Potter T.D., Colman B.R. (co-chief editors). The handbook of weather, climate, and water: dynamics, climate physical meteorology, weather systems, and measurements
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SECTION 2
THE CLIMATE SYSTEM
Contributing Editor: Robert E. Dickinson
CHAPTER 9
OVERVIEW: THE CLIMATE SYSTEM
ROBERT E. DICKINSON
The climate system consists of the atmosphere, cryosphere, oceans, and land
interacting through physical, chemical, and biological processes. Key ingredients
are the hydrological and energy exchanges between subsystems through radiative,
convective, and fluid dynamical mechanisms. Climate involves changes on seasonal,
year-to-year, and decadal or longer periods in contrast to day-to-day weather
changes. However, extreme events and other statistical measures are as, or more,
important than simple averages. Climate is seen to impact human activities most
directly through the occurrence of extremes. The frequency of particular threshold
extremes, as, for example, the number of days with maximum temperatures above
100
F, can change substantially with shifts in climate averages.
1 THE ATMOSPHERE
The atmospher e is described by winds, pressures, temperatures, and the distribution
of various substances in gaseous, liquid, and solid forms. Water is the most impor-
tant of these substances. Also important are the various other radiatively active
(‘‘greenhouse’’) gases, including carbon dioxide and liquid or solid aerosol particu-
lates. Most of the mass of the atmosphere is in the troposphere, which is comprised
of the layers from the surface to about 12 km (8 km in high latitudes to 16 km at the
equator) where the temperature decreas es with altitude. The top of the troposphere is
called the tropopause. Overlying this is the stratosphere, where temperatures increase
with altitude to about 50 km or so (Fig. 1). The tropospheric temperature decreases
with altitude are maintained by vertical mixing driven by moist and dry convection.
Handbook of Weather, Climate, and Water: Dynamics, Climate, Physical Meteorology, Weather Systems,
and Measurements, Edited by Thomas D. Potter and Bradley R. Colman.
ISBN 0-471-21490-6 # 2003 John Wiley & Sons, Inc.
119
The temperature increases with altitude in the stratosphere in response to increasing
heating per the unit mass by ozone absorption of ultr aviolet radiation. The variation
of temperature structure with latitude is indicated in Figure 2. The troposphere is
deepest in the tropics because most thunderstorms occur there. Because of this depth
and stirring by thunderstorms, the coldest part of the atmosphere is the tropical
tropopause. In the lower tropo sphere temperatures generally decrease from the
equator to pole, but warmest temperatures shift toward the summer hemisphere,
especially in July. Longitudinally averaged winds are shown in Figure 3. Because
of the geostrophic balance between wind and pressu res, winds increase with altitude
where temperature decreases with latitude. Conversely, above about 8 km, where
temperatures decrease toward the tropical tropopause, the zonal winds decrease
with altitude. The core of maximum winds is referred to as the jet stream. The jet
stream undergoes large wavelike oscillations in longitude and so is usually stronger
at a given latitude than in its longitudinal average. These waves are especially
noticeable in the winter hemisphere as illustrated in Figure 4.
2 GLOBAL AVERAGE ENERGY BALANCE
Solar radiation of about 342 W=m
2
entering Earth’s atmosphere is absorbed and
scattered by molecules. The major gaseous absorbers of solar radiation are water
vapor in the troposphere and ozone in the stratosphere. Clouds and aerosols likewise
Figure 1 Main zones of the atmosphere defined according to the temperature profile of the
standard atmosphere profile at 15
N for annual-mean conditions (Hartmann, 1994).
120
OVERVIEW: THE CLIMATE SYSTEM
scatter and absorb. Clouds are the dominant scatterer and so substantially enhance
the overall planetary reflected radiation, whose ratio to incident solar radiation, about
0.31, is referred to as albedo. Thermal infrared radiation, referred to as longwave,is
controlled by clouds, water vapor, and other greenhouse gases. Figure 5 (Kiehl and
Trenberth, 1997) illustrates a recent estimate of the various terms contributing to the
global energy balance. The latent heat from global average precipitation of about
1.0 m per year is the dominant nonradiative heating term in the atmosphere.
Because of the seasonally varying geometry of Earth relative to the sun, and the
differences in cloudiness and surface albedos, there are substantial variations in the
distribution of absorbed solar radiation at the surface and in the atmosphere, as
likewise in the transfer of latent heat from the surface to the atmosphere. This
heterogeneous distribution of atmospheric heating drives atmospheric wind systems,
either directly or through the creation of available potential energy, which is utilized
Figure 2 Zonal mean meridional cross sections of temperature during two seasons. (a)
December 1978–February 1979 and (b) June–August 1979. Height scale is approximate and
the contour interval is 5 K (Grotjahn, 1993).
2 GLOBAL AVERAGE ENERGY BALANCE 121
Figure 3 Meridional cross sections of longitudinally averaged zonal wind (top panels, m=s)
for DJF (Holton, 1992).
Figure 4 Mean 500-mb contours in January, Northern Hemisphere. Heights shown in tens
of meters (Holton, 1992).
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OVERVIEW: THE CLIMATE SYSTEM
to maintain random occurrences of various kinds of instabilities, such as thunder-
storms and wintertime cyclonic storm systems. These dynamical systems hence act to
redistribute energy within the atmosphere and so determine the distributions of
temperature and water vapor. Likewise, the balances at the surface between fluxes
of radiative, latent, and thermal energies determine surface temperatures and soil
moistures. The properties of the near-surface air we live in are determined by a
combination of surface and atmospheric properties, according to processes of the
atmospheric boundary layer. Thus climatic anomalies in surface air may occur either
because of some shift in atmospheric circulation patter ns or through some modifica-
tion of surface properties such as those accompanying deforestation or the
development of an urban area.
3 THE ATMOSPHERIC BOUNDARY LAYER
The term boundary layer is applied in fluid dynamics to layers of fluid or gas,
usually relatively thin, determining the transition between some boundary and the
rest of the fluid. The atm ospheric boundary layer is the extent of atmosphere that is
mixed by convective and mechanical stirring originating at Earth’s surface. Such
stirring is commonly experienced by airplane travelers as the bumps that occur
during takeoff or landing, especially in the afternoon, or as waves at higher levels
in flying over mountainous regions. The daytime continental boundary layer, extend-
ing up to several kilometers in height, is most developed and vigorously mixed,
being the extent to which the daytime heating of the surface drives convective
overturning of the atmosphere. The land cools radiatively at night, strongly stabiliz-
Figure 5 Earth’s annual global mean energy budget based on the present study. Units are
W=m
2
(Kiehl and Trenberth, 1997).
3 THE ATMOSPHERIC BOUNDARY LAYER 123
ing the atmosphere against convection, but a residual boundary layer extends up to
about 100 m stirred by the flow of air over the underlying rough surface. This diurnal
variation of fluxes over the ocean is much weaker and the boundary layer is of
intermedi ate height. The temperature of the atmosphere, when stirred by dry
mixing, decreases at a rate of 9.8 K=km. Above the boundary layer, temperatures
decrease less rapidly with height, so that the atmosphere is stable to dry convection.
A layer of clouds commonly forms at the top of the daytime and oceanic boundary
layers and contributes to the convection creating the boundar y layer through its
radiative cooling (convection results from either heating at the bottom of a fluid
or cooling at its top). Also, at times the clouds forming near the top of the boundary
layer can be unstable to moist convection, and so convect upward through a deep
column such as in a thunderstorm.
4 ATMOSPHERIC HYDROLOGICAL CYCLE
The storage, transport, and phase changes of water at the surface and in the atmo-
sphere are referred to as the hydrological cycle. As already alluded to, the hydro-
logical cycle is closely linked to and driven by various energy exchange processes at
the surface and in the atmosphere. On the scale of continents, water is moved from
oceans to land by atmospheric winds, to be carried back to the oceans by streams
and rivers as elements of the land hydrological cycle. Most of the water in the
atmosphere is in its vapor phase, but water that is near saturation vapor pressure
(relative humidity of 100%) converts to droplets or ice crystals depending on
temperature and details of cloud physics. These droplets and crystals fall out
of the atmosphere as precipitation. The water lost is replenished by evapora-
tion of water at the surface and by vertical and horizontal transport within the
atmosphere. Consequently, much of the troposphere has humidities not much
below saturation. Saturation vapor pressure increases rapidly with temperature
(about 10% per kelvin of change). Hence, as illustrated in Figure 6, the climatolo-
gical concentrations of water vapor vary from several percent or more when going
from near-surface air to a few parts per million near the tropical tropopause. Water
vapor concentrations in the stratosphere are close to that of the tropical tropopause,
probably because much of the air in the lower stratosphere has been pumped through
the tropical tropopause by moist convection.
5 CLIMATE OF THE STRATOSPHERE
The dominant radiative processes in the stratosphere are the heating by absorption of
solar ultra violet (UV) radiation and cooling by thermal infrared emi ssion from
carbon dioxide and, to a lesser extent, ozone molecules. The stratospheric absorption
of UV largely determines how much harmful UV reaches the surface. Ozone in the
upper troposphere and lower stratosphere additionally adds heat by absorption of
thermal emission from the warmer surface and lower layers. The stratosphere,
124 OVERVIEW: THE CLIMATE SYSTEM
Figure 6 Zonal mean cross sections of the specific humidity in g=kg for annual, DJF, and JJA mean
conditions. Vertical profiles of hemispheric and global mean values are shown on the right (Peixoto and Oort,
1992).
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