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Chapter 2
CHAPTER 2
2.1
2.2
2.3
2.4
2.5
2.6
2.7
2.8
2.9
Global Climatology
GLOBAL CLIMATOLOGY .......................................................................................... 1
INTRODUCTION............................................................................................................................. 2
ANNUAL GLOBAL MEAN ENERGY BUDGET .................................................................................... 2
ATMOSPHERIC GENERAL CIRCULATION ........................................................................................ 4
CONTINENTS AND OCEANS ......................................................................................................... 10
SEASONS .................................................................................................................................... 13
CLIMATE ZONES ......................................................................................................................... 18
CLIMATE AND VEGETATION........................................................................................................ 21
TABLES ...................................................................................................................................... 23
FIGURE LEGENDS ....................................................................................................................... 24
Ecological Climatology
2.1 Introduction
Climate is the average weather in any particular region during a given month or season. It is often
characterized in terms of long-term (e.g., 30-year) mean temperature, precipitation, wind, humidity, and
cloudiness. However, climate can change from year to year. Some years are hotter than normal; some years
are colder than normal. Some years are wet while others are dry. Climate also changes over longer
timescales of centuries and millennia. Some 18 000 years ago much of the world was in the grips of a
prolonged ice age in which temperatures were several degrees colder than present and ice sheets several
kilometers thick covered much of the northern regions of the world. This chapter and the subsequent two
chapters introduce climate, climate variability, and climate change. Earth’s surface has distinct climate
zones that range from hot and wet (tropical), hot and dry (desert), cool and moist (temperate), and cold
(polar). This chapter focuses on the geographic and seasonal differences in climate, which arise primarily
due to the uneven geographic heating of Earth by the Sun, the resultant transport of heat by the atmosphere
and oceans, and the nature of Earth’s surface (Hartmann 1994; Ahrens 1998, 2000). Chapter 3 examines
climate variability at seasonal-to-interannual timescales. Chapter 4 considers climate change over longer
timescales and the internal feedbacks within the climate system that accentuate and mitigate climate
change.
2.2 Annual global mean energy budget
Electromagnetic radiation from the Sun is the source of energy that drives Earth’s climate. All
matter with a temperature above -273.15 °C radiates energy in the form of electromagnetic radiation. The
Stefan-Boltzmann Law relates emitted energy to temperature as
E = εσ (T + 273.15)
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where E is the rate of emitted energy per unit area (J s-1 m-2 = W m-2), T is the temperature of the object in
°C, σ = 5.67 × 10-8 W m-2 K-4 is a constant, and ε is emissivity. For a blackbody, ε = 1. Most objects are
not blackbodies and have emissivity of about 0.95 to 0.98. The term (T + 273.15) converts temperature in
degrees Celsius to the Kelvin scale. Objects with a higher temperature emit radiation at a greater rate than
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objects with a lower temperature. This radiation travels at the speed of light in the form of waves with
peaks and troughs. The distance between successive peaks (or troughs) is called the wavelength. The
wavelength of emitted energy depends on the temperature of the material. The higher the object’s
temperature, the shorter the wavelength of emitted radiation. The Sun, with a temperature of 6000 K, emits
radiation in short wavelengths between 0.2 × 10-6 and 4 × 10-6 m (0.2 to 4 µm). Maximum emission occurs
at a wavelength of about 0.5 µm. Solar radiation is divided into ultraviolet radiation with a wavelength less
than 0.4 µm (containing 10% of the Sun’s energy), visible radiation between 0.4 and 0.7 µm (40% of the
Sun’s energy), and near-infrared radiation at wavelengths greater than 0.7 µm (50% of the Sun’s energy).
Visible radiation is further divided into violet, blue, green, yellow, orange, and red (in order of increasing
wavelength). Earth, with an effective temperature of about 15 °C, emit less radiation than the Sun and in
longer wavelengths from 4 µm to 100 µm. Peak emission occurs at a wavelength of 10 µm. This radiation
is called longwave, terrestrial, or infrared radiation.
Annually, approximately 342 W m-2 of solar radiation is received at the top of the atmosphere
(Figure 2.1). Twenty percent (67 W m-2) is absorbed by the clouds, gases, and aerosols in the atmosphere.
An additional 168 W m-2 (49%) is absorbed by Earth’s surface. The remainder, 107 W m-2 (31%), is
reflected back to space by materials in the atmosphere and on the surface. The absorbed solar radiation
warms Earth, which emits longwave radiation. At the surface, Earth emits 390 W m-2 of longwave
radiation. Most of this radiation (350 W m-2) is absorbed by clouds, water vapor, carbon dioxide, and other
gases in the atmosphere; only 40 W m-2 escape to space. The gases, particles, and other material suspended
in the atmosphere emit longwave radiation. This radiation travels in all directions with some being lost to
space and some reaching Earth’s surface. A total of 195 W m-2 is lost to space, which together with the 40
W m-2 from the surface balances the 235 W m-2 solar radiation absorbed by the atmosphere and surface.
That is, the net radiation absorbed by Earth equals zero; the solar radiation absorbed by Earth is equal to
the longwave radiation emitted to space.
Although Earth as a whole balances solar and longwave radiation at the top of the atmosphere, the
geographic distribution of radiation is unequal. Latitudes near the tropics, between 30° S and 30° N,
generally absorb more than 275 W m-2 solar radiation annually; latitudes closer to the poles absorb less
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Ecological Climatology
radiation (Figure 2.2). The major exception to this is North Africa, where a large portion of the solar
radiation is reflected back to space by the bright desert soils. Tropical latitudes, because they are warmer,
generally emit more longwave radiation than high latitudes. Annual outgoing longwave radiation at the top
of the atmosphere is generally greater than 250 W m-2 between latitudes 30° S and 30° N (Figure 2.2).
Regions of high precipitation, which have low longwave fluxes because of the deep, cold clouds associated
with precipitation, are an exception. Between latitudes 30° S and 30° N, this heat loss is less than the heat
gained from solar radiation. In general, there is an annual excess of solar radiation gain over longwave
radiation loss in the tropics and a deficit at latitudes polewards of 35° to 40° (Figure 2.2).
The latitudinal gradient in net radiation results in an equator-to-pole temperature gradient. Low
latitudes, which gain radiation, are warmer than high latitudes, which lose radiation. However, if only
radiative processes determined temperatures, the tropics would be 30 °C warmer than they actually are and
polar regions would be 40 °C colder than they actually are (Kiehl 1994). Instead, the uneven geographic
distribution of the Sun’s radiation on Earth produces winds, set in motion by differences in air pressure,
that reduce the poleward temperature gradient by carrying heat from tropical latitudes to polar regions.
2.3 Atmospheric general circulation
Air pressure is the force exerted over a given area by the movement of air molecules. Air pressure
is also a measure of the mass of air above a given point. Change in mass per unit area (∆m, kg m-2) is
related to change in pressure (∆P, pascals [1 Pa = 1 Newton m-2 = 1 kg m-1 s-2]) as
∆m = −∆P / g
where g = 9.81 m s-2 is gravitational acceleration. At sea level, where the standard atmospheric pressure is
1013.25 hectopascals (hPa, 1 hPa = 100 Pa = 1 millibar), a column of air covering one square meter of
surface area and extending upwards to the top of the atmosphere (i.e., to a pressure of 0 Pa) has a mass of
10 329 kg. Most air molecules are held near the surface by the force of gravity. With greater height in the
atmosphere, the number of air molecules, and pressure, decreases (Figure 2.3). Fifty percent of the mass of
the atmosphere is below a height of about 5500 m.
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Winds are created by horizontal gradients in air pressure. For example, a column of air with a
surface pressure of 1020 hPa has a mass of 10 400 kg per square meter of surface area. Suppose a nearby
column of air, separated by a distance of 100 km, has a lower surface pressure of 1012 hPa (10 320 kg m2
). The difference in mass between the two columns is 80 kg m-2, which creates a pressure gradient of 8
hPa per 100 km. The resulting pressure gradient force causes wind to blow from high pressure to low
pressure. This force is directed from high to low pressure at right angle to lines of equal pressure or
isobars. Closely spaced isobars indicate steep pressure gradients, strong forces, and strong winds; widely
spaced isobars indicate weak pressure gradients and weak winds.
Differences in air temperature can create horizontal pressure gradients that initiate air movement.
Molecules in warm air move fast and spread apart; the air becomes less dense and the column expands
vertically. Conversely, molecules in a column of cold air move slowly and become dense; the column
shrinks. This is an expression of the ideal gas law, which relates pressure (P) and volume (V) to
temperature (T) and mass (M) as
PV = MRT
where R is a constant. If pressure remains constant, any increase in temperature must result in a
corresponding decrease in density ( M / V ). Likewise, a decrease in temperature results in an increase in
density. In other words, a short column of cold, dense air exerts the same surface pressure as a tall column
of warm air.
The result is that the mass of air – its pressure – decreases more rapidly with height in a cold
column of air than in a warm column of air. Consider the two columns of air in Figure 2.4 (left). Both have
a total mass of 10 400 kg m-2 and a surface pressure of 1020 hPa. However, because the cold air column is
denser than the warm air, more air molecules are closer to the ground than in the warm air column. For
example, in the cold column air molecules below height A have a combined mass of 2000 kg m-2. In the
warm column, where air molecules spread farther apart, the mass is only 1500 kg m-2. Because both
columns of air have the same total mass, the number of molecules (i.e., the mass of air) above height A is
less in the cold air (8400 kg m-2) than in the warm air (8900 kg m-2). And because there is less mass above
height A in the cold column, air pressure (i.e., the mass of air above height A) is lower than in the warm
air column. This is true at other heights (B, C). The mass of air above height B is only 4400 kg m-2 in the
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Ecological Climatology
cold air but 6400 kg m-2 in the warm air. At this height, the air pressure increases from 431 hPa in the cold
air to 628 hPa in the warm air. In other words, the mass of air and pressure decrease more rapidly with
height in a column of cold air than in a column of warm air.
For any given height above the surface, a column of cold air has fewer air molecules, less mass,
and lower air pressure than a column of warm air. Hence, warm air is generally associated with high air
pressure aloft and cold air is associated with low air pressure aloft. Differences in temperature, by creating
a horizontal pressure gradient, initiate wind flow from high to low pressure. In Figure 2.4 (right), both
columns of air initially have the same surface pressure. However, the horizontal difference in temperature
between the two columns creates a horizontal pressure difference aloft. The pressure gradient force then
causes air to flow from high pressure (warm column) to low pressure (cold column). As air aloft leaves the
warm column, the mass of air in the column decreases and its surface pressure decreases. Conversely, the
addition of air to the cold column increases its mass and surface pressure. As a result, high surface pressure
develops under the cold column and low surface pressure develops under the warm column. Surface winds
blow from high pressure (cold air) to low pressure (warm air), closing the atmospheric circulation.
This simple model of thermally driven atmospheric circulation begins to explain the geographic
redistribution of heat by atmospheric winds. Tropical regions, because they gain radiation, are hot and
develop high pressure aloft. Polar regions, because they lose radiation, are cold and develop low pressure
aloft. In response to this pressure gradient force, warm tropical air flows towards the poles aloft. A broad
band of surface low pressure develops in the tropics, while the poles, where air converges aloft, develop
high surface pressure. In response to this surface pressure gradient, cold polar air flows over the surface
towards the equator and completes the atmospheric circulation.
This single cell, thermally driven circulation does not fully explain atmospheric circulation. Winds
are the balance of three forces acting simultaneously: the pressure gradient force, the Coriolis force, and
friction. The pressure gradient force is the primary cause of air movement. If the pressure gradient force
were the only force acting on air, winds would always blow directly from high to low pressure. However,
the Coriolis force deflects air as it moves. The Coriolis force is the apparent motion caused by Earth’s
rotation, which pushes winds to the right of their intended path in the Northern Hemisphere and to the left
in the Southern Hemisphere. The Coriolis force acts at right angle to the wind’s intended direction,
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influencing its direction but not its speed. The Coriolis force is greater for strong winds than for weak
winds and is zero along the equator and greatest at the poles.
The top panel in Figure 2.5 shows how these two forces determine wind speed and direction aloft
at heights typically greater than 1000 m, where the influence of surface friction is negligible. The pressure
gradient force is perpendicular to the isobars in the direction of low pressure. This force accelerates a
parcel of air northwards towards the low pressure. As it begins to move, the Coriolis force deflects the air
towards the right, curving its path. As the path changes, so does the Coriolis force, which is always
directed at right angle to the direction of motion. As the parcel of air increases in speed, the magnitude of
the Coriolis force increases. Eventually wind speed and direction are such that the pressure gradient force,
acting from high to low pressure, is balanced by the Coriolis force acting in the opposite direction. At this
point, the net force acting on the air is zero and the wind flows parallel to the isobars at a constant speed.
This flow of air is called geostrophic wind. In the Northern Hemisphere, geostrophic wind blows with low
pressure to the left and high pressure to the right with a speed directly related to the pressure gradient. In
the Southern Hemisphere, the flow is reversed because the Coriolis force deflects winds to the left of their
intended path.
Near the surface, winds encounter trees, buildings, hills, and other objects. The frictional drag of
these objects slows wind speed and in doing so deflects the direction of motion. Surface winds do not blow
parallel to the isobars but rather cross them moving from high to low pressure. Figure 2.5 illustrates the
difference in wind between the surface, where friction is important, and aloft, where friction is negligible.
The upper air wind is geostrophic, flowing parallel to the isobars with the pressure gradient force balanced
by the Coriolis force. At the surface, friction, which acts counter to the direction of motion, reduces wind
speed. Because of this, the same pressure gradient produces slower winds at the surface than aloft. With
reduced speed, the Coriolis force is also reduced, and the weaker Coriolis force no longer balances the
pressure gradient force. The winds are not deflected as much towards the right of their intended path and
winds blow across isobars towards low pressure.
The deflection of surface winds by friction creates vertical motions in the atmosphere. In the
Northern Hemisphere, winds blow in a clockwise direction around high pressure cells and
counterclockwise around low pressure cells. In the Southern Hemisphere, where the Coriolis force deflects
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winds to the left, winds blow clockwise around low pressure cells and counterclockwise around high
pressures. Friction deflects surface winds in towards the center of the low pressure (Figure 2.6). As the
surface air moves inward, the converging air slowly rises, typically to a height of several thousand meters,
and diverges aloft. So long as the upper-level outflow of air balances the inflow of surface air, surface
pressure remains unchanged. However, if the upper-level divergence exceeds the surface convergence,
surface pressure will decrease, the pressure gradient will increase, and surface winds will strengthen. In
contrast, surface winds blow outward from the center of a high pressure cell. Air from above converges
and descends to replace the diverging surface air.
The preceding discussion has shown that: differences in air pressure create winds; winds blow
from high to low pressure; the Coriolis force deflects winds to the right of their intended path in the
Northern Hemisphere and to the left in the Southern Hemisphere; and surface winds spiral into low
pressure cells, creating upward vertical motions, but spiral out of high pressure cells, creating downward
air motions. These facts can now be put together to understand atmospheric winds.
The uneven geographic heating of Earth by the Sun, with warmer temperatures in the tropics than
in polar regions, creates a pressure gradient that initiates winds aloft and along the surface. Figure 2.7
illustrates the general circulation of the atmosphere as a system of high and low surface pressure regions
arising from the unequal heating of the surface. In contrast to the single cell model, each hemisphere has
three cells that redistribute heat. In the Northern Hemisphere, warm tropical air with high pressure aloft
flows as upper-level wind from the equator towards the North Pole. As it moves north, the Coriolis force
deflects the wind to the east. This poleward moving air cools as it moves northwards. The column shrinks,
the air becomes denser, and surface pressure increases. The converging masses of air moving from the
tropics to middle latitudes further increases surface pressure. This produces a belt of high surface pressure
at about latitude 30° N. Air flows along the surface back towards the equator from this high pressure, being
deflected to the west by the Coriolis force. These are the northeasterly trade winds. This tropical cell is
known as the Hadley circulation. At the North Pole, cold air causes low pressure aloft while warmer air at
latitude 60° N creates high pressure aloft. This warm air flows towards the pole as southwesterly wind aloft
(northward moving air deflected to the east by the Coriolis force) and back as the polar easterlies along the
surface (southward moving air deflected to the west by the Coriolis force). A third cell in the middle
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latitudes between latitudes 30° N and 60° N connects these tropical and polar cells. The surface high
pressure at latitude 30° N creates subsiding motion and surface divergence. One branch of this air spreads
southwards as trade winds. The other branch flows northwards, being deflected to the east in middle
latitudes. These are the mid-latitude westerlies. At about latitude 60° N, this relatively warm air clashes
along the polar front with cold air moving towards the equator. Similar circulations develop in the
Southern Hemisphere.
Figure 2.7 shows several prominent features characterize the circulation of the atmosphere. Semipermanent low surface pressure systems develop at the equator and latitude 60°, where air rises. Semipermanent high surface pressure systems develop at the poles and the subtropics at latitude 30°, where air
subsides. Between latitudes 30° N and 30° S, surface trade winds converge along the equator from the
northeast in the Northern Hemisphere and the southeast in the Southern Hemisphere in a broad region
known as the intertropical convergence zone. Between latitudes 30° and 60° in both hemispheres, surface
winds are westerly. At about latitude 60°, these winds clash with polar easterlies along the polar front.
Figure 2.8 shows the actual geographic distribution of surface pressure and surface wind for
January and July. The three principal features are: high pressure cells at about 30° latitude in both
hemispheres; lower pressure along the equator between these subtropical high pressure centers; and lower
pressure at high latitudes (60° N, 60° S). Wind flows clockwise around high pressure systems in the
Northern Hemisphere. Thus, in the Northern Hemisphere the northern edges of the subtropical high
pressures drive the mid-latitude westerlies while the southern edges drive the westwardly flowing trade
winds. Conversely, winds flow counterclockwise around low pressure centers, driving the polar easterlies
to the north and the mid-latitude westerlies to the south. In the Southern Hemisphere, flows are in the
opposite direction (counterclockwise around high pressures and clockwise around lows).
The general circulation of the atmosphere accounts for the major climate zones in the world.
Rainfall is abundant where air rises – along the equator – and lacking where air sinks – near latitudes 30° N
and 30° S (Figure 2.9, color plate). In the tropics, the trade winds converge from the northeast in the
Northern Hemisphere and the southeast in the Southern Hemisphere on the equator in the intertropical
convergence zone. This convergence and lifting of warm, moist air leads to high annual rainfall. Many
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tropical regions have wet and dry seasons. This occurs because of seasonal variation in the geographic
location of the convergence zone. Regions of subsidence, as occurs in high pressure systems, generally
have low rainfall. Many of the world’s major deserts – in southwestern United States, North Africa,
southern South America, South Africa, and western Australia – are located on the eastern flanks of the
subtropical high pressures near latitudes 30° N and 30° S. Rainfall is also high, though not as high as in the
tropics, in the middle latitudes between 40° and 60° where warm moist air clashes with cold air along the
polar front. Temperatures are generally distributed in latitudinal bands with warmest temperatures in the
tropics and progressively colder temperatures towards the poles (Figure 2.10, color plate). This reflects
latitudinal variation in net radiation and redistribution of energy from atmospheric circulations.
2.4 Continents and oceans
It is evident from Figure 2.8 that the Northern and Southern Hemispheres differ in circulation.
The Southern Hemisphere has contiguous bands of high pressure at latitude 30° S and low pressure at
latitude 60° S, as expected from Figure 2.7. These bands are intermingled in the Northern Hemisphere.
This difference arises because of the different distribution of land in the two hemispheres: 70% of all land
is in the Northern Hemisphere. Maximum land area is between latitudes 40° N and 75° N, where more than
50% of Earth’s surface area is land. In the Southern Hemisphere, land is generally less than 25% of the
surface area. Between latitudes 40° S and 65° S, there is little land. Continents heat and cool faster than
oceans. In January, when northern continents are colder than oceans, high pressure systems form over land;
low pressure systems are most pronounced over the northern regions of the Pacific and Atlantic Oceans.
The opposite pattern occurs in summer when northern continents are warmer than oceans. Strong high
pressures develop in the North Pacific and North Atlantic; low pressures develop over Asia and North
America.
Continents also affect climate through the presence of high mountain ranges. The Rocky
Mountains in western North America and the Andes in western South America form major north-south
mountain ranges (Figure 2.11). Prominent west-east mountains are the Alps in Europe and the Himalayas
in Asia. Mountains affect regional climate because air cools at a rate of 0.5 to 1.0 °C for every 100 m
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increase in elevation. As air rises over the mountains and cools, the amount of water it can hold decreases.
Clouds form and rain falls. Consequently, windward sides of large mountains often receive much more
rainfall than leeward sides.
Proximity to oceans also affects temperature. Water has a moderating influence on climate,
preventing extreme fluctuations in temperature that arise in interior regions of the continents. The large
heat capacity of oceans creates a thermal inertia. Heat is stored in the summer and released in winter,
damping summertime warming and wintertime cooling. In addition, water is a fluid and carries heat from
one region to another. Like the atmosphere, the general circulation of the ocean transports heat from the
tropics to polar regions. In the Northern Hemisphere, the maximum energy transported by oceans is
comparable to atmospheric heat transport (Trenberth and Solomon 1994). In the Southern Hemisphere,
oceanic heat transport is reduced by about one-half and is only about one-half of the atmospheric heat
transport. Oceanic heat transport is the result of two types of circulations: wind-driven surface currents and
the density-driven thermohaline circulation.
As winds blow over the oceans, they push surface water in the same direction as the air. The trade
winds blowing from the southeast in the Southern Hemisphere and northeast in the Northern Hemisphere
set in motion westward-flowing equatorial currents to the north and south of the equator. When these
currents reach the western edge of the Pacific and Atlantic basins, they are deflected and flow as western
boundary currents to the north in the Northern Hemisphere and to the south in the Southern Hemisphere.
Between latitudes 30° and 60° in both hemispheres, the prevailing westerly winds blow surface water east
across the ocean basins. At the eastern edge of the oceans, this water is deflected as eastern boundary
currents to the south in the Northern Hemisphere and to the north in the Southern Hemisphere. This
circulation of surface water, from equatorial current to western boundary current to westerly current to
eastern boundary current, creates circular flows in ocean basins known as gyres (Figure 2.12). These gyres
are centered in the subtropics at latitudes 30° N in the North Atlantic and the North Pacific and at 30° S in
the South Atlantic, South Pacific, and Indian Ocean. These gyres rotate clockwise in the Northern
Hemisphere and counterclockwise in the Southern Hemisphere. At high latitudes, polar easterlies drive
surface currents in a westward direction to produce subpolar gyres that rotate in a direction opposite to the
adjacent subtropical gyres. These subpolar gyres are best developed in the Atlantic east of Greenland and
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Ecological Climatology
in the Weddell Sea off Antarctica. In the Southern Hemisphere, the strong surface westerlies near latitude
60° S create an eastward-flowing Antarctic circumpolar gyre.
Western boundary currents are strong poleward-flowing currents that carry warm water from the
tropics to middle latitudes. The most prominent are the Kuroshio current along the western boundary of the
North Pacific and the Gulf Stream current along the western boundary of the Atlantic off the United States
coast. These currents push warm water to the northeast across the Pacific and Atlantic basins, respectively.
As a result, water off western and northern Europe is much warmer than water off eastern North America
at similar latitudes (Figure 2.10, color plate). Likewise, water off southeast Alaska and western Canada is
warmer than corresponding water north of Japan. Proximity to these warm currents creates a milder climate
than would be expected from latitude. In contrast, eastern boundary currents transport cold water towards
the tropics. Prominent eastern boundary currents are the California current off the United States west
coast, the Peru current off the west coast of South America, the Canary current off northern Africa, and the
Benguela current off southern Africa.
In addition to surface currents, deep ocean currents arise from density differences in water. This
circulation, called the thermohaline circulation, arises because variations in the salinity and temperature of
water create differences in density. The density of water increases as temperature decreases or salinity
increases. Cold, salty water is more dense than warm, fresh water. Water involved in the thermohaline
circulation is initially formed in polar regions at the surface. In the Antarctic Ocean, surface water is cooled
to below freezing by the cold overlying air. When the water freezes, it forms a layer of sea ice several
meters thick. Salt is excluded from the ice as it freezes, and the unfrozen water under the sea ice becomes
saltier. The combination of cold, salty water results in very dense water that sinks, encircling Antarctica
and flowing northwards along the ocean bottom. Similar masses of cold, dense deep water form in the
Arctic Ocean off the coast of Greenland, where it flows southwards along the ocean bottom into the North
Atlantic. Figure 2.13 illustrates the full thermohaline circulation. In the North Atlantic, warm surface water
flowing northwards becomes more dense as it cools and evaporation increases surface salinity. The dense
water sinks and returns southward along the ocean bottom. Near the southern tip of Africa, this deep water
joins the Antarctic bottom water. This river of bottom water spreads into the Indian and Pacific Oceans,
where it is the dominant supply of deep water. In these oceans, surface waters are too warm (Indian) or not
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salty enough (Pacific) to allow for sinking of surface water. The deep water slowly rises and begins its long
journey along the surface back to the North Atlantic. Changes in the thermohaline circulation, by altering
the transport of heat by oceans, can cause large and rapid climate changes.
2.5 Seasons
Seasonal variation in climate arises from the geometry of Earth’s annual orbit around the Sun and
its daily rotation on its axis. As Earth rotates over the course of a day, the Sun appears to sweep a broad arc
through the sky. Its position is defined by two angles (Figure 2.14): its altitude above the horizon; and its
bearing on the horizon, which is called the azimuth angle. A third angle, the zenith angle, is often used
instead of altitude. Whereas altitude is the angular distance above the horizon, zenith angle is the angular
distance from a line perpendicular to the surface (Figure 2.14). Altitude is zero at sunrise and sunset and is
maximum at solar noon. At solar noon, the Sun is due south on the horizon in the Northern Hemisphere; in
the Southern Hemisphere, the Sun is due north. In the morning, the Sun is east of south; it is west of south
in the afternoon.
The Sun’s position in the sky is important because it determines the intensity of radiation. A
surface receives the most solar radiation when it is perpendicular to the Sun’s rays. At other angles, the
Sun’s radiation is spread over a larger surface area leading to less radiation per unit area. For example,
when the Sun is directly overhead a unit beam of radiation is spread over a unit surface area (Figure 2.15).
But when the zenith angle is 60° (i.e., the altitude angle is 30°), the same unit beam of radiation is spread
over twice as much surface area; the energy per unit area has been reduced by one-half. The ratio of
radiation received on a horizontal surface (SH) to that received on a surface perpendicular to the Sun’s rays
(SP) decreases with greater zenith angle (Z) as
S
H
S
P
= cos Z
The altitude and azimuth angles vary depending on latitude, time of day, and day of year. Earth
rotates over the course of a day on an axis that is tilted at an angle of 66.5° to an imaginary line connecting
Earth and the Sun (Figure 2.16). This axis is always pointed at the same point in space. Consequently, as
Earth moves in its orbit around the Sun, the direction of tilt relative to the Sun varies. In boreal winter
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(austral summer), the Northern Hemisphere is tilted away from the Sun while the Southern Hemisphere is
tilted towards the Sun. In boreal summer (austral winter), the Northern Hemisphere is tilted towards the
Sun while the Southern Hemisphere is tilted away from the Sun. Seasonal changes in orientation towards
or away from the Sun are seen in the declination angle. Declination is the angle between a line connecting
the centers of the Sun and Earth and the plane of the equator (Figure 2.16). It is the latitude where the Sun
is directly overhead at solar noon.
Declination angle varies through a year (Figure 2.17). Four days of the year have special
significance. On the winter solstice, which occurs on or about December 21, the declination angle is –
23.5o. The Sun is directly overhead (i.e., has a zenith angle of zero) at noon at latitude 23.5° S. This
latitude is the farthest point south of the equator where the Sun is directly overhead and is called the Tropic
of Capricorn. On this day, which is the shortest day of the year in the Northern Hemisphere, no solar
radiation is received north of latitude 66.5° N. This latitude defines the Arctic Circle. South of latitude
66.5° S, daylight is continuous. This latitude defines the Antarctic Circle. The Sun is at its lowest point on
the horizon in the Northern Hemisphere, beginning to rise again on December 25. On the summer solstice,
on or about June 21, the noon Sun is directly overhead at latitude 23.5° N. This latitude defines the Tropic
of Cancer – the farthest point north of the equator where the Sun is directly overhead at noon. This day is
the longest day of the year in the Northern Hemisphere. Daylight is continuous north of 66.5° N and absent
south of 66.5° S. On the equinoxes (on or about March 21 and September 22), the noon sun is directly
overhead at the equator (i.e., the declination is zero). All latitudes receive 12 hours of light. The Sun rises
due east at all latitudes on this day.
Zenith angle and altitude angle vary with latitude, time of year, and time of day as
cos Z = sin a = sin φ sin δ + cos φ cos δ cos h
where Z is zenith angle, a is altitude angle, φ is latitude, δ is solar declination, and h is the solar hour angle.
The hour angle measures time of day. Because Earth rotates on its axis 360° in 24 hours, one hour is
equivalent to 15° of longitude, and the hour angle varies by 15° for each hour before or after solar noon.
For example, the hour angle is 15° at 1100 hours and -15° at 1300 hours; it is 60° at 0800 hours and -60° at
1600 hours. At solar noon, when h = 0°and cos h = 1 , the zenith angle is at a minimum and the altitude
14
Chapter 2 – Global Climatology
angle is at a maximum. Hence, the Sun’s radiation onto a horizontal surface, ignoring for the moment
attenuation for clouds, dust, and the atmosphere, is at a maximum at solar noon. At this time, the zenith
angle is simply latitude minus declination, Z = φ − δ . Solar noon occurs at 1200 hours only for the
longitude that defines the time zone within which the observer is located. East of this longitude, solar noon
occurs earlier in the day; west of this longitude, solar noon occurs later.
The solar radiation reaching Earth varies over the course of a day and the course of a year due to
orbital geometry and the Sun’s apparent motion in the sky. The amount of solar radiation on a horizontal
surface, not accounting for atmospheric attenuation, is
S
H
(
)
2
= S / r cos Z
c v
where Sc = 1367 W m-2 is the solar constant, rv is the radius vector squared, and Z is the solar zenith
2
angle. The cosine of the zenith angle adjusts the radiation that would be received on a surface
perpendicular to the solar beam to that received on a horizontal surface (Figure 2.15). The radius vector
accounts for the fact that as Earth moves in an elliptic orbit around the Sun, taking about 365¼ days to
complete one revolution, Earth is slightly closer to the Sun in January than in July (Figure 2.16). Since the
strength of the Sun’s radiation decreases with the square of the distance from the Sun, Earth receives
slightly more radiation in winter than summer. This distance, expressed as a fraction of the mean Earth-Sun
distance, is called the radius vector and is at a minimum in early January. At this time, the Earth receives
about 3.5% more solar radiation relative to the average Earth-Sun distance. It is at a maximum in early
July, when Earth receives about 3.25% less solar radiation. However, seasonal changes in radiation due to
Earth’s tilt are much larger. Indeed, if Earth’s axis of rotation were perpendicular to the plane of rotation
(i.e., no tilt), there would be no seasons since all points would be illuminated an equal amount of time
throughout the year.
Several important geographic patterns arise in the Northern Hemisphere as a result of Earth-Sun
geometry. First, the Sun is higher in the sky and rises and sets further from south in summer than in winter
(Figure 2.18). Second, altitude angle decreases with northern latitudes. Third, daylength, defined as the
period during which the Sun is above the horizon, decreases with northern latitudes in winter and increases
with northern latitudes in summer (Figure 2.19). Tropical latitudes have relatively little seasonal variation
15
Ecological Climatology
in daylength. Seasonal variation increases with higher latitudes in the Northern and Southern Hemispheres.
The most extreme seasonal variation occurs in polar regions, where the Sun is below the horizon in winter
and never sets in summer.
Temporal change in solar radiation over the course of a year drives seasonal changes in climate
(Figure 2.20). From December 21 to June 21, as declination increases from –23.5° to 23.5°, zenith angle
decreases, days become longer days, and more solar radiation is received in the Northern Hemisphere.
From June 21 to December 21, declination decreases, days get shorter, and radiation decreases in the
Northern Hemisphere. Latitudes near the equator have relatively little seasonal variation in solar radiation.
At higher latitudes, the amplitude of the seasonal cycle increases. The most extreme seasonal cycle occurs
near the poles, where the Sun is below the horizon until the equinoxes (March 21, September 22) and
where the long summer days result in much radiation. Figure 2.20 also shows a strong poleward decrease
in solar radiation during winter, and a much weaker equator-to-pole temperature gradient in summer.
These seasonal changes in solar radiation drive the seasonal heating and cooling of the
atmosphere. Air temperatures generally lag by one month compared with solar radiation so that although
the least solar radiation is received in December the coldest month is typically January (Figure 2.21, color
plate). Temperatures generally follow latitudinal bands, with warmest temperatures in the tropics and
colder temperatures at high latitudes. In the Northern Hemisphere, which is experiencing winter, the
poleward temperature gradient is particularly strong. For example, temperatures over land decrease from
20 °C along latitude 20° N to 0 °C along latitude 40° N. At latitude 60° N, temperatures are extremely
cold, ranging from -20 to –30 °C. July is typically the warmest month. In July, the Northern Hemisphere
south-to-north temperature gradient is weakened as the Sun’s declination is north of the equator and the
latitudinal gradient in solar radiation diminishes. The tropics and subtropics, between latitudes 30° S and
30° N, show the least seasonal variation in temperatures. Temperatures in this region are generally between
20 °C and 30 °C throughout the year. In contrast, middle and high latitudes have pronounced seasons, with
temperatures below freezing in winter and generally from 10 °C to 20°C in summer.
The moderating influence of oceans on temperature is evident in Figure 2.21 (color plate). This is
particularly clear in January, when temperatures are colder in the interior regions of North America and
16
Chapter 2 – Global Climatology
Eurasia than at similar latitudes near oceans. Moreover, the continents have a larger seasonal temperature
range than corresponding oceans. For example, while the monthly temperature range in interior North
America is about 25 °C, the temperature range in the North Atlantic Ocean is only about 10 °C. Land has a
much lower heat capacity than water so that the continents cool faster than oceans in winter and warm
more in summer.
As a result of these seasonal changes in solar radiation and temperature, the major surface high
and low pressure systems migrate to different locations. In winter, there are two massive subarctic low
pressure systems centered in the northern regions of the Atlantic and Pacific Oceans (Figure 2.8) – the
Icelandic and Aleutian lows, respectively. Two high pressure systems – the Bermuda high in the Atlantic
and the Pacific high off the California coast – are found in subtropical latitudes near 30° N. High pressure
cells form over Asia and North America due to the intense cooling of land. During summer, land warms
and the continental high pressure systems are replaced by low pressure. The subpolar Icelandic and
Aleutian lows, so strong in winter, are barely noticeable. Instead, the subtropical Bermuda and Pacific
highs dominate.
Seasonal changes in solar radiation and temperature also alter winds aloft. Winds can be
partitioned into their south-to-north (meridional) and west-to-east (zonal) components. Figure 2.22 shows
the zonal wind as a function of height. The high winds in the middle latitudes of each hemisphere are the
jet streams. Jet streams are fast moving westerly air in the upper atmosphere between 10 km to 15 km aloft.
They are located in middle latitudes along the fronts that separate cold polar air from warm tropical air,
where the horizontal temperature and pressure gradients are greatest. The subtropical jet stream is centered
at latitude 30o, near the polar extremity of the tropical cell. A second jet – the polar jet – is located farther
poleward at the edge of the polar cell. The sharp poleward contrast in temperature produces a rapid
horizontal pressure gradient that intensifies winds along the front. Because the poleward temperature
gradient is strongest in winter and weakest in summer, the jet streams show seasonal variation. Winds are
strongest (greater than 40 m s-1) and dip further from the poles in winter. In summer, the jets are weaker
and centered at higher latitudes.
These seasonal changes in surface pressure and atmospheric circulation drive seasonal changes in
precipitation (Figure 2.23, color plate). In the tropics the intertropical convergence zone, where the
17
Ecological Climatology
convergence of the trade winds from the northeast in the Northern Hemisphere and southeast in the
Southern Hemisphere produces rising motion and heavy rainfall, migrates from south of the equator in
boreal winter (austral summer) to north of the equator in boreal summer (austral winter). For example,
tropical regions of South America and Africa between the equator and latitude 10° S receive more rainfall
from November through April than from May through October. Southeast Asia has a strong summer
monsoon created by the thermal difference between land and adjacent oceans. In winter, the high pressure
over Asia drives air out towards the oceans (Figure 2.8). In summer, as the continental low pressure
intensifies, warm, moist ocean air is drawn into the continent producing the wet summer monsoon.
Seasonal changes in precipitation related to the seasonal migration of the semi-permanent high
and low pressure systems are also evident in the United States (Figure 2.23, color plate). The East Coast
tends to receive equal amounts of precipitation throughout the year. However, Florida, with its subtropical
climate, receives more than 1000 mm of rain from May through October, but less than 500 mm from
November to April. The Plains states tend to receive more precipitation in the warm months of May
through October than in November to April. This is related to the westward displacement of the Bermuda
high, which brings in moist air out of the Gulf of Mexico. The West Coast has a pronounced seasonality to
its precipitation. From Washington to Southern California, much more rain is received from November to
April than in the rest of the year. Much of California receives less than 125 mm of rain in the dry season.
This is related to the wintertime influence of the Aleutian low, which brings in moisture-laden air from the
Pacific. In summer, the strong high pressure cell off the West Coast leads to subsidence, which inhibits
rainfall, and to northerly flow along the coastline. The dominating influence of the Bermuda high on the
precipitation climatology of eastern United States is clearly seen in the bands of precipitation that decrease
with distance from the Gulf of Mexico (Figure 2.9, Figure 2.23, color plates).
2.6 Climate zones
Although no two places experience exactly the same climate, geographically similar regions have
similar climates. For example, there are four broad temperature zones (Figure 2.24, color plate). Tropical
temperatures are found near the equator where monthly temperatures are warmer than 18 °C. Polar
18
Chapter 2 – Global Climatology
temperatures are found near the poles, where the warmest month of the year is colder than 10 °C. In
between are the middle latitude climates with both warm and cold seasons. Temperate regions have more
than four months with temperatures warmer than 10 °C; subarctic regions have less than four months of
warm temperatures. Similarly, there are five major precipitation regions (Figure 2.24, color plate). Tropical
regions along the equator receive abundant rainfall year-round. Other tropical regions have pronounced
wet and dry seasons caused by seasonal migration in the intertropical convergence zone. Middle latitudes
are generally moist. Arid regions are located along the subtropical high pressures at latitude 30° in both
hemispheres and in regions far removed from sources of atmospheric moisture. Polar regions are generally
dry because the cold air holds little moisture.
These temperature and precipitation zones can be combined into generalized climate zones. Figure
2.25 (color plate) illustrates one such climate classification – the Köppen classification as modified by
Trewartha (Ahrens 1998, pp. 353-375; Bailey 1996, 1998). This scheme utilizes five major climate zones
with dominant characteristics:
1.
Humid Tropical Climate: warm year-round; coldest month 18 °C or warmer.
2.
Dry Climate: deficient precipitation throughout the year.
3.
Moist Subtropical Mid-Latitude Climate: warm to hot summers with mild winters; coldest
month above 0 °C but below 18 °C; warmest month above 10 °C.
4.
Moist Continental Climate: warm summers and cold winters; coldest month below 0 °C;
warmest month above 10 °C.
5.
Polar Climate: extremely cold winters and cold summers; warmest month below 10 °C.
Each of these climate zones has sub-zones defined by temperature and precipitation (Table 2.1).
Tropical humid climates occur where temperatures are warm throughout the year. A tropical
rainforest climate occurs where rainfall is abundant throughout the year; a tropical savanna climate is found
where there is a pronounced dry season (Figure 2.26). In both climate zones, seasonal variation in air
temperature is minimal – less than a few degrees. In the tropical rainforest zone, rainfall is abundant
throughout the year but may vary seasonally with the position of the intertropical convergence zone.
Tropical rainforest climates occur in hot, wet equatorial regions of South America, Africa, southeast Asia,
19
Ecological Climatology
and Indonesia (Figure 2.25, color plate). The tropical savanna climate is found in tropical regions that are
warm year-round but have a pronounced dry season An example climate is given in Figure 2.26, which
illustrates the distinguishing feature between tropical rainforest and tropical savanna climates: a
pronounced dry season. At this site, the rainy season is from April to October. Little rainfall occurs in
December, January, and February, when the region is under the influence of a subtropical high pressure.
The double peak in summer precipitation reflects the seasonal migration of the intertropical convergence
zone across the equator. Temperatures are warmer during the dry season than in the rainy season when
clouds cool the surface. Tropical savanna climates occur in Central America, to the north and south of the
Amazon Basin in South America, to the north and south of the Congo Basin in Africa, east Africa, parts of
India and Southeast Asia, and northern Australia (Figure 2.25, color plate).
Dry climates, which occur where rainfall is sparse throughout the year, are divided into semi-arid
and arid climates based on moisture deficiency (Figure 2.27). The semi-arid climate occurs in temperate
regions – most prominently in the Great Plains of the United States, the steppes of Central Asia, and parts
of southern South America, southern Africa, and Australia (Figure 2.25, color plate). An example climate
is given in Figure 2.27 (top). In this particular location, most precipitation falls in April and May.
Temperatures are hot during June through August when the clear sky and intense solar radiation heat the
surface. The arid climate of deserts is not only dry, but also hot as solar radiation readily penetrates the
clear, dry skies (Figure 2.27). Desert climates occur on the eastern flanks of the subtropical high pressures
near latitudes 30° N and 30° S – in southwestern United States, North Africa, southern South America,
South Africa, and western Australia (Figure 2.25, color plate). They are also found in continental areas of
the middle latitudes that are far removed from sources of atmospheric moisture – central Asia, central
Australia, and the Great Basin of western United States.
Moist subtropical middle latitude climates occur in regions with distinct summer and winter
seasons, where summers are warm to hot and winters are mild. They are divided into Mediterranean, humid
subtropical, and marine zones (Figure 2.28). Mediterranean climates occur where a summer dry season is
pronounced – in southern California and along coastal areas of the Mediterranean Sea (Figure 2.25, color
plate). These climates are characterized by mild, moist winters and hot, dry summers (Figure 2.28, top).
Humid subtropical climates occur in southeastern United States, eastern China, Japan, and along the
20
Chapter 2 – Global Climatology
southeastern coasts of South America, Africa, and Australia (Figure 2.25, color plate). These climates are
located on the western edge of subtropical high pressure areas, which drives warm, moist tropical air
towards the middle latitudes, and consequently are characterized by hot, muggy summers (Figure 2.28,
middle). Winters are mild and precipitation is abundant throughout the year. Moderate to pronounced
seasonality is a dominant feature of climate. Marine climates occur in the Pacific Northwest region of the
United States, western Europe, and western South America in middle latitudes where oceans moderate
climate (Figure 2.25, color plate). Marine climates are characterized by mild winters, with temperatures
rarely below freezing, cool summers, and abundant precipitation year-round (Figure 2.28, bottom).
Moist continental climates occur in the northern regions of North America, Europe, and Asia
(Figure 2.25, color plate). Large seasonal variation in temperature, with moderate to cool summers and
cold winters, characterizes the climate (Figure 2.29). The humid continental subzone is divided into warm
summer and cool summer regions based on whether the warmest month is above or below 22 °C. Farther
north, in Alaska, northern Canada, northern Europe, and northern Russia, where the winters are bitterly
cold and the summers are cool and short, the climate is subarctic. Precipitation is generally light.
Polar climates are found in high latitudes or mountain tops where the warmest month is below 10
°C (Figure 2.30). Tundra climates are found where plants can still survive in the short summers and long,
cold winters. In the extreme cold of Greenland and Antarctica, little vegetation grows and permanent
glaciers cover the land.
2.7 Climate and vegetation
The natural vegetation of Earth has a distinct geographic pattern that corresponds to climate
zones. Forests grow in tropical rainforest, humid subtropical, marine, and moist continental climates. In
these regions, precipitation is abundant year-round. Trees cannot survive in the bitter cold of tundra
climates. Instead, small shrubs, herbaceous plants, and mosses grow in the short summers. Extensive
grasslands occur in semi-arid and savanna climates of central North America, northern and central South
America, central and southern Africa, central Asia, and Australia. Here climate is hot and dry. Short, dense
woody bushes form chaparral or ‘Mediterranean’ vegetation in the Mediterranean climate where summers
21
Ecological Climatology
are hot and dry and winters are mild and moist. Deserts, with sparse or widely spaced scrubby plants, are
found in arid climates. The close correspondence between climate zones and major vegetation zones is
readily apparent. Many climate zones such as tropical savanna, tropical rainforest, and tundra are named
after vegetation.
The traditional ecological view is one in which climate determines the broad geographic
distribution of vegetation. The traditional atmospheric view is one in which transport of heat drives the
circulation of the atmosphere and oceans and creates climate zones. However, observational studies and the
advent of sophisticated mathematical models of atmospheric physics, atmospheric dynamics, oceanic heat
transport, and surface ecological and hydrological processes have allowed scientists to examine how
different types of vegetation affect climate. Numerous studies beginning in the 1970s all point to the same
conclusion: the geographic distribution of vegetation is an important determinant of regional and global
climate. The correlation between climate zones and biogeography is not merely passive but rather reflects
the outcome of coupled climate-vegetation dynamics in which the geographic distribution of vegetation
affects and is affected by climate. Consequently, natural and human-induced changes in land cover can
alter climate.
22
Chapter 2 – Global Climatology
2.8 Tables
Table 2.1. Definition of climate zones in Figure 2.25 (color plate)
Climate zone
Major characteristics
Humid tropical climate
warm year-round; coldest month 18 °C or warmer
Tropical rainforest
wet all seasons; all months more than 60 mm rain
Tropical savanna
dry season; rainfall in driest month less than 60 mm
Dry climate
deficient precipitation throughout year
Semi-arid
steppe vegetation
Arid
desert
Moist subtropical mid-latitude climate
warm to hot summers with mild winters; coldest
month above 0 °C but below 18 °C; warmest month
above 10 °C
Mediterranean
dry season in summer
Humid subtropical
warmest month above 22 °C; at least 4 months
above 10 °C
Marine
all months below 22 °C; at least 4 months above 10
°C
Moist continental climate
warm summers and cold winters; coldest month
below 0 °C; warmest month above 10 °C
Humid, warm summer
warmest month above 22 °C
Humid, cool summer
warmest month below 22 °C
Subarctic
less than 4 months above 10 °C
Polar climate
extremely cold winters and cold summers; warmest
month below 10 °C
Tundra
warmest month below 10 °C but above 0 °C
Ice cap
warmest month below 0 °C
23
Ecological Climatology
2.9 Figure Legends
Figure 2.1. Earth’s annual mean global energy budget showing solar radiation (left), sensible and latent
heat fluxes (middle), and longwave radiation (right). Data from Kiehl and Trenberth (1997).
Figure 2.2. Annually averaged radiative fluxes as observed by satellite. Fluxes poleward of latitude 60° are
unreliable and are not shown. Top: Annual absorbed solar radiation. Regions absorbing more than 275 W
m-2 are shaded. Middle: Annual outgoing longwave radiation. Regions losing more than 250 W m-2 are
shaded. Bottom: Annual net radiation (absorbed solar minus outgoing longwave). Regions with a net loss
of radiation are shaded. Data from Hurrell and Campbell (1992) and provided by the National Center for
Atmospheric Research (Boulder, Colorado).
Figure 2.3. Relationship between height and air pressure.
Figure 2.4. Pressure gradient and atmospheric circulation resulting from temperature differences. In this
figure, each dot depicts many air molecules with a combined mass per unit area of 100 kg m-2. The
atmosphere is depicted uniformly with height. In fact, mass and pressure decrease rapidly with height. This
is omitted from the figure for simplicity and does not invalidate the general conclusions. Left: Mass and
pressure at several heights for a column of cold air (left) and a column of warm air (right). Mass at the
surface and at heights A, B, and C are the mass of air above each height. Pressure is the corresponding air
pressure. Right: Resulting thermally driven circulation of winds aloft (top) as a result of the pressure
gradient aloft and along the surface (bottom) in response to changing surface pressure.
Figure 2.5. Difference in winds at the surface and aloft in the Northern Hemisphere. Top: Geostrophic
wind aloft is balanced by the pressure gradient force and Coriolis force with motion parallel to isobars.
Bottom: Surface wind is balanced by the pressure gradient force, the Coriolis force, and friction. Friction
slows wind speed and causes wind to cross isobars typically at a 30° angle.
Figure 2.6. Surface winds and vertical air motions in the Northern Hemisphere. Left: Low pressure. Right:
High pressure.
24
Chapter 2 – Global Climatology
Figure 2.7. Idealized atmospheric circulation with resulting surface winds and surface pressure.
Figure 2.8. Mean January (top) and July (bottom) sea level pressure in hectopascals (hPa). High pressure
regions greater than 1015 hPa are shaded in grey. Low pressure regions less than 1015 hPa are unshaded.
Arrows indicate the direction and magnitude of surface wind in meters per second. Data from Trenberth
(1992b) and provided by the National Center for Atmospheric Research (Boulder, Colorado).
Figure 2.9. Mean annual precipitation. The chart on the right shows precipitation averaged around the
world, from longitude 180° W to 180° E, as a function of latitude. Data from Legates and Willmott
(1990b) and provided by the National Center for Atmospheric Research (Boulder, Colorado)..
Figure 2.10. Mean annual air temperature. The chart on the right shows temperature averaged around the
world, from longitude 180° W to 180° E, as a function of latitude. Data from Legates and Willmott (1990a)
and provided by the National Center for Atmospheric Research (Boulder, Colorado)..
Figure 2.11. Map of elevation illustrating the major mountain ranges and highlands at ½ degree spatial
resolution. Data provided by the National Center for Atmospheric Research (Boulder, Colorado).
Figure 2.12. Wind-driven ocean surface currents.
Figure 2.13. Thermohaline circulation showing cold, salty deep water (solid lines) and warm shallow water
(dashed lines).
Figure 2.14. Solar geometry illustrating azimuth, altitude, and zenith angles. The boxed inset shows the
difference between altitude and zenith angles.
Figure 2.15. Effect of zenith angle on solar radiation. Top: A beam of solar radiation with unit width is
spread over twice as much surface area with a zenith angle of 60° than for a zenith angle of 0°. Bottom:
Reduction in solar radiation as a function of solar zenith angle.
25
Ecological Climatology
Figure 2.16. The revolution of Earth around the Sun and the resulting astronomical seasons. The boxed
inset illustrates the declination angle. Variation in the time of the solstices and equinoxes occurs because
Earth’s 365¼ day orbit around the Sun is approximated by 365 days with an extra day every four years.
Figure 2.17. Solar declination angle as a function of day of year.
Figure 2.18. Sun path diagram for latitude 40° N showing azimuth angles at sunrise and sunset and altitude
angle at solar noon. Top: Winter solstice. Bottom: Summer solstice.
Figure 2.19. Daylength as a function of latitude (vertical axis) and day of year (horizontal axis). Dark
shading shows when the Sun never rises. Light shading shows when the Sun is above the horizon for 24
hours. Daylength in hours is equal to
24
π
cos
−1
( − tan φ
tan δ )
Figure 2.20. Daily solar radiation at the top of the atmosphere in relation to latitude (vertical axis) and day
of year (horizontal axis). Units are in megajoules (106 joules) per square meter.
Figure 2.21. Mean air temperature for January (top) and July (bottom). Data from Legates and Willmott
(1990a) and provided by the National Center for Atmospheric Research (Boulder, Colorado).
Figure 2.22. Mean monthly zonal (west-to-east) wind as a function of latitude and height. Height is
expressed in terms of pressure (in millibars) on the left axis and altitude (in kilometers) on the right axis.
One millibar equals 1 hPA. East-to-west winds (negative zonal wind) are lightly shaded. Strong westerly
winds are darkly shaded. Winds are averaged around the world, from longitude 180° W to 180° E. As a
result, the jet streams are seen as single solid cores. Top: January. Bottom: July. Data from Trenberth
(1992b) and provided by the National Center for Atmospheric Research (Boulder, Colorado).
Figure 2.23. Precipitation for November through April (top) and May through October (bottom). Data
from Legates and Willmott (1990b) and provided by the National Center for Atmospheric Research
(Boulder, Colorado).
26
Chapter 2 – Global Climatology
Figure 2.24. Generalized temperature (top) and precipitation (bottom) zones.
Figure 2.25. Earth’s major climate zones. This figure uses Legates and Willmott’s (1990a,b) ½° monthly
temperature and precipitation climatology with the rules of Köppen (Ahrens 1998, p. 413) and Trewartha
(Espenshade 1990, pp. 8-9).
Figure 2.26. Representative humid tropical climates. Top: Tropical savanna, Kintampo, Ghana. Bottom:
Tropical rainforest, Uaupes, Brazil. Data from Griffiths (1972, p. 248) and Schwerdtfeger (1976, p. 270).
Figure 2.27. Representative dry climates. Top: Semi-arid, Denver, Colorado, USA. Bottom: Arid, Phoenix,
Arizona, USA. Data from Bryson and Hare (1974, pp. 277, 268).
Figure 2.28. Representative moist subtropical mid-latitude climates. Top: Mediterranean, Barcelona, Spain.
Middle: Humid subtropical, Atlanta, Georgia, USA. Bottom: Marine, Dublin, Ireland. Data from Wallén
(1970, p. 230), Bryson and Hare (1974, p. 282), and Wallén (1970, pp. 113-114).
Figure 2.29. Representative moist continental climates. Top: Humid continental, warm summer, New York
City. Middle: Humid continental, cool summer, Moscow. Bottom: Subarctic, Fort Smith, Canada. Data
from Bryson and Hare (1974, p. 312), Lydolph 1977, p. 394), and Bryson and Hare (1974, p. 158).
Figure 2.30. Representative polar climates. Top: Tundra, Barrow, Alaska, USA. Bottom: Ice cap,
McMurdo, Antarctica. Data from Bryson and Hare (1974, p. 144) and Orvig (1970, p. 337).
27