Download Atmosphere-Ocean Coupling and Surface Circulation of the Ocean

Atmosphere-Ocean Coupling and Surface
Circulation of the Ocean
JAMES C. INGLE, JR.
WINDS
AND
WAVES
Ancient Greek philosophers viewed the ocean beyond the
ocean constitute a single dynamic circulatory system under-
Mediterranean Sea as a great river, Oceanus Fluvius, which
they considered to be directly related to the Earth (Ge) and
the sky (Uranus). Today, we recognize that surface circulation
going constant convective motion driven by radiant heat from
the Sun. The importance of this system to the well-being of
of the
winds,
water,
direct
global ocean is largely the product of so-called zonal
the Earth’s rotation, differences in the density of seaand the spatial constraints imposed by continents. The
link between the motion of the atmosphere and the
our planet cannot be overemphasized. The constant physical
and chemical exchanges between the ocean and the atmosphere govern the tempo and mode of both marine and terrestrial environments on all scales, and thus control the fundamental character of the surface of the Earth. The goal of this
surface ocean is apparent to anyone watching wind-whipped
storm waves crash on a beach. Even when local winds are
calm, drifting objects offer direct evidence that the surface of
chapter is to provide an understanding of the dynamic interplay between the atmosphere and the ocean, with an emphasis on the basic processes controlling the surface circulation
the sea is in constant motion. Viewed on a global scale, the
Earth’s intimately coupled gaseous atmosphere and liquid
of the ocean. The following chapter reviews circulation of the
deep ocean and considers ocean circulation in its entirety.
A
lthough the ocean covers 71 percent of the Earth’s
surface and constitutes the dominant and defining
environment of our planet, we are only now becoming familiar with how it circulates and what lies
beneath its surface. The very dimensions of the ocean have
until very recently stood as a primary obstacle to understanding the circulation of this immense body of water, despite its
obvious importance to human affairs ranging from sustenance and commerce to climate prediction. The earliest written records make it clear that the seagoing Phoenicians applied their knowledge of wind and current patterns in the
Mediterranean Sea and elsewhere for both war and trade,
information no doubt hard won at sea and patiently accumulated to their advantage over the centuries. Prior to the
launching of Earth-observing satellites, any attempt to gain
an overview of the surface circulation of the ocean involved
the same strategy presumably employed by the Phoenicians
some 3000 years ago - laborious compilation of observations
made aboard individual ships at sea in order to obtain the
larger picture. In fact, knowledge gained in this manner by
American sailors allowed Benjamin Franklin to publish the
first map of a major ocean current, the Gulf Stream, in 1769,
aiding merchant shipping from North America to England in
the process.
By 1837-40, the German geographer Heinrich Berghaus
had assembled the first truly comprehensive maps depicting
major surface currents of the world oceans aided by the observations of other scientists including Alexander von Humboldt. At approximately this same time, the importance of
commercial whaling led the government of the United States
to support an “exploring expedition.” This global voyage was
led by Commander Charles Wilkes, who published a fivevolume report in 1845. This study included a remarkably
accurate map of global surface circulation, which, for various
reasons, received little attention. Subsequently, one of
Wilkes’s younger cohorts, Lieutenant Matthew F. Maury, set
about synthesizing information on winds, currents, and
ocean temperatures recorded in a huge collection of ships’
logbooks that had been stored, but little used, by the U.S.
Navy. Maury’s efforts yielded a fresh set of maps and track
charts depicting global patterns of winds and surface currents. These maps and charts were widely distributed and had
an immediate positive impact on maritime trade and selection of sailing routes (Fig. 10.1). The usefulness of Maury’s
W. G. Ernst (ed.), Erzrth Systems: Procr.ssc.s rwrl fwrc~~. Printed in the United States of America. Copyright 0
2000 Cambridge University Press. All rights reserved.
152
Figure 10.2. Satellite view of the Earth centered over Africa and
emphasizing the dominance of the ocean environment. Thick
swirls of clouds in the lower half of this photograph mark storm
systems and winds driving the West Wind Drift (also known as
the Circum-Antarctic Current) eastward around the ice-covered
continent of Antarctica (clearly visible at the bottom of the
gwiR). The a’lc d 5pott-r’ ,&xl& ranging Ttcx65 the AeiYiic
Ocean, central Africa, and the Indian Ocean just above the midline of the photograph represent the equatorial zone of rising
warm, moist air and low-pressure storm cells associated with the
intertropical convergence zone (ITCZ). Areas of clear sky over
southwest Africa and the desert areas of North Africa, the Red
Sea, and the Persian Gulf constitute mid-latitude zones of descending dry air, which increases in temperature through conduction and high pressure (adiabatic compression) as it approaches Earth’s surface. Clear air over Antarctica is due to
extremely cold, dense, dry air descending over the polar region.
(Source: Photograph courtesy of NASA.)
charts led to the first attempt to standardize oceanographic
observations aboard all sailing vessels. Among his other accomplishments, Maury also produced the first bathymetric
(depth) chart of a portion of the Atlantic Ocean based on
primitive sounding by U.S. Navy vessels under his direction.
In 1855, Maury summarized his work in a volume entitled
the Physical Geography of
Sea, which was reprinted several
times to meet demand and translated into several languages.
However, many of his attempts to explain ocean and atmospheric circulation were naive and speculative, and therefore
were much criticized by the scientific community, despite
the value of his charts for improved sailing.
Although progress continued in amassing data on surface
circulation, little was known about the deep sea, and another
century passed before the first modern maps depicting the
floor of the ocean appeared. In the intervening years, the
modern science of oceanography took shape, spurred by
the results of the HMS Challenger expedition, which circled
the globe from 1872 to 1876. Among the manp
scientific firsts accomplished during the Ch&
Ienger expedition were measurements of terns _
perature, salinity, and density at 362 deep-sea
stations, yielding some of the first insights into
.$
the layered structure and circulation, of the 3
deep sea. This circumglobal cruise was the irn- -<
petus for a number of later expeditions, and !
the collection of many additional measure8 e:
4
ments.
However, it was the advent of surface and
submarine warfare during World Wars I and 11
that accelerated research into all aspects of
physical oceanography and ocean circulation,
This latter research peaked during the Cold
War years of 1950-90, as ever more sophisticated acoustic technology was employed to determine the detailed density structure of the
ocean as a means to detect enemy vessels and
operate effectively in the submarine environment (as dramatically fictionalized in Tom
Clancey’s 1984 novel The Hunt forRed October}.
Despite all the advances in sampling and
measuring properties of the surface and deep ocean from
the 1950s to the present, it is the synoptic observations of
the Earth from satellites that have allowed a quantum leap
in QUK un&~stanting of c~upkd atmosphere-ocean &a&tion (Fig. 10.2). Satellite observations became routinely available in 1960 with the launching of 77ROS, which photographed the Earth’s weather patterns and forever changed
our view of the ocean and the world. Monitoring of electromagnetic radiation from various Earth environments was initiated by the LANDSAT satellites beginning in 1972, followed by the NIMBUS satellites from 1978 through 1984.
NIMBUS instruments provided quantitative measurements of
ocean productivity, water vapor in the atmosphere, ice coverage, and a host of other parameters reflecting aspects of
ocean circulation and climate. SEASAT, the first satellite devoted exclusively to oceanographic observations, was also
launched in 1978 but operated for only three months. Despite the premature failure of SEASAT, its radar provided the
first synoptic observations of ocean wave patterns and surface winds along with measurements of sea surface temperature and elevation.
In 1992, the French-U.S. TOPEXPoseidon
satellite was
launched into orbit for the purpose of observing and sensing
the ocean - the first such satellite since ill-fated SEASAT. The
high-altitude orbit of TOPEXLF’oseidon allows it to observe
over 95 percent of the ice-free ocean every ten days. Instruments aboard the satellite are capable of measuring a range
of ocean parameters with unprecedented accuracy, including
wind speed and direction, surface currents, and minute variations in the height of the sea surface reflecting both aspects
of wind-driven ocean circulation and the effect of the Earth’s
rotation and gravity. Significantly, the first analyses of TOPEX/Poseidon data are only recently appearing in the scien-
ATMOSPHERE-OCEAN
tific literature; it is fair to say that 1992 marked a watershed
moment in our ability to observe and understand ocean circulation, comparable to the scientific threshold crossed when
the C?zaZZenger expedition set sail over a century ago.
THE AIR AND SEA IN MOTION
The most important role of the atmosphere-ocean system is
the redistribution of excess solar heat that the Earth receives
in the equatorial and mid-latitude regions (Fig. 10.3). The
atmosphere carries heat in the form of latent heat of evaporation (e.g., water vapor) as well as sensible heat (the sort
measured with a thermometer), whereas the ocean carries
heat only in the sensible form. Most of the Sun’s radiant heat
energy arriving at the sea surface goes into breaking the weak
hydrogen bonds between individual water molecules and
evaporating seawater, not into raising the temperature of the
water. This process is a function of the polar character of the
water molecule and reflects the fact that the heat capacity of
water (e.g., heat absorbed divided by temperature rise) is the
highest of ail solids and liquids, with the exception of liquid
ammonia. Thus, the great volume and heat capacity of seawater allows the ocean to store enormous amounts of solar
heat and release it slowly back to the atmosphere by conduction at the air-sea interface. This process, along with the
continuous evaporation and condensation of seawater by the
atmosphere, serve to w the surface tenspcrrature af the
Earth within a range allowing life as we know it to thrive.
Clearly, the ocean plays a critical role in maintaining a lifefriendly climate on our planet.
Circulation of the atmosphere and ocean represents the
never-ending quest nature of these two systems collectively
to establish thermal equilibrium between the poles and the
equator - a goal they will never reach thanks to unequal
distribution of the Sun’s heat over the curved surface of the
Earth, the Earth’s constantly changing climate, and the slow
but unceasing tectonic rearrangement of continents and
ocean basins. Air and water heated in the tropics and subtropics are transported poleward, while water and air cooled
N 90”
60”
30”
155
at high latitudes move equatorward. Complex and littleunderstood feedback loops characterize these processes and
assure that any change in the behavior of the atmosphere or
ocean will have profound consequences for the circulation
of both systems. The hypersensitive interrelationships between the atmosphere and the ocean are dramatically illustrated by the global weather extremes associated with El
Nifio/Southem Oscillation (ENSO) conditions, which appear
approximately every seven years in the Pacific region and
elsewhere.
At the initiation of an El Niiio event, the large, atmospheric high-pressure zone normally present over the South
Pacific weakens while the large, low-pressure system operating over the Indian Ocean becomes stronger. In turn, equatorial trade winds weaken, and the thick mound or wedge of
warm water normally maintained by these winds in the westem Pacific is allowed to flow eastward toward the Pacific
coasts of the Americas. The arrival of the warm surface water
in the eastern Pacific disrupts normal upwelling of cold
nutrient-rich waters, causing the temporary collapse of fisheries, exceptional storms, high rainfall and flooding, and
even a rise in local sea level. The consequences of a sustained
El Nifio event include associated changes in the positions of
the atmospheric jet streams and disruption of normal
weather patterns on a global scale, translating to billions of
dollars in storm, drought, and agricultural damage, as well as
lost lives and subsequent Yeats af recovery, There is cl.ea.rLy a
need to understand the details of ocean-atmosphere interactions, and these processes represent frontier areas of ongoing
ocean research.
Aspects of atmosphere-ocean coupling are intuitively
straightforward; the wind blows and the sea’s surface is
moved, resulting in wind-driven surface circulation. However,
the wind effectively stirs only a relatively thin layer of surface
water, commonly no more than 100 meters in thickness.
Contrary to common thinking, the bulk of the water in the
oceanic bowl is not stirred by the wind but is gravity-induced
and circulates as a function of variations in the temperature,
salinity, and density of individual water masses. Therefore,
f-1
Heat
loss
Elgnre 10.3. Net gain and loss of solar
energy (heat) at the top of the Earth’s
atmosphere and at the surface of the
Earth. Solar energy received at the sur-
(4
Heat
gain
face (insolation) is unevenly distributed
across latitude as a function of the curvature of the Earth and atmospheric effects. See text for discussion.
30”
60”
s 90”
COUPLING
(4
Heat
loss 0
300
600
Solar energy (heat) - cal./cm2/min
156
JAMES C. INGLE, JR.
circulation of the deep ocean is referred to as thermohaline
circuZation (“thenno” means “temperature,” and “haline”
means “related to salinity or salt content”), a subject considered in Chapter 11. Although only a relatively thin layer of
the surface ocean is directly moved by the wind, global patterns of surface circulation and air-sea interaction control
much of the ocean’s physical, chemical, and biologic character at all depths.
Convective and advective processes are responsible for
heat transport and motion of discrete air and water masses
within the atmosphere and the ocean (Fig. 10.2). Advection
refers to changes in the property of an air mass or water mass
by virtue of bodily motion. The term is often restricted to the
horizontal motion of air or ocean water but can also apply to
vertical motion. Water or air in motion is advectively transporting heat or other properties of the fluid regardless of how
the motion was initiated, whether through mechanical
means (e.g., one moving air mass pushing another) or by
gravity-driven convection. A simple analogy is the initiation
of motion by stirring water in a pot with a spoon. In contrast,
convective motion is self-initiated whenever a fluid or air
mass experiences changes in density (e.g., change in mass
per unit volume, measured in grams, per cubic centimeter) as
a function of variations in temperature, composition, or pressure and is a response to gravity acting upon these changes.
As a given mass of air or water is heated, molecular activity
increases and it expands, taking up more space (e.g., specific
volume or volume per unit mass), with a consequent decrease in its density. The change to lower density causes the
air or water mass to rise within the density-stratified column
of the atmosphere or ocean to a level commensurate with its
new density. Conversely, a parcel of cold, dense air or water
will sink in the density-stratified column to a level where it
is surrounded by air or water of similar density and underlain
by fluid of greater density.
The high angle of the Sun’s rays or solar beam with Earth
and the round shape of the Earth ensure that most of the
radiant solar heat or energy is received between 30” N and S
of the equator (Fig. 10.3). In contrast, the polar regions experience a constant heat deficit due to the low angle of the
Sun’s rays at high latitudes and heat loss through back radiation to space. The unequal distribution of heat results in a
significant difference in temperature, or thermal gradient,
between the equator and the polar regions regardless of
whether the Earth is in a glacial or nonglacial climatic mode.
The greater the difference in temperature between the poles
and the equator, the steeper the thermal gradient. The ever
changing pole-to-equator thermal gradient, together with
subtle but critical changes in the density layered structure of
the atmosphere and the ocean, maintain constant convective
motion in both systems. The steeper the pole-to-equator
thermal gradient, the faster the rate of atmosphere-ocean
convection and circulation, and vice versa. (See the box entitled The Dynamic Energy Balance of the Earth.)
The overall pattern of atmosphere-ocean convective circulation thus mimics the familiar convective motion seen in
a pot of heated water on a stove. Water heated at the bottom
of the pot expands as molecules become more active, decreases in density, and rises to the surface where it cools,
increases in density, and sinks to the bottom of the pot,
where the process begins again. With sufficient heat and
time, the result of this process is a rolling boil - vigorous
circulation will have been induced through changes in the ternperature and density of the water without any mechanical help from
a spoon. Just as the rate of boil in a pot of water can be
modulated by adjusting the stove-top flame, any changes in
the pole-to-equator thermal gradient over time result in increasing or decreasing rates of convective circulation within
the atmosphere and ocean. The steepest pole-to-equator thermal gradients occur when the Earth is in a glacial mode, and
evidence suggests that these periods are indeed characterized
by increased rates of atmospheric and ocean circulation.
Conversely, periods in Earth history when the polar regions
remain ice-free and relatively warm are marked by shallow
pole-to-equator thermal gradients and relatively slow oceanatmosphere circulation.
The convective and advective processes operating in the
atmosphere cause air masses to rise and sink, generating variations in atmospheric pressure at the surface of the Earth. In
turn, winds are generated as air rushes from areas of high
pressure to areas of low pressure. Winds developed in the
lower atmosphere transfer their energy and momentum to
the surface layer of the ocean through friction at the air-sea
interface. Not surprisingly, global patterns of surface circulation in the ocean reflect the direction and strength of winds
(i.e., wind stress) in the lower atmosphere. This is the case no
matter what the configuration of continental masses might
ATMOSPHERE-OCEAN
be over geologic time. However, tectonically induced
changes in the number, shapes, and locations of continents
and ocean basins over geologic time play an equally large
role in modulating ocean circulation, in turn emphasizing
the intimate link between the behavior of internal (endogenie) and external (exogenic) Earth systems over periods of
millions of years.
THE LAYERED STRUCTURE OF THE
ATMOSPHERE AND OCEAN
Many people enjoying a summer swim in a lake or the ocean
have experienced the strange sensation of having their upper
body in comfortably warm surface water while their legs dangle in cold water below - a clear example of the thermal and
density stratification of a water column. The warmer, less
dense water floats on top of denser, cold water with a sharp
temperature and density boundary dividing the two layers.
Convective circulation of the atmosphere and ocean is dependent upon variations in the density of the various air and
water masses. Both systems exhibit a layered structure, with
the heaviest or densest air and water residing in the bottom
of the atmosphere and the bottom of the ocean, respectively.
Because of density stratification, horizontal motion dominates both systems. Areas of vertical motion are limited to
zones where advective or convective processes introduce instabilities in the air or water column through rapjd changes
in temperature or other parameters, causing the air or water
to rise or sink. Although zones of vertical motion in the
atmosphere and ocean are limited in area1 extent, they are
critically important.
The Earth’s gravity maintains the highest density of atmospheric gases immediately adjacent to the Earth’s surface,
accompanied by high atmospheric pressures (this phenomenon is manifest in the annoying ear pressure change felt
during an airplane landing or a rapid descent down a mountain road). Atmospheric gases thin and pressure and density
decrease with increasing altitude above the Earth’s surface.
The atmosphere can be divided into four major layers based
on the systematic changes in its temperature, beginning with
the near-surface troposphere and followed by the stratosphere, mesosphere, and thermosphere. The lowest layer, the
troposphere, contains 75 percent by mass of all the gases in
the atmosphere, along with most of the water vapor and the
majority of clouds, dust, and so forth. The lower atmosphere
is heated both by release of latent heat when water vapor
condenses into clouds and rain, and by conduction from the
surface of the Earth (in the same manner that your hand is
heated when you place it on a hot surface). Latitudinal variations in the surface temperature of the Earth as well as
seasonal variations in temperature at any given location
maintain constant thermal instability in the troposphere.
Convective motion in the troposphere spawns the Earth’s
weather, and this layer contains the major surface wind systems of primary importance to oceanographers studying
ocean circulation.
157
COUPLING
Density stratification in the ocean is the product of variations in temperature, salinity, and pressure with depth. Density (measured in grams per cubic centimeter) increases with
decreasing temperature, increasing salinity, and increasing
pressure; density decreases with increasing temperature, decreasing salinity, and decreasing pressure. A plot illustrating
the range of variation in temperature and salinity demonstrates that the modern ocean has an average temperature of
only approximately 3.5X, reflecting the high-latitude origin
of most of the water filling the ocean basins (Fig. 10.4). Temperatures in the open ocean range from a low of -2°C to a
high of approximately 32°C exceeding the much narrower
range of salinity. Salinity represents the amount of salt in a
given volume of water and is routinely measured in terms of
parts per thousand (O/O). The ocean exhibits an average salinity of approximately 35 parts per thousand. Pressure due to
the weight of the overlying water column becomes a significant factor affecting the calculation of density only at depths
greater than 1000 meters and is generally discounted at shallower depths. Thus,
parameters governing the density of seawater, and both vary
horizontally and vertically in the ocean (Fig. 10.5). Even a
minute change in either or both of these parameters in a
given parcel of water translates to a significant change in its
density. The relatively narrow range of temperatures and salinities in the ocean means that scientists must precisely calculate very small differences in density, a calculation typically carried out to five places and then converted to a
density factor termed sigma-t for convenience. For example,
a calculated density of 1.02532 grams per centimeter would
be converted to a sigma-t value of 25.32 for ease of plotting
and manipulation.
Clearly, any process controlling the temperature or salinity of seawater has the capacity to change the density of the
Range of temperature and salinity in the global
ocean, as illustrated by contours enclosing values for 99 and 75
percent of all the water in the ocean. The range of salinity is
relatively narrow compared with that of temperature. The very
cold average temperature of the ocean reflects the high-latitude
origin of most of the water in the deep ocean. (Source: Adapted
from M. G. Gross and E. Gross, Oceanography, 7th edition. Englewood Cliffs, NJ: Prentice-Hall, 1996.)
Average salinity = 34.7%0
Figure 10.4.
Average
surface
1( temperature
= 17S”C
Average
d temperature
= 3S”C
35
Salinity (%0)
158
JAMES C. INGLE, JR.
Sea Surface Temperature (“C)
I
60
Sea Surface Salinity (o/o)
Figure 10.5. Generalized patterns of ocean surface (A) temperatures and (B)salinities in August (Northern Hemisphere summer).
Lines of equal temperature (isotherms) and equal salinity (isohalines) tend to parallel latitude in the open ocean. Alterna-
tively, isotherms are torqued north and south along continental
margins, reflecting displacement by surface currents in these
regions. (Sotrrce: Goode Base Maps, courtesy University of Chicage.)
water and in turn, its position in the oceanic water column.
Significantly, most variations in temperature and salinity in
the ocean as a whole initially occur at the sea surface through
heating, cooling, evaporation, precipitation, and freezing
(Figs. 10.5 and 10.6). The distribution of surface temperature
reflects the pole-to-equator thermal gradient and patterns of
wind-driven surface circulation. The temperature of a parcel
of surface water can vary as a function of its latitudinal position, mixing with water of different temperature, or residence time in a given locality. Warm, light water character-
izes the equatorial areas, cold, dense water is formed in polar
regions as surface water arriving from lower latitudes is
cooled. Variations in salinity also play a major role in creating differences in density of individual water masses, In the
highest latitudes, seawater is frozen; sea ice begins to form at
a temperature of approximately -1°C. Because salt is excluded during the formation of ice crystals, the salinity of
the remaining unfrozen water increases, in turn lowering its
freezing temperature and increasing its density, due to this
increase in density, the unfrozen water sinks. The salinity of
ATMOSPHERE-OCEAN
COUPLING
1 5 9
Variations in temperature, salinity, and density with
depth in a hypothetical mid-latitude column of ocean water
clearly define the three basic layers or zones common to much
of the world ocean: the surface, intermediate, and deep layers
(Fig. 10.7). In some areas, a fourth layer is present in the form
of Antarctic Bottom Water, representing the very cold and
relatively saline, high-density water derived from the freezing
of seawater around the margins of Antarctica. The surface or
36
mixed Zayer is the product of stirring and turbulent mixing by
1.0280
the wind. Although the surface layer can vary between 50
1.0260
and 500 meters in thickness, frictional decrease in velocity
G
with depth on average limits wind-induced motion to only
g 1.0240
3 4 W) the upper 100 meters of the ocean and in the process defines
53
the limits of the mixed layer. The range of temperature and
1.0220
salinity in the surface layer are relatively constant at any
given location, reflecting latitudinal location, mixing with
32
the atmosphere, wetness of the overlying atmosphere (i.e.,
humid versus dry air), and seasonal
_ --- variations. For example,
a shallow seasonal thermocline or temperature gradient commonly develops within the surface layer as summer warming
heats near-surface water. Although the surface layer contains
only approximately 2 percent of all the water in the ocean, it
is arguably the most important part of the ocean in terms of
the physical, chemical, and biologic processes that control
the activity and character of the ocean as a whole.
N- 7 0 ” 4 0 ” 2 0 ”
0”
20”
In contrast to seasonal variations in temperature within
Latitude
the se layer, the pm themuKline (Ykwmlo”means
Figure 10.6. Variation of sea surface temperature, salinity, and “temperature,” and “cline” means “gradient”) begins at the
density with latitude and average annual patterns of precipitabase of the surface layer and extends on average to a depth
tion (rain) and evaporation at the sea surface. A correlation exists
of approximately 1000 meters (Fig. 10.7). The top of the
between areas of excess precipitation beneath the polar frontal
thermocline is commonly associated with the 15°C isotherm
zones (approximately 50 to 60” N and S) and the equatorial
(i.e., the line of equal temperature) in mid-latitude locations,
region with areas of depressed surface salinities. Similarly, there
with
the base of this zone generally marked by the 4 or 5°C
is a clear relationship between areas of excess evaporation associated with mid-latitude high-pressure zones (approximately 30”
isotherm. A well-developed permanent thermocline is presN and S) and relatively higher surface salinities. The slight offset
ent in most of the ocean but is absent in polar regions where
of the equatorial zone of excess precipitation to the north of the
surface temperatures remain very cold throughout the year.
geographic equator reflects the northward displacement of
The permanent thermocline is commonly accompanied by
the so-called meteorologic equator (e.g., the east-west line of
hypothetical thermal equilibrium between hemispheres located an equally dramatic hdocline representing a significant inapproximately 5” north of the geographic equator) and the intercrease in the salinity of water across these same depths. The
tropical convergence zone (see Fig. 10.8). g/cm3 = grams per
rapid changes in temperature and salinity associated with the
cubic centimeter; cm/yr = centimeters per year; 9;60 = parts per
thermocline and halocline combine to produce an accompathousand.
nying gradient in the density of the intermediate water
termed the pycnocline. The pycnocline layer or zone contains
approximately 18 percent of the water in the ocean and
surface water is also increased in the mid-latitudes as a funcserves to separate the relatively dynamic surface layer from
tion of evaporation. Conversely, high precipitation in equa- the very cold, dense, and relatively stable water residing in
torial and subpolar areas decreases the salinity of surface wathe deep ocean.
ters in these regions. The mixing of water masses of different
The deepest of the three primary zones constituting the
temperature and salinity can also alter the character of the oceanic water column is appropriately termed the deeq layer
newly formed water mass (as discussed in Chap. 11). Finally,
or deep zone, which is characterized by very cold, highit is important to note that salinity can be the dominant density water resulting from its origin in higher-latitude
factor controlling the density of seawater in some shallowregions (Fig. 10.7). Water in the deep layer has an average
marine settings, such as over continental shelves or within temperature of less than 4°C. The deep layer includes 80
semienclosed coastal lagoons and estuaries, and in larger enpercent of the water in the ocean and consequently plays a
closed bodies of water, such as the Mediterranean Sea and major role in global heat distribution (as discussed in Chap.
the Persian Gulf.
11). That water in contact with the deep-sea floor is termed
160
JAMES C. INGLE, JR.
Temperature
10
Salinity
20°C 32
33
Density
after the French mathematician Gaspard
G. de Coriolis, who first auantitativelv ex\ Surface
plained the effect of a rotating frame of
or mixed
reference in 1835. For example, if a person
layer
riding on the inside ring of a rotating carousel throws a ball to another rider traveling on the outside edge of the carousel,
the ball appears to travel in a curved path
from the perspective of the ball thrower.
In fact, the ball travels in a straight line
between the two riders, as observed by a
person standing next to the carousel and
viewing the action from a fixed frame of
reference. The curved path of the ball observed by the thrower results from the ball
catcher moving at the same time the ball
is in flight - an apparent deflection of the ball’s path imparted by the rotating frame of reference of the thrower. No
“force” (or acceleration) is involved in creating this apparent
deflection of motion.
Any object with mass moving horizontally and freely over
the surface of the rotating Earth is subject to the apparent
deflection of its motion due to the Coriolis effect, although
no force has been applied. The apparent deflection of motion
occurs when the speed and direction of the object are viewed
or measured in reference to the underlying surface of the
rotating Earth. Hence, it is more correct to speak of the Coriolis effect rather than the Coriolis force. However, the Coriolis effect or “force” is relatively weak and typically does not
influence the motion of small masses over short distances
where other forces are dominant. For example, in a bathtub,
the water does not always swirl to the right as it exits down
a drain in the Northern Hemisphere. Another example is the
fact that you do not have to compensate for the Coriolis
effect as you walk down the street or toss a Frisbee to a friend.
In contrast, as huge masses of air in the atmosphere or water
in the ocean travel great distances over the rotating solid
Earth, the Coriolis effect is very significant relative to other
forces acting on these masses.
The differential velocity of the eastward spinning Earth
increases with increasing latitude; therefore, the amount of
deflection imparted by the Coriolis effect or “force” and experienced by a moving air mass (or water mass in the ocean)
is dependent upon its velocity and latitude. The Coriolis effect is zero at the equator and increases with increasing latitude toward the poles, while the magnitude of deflection
increases with the increasing velocity of the object in motion. Hence, both the circulation of the atmosphere and the
surface ocean are profoundly influenced by the Coriolis effect. The phenomenon is clearly seen in the tendency of the
surface winds to turn to the right of their motion in the
Northern Hemisphere and to the left of their motion in the
Southern Hemisphere (Figs. 10.2 and 10.8). These deflections, together with convection and variation in the velocity
of air masses as they travel poleward, result in six cells or
“tubes” of rotating air that encircle the Earth and define
34 35%0 1.023 1.025 1.027 g/cm3
\
Deep
layer
10.7. Generalized patterns of temperature, salinity, and
density through a mid-latitude water column, emphasizing the
basic three-layer density-stratified character of the ocean. The
thermocline, halocline, and pycnocline zones mark steep gradients in these parameters within the intermediate layer of the
ocean. The depth to the base of thermocline, halocline, and
pycnocline can vary between 700 and 1500 meters, but is shown
here as an idealized 1000 meters. The surface or mixed layer
extends to only approximately 100 meters due to the frictional
decrease in wind-induced motion with depth. g/cm3 = grams per
cubic centimeter; km = kilograms; %o = parts per thousand.
Figure
bottom
Discrete water masses traveling along the bottom of the deep sea include the Antarctic Bottom Water, the
coldest and densest water in the world ocean, with an average temperature of -0.4”C.
ZONAL WINDS AND CIRCULATION OF THE
LOWER ATMOSPHERE
Convection of the atmosphere is driven by the unequal distribution of the Sun’s heat over the surface of the Earth.
Given a nonrotating Earth, air heated in the equatorial region would expand and rise, creating a low-pressure zone,
and would flow toward the poles. As the air approached the
polar regions, it would cool, contract, increase in density,
and sink, creating high pressure and flow equatorward along
the Earth’s surface, completing a simple convective loop. Under these conditions, surface winds would blow from the
polar regions of high pressure to the equatorial region of low
pressure. Moreover, the flow of surface winds would be oriented due north-south, at right angle to the lines of equal
atmospheric pressure or isobars.
In the real world, the Earth is rotating from west to east,
masses of air moving freely over the Earth’s surface must be
viewed in a rotating frame of reference that in turn imparts
an apparent torque to their motion. One can think of the
solid Earth as rotating out from beneath a moving air mass
and in the process creating an apparent clockwise deflection
of its motion (to the right) in the Northern Hemisphere and
a counterclockwise deflection (to the left) in the Southern
Hemisphere. The right-and left-handed hemispheric deflections are due to the Coriolis effect or “force,” named
ATMOSPHERE-OCEAN
COUPLING
161
High pressure
LOW
pressure\
90"N
30”N
pressure--
Intertropical convergence zone
:eizi\
cell
z
\h
\ilI:::-oo
1
30”N
pressure
Ferrell
cell *
90”N
pressure
F4gare 10.8. Schematic illustration of global atmospheric circulation and surface wind patterns
(arrows on the Earth’s surface). Three large convecting cells of air (shown in cross section on the
left-hand side of the globe) define circulation of the lower atmosphere in each hemisphere. The
surface components of each atmospheric cell form the zonal wind belts that drive surface circulation
of the ocean. The limbs of the atmospheric cells include: (1) zones of rising moist air, low pressure
(L), and high rainfall in the equatorial zone and along the polar fronts (at 50 to 60” N and S), and
(2) zones of descending dry air, high pressure (H), and low rainfall over the polar regions and in the
mid-latitudes (at approximately 30” N and S). Eastward rotation of the Earth and the Coriolis effect
cause surface winds to veer to the right of their motion in the Northern Hemisphere and to the left
of their motion in the Southern Hemisphere. The polar jet streams are not surface winds but
rather
flow eastward at high tropospheric altitude along the polar fronts in wave-like patterns
and influence
the position of individual high- and low-pressure systems north and south of these fronts.
.
global atmospheric circulation. The number of convecting
cells is in part a function of how fast the Earth is spinning,
and (as noted earlier) only two large convection cells would
be in operation on a hypothetical nonrotating Earth.
A brief description of the major atmospheric cells will
assist in explaining the formation and pattern of surface
winds responsible for wind-driven surface circulation of the
ocean. Solar radiation is at a maximum in the equatorial
region, resulting in high sea surface temperatures and high
rates of evaporation of seawater. As the warm, moist air expands and rises, it ultimately reaches an elevation where it
cools to the point that water vapor condenses to form clouds
and rain, releasing latent heat in the process. The result is a
persistent band of clouds and low pressure along the equatorial region. Small, low-pressure cells form continually in
this region, resulting in heavy rainfall throughout the year
and hot and humid weather with irregular breezes (Fig. 10.8).
The high rainfall associated with this zone offsets the high
rates of evaporation and depresses the salinity of equatorial
surface waters (Fig. 10.6). In short, heat-driven vertical motion dominates the equatorial atmosphere, resulting in a
zone of weak and variable surface winds termed the doldrums,
a phenomenon that presented a major obstacle to ancient
sailing vessels and still hinders modem sailing vessels attempting to cross the equator. At the same time the cooler
and drier air aloft in the upper troposphere moves north and
162
JAMES C. INGLE, JR.
south away from the equator and continues to cool as it
travels poleward, gaining density as its temperature drops.
At approximately 30” N and S of the equator, the now
cold and relatively dense air derived from equatorial areas
sinks toward the surface of the Earth, creating a zone or belt
of high pressure (Figs. 10.2 and 10.8). As the air descends, it
is heated by conduction (i.e., by transfer of heat from the
relatively warm surface of the Earth) and by compression
(i.e., increasing adiabatic pressure). As the temperature of the
air rises, its capacity to hold moisture increases, with the
result that the mid-latitude high-pressure belts are characterized by cloudless skies and low rainfall. The high rates of
evaporation associated with the 30” high-pressure zones increase the salinity of underlying surface water and create
deserts on land. Upon reaching the Earth’s surface, some of
this air flows toward the low-pressure belt of the equatorial
region. The Coriolis effect deflects these winds to the right of
their motion in the Northern Hemisphere and to the left in
the Southern Hemisphere, forming the northeast and southeast trade winds, which are separated by the
vergence zone (ITCZ) and the doldrums. The trade winds blow
continually westward except during the unusual conditions
associated with El Nifio events, when they slow, stop, or even
reverse their direction.
Commonly during the late summer, isolated low-pressure
disturbances within the tropical trade wind belts between 5”
and 2W latitude graw into incre&ngLy
large and violent
storms termed hurricunes, typhoons, or cyclones (except in the
equatorial South Atlantic Ocean). These storms rapidly transport large amounts of latent heat into higher latitudes accompanied by winds in excess of 118 kilometers per hour
and heavy rains, often with tragic consequences where they
meet land. A hurricane derives its energy from the latent heat
released as water vapor, rising off the tropical ocean, condenses into clouds and rain around the low-pressure center
of rapidly rising warm air marking the eye of a storm. Thus,
ocean surface temperatures play a key role in the formation,
travel, and ultimate death of these storms. Evidence indicates
that sea surface temperatures between 26 and 29°C are necessary to initiate the rapid vertical convection characteristic
of a hurricane and that this process cannot be sustained
when a storm arrives over water of less than 20°C.
At the same time the trade winds are blowing westward
and equatorward, some of the air descending at 30” N and S
flows poleward and is deflected eastward by the Coriolis effect, forming the prevailing westerlies* in both hemispheres
(Fig. 10.8). The warm, dry air of the westerlies aggressively
evaporates seawater and increases the humidity of these air
masses as they sweep poleward. Meanwhile, the very cold
and dense air formed at higher altitudes over the north and
south poles sinks in these regions, forming high-pressure
zones marked by cold, dry air that flows westward and equatorward, constituting the polar easterlies. The warm prevailing
l
Winds are labeled according to the direction porn which they blow;
hence, westerlies arrive from the west and travel eastward.
westerlies and cold polar easterlies collide at appro,ximately
50 to 60” N and S, forming the wave-like polar front in the
Northern Hemisphere and the Ankmtic front in the Southern
Hemisphere.
The polar frontal zones include associated polar jet streams
of high-velocity winds in the upper troposphere that travel
eastward around the world in ever changing sinusoidal patterns. These convergences
or collisions result in the advection of the relatively warm and lower-density air of the prevailing westerlies up and over the cold, dense air of the polar
easterlies. The rising, warm, moist air is rapidly cooled and
water vapor condenses, forming prominent zones of clouds
and high rainfall, and continually producing low-pressure
storm systems. The high rainfall associated with the polar
frontal zones dilutes underlying surface water, imparting
characteristic lower salinities to surface currents formed in
these regions (Fig. 10.6). The low-pressure storm systems
formed in these zones move from west to east along the polar
fronts, guided by the associated jet streams. As is the case in
the mid-latitudes, the drier air at higher elevations flows
equatorward and poleward away from 60” N and S latitude,
closing the convective cells on either side of the polar front.
Thus, six global atmospheric cells are responsible for the
basic pattern of surface winds no matter what the Earth’s
climatic state, and regardless of the geologically transient
locations of the continents and oceans (Fig. 10.8). Although
horizontal motion prevails in these convecting c&s, it is the
relatively narrow zones of vertical motion marking the limbs
of the cells that are of special importance. Rising warm, moist
air, condensation, and high rainfall are associated with the
belts or zones of low pressure at 0 to 20” and 50 to 60” N and
S, whereas zones of descending air and high pressure mark
zones of little precipitation at 30 to 40” and 80 to 90” N and
S. As the surface winds blow from areas of high to low pressure, they set in motion the large-scale surface circulation of
the ocean and transfer a portion of their energy to the ocean.
WIND-INDUCED MOTION OF THE SEA
SURFACE
Small capillary waves created by surface tension constantly
roughen the surface of the ocean and allow the wind to grip
the sea surface, transfer energy and momentum via frictional
drag, form waves, and sustain wind-driven surface circulation
and major surface currents of the global ocean. However,
wind waves per se simply represent the transfer of energy
along the air-sea interface via orbital motion and involve
very little transport of water. This section is concerned with
the long-term momentum imparted to the surface waters of
the ocean by the combined effects of wind, the rotation of
the Earth, and gravity; these effects are responsible for initiating and sustaining the major surface currents of the ocean
and their transport of truly enormous volumes of water on a
global scale. Once set in motion by the wind, momentum
carries the surface ocean forward in the direction of the dominant wind pattern even after local winds have slackened or
ATMOSPHERE-OCEAN
COUPLING
163
60”
South Pacific
60”
S
60”
Lx?”
180”
~20”
60”
Generalized surface circulation of the ocean and major surface currents and gyres.
Compare surface currents shown on this map with general patterns of surface winds shown on
Figure 10.8.
Figure
10.9.
died. However, frictional effects cause a rapid decrease in
current velocity with depth, and they restrict wind-induced
motion on average to the uppermost 100 meters of the water
column - the so-called surface layer. As with moving air
masses, surface water masses in the ocean are subject to the
Coriolis effect and drift to the right of their motion in the
Northern Hemisphere and to the left of their motion in the
Southern Hemisphere.
GYRES AND BOUNDARY CURRENTS
A quick comparison of maps depicting generalized patterns
of surface winds and major surface currents clearly illustrates
that zonal winds (e.g., latitudinal wind belts) are a primary
factor controlling surface circulation of the ocean (Fig. 10.8
and 10.9). Other factors also affect surface circulation, including differences in the temperature, salinity, and density
of individual water masses, variations in the elevation of the
sea surface from place to place, and the Coriolis effect. Surface circulation is also affected by the size and shape of individual ocean basins and the positions and configuration of
gateways between oceans, such as Drake’s Passage between
South America and Antarctica, which allows free communication between the Atlantic and Pacific oceans.
On an ocean-covered Earth without continents, surface
winds would produce a series of six east- and west-flowing
surface currents, each of which would continually circle the
world beneath the six zonal wind belts. In reality, the only
region of the ocean displaying this laboratory-like configuration lies below 40” S and constitutes the Southern Ocean
surrounding Antarctica, where no continents are present to
block or divert wind or surface water motion. The result is
the dreaded “Roaring Forties, ” where winds have an infinite
fetch and the West Wind Drift (also known as the CircumAntarctic Current) continuously transports water around the
globe driven by the unimpeded Southern Hemisphere westerlies. Elsewhere, continents act like walls blocking winddriven east-west motion and force the surface ocean to move
north and south along continental margins.
The combined result of zonal winds and flow constraints
imposed by continents is the formation of large surface 8yres
in each ocean basin, representing essentially closed current
loops or rings. Wind motion and the Coriolis effect produce
clockwise subtropical gyres in the Northern Hemisphere and
counterclockwise subtropical gyres in the Southern Hemisphere (Fig. 10.9). Each gyre includes four major surface currents: two east-west currents driven by the zonal wind belts
(e.g., the trades, westerlies, and polar easterlies) forming the
northern and southern limbs of a given gyre, and two northsouth boundary currents that flow parallel or subparallel to
164
JAMES C. INGLE, JR.
the adjacent continental margins. Surface currents in the Indian Ocean are more complex due to changes in seasonal
wind patterns associated with the monsoonal climate in its
northern reaches, and due to the fact that this ocean is located largely in the Southern Hemisphere and bounded by
the West Wind Drift. Nevertheless, north and south equatorial currents and a subtropical gyre also characterize the Indian Ocean. The major exception to this general pattern of
closed surface gyres is the West Wind Drift, or CircumAntarctic Current, which represents unconstrained surface
flow around Antarctica.
Polar regions represent special cases because of their unusual geography and freezing temperatures. A slow but sustained west-flowing gyre prevails in the Arctic Ocean under
the infiuence of the polar easterlies, as demonstrated by studies of drifting ice. Polar easterlies in the Southern Hemisphere
also drive a west-flowing current around the margin of Antarctica, the East Wind Drift in contrast to the dominant and
much larger eastward-flowing West Wind Drift (Figs. 10.2
and 10.9).
The trade winds, prevailing westerlies, and polar easterlies
in both hemispheres are all responsible for sustaining major
east- and west-flowing surface curr nts. The north and south
equutoriuZ~cumnts in the Pacific, A ffantic, and Indian oceans
are clearly the product of the trade winds in these regions.
These latter wind systems drive’ equatorial water westward
until it reaches a continental margin, where it is deflected
north or south and eventually encounters the prevailing
westerlies. Sigrrincantly, some of this water returns eastward
in the form of equatorial cquntercur~ents that flow in the narrow zone between the prevailing trade wind belts and beneath the doldrums, driven by west-to-east pressure gradients
rather than by wind (Fig. 10.9). In addition, some equatorial
water moves eastward within submerged undercurrents such
as the Pacinc Equatorial Undercurrent located just beneath
the North Equatorial Current. Although there is some “leakage” of surface waters across the equatorial regions of the
major ocean basins, most significantly in the Atlantic Ocean,
the Coriolis-controlled clockwise and counterclockwise circulation of northern and southern hemispheric gyres generally separates Northern and Southern Hemisphere surface circulation regimes. This is not the case in the deeper ocean (as
discussed in Chap. 11).
East-west surface currents are characterized by relatively
slow and steady velocities of between 3 to 6 kilometers per
day. However, the dimensions of the ocean’s surface currents
dictate that they transport enormous volumes of water regardless of whether they are moving fast or slow. Indeed, the
volumes moved by these currents are so large that a special
flow unit named the swdrup (after the famous oceanographer Harold U. Sverdrup) is applied to their measurement. A
sverdrup (Sv) represents a flow of 1 million cubic meters of
seawater per second. Although relatively slow, east-west currents such as the North Pacific Current and the North Atlantic Current flow at rates of 10 to 16 Sv.
Eastern and western boundary currents represent the north-
south limbs of the principal surface gyres and display subtle
to exaggerated western intensification due to the Earth’s eastward rotation, the conservation of angular momentum, the
effect of west-blowing trade winds, and the consequent
pileup of water against continental “walls” forming the westem margins of the ocean basins. These currents are responsible for transporting enormous volumes of warm tropical
water poleward and bringing cool water equatorward from
higher latitudes, as emphasized by the position of the 20°C
isotherm on opposite sides of the Pacific and Atlantic oceans
(Fig. 10.5). One need only compare the hot and muggy summers experienced by the inhabitants of Tokyo (35’41’ N) with
the cool and foggy summers of San Francisco (37”47’ N) to
grasp the impact of boundary current asymmetry on the climate of adjacent continental margins. Because of the westward intensification of flow within individual current gyres,
western boundary currents are characterized by high velocities (2 to 5 kilometers per hour) and relatively narrow and
deep profiles, as typified by the Gulf Stream in the western
North Atlantic Ocean and the Kuroshio Current in the westem North Pacific Ocean. In contrast, eastern boundary currents such as the California Current in the Pacific and the
Canary Current in the Atlantic, are commonly slow moving
(0.1 to 2 kilometers per hour) and have wide, shallow profiles. Despite the differences in speeds and cross-sectional
geometries of western and eastern boundary currents, they
transport approximately the same amount of water, thus
maintaining continuity of flow within a given gyre.
The temptation is to view the eastern and western boundary currents as gigantic rivers; however, these flows are not
confined to rigid and fixed channels. Although they are
bounded on one side by a solid continental margin, their
seaward and subsurface boundaries are simply other water
masses. Hence, the tracks of these currents change in shape
and position with variations in speed and volume of flow.
For example, the Gulf Stream experiences exaggerated meandering as it jets past Cape Hatteras, in turn creating is+
lated rings or patches of warm Gulf Stream water that sephrate from the main current and take on a life of their own for
up to several months or a year before mixing with surrounding water. Similarly, studies of the Kuroshio and California
currents demonstrate that they also meander and form large
eddies in response to seasonal and long-term changes in
winds and climate. Satellite monitoring has revealed that
these m e s o s c a l e features of surface currents are far more dynamic than previously understood and have practical importance for both local weather forecasting and fisheries predictions.
Although surface wind stress is the primary driving force
maintaining surface circulation, other factors enter into this
process. No one would argue with the concept that water
runs downhill, and that the interaction of wind, gravity, variations in density of seawater, and the Coriolis effect combine to enhance horizontal surface circulation through the
creation of “hills” and “valleys” on the ocean’s surface. In
addition, surface winds in some areas force the density-
ATMOSPHERE-OCEAN
stratified ocean to do something it fiercely resists - move
vertically. There is more to wind-driven surface circulation
than wind alone.
EKMAN SPIRAL AND EKMAN TRANSPORT
We have already noted that as the wind blows across the sea
surface, there is a decrease in velocity of the water with depth
due to frictional effects. As the wind drags the thin veneer of
water at the air-sea interface, momentum is lost in transferring energy and motion to the next layer below and so on
down through the water column until a point of essentially
zero motion is reached, on average a depth of approximately
100 meters at mid-latitude locations. At the same time that a
given layer is moving horizontally, it is also under the influence of the Coriolis effect and therefore is deflected slightly
to the right of its motion (in the Northern Hemisphere),
leading to a systematic change in the direction of flow with
depth. The combined result of these two processes is the
Ekman
named after the Swedish scientist V. W. Ekman,
who first quantitatively described this important interaction
in 1905 and initiated modern concepts of wind-driven ocean
circulation (Fig. 10.10).
Ekman demonstrated that under the influence of a steadily blowing wind and given a homogeneous column of water
(e.g., an ocean of uniform density and viscosity), surface wa-
Idealized view of the wind-driven Ekman spiral
and Ekman transport within the surface layer of the ocean (in
the Northern Hemisphere). T h e lengths of the solid arrows depict the frictional decrease in velocity with depth; the directions
of the arrows illustrate the results of the Coriolis effect on the
motion of each succeeding layer down through the column. The
Coriolis effect accounts for a 45” angle (to the right of motion in
the Northern Hemisphere) between the direction of the wind
and the direction of the wind-driven surface current, whereas
the net drift of the entire surface layer is at 90” to the right of
the wind. m = meters. (Source: Adapted from P. R. Pinet, Oceanography, West Publishing Company, 1992.)
COUPLING
165
ter moves at an angle of 45” to the right of the wind’s motion
in the Northern Hemisphere and 45” to the left of the wind
in the Southern Hemisphere. As momentum is lost through
friction with depth and as each layer is deflected farther to
the right or left, an end point is reached at which an extremely weak current is moving in the opposite direction to
that of the surface motion. This latter depth averages approximately 100 meters and marks the base of the surface or
mixed layer. Summing the individual vectors of each layer
yields a net direction of motion for the column of 90” to the
right or left of the prevailing wind. Thus, under ideal conditions, the net horizontal motion of the entire wind-driven surface
layer (approximately 0 to 100 meters) is perpendicular to the
direction of the wind (Fig. 10.10). This motion is commonly
referred to as Ekman transport or Ekman drift. Although actual
measurements of wind and surface current vectors deviate
from these ideal or theoretically constant angles, the Ekman
spiral relationship offers a powerful predictive tool for dealing with the dynamics of the surface ocean and has special
significance for understanding gyre circulation and upwelling. The largest deviations from the ideal Ekman spiral relationship occur in shallow areas of the ocean, such as over
continental shelves where frictional dissipation along the sea
bed occurs. On the other hand, angular relationships predicted by the Ekman spiral effect commonly prevail in open
ocean areas, allowing oceanographers confidently to forecast
and hindcast motion of the surface layer based on wind direction and speed.
Figure 10.10.
Windw
0
r
Surface
layer
100 m
1
--+-----/
/ / No wind-driven motion
I /
GEOSTROPHIC CURRENTS AND DYNAMIC
TOPOGRAPHY
As the trade winds and westerlies blow across the North Pacific Ocean, they not only set in motion the North Equatorial
Current and the North Pacific Current, but also result in the
net drift of the surface layer at 90” to the right of their motion, representing a clear example of Ekman transport. This
process slowly moves warmer, less
Map view
dense surface water toward the center
of the North Pacific subtropical surWind
face gyre, creating an area of converdirection
gence. Keeping in mind that specific
volume (measured in cubic centimeters per gram) is the inverse of density
(measured in grams per cubic centimeter), the less dense water takes up
more space or volume than relatively
transport
Ll
higher-density water (Fig. 10.11).
(Northern
Hence, as Ekman transport forces
hemisphere)
warm, less dense water into the center
of the gyre, it stands at a slightly
higher elevation than the surrounding sea surface, forming a large, low
hill or mound. One can easily envision the relationship between den-
JAMES C. INGLE, JR.
166
Steep dynamic
(Northern
1.0260
Figure 10.11. Aspects of dynamic topography of
the (A) sea surface, geostrophic flow, and geohemisphere)
strophicsurface circulation (in the Northern Hemisphere). (A) The idealized motion of a particle of
(A)
water on a large dynamic “hill” in the-center of an
oceanic gyre as it attains a position of perfect bal“Hill” composed of “lightest” or lowest ance between (1) wind-driven uphill motion imdensity water which takesup more
parted by Ekman transport (Coriolis effect) to the
right of west-blowing trade winds, and (2) downhill
motion induced by gravity (pressure) and the slope
of the “hill.” When a position of balance is
achieved, the particle (and others of similar density)
Reference
travels to the right and around the dynamic “hill”
level (surface
representing a geostrophic flow or current. Dynamic topography of the sea surface is created when
1.0250
1.02 0
1.0250
1.0255
1.0260
Average density of column(g/cm3) determined
from data on temperature& salinity
m
wind-driven Ekman transport pushes warmer and
less dense water toward the centers of an oceanic
gyre. (B) As shown in this hypothetical cross section, the higher specifk volume of the less dense
water causes these water masses to stand at a
slightly higher elevation above a given reference
level than surrounding waters of lower density. g/
cm3 = g-rams per cubic centimeter. (C) An example
of the dynamic topography of the North Pacific
Ocean in terms of height differences measured in
dynamic centimeters above a reference level of 500
decibars (db) (the decibar represents the standard
measure of pressure in the ocean and is defined as
100,000 dynes per-square centimeter). The eastward
rotation of the Earth has caused the dynamic “hill”
of warm surface water to shift westward, in turn
dictating that the steepest dynamic slopes and
highest velocities of geostrophic flow occur along
the western Pacific margin.
Dynamic sea surface height
in dynamic cm relative to 500 db reference level
0
sity, specific volume, and sea surface elevation by selecting
an arbitrary reference depth (i.e., a level or depth of equal
pressure) and comparing water columns of different character above the given reference level (Fig. 10.11). A column of
warm, lower-density water will obviously take up more space
(volume) and hence stand higher above the reference level
than will an adjacent column of cooler, denser water.
Satellite-based measurements have coniirmed that the
mounds of water marking the centroids of gyre circulation in
fact stand as much as 2 meters higher than the level of the
ocean forming the margins of a given gyre. Although the
slopes of these giant “hills” are very gentle, they result in
pressure gradients with water moving downhill toward areas
of lower pressure. In effect, the high-standing mound of water responds to the horizontal pressure gradient and attempts
to “flatten out” the sea surface. Individual water particles are
acted upon by gravity pulling them down the “hiI1” at the
same time that the wind-induced Coriolis “force” is pushing
them uphill toward the center of the gyre as a result of Ekman transport (Fig. 10.11). As gravity, density, and the horizontal pressure gradient cause a particle to move downhill,
the Coriolis effect again acts to deflect motion to the right
(in the Northern Hemisphere). The particle thus moves down
and around the hill until reaching an elevation where the
effects of gravity (and density) and the uphill Coriolis “force”
are in precise balance. When balance is achieved at a given
elevation or position on the slope, the particle then travels
continually around the hill with other particles of similar
density, forming an integral part of an ensuing surface current.
Currents and gyre motion generated in this manner are
termed geostrophic currents and geostrophic circulation, respec-
ATMOSPHERE-OCEAN
tively. The word “geostrophic” literally means “Earth
turned” and refers to the fact that the motion of the water is
largely controlled by the Earth’s rotation and the Coriolis
effect, in balance with the effects of density and gravity. The
variations in the topography or elevation of the sea surface
brought about by these interactions, and which in turn govern geostrophic currents, are logically termed dynamic topography. Using data on variations in temperature, salinity, and
density, and an arbitrary reference level (commonly a depth
representing an equal pressure of 500, 1000, or 1500 decibars), oceanographers routinely construct contour maps of
the dynamic topography of the sea surface (Fig. 10.11). The
resulting patterns of dynamic topography can then be used
with great confidence to predict current flow from the orientation and shape of contour lines and the appropriate Coriolis effect (left or right), and the estimated velocity of a given
current can be derived by knowing the slope of the sea surface - the steeper the slope, the faster the current. Areas of
closely spaced dynamic contours depict steep slopes and
hence high-velocity currents, whereas areas of widely spaced
contours correspond to low slopes and hence slow-current
motion.
Because of the Earth’s rotation from west to east, the centers of major surface gyres and the associated elevated
mounds of less dense water are shifted toward the western
sides of individual ocean basins. Because of this shift of mass
and momentum within the gyres, the steepest dynamic topographies are found along the western margins of the surface gyres leading to the so-called western intensification of
currents and typified by the Gulf Stream and Kuroshio currents (Figs. 10.9 and 10.11). In contrast, the eastern margins
of the gyres exhibit low dynamic slopes and relatively sluggish current flow.
Surface winds not only maintain the ocean’s dynamic topography and the resulting geostrophic circulation but also
govern the location and magnitude of one the most important physical processes of the surface ocean - upwelling.
UPWELLING
AND
DOWNWELLING
Upwelling represents the vertical movement of subsurface wa-
ter to the surface of the ocean, commonly from depths
within the upper thermocline or pycnocline layers. This process is critical to the recycling of key nutrients in the ocean
(e.g., phosphorous, nitrogen, and silicon) and their transport
to the surface of the sea, where they can be utilized by phytoplankton through photosynthesis. Upwelling areas are
characterized by exceptionally high rates of primary biologic
productivity along with secondary productivity by the grazers through top carnivores - in fact, the entire food chain drawn to these areas. Predictably, upwelling areas support
rich and important fishing industries. A 1969 study estimated
that upwelling zones account for 50 percent of all marine
fish production in the world despite the fact that they constitute less than 1 percent of the surface area of the global
ocean. Although more recent studies have lowered this esti-
COUPLING
167
mate to 25 percent, these remarkable numbers emphasize the
importance of upwelling to the biologic health of the ocean.
Coastal upwelling can be triggered when local winds blow
in an offshore direction for a long enough period to push
surface water away from the coast, allowing deeper water to
move upward to replace it. In some areas, advective collision
of two surface water masses causes upwelling, such as occurs
where the Oyashio Current meets the Kuroshio Current off
northern Japan (Fig. 10.9). Vertical motion and upwelling
can also occur where a surface current flows over a shallow
submerged bank or seamount or where a surface current
flows past a large coastal prominence. The most significant
upwelling processes in both coastal zones and the open
ocean involve wind-driven Ekman transport (Fig. 10.10).
Where winds blow parallel with a coastline, Ekman transport can induce either upwelling or downwelling (Fig. 10.12).
As water in the surface layer is moved horizontally (at 90” to
the direction of the wind) it is replaced by water from below,
commonly from depths of 100 to 300 meters, within the
upper part of the thermocline or intermediate layer. If Ekman
transport is in the offshore direction, upwelling results. Alternatively, if transport of the surface layer is toward the coast,
downwelling occurs due to the wall-like effect of the continental margin. Ekman coastal upwelling and associated
zones of high biologic productivity are common along the
western coasts of continents (i.e., the eastern sides of ocean
basins) where sustained seasonal winds blow north and
south. Clear examples of these settings include the Pacific
coasts of North and South America and the Atlantic coast of
South Africa and adjacent Namibia. All three regions experience seasonal winds that cause vigorous upwelling of cold,
nutrient-rich water during the spring and early summer seasons, in the turn producing fog and cool weather during
these periods. Mark Twain’s much paraphrased statement
that “the coldest winter he ever experienced was a summer
in San Francisco” neatly sums up this interplay between upwelled cold water and the overlying atmosphere.
Upwelling also takes place away from coastal regions in
the open ocean where the directions of wind and current
motion together with the Coriolis effect cause Ekman drift in
opposing directions, allowing water to well up from below.
This type of wind-induced vertical motion is termed divergent
upwelhng, a process that characterizes the equatorial region
between the northern and southern trade wind belts (Fig.
10.12). Ekman transport forces surface waters to the north
(to the right of the west-blowing trade winds) in the
Northern Hemisphere and to the south (to the left of westblowing trade winds) in the Southern Hemisphere. The result
is a north-south divergence of surface layer motion away
from the equator and the upwelling of subsurface water in
the intervening area.
Divergent upwelling also takes place in the CircumAntarctic region in the area between the east-flowing West
Wind Drift and the west-flowing East Wind Drift (Polar Current). In this case, northward Ekman transport associated
with the east-flowing West Wind Drift is in opposition to the
166
JAMES C. INGLE, JR.
southward Ekman drift associated with the west-flowing East
Wind Drift, resulting in the so-called Antarctic Divergence, a
prominent zone of high primary productivity. Upwelling in
this region is also assisted by the offshore flow of winds from
the Antarctic continent. In addition, upwelling of intermediate water takes place in this area to compensate for the
sinking of the extremely cold and dense surface water produced by the freezing of seawater around the Antarctic margin, a density-driven process.
Surface
Layer
Coastal upwelling
Wind
QUESTIONS
1. Humans have been using the ocean for exploration, trade,
and harvesting of marine resources for thousands of years (not to
speak of the ocean’s past and future role in global politics and
war). What advantages to human well-being can
Coastal
downwelling
you ascribe to
an increased knowledge of ocean circulation?
m
2. List as many fundamental differences as you can think of
between eastern and western boundary currents, and their effects on the adjacent continental margins, including their cli-
Trade winds
mates and cultures.
quator
3. Describe the surficial conditions of the Earth, assuming
the ocean (and the rest of the hydrosphere) had been removed
some 100 million years ago.
Divergent
upwelling
(0
Figure 10.12. Schematic illustrations of (A) coastal upwelling,
(B) coastal downwelling, and (C) open ocean divergent upwelling (in the Northern Hemisphere). All three types of vertical
motion occur as a function of wind-driven Ekman transport (see
Fig. 10.10). As winds blow equatorward (south) along the western side of a continent in the Northern Hemisphere, Ekman
transport forces surface water seaward away from the coast (at
90” to the right of the wind direction). This latter water is in turn
replaced by nutrient-rich intermediate water upwelled from below the surface layer, triggering high primary and secondary
productivity. Conversely, the direction of Ekman transport is
reversed if winds blow poleward in the same area, with the result
that surface water undergoes downwelling (sinking) as it is forced
against the coast. Equatorial divergent upwelling results where
west-blowing northern trade winds induce northward Ekman
transport at the same time that parallel southern trade winds (in
the Southern Hemisphere) are inducing southerly Ekman transport. Thus, subsurface water is upwelled in the zone between the
two opposing surface flows.
4. What changes in global ocean surface circulation would
you predict might take place if the Isthmus of Panama were
removed, allowing Pacific and Atlantic-Caribbean water and circulation to be connected along the equator? This is not a moot
question; the Isthmus of Panama was in fact not present prior to
3 million years ago.
5. What do you think the surface circulation of the ocean
would look like if the Earth were rotating westward rather than
eastward?
F U R T H E R
Bearman,
READINO
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G. (ed.) 1989. Ocean circulation. Oxford: The Open
University and Pergamon Press.
Garrison, T. 1999. Oceanography, an
ence,
invitation to marine wia
3rd edition. Belmont, CA: Wadsworth Publishing Corn*
PanY.
McLeisch, W. H. 1989. The blue god. Smithsonian 19(2):
4rl-
58.
Pickard, G. L., and Emery, W. J. 1990. Descriptive physicvl
oceanography, 5th (SI)
.
enlarged edition. Oxford. Perganttrrt
Press.
Wunsch, C. 1992. Observing ocean circulation from spice
Oceanus 35(2) 9-17.
:
.Js
J:._