CHAPTER
7
Ocean Structure
and Circulation
Learning Outcomes
Chapter Outline
After studying the information in this chapter students
should be able to:
7.1 Ocean Structure 196
7.2 Thermohaline Circulation and
Water Masses 200
7.3 The Layered Oceans 203
7.4 What Drives the Surface Currents? 205
7.5 Ocean Surface Currents 207
7.6 Current Characteristics 211
7.7 Eddies 213
7.8 Convergence and Divergence 214
7.9 The Great Ocean Conveyor Belt 218
7.10 Changing Circulation Patterns 218
7.11 Measuring the Currents 220
Diving In: Ocean Drifters 221
Summary 224
Key Terms 225
Study Problems 225
1. estimate the density of a mixture of two samples of seawater
that have the same density but different temperatures and
salinities,
2. describe and sketch changes in the seasonal thermocline at
midlatitudes through the year,
3. plot temperature and salinity as a function of depth and
identify the thermocline and halocline,
4. list five different water masses and describe how they form,
5. relate surface convergence and divergence to downwelling
and upwelling,
6. describe the properties of water masses in each ocean
basin,
7. describe and sketch the motion of water in the Ekman layer,
8. diagram the formation of surface current gyres,
9. locate the major surface currents on a map of the ocean
basins,
10. explain the process of western intensification, and
11. sketch the Great Ocean Conveyor Belt.
Sea surface temperature (SST) simulation created by scientists at the
National Oceanic and Atmospheric Administration’s Geophysical Fluid
Dynamics Laboratory using a coupled atmosphere-ocean model. Currents
and eddies off the southern tip of Africa are evident.
Source: National Oceanic and Atmospheric Administration (NOAA), Geophysical Fluid Dynamics
Laboratory
195
E
∙∙ a surface layer tens to a few hundreds of meters thick,
called the mixed layer;
∙∙ a region called the thermocline, extending from the
bottom of the mixed layer to a depth of about 1000 m
(3280 ft); and
∙∙ the region from the base of the thermocline to the sea floor.
arth is surrounded by two great oceans: an ocean
of air and an ocean of water. Both are in constant
motion, driven by the energy of the Sun and the
gravity of Earth. Hidden below the ocean’s surface is its
structure. If we could remove a slice of ocean water in the
same way we might cut a slice of cake, we would find that,
like a cake, the ocean is a layered system. The layers are
invisible to us, but they can be detected by measuring the
changing temperature and salt content, and by calculating
the density of the water from the surface to the ocean
floor. This layered structure is a dynamic response to
processes that occur at the surface: the gain and loss of
heat, the evaporation and addition of water, the freezing
and thawing of ice, and the movement of water in response
to wind. These surface processes produce a series of
horizontally moving layers of water, as well as local areas
of vertical motion. Surface currents carry heat from one
location to another, altering Earth’s surface temperature
patterns and modifying the air above. The interaction
between the atmosphere and the ocean is dynamic; as one
system drives the other, the driven system acts to alter the
properties of the driving system.
In this chapter, we will study both the surface processes
and their below-the-surface results in order to understand why
the ocean is structured and how its structure is maintained. We
will also explore the formation of the ocean’s surface currents.
We follow these currents as they flow, merge, and move away
from each other. We examine both horizontal and vertical
circulation, and consider ways in which they are linked to the
overall interaction between the atmosphere and the ocean.
The mixed layer is an isothermal layer—a layer of constant
temperature. The thickness of the mixed layer is variable. It depends on the depth to which the surface water is mixed by turbulence caused by waves and wind. The mixed layer can be as thick
as 200 to 300 m (~650 to 1000 ft) at midlatitudes in the open
ocean, whereas in protected coastal waters in the summer, it can
be as little as 10 m (~33 ft) thick. Between about 200 and 300 m
and 1000 m depth is the thermocline, where temperature decreases rapidly throughout much of the ocean. This layer is
also frequently known as the “permanent thermocline” because
seasonal changes in climate at the surface do not influence water
temperature at these depths. Below 1000 m depth is the third,
and largest in volume, layer, which extends to the sea floor. This
deep water is nearly isothermal, temperature decreases very
slowly with depth, and it is uniformly cold everywhere. Roughly
75% of the water in the ocean has a temperature between 0 and
4°C (table 7.1 and fig. 7.3). The discovery that the deep water
of the ocean, even in tropical regions, is very cold was made in
the eighteenth century. The obvious conclusion that followed
was that deep seawater must originate in polar regions, where
cold, dense surface water sinks and flows toward the equator
along the ocean floor.
The details of actual temperature-versus-depth profiles
vary considerably depending on latitude and season of the year
(fig. 7.4). At midlatitudes, the temperature and depth of the
mixed layer undergo seasonal changes with the formation of a
shallow seasonal thermocline in the summer and its disappearance in the winter (fig. 7.4a). Throughout the winter, when strong
400
Variation of Temperature
with Depth
With very few exceptions, the temperature of
seawater decreases with depth. Effectively all
of the energy available to heat the ocean comes
from incoming solar radiation. Consequently,
only a thin surface layer of the ocean is heated
directly because of how rapidly solar radiation
is absorbed with depth (fig. 7.1). Nearly half
of the total solar energy at the sea surface is
absorbed within 10 cm of the surface, and
all of the infrared energy is absorbed within
about a meter of the surface. A typical seawater
temperature-versus-depth profile consists of
three “temperature layers” (fig. 7.2):
196
Visible part
Energy (arbitrary units)
7.1 Ocean Structure
Infrared
300
200
100
60%
45%
20% 1 m
10 m
Surface
0.1 m
100 m
3%
0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0
100%
Total Energy
1.1 1.2 1.3 1.4 1.5 1.6 1.7 1.8 1.9 2.0 2.1 2.2 2.3 2.4
Wavelength (10–4cm)
Figure 7.1 A simplified plot of total solar energy in seawater as a function of depth. Area
below each curve is representative of the percent of available solar energy at the surface
that reaches that depth. A little over half of the total solar energy at the surface is absorbed
in the upper 1 m of water (only 45% penetrates to a depth of 1 m). Infrared energy is
absorbed particularly rapidly, with nearly all of it absorbed in the upper 1 m of water.
7.1 Ocean Structure
0
0
4
Temperature (°C)
8
12
16
60,000
20
24
Mixed layer
50,000
Thermocline
197
Figure 7.3 Distribution of seawater
temperature. Histograms indicate
the number of cubic kilometers of
seawater in each 2°C range.
1000
2000
km3
Depth (m)
40,000
30,000
3000
20,000
4000
10,000
Figure 7.2 Simplified profile of temperature with depth in the
ocean. A shallow mixed layer with relatively constant temperature
overlies the thermocline, where temperature decreases rapidly
with depth. Beneath the thermocline, below a depth of about
1000 m, temperature is fairly constant and cold.
Table 7.1 Distribution of Seawater Temperature
Temperature
Range (°C)
Seawater Volume
(km3)
Percent of
Ocean Volume
–2–0
6409
4.7
0–2
57,990
42.4
2–4
40,065
29.3
4–6
11,849
8.7
6–8
6059
4.4
8–10
4222
3.1
10–12
2632
1.9
12–14
2358
1.7
14–16
1340
1.0
16–18
985
0.7
18–20
685
0.5
20–22
638
0.5
22–24
475
0.3
24–26
426
0.3
26–28
489
0.4
28–30
269
0.2
30–32
2
0.001
0
–2 0 2 4 6 8 10 12 14 16 18 20 22 24 26 28 30 32
Temperature (°C)
winds produce deep mixing and surface-water temperatures
are cold, the mixed layer may extend all the way to the top of
the permanent thermocline, producing an essentially vertical temperature profile in the upper 200 to 300 m (fig. 7.5,
March). Throughout the summer, as surface temperatures
rise and winds decrease, the mixed layer will become more
shallow and a strong (steep temperature gradient) seasonal
thermocline can develop above the permanent thermocline
(fig. 7.5, August). At low latitudes, surface temperatures
are warm and constant throughout the year. Consequently,
there is no development of seasonal thermoclines and the
three distinct temperature layers are quite stable (see fig. 7.4b).
At high latitudes above about 60°, there is no permanent
thermocline (see fig. 7.4c). However, weak seasonal thermoclines can develop in the summer. At high latitudes
there is often a layer of cold water 50 to 100 m below the
surface. Overall, the presence of three temperature layers in
the ocean is illustrated in the plot of seawater temperature in the
Pacific Ocean versus depth and latitude shown in figure 7.6.
Variation of Salinity with Depth
As discussed in chapter 5, the salinity of surface seawater varies
as a function of latitude in a relatively stable pattern (review figs.
5.2 and 5.3). However, it is more difficult to draw general conclusions about vertical profiles of salinity versus depth. Depending on
latitude, salinity may be relatively constant, decrease, or increase
with depth down to about 1000 m. Below about 1000 m depth, the
influence of surface processes is minor and salinity is fairly constant. The total salinity range of 75% of the ocean is between 34.5
198
Chapter 7
Ocean Structure and Circulation
Temperature (°C)
5
10
0
15
20
0
Temperature (°C)
10
15
5
20
25
Temperature (°C)
–5
0
5
0
Seasonal
thermocline
500
Permanent
thermocline
Permanent
thermocline
Depth (m)
1000
1500
2000
2500
Midlatitudes
3000
High
latitudes
Low latitudes
(a)
(b)
(c)
Figure 7.4 Simple temperature-versus-depth profiles for three latitude zones: (a) midlatitudes with significant seasonal variation,
(b) low latitudes where climate tends to be uniformly warm through the year, and (c) high latitudes where climate tends to be uniformly
cold through the year.
2
4
6
Temperature (°C)
8
10
12
14
16
0
March
May
July
August
20
Variation of Density with Depth
Depth (m)
September
40
60
change in salinity with depth is called the halocline
(fig. 7.8). The general distribution of seawater salinity in the Pacific Ocean is illustrated in the plot of
salinity versus depth and latitude shown in figure 7.9.
November
Variations in temperature, salinity, and pressure
(depth) combine to control the density of seawater.
Seawater density is inversely proportional to temperature and directly proportional to salinity and
pressure. This can be expressed as (↑ = increases,
and ↓ = decreases):
Density ↑ as temperature ↓
Density ↑ as salinity ↑
Density ↑ as pressure ↑
80
In general, variations in temperature and salinity
are more influential in determining seawater
100
density than are variations in pressure (or depth).
Figure 7.5 Detailed variation in the depth of the seasonal thermocline and the
Consequently, we can consider density to be a
temperature of the surface mixed layer at midlatitudes. Solid lines indicate the
function of temperature and salinity. The stratifigrowth of the seasonal thermocline as it increases in strength and shoals during
cation of the ocean by temperature and salinity that
the summer, and dashed lines indicate its decay as it becomes deeper and
we have just discussed results in a stratification
weaker in winter. Note the different scales compared to figure 7.4.
of the ocean into density layers also. A typical
­seawater density-versus-depth profile consists of
and 35.0 ppt (table 7.2 and fig. 7.7). In the upper 1000 m, areas
three “­density layers” (fig. 7.10):
where salinity decreases with depth are typically found at low and
middle latitudes, whereas areas where salinity increases with depth
∙∙ the mixed layer, a surface layer tens to a few hundreds of
are typically found at high latitudes. In either case, this marked
meters thick;
January
199
7.1 Ocean Structure
60°S
55°
50°
45°
40°
35°
30°
25°
20°
15°
0
10°
25
5°
0°
5°
15°
20°
25°
30°
35°
40°
45°
50°
55°N
10
15
500
10°
28
5
1000
1500
2000
2.0
Depth (m)
2500
0.0
3000
3500
4000
1.0
4500
5000
5500
6000
6500
0
500
1000
1500
2000
2500
3000
3500
4000
4500
5000
5500
6000
6500
7000
7500
8000
8500
9000
9500
10,000 10,500 11,000
11,500 12,000 12,500 13,000 13,500
Distance (km)
Figure 7.6 Ocean temperature versus depth and latitude. North–south profile through the Pacific Ocean along the 150°W meridian.
Temperature decreases with depth. The depth of the surface mixed layer and the permanent thermocline are less at low latitudes than
at midlatitudes because winds are generally weaker and seasonal temperature contrasts are less at low latitudes. The black represents
seafloor bathymetry.
∙∙ a region called the pycnocline (which closely follows the
thermocline), extending from the bottom of the mixed
layer to a depth of about 1000 m (3280 ft); and
∙∙ the region from the base of the pycnocline to the sea floor.
40,000
QUICK REVIEW
km3
30,000
1. What processes control the depth of the mixed layer?
2. Explain the difference between a seasonal thermocline
and the permanent thermocline.
3. How does the seasonal thermocline change during
the year?
4. Describe the basic characteristics of a temperatureversus-depth profile at high latitudes, midlatitudes,
and low latitudes.
5. Describe the total salinity range of seawater.
6. What is the halocline?
7. How does seawater density depend on temperature
and salinity?
20,000
10,000
1 2 3 4 5 6 7 8 9
1 2 3 4 5 6 7 8 9
1 2 3 4 5
1 2 3 4 5 6 7 8 9
35.0
36.0
33.0
34.0
Salinity
> 36.5
< 33.0
0
Figure 7.7 Distribution of seawater salinity. Histograms indicate the number of cubic kilometers
of seawater in each 0.1 ppt range.
200
Chapter 7
Ocean Structure and Circulation
33
0
Table 7.2 Distribution of Seawater Salinity
Percent of
Ocean Volume
36.5
37
S/L
Halocline
1000
High
latitudes
Depth (m)
Seawater Volume
(km3)
Salinity Range
(ppt)
Increasing salinity (‰)
34
35
36
Low
latitudes
2000
3000
4000
Figure 7.8 Simplified profiles of salinity with depth in the ocean
for high and low latitudes. In each case, there is a surface
mixed layer of relatively constant salinity. Beneath the mixed
layer at high latitudes, salinity decreases rapidly to a depth of
about 1000 m. Beneath the mixed layer at low latitudes, salinity
increases rapidly to a depth of about 1000 m. The region of rapid
change in salinity is called the halocline.
7.2 Thermohaline Circulation
and Water Masses
Thermohaline Circulation
If the density of the water increases with depth, the water column
from the surface to that depth is stable. If there is higher-density
water on top of lower-density water, the water column is
unstable. An unstable water column cannot persist; the denser
surface water sinks and the less-dense water at depth rises to
replace the water above it. When dense water from the surface
sinks and reaches a level at which it is denser than the water
above but less dense than the water below, it spreads horizontally as more water descends. At the surface, water moves
horizontally into the region where sinking is occurring. The
dense water that has descended displaces deeper water upward,
completing the cycle. Because water is a fixed quantity in the
oceans, it cannot be accumulated at one location or removed at
another location without movement of water between those locations. This concept is called continuity of flow. This motion,
caused by variations in density due to differences in temperature
and salinity, is called thermohaline circulation. Areas of thermohaline circulation where water sinks are called downwelling
zones; areas of rising waters are upwelling zones. Downwelling
is a mechanism that transports oxygen-rich surface water to
depth, where it is needed for deep-living animals. Upwelling returns low oxygen-content water with dissolved, decay-produced
nutrients to the surface, where the nutrients act as fertilizers to
60°S
55°
0
50°
45°
40°
35°
34.0
30°
25°
20°
15°
10°
0°
5°
10°
15°
20°
25°
30°
35°
40°
45°
50°
55°N
34.0
35.0
500
5°
35.5
34.7
1000
34.0
34.6
34.3
34.3
34.5
34.5
34.60
1500
2000
Depth (m)
2500
34.73
3000
34.70
3500
34.68
4000
4500
5000
5500
6000
6500
0
500
1000
1500
2000 2500
3000
3500 4000 4500 5000 5500
6000 6500 7000
7500 8000 8500 9000 9500 10,000 10,500 11,000 11,500 12,000 12,500 13,000 13,500
Distance (km)
Figure 7.9 Seawater salinity versus depth and latitude. North–south profile through the Pacific Ocean along the 150°W meridian. Most
of the variation in salinity is found in surface water with high values at low and middle latitudes and low values at high latitudes. Beneath
about 1000 to 1500 m, salinity shows only minor variations. The black represents seafloor bathymetry.
3000
promote photosynthesis and the production of more oxygen in
the sunlit surface waters. Upwelling and downwelling can also
be caused by wind-driven surface currents. When the surface
waters are driven together by the wind or against a coast, a
surface convergence is formed. Water at a surface convergence
sinks, or downwells. When the wind blows surface waters away
from an area or a coast, a surface divergence occurs and water
upwells from below. Surface convergences and divergences are
discussed further later in this chapter.
The speed of upwelling and downwelling water is about
0.1 to 1.5 m (0.3 to 5 ft) per day. Compare these flows to ocean
surface currents, which reach speeds of 1.5 m (5 ft) per second.
Horizontal movement at mid-ocean depth due to thermohaline
flow is about 0.01 cm (0.004 in) per second. Water caught in this
slow but relentlessly moving cycle can spend 1000 years at the
greater ocean depths before it again reaches the surface.
If the water column has the same density over depth, it has
neutral stability and is termed isopycnal. A neutrally stable
water column can easily be mixed vertically by wind, wave
action, and currents. If the water temperature is unchanging over
depth, the water column is isothermal; if the salinity is constant
over depth, it is isohaline.
3500
Water Masses
1.023
1.024
0
Mixed layer
Density (g/cm3)
1.025
1.026
1.027
1.028
Pycnocline
500
Depth (m)
1000
1500
2000
2500
Figure 7.10 Simplified profile of density with depth in the ocean.
A shallow mixed layer with relatively constant density overlies
the pycnocline, where density increases rapidly with depth (this
layer corresponds closely to the thermocline shown in figure 7.2).
Beneath the pycnocline, density increases slowly to the sea floor.
There are many different combinations of temperature and
salinity that produce the same seawater density. This can be
illustrated in a temperature-salinity diagram, or T-S diagram,
where lines of equal density, or isopycnals, are drawn for
many combinations of temperature and salinity (fig. 7.11).
201
202
Chapter 7
Float ID# 1900036
Temperature-Salinity Diagram
(lines of equal density in gcm–3)
20
5
A
0
33.5
55
1.02
ρ=
60
B
1.02
ρ=
5
6
2
1.0
ρ=
70
1.02
ρ=
75 25% A – 75% B Mixture
1.02
0
28
ρ=
1.0
=
ρ
5
50% A – 50% B
28
1.0
Mixture
ρ=
75% A – 25% B
Mixture
34.0
34.5
35.0
35.5
Salinity ‰
36.0
1
0
5
10
15
20
25
20
25
500
~Depth (m)
10
ρ=
50
1.02
1000
1500
1.0
29
0
15
Cycle number/relative temperature
45
1.02
ρ=
ρ=
Temperature (°C)
Ocean Structure and Circulation
2000
0
(a)
5
10
15
Temperature (°C)
36.5
Figure 7.11 The density of seawater, measured in grams per
As the salinity increases, the density increases; as the temperature increases, the density decreases. Salinity can be increased
by evaporation or by the formation of sea ice; it can be decreased
by precipitation, the inflow of river water, the melting of ice, or
a combination of these factors.
Any seawater sample can be plotted on a T-S diagram as a
single point and the density of the sample can be determined from
the isopycnals. Seawater samples from a common source that all
plot very near each other on a T-S diagram, with a narrow range
of temperature and salinity, define a water type. When two water types (e.g., A and B in fig. 7.11) having the same density, but
different values of salinity and temperature, are mixed, they form
a new water type that lies on a mixing line (a straight line drawn
from point A to point B). The location of the new water type on
the mixing line will depend on the relative amounts of the two
original water types that were combined to form the mix. Notice
that the isopycnals are curved. Consequently, the mixture of two
water types with the same density will produce a new water type
with a density greater than either of the two original water types,
and the newly created mixed water will sink. This mixing and
sinking process is known as caballing and it occurs wherever
surface waters converge and mix.
A water mass is a large body of water that has similar values of temperature and salinity throughout. Water masses can
be identified by measuring water temperature and salinity along
vertical profiles in the ocean and plotting these data on a T-S
diagram (fig. 7.12a and b). A water mass will be represented
on a T-S diagram by closely clustered data points (fig. 7.12c).
Surface-water masses typically have greater variability in temperature and salinity than do deeper water masses.
Cycle number/relative salinity
0
1
5
10
15
20
25
~Depth (m)
500
1000
1500
2000
(b)
34.5
35.0
35.5
36.0
36.5
37.0
Salinity (PSU)
37.5
25
Surface waters
20
Temperature (°C)
cubic centimeter, is abbreviated as ρ (rho) and varies with
temperature and salinity. Many combinations of salinity and
temperature produce the same density. Low densities are at the
upper left, and high densities are at the lower right. The straight
line is the mixing line for waters A and B, both with the same
density. A mixture of A and B lies on the mixing line and is more
dense than either A or B.
North Atlantic
central waters
15
10
Mediterranean water
5
Upper North Atlantic deep water
0
34.5
35.0
2000
(c)
35.5
36.0
36.5
Salinity (PSU)
1500
1000
Pressure (decibars)
500
37.0
37.5
0
Figure 7.12 ARGO float measurements north of the Canary
Islands in the North Atlantic (about 31° to 37°N, 25° to 27°W)
along vertical profiles of (a) temperature and (b) salinity. (c) T-S
plot of the data collected can be used to identify specific water
masses. Notice the greater range of temperature and salinity
found in surface waters (red) compared to deep waters (violet).
7.3 The Layered Oceans
zone extends to 200 m (650 ft), and the ­intermediate zone lies
between 200 and 2000 m (1000 and 6500 ft). Deep water is
found between 2000 and 4000 m (6500 and 13,000 ft), and
bottom water is below 4000 m (13,000 ft).
QUICK REVIEW
1. Describe the concept of continuity of flow.
2. How do stable and unstable water columns differ?
3. What is the relationship between upwelling and
downwelling? between surface divergence and
convergence?
4. What is the difference between a water type and a
water mass?
5. Explain how water types with different temperature
and salinity can have the same density.
The Atlantic Ocean
The properties of the layers of water making up the Atlantic
Ocean are shown in figure 7.13. At the surface in the North
Atlantic, water from high northern latitudes moves southward, while water from low latitudes moves northward along
the coast of North America and then east across the North
­Atlantic. These waters converge in areas of cool temperatures
and high precipitation at approximately 50° to 60°N in the
Norwegian Sea and at the boundaries of the Gulf Stream and
the Greenland and Labrador Currents. The resulting mixed
water has a salinity of about 34.9‰ and a temperature of
2° to 4°C. This water, known as North Atlantic deep water
(NADW), sinks and moves southward. North Atlantic deep
water from the Norwegian Sea moves south along the east
side of the Atlantic, while water formed at the boundary of the
Labrador Current and the Gulf Stream flows along the western
side. Above this water at 30°N, a low-density lens of very
salty (36.5‰) but very warm (25°C) surface water remains
trapped by the circular movement of the major oceanic surface
currents. Between this surface water and the North Atlantic
deep water lies water of intermediate temperature (10°C) and
salinity (35.5‰). This water is a mixture of surface water and
the upwelled colder, saltier water from the subtropical regions.
It moves northward to reappear at the surface south of the
convergence in the North Atlantic.
7.3 The Layered Oceans
Oceanographers have taken salinity and temperature measurements
with depth at many locations and for many years. Gradually, they
accumulated sufficient data to identify the layers of water that
make up each ocean and the surface source of the water forming
each layer. The structure of an ocean is determined by the properties of these layers. Each layer received its characteristic salinity,
temperature, and density at the surface. The water’s density controls
the depth to which the water sinks; the thickness and horizontal
extent of each layer are related to the rate of its formation and the
size of the surface source region. Water that sinks from the surface
to spread out at depth and slowly mixes with adjacent layers eventually rises at another location. In all cases, water that sinks displaces
an equivalent volume of water upward toward the surface at some
other location so that the oceans’ vertical circulation is continuous.
The layers of water just described are associated with depth
zones: surface, intermediate, deep, and bottom. The surface
Antarctic
intermediate water
Surface water
36.5‰
25°C
Mediterranean
outflow water
203
Subantarctic
front
Antarctic
front
An
tar
c
tic
a
Africa
1
35.5‰
10°C
34.4‰
2
34.2‰
5°C
34.9‰
2–4°C
3
4
5
Depth (km)
6
80°N
60°N
34.6‰
2°C
North Atlantic
deep water
34.8‰
–0.5°C
Antarctic
bottom water
40°N
Surface convergence
Surface divergence
20°N
Figure 7.13 North–south cross section of the
0°
20°S
40°S
60°S
80°S
Atlantic Ocean illustrating major water masses
and their general movement. Areas of surface
convergence and divergence associated with
downwelling and upwelling are also shown.
204
Chapter 7
Ocean Structure and Circulation
Near the equator, the upper boundary of the North Atlantic deep water is formed by water produced at the convergence
centered about 40°S. This is Antarctic intermediate water
­
(AAIW). Because it is warmer (5°C) and less salty (34.4‰)
than the North Atlantic deep water, it is less dense and remains
above the denser and saltier water below. Along the edge of Antarctica, very cold (–0.5°C), salty (34.8‰), and dense water is
produced at the surface by sea ice formation during the Southern
Hemisphere’s winter. This is Antarctic bottom water (AABW),
the densest water in the oceans. This water sinks to the ocean floor
and flows slowly northward, creeping beneath North Atlantic deep
water, as it continues on through the deep South Atlantic Ocean
basins west of the Mid-Atlantic Ridge. Antarctic bottom water
does not accumulate enough thickness to be able to flow over
the mid-ocean ridge system into the basins on the African side
of the ridge; it is confined to the deep basins on the west side of
the South Atlantic and has been found as far north as the equator.
At the same time, the North Atlantic deep water between
the Antarctic bottom and intermediate waters rises to the ocean’s
surface in the area of the 60°S divergence. As it reaches the surface, it splits; part moves northward as South Atlantic surface
water and Antarctic intermediate water; part moves southward
toward Antarctica, to be cooled and modified to form Antarctic bottom water. A mixture of North Atlantic deep water and
Antarctic bottom water becomes the circumpolar water for the
Southern Ocean that flows around Antarctica. The Antarctic
circumpolar water becomes the source of the deep water found
in the Indian and Pacific Oceans. In this way, the Atlantic Ocean
and its circulation play a defining role in the structure and circulation of all the oceans.
Warm (25°C), salty (36.5‰) surface water in the South
­Atlantic is also caught by the circular current pattern at the
surface and is centered about 30°S. South of the southern tips
of South America and Africa, the water flows eastward, driven
by the prevailing westerly winds, which move the water around
and around Antarctica.
Water from the Mediterranean Sea has a temperature of about
13°C and a salinity of 37.3‰ as it leaves the Strait of Gibraltar.
This water, mixing with Atlantic Ocean water, forms an intermediate density water, Mediterranean intermediate water (MIW),
also known as Mediterranean outflow water (MOW), that sinks in
the North Atlantic to a depth of approximately 1000 m (3300 ft).
The influence of Mediterranean water can be traced 2500 km
(1500 mi) from the Strait of Gibraltar before it is lost through
­modification and mixing.
Because the Atlantic Ocean is a narrow, confined ocean of
relatively small volume but great north–south extent, the water
types are readily identifiable and the movement of the layers can
be followed quite easily. In addition, the bordering nations of the
Atlantic have had a long-standing interest in oceanography, so
the vertical circulation and layering of the Atlantic are the most
studied and the best understood of all the oceans.
The Pacific Ocean
In the vast Pacific Ocean, waters that sink from relatively small
areas of surface convergences lose their identity rapidly, making
the layers difficult to distinguish. Antarctic bottom water forms
in small amounts along the Pacific rim of Antarctica, but it
is quickly lost in the great volume of the Pacific Ocean. The
deeper water of the South Pacific Ocean is the water of the Antarctic circumpolar flow. Because the North Pacific is isolated
from the Arctic Ocean, only a small amount of water comparable to North Atlantic deep water can be formed. In the extreme
western North Pacific, convergence of the southward-flowing
cold water from the Bering Sea and the Sea of Okhotsk and
the northward-­moving water from the lower latitudes produces
only a small volume of water that sinks to mid-depths. There is
no large source of deep water similar to that found in the North
Atlantic. Warm, salty surface water occurs at subtropical latitudes
(30°N and 30°S) in each hemisphere, and Antarctic intermediate
water is produced in small quantities, but its influence is small.
Deep-water flows in the Pacific are sluggish, and conditions are
very uniform below 2000 m (6600 ft). The slow circulation of
the Pacific means that it has the oldest water at depths where
age is measured as time from the water’s last contact with the
surface. Residence time for deep water in the Pacific is about
twice that of deep water in the Atlantic.
The Indian Ocean
The Indian Ocean is principally an ocean of the Southern Hemisphere and has no counterpart of the North Atlantic deep water.
Small amounts of Antarctic bottom water are soon mixed with
the deeper waters to form a fairly uniform mixture of Antarctic
circumpolar water brought into the Indian Ocean by the Antarctic circumpolar current. There is a small amount of Antarctic
intermediate water, and in the subtropics, a lens of warm, salty
water occurs at the surface.
The Arctic Ocean
The Arctic Ocean basin is unique. About one-third of its area
is covered by extensive continental shelves, the widest of any
ocean basin. Two basins, the Eurasian to the east and the Canadian to the west, occupy the central portion of the ocean; they
are separated by the Lomonosov Ridge extending due north
from Greenland.
The density of Arctic Ocean water is controlled more by
salinity than by temperature. Its surface layer is formed from
low-salinity water entering from the Bering Sea, fresh water
from Siberian and Canadian rivers, and seasonal melting of sea
ice. The surface layer from these combined sources is about
80 m (250 ft) deep and has a low salinity (32.5‰) and a low
temperature (–1.5°C). Below the surface layer, salinity increases
with depth in the halocline layer, 200 m (650 ft) thick, to reach
34.5‰ at its base. The cold, salty water of the halocline layer
is produced by annual freezing and formation of sea ice over
the continental shelves bordering the ocean. This water sinks
and moves across the shelves to spread out in the central ocean
basins. West of Spitsbergen, Norway, North Atlantic water (2°C
and 35‰) enters the Arctic Ocean and is cooled as it flows
under the halocline and fills the Arctic Ocean basins. This water
205
7.4 What Drives the Surface Currents?
upwells along the edge of the continental shelves, mixing with
the water formed during freezing, and exits the Arctic as water
of 0.5°C temperature and 34.9‰ salinity along the edge of the
shelf adjacent to Greenland. This exiting water moves south
along the coast of Greenland and enters the North Atlantic south
of Greenland and Iceland, where it combines with Gulf Stream
water to form North Atlantic deep water.
Internal Mixing
Mixing between waters in the ocean is most active when turbulence and energy of motion are available to stir the waters and
blend their properties. At the sea surface, wind-driven waves and
currents supply energy for mixing, and the tides create currents
at all depths. The large eddies that may form at the boundaries of currents also stir together dissimilar waters, acting to
homogenize them. When surface currents converge, mixing at
current boundaries may produce caballing of the mixed water.
When currents and their associated turbulence are weak, mixing is reduced. Mixing by diffusion occurs continually at the
molecular level, but diffusion is much weaker than mixing by
turbulent processes.
If a parcel of water is displaced vertically by turbulence,
buoyancy forces tend to return the parcel to its original density
level. Therefore, vertical mixing between the water types that
form the oceans’ internal layers is weak. Horizontal mixing is
more efficient because it requires less energy than vertical mixing.
A parcel of water displaced horizontally along a surface of constant density remains at its new position and shares its properties
with the surrounding water.
In areas under warm, high-salinity surface water with
an ­appreciable salinity and temperature decrease with depth,
­internal vertical mixing processes occur despite the stability of
the water column. Vertical columnar flows, approximately 3 cm
(5 in) in diameter, are called salt ­fingers; they develop and mix
the w
­ ater vertically, causing a stair-step salinity and ­temperature
change with depth. This phenomenon is caused by the ability of
seawater to gain or lose heat faster by conduction than it gains or
loses salt by diffusion. This causes the d­ ensity of the vertically
moving ­water to change relative to that of the surrounding water,
and the moving water is propelled either up or down. Salt fingers
mix water over limited depths, creating h­ omogeneous layers
30 m (100 ft) thick. These layers e­ xist from about 150 to 700 m
(500 to 2300 ft) deep and are estimated to occur over large areas
of the oceans when the required ­conditions are present.
7.4 What Drives the
Surface Currents?
When the winds blow over the oceans, they set the surface w
­ ater
in motion, driving the large-scale surface currents in nearly
constant patterns. The density of water is about 1000 times
greater than the density of air, and once in motion, the mass of
the moving water is so great that its inertia keeps it flowing.
The currents flow more in response to the average ­atmospheric
circulation than to the daily weather and its short-term changes;
however, the major currents do shift slightly in response to seasonal changes in the winds. The currents are further modified by
interactions between the currents and along zones of converging
and diverging water. The major surface currents have been called
the rivers of the sea; they have no banks to contain them, but
they maintain their ­average course.
Because the frictional coupling between the ocean water
and Earth’s surface is small, the moving water is deflected by
the Coriolis effect in the same way that moving air is deflected
(see chapter 6). But because water moves more slowly than air,
it takes longer to move water the same distance as air. During
this longer time period, Earth rotates farther out from under the
water than from under the wind. Therefore, the slower-moving
water appears to be deflected to a greater ­degree than the overlying air. The surface-current acted upon by the Coriolis effect is
deflected to the right of the driving wind direction in the Northern Hemisphere and to the left in the Southern Hemisphere. In
the open sea, the surface flow is deflected at a 45° angle from
the wind direction, as shown in figure 7.14.
Current speed
and direction
Wind stress
45°
10 m
20 m
30 m
40 m
Each layer
drags the
layer below
50 m
Mean direction
and flow
60 m
QUICK REVIEW
1. Identify the water masses of the Atlantic Ocean. What is
the origin of each? In which direction does each flow?
2. Why is the layering of the Pacific Ocean less dramatic
than the layering of the Atlantic Ocean?
3. In what ways is the water of the Indian Ocean similar to
the water of the South Atlantic Ocean?
4. What is the origin of Arctic Ocean deep water?
70 m
80 m
90 m
100 m
Figure 7.14 The Ekman spiral and Ekman transport. Water
motion in the surface “Ekman layer” is due to wind stress.
206
Chapter 7
Ocean Structure and Circulation
N
Westerlies
The Ekman Spiral and Ekman Transport
Wind-driven surface water sets the water immediately below it
in motion. But because of low-friction coupling in the water, this
next deeper layer moves more slowly than the surface layer and
is deflected to the right (Northern Hemisphere) or left (Southern
Hemisphere) of the surface-layer direction. The same is true for
the next layer down and the next. The result is a spiral in which
each deeper layer moves more slowly and with a greater angle
of deflection to the surface flow (fig. 7.14). This current spiral
is called the Ekman spiral, after the physicist Vagn Walfrid
Ekman, who developed its mathematical relationship. The
spiral extends to a depth of approximately 100 to 150 m (330 to
500 ft), where the much-reduced current will be moving in the
opposite direction to the surface current. The surface layer of
water corresponding to the Ekman spiral is known as the Ekman
layer. Over the depth of the spiral, the average flow of the water
set in motion by the wind, or the net flow, moves 90° to the right
(Northern Hemisphere) or left (Southern Hemisphere) of the
surface wind. This is known as Ekman transport. This relationship is in contrast to the surface water, which moves at an angle
of 45° to the wind direction.
40°N
Elevated convergence
30°N
Northeast
trades
Doldrums
Tropical
E divergence
0°
W
Southeast
trades
Elevated convergence
30°S
Westerlies
40°S
S
Winds
Ekman
transport
Resulting
surface currents
Figure 7.15 Wind-driven transport and resulting surface currents in
Ocean Gyres
an ocean bounded by land to the east and to the west. The currents
form large oceanic gyres that rotate clockwise in the Northern
The major surface currents in the ocean are driven primarily
Hemisphere and counterclockwise in the Southern Hemisphere.
by the trade winds, blowing in a westerly direction toward the
equator, and the westerlies, blowing in an easterly direction away
from the equator (review figure 6.20). Surface water is driven at
(fig. 7.16). A convergent lens of ­surface ­water is elevated as
a 45° angle to the direction of these winds (fig. 7.15). Thus, the
much as 2 m (6.5 ft) above the ­equilibrium sea level, and this
trade winds drive surface currents that flow from east to west on
lens depresses the underlying denser water.
either side of the equator. When these currents reach the westThe thickness of the surface lens is about 1000 times greater
ern boundaries of ocean basins, they are deflected away from
than the elevation of the lens above sea level. This is because the
the equator and move to higher latitudes where they
enter the region of the westerlies. Once under the
Figure 7.16 There are five major ocean gyres: (a) Indian Ocean, (b) North Pacific,
influence of the westerlies, the surface currents are (c) South Pacific, (d) North Atlantic, and (e) South Atlantic gyres. Each has a strong and
then driven back across the ocean basins from west narrow “western boundary current” and a weak and broad “eastern boundary current.”
to east. When these eastward-flowing currents reach
the eastern boundaries of the ocean basins, they are
largely deflected back toward the equator, where they
again come under the influence of the trade winds.
This completes a full cycle of surface currents that
rotate clockwise in the Northern Hemisphere and
counterclockwise in the Southern Hemisphere. These
large, circular-motion, wind-driven current systems
are known as gyres (fig. 7.16). In high southern
d
latitudes, no land separates the Atlantic, Pacific, and
b
Indian Oceans; here, the surface currents, driven by
the westerlies, continue around Earth in a circumpolar flow around Antarctica.
Geostrophic Flow
Because of Ekman transport, a portion of the winddriven surface w
­ ater is d­ eflected toward the center of
each of the large, circular ­current gyres just described
a
c
e
207
7.5 Ocean Surface Currents
Figure 7.17 Geostrophic flow (V) exists around a
gyre when Fc , the inward deflection force due to
the Coriolis effect, is balanced by Fg , the outwardacting pressure force created by the elevated water
and gravity. This example is of a clockwise gyre in
the Northern Hemisphere.
V
V
V
Fc
Fc
V
Fg
Fg
Fg
V
Fc
Fc
Fg
Fc
V
Fc
Lens of accumulated
surface water
V
Fc
Deeper water
Fg
difference in density between the surface water and the deeper
water is only about 1/1000 of the density difference between air
and water at the sea surface. The surface slope of the mound
increases as deflected water moves inward until the outward
pressure driving the water away from the gyre center equals
the Coriolis effect, acting to deflect the moving water into the
raised central mound. At this balance point, geostrophic flow
is said to exist, and no further deflection of the moving water
occurs. Instead, the currents flow smoothly around the gyre
parallel to its elevation contours. See figure 7.17 for a diagram
of this process.
7.5 Ocean Surface Currents
The currents that make up the large oceanic gyre systems and
other major currents have been given names and descriptions
based on their average positions. These are presented here ocean
by ocean and can be followed on figure 7.18. As you follow
these current paths, review their associations with the large gyre
systems and their overlying wind belts.
Pacific Ocean Currents
In the North Pacific Ocean, the northeast trade winds push
the water toward the west and northwest; this is the North
QUICK REVIEW
­Equatorial Current. The westerlies create the North Pacific
1. If a north wind blows across the sea surface, what
Current, or North Pacific Drift, moving from west to east.
direction does the surface current flow in the Northern
Note that the trade winds move the water away from Central
Hemisphere? in the Southern Hemisphere?
and South America and pile it up against Asia, whereas the
2. What is the direction of Ekman transport in question 1?
­westerlies move the water away from Asia and push it against
3. Why is the sea surface elevated in the interior of the
the west coast of North America. The water that accumulates in
major current gyres?
one area must flow toward areas from which the water has been
4. Why do gyres rotate in opposite directions in the two
removed. This movement forms two currents: the ­California
hemispheres?
Current, moving from north to south along the western coast
of North America, and the Kuroshio Current, moving from
south to north along the east coast of Japan. The Kuroshio
and California Currents are not winddriven currents; they provide continuity
of flow and complete a circular motion
East
centered around 30°N latitude. This cirGreenland
cular, clockwise flow of water is called
Norwegian
the North Pacific gyre. Other major North
Labrador
Pacific currents include the O
­ yashio
North Atlantic
Alaska
Oyashio Bering
Current, driven by the polar easterlies,
and the Alaska Current, fed by water
North Pacific
Canary
Gulf Stream
from the North Pacific Current and movCalifornia
Kuroshio
North Equatorial
ing in a counterclockwise gyre in the
North Equatorial
Somali
Gulf of Alaska. Little exchange of water
Equatorial CounterEquatorial CounterEquatorial Countercurrent
occurs through the Bering Strait between
current
current
South Equatorial
the North Pacific and the Arctic Ocean;
South Equatorial
South
Equatorial
no current exists that is comparable to the
Benguela
East Australian
Peru
Brazil
Agulhas
West
Australian
Figure 7.18 The long-term average flow
of the major wind-driven surface currents.
Antarctic Circumpolar
Antarctic Circumpolar
208
Chapter 7
Ocean Structure and Circulation
­ tlantic Ocean’s Norwegian Current, which moves warm ­water
A
to the Arctic Ocean.
In the South Pacific Ocean, the southeast trade winds
move the water to the left of the wind and westward, forming
the South Equatorial Current. The westerly winds push the
water to the east. The current formed, the Antarctic Circumpolar
Current, moves continuously around Earth. The tips of South
­America and Africa deflect a portion of this flow northward on
the east sides of the South Pacific and South Atlantic Oceans. As
in the North Pacific, continuity currents form between the South
Equatorial Current and the Antarctic Circumpolar Current. The
Peru ­Current, or Humbolt Current, flows from south to north
along the coast of South America, while the East Australia
Current can be seen moving weakly from north to south on the
west side of the ocean. These four currents form the counterclockwise South Pacific gyre.
The North Pacific and South Pacific gyres form on either
side of 5°N because the meteorological equator, or doldrums
belt, is displaced northward from the geographic equator (0°),
owing­to the unequal heating of the Northern and Southern
Hemispheres. Also between the North and South Equatorial
Currents, in the zone of the doldrums is a current moving in the
opposite direction, from west to east. This is a continuity current known as the Equatorial Countercurrent, which helps to
return accumulated surface water eastward across the Pacific.
Under the South Equatorial Current is a subsurface current
flowing from west to east called the Cromwell Current. This
cold-water ­continuity current also returns water accumulated in
the western Pacific.
Atlantic Ocean Currents
The North Atlantic westerly winds move the water eastward
as the North Atlantic Current, or North Atlantic Drift. The
northeast trade winds push the water to the west, forming the
North Equatorial Current. The north–south continuity currents
are the Gulf Stream, flowing northward along the coast of North
America, and the Canary Current, moving to the south on the
eastern side of the North Atlantic. The Gulf Stream is fed by the
Florida Current and the North Equatorial Current. The North
Atlantic gyre rotates clockwise. The polar easterlies provide the
driving force for the Labrador and East Greenland Currents,
which balance water flowing into the Arctic Ocean from the
Norwegian Current.
In the South Atlantic, the westerlies continue the West Wind
Drift. The southeast trade winds move the water to the west, but
the bulge of Brazil deflects part of the South Equatorial Current
northward into the Caribbean Sea and eventually into the Gulf of
Mexico, where it exits as the Florida Current and joins the Gulf
Stream. A portion of the South ­Equatorial Current moves south
of the Brazilian bulge along the western side of the South Atlantic to form the Brazil Current. The Benguela Current moves
­northward along the African coast. The South Atlantic gyre is
complete, and it rotates counter­clockwise.
Because much of the South Equatorial Current is deflected
across the equator, the Equatorial Countercurrent appears only
weakly in the eastern portion of the mid-Atlantic. The northward movement of South Atlantic surface water across the
equator ­results in a net flow of surface water from the Southern
­Hemisphere to the Northern Hemisphere. This flow is balanced
by a flow of water at depth from the Northern Hemisphere to the
Southern Hemisphere. This deep-water return flow is the North
Atlantic deep water. Again, the e­ quatorial currents are displaced
northward, although not as markedly as in the Pacific Ocean.
The Sargasso Sea marks the middle of an ocean gyre. It is
located in the central North Atlantic Ocean, and its boundaries
are the Gulf Stream on the west, the North Atlantic Current
to the north, the Canary Current on the east, and the North
­Equatorial Current to the south. The circular motion of the
gyre ­currents ­isolates a lens of clear, warm, downwelling water
1000 m (3000 ft) deep. The region is famous for the floating
mats of ­
Sargassum, a brown seaweed, stretching across its
surface. The extent of the floating seaweed frightened early
sailors, who told stories of ships imprisoned by the weed and
sea monsters lurking below the surface. Except for the floating
­Sargassum, with its rich and s­ pecialized ecological community,
the clear water is nearly a ­biological desert.
Indian Ocean Currents
The Indian Ocean is mainly a Southern Hemisphere ocean. The
southeast trade winds push the water to the west, creating the
South Equatorial Current. The Southern Hemisphere westerlies still move the water eastward in the West Wind Drift.
The gyre is completed by the West Australia Current moving
northward and the Agulhas Current moving southward along
the east coast of Africa. Because this is a Southern Hemisphere
ocean, the currents are deflected left of the wind direction, and
the gyre rotates counterclockwise. The northeast trade winds
in winter drive the North Equatorial Current to the west,
and the E
­ quatorial Countercurrent ­
returns water eastward
toward Australia. Again, these equatorial currents are displaced
­approximately 5°N. With the coming of the wet monsoon s­ eason
and its west winds, these currents are reduced. The strong
­seasonal monsoon effect controls the surface flow of the ­Northern
Hemisphere portion of the Indian Ocean. In the ­summer, the
winds blow the surface water eastward, and in the winter, they
blow it westward (fig. 7.19). This strong seasonal shift is ­unlike
anything found in the Atlantic or the Pacific Ocean.
Arctic Ocean Currents
The relentless drift of water and ice in the Arctic Ocean moves
in a large, clockwise gyre driven by the polar easterly winds.
This gyre is not centered on the North Pole, as early explorers
­expected, but is offset over the Canadian basin at 150°W and
80°N (fig. 7.20). Although the currents and the winds move the
ice slowly at 0.1 knot (2 mi/day), Arctic explorers trying to reach
the North Pole found that they traveled south with the drifting ice
and water at speeds almost equal to their difficult progress north.
The Arctic Ocean is supplied from the North Atlantic
by the Norwegian Current; some of this flow enters west of
Northeast Monsoon (January)
H
H
OMAN
INDIA
INDIA
SOMALIA
East African
Coastal
Current
North Equatorial Current
SUMATRA
Equatorial Countercurrent
MOZAMBIQUE
L
South Equatorial Current
MADAGASCAR
Southwest Monsoon (July)
L
L
OMAN
INDIA
East Arabian
Current
INDIA
SOMALIA
Somali Current
Southwest Monsoon Current
SUMATRA
Weak Equatorial
Countercurrent
MOZAMBIQUE
South Equatorial Current
H
MADAGASCAR
Figure 7.19 Indian Ocean monsoonal circulation. In the winter, high pressure over the continent creates dry monsoon winds roughly out of the
northeast that drive water in the Northern Hemisphere portion of the ocean to the west. In the summer, low pressure over the continent creates
wet monsoon winds roughly out of the southwest that reverse the surface currents in the Northern Hemisphere portion of the ocean.
W180°E
ALASKA
(U.S.)
CANADA
SIBERIA
90°W
90°N
90°E
.
C.
80°
GREENLAND
Spitsbergen
60
°
ICELAND
Figure 7.20 The circulation in the Arctic Ocean is driven by
NORWAY
.
70°
g ian C
reenland C .
ast G
W0°E
No
rwe
E
dor C
L abra
nd
la
en
re
tG
es
W
Spitsbergen, but most flows along the coast of Norway and
moves eastward along the Siberian coast into the Chukchi Sea.
A small inflow of water entering the Arctic through the Bering
Strait brings water from the Bering Sea to join the eastward
flow along Siberia and the large Arctic gyre. The western side
of the gyre crosses the center of the Arctic Ocean to split north
of Greenland. Here, the larger flow forms the East Greenland
Current flowing south and taking Arctic Ocean water into the
North Atlantic. The lesser flow moves along the west side of
­Greenland to join the Labrador Current and move south along
the Canadian coast.
Outflow from Siberian rivers is caught in the eastward
flow of water and ice along Siberia. Eventually, this discharge
joins the gyre, distributing sediments and pollutants throughout the Arctic.
FINLAND
SWEDEN
the polar easterlies, which produce a large, clockwise gyre.
Water enters the Arctic Ocean from the North Atlantic by way
of the Norwegian Current and exits to the Atlantic by the East
Greenland Current and the Labrador Current.
209
210
Chapter 7
Ocean Structure and Circulation
Figure 7.21 The Antarctic Circumpolar
20°S
Current, also known as the West Wind
Drift, provides exchange of water among
the Pacific, Atlantic, and Indian Ocean
basins.
Antarctic Circumpolar Current boundary
40°S
Subantarctic front
60°S
Antarctic circle
Antarctic front
Ross Sea gyre
Weddell
gyre
80°S
160°E
160°W
120°W
80°W
40°W
0°
40°E
Antarctic Currents
80°E
120°E
Throughflow transports large amounts of relatively warm and fresh
water to the Indian Ocean.
The volume of water moving in the Indonesian Throughflow can vary with changing conditions in the Pacific. In particular, the amount of flow of Pacific water through the Indonesian
archipelago is less during an El Niño event.
At high southern latitudes, the tips of South America and Africa
deflect some of the eastward surface flow back toward the equator,
completing the Southern Hemisphere gyres. However, much of this
eastward-moving water continues its journey, circling Antarctica
as the Antarctic Circumpolar Current (AAC), also known as
the West Wind Drift (fig. 7.21). The AAC is the largest windQUICK REVIEW
driven current on Earth and the only ocean current to travel all the
way around the planet. It is also an important current because it
1. Follow the major ocean gyres, and identify the major
provides a mechanism for sharing, or mixing, water among the
currents in each.
Pacific, Atlantic, and Indian Ocean basins. The AAC circles about
2. Why are the equatorial countercurrents located under
24,000 km (15,000 mi) around Antarctica. It extends from the sea
the doldrums?
surface to depths of 2000–4000 m (6500–13,000 ft) and can be up
3. Where does water enter and exit the Arctic Ocean?
to 2000 km (~1200 mi) wide. Because of its huge cross-sectional
4. Why is it possible for the Antarctic Circumpolar
area, it transports an enormous volume of water despite a relatively
Current to flow all the way around Antarctica without
slow average speed of less than 20 cm s-1. The AAC has very
interruption?
consistent characteristics. Generally, its path doesn’t differ much
from that shown in figure 7.21, but in the
central and western Pacific, it can shift by
North Pacific Inflow
more than 10° of latitude between summer
Pacific Ocean
South Pacific Inflow
and winter. Two clockwise gyres associated
Java Sea surface water
(seasonally varying)
with the current are found in the Ross Sea
Mindanao Eddy
and the Weddell Sea.
Sulawesi
Halmahera Eddy
The Indonesian
Throughflow
The western border of the South Pacific
Ocean basin is not a solid landmass.
Rather, it consists of a number of islands
of varying size that constitute the Indonesian archipelago. When the Pacific
Ocean’s westward-moving Equatorial
Currents reach the western side of the ocean
basin, some of their water is deflected away
from the equator as the Kuroshio and East
Australia Currents. The remainder, known
as the I­ ndonesian Throughflow, continues
westward as a complicated series of currents
winding through the Indonesian archipelago
(fig. 7.22). Because the water in the western equatorial Pacific Ocean has a higher
temperature and lower salinity than the
water in the Indian Ocean, the Indonesian
Borneo
Makassar
Straight
Sulawesi
Java Sea
Java
Lombok
Maluku
La Niña
Dewakang
Sill
Halmahera
Seram
Banda
Flores
Timor
Timor
Indian Ocean
Average flow of Indonesian Throughflow = 10 Sverdrups
Australia
Figure 7.22 The complicated current system through the Indonesian archipelago that
comprises the Indonesian Throughflow, carrying Pacific water into the Indian Ocean.
7.6 Current Characteristics
211
Major ocean currents transport enormous volumes of water. A convenient unit to report transport volume is the Sverdrup (Sv) (named
after Harald Sverdrup, a leading oceanographer of the last century
and former director of the Scripps Institution of Oceanography). A
Sverdrup equals 1 million cubic meters (~3.5 × 107 ft3) per second.
The transport rate of fresh water in all of the world’s rivers into the
ocean is about 1 Sv. Transport rates of ocean currents are difficult to
measure accurately and can vary by both location in the current and
time of the year. The Gulf Stream transports about 30 Sv passing
through the Strait of Florida as the Florida Current. This increases
steadily as it moves north along the coast until it transports about
80 Sv near Cape Hatteras. The transport of the Gulf Stream continues to increase downstream of Cape Hatteras at a rate of 8 Sv
every 100 km, reaching a maximum transport of about 150 Sv at
55°W. The downstream increase in transport between Cape Hatteras
and 55°W is thought to be caused by increased velocities in the
deep waters of the Gulf Stream. The current transports a maximum
amount of water in the fall and a minimum in the spring.
The largest current in the oceans, the Antarctic Circumpolar
Current, transports an estimated 173 Sv of water.
displacement of the mounds of water in the North Pacific and
North Atlantic gyres is particularly easy to see in the satellite sea
surface elevation data shown in figure 7.24. The geostrophic currents that flow around the mound have very different properties on
the two sides. Currents flowing on the western side of these gyres,
where the slope is steeper, tend to be much faster, deeper, and narrower. Currents flowing on the eastern side of these gyres, where
the slope is gentler, tend to be slower, shallower, and broader.
This phenomenon is known as the western intensification of
currents. The Gulf Stream and Kuroshio ­Currents are faster and
narrower than the Canary and ­California Currents, although both
the eastern and western boundary currents transport about the
same amount of water to preserve continuity of flow around the
gyres. Western intensification of currents traveling from low to
high latitudes is related to (1) the eastward turning of Earth, (2)
the increase in the C
­ oriolis effect with increasing latitude, (3) the
changing strength and direction of the east–west wind field (trade
winds and westerlies) with latitude, and (4) the ­friction ­between
landmasses and ocean water currents. These factors cause a compression of the currents on the western side of the oceans, where
water is moving from lower to higher ­latitudes. This compression
requires that the current speed ­increase to transport the water
circulating about the gyre. On the eastern side of the gyre, where
currents are moving from higher to lower latitudes, the currents
are stretched in the east–west direction. Here, the current’s speed
is reduced, but it still transports the required volume of water.
The changing speed of flow around the gyre causes the Coriolis
effect to vary. Where the current speed is high, the Coriolis effect is large, and a steeper surface slope is r­equired to create a
geostrophic flow balance.
Fast-flowing, western-boundary currents move warm equatorial surface water to higher latitudes. Both the Gulf Stream
and the Kuroshio Current bring heat from equatorial latitudes to
moderate the climates of Japan and northern Asia (in the case of
the Kuroshio) and the British Isles and northern ­Europe (in the
case of the Gulf Stream, via the North Atlantic and Norwegian
Currents). Western intensification is obscured in the South Pacific and South Atlantic, because both Africa and South America
deflect portions of the West Wind Drift and ­create strong currents on the eastern side of these oceans. The deflection of w
­ ater
from the Atlantic’s South Equatorial Current to the Northern
Hemisphere r­emoves water from the South ­Atlantic gyre and
strengthens the Gulf Stream. The flow of s­ urface water from the
Pacific to the Indian Ocean through the islands of I­ ndonesia also
helps to prevent the development of strongly flowing currents on
the west side of Southern Hemisphere oceans.
Western Intensification
QUICK REVIEW
7.6 Current Characteristics
Current Speed
Wind-driven open-ocean surface currents move at speeds that are
about 1/100 of the wind speed measured 10 m (30 ft) above the sea
surface. The water moves between 0.25 and 1.0 knot, or 0.1 to 0.5 m
(0.3 to 1.5 ft) per second. Currents flow faster when a large volume
of water is forced to flow through a narrow gap. For example, the
North and part of the South Atlantic Equatorial Currents flow into
the Caribbean Sea, then flow into the Gulf of Mexico, and finally
exit to the North Atlantic as the Florida Current through the narrow
gap between Florida and Cuba. The Florida Current’s speed may
exceed 3 knots, or 1.5 m (5 ft) per second. Once into the Atlantic
Ocean, this current turns north and becomes the Gulf Stream.
The flow is distributed over the width and depth of the current.
When the cross-sectional area of the current expands, the current slows down; when the cross-sectional area decreases, the
current speeds up. Speed of flow may not be directly related to
surface wind speed but can be affected by the depth and width
of the ­current as determined by land barriers, by the presence of
another current, or by the rotation of Earth, as explained in the
“Western Intensification” section of this chapter.
Current Volume Transport
When we look at the shape of the actual mounds of water created by Ekman transport in the major subtropical gyres (review
figs. 7.16 and 7.17), it is clear that the peaks of the mounds are
not in the center of the gyres. The highest point of the mounds is
displaced to the western side of the gyres (fig. 7.23). This creates
a more gently sloping sea surface, over a longer distance, on the
eastern side of the mound and a more steeply sloping surface,
over a shorter distance, on the western side (fig. 7.23a). The
1. Explain in general terms differences in the
characteristics of the Kuroshio and California Currents.
2. What forces balance to create geostrophic currents?
3. What do surface currents transport?
4. How does changing the width of a current affect its
speed?
5. What is western intensification?
212
Chapter 7
Ocean Structure and Circulation
Northern Hemisphere Subtropical Gyre
Path of ideal
geostrophic flow
Raised sea surface
Corio
li
s
Grav
ity
Path of actual
geostrophic flow
Rotation
of Earth
Geostrophic flow
Eastern side
of ocean basin
Western side
of ocean basin
(a) Profile view
60°N Latitude
lines
flow
c
i
h
op
rly
te
str
o
es
e
Eastern side
of ocean basin
W
W
G
Western side
of ocean basin
d
in
s
Wide and weak
flow to south
Center
30°N Latitude
ind
s
of gyre
Rotation
of Earth
W
Tr
ad
e
Narrow and strong
flow to north
0° Latitude
(b) Map view
Figure 7.23 (a) A cross-sectional view of a subtropical gyre illustrating the accumulation of water in the center, elevating the sea surface by as
much as 2 m (6.6 ft). The peak of the mound is displaced to the west. Gravity and the Coriolis effect balance to create geostrophic flow around
the hill. Friction makes the current gradually spiral downslope. (b) A map view of the same gyre showing the concentration of the currents on the
western side of the gyre compared to the currents on the eastern side, creating the phenomenon called western intensification.
Figure 7.24 Variations in the average
height of the sea surface are shown in
this image constructed from data taken
in 1992 by the TOPEX/Poseidon satellite.
Red indicates the highest elevations,
whereas green and blue indicate the
lowest. Note that the highest elevations
within the subtropical gyres are found on
the western sides of the ocean basins.
These are particularly easy to see in the
western Pacific and western Atlantic in
the Northern Hemisphere. Source of data:
National Oceanic and Atmospheric Administration
(NOAA).
Courtesy of NASA/JPL-Caltech
7.7 Eddies
213
indentations pinch off, they become counterclockwise-rotating,
cold-water eddies that are displaced eastward through the Gulf
When a narrow, fast-moving current moves into or through
Stream and into the warm-water core of the gyre. Bulges at
slower-moving ­water, the force of its flow displaces the quieter
the western edge of the Gulf Stream are filled with warm Gulf
water and captures additional ­water as it does so. The current
Stream water. When these bulges are cut off, they become warmoscillates and d­ evelops waves along its boundary that are known
water, clockwise-rotating eddies drifting into cold water to the
as meanders. These ­meanders break off to form eddies, or pockets
west and north of the Gulf Stream. ­Follow these processes in
of ­water moving with a circular motion; eddies take with them
figures­7.25 and 7.26. These eddies may maintain their physical
­energy of motion from the main flow and gradually dissipate this
identity for weeks as they wander about the oceans; they are
energy through friction. ­Eddies also act to mix and blend water.
especially numerous in the area north of the ­Sargasso Sea. CurAs the Gulf Stream moves away from the North American
rent meanders and eddy f­ ormation produce surface flow patterns
coast, it is likely to ­develop a meandering path. The western edge
that differ markedly from the uniform current flows shown on
of the Gulf Stream develops oscillations, and the indentations are
­current charts. These charts show average current flow, not daily
filled by cold water from the Labrador Current side. When these
or weekly variations.
Large and small eddies generated
by horizontal flows or currents exist in
Cold
all parts of the oceans; these eddies are
water
W
of varying sizes, ranging from tens to
several hundred kilometers in ­diameter.
Each eddy contains water with specific
chemical and physical properties and
maintains its identity and ­rotational inC
W
ertia as it wanders through the oceans.
Eddies may appear at the sea surface or
be embedded in waters at any depth.
C
Eddies rotate in a clockwise or counterclockwise direction. They stir the ocean
until they gradually dissipate because
of fluid friction, losing their chemical
and thermal identity and their energy of
motion. By testing the water properties
of an eddy, oceanographers are able to
C
determine the eddy’s place of origin.
0
25
50 Miles
Small surface eddies encountered 800 km
0
25 50 Kilometers
Cape
(500 mi) southeast of Cape Hatteras in
Hatteras
the North Atlantic have been found with
water properties of the eastern AtlanWarm
tic near Gibraltar, more than 4000 km
water
(2500 mi) away. Eddies from the Strait
of Gibraltar are formed in the salty water
of the Mediterranean as it sinks and spreads
out into the Atlantic 500 to 1000 m (1600
Cold
Cold
Cold
to 3300 ft) down; these eddies have been
nicknamed “Meddies.” Deep-water e­ ddies
near Cape Hatteras may come from the
C
eastern and western Atlantic, the CaribC
W
W
bean, or Iceland. Researchers estimate that
C
W
some of these eddies are several years
old; age determination is based on drift
Warm
Warm
rates, distance from source, and biological
(a)
(b)
(c)
(d)
­consumption of oxygen.
Figure 7.25 The western boundary of the Gulf Stream is defined by sharp changes in
The rotational water speed in the
current velocity and direction. Meanders form at this boundary after the Gulf Stream leaves
large
eddies that form at the western
the U.S. coast at Cape Hatteras. The amplitude of the meanders increases as they move
boundary
of the Gulf Stream is about
downstream (a and b). In time, the current flow pinches off the meander (c). The current
0.51
m/s
(1
knot), but because of the
boundary re-forms, and isolated rotating cells of warm water (W) wander into the cold water,
water’s
density,
the force of the flow is
whereas cells of cold water (C) drift through the Gulf Stream into the warm water (d).
ns
ide
Easte
r
Western boundar
y
7.7 Eddies
214
Chapter 7
Ocean Structure and Circulation
Figure 7.26 A composite satellite image of the sea surface
reveals the warm (orange and yellow) and cold (green and blue)
eddies that form along the Gulf Stream. (Reddish blue areas at the
top are the coldest waters.) These eddies may stir the water column
right down to the ocean floor, kicking up blizzards of sediment.
Figure 7.27 Space shuttle view. Sunlight reflected off the
Mediterranean reveals spiral eddies; their effects on climate are
being monitored. Dimensions of this image are 500 km × 500 km
(310 mi × 310 mi).
©GRANGER/GRANGER
Courtesy Otis Brown, University of Miami/RSMAS
similar to that generated by a 35-knot wind. The ­diameter of the
eddies may be as much as 325 km (200 mi), and their effect may
reach to the sea floor. At the sea floor, the rotation rate is zero;
therefore, a few meters above the bottom the speed of r­otation
diminishes very rapidly, and considerable turbulence is generated
as the energy of the eddy is dissipated. These eddies are similar
in many ways to the winds rotating about atmospheric pressure
cells, and they are sometimes called abyssal storms. As the eddies wander through the oceans, they stir up bottom sediments,
producing ripples and sand waves in their wakes; they also mix
the water, creating homogeneous water properties over large areas.
Eventually, the eddies lose their energy to turbulence and blend
into the surrounding water.
Eddies constantly form, migrate, and dissipate at all
depths. Eddy motion is superimposed on the mean flow of the
oceans. To understand the role of eddies in mixing the oceans,
we need more data and better tracking of eddy size, position,
and rate of dissipation. Satellites are important tools for detecting surface eddies because they can precisely measure temperature, increased­elevation, and light reflection of the sea surface
(fig. 7.27). Deep-water eddies are monitored by using instruments designed to float at a mid-depth density layer. The instruments are caught up in the eddies, moving with them and
sending out acoustic signals that are monitored through the
sofar channel. In this way, the rate of deep-water eddy formation, the numbers of major eddies, their movements, and their
life spans can be observed.
QUICK REVIEW
1. What are eddies, and where do they form?
2. How are parcels of warm water transferred across the
Gulf Stream and embedded in regions of cold water?
3. How can we monitor the movement of deep-water
eddies?
7.8 Convergence and Divergence
Changes in density and the accompanying concepts of upwelling,
downwelling, convergence, divergence, and continuity of flow
were introduced in section 7.2. The convergence and divergence
zones discussed here are the product of wind-driven surface
currents that produce large-scale areas of convergence and divergence at the sea ­surface. For ­example, convergence zones are at
the centers of the large oceanic gyres, and when wind-driven
surface currents collide or are forced against landmasses, they produce convergences. When surface currents move away from each
other or away from a landmass, they produce a surface ­divergence.
Upwellings and downwellings of this type are nearly permanent
but do react to seasonal changes in Earth’s surface winds.
Langmuir Cells
A strong wind blowing across the sea surface often causes
streaks of foam and surface debris that are seen trailing off
7.8 Convergence and Divergence
215
the windrows varies between 5 and 50 m (16 and 160 ft), and
the rows are closer together if the thermocline depth is shallow. The vertical extent of these Langmuir cells is 4 to 10 m
(12 to 30 ft). The distance between rows becomes larger at higher
wind speeds. Langmuir cells are not long-lasting but help to mix
the surface waters and organize the distribution of suspended
organic matter in convergence and divergence zones. Sinking particles and organisms tend to congregate at depth in the ­regions of
rising currents, while floating particles and organisms accumulate
at the surface where currents ­converge and descend.
Permanent Zones
g (d
ellin
Upw
Win
d
iver
g
enc
e)
Areas of surface-water convergence and divergence occur in association with high- and low-pressure systems over the ocean. In the
Northern Hemisphere, a low-pressure system produces cyclonic
Figure 7.28 Windrows in the water of the Potomac River
winds that rotate counterclockwise. These winds result in Ekman
created by Langmuir circulation.
transport of water away from the center of the system, creating
©Keith A. Sverdrup
an area of surface divergence. This surface divergence produces
in the direction the wind is blowing. These streaks are called
upwelling of deep water (fig. 7.30a). High-pressure systems in
windrows and may be 100 m (330 ft) in length. Windrows mark
the Northern Hemisphere produce anticyclonic winds that rotate
the convergence zones of shallow circulation cells known as
clockwise. These winds drive Ekman transport of water toward
Langmuir cells (fig. 7.28). These cells are composed of paired
the center of the system, creating an area of surface convergence.
right- and left-handed helixes (fig. 7.29). The spacing between
This surface convergence produces downwelling of surface water (fig. 7.30b).
w
w
)
ro
ro
There are also zones of permanent
e
d
d
c
n
n
en
wi )
wi )
convergence
and divergence, and associg
s
s
r
i
i
e
e
br nc
br nc
ive
ated downwelling and upwelling, related
de rge
de rge
(d
d e
d e
g
lin
an nv
an nv
to the large ocean gyres and the Ekman
el
m (co
m (co
w
a
a
p
o
o
transport occurring in them (figs. 7.13
F
F
U
and 7.31). There are five major zones of
con­vergence: the tropical convergence
at the equator and the two subtropical
­convergences at approximately 30° to
40°N and S. These convergences mark the
centers of the large ocean gyres. The Arctic and Antarctic convergences are found
at about 50°N and S. Surface convergence
zones are regions of downwelling. These
areas are low in nutrients and biological productivity. There are three major
divergence zones: the two tropical divergences and the Antarctic divergence.
Upwelling associated with divergences
delivers nutrients to the surface waters to
supply the food chains that support the
anchovy and tropical tuna fisheries and
the richly productive w
­ aters of Antarctica.
Surface
Figure 7.32 illustrates global patterns of
wind-induced upwelling and downwelling.
The equatorial upwelling is particularly easy
Thermocline
to see in this image as well as slower rates of
Accumulating sinking material
upwelling around Antarctica and along the
eastern coasts of continents.
Figure 7.29 Langmuir cells are helixes of near-surface, wind-driven water motion.
In coastal areas where trade winds
The cells extend downwind to form windrows of surface debris along convergences.
At the same time, sinking materials are swept into zones of upwelling water.
move the surface waters away from the
216
Chapter 7
Ocean Structure and Circulation
Northern Hemisphere
Ekman transport
Surface
convergence
Surface
divergence
Surface current
Wind
Upwelling
Downwelling
Thermocline
Thermocline
Cyclonic wind
Anticyclonic wind
(a)
(b)
Figure 7.30 Surface convergence and surface divergence occur in association with low- and high-pressure systems. (a) In the Northern
Hemisphere, cyclonic winds (low-pressure systems) result in Ekman transport of water away from the center of the low pressure, producing
surface divergence and upwelling. (b) Anticyclonic winds (high-pressure systems) result in Ekman transport of water toward the center of
the high pressure, producing surface convergence and downwelling.
60°N
Arctic convergence
o
Subtr
e
l
ica
Subtrop enc
converg
ce
gen
er
nv
o
c
ic
nce
Arct
rge
nve
o
c
l
pica
Tropical diverg
ence
Tropical convergence
Tropical co
nverg
enc
Tropical divergence
o
c
cal
ropi
Subt
40°N
cal
Tropi nce
e
diverg
Tr o
dive pical
rgenc
e
al convergence
Tropic
e
vergenc
Tropical di
e
e nc
erg
nv
Subtr
20°N
e
0°
Tropical diverge
nce
opical co
nvergence
20°S
Subtropical converg
ence
40°S
ctic convergence
Antar
Antarct
Anta
rctic
150°E
180°
ic convergence
60°S
Antarctic divergence
divergence
150°W
120°W
90°W
60°W
30°W
0°
30°E
60°E
90°E
120°E
Figure 7.31 The principal zones of open-ocean surface convergence and divergence associated with wind-driven and
thermohaline circulation.
western side of continents, upwelling occurs nearly continuously
throughout the year. For example, the trade winds drive upwellings off the west coasts of Africa and South America that are very
productive and yield large fish catches.
Seasonal Zones
Off the west coast of North America, downwelling and upwelling
occur seasonally as the Northern Hemisphere temperate wind
pattern changes from southerly in winter to northerly in summer.
The downwelling and upwelling occur because of the change
in direction of Ekman transport (fig. 7.33). R
­ emember that the
wind-driven Ekman transport moves at an angle of 90° to the
right or left of the wind direction, depending on the hemisphere.
Along this coast, the average wind blows from the north in the
summer, and the net movement of the water is to the west, or 90° to
the right of the wind. This action results in the offshore transport of
the surface water and the upwelling of deeper w
­ ater along the coast
7.8 Convergence and Divergence
217
July Global Wind-induced Upwelling (cm/day)
60
60
40
40
20
20
EQ
EQ
20
20
40
40
60
60
0
20
40
60
80 100 120 140
160 180 200 220 240 260 280 300 320 340 360
–50 –45 –40 –35 –30 –25 –20 –15 –10 –5
0
5
10
15
20 25
30 35 40 45 50
Figure 7.32 Equatorial upwelling occurs in all of the ocean basins but is particularly well developed in the Pacific. Positive values
indicate upwelling associated with surface divergence. Negative values indicate downwelling associated with surface divergence.
Courtesy of Liesen Xie and William Hsieh, University of British Columbia
Sea surface
slopes upwards
offshore
Wind from north
Upwelled water from
depths between 100 and 300 m
Ekman transport causes
water to move offshore
Upwelling
(a)
Sea surface
slopes upwards
onshore
Wind from south
Ekman transport causes
water to move onshore
to replace it. The upwelling zone is evident from central California
to Vancouver Island (fig. 7.34). In w
­ inter, the winds blow from the
south; the wind-driven surface w
­ aters move to the east, onshore
against the coast; and downwelling occurs. The summer upwelling pattern is what produces the band of cold coastal water at San
­Francisco that helps cause the frequent summer fogs. ­Because of
the lack of land at the middle latitudes in the Southern Hemisphere,
this type of seasonal ­upwelling is less common.
Convergence and divergence of surface currents result not
only in upwellings and downwellings but also in the mixing of
water from different geographic areas. Waters carried by the
surface currents converge, share properties due to mixing, and
form new water mixtures with specific ranges of temperature
and salinity, which then sink to their appropriate density levels.
After sinking, the water moves horizontally, blending and sharing its properties with adjacent water, and eventually it rises to
the surface at a new location. This process, known as caballing,
is discussed as thermohaline circulation in section 7.2. Thermohaline circulation and wind-driven surface currents are closely
related. The connections are so close among formation of water
Figure 7.33 Winds blowing along the coastline can generate
Downwelling
(b)
Ekman transport of water toward, or away from, the coast. Along
the northwest coast of North America (a) northerly summer winds
transport water away from the coast, depressing the sea surface
and producing upwelling, and (b) southerly winter winds transport
water toward the coast, elevating the sea surface and producing
downwelling.
218
Chapter 7
Ocean Structure and Circulation
Hemisphere around Antarctica; this is the Antarctic bottom water
(AABW) discussed earlier in the chapter. AABW is the densest
water in the oceans. Deep water eventually wells up in the centers
of the ocean basins and moves toward regions of dense, sinking
surface water. This transport of water through the entire conveyor
belt can take 1000 years or more. As water is transported, so is heat.
Tropical water, moving near the surface toward the poles, carries
heat with it. As this water moves to higher latitudes, it releases its
heat to the environment. Seawater temperature beneath the thermocline is uniformly cold everywhere in the oceans because all deep
water has its origin at polar latitudes where low environmental
temperatures contribute to high seawater density.
7.10 Changing Circulation Patterns
North Pacific Oscillations
Figure 7.34 Coastal upwelling (shades of purple) appears as
patches of cold water extending from the coast of California.
Courtesy of Mark Abbott/NASA
mixtures, upwelling and downwelling, and convergence and
divergence that it is difficult to assign a priority of importance
to one process over another. Changes in these flows and events
that might cause such changes are discussed in the next section.
7.9 The Great Ocean Conveyor Belt
Large-scale climate variability in the North Pacific exhibits characteristics similar to those in tropical and equatorial regions due to
El Niño–Southern Oscillation (ENSO) events. This North Pacific
variability is called the Pacific Decadal Oscillation (PDO). There
are two major differences between PDO and ENSO events. First,
PDO events persist for decades, while typical ENSO events persist
for six to eighteen months on average. Second, the effects of PDO
events are greatest in the Northeast Pacific and have less influence
in the tropics, while the opposite is true for ENSO events. There
are two phases of PDO events––a “cold, wet” phase characterized
by southerly flowing currents and a “warm, dry” phase characterized by northerly flowing currents (fig. 7.36). This north–south
current oscillation has two general periodicities, one from fifteen
to twenty-five years and one from fifty to seventy years. Historic
PDO events can be inferred from their impact on the growth of
western juniper trees in Oregon. These trees show wide rings
during PDO warm, dry phases when currents are displaced to
It was first suggested in the 1980s that the movement of heat around
the globe in ocean currents and water masses could be modeled
as a thermohaline conveyor belt (fig. 7.35). This model is not
meant to be a literal map of actual warm and cold currents—it is
a representation of the overall effect of
the mass movement of warm and cold Figure 7.35 The Great Ocean Conveyor Belt of thermohaline circulation within and between
water on the vertical circulation in the oceans circulates at the surface (red arrows) and at depth (blue arrows). Dark-blue dots mark
ocean. We can think of the overall driv- regions of deep-water formation.
ing mechanism of this conveyor belt
Surface flow
as the formation and sinking of dense
Deep flow
water at polar latitudes. The formation of deep-water masses depends on
the production of dense surface water,
either by significant cooling and/or an
increase in salinity. In the Northern
Hemisphere, deep water is created in
the North Atlantic; this is the North
Atlantic deep water (NADW) discussed
earlier in the chapter. NADW is dense
because it is cold and has a relatively
high salinity. NADW sinks and moves
slowly southward, extending all the way
to the southern tip of Africa, where it
joins the deep water circling Antarctica. From there it can move into the
Atlantic and Indian Ocean basins. Deep
water is also created in the Southern
219
7.10 Changing Circulation Patterns
1978–79
GREENLAND
Baffin
Bay
1981–82
1968
56°N
1969–70
W
152°
6
12
°W
Hudson
Bay
ICELAND
1976
Labrador
Sea
1975
CANADA
49°N
Ocean
weather
station
“PAPA”
139
°W
Area of northerly flowing currents
Area of southerly flowing currents
1971–72
UNITED
STATES
1977–78
1974
ATLANTIC
OCEAN
EUROPE
AFRICA
Figure 7.36 North–south current shifts in the Northeast Pacific
Figure 7.37 Purple arrows follow the path of the North Atlantic
are associated with changes in atmospheric pressure, winds,
precipitation, and water temperature. This phenomenon is known
as the Pacific Decadal Oscillation (PDO).
cool pool. Red arrows show warm-water flow from the Gulf
Stream.
the north and narrow rings during PDO cold, wet phases when
currents are displaced to the south. Tree-ring data cover many
more years than oceanographic and meteorological data, and from
the tree rings, it is calculated that 34 PDO events have occurred
since the time of Columbus. The PDO pattern is presently warm
and dry with a northerly current flow that has not changed since
1967––one of the longest periods without a reversal that has been
documented during the last 500 years.
The PDO impacts marine ecosystems in the Northeast
Pacific because it affects coastal surface temperatures from
California to Alaska. During PDO warm, dry periods associated with northerly flowing currents, there is enhanced coastal
ocean biological productivity in Alaskan waters and a drop in
biological productivity off the coast of the contiguous United
States. This pattern is reversed during PDO cold, wet periods
associated with southerly flowing currents.
pieces of the puzzle that are associated with NAO but have not yet
found the factors that control the system. A counterclockwise wind
circulation centered over Iceland and a high-pressure, clockwise
circulation residing near the Azores are the usual situation. If the
air-pressure differences between these two locations are large, then
strong westerly winds supply Europe with heat from the North
Atlantic Current. If the air-­pressure difference decreases, then
weaker-than-normal westerlies drive less warm water into the Norwegian Current and less heat is delivered to Europe. The periods of
1950–71 and 1976–80 were recognized as prolonged cool periods
for Europe. The winds may also drive cold, low-salinity water from
the Arctic into the area where North Atlantic deep water is formed.
The i­nflux of this water may cause a small-scale reduction in the
­formation of North Atlantic deep water and an accompanying
­reduction of warm surface water moving northward. How much
freshening is required to change ocean circulation? This question is
unanswered, but if the current melting of Arctic sea ice continues,
more information may soon be available.
The surface waters of the North Atlantic Ocean have also
undergone a series of warmer and cooler phases with a periodicity of about twenty to forty years. The maximum difference in
surface water temperature is on the order of 0.5°C (0.9°F). This
phenomenon is known as the Atlantic Multidecadal Oscillation
(AMO). The warm phase, which we are currently in, is associated
with an increase in the number of Atlantic hurricanes by at least
a factor of two. The AMO also impacts North American and European summer climates. During warm phases of the AMO, there
are more frequent or prolonged droughts in the U.S. Midwest and
Southwest, while Florida and the Pacific Northwest experience
higher rainfall. Although the existence of the AMO is very well
documented, its cause is still a subject of research.
Natural cycles operate at a variety of time scales and magnitudes, and they interact and react in ways we do not fully
comprehend. In the enormously complex interactions that they
are observing, oceanographers and meteorologists are beginning
North Atlantic Oscillations
A large pool of cold, low-salinity surface water appeared off
Greenland, north of Iceland, in 1968. It was about 0.5‰ less
salty and 1° and 2°C colder than usual. Within two years, this
cool pool had moved west into the Labrador Sea off eastern
­Canada; then it crossed the Atlantic and, in the mid-1970s,
moved north into the Norwegian Sea. It had returned to its place
of origin by the early 1980s. Follow the path of this pool in
figure­7.37. During this period, harsh winters plagued Europe,
and the entire Northern Hemisphere had cooler-than-average
temperatures for more than ten years.
Recent studies have further investigated the North Atlantic,
seeking to improve our ability to predict climate. The new analysis
shows pools of warm and cool surface water that circle the North
Atlantic; the pools seem to have life spans of four to ten years.
Because of the alternation of climatic conditions, this is referred to
as the North Atlantic Oscillation (NAO). Investigators are finding
220
Chapter 7
Ocean Structure and Circulation
to identify the strands that united the ocean-atmosphere and the
world current systems. Scientists use mathematical equations
­assembled into a model to describe interactions between the
oceans and the atmosphere. Supercomputers are needed to make
the millions of calculations that a model requires to predict
changes in environmental conditions over a given time period.
This pro­cess is repeated many times to predict conditions. Models are verified and tuned by adjusting the equations so that the
predicted model changes closely agree with changes that have
been observed in the past. Modeling is nearly as much an art as
a science. Sometimes predictions work; sometimes they don’t. It
depends on how well the model has been conceived, how much
data are available, and the capability of the computers to process
the data. The model may be constructed to approximate the
whole ocean-atmosphere­system or to apply only to part of Earth.
Eventually, long-term predictions concerning the oceans and the
atmosphere will become possible; how we use such information
and whether it will benefit ocean resources are not yet clear.
QUICK REVIEW
1. Describe the Pacific Decadal Oscillation and its causes.
2. Describe the North Atlantic Oscillation.
3. What effect does the North Atlantic Oscillation have on
Europe’s climate?
7.11 Measuring the Currents
Direct measurements of currents fall into two groups: (1) those
that follow a parcel of the moving water and (2) those that measure the speed and direction of the water as it passes a fixed
point. Moving waters may be followed with buoys designed to
float at predetermined depths. These buoys ­signal their positions
acoustically to a research vessel or shore station; their paths are
followed and their speed and displacement due to the current are
calculated. Autonomous profilers also measure currents. Surface
Figure 7.38 Surface buoys (red) carrying instruments are
deployed from the research vessel Thomas G. Thompson.
Courtesy of Will Peterson, School of Oceanography, University of Washington
­ ater may be labeled with buoys or with dye that can be phow
tographed from the air. Buoy positions may also be tracked by
satellites using GPS. A series of pictures or position fixes may
be used to calculate the speed and direction of a current from the
buoys’ drift rates. Buoys can be instrumented to measure other
water properties such as temperature and salinity (fig. 7.38).
A variation of this technique uses drift bottles. Thousands
of sealed bottles, each containing a postcard, are released at a
known position. When the bottles are washed ashore, the finders are requested to record the time and location of the find and
return the card. In this case, only the release point, the recovery
points, and the elapsed time are known; the actual path of motion is assumed. See the Diving In box titled “Ocean Drifters.”
Sensors used to measure current speed and direction at
fixed locations are called current meters. The current meters
used by oceanographers over the last twenty years include a
rotor to measure speed and a vane to measure direction of flow
(fig. 7.39a). If the current meter is lowered from a stationary
vessel, the measurements can be returned to the ship by a cable
Figure 7.39 (a) An internally recording
Aanderaa current meter. The vane
orients the meter to the current while
the rotor determines current speed.
(b) A Doppler current meter sends out
sound pulses in four directions. The
frequency shift of the returning echoes
allows the detection of the current.
These meters are also equipped with
salinity-temperature-depth sensors
as well as instruments for measuring
water turbidity and oxygen. The data
may be stored internally and collected
at another time.
(a)
(b)
Courtesy of Kathy Newell, School of Oceanography,
University of Washington
Diving
in
Ocean Drifters
ns
tr e
am
rrent
In May 1990, a severe storm in the North Pacific caused the Korean
container ship Hansa Carrier to lose overboard twenty-one deck-cargo
containers, each approximately 40 feet long. Among the items lost
were 39,466 pairs of NIKE brand athletic shoes on their way to the
United States. Six months to a year later, those shoes began washing
up along the beaches of Washington, Oregon, and British Columbia
(box fig. 1). They were wearable after washing and having the barnacles
and oil removed, but the two shoes of a pair had not been tied together
Box Figure 1 Shoes and toys after their rescues from Pacific cargo spills.
©Rowland Studio, Seattle, WA
for shipping, and pairs did not come ashore together. As beach residents recovered the shoes (some with a retail value of $100 a pair),
swap meets were held in coastal communities to match the pairs.
Other Scenarios
In May of 1991, oceanographer Curtis Ebbesmeyer of Seattle,
Ebbesmeyer was interested in seeing where the shoes might have
Washington, read a news article on the beached NIKEs. He was ingone if they had been lost on the same date but under different conditrigued and realized that 78,932 was a very large number of drifting
tions in other years. The computer allowed simulations for May 27 of
objects compared to the 33,869 drift bottles used in a 1956–59 study
each year from 1946 to 1991. Box figure 3 shows the wide variation in
of North Pacific currents. He contacted Steve McLeod, an Oregon
model-­predicted drift routes. If the shoes had been lost in 1951, they
artist and shoe collector, who had information on locations and dates
would have traveled in the loop of the Alaska Current. If they had
for some 1600 shoes that had been found between northern Califorbeen lost in 1982, they would have been carried far to the north durnia and the Queen Charlotte Islands in British Columbia. Additional
ing the very strong El Niño of 1982–83, and if lost in 1973, they would
beachcombers were asked for information, and Ebbesmeyer mapped
have come ashore at the Columbia River.
the times and locations where batches of 100 or more shoes had
On December 15, 2002, a ship carrying cargo containers bebeen found (box fig. 2). Next Ebbesmeyer visited Jim Ingraham at
tween Los Angeles and Tacoma, Washington, ran into 25-foot seas off
NOAA’s National Marine Fisheries Service’s offices in Seattle to study
Cape Mendocino in northern California. Several containers fell overhis computer model of Pacific Ocean currents and wind s­ ystems north
board, one containing 33,000 NIKE athletic shoes. By mid-January
of 30°N latitude. Using the spill date (May 27, 1990), the spill location
2003, the northward-flowing Davidson Current had taken the shoes
(161°W, 48°N), and the dates of the first shoe landings on Vancouver
833 km (450 nautical miles) to the Washington coast. Once again, the
Island and Washington State beaches between Thanksgiving and
Christmas 1990, Ebbesmeyer found that the shoe drift rates
agreed with the computer model’s predicted currents.
News of Ebbesmeyer and Ingraham’s interest in the shoe
Recovery locations
spill reached an Oregon news reporter and was then circulated
250 Shoes
100 Shoes
nationally. Readers sent letters describing their own shoe finds.
26 March
Jan./Feb.
63°N
200 Shoes
Reports of single shoes were valuable, because each shoe had
200 Shoes
Alaska c
18 May
u
Nov./Dec.
within it a NIKE purchase order number that could be traced to
200 Shoes
a specific cargo container. Ebbesmeyer was able to determine
150 Shoes
4 April
from these numbers that only four of the five shoe containers
200 Shoes
9 –10 May
had broken open, with only 61,820 shoes set afloat.
Ca
ka
l
c
s
i
ur f
a
l
The computer model and previous experiments with
56°N
A
r
­satellite-tracked drifters showed that there would have been
little scattering of the shoes as the ocean currents carried them
12
t
6°
W
re n
eastward and approximately 1500 miles from the spill site to
c ur
c
i
t
rc
a
b
Su
shore, but the shoes were found scattered from California to
northern British Columbia. …
Purchase answer to see full
attachment