Download 1 Oceanic Processes: Introduction

1 Oceanic Processes: Introduction
There is no good textbook that covers all of solid Earth geology, meteorology, climatology,
oceanography, glaciology, and life sciences. In fact, all the books on the reading list for this
course are rather short on oceanography.
In addition, this lecture is just a brief two-hour introduction to the oceans. Mostly it is
an introduction to the dynamics of the oceans. So it doesn’t make sense to have a separate
textbook for this part alone. Instead, I provide these notes to go along with the course.
Much of the information (and many of the figures) are taken from a free online textbook
available at http://oceanworld.tamu.edu/ocean410/ocng410 text book.html
This is a nice textbook and well worth reading. Especially if something wasn’t clear in
class, you may find the answers there.
Other good books for physical oceanography are:
G.L. Pickard and W.J. Emery, Descriptive Physical Oceanography: An Introduction, Pergamon Press.
S. Pond and G.L. Pickard, Introductory Dynamical Oceanography, Pergamon Press.
And for geological oceanography: J. Kennett, Marine Geology, Prentice-Hall.
Suggested Reading from the Online Textbook
Topic
Brief Overview and Historical Setting
Depth and Size of Oceans
Atmospheric Influences
The Heat Budget
Temperature, Density, and Salinity
Equations of Motion
Effects of Viscosity
Response of Upper Ocean to Wind
Geostrophic Currents
Wind Driven Currents
Deep Circulation
Chapter
All of chapters 1 & 2
Chapter 3, except 3.6
4.1, 4.6, 4.7
5.1, 5.2, 5.6, 5.7, 5.9
6.3, 6.4, 6.5, 6.11
All of chapter 7; much of this will be covered in the lectures about energy and dynamics
8.1, 8.2
9.1, 9.2 (skip equations)
10.1, 10.2, (10.3) (skip equations)
All of chapter 11
All of chapter 13
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2 Basics about the Oceans
F IGURE 1: The oceans and seas cover 70.8% of the surface of the Earth.
How deep are the oceans, and how high are the continents?
• This question is not as silly as it might seem. Looking carefully at Fig. 1, one finds that
very little of the water-covered area is shallower than 2000 m.
• Most of the oceans are between 3000 and 5000 m deep.
• Most of the continental areas are between 0 m and 1000 m.
• The very shallow water-covered regions (typical depth of 100 m or less) are not oceanic
at all. They are low-lying, flooded parts of continents such as the Baltic Sea or the North
Sea.
There are a couple of nice online tools for plotting bathymetry (and topography) if you are
interested:
http://maps.ngdc.noaa.gov/viewers/bathymetry/
http://cer.ori.u-tokyo.ac.jp/toolmap/3d-bottom.html
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F IGURE 2: East-west section across the Atlantic ocean at 25◦ S. Mid-Atlantic Ridge near 14◦ W,
Walvis Ridge near 6◦ E, a seamount near 35◦ W. Vertical exaggeration 180:1 in upper figure; 30:1
in lower figure. If shown in true aspect ratio, the ocean depth would be as thin as a line on the plot.
F IGURE 3: Terminology of oceanic features.
Notice that the oceans are very shallow; the above plots are strongly vertically exaggerated. Most oceans are about 1000 times wider than they are deep. A direct consequence is
that vertical velocities in the oceans are much lower than horizontal velocities. Horizontal
velocities are typically less than 1% of the vertical velocities.
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F IGURE 4: Hypsometric curve, illustrating global topography/bathymetry in form of a histogram.
Notice two peaks: (1) Just over 0 m are the continents; (2) around 4500-5000 m are the oceans.
The bimodal distribution of global topography and bathymetry is a consequence of geology:
• Oceanic crust is composed of basalt, a volcanic rock formed by partial melting of the
mantle. Basalt is relatively dense (ρ ≈ 2800 kg/m3 ).
• The oceanic crust is also fairly thin, with very uniform thickness of 5-9 km (typically
about 7 km).
• Continental crust is “granitic” or “felsic” (this doesn’t mean it’s all granite, but especially at shallow levels a lot of it is). It is lighter than oceanic crust (ρ ≈ 2300 kg/m3).
• The continental crust is thick, with a fairly uniform thickness of about 30-35 km.
• Putting all the above together: The continental plates float high in the mantle because
they are light and thick. The oceanic plates do not float as high because they are heavier
and thinner.
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b)
a)
F IGURE 5: a) A seamount is heavier than the water it displaces; there is excess mass at the base
of the ocean relative to a flat bottom. This extra mass gives rise to a perturbation in gravity, which
in turn causes a small “mountain” of water at the sea surface (strongly exaggerated in the figure).
b) A satellite measures the distance from the orbit to the sea surface and detects the mountain of
water.
How do we know the ocean bathymetry?
• We have some data from direct depth soundings from ships, but these don’t cover the
ocean basins very well. Large areas are unsampled.
• Satellites sample all of the oceans’ surface, but cannot see the ocean floor.
• Extra mass on the seafloor causes uplift of the sea surface. This can be measured by
satellites.
• The equilibrium sea surface elevation (if there were no ocean currents) is an equipotential surface known as the geoid. It is very close to an ellipsoid, but with lots of small
perturbations due to mass anomalies at the solid Earth surface and below.
• From the elevation of the sea surface, we can calculate (more or less accurately) the
ocean floor bathymetry.
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F IGURE 6: Submarine canyon (offshore California). Formed by underwater turbidity currents,
which are made up of a mix of sea water and mud, sand, and gravel. Turbidity currents are essentially underwater landslides and constitute a significant portion of near shore sediment transport.
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3 Properties of Sea Water
F IGURE 7: Density of sea water depends on temperature, salinity, and pressure (effect of pressure
is not shown here). “Salinity” is the total amount of solids dissolved in the water. For most oceans,
salinity is about 3.5% or 35h. “Density Index” in the figure is defined as density difference relative
to 1000 kg/m3 , multiplied times 1000; or more simply, a density index of 27 means a density of
1027 kg/m3 .
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F IGURE 8: Another way of graphing density.
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F IGURE 9: Temperature and salinity of the upper ocean in different parts of the world: AAC
= Antarctic Circumpolar Current; “Warm Pool” = Tropical West Pacific; BATS = near Bermuda.
The upper 100-200 m are typically warm (except in arctic regions). Temperatures decrease rapidly
between 200 and 1000 m, this is called the thermocline. The entire deep ocean below 1000-1500 m
is cold, about 0◦ C. Note: 1 decibar corresponds to approximately 1 m in depth. The seasonal effect
is generally low.
The wind-driven circulation reaches a depth of about 200 m. This upper layer is wellmixed due to turbulence in the rapidly moving water. The upper layer is also relatively warm
(of course more so in the tropics). The water below the mixed layer is relatively cold, typically
-4 to 0◦ C. At the base of the mixed layer, the temperature rapidly decreases from the temperature of the upper layer to that of the lower layer: this steep temperature gradient is called the
thermocline.
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F IGURE 10: Growth and decay of the mixed layer and seasonal thermocline (near Bermuda).
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4 The Coriolis Effect
• All physics is properly done in an inertial (non-accelerating) reference frame.
• A natural and customary reference frame for people on Earth is of course the Earth’s
surface. Your reference frame is a fixed position on Earth.
• However, the Earth rotates. All points at the Earth’s surface are constantly accelerated
towards the Earth’s rotation axis.
• The Earth’s surface is NOT an inertial reference frame.
• Two possible ways out:
[1] Use an inertial reference frame (e.g., fixed relative to the solar system). This is
very inconvenient because all objects on Earth move around and accelerate in this
frame.
[2] Use the Earth surface as reference frame, but correct for the accelerations. This
introduces a “fictional” force: the Coriolis force or Coriolis effect.
• The Coriolis force acts on objects moving on the Earth’s surface. It is proportional to
the object’s velocity.
• The Coriolis force acts perpendicular to the velocity: to the right in the northern hemisphere, and to the left in the southern hemisphere. It is zero at the equator.
• More details on the Coriolis force in the lectures about dynamics.
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North Pole
equator
South Pole
F IGURE 11: Particle paths for frictionless motion on a rotating Earth (ellipsoid). The paths
are relative to an observer stationary on the Earth’s surface. In the northern hemisphere, moving
particles are always deflected towards the right, leading to clockwise motion. The effect is the
opposite on the southern hemisphere.
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5 Geostrophic Motion
Taken from Wikipedia (geostrophic wind) for atmospheric motion:
Air naturally moves from areas of high pressure to areas of low pressure, due to
the pressure gradient force. As soon as the air starts to move, however, the Coriolis
force deflects it due to the rotation of the earth. The deflection is to the right in the
northern hemisphere, and to the left in the southern hemisphere. As the air moves
from the high pressure area, its speed increases, and so does the deflection from
the Coriolis force. The deflection increases until the Coriolis and pressure gradient
forces are in geostrophic balance, at which point the air is no longer moving from
high to low pressure, but instead moves along an isobar, a line of equal pressure
(note that this explanation assumes that the atmosphere starts in a geostrophically
unbalanced state and describes how such a state would evolve into a balanced flow.
In practice, the flow is nearly always balanced. Similarly, when evaluating an eastwest gradient in the northern hemisphere, eastward movements turn south, and
westward movements turn north by virtue of the Coriolis force. The geostrophic
approximation has no predictive value since it does not contain any expression for
change: it is purely diagnostic). The geostrophic balance helps to explain why low
pressure systems spin counterclockwise and high pressure systems spin clockwise
in the northern hemisphere (and the opposite in the southern hemisphere).
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F IGURE 12: Snap-shot of large-scale high- and low-pressure systems in the Earth’s atmosphere.
The air tries to flow from high to low pressure, but is deflected by the Coriolis force. The resulting
flow largely follows contours of constant pressure. The force balance in these rotating systems is
mostly geostrophic: the pressure gradient force is balanced by the Coriolis force.
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6 General Observations of Ocean Circulation
Conceptually, the flow of the world’s ocean currents can be divided into two parts: a shallow,
wind-driven circulation, and a deeper, convectively driven thermohaline circulation. Dynamically, there is quite a bit more to it and we will need to understand the individual terms of the
Navier-Stokes equation and the dominating force balances in the various parts of the ocean.
F IGURE 13: The shallow, wind-driven circulation of the Earth’s oceans. The dominant patterns
are the “gyre” of the major ocean basins with water spinning around in a circle. The gyre are
(mostly) clockwise in the northern hemisphere and counter-clockwise in the southern hemisphere.
Additional currents are various E-W trending currents along the equator, and the Antarctic Circumpolar Current. The strongest currents are usually found at the western edge of each ocean,
most notably the Gulf Stream (W of N Atlantic), the Agulhas Current (W of Indian Ocean), and
the Kuroshio Current (W of N Pacific). The Humbold Current (E of S Pacific) and the California
Current (E of N Pacific) are also fairly strong, but the arrow length and the arrow density do not
correlate with current strength.
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F IGURE 14: Schematic of the oceans’ deep current pattern. Observations of deep currents are
difficult and data is scarce. Note, however, the two points labeled S1 and S2 . These are the
only two places on Earth where deep water is “formed”, i.e., where surface waters become dense
enough to sink into the deep oceans. S1 is in the Norwegian Sea, S2 is in the Weddell Sea. In these
regions, the surface waters are strongly cooled in winter. When sea ice forms, the concentration of
salt in the ice is very low, and hence the salt concentration in the remaining liquid water increases.
Salt increases the density of water, so the water in the Norwegian and Weddell Seas is very dense
because it is both cold and salty. Note that deep waters form only in winter, and probably not
even every winter. When a lot of fresh water enters the surface waters here (e.g., due to melting
glaciers), no deep water is formed at all. The localized downwellings are balanced by general slow
upwelling throughout the oceans.
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F IGURE 15: A strongly simplified schematic of deep currents and the corresponding surface
counterflow. The wind-driven circulation moves much faster than the thermohaline circulation.
The counterflow drawn in red thus does not show up in the surface currents.
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7 Wind-Driven Currents and the Geostrophic Balance
The wind field ultimately drives the surface water circulation. We’re dealing with oceanography now, so we just take the wind field for granted. Obviously, there is a lot of variability in
the wind field. On average, however, there is a clear pattern. It is much less obvious how this
wind field actually moves the oceans.
F IGURE 16: Average global wind pattern in northern hemisphere summer (June-August).
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F IGURE 17: Average global wind pattern in northern hemisphere winter (December-February).
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Dynamics of the Oceans:
In the lectures on energetics and fluid mechanics, we will treat the governing equations in
more detail. A few short points here:
• All fluids (including glaciers and the Earth’s mantle) are governed by the same general
equations.
• These are the conservation equations of momentum, energy, and mass. We know these
equations very well.
• In addition to the governing equations, we also need to know what is known as “boundary conditions”. We need to specify what happens at all the boundaries of the fluid we
are interested in. One boundary condition for the oceans is that the ocean floor does not
move, and hence the velocity at the bottom of the ocean is zero.
• A more complicated boundary condition is the ocean surface: How do the atmosphere
and the ocean link together?
• One condition at the surface is that the air speed and the water speed right at the interface
are equal. But this boundary condition is not sufficient.
• The second boundary condition specifies how much stress (stress = force/area) wind
exerts on the oceans.
• We know the wind speeds fairly well, but calculating the forcing is very difficult. We
have no fundamental theory for this. Instead, we parameterize it using measurements.
• A good approximation is
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τ = ρair CD U10
(1)
where τ is the stress, ρair is the air density, CD is an experimentally determined drag
coefficient, and U10 is the wind speed at 10 m elevation.
• Measurements for CD are shown below.
• The point of this is not for you to memorize how to calculate wind forcing for the
oceans; but I want you to understand that the interactions between different materials
are typically very poorly understood. This is where much of the uncertainty comes from
in weather and climate models, etc.
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F IGURE 18: Wind drag coefficient CD as a function of wind speed 10 m above the sea. Circles
and triangles are different measurements. Solid line and dashed line are different models (i.e.,
different fits to the data) to be used in ocean modeling.
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The dominant force balance in the shallow ocean away from the coasts is the geostrophic
balance, which is a balance between the pressure gradient force and the Coriolis force:
~ = −ρΩ
~ × ~v
∇P
(2)
Note that on an ellipsoidal Earth, the centrifugal force (or centripetal acceleration) due to
rotation of the Earth is not important for the dynamics: a stationary object does not experience
a net force in the horizontal plane. Only a moving object (~v 6= 0) experiences a force in the
rotating frame.
F IGURE 19: What causes a pressure gradient force? There actually are bumps on the ocean.
Regions of high elevation have high pressure, regions of low elevation have low pressure. Pressure
gradient and coriolis effect balance each other; the particle motion is always at right angles to both.
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F IGURE 20: Detailed picture of global ocean topography and resulting geostrophic currents.
Maximum topography is about 1 m. As we will see soon, the topography is set up by the global
wind patterns.
F IGURE 21: Perfect geostrophy makes currents go around in circles and the pressure gradients are
never removed (the particles never reach the low pressure regions). Viscous dissipation (friction)
makes the trajectories go slightly towards the low pressure.
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F IGURE 22: Water always tries to go from high to low pressure, or opposite the pressure gradient.
In the northern hemisphere, a pressure low (a depression in the water surface) is characterized by
counter-clockwise rotating currents. A pressure high causes clockwise rotation. The orientations
are reversed in the southern hemisphere. Winds in atmospheric high and low pressure regions
follow the same laws.
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The key to understanding the way in which winds force ocean currents is of course once
again the Coriolis effect. Norwegian oceanographer Fridtjof Nansen discovered around 1898
that icebergs in the oceans are not simply pushed along with the wind, but always drift towards
the right at an angle of about 20-40◦ . He gave a qualitative explanation, balancing the driving
stress, the resisting viscous force of the surrounding water, and the Coriolis force. Swedish
oceanographer Vagn Walfrid Ekman shortly therafter presented a quantitative description of
the same general phenomenon. Ekman’s more detailed calculations predicted the so-called
Ekman spiral. The top layer of the ocean is driven by the winds at an angle of 45◦ to the
right (in the N hemisphere). The next deeper layer is driven by the water layer above it. It is
again deflected further to the right. The velocities decrease with depth, so that the net water
transport is at 90◦ to the right of the wind direction.
F IGURE 23: The theoretical Ekman spiral predicts water transport at right angles to the wind
direction. Deflection due to Coriolis effect is to the right in the N and to the left in the S. A perfect
Ekman spiral has not been observed, but the general idea has been verified experimentally. A
similar Ekman spiral exists at the ocean bottom, where moving water interacts with the stationary
seafloor.
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8 The Sverdrup Balance
Harald Sverdrup, another Norwegian oceanographer, combined the geostrophic balance with
a simple approximation of global wind forcing (westerlies at high latitudes and easterlies at
low latitudes). This work was done in the late 1930’s to early 1940’s when no computers or
numerical methods existed. Only very simple systems could be solved at that time.
F IGURE 24: Sverdrup’s solution. He integrated the equations from the eastern boundary. The general pattern of currents is well explained. However, there is no solution for the western boundary
of the ocean.
F IGURE 25: Predicted sea-surface topography from Sverdrup’s solution.
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9 Stommel’s Solution
At about the same time as Sverdrup deriving his main circulation balance, American oceanographer Henry Stommel was beginning to understand why the ocean gyre are always stronger
at the western edge of the oceans (“westward intensification”). Two essential ingredients were
missing in Sverdrup’s work: viscous stresses against the ocean bottom, which are important
near the western boundary where currents are strong and velocities are high; and a Coriolis
force that changes with latitude (Sverdrup assumed for simplicity that the Coriolis force is
independent of latitude).
a)
b)
c)
F IGURE 26: a) Wind-forcing assumed by Stommel (same as Sverdrup); b) Gyre predicted if
viscous stress is taken into account, but Coriolis force is constant with latitude (it isn’t, but this assumption makes calculations much simpler, especially if you don’t have computers); c) Predicted
circulation with westward intensification which results from a Coliolis force that varies with latitude (the Coriolis force is zero at the equator and maximum at the poles).
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10 Munk’s Solution
Walter Munk is an Austrian born oceanographer. His main contribution combines the works
of Sverdrup and Stommel. Stommel’s solution was an illustration of a physical principle, not a
model to predict current strenghts. Munk used actual wind-stress data, limited the flow to the
upper ocean (Sverdrup and Stommel used vertically integrated circulations), and made some
other small changes. The net outcome was a fairly realistic model of the main features of the
oceans’ surface currents.
F IGURE 27: Munk’s solution of global circulation.
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F IGURE 28: Current patterns and strengths in the N Atlantic. Units are in sverdrup, which is
106 m3 /s.
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a)
b)
F IGURE 29: How do we know about currents? One way is to release a lot of buoys that float
along with the currents. a) Buoy tracks; b) calculated flow field. Another way is to use stationary
current meters.
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11 Other Currents: Antarctic Circumpolar Current
F IGURE 30: The Antarctic Circumpolar Current is one of the most important and strongest currents on Earth. It is largely geostrophic. It is different from the main oceanic gyre because it has
no eastern or western boundaries.
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F IGURE 31: Density contours across the Antarctic Circumpolar Current. The oceans are generally
stratified with density increasing with depth. A steep slope in the density contours indicates a
strong horizontal pressure gradient. The Antarctic Circumpolar Current has 3 major streams as
indicated by the shaded zones.
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12 Equatorial Currents
F IGURE 32: Equatorial surface currents are similar to those predicted by Munk. They are 1) the
North Equatorial Current flowing to the west (about 10◦ N), 2) the Equatorial Counter Current
(flowing east between 3 and 10◦ N), and 3) the South Equatorial Current flowing west (on and just
S of the equator). The Equatorial Counter Current really should be centered on the equator (or
so the theory says), but all currents and weather patterns are shifted a bit to the north, probably
because of the asymmetry of the ocean/land ratio between N and S hemispheres.
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F IGURE 33: Cross-section across the equator, showing the Equatorial Undercurrent flowing west
just below the surface. This current was not discovered until the 1950’s. The physical theory
governing the dynamics is not yet fully understood. It should be clear, however, that the dynamics
here are different from the rest of the oceans because the Coriolis force is zero at the equator.
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13 Net Heat Flux Resulting From Surface Currents
F IGURE 34: Heat is taken up by the oceans near the equator (hence, cooling of the atmosphere,
blue colors). It is released again at high latitude (orange color; notice the heating around Scandinavia!).
• For the coupling between oceans and atmosphere, the heat transport by the oceans is of
great importance. Surface waters heat up in the tropics and cool down at high latitude.
The north-south component of the general circulation transports warm water towards
the arctic and cold water towards the equator. The net uptake and release of energy by
the oceans is rather significant.
• Heating and cooling of the oceans is determined by “insolation” (radiation from the sun),
and “radiative” (infrared radiation from the water) “latent” (evaporation) and “sensible”
(conductive) heat fluxes:
[1] Insolation: Depends on height of sun above horizon, and on length of day. Also
depends on transmittance of atmosphere (clouds, water vapor, CO2 ).
30 W/m2 < QSW < 260 W/m2
[2] Radiative heat flux: Atmosphere is relatively transparent to solar radiation (sun
has high surface temperature: “black body radiation” from sun has high frequency
with peak in the visible spectrum). The atmosphere is more opaque to black body
radiation from relatively cool oceans. Much of the radiation is reflected by clouds.
Radiative flux depends on cloud cover, cloud thickness, water vapor and CO2 in
the atmosphere, ice cover, and on ocean surface temperature.
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[3] Latent heat flux: Latent heat flux is influenced primarily by wind speed and relative
humidity. High winds and dry air evaporate much more water than weak winds
with relative humidity near 100%. In polar regions, evaporation from ice covered
ocean is much less than from open water. In the arctic, most of the heat lost from
the sea is through leads (ice-free areas). Hence the percent open water is very
important for the arctic heat budget.
The average annual value for latent-heat flux is in the range:
-130 W/m2 < QL < -10 W/m2
[4] Sensible heat flux is influenced by wind speed and air-sea temperature difference.
High winds and large temperature differences cause high fluxes. Think of this as a
wind-chill factor for the ocean.
The average annual value for sensible-heat flux is in the range:
-42 W/m2 < QS < -2 W/m2
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Name
Abbreviation Temperature (◦ C)
Antarctic Intermediate Water
Antarctic Bottom Water
North Atlantic Deep Water
North Atlantic Bottom Water
Mediterranean Intermediate Water
Subtropical Lower Water
AIW
ABW
NADW
NABW
MIW
U
Salinity (h)
3.3
0.4
4.0
2.5
12.6
18.0
34.15
34.67
35.00
34.90
38.4
35.94
TABLE 1: Major deep water masses and their temperatures and salinities.
14 Deep Circulation: Water Masses
Not nearly as much data has been collected from the deep oceans as from the shallow oceans.
As a result, the dynamics are more poorly understood. Most of our knowledge comes from
the determination of different water masses. Depending on the processes under which a water
mass forms, that water will have a certain temperature and salinity. This combination of T
and S is essentially a fingerprint that tells us where the water comes from. The water masses
are connected in T − S-space due to mixing.
F IGURE 35: Water masses in the Atlantic. Clearly seen is the North Atlantic Deep Water extending like a tongue down and south from the surface in the Greenland and Norwegian Sea. Also
clearly seen is the Antarctic Intermediate Water.
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F IGURE 36: Water masses in different oceans.
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F IGURE 37: Temperature-Salinity plot of water masses in the various oceans. Since mixing is
very inefficient, water masses can be traced for long distances and times. Note that areas on this
plot have nothing to do with volumes of the respective masses: A long line typically represents
mixing of limited volumes of different water types. Large homogeneous volumes of water such as
the Antarctic Bottom Water are represented by a single point.
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Importance of the Deep Circulation:
• The deep circulation carries heat, salinity, oxygen and CO2 between high and low latitudes.
• Deep-sea organisms use up oxygen. The oxygen content of deep water thus gets lower
and lower the further it moves from its source. The water “gets older”.
• Deep currents are slow, but the volumes at depth are much larger than the mixed layer.
Total deep transports are thus comparable to surface transports.
• The deep oceans are a large thermal reservoir which modulates the Earth’s climate.
Deep waters change over timescales of centuries to millenia, and they stabilize climate
over these timescales.
What drives the deep circulation?
• Surface waters cool in arctic regions. Formation of ice (ice contains almost no salt!)
causes residual water to become more saline. Cold, saline water is heavy and sinks.
• Deep water forms in two locations: the Greenland Sea (in the North Atlantic) and the
Wedell Sea (in the South Atlantic).
• Uniform mixing in the oceans brings the cold deep water back to the surface.
• The deep currents are purely geostrophic.
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15 Other Types of Flow: Turbulence and Eddies
F IGURE 38: The major currents often develop instabilities which may grow and eventually
“pinch off” from the main current, thus forming separate rotating cells (“eddies”) of 100-500 km
size. These are essentially very large patterns of turbulence, but the main balance for the oceanic
mesoscale eddies is geostrophic.
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F IGURE 39: More eddies on the Gulf Stream.
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16 Upwellings and Downwellings
F IGURE 40: Wind stress causes net water transport (Ekman spiral!) to the right (in the N hemisphere). If the wind blows along a coastline, it either tries to pull water away from the coast or
pushes it against the coast. Depending on which way the wind blows relative to the coast, a) upwelling or b) downwelling results. Famous examples are the upwellings along the western coast of
the Americas. Generally southward winds N of the equator and northward winds S of the equator
result in upwellings in both locations. The upwelling water comes from just below the thermocline
and is thus very cold (e.g., swimming is not very pleasant around San Francisco). The deeper water
is also very rich in nutrients, leading to a high productivity in algae and fish along these coasts.
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17 The El Niño – Southern Oscillation (ENSO) Event
The observations regarding El Niño are closely related to the upwelling of deep water off the
S American west coast. Surface waters are heated in the tropical regions. Current transport is
to the west, making the surface waters warmer there. Episodically, stronger trade winds pile
up more warm water in the west. When the winds decrease (or often even reverse!), this thick
layer of warm water is transported eastward along the equator. A thick layer of warm water
now resides off the S American coast. Upwelling continues as usual, but the surface layer is
so thick that the upwelling only brings deeper surface waters back to the surface; the deep
water remains untapped. Without the nutrients, the fisheries collapse. This usually happens
around Christmas, and hence the Spanish name “El Niño”, relating to “the boy” or Jesus. It is
now realized that the event is not localized to the S Pacific, but it is a global event that is most
pronounced on the southern hemisphere.
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F IGURE 41: Sea surface temperature anomalies during an El Niño year.
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