ocean from the atmosphere.
Both effects tend to cool the ocean surface in different ways. The increased albedo leads to a reduction
of the absorption of incoming solar radiation at the surface. The effect is largest in Southern Hemisphere
summer, when the incoming solar radiation is largest (fig. 20a), while in the Southern Hemisphere winter
short wave radiative heat flux is near zero at the relevant latitudes. The insulation effect tends to reduce
turbulent heat flux throughout the year (fig. 20b) but is strongest during the southern hemispheric
winter when the ocean is much warmer than the atmosphere. Thus, the enhanced sea ice cover results in
a reduction of the total amount of heat received by the atmosphere throughout the year (figure 18c, red
circles) and consequently yields an SAT reduction for low SO wind stress (figure 18c, black squares).
6.5
Discussion
In contrast to the classic bipolar see-saw pattern, we observe a cooling in both hemispheres for an AMOC
weakening in response to a reduction in SO wind stress. This effect is explained by increased sea ice
cover in the SO due to reduced sea ice export into warming northern latitudes. This increases surface
albedo and insulates the atmosphere from the ocean. Thus both radiative and turbulent heat loss over
the SO are enhanced, which overcompensates the net heat gain due to reduced northward oceanic heat
36
6
LACK OF BIPOLAR SEE-SAW
Heat flux anomaly (W/m2)
20
winter
0
−20
summer
RADIATIVE
−40
a
summer
2
Heat flux anomaly (W/m )
−60
10
0
−10
−20
TURBULENT
winter
−30
b
−40
0
10
20
30
−8
SO wind stress gradient (10
3
N/m )
Figure 20: Changes of (a) radiative and (b) turbulent ocean-atmosphere heat flux
in response to changed SO winds. Averages for Southern Hemisphere winter (JuneAugust) are given by open circles and dashed lines, and for Southern Hemisphere
summer (December-February) by dots and dotted lines. Solid lines give annual means.
The effect of sea ice albedo on radiative heat uptake plays a role only in summer when
there is significant insolation at high southern latitudes. The insulating effect of sea
ice on turbulent heat exchange is largest in winter but occurs throughout the year.
transport by the AMOC weakening. The latter effect is nevertheless found in subsurface.
Albedo and insulation effect might be overestimated in CLIMBER-3α due to the larger SO sea ice
extent in the standard simulation compared to pre-industrial reconstructions. On the other hand, changes
in turbulent heat exchange might be underestimated because they are increasing with wind speed but are
not affected by the changes in wind-stress that were prescribed in our experiments. The enhancement of
turbulent heat flux with increasing wind speed would enhance the effect described here. In any case, the
qualitative behaviour is quite straightforward and should be at play whenever wind stress changes in the
SO.
Recently, Fyfe and Saenko (2006) showed that SO winds increase in AOGCM projections of future
climate change by 25 % in 2100 under the A2 scenario. Their results indicated that an increase in SO
winds by 20 % for a doubling of atmospheric CO2 concentration from 280 ppm to 560 ppm is in the realm
6.5 Discussion
37
of possibilities. In our model this would correspond to a global warming of about 0.15 K - a small but
not negligible contribution to climate sensitivity.
As mentioned in the introduction, further implications might arise for abrupt climatic events during
the last glacial period and deglaciation. If these events were indeed triggered from the SO as opposed
to the North Atlantic, a lack of the bipolar see-saw pattern might be necessary to explain the temporal
evolution of events seen in the ice cores. The main result of our study, however, is that modifications in
AMOC strength are not necessarily associated with a bipolar SAT and sea surface temperature response.
AMOC changes due to SO wind stress variations can be associated with equal-sign temperature changes
in both hemispheres.
38
7
7
OVERALL DISCUSSION AND CONCLUSIONS
Overall discussion and conclusions
The aim of this study was to investigate the role of Southern Ocean wind stress for the global meridional
overturning circulation (MOC). The MOC is a large-scale oceanic deep circulation spanning the entire
globe. It has important implications for global climate due to its various interactions with the atmosphere,
cryosphere, and biogeochemical cycles; among others, it serves as a giant interhemispheric heat pump
and a sink for atmospheric carbon, and affects the global distribution of sea surface elevation. It also
interacts directly with the biosphere by setting the living conditions of numerous marine ecosystems,
many of which are themselves of great socio-economic importance. Evidence from the geological past
suggests that abrupt reorganizations of the MOC were involved in high-impact climate transitions. At
the same time, recent studies have raised concern about the future of the MOC in a climate that is rapidly
changing due to anthropogenic greenhouse gas emissions.
In order to fully understand the dynamics of the MOC and to realistically assess its possible future
evolution, knowledge of its principal driving mechanism is indispensable. There are currently two candidates under discussion: Turbulent mixing processes in low latitudes, and wind-driven upwelling in the
Southern Ocean due to the Drake Passage effect. We have found that in our model the Atlantic MOC,
defined as the Southern Ocean outflow from the Atlantic basin, is completely sustained by wind-driven
Southern Ocean upwelling under present-day boundary conditions. Mixing-induced upwelling in low latitudes is negligible due to the low background vertical diffusivity, which is consistent with observations.
For higher wind stress in the SO, part of the additional upwelling is drawn out of the Pacific and Indian
basins. For low wind stress, on the other hand, there is a net inflow into the Pacific and Indian oceans
that partly balances the remaining outflow from the Atlantic.
We have shown in section 5 that the strength - and direction - of the meridional overturning circulation
in the different basins is set by meridional density gradients. The Southern Ocean outflow rates of all
basins obey a universal linear relation with respect to the meridional density gradient along the Western
Boundary of the respective basin. An explanation has been offered on the grounds of classical boundary
layer theory, and the slope of the relation could be confirmed to a degree of accuracy imposed by the
model resolution. We conclude that both wind-driven Southern Ocean upwelling and meridional density
gradients combine to set up a coherent system of conditions that determines the magnitude and structure
of the global MOC. This system of conditions is independent of the mechanism of density transformation
and even holds for reversed direction of circulation in the Indian and Padific basins. Nonetheless, detailed
description and explanation of the possible mechanisms of sinking and upwelling in these basins are
certainly necessary for a comprehensive understanding of the circulation, and remain subject to further
investigations. Also, detailed investigation of the circulation dynamics of the Southern Ocean, including
the link between wind-driven upwelling and the formation of AABW, is beyond the scope of this study.
39
The results presented in section 5 are in preparation for publication (Schewe and Levermann, 2008).
Concerning the implications of a possible slow-down of the AMOC for global surface temperatures,
a novel effect was presented in section 6. In contrast to the classic bipolar see-saw pattern of northern
cooling and southern warming, we observe a cooling in both hemispheres for an AMOC weakening in
response to a reduction in SO wind stress. This effect is explained by increased sea ice cover in the SO due
to reduced sea ice export into warming northern latitudes. This increases surface albedo and insulates
the atmosphere from the ocean. Thus both radiative and turbulent heat loss over the SO are enhanced,
which overcompensates the net heat gain due to reduced northward oceanic heat transport by the AMOC
weakening. Projections of future climate change show an increase in SO winds within the next century
and thus suggest that the effect contributes to climate sensitivity. Implications might also arise for abrupt
climatic events during the last glacial period and deglaciation. We conclude that modifications in AMOC
strength are not necessarily associated with a bipolar SAT and sea surface temperature response, but can
be associated with equal-sign temperature changes in both hemispheres. The results presented in section
6 have been published in Geophysical Research Letters (Levermann et al., 2007).
40
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47
LIST OF FIGURES
List of Figures
1
Strongly simplified sketch of the global meridional overturning circulation (MOC), from
Kuhlbrodt et al. (2007), after Rahmstorf (2002). In the Atlantic, warm, saline surface
waters from the tropics are exported northwards via the Gulf stream and the North Atlantic
current (red). Due to surface cooling, they sink in the Nordic Seas and the Labrador Sea
to form North Atlantic Deep Water (NADW), which returns southward at depth (light
blue). In the Southern Ocean, the Antarctic Circumpolar current (ACC) spans the entire
globe and connects all other basins. Sinking regions also exist south of the ACC; the deep
water masses formed here are denser than the NADW and thus circulate at greater depths
(dark blue). The upwelling which must balance the sinking at these localised sites could
in principle be provided by two processes, low-latitude vertical mixing and wind-driven
Southern Ocean upwelling, both of which are indicated here; their relative importance is
a main subject of this study.
2
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3
Coastlines and bottom topography of the oceanic model grid of CLIMBER-3α. Note that
the 24 depth levels are variably spaced, see table 1. For a detailed model description see
Montoya et al. (2005). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3
5
Idealized sketch representing a zonally averaged picture of the Atlantic and the Southern Ocean, from Kuhlbrodt et al. (2007). The color shading depicts a zonally averaged
density profile derived from observational data (Levitus, 1982), where blue is light and
yellow is dense. Straight arrows indicate the main volume transport associated with the
AMOC. Two principal water masses circulate in the deep ocean: North Atlantic Deep
Water (NADW), which forms in the Nordic Seas due to atmospheric cooling of salty surface waters and, after flowing over a shallow sill between Greenland and Scotland, sinks
to greater depths in the Atlantic basin; and Antarctic Bottom Water (AABW), which
forms in the Weddell and Ross Seas in the Southern Ocean. These deep water masses are
separated from the lighter and warmer surface waters by a region of large temperature
gradient, the thermocline. Breaking of internal waves, generated at the surface as well as
over bottom topography, leads to upwelling through the thermocline via turbulent mixing
along the density gradient (diapycnal mixing). In the Southern Ocean, deep waters are
upwelled by surface wind stress divergence due to the Drake Passage effect (section 4).
Part of these waters recirculate within the so-called Deacon cell. Note that this sketch
is only meant to illustrate some terms and concepts mentioned in the text; the processes
depicted are highly simplified. Note also that the typical ratio of meridional to vertical
extent of the real ocean is of the order of 5000. . . . . . . . . . . . . . . . . . . . . . . . .
8
48
LIST OF FIGURES
4
Meridional overturning streamfunction of the Atlantic (top) and Indopacific (bottom)
basins, in Sv, for the standard configuration of CLIMBER-3α. Positive (negative) values and red (blue) shading indicate clockwise (anticlockwise) circulation. Spacing between
contours is 1 Sv. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5
9
Simplified sketch of a south-north section through an ocean basin, illustrating the occurance
of (a) shallow Ekman cells and (b) deep upwelling due to divergent northward Ekman
transport at the surface. The opening labelled ’DP’ represents the zonal gap of Drake
Passage.
6
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
15
(a) Global upwelling budget through the 780m depth level for the standard configuration
(solid line) and for 0.5 (dot and dash) and 1.5 (dashed) times the observed wind stress field.
The main upwelling region is located in the Southern Ocean around 60◦ S. The location
of Drake Passage is indicated by blue lines. (b) Global meridional streamfunction for the
standard configuration. The DP opening (blue) and the 780m depth level (dashed blue)
are indicated.
7
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
16
(a) Mean annual zonal mean wind stress in Southern Ocean latitudes, for the experiments
with winds varying from 0.2 to 1.5 times the observed field. (b) Global meridional overturning streamfunction in corresponding latitudes, for wind stress forcing according to
observations. Spacing between contours is 2 Sv, dashed contours mean anticlockwise circulation. The latitudinal extent and depth of the Drake Passage gap are indicated (blue).
At the very surface, the gap is somewhat wider (compare fig. 2). . . . . . . . . . . . . . .
8
20
The upwelling south of 30◦ S is a linear function of the Southern Ocean wind stress gradient. The total upwelling (diamonds) is proportional to the SO wind stress gradient; it
extrapolates approximately to the origin (dot-and-dash line). Because part of this water
sinks within the SO to form Antarctic Bottom Water (AABW), the net upwelling across
the SO (triangles) is offset from the total upwelling by the sinking rate of AABW. In the
standard configuration (indicated by the vertical dashed line), all of the net SO upwelling
is drawn out of the Atlantic (circles). For high wind stress, the Indopacific basin (squares)
supplies additional outflow into the SO. For low winds, on the other hand, there is a net
upwelling in the Indopacific, which is balanced by comparatively high outflow from the Atlantic. Solid and dashed lines are linear interpolations. The vertical dashed line indicates
the standard configuration. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
9
21
Southern ocean outflow from the Pacific (filled squares) and the Indian basin (empty
squares). Net Southern Ocean upwelling is plotted again as in fig. 8 for comparison
(triangles). The vertical dashed line indicates the standard configuration. . . . . . . . . .
22
49
LIST OF FIGURES
10
Ocean regions used to compute meridional density gradients. Densities are taken at 1100m
depth, i.e. within the lower branches of the MOC. The regions are chosen to match the
Western Boundary currents (WBCs) responsible for the main meridional volume transport.
Horizontal velocities of the α = 1.5 experiment, averaged between 1000m and 2000m depth,
are overlaid as vectors (velocities south of 30◦ S are omitted for the sake of clarity).
11
. . .
23
Zonal section of meridional velocity at 10◦ S in the α = 1.0 experiment, in cm/s. Principal
deep currents flow along the western boundaries of the African, Australian, and American
continents.
12
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
23
Densities σsouth (solid lines) and σnorth (dashed lines), normalised to σsouth at standard
wind conditions (σ0 ), for the Atlantic (top), Pacific (center) and Indian ocean (bottom).
Panels on the right show the FWF experiments. Densities in the south are more strongly
affected by changing SO winds. Vertical dashed lines denote the standard wind stress
configuration.
13
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
24
Average density difference between α = 0.2 and α = 1.5 (colour shading, 1.5 minus 0.2, in
kg/m3 ) in the southern Atlantic (top), Pacific (center), and Indian ocean (bottom). The
green lines indicate the latitudes and depth of the three Southern Ocean regions depicted
in fig. 10. Five selected, corresponding isopycnals are overlaid each for α = 1.5 (black)
and α = 0.2 (red) to illustrate the change in outcropping and stratification due to changed
wind stress.
14
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
25
As figure 8, but with anomalous freshwater fluxes (FWF) applied in the North Atlantic,
and compensated for in the North Pacific. SO upwelling (filled triangles) does not change
significally compared to the runs without FWF (empty triangles), but the distribution of
the upwelled water among the ocean basins changes: The overturning circulation in the
Atlantic is reduced (circles), while the outflow from the Indopacific increases (squares).
However, slopes with respect to Southern Ocean wind stress are approximately the same
as in figure 8. Lines are linear fits to the data. The vertical dashed line indicates the
standard configuration.
15
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
26
Throughout all experiments, southern outflow from the Atlantic, Indian, and Pacific oceans
scales linearly with the respective meridional density gradient at 1100m depth. The slope
of this relation is universal for all basins.
16
. . . . . . . . . . . . . . . . . . . . . . . . . . .
27
Sketch of an idealized zonal profile of the southward velocity within a Western Boundary
current. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
28
50
LIST OF FIGURES
17
Annual SAT difference (in K, on- minus off-state): AMOC cessation is obtained by (a)
application of 0.35 Sv of anomalous freshwater flux (FWF) to the northern North Atlantic
(b) reduction in SO wind stress (α = 0.01). Note the asymmetric color scale. The thick
contours in the Southern Ocean in (b) mark the level of 70% average winter sea ice cover
for normal wind conditions (α = 1, black), for collapsed AMOC (α = 0.01, red), and for
pre-industrial climate reconstructions (Rayner et al., 2002) (year 1870, green).
18
. . . . . .
32
◦
Changes with varying SO winds: (a) SO outflow (at 33.75 S). The line gives a linear
approximation. For comparison, the AMOC on- and off-states of the anomalous freshwater
flux (FWF) experiments are indicated by open circles. (b) Sea ice fraction (red dots) and
surface albedo (black squares), averaged south of 40◦ S. Sea ice fraction increases as wind
stress is reduced, enhancing surface albedo. The lines give exponential approximations
with saturation at 0.24 and 0.33 for the sea ice fraction and surface albedo, respectively, as
Ty → ∞. (c) Ocean-atmosphere heat flux anomalies (red dots) and SAT (black squares),
averaged south of 40◦ S. For decreasing wind stress, reduced heat flux into the atmospheric
surface layer reduces surface temperature. The lines gives exponential approximations
with saturation at 4W/m2 and −2.6◦ C for heat flux anomaly and for SAT, respectively, as
Ty → ∞. In (b) and (c), for sea ice fraction, albedo, and SAT, the AMOC off-states-values
of the FWF experiment are given for comparison. FWF-on-state-values are equivalent to
the α = 1 experiment. See figure 20 for a decomposition of the ocean-atmosphere heat
flux into its radiative and turbulent components. The x-axis shows the maximum of the
zonally averaged wind stress gradient Ty between 63.75◦S und 52.5◦S. . . . . . . . . . . .
34
19
Zonally integrated annual mean northward sea ice transport (Sv) for α = 0.5, 1.0, and 2.0.
35
20
Changes of (a) radiative and (b) turbulent ocean-atmosphere heat flux in response to
changed SO winds. Averages for Southern Hemisphere winter (June-August) are given by
open circles and dashed lines, and for Southern Hemisphere summer (December-February)
by dots and dotted lines. Solid lines give annual means. The effect of sea ice albedo on
radiative heat uptake plays a role only in summer when there is significant insolation at
high southern latitudes. The insulating effect of sea ice on turbulent heat exchange is
largest in winter but occurs throughout the year.
. . . . . . . . . . . . . . . . . . . . . .
36
Danksagung
Mein Dank gilt zuallererst Prof. Dr. Anders Levermann für die hervorragende Betreuung, die für
den Erfolg meiner Arbeit und nicht zuletzt die Freude an derselben von unschätzbarem Wert war und
mich stets sehr motiviert hat. Auch über wissenschaftliche Belange hinaus habe ich von ihm vielfältige
Unterstützung erfahren.
Für die Einarbeitung und zuvorkommende Hilfestellung in technischen wie inhaltlichen Fragen bedanke ich mich bei Heiko Gölzer, Maik Stöckmann und Andreas Born; für Diskussionen und Kommentare
zum Manuskript bei Silvia Volkwardt, Daria Schönemann, Johannes Fürst und Tore Hattermann; und
bei allen zusammen für die heitere und motivierende Arbeitsatmosphäre. Auch danke ich allen weiteren
Kolleginnen und Kollegen am PIK, von denen ich bei zahlreichen Besprechungen und Vorträgen viel lernen konnte. Dr. Marisa Montoya und Dr. Juliette Mignot danke ich für die gute Zusammenarbeit über
große Distanzen hinweg; Prof. Dr. Nacho Pascual für seine Beratung und die Übernahme der Betreuung
am Fachbereich; Silke Graser für die abschließende Fehlerkorrektur.
Besonderer Dank schließlich gebührt meiner Familie, die mich während des gesamten Studiums
unterstützt und mir diese Arbeit erst ermöglicht hat.
Diese Arbeit wurde durchgeführt am
Potsdam-Institut für Klimafolgenforschung
Telegrafenberg A62, 14473 Potsdam

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