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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which is transferred from the surface to the atmosphere as latent heat. The net
radiation at the surface that is not returned to the atmosphere as sensible or latent
heat is stored at the surface. Heat storage is very important for the diurnal cycle of
temperature over land. Soils have a much smaller heat capacity than water.
Consequently, land heats and cools faster than water.
Surface properties determine these energy fluxes and the resulting surface
climate. Forests are ‘‘darker’’, hence absorbing more solar radiation, than grasslands.
Forests are also taller—that is ‘‘rougher’’—than shrubs or grasses and exert more
stress on the fluid motion of the atmosphere. Deserts and shrublands, with their dry
soils, have less evapotranspiration than well-watered forests or crops. Observational
studies and the advent of sophisticated mathematical models of atmospheric physics,
atmospheric dynamics, and surface ecological and hydrological processes have
allowed scientists to examine how these different surfaces affect climate. Numerous
studies all point to the same conclusion: The distribution of vegetation on Earth’s
surface is an important determinant of regional and global climates; consequently,
natural and human-induced changes in Earth’s surface can alter the climate.
Tropical deforestation is one example of the way in which human alterations of
the natural landscape are changing climate. Climate model experiments show that
the replacement of tropical forests with pastures causes a warmer, drier climate.
Desertifica tion, in which deserts encroach into forest landscapes, also results in a
warmer, drier climate. Conversely, vegetation expansion into deserts, as happened
6000 years ago in North Africa, causes a cooler, wetter climate. By masking the high
albedo of snow, the boreal forest creates a warmer climate compared to simulations
in which the boreal forest is replaced with tundra vegetation. In Asia, monsoon
circulations are created by land–sea temperature contrasts. A high land albedo,
such as occurs when there is increased snow cover in Tibet, cools the surface,
decreasing the land–sea temperature contrast and causing less rainfall.
2 SURFACE ENERGY FLUXES AND TEMPERATURE
The radiation that impinges on a surface or object must be balanced by the energy
re-radiated back to the atmosphere, energy lost or gained as sensible and latent heat,
and heat storage. More formally, the energy balance at the surface is
ð1 rÞS#þa L#¼L"þH þ l E þ G
The first term in this equation, ð1 rÞS#, is the solar radiation absorbed by the
surface. S# is the radiation onto the surface and r is the albedo, which is defined as
the fraction of S# that is reflected by the surface. The remainder, ð1 rÞ, is absorbed
by the surface. The second term, a L#, is the atmospheric longwave radiation
absorbed by the surface, where a is the fraction of the incoming radiation L# that
is absorbed. Together, the absorbed solar radiation and longwave radiation comprise
the radiative forcing, Q
a
. This must be balanced by:
136 PROCESSES DETERMINING LAND SURFACE CLIMATE
1. Longwave radiation emitted by the surface ðL"Þ in proportion to its absolute
temperature, in kelvins, raised to the fourth power.
2. Energy transferred to or from the surface by convection (H). This sensible heat
flux is directly proportional to the temperature difference between the surface
and air and inversely proportional to a transfer resistance.
3. Energy used to evaporate water (l E). The latent heat flux is directly
proportional to the vapor pressure difference between the surface and air
and is inversely proportional to a transfer resistance.
4. Energy stored in the soil via conduction (G). When the ground beneath the
surface is colder than the surface, heat is conducted into the ground. At night,
when the surface is colder than the ground, heat is transferred from the ground
to warm the surface.
These energy flows at the surface must be balanced. This balance is maintained
by changing the surface temperature. For example, the temperature of a surface will
rise as more radiation is received on the surface. As a result, more energy is returned
to the atmosphere as longwave radiation and as sensible heat. More energy is stored
in the ground via conduct ion. As the latent heat flux increases, the surface tempera-
ture cools. During the day, when the surface is warmer than the air, this decreases the
sensible heat flux. If the underlying soil is a good conductor of heat, and there is a
large soil heat flux, the surface will not be as hot and the sensible heat flu x will
decrease to compensate for the increased heat storage. Conversely, if the soil is a
poor conductor of heat, little heat will be trans ferred from the surface to the soil and
the surface will be hot.
The importance of these energy fluxes in determining surface temperature can be
illustrated with a simple example. Suppose a leaf has a radiative forci ng of 1000,
700, and 400 W=m
2
, which are representative of values for a clear sky at mid-day, a
cloudy sky at mid-day, and at night, when solar radiation is zero and the surface
receives only longwave radiation. If the only means to diss ipate this energy is
through re-radiation (i.e., H ¼0 and l E ¼0) and there is no heat storage
(G ¼0), the leaf surface would attain temperatures of 91, 60 and 17
C with the
high, moderate, and low radiative forcings (Table 1). When heat loss by convection
TABLE 1 Temperatures of Well-Watered Leaf for Radiative Forcings
a
Temperature (
C)
L: þ HL: þ H þ lE
Q
a
(W=m
2
) L: 0.1 m=s 0.9 m=s 4.5 m=s 0.1 m=s 0.9 m=s 4.5 m=s
1000 91 53 39 34 39 33 31
700 60 40 34 32 32 29 29
400 17 26 28 29 23 26 27
a
Air temperature is 29
C, relative humidity is 50%, and wind speeds are 0.1, 0.9, and 4.5 m=s.
2 SURFACE ENERGY FLUXES AND TEMPERATURE 137
is included, leaf temperature depends on wind speed because the transfer resistance
decreases with high wind speeds. Under calm conditions, with a wind speed of
0.1 m=s, leaf temperature decreases by 38
C with the high radiative forcing and
by 20
C with the moderate forcing (Table 1). Higher wind speeds lead to even
cooler temperatures. At the low radiative forcing, convection warms the leaf because
it is colder than the surrounding air and heat is transferred from the air to the leaf.
Latent heat exchange is also a powerful means to cool a surface because of the large
amount of energy required to evaporate water. For a well-watered leaf, under
calm conditions and high radiative forcing, evapotranspiration cools the leaf an
additional 14
C to a temperature of 39
C (Table 1). Higher winds result in even
lower temperatures.
3 HYDROLOGIC CYCLE
As the preceding example shows, evapotranspiration is an effective means to cool a
surface, particularly at high radiative forcing and low wind speed. The rate of latent
heat loss depends on the amount of water present. A well-watered site has more
water to evaporate than a dry site. Because more energy goes into latent heat rather
than sensible heat, the lower atmosphere is likely to be cool and moist. A dry
surface, on the other hand, has low latent heat flux, high sensible heat flux, and
the air is likely to be warm and dry. Typical values of the Bowen ratio (the ratio of
sensible to latent heat) are: 0.1 to 0.3 for tropical rain forests, where high annual
rainfall keeps the soil wet; 0.4 to 0.8 for temperate forests and grasslands, where less
rainfall causes drier soils; 2.0 to 6.0 for semiarid regions; and greater than 10.0 for
deserts.
The water stored on land (DW) is the difference between water input as precipita-
tion (P) and water loss via evapotranspiration (E) and r unoff (R):
DW ¼ P E R
In many regions, water is stored as snow in winter and not released to the soil
until the following spring. Snow is climatologically important because its high
albedo reflects a majority of the solar radiation. Snow has a low thermal conductiv-
ity, and a thick snow pack insulates the soil. In spring, a large portion of the net
radiation at the surface is used for snow melt, preventing the surface from warming.
Precipitation is greatly modified by vegetation. Some of the rainfall is intercepted by
foliage, branches, and stems. The remainder reaches the ground as throughfall,
stemflow, and snowmelt.
Only a portion of the liquid water reaching the ground surface infiltrates into the
soil. The actual amount depends on soil wetness, soil type, and the intensity of the
water flux. Table 2 shows representative hydraulic conductivity when the soil is
saturated. Sandy soils, because of their large pore sizes, can absorb water at fast
rates. Loamy soils absorb water at slower rates. Clay soils, because of their small
pores, have the lowest hydraulic conductivity. The water that does not infiltrate into
138 PROCESSES DETERMINING LAND SURFACE CLIMATE
the soil accumulates in small depressions or is lost as surface runoff, which flows
overland to streams, rivers, and lakes.
The water balance of the soil is the difference between water input via infiltration
and water loss from evapotranspiration and subsurface drainage. Since the latent heat
flux decreases as soil becomes drier, the amount of water in the soil is a crucial
determinant of the surface climate. Two hydraulic parameters determine the amount
of water a soil can hold. Field capacity is the amount of water after gravitational
drainage. Sandy soils, because of their large pores, hold less water at field capacity
than clay soils, with their small pores (Table 2). Wilting point is the amount of water
in the soil when evapotranspiration ceases. Because water is tightly bound to the
soil matrix, clay soils have very high wilting points. The difference between field
capacity and wilting point is the available water holding capacity. Loamy soils hold
the most water; sands and clays hold the least amount of water.
4 VEGETATION
The ecological characteristics of the surface vary greatly among vegetation types.
Some plants absorb more solar radiation than others; some plants are taller than
others; some have more leaves; some have deeper roots. For example, the albedo of
coniferous forests generally ranges from 0.10 to 0.15; deciduous forests have
albedos of 0.15 to 0.20; grasslands have albedos of 0.20 to 0.25. As albedo
increases, the amount of solar radiation absorbed at the surface decreases and if
all other factors were equal the surface temperature would decrease. However, plants
also vary in height. Trees are taller than shrubs, which are taller than grasses. Taller
surfaces are ‘‘rougher’’ than shorter surfaces and exert more stress on atmospheric
motions. This creates more turbulence, increasing the transfer of sensible and latent
heat away from the surface. Plants also increase the surface area from which sensible
and latent heat can be transferred to the atmosphere. The leaf area of plant commu-
nities is several times that of the underlying ground. A typical ratio of leaf area to
ground area (called the leaf area index) is 5 to 8. Plant communities differ greatly in
leaf area index. A dry, unproductive site supports much less foliage than a moist,
TABLE 2 Hydraulic Conductivity at Saturation and Water Contents
Volumetric Water Content (mm
3
=mm
3
)
Hydraulic
Wilting Point Field Capacity Saturation Conductivity (mm=s)
Sand 0.07 0.23 0.40 0.176
Sandy loam 0.11 0.32 0.44 0.035
Loam 0.15 0.39 0.45 0.007
Silty clay loam 0.22 0.42 0.48 0.002
Clay 0.29 0.45 0.48 0.001
4 VEGETATION 139
nutrient-rich grassland or forest. The rooting depth of plants is important because
this determines how deep in the soil water can be obtained for transpiration. Shallow
rooted plants have a much smaller volume of water for evapotranspiration than deep
rooted plants.
Leaves have microscopic pores, called stomata, through which they a bsorb
carbon dioxide from the atmosphere during photosynthesis. These pores open and
close in response to environmental factors such as light, temperature, atmospheric
CO
2
concentration, and soil water. When they are open, the plant absorbs CO
2
; but
water also diffuses out of the leaf to the surrounding air—a process known as
transpiration. Plants differ greatly in their stomatal physiology, especially responses
to environmental factors. Some plants photosynthesize, and hence have open
stomata, at lower light levels than others. Plants differ in the water content at
which stomata close. They differ in optimum temperatures for photosynthesis, and
they differ in their responses to increasing CO
2
concentration. These different stoma-
tal physiologies contribute to variations in latent heat flux, and hence surface
temperature, among vegetation types.
5 COUPLING TO ATMOSPHERIC MODELS
Different surfaces, with different soil and vegetation types, can create vastly different
climates. These differences can be examined by coupling models of surface energy
fluxes, which depend on the hydrological and ecological state of the land, with
atmospheric numerical models. The surface model provides to the atmospheric
model albedo and emitted longwave radiation, which determine the net radiative
heating of the atmosphere; sensible and latent heat flu xes, which determine
atmospheric temperature and humidity; and surface stresses, which determine atmo-
spheric winds. In turn, the atmospheric model provides the atmospheric conditions
required to calculate these fluxes: temperature, humidity, winds, precipitation, and
incident solar and longwave radia tion.
Land surface models account for the ecological effects of different vegetation
types and the hydrological and thermal effects of different soil types. Although the
equations needed to model these processes at a single point are well understood, the
scaling over large, heterogeneous areas comprised of many soils and many plant
communities is much less exact. Moreover, when coupled to a global climate model
that may be used to simulate the climate of thousand of points on the surface for tens
to hundreds of years, there is a need to balance model complexity with computa-
tional efficiency.
140 PROCESSES DETERMINING LAND SURFACE CLIMATE
CHAPTER 12
OBSERVATIONS OF CLIMATE AND
GLOBAL CHANGE FROM REAL-TIME
MEASUREMENTS
DAVID R. EASTERLING AND THOMAS R. KARL
1 INTRODUCTION
Is the planet getting warmer?
Is the hydrologic cycle changing?
Is the atmospheric=oceanic circulation changing?
Is the weather and climate becoming more extreme or variable?
Is the radiative forcing of the climate changing?
These are the fundamental questions that must be answered to determine if
climate change is occurring. However, providing answers is difficult due to an
inadequate or nonexistent worldwide climate observing system. Each of these appar-
ently simple questions are quite complex because of the multivariate aspects of each
question and because the spatial and temporally sampling required to address
adequately each question must be considered on a global scale. A brief review of
our ability to answer these questions reveals many successes, but points to some
glaring inadequacies that must be addressed in any attempt to understand, predict, or
assess issues related to climate and global change.
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.
141
2 IS THE PLANET GETTING WARMER?
There is no doubt that measurements show that near-surface air temperatures are
increasing. Best estimates suggest that the warming is around 0.6
C(þ0.2
C)
since the late nineteenth century (IPCC, 2001). Furthermore, it appears that the
decade of the 1990s was the warmest decade since the 1860s, and possibly for
the last 1000 years. Although there remain questions regarding the adequacy of
this estimate, confidence in the robustness of this warming trend is increasing
(IPCC, 2001). Some of the problems that must be accounted for include chan ges
in the method of measuring land and marine surface air temperatures from ships,
buoys, land surface stations as well as changes in instrumentation, instrument expo-
sures and sampling times, and urbanization effects. However, recent work evaluating
the effectiveness of corrections of sea surface temperatures for time-dep endent
biases, and further evaluation of urban warming effects on the global temperature
record have increased confidence in these results. Furthermore, by consideration of
other temperature-sensitive variables, e.g., snow cover, glaciers, sea level and even
some proxy non-real-time measurements such as ground temperatures from bore-
holes, increases our confidence in the conclusion that the planet has indeed warmed.
However, one problem that must be addressed is that the measurements we rely upon
to calculate global changes of temperature have never been collected for that purpose,
but rather to aid in navigation, agriculture, commerce, and in recent decades for
weather forecasting. For this reason there remain uncertainties about important
details of the past temperature increase and our capabilities for future monitoring
of the climate. The IPCC (2001) has summarized latest known changes in the
temperature record, which are summarized in Figure 1.
Global-scale measurements of layer averaged atmospheric temperatures and sea
surface temperatures from instruments aboard satellites have greatly aided our ability
to monitor global temperature change (Spencer and Christy, 1992a,b; Reynolds,
1988), but the situation is far from satisfactory (Hurrell and Trenberth, 1996).
Changes in satellite temporal sampling (e.g., orbital drif t), changes in atmospheric
composition (e.g., volcanic emissions), and technical difficulties related to overcom-
ing surface emissivity variability are some of the problems that must be accounted
for, and reduce the confidence that can be placed on these measurements (NRC,
2000). Nonetheless, the space-based measurements have shown, with high confi-
dence, that stratospheric temperatures have decreased over the past two decades.
Although perhaps not as much as suggested by the measurements from weather
balloons, since it is now known that the data from these balloons high in the atmo-
sphere have an inadver tent temporal bias due to improvements in shielding from
direct and reflected solar radiation (Luers and Eskridge, 1995).
3 IS THE HYDROLOGIC CYCLE CHANGING?
The source term for the hydrologic water balance, precipitation, has been measured
for over two centuries in some locations, but even today it is acknowledged that in
142 OBSERVATIONS OF CLIMATE AND GLOBAL CHANGE FROM REAL-TIME MEASUREMENTS
Figure 1 Schematic of observed variations of selected indictors regarding (a) temperature and (b)
the hydrologic cycle (based on IPCC, 2001). See ftp site for color image.
143
many parts of the world we still cannot reliably measure true precipitation (Sevruk,
1982). For example, annual biases of more than 50% due to rain gauge undercatch
are not uncommon in cold climates (Karl et al., 1995), and, even for more moderate
climates, precipitation is believed to be underes timated by 10 to 15% (IPCC, 1992).
Progressive improvements in instrumentation, such as the introduction of wind
shields on rain gauges, have also introduced time-varying biases (Karl et al.,
1995). Satellite-derived measurements of precipitation have provided the only
large-scale ocean coverage of precipitation. Although they are comprehensive esti-
mates of large-scale spatial precipitation variability over the oceans where few
measurements exist, problems inherent in developing precipitation estimates
hinder our ability to have much confidence in global-scale decadal changes. For
example, even the landmark work of Spencer (1993) in estimating worldwide ocean
precipitation using the microwave sounding unit aboard the National Oceanic and
Atmospheric Administration (NOAA) polar orbiting satellites h as several limita-
tions. The obser vations are limited to ocean coverage and hindered by the require-
ment of an unfrozen ocean. They do not adequately measure soli d precipitation, have
low spatial resolution, and are affected by the diurnal sampling inadequacies asso-
ciated with polar orbiters, e.g., limited overflight capability. Blended satellite=in situ
estimates also show promise (Huffman et al., 1997); however, there are still limita-
tions, including a lack of long-term measurements necessary for climate change
studies.
Information about past changes in land surface precipitation, similar to tempera-
ture, has been compared with other hydrologic data, such as changes in stream flow,
to ascertain the robustness of the documented changes of precipitation. Figure 1
summarizes some of the more important changes of precipitation, such as the
increase in the mid to high latitude precipitation and the decrease in subtropical
precipitation. Evidence also suggests that much of the increase of precipitation in
mid to high latitudes arises from increased autumn and early winter precipitation in
much of North America and Europe. Figure 2 depicts the spatial aspects of this
change, reflecting rather large-scale coherent patterns of change during the twentieth
century.
Other changes related to the hydrologic cycle are summarized in Figure 1. The
confidence is low for many of the changes, and it is particularly disconcerting
relative to the role of clouds and water vapor in climate feedback effects.* Observa-
tions of cloud amount long have been made by surface-based human observations
and more recent ly by satellite. In the United States, however, human observers have
been replaced by automated measurements, and neither surface-based or spaced-
based data sets have proven to be entirely satisfactory for detecting changes in
clouds. Polar orbiting satellites have an enormous difficulty to overcome related to
sampling aliasing and satellite drift (Rossow and Cairns, 1995). For human obser-
vers changes in observer schedules, observing biases, and incomplete sampling have
created major problems in data interpretations, now compounded by a change to new
automated measurements at many stations. Nonetheless, there is still some confi-
*An enhancement or diminution of global temperature increases or decreases due to other causes.
144 OBSERVATIONS OF CLIMATE AND GLOBAL CHANGE FROM REAL-TIME MEASUREMENTS
Figure 2 Precipitation trends over land 1900–1999. Trend is expressed in percent per century (relative
to the mean precipitation from 1961–1990) and magnitude of trend is represented by area of circle with
green reflecting increases and brown decreases of precipitation. See ftp site for color image.
145