Download Chapter 8 Surface Climates CHAPTER 8 SURFACE CLIMATES

Chapter 8
CHAPTER 8
Surface Climates
SURFACE CLIMATES ................................................................................................... 1
8.1
INTRODUCTION ................................................................................................................................ 2
8.2
HILLSLOPES AND MOUNTAINS ......................................................................................................... 2
8.3
LAKES AND OCEANS ........................................................................................................................ 9
8.4
LAND COVER................................................................................................................................. 12
8.4.1
Forests and clearings ........................................................................................................... 15
8.4.2
Land surface processes in climate models............................................................................ 17
8.4.3
Landscape heterogeneity and mesoscale circulations.......................................................... 19
8.5
HYDROLOGY ................................................................................................................................. 21
8.5.1
Snow ..................................................................................................................................... 21
8.5.2
Soil water.............................................................................................................................. 22
8.6
MICROCLIMATES AND LAND USE PLANNING .................................................................................. 24
8.7
TABLES ......................................................................................................................................... 26
8.8
FIGURE LEGENDS .......................................................................................................................... 31
Ecological Climatology
8.1 Introduction
Macroclimate is the large-scale climate over 2000 km or more resulting from geographic variation
in net radiation, the resultant transport of heat by the atmosphere and oceans, and high and low surface
pressure belts (Chapter 2). In contrast, mesoclimates and microclimates are regional and local climates,
respectively. Microclimates are climatic features typically smaller than 2 km. A forest has a different
microclimate than an adjacent clearing. Microclimates can be as small as a few centimeters; soil covered
with litter has different temperature and moisture than an adjacent patch of bare soil. Mesoscale is between
microscale and macroscale, covering atmospheric processes at scales of 2 km to 2000 km. At the local and
regional scale, the landscape is a mosaic of patches created by spatial variation in topography, soils, and
land cover (Figure 8.1). By altering the cycling of energy and water between land and atmosphere,
landscape heterogeneity can create unique microclimates and mesoclimates that differ from the prevailing
macroclimate. Mountains, with variation in elevation and slope of terrain, have a different mesoclimate
from plains. Within mountains, valleys have a different microclimate from ridgetops. Forests have a
different microclimate from open rangeland. The previous chapter examined the cycling of energy between
land and atmosphere, especially heat storage and the partitioning of net radiation into sensible and latent
heat. This chapter examines how local landscape features – hills, valleys, lakes, vegetation, snow, soil
water – alter these energy fluxes and create climates in the surface layer (Geiger 1965; Rosenberg et al.
1983; Oke 1987).
8.2 Hillslopes and mountains
Topographic variation in slope of terrain can produce local differences in solar radiation
equivalent to tens of degrees of latitude. This is most evident in mountains, but also occurs along
hillslopes. In addition, changes in elevation alter temperature, precipitation, and winds. As a result,
mountain climates are quite different from low elevations (Barry 1992; Whiteman 2000).
The direct beam radiation on a sloped surface depends on the Sun’s zenith angle, the angle of
slope, the direction of the Sun, and the direction of the sloping surface (Figure 8.2). Maximum intensity is
received when the surface is perpendicular to the Sun’s rays. For a horizontal surface, the solar beam is
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perpendicular to the surface only when the Sun is directly overhead with an altitude angle of 90°. At other
times, the angular deviation from perpendicular is given by the zenith angle (Z), and the amount of
radiation on the surface is S ↓ cos Z , where S↓ is the incoming solar radiation. On a sloped surface, the
angular deviation from perpendicular must be adjusted for the tilt of the surface. The incidence angle is
defined as the angle between the Sun’s beam and an imaginary line perpendicular to the slope. This angle
is given by
cos i = cos s cos Z + sin s sin Z cos ( A
)
−A
sun
slope
where s is the angle of slope, Asun is the azimuth of the Sun, and Aslope is the azimuth of the slope. The
radiation onto the surface is S ↓ cos i . On a horizontal surface, s = 0° and the incidence angle is the zenith
angle. On a sloped surface, the incidence angle can be more or less than the zenith angle depending on the
angle and orientation of the slope.
The azimuth angle of the slope is the direction to which the slope is oriented (north = 0°, east =
90°, south = 180°, west = 270°). The azimuth of the Sun is the compass bearing of the Sun on the horizon.
This direction is in the east in morning, due south in the Northern Hemisphere at solar noon, and in the
west after noon. The azimuth angle, measured as the angular distance from north, varies with time of year,
time of day, and latitude as
⎡ sin δ cos φ −cos δ sin φ cos h ⎤
cos A
=
⎥⎦
sun ⎢⎣
sin Z
where δ is solar declination (Chapter 2), φ is latitude, h is hour angle (Chapter 2), and Z is zenith angle.
This formula gives the angular deviation, up to 180°, from north. It is more convenient to consider azimuth
angles in terms of compass bearing. In the morning, when the Sun is east of north, the computed angle is
also the compass bearing ranging from 0° (north) to 180° (south). In the afternoon, when the Sun is west of
north, the compass bearing is equal to 360° minus the computed angle.
The diffuse radiation on a sloped surface also depends on the angle of slope. The sky forms an
inverted bowl, or half sphere, above and around a point in the landscape (Figure 8.3). On a horizontal
surface, diffuse radiation emanates from all portions of the sky. As the surface is tilted at an angle, less of
the sky hemisphere is viewed from a point on the surface. A portion of the sky is blocked by the terrain,
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Ecological Climatology
from which no sky diffuse radiation is received. With a vertical wall, the sky hemisphere is cut in half and
each side of the wall receives diffuse radiation from only one-half of the sky. A sloped surface, therefore,
sees less of the sky as the angle of slope increases. The fraction of the sky seen, known as the sky view
factor, is given by
ψ sky = (1 + cos s ) / 2
The portion of the sky that is not blocked, ( 1 − ψ sky ), is the fraction of the hemisphere composed of terrain.
This terrain is also a source of diffuse radiation as some of the solar radiation incident on the slope is
reflected. As the angle of slope increase, less of the sky contributes diffuse radiation and more of the
terrain is viewed.
Figure 8.4 illustrates direct and diffuse radiation on various slopes, ignoring terrain radiation. For
all diagrams, the Sun is due south with a zenith angle of 27° (63° elevation above the horizon). On the 27°
(51%) south-facing slope, a line perpendicular to the slope is oriented 63° above the horizon. This is the
same angle at which the solar beam strikes the surface so that the incidence angle is 0° and the surface
receives 1.00 S↓ units of direct beam radiation. Diffuse radiation is received from 95% of the sky. On the
27° north-facing slope, diffuse radiation is still received from 95% of the sky. Now, however, the solar
beam strikes the surface at an angle of 54° from local perpendicular, and the direct beam radiation is only
0.59 S↓. A greater portion of the sky is blocked on the 45° (100%) slopes; the surfaces receive diffuse
radiation from only 85% of the sky. Both 45° slopes receive less direct beam radiation than the comparable
27° slopes. The incidence angle on the south-facing slope is 18° while that of the north-facing slope is 72°.
In these examples, diffuse radiation is little affected by angle of slope. The greatest reduction in radiation
from horizontal is only 15%. Angle of slope has relatively minor effect on direct beam radiation for the
south-facing slopes. The greatest reduction in radiation comes from the direction of slope. The 27° northfacing slope receives only 59% of the direct beam radiation on the south-facing slope. The 45° north-facing
slope receives one-third of the direct beam radiation of the south-facing slope.
The angle of slope, its direction of tilt, and latitude interact with time of year and time of day to
produce complex patterns of solar radiation on a surface. For example, midday direct beam radiation varies
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greatly throughout the year on a 42% (23°) north slope at latitude 45 °N, but is relatively constant
throughout the year on a 20% (11°) southeast slope located at latitude 30 °N (Figure 8.5). The length of
day the slope is illuminated by the solar beam varies by less than three hours over the year on the southeast
slope. In contrast, the north slope receives no direct beam radiation from December through mid-January,
and daylength increases to 15 hours in June. Figure 8.6 illustrates diurnal patterns of solar radiation on
vertical walls at latitude 40 °N. On the winter solstice, a south wall receives more radiation than a
horizontal surface. This is because the Sun is low on the horizon and its rays strike the wall at an angle
closer to perpendicular than on the horizontal surface. The east and west walls receive direct beam
radiation before noon and after noon, respectively. During other times of the day, they receive only diffuse
radiation. The north wall is never illuminated by the Sun, but receives a small amount of diffuse radiation
from the sky. Similar patterns occur in summer with three notable differences. First, the south wall receives
direct beam radiation for a shorter period than the Sun is above the horizon because the Sun rises north of
east in the morning and sets north of west in the evening. The Sun is not south of the east-west plane
created by the wall until 0800 and remains so until 1600 hours. Second, the north wall is illuminated by
direct beam radiation in early morning (0500 to 0800 hours) and late afternoon (1600 to 1900 hours) when
the Sun is north of the east-west plane. Third, all walls receive less radiation at midday than the horizontal
surface. In summer, high solar altitude angles cause the solar beam to strike vertical surfaces at a more
oblique angle.
There is a complex relationship among slope, aspect, latitude, time of year, time of day, and solar
radiation that precludes simple statements that north slopes receive less radiation than south slopes and
gentle slopes receive more radiation than steep slopes (Figure 8.7, Figure 8.8). In general, however, the
effect of slope and aspect on direct beam radiation is greater at high latitudes, where the Sun is low on the
horizon, than in the tropics, where the Sun is high above the horizon throughout the year. In middle to high
latitudes, the direction of slope is critical to the thermal balance. For example, north-facing slopes in
interior Alaska receive less solar radiation than south-facing slopes and therefore have cold, perennially
frozen soil (permafrost) (Viereck et al. 1983). Slope and aspect have greater effect on direct beam radiation
than on diffuse radiation. Hence, slope and aspect are more important in summer, when the Sun is high in
the sky, than in winter, when the low sun angle increases the ratio of diffuse radiation to direct beam
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Ecological Climatology
radiation. Slope and aspect are more important on clear days, when much of the radiation is direct beam,
than on cloudy days, when the radiation is diffuse.
The effect of radiation loading on air temperature is illustrated in Figure 8.9, which shows air
temperature and elevation for a forested site in summer. At this site, elevation ranges from 340 m to 540 m
with slopes inclined at 27% (15°) to 84% (40°). On the particular summer day studied, air temperatures
varied by as much as 3.5 °C depending on aspect. The southwest-facing slope was warmest, with
temperatures up to 22.5 °C. The northeast-facing slope, with temperatures as low as 19.0 °C, was coldest.
On the north side of the hill, temperatures increased from 19.0 °C in the east to 21.0 °C in the west. This is
because the Sun was high above the horizon and illuminated the north-facing slopes late in the afternoon.
Tall mountains also affect incident solar radiation by reducing the path length that solar radiation
travels through the atmosphere. With greater height in the atmosphere, fewer air molecules are found and
pressure decreases (Figure 2.3). At a height of 1.5 kilometers, the pressure is 846 hPa – 83% of that at sea
level. At a height of 3 km, air pressure is 69% that at sea level, and at a height of 5.5 km air pressure is
about 50% that at sea level. Because there is less air mass at higher elevation, solar radiation is less likely
to be scattered or absorbed as it passes through the atmosphere and more radiation reaches the ground.
Mathematically, the attenuation of solar radiation is
S = ( S cos i )τ
o
m
where So is solar radiation on a horizontal surface at the top of the atmosphere, i is incidence angle, and τm
is atmospheric transmittance. A typical transmittance ranges from 0.6 to 0.7 for clear skies. The optical air
mass m is defined as
m=
1 P
cos Z Ps
where Z is zenith angle, P is pressure, and Ps is pressure at sea level. Optical air mass increases so that
transmittance decreases with greater zenith angle. The longer path length through the atmosphere causes
less radiation to reach the ground. For a given zenith angle, optical air mass decreases with elevation (as
pressure decreases). Consequently, atmospheric transmittance increases with higher elevation. For
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Chapter 8 – Surface Climates
example, with a zenith angle of 20°, solar radiation is 0.68 So at sea level (assuming τ = 0.7). This increases
to 0.73 So at 1.5 km, 0.77 So at 3 km, and 0.83 So at 5.5 km.
Another noticeable feature of mountains is the cool air. In part, this is because there is less heating
of the atmosphere from the surface below. In addition, the amount of water vapor in the atmosphere
decreases with higher altitude and consequently the absorption and re-radiation of surface longwave
radiation by the atmosphere also decreases. So long as the air is not saturated with water, it cools at a rate
of about 1 °C per 100 m. It regains heat at the same rate when it descends. This temperature change is
called the dry adiabatic lapse rate. If a low-lying site has an air temperature of 30 °C, a site located 300 m
higher will be exposed to 27 °C air. At the top of a 1500 m mountain, the temperature is only 15 °C (Figure
8.10).
The amount of water vapor that air can hold without becoming saturated depends on its
temperature (Figure 7.4). As moist air rises up a mountain and cools, the amount of water vapor it can hold
decreases. So long as the air is not saturated, it cools at the dry adiabatic lapse rate. When the air becomes
saturated (i.e., relative humidity is 100%), some of the water vapor condenses into tiny droplets, forming
fog or clouds. This condensation releases heat (the stored latent heat of vaporization), and the cooling of air
as it rises is reduced to about 0.5 °C per 100 m. This is called the moist adiabatic lapse rate. The bottom
panel in Figure 8.10 shows changes in air temperature as moist air moves over a mountain. The rising air
on the windward slope cools at the dry adiabatic lapse rate until the air becomes saturated, in this example
at about 900 m. When saturated, clouds form and the cooling is reduced to the moist adiabatic lapse rate.
At the summit, the air is 18 °C, which is 3 °C warmer than if it cooled at the dry adiabatic lapse rate. If, as
in this example, water is lost by precipitation, the air becomes unsaturated and descends on the leeward
slope at the dry adiabatic lapse rate. Because air is warmed by latent heat of condensation as it moves
upslope and by adiabatic heating as it moves downslope, it reaches the bottom warmer than it started on
the other side – in this case with a temperature of 33 °C.
Under the right conditions, temperatures increase with elevation rather than decrease. Mountains
and hillslopes develop local wind circulations in response to spatial variation in surface heating. Under
calm conditions and clear sky, light winds often blow upslope during the day and downslope at night
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(Figure 8.11). During the day, mountain slopes absorb solar radiation. Air is heated by these warm
surfaces, becomes less dense, and rises. Air flows upslope from low-lying valleys, ravines, or plains to
replace the ascending mountain air. These upslope circulations depend on a temperature contrast and
develop most strongly on slopes receiving the greatest amount of solar radiation. For example, an eastfacing slope heats up from early morning solar radiation and may develop upslope winds before a westfacing slope. At night, the slopes cool, and cold air near the surface flows downhill and collects in lowlying areas, often forming frost pockets. Cold air drainage is seen in Figure 8.12, which shows earlymorning air temperature in relation to elevation in a broad valley in central Pennsylvania (Hocevar and
Martsolf 1971). Temperature over a 20-km distance varied by as much as 9 °C with low elevations colder
than high elevations. For this particular night, air temperature increased 3.4 °C per 100 m. On average for
all nights studied, air temperature increased at a rate of 6.2 °C per 100 m.
Warmest nighttime temperatures are often found at mid-slope. Ridgetops are colder due to their
high elevation; valleys are cold because of cold air drainage. Figure 8.13 shows the characteristics of this
mid-slope thermal belt for Austrian mountains. On the particular mountain studied, daytime temperatures
decreased with elevation as expected. Nighttime temperatures increased with elevation up to 800 m;
thereafter temperatures decreased with elevation. The warm region at 800 m is the mid-slope thermal belt.
Many meteorological factors influence the location of this belt. The thermal belt is best developed during
clear, calm nights. With high winds or rain, the normal lapse rate occurs. As a result, the elevation with
warmest temperature is not constant but varies over time. The middle panel in Figure 8.13 shows the
statistical distribution of the thermal belt. The warmest temperature was most often found at an elevation of
800 m. A second weaker maximum occurred at the bottom of the valley, reflecting conditions in which
temperature decreased regularly with height.
As shown in Figure 8.10, the cooling of air with higher elevation often causes condensation and
precipitation on the windward side of mountains and dry conditions on the leeward side. This mountaininduced precipitation is called orographic precipitation. The influence of mountains on precipitation can be
seen in annual precipitation along a west-to-east transect in the United States from the Pacific Ocean to the
Atlantic Ocean (Figure 8.14). Annual precipitation drops from 1000 mm west of the Pacific Coastal
Ranges to about 500 mm on the east side of the mountains. Precipitation on the west side of the Sierra
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Nevada is 1000 mm, but less than 200 mm on the east side. Similar orographic precipitation is seen in the
Rocky Mountains.
Another feature of mountain climates is that mountaintops can often be exceedingly windy. These
winds are evident in the twisted, gnarled shapes of exposed trees growing on or near mountaintops. In the
United States, the strongest wind recorded (104 m s-1) occurred on the summit of Mount Washington in
New Hampshire (Williams 1994, p. 48). As air flows over the land surface, frictional resistance from
vegetation, buildings, and the ground slows the wind. This resistance decreases rapidly with height in the
atmosphere so that high altitude winds are stronger than surface winds. In addition, air flowing over a
mountain is constricted because it has less atmosphere to flow through. This accelerates the wind, much as
water flows slowly in a wide river but rushes rapidly through a narrow gorge.
8.3 Lakes and oceans
Lakes and oceans are important determinants of local and regional climate. In general, milder
temperatures with smaller diurnal and seasonal temperature variability characterize seashore and lakeside
climates compared with inland climates. Even small lakes in a city can create a locally cooler climate (Ishii
et al. 1990/91). In winter, temperatures are warmer along the shoreline than inland. Rivers can also locally
warm climate in winter. Morning air temperatures within 1 km of the Connecticut River near Piermont,
New Hampshire, have been found to be 4 °C warmer than further inland while the river was freezing,
perhaps due to latent heat release as ice cover grows (Hogan and Ferrick 1998). Under the right conditions,
the climatic effects of lakes and oceans extend far from shoreline. Simulations with a mesoscale climate
model run once with the Great Lakes and once without the lakes show a 2 °C warming due to the lakes as
far away as Philadelphia, Pennsylvania, during a 2-day cold period in November 1982 (Sousounis and
Fritsch 1994; Sousounis 1997, 1998). Large paleolakes in western United States (Hostetler et al. 1994) and
paleolakes and wetlands in northern Africa (Coe and Bonan 1997; Broström et al. 1998; Carrington et al.
2001) during wet climates increased precipitation compared with present day. Large cold lakes along the
margin of the Laurentide Ice Sheet 11 000 years ago reduced precipitation on the glacier (Hostetler et al.
2000).
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Ecological Climatology
These differences arise from differences in surface radiation and heat capacity between water and
land. Lakes and oceans, when ice-free, generally have a lower albedo than land (0.03 to 0.10 versus greater
than 0.10). As a result, they absorb more solar radiation than land. However, solar radiation penetrates
through the water to considerable depths so that it heats the upper 10 m or so of water rather than merely
the surface. In addition, lakes and oceans have a large capacity to store heat. A typical heat capacity for soil
is 2 × 106 J m-3 °C-1 while that of water is 4.2 × 106 J m-3 °C-1 (Table 6.5). In other words, a cubic meter of
water requires twice as much energy to warm 1 °C as does the same volume of soil. Conversely, it must
lose twice as much energy to cool by 1 °C. The large heat capacity of water reduces daytime warming and
nighttime cooling compared with land. Likewise, the annual cycle of air temperature is moderated as heat
is stored in water during summer and released in winter.
The heat stored in lakes and oceans is mixed vertically as a result of density differences in water.
For freshwater, greatest density occurs at about 4 °C. Water warmer than 4 °C is less dense. Hence, cold,
dense water sinks to the bottom of lakes while warm, light water rises to the surface. This is most evident
in summer months, when temperate lakes are stratified into three distinct layers: an upper layer formed by
warm, light water (the epilimnion); a deep layer formed by cold, dense water (the hypolimnion); and a
transition zone in which temperature changes rapidly with depth (the thermocline) (Horne and Goldman
1994).
Figure 8.15 illustrates this thermal stratification for a typical temperate lake. Mirror Lake is a
small lake in the White Mountains of New Hampshire, with a surface area of 15 ha (1 ha = 10 000 m2) and
an average depth of about 6 m. The lake is thermally stratified in summer. The surface water is warm, with
a temperature of 23 °C. The water is well mixed by winds to a depth of about 4-5 m. Below this depth,
temperatures decrease rapidly so that water deeper than 7-8 m is 10 °C or colder. Towards the end of
summer and into autumn, surface water cools and sinks to deeper depths. As this warm water mixes with
deeper cold water, heat is distributed throughout the vertical profile in a process known as overturning, and
the lake develops a uniform temperature profile from surface to bottom. Autumn overturning begins in
August and progressively penetrates downward. Mixing occurs rapidly and by October the lake has a
uniform temperature profile. This isothermal period lasts until ice forms on the lake surface. Ice typically
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Chapter 8 – Surface Climates
forms in late November or early December and lasts for 138 days on average. The ice is generally 40-75
cm thick by late February. During this time, the lake is inversely stratified. Temperatures remain relatively
constant between 3 °C and 4 °C, with surface water directly under ice warmer than deeper water. Ice
usually begins to break up in mid- to late April. This is followed by spring convective mixing, when the
lake again overturns and develops a uniform temperature profile. Spring mixing is rapid and occurs over a
period of a few days to a week. Surface water then warms, leading to summer thermal stratification.
In addition to modulating temperature, lakes and oceans can develop local atmospheric
circulations. Land and sea breezes develop when the temperature contrast between land and water produces
slight air pressure differences that result in surface breezes onshore during the day and offshore at night
(Figure 8.16). During the day, land warms more than water. The air overlying land warms, and the warm
surface air rises. The expansion of the air column over land means that pressure aloft is higher over land
than over water. The horizontal pressure gradient causes winds aloft to blow offshore from land to water.
The offshore flow aloft adds mass to the column of air over water, increasing the pressure at the ocean
surface. Surface air flows from high pressure over ocean to the low pressure over land. The end result is a
local circulation system in which warm air rising over land is replaced with cool ocean air at the surface.
These sea breezes generally develop by mid-morning on hot, sunny days with calm weather. At night, land
cools more than water. The result is a system of descending air over land and rising air over water driven
by offshore surface breezes. Similar lake breezes can develop for lakes of 200 km2 or so (Sun et al. 1997).
While sea breezes are local in extent, monsoons are seasonal large-scale atmospheric circulations
generated by land-sea temperature contrasts. The Asian summer monsoon is one example of this. In
summer, the continent warms more than oceans. Low surface pressure over the continent draws cool, moist
air inland from adjacent oceans (Figure 2.8), triggering heavy rainfall. The strength of the monsoon is
determined by differences in the surface temperature of the Asian continent and surrounding water. Snow
cover is an important determinant of this temperature contrast (Barnett et al. 1988, 1989; Meehl 1994).
Low snow cover in the Tibetan Plateau results in a low land albedo, causing warmer land temperatures,
greater land-sea temperature contrast, and a stronger summer monsoon. Soil moisture is also important
(Meehl 1994). Wet soils cool the surface due to greater evapotranspiration, reducing the land-sea
temperature contrast, but also provide moisture for further precipitation. In a similar process, heavy spring
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Ecological Climatology
snow cover and wet soil in the southern Rocky Mountains diminish the summer monsoon in Southwest
United States (Gutzler and Preston 1997; Small 2001).
Another more local influence of water bodies on precipitation is the lake effect storms that dump
heavy snowfall on the southern and eastern shores of the Great Lakes. Snowfall at lakeside may be twice as
deep as inland sites. Lake effect snow arises when cold, dry Arctic air moves over the warmer Great Lakes.
The cold air is heated and moistened as it travels over the lakes. Heat and moisture makes the air unstable.
The warm, moist surface air rises, forming clouds that precipitate snow. In addition, the warmer, more
humid air encounters a rough surface as it flows on shore. Winds slow and the convergence of air forces air
upwards where it cools, condenses, and precipitates as snow.
8.4 Land Cover
Land cover, especially the type of vegetation, also affects local and regional climates due to
variation in albedo, soil water, surface roughness, plant physiology, the amount of leaf area from which
heat can be exchanged, and rooting depth (Table 7.2). One important factor is albedo. High albedos result
in less solar radiation absorbed at the surface and a cooler surface, all other factors being equal; low
albedos result in more absorbed solar radiation and a warmer surface. Albedos range from about 0.80 to
0.95 for fresh snow to as little as 0.03 to 0.10 for water at low solar zenith angle (Table 8.1). Snow, deserts,
and glaciers have the highest albedos. Urban surfaces have low albedos. Soil albedo generally decreases
with coarser particle size. Coarse soil particles trap radiation through multiple reflections among adjacent
particles. In contrast, fine soil particles expose a relatively uniform surface, trapping less radiation. Soil
albedo also decreases with soil wetness because radiation is trapped by internal reflection.
Vegetation has low albedos, typically ranging from 0.05 to 0.25, with forests absorbing more solar
radiation than grasslands or croplands. This overall surface albedo is the combined reflection of all plant
material (leaves and stems) and soil. The solar radiation striking an individual leaf is absorbed, reflected, or
transmitted by the leaf. These optical properties have a strong wavelength dependence. For example, the
leaf of a broadleaf tree typically absorbs 85% of the radiation in wavelengths between 0.4 µm and 0.7 µm
(Table 8.2). This radiation, referred to as the visible waveband, is used during photosynthesis to convert
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Chapter 8 – Surface Climates
CO2 into biomass. Absorption decreases to 30% at wavelengths above 0.7 µm (the near-infrared
waveband). Lower absorption ensures that this light, which is not utilized during photosynthesis, does not
overheat the leaf. Different plant types differ in optical properties. Needleleaf trees reflect less solar
radiation than broadleaf trees or herbaceous plants (Table 8.2).
The albedo of a canopy of leaves is the integrated transfer of radiation among all leaves, stems,
and the ground. With low leaf area index, plants absorb little solar radiation and the overall albedo is
largely is that of soil. However, the absorption of radiation increases greatly with higher leaf area (Figure
8.17). With a leaf area index of 4 m2 m-2, 90% of the visible radiation is absorbed. Absorption saturates at
95% of the incoming radiation with a leaf area index of 6 m2 m-2. Even when the ground is covered by
snow, which has a high albedo, plant material effectively masks the underlying snow (Figure 8.17). At low
leaf area index, the overall albedo is that of snow, but albedo decreases markedly with higher leaf area
index. For leaf area index greater than 3 m2 m-2, the overall albedo is similar to that of vegetation without
snow. Reflection of solar radiation by plants depends on solar zenith angle. In general, surface albedo
increases greatly when the Sun is near the horizon (Figure 8.17). Plants differ in leaf orientation, which
affects the dependence of albedo on zenith angle. Needles tend to be arranged randomly, evergreen and
deciduous broadleaves tend to be semi-horizontal, and grasses and crops have semi-vertical foliage (Table
8.2). The albedo of a canopy of semi-horizontal leaves varies little with solar zenith angle while that of
random and semi-vertical leaves increases with zenith angle (Figure 8.17).
Another important factor is surface roughness. Aerodynamic resistance decreases with rougher
surfaces. Rough surfaces generate more turbulence and have higher sensible and latent heat fluxes than
smooth surfaces, all other factors being equal. Surface roughness varies greatly with land cover (Table
8.3). Bare soil is a smooth surface. Suburban and urban surfaces are rough surfaces depending on the
density and height of buildings. The presence of vegetation increases surface roughness in relation to
canopy height and density. Short grass is a smoother surface than tall grass. Forests are aerodynamically
rougher than crops or grasses. It is often assumed that roughness length for vegetation is one-tenth canopy
height and displacement height is seven-tenths canopy height (Grace 1983, p. 44; Monteith and Unsworth
1990, p. 117; Campbell and Norman 1998, p. 71). With these assumptions, aerodynamic resistance
decreases with increasing vegetation height (Figure 7.7).
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Ecological Climatology
The depth to which roots penetrate in soil affects the overall supply of water for transpiration.
Deeper roots provide a greater volume of soil from which water can be extracted. Consider, for example, a
loam, which has a volumetric available water-holding capacity of 0.239 (0.53 × 0.451, Table 6.1). Plants
with roots in only the upper 20 cm of soil have 47.8 mm of water available for transpiration. This water is
likely to be quickly depleted unless replenished by rainfall. In contrast, plants with roots extending to 1 m
have 239 mm of available water. This much larger pool of water protects plants from lack of rainfall. Root
profiles vary greatly among vegetation types (Figure 10.43).
Yet another important difference among vegetation is canopy resistance to heat and moisture
fluxes. This depends on the size of leaves and the physiology of stomata (Table 5.3, Figure 7.15). Plants
differ in the physiology of stomata (Chapter 9). Some plants open stomata at low light levels; others need
high light levels before stomata fully open. Some plants prefer warm temperatures while others thrive in
cold temperatures. Plants also differ in their ability to tolerate drought. Some plants can extract water and
grow on extremely dry soils whereas other plants close stomata and cease growth when the soil is only
moderately dry. Canopy resistance also depends on leaf area index (Figure 7.15). Higher leaf area provides
a greater surface area from which heat and moisture are exchanged with the atmosphere.
The physical properties of soils determine how much water is available for evapotranspiration. In
general, clay and sand have less available water than loam (Table 6.1). The large pores in sand limit
moisture retention. Conversely, clay retains considerable water at wilting point, but this water is
unavailable for evapotranspiration. Thus, sand and clay are likely to sustain less evapotranspiration than
loam. Moreover, sand drains faster than loam or clay (Figure 6.18) and is therefore likely to store less
water following a rainstorm than other soils. Conversely, infiltration is low and runoff is high in clay soils
(Figure 6.20).
The use of paper, hay, and black plastic mulches to cover soil illustrates the effect of surface
properties on energy fluxes and microclimate (Table 8.4). These mulches alter surface albedo, surface
resistance to evaporation, and thermal conductivity. At midday on the warm summer day when
measurements were made, the uncovered soil had a surface temperature of 40 °C. Almost one-third of the
net radiation was dissipated as latent heat. All mulches reduced evaporation compared with bare soil, but
they differed in their effect on surface temperature. The soil covered with paper was similar in temperature
14
Chapter 8 – Surface Climates
to the bare soil. Although evaporation was reduced compared with the bare soil, much less radiation was
absorbed at the surface due to the high albedo of the light-colored surface. In contrast, soil covered with
black plastic had a temperature of 52 °C. This material was hot because of its low albedo and because the
plastic barrier prevented evaporation. Instead, most of the net radiation was dissipated as sensible heat. The
hay mulch also was hot (51 °C). In this case, the surface absorbed radiation similar to the bare soil.
However, the mulch reduced evaporation and hindered heat transfer to the underlying soil due to low
thermal conductivity.
Surface albedo can be purposely altered to change soil temperature. In one such study, a white
powder was applied to cool soil temperature (Table 8.5). In the untreated soil, 30% of incoming solar
radiation was reflected. Net radiation (6.2 MJ m-2 day-1) was partitioned primarily as latent heat. Whitening
the soil surface increased albedo so that 60% of incoming solar radiation was reflected. This reduced net
radiation at the surface despite a reduction in outgoing longwave radiation brought about by the cooler
surface. Less radiation was available to warm the soil or evaporate water. The soil cooled by 5 °C and
evaporation decreased by 19%. Because the treated surface was colder than the air, sensible heat flux
changed direction and was transferred to the colder surface.
8.4.1 Forests and clearings
The effect of vegetation on microclimates is seen by contrasting forested and open locations
(Kittredge 1948; Geiger 1965). In general, daytime air temperature during summer is cooler in a forest than
above canopy or in clearings (Figure 8.18). For example, one study found daytime air temperature above
the canopy was greater than 22 °C while air near the forest floor was 4 °C cooler. Another study found
daytime air temperature within a forest (18 °C) was 2-4 °C cooler than temperatures in an adjacent
clearing. At night, however, the dense foliage blocks the loss of longwave radiation from the surface and
temperatures within the forest were up to 3 °C warmer than in the clearing.
Many factors such as the orientation of the forest edge, the height of trees, their leaf area, and
prevailing winds influence the microclimates of a clearing. The size of a clearing relative to the height of
surrounding trees is one critical factor. The bottom panel in Figure 8.18 shows that midday air temperature
measured at the center of a clearing is much warmer than that measured in the surrounding forest. This
15
Ecological Climatology
temperature excess increases with clearing size. This is due to less shading by surrounding trees as the
clearing increases in size. The trees surrounding a clearing cast a shadow into the clearing with length
L = H / tan( a )
where H is the height of the trees and a is the solar altitude angle (Figure 8.19). With larger clearing size,
however, radiation heating is less important as winds effectively carry heat away from the surface.
The size of a clearing also affects the nighttime radiation balance. At night, net radiation is the
difference between atmospheric longwave radiation impinging on the surface and longwave radiation
emitted by the surface. However, the surrounding trees obscure a portion of the sky. For a circular clearing
of diameter D surrounded by trees of height H, the portion of the sky seen at the center of the clearing is
ψ
sky
= cos 2 ( β )
where β is the angle formed between the center of the clearing and treetop (Figure 8.20). The remainder
( 1 − ψ sky ) is the terrain view factor. The total longwave radiation incident at the center of the clearing is
the sum of two streams of radiation – that from the sky (L↓) and that from the surrounding trees (Lv) – and
is given by L ↓ ψ sky + Lv (1 − ψ sky ) . As the diameter of a clearing increases in size, a greater fraction of
the longwave radiation comes from the sky and less comes from the surrounding trees. Since atmospheric
radiation is considerably less than terrestrial radiation at night, there is less energy at the surface as the
clearing size increases (Figure 8.21). Net radiation is an important determinant of nighttime surface air
temperature. Hence, one would expect smaller clearings to be warmer than larger clearings at night. In
early growing season, when sudden frosts can kill new growth, this can be a critical factor in successful
stand regeneration.
The presence of trees can also alter local winds. Patches of trees arranged closely together
generally decrease wind to a distance of 20 to 30 times their height. This windbreak, or shelterbelt, effect
has been widely used in rural settings to reduce winds blowing across open farmlands, thereby reducing
erosion of soil and heat loss from isolated buildings (Rosenberg et al. 1983; Oke 1987). Figure 8.22
illustrates the general aspects of windbreaks. The greatest reduction in wind occurs within a distance that is
less than five times tree height, where winds are reduced to 20-40% of those in open areas. The
16
Chapter 8 – Surface Climates
penetrability of barriers greatly affects wind reduction. A dense barrier has the largest reduction in wind
speed, but the abrupt blockage of wind produces strong turbulence downwind so that the shelter effect is
limited to only a short distance. The shelter effect is also small in sparse vegetation, where the winds
readily penetrate the barrier. In contrast, the shelter effect extends further downwind in medium density
windbreaks. Phenology influences the shelter effect. Winds are generally 20-30% lower when deciduous
trees are in leaf than when leaves are off the trees (Landsberg 1981, p. 130).
8.4.2
Land surface processes in climate models
Observations such as those in Table 8.4 and Table 8.5 demonstrate microscale variation in climate
related to surface energy fluxes. At larger scales such as a forest clearing, careful observations can
demonstrate different microclimates related to surface properties. However, much of our understanding of
how land affects weather and climate comes from numerical models of surface fluxes and the hydrologic
cycle coupled to climate models (Figure 8.23). Biogeophysics is the physical interactions of the biosphere
and geosphere with the atmosphere in terms of fluxes of energy, water, and momentum. Biogeophysical
processes represented in these models include: absorption, reflection, and transmission of solar radiation;
absorption and emission of longwave radiation; sensible and latent heat fluxes, partitioning latent heat into
evaporation of intercepted water, soil evaporation, and transpiration; momentum fluxes; heat transfer in
multilayered soil; and stomatal physiology. The hydrologic cycle is represented in terms of interception,
throughfall, stemflow, infiltration, runoff, soil water, snow, evaporation, and transpiration. Watersheds are
represented by separate calculation of runoff from upslope and saturated areas and by routing runoff
through river networks to oceans.
In climate models, land and atmosphere represent a coupled system. Atmospheric radiation,
temperature, humidity, wind, and precipitation depend on surface fluxes of energy, moisture, and
momentum. In turn, these fluxes depend on atmospheric radiation, temperature, humidity, wind, and
precipitation. Hence, there is a coupling between the atmospheric state and surface conditions. Changes in
albedo, roughness, leaf area index, stomatal physiology, rooting depth, soil texture and soil water alter
surface fluxes and the hydrologic cycle and in doing so alter climate (Table 7.2). These properties vary
spatially and are input to the model from spatial maps of ecosystem geography and soil texture. Paired
17
Ecological Climatology
climate model simulations that alter surface properties can be used to determine the effect of a particular
vegetation change on climate. For example, one simulation might simulate climate with the current
vegetation cover (Figure 8.24). A second simulation might simulate climate with an altered vegetation
cover. The difference between the two simulated climates is the effect of the altered vegetation on climate.
Such paired simulations routinely demonstrate the importance of vegetation in determining
regional and global climate (Chapters 12, 13). For example, higher leaf area index generally increases
evapotranspiration over vegetated regions in summer provided there is sufficient soil water (Chase et al.
1996; Dickinson et al. 1998; Eastman 1999; Bounoua et al. 2000; Lu et al. 2001; Tsvetsinskaya et al.
2001a,b; Eastman et al. 2001a,b; Buermann et al. 2001). As a result, surface temperature cools and
precipitation may increase. Lower leaf area has the opposite effect. Other studies show that by providing
reliable sources of water, deep roots result in more sustained evapotranspiration during dry seasons than
shallow roots, creating a cooler, moister climate (Nepstad et al. 1994; Desborough 1997; Kleidon and
Heimann 1998a,b, 2000; Zeng et al. 1998). Climate model simulations in which stomatal conductance
decreases in response to higher atmospheric CO2 routinely show decreased evapotranspiration, increased
sensible heat, and surface warming over large vegetated regions in summer (Henderson-Sellers et al. 1995;
Pollard and Thompson 1995; Sellers et al. 1996a; Bounoua et al. 1999; Douville et al. 2000; Levis et al.
2000).
Other studies have examined the impact of deforestation on climate (Chapters 12, 13). In the
tropics, clearing of forest for pastureland generally results in a warmer and drier climate owing to higher
albedo, decreased surface roughness, and reduced evapotranspiration (Dickinson and Henderson-Sellers
1988; Lean and Warrilow 1989; Nobre et al. 1991; Dickinson and Kennedy 1992; Mylne and Rowntree
1992; Henderson-Sellers et al. 1993; Lean and Rowntree 1993; Pitman et al. 1993; Polcher and Laval
1994a,b; Sud et al. 1996; McGuffie et al. 1995; Lean and Rowntree 1997; Hahmann and Dickinson 1997;
Costa and Foley 2000). Replacement of dryland vegetation with desert also reduces rainfall by increasing
albedo and reducing evapotranspiration (Charney et al. 1975; Chervin 1979; Sud and Fennessy 1982,
1984; Laval and Picon 1986; Sud and Molod 1988; Xue and Shukla 1993, 1996; Dirmeyer and Shukla
1996; Xue 1996, 1997; Clark et al. 2001). Conversely, greening of desert landscapes increases rainfall
(Kutzbach et al. 1996; Claussen and Gayler 1997; Texier et al. 1997; Broström et al. 1998; Ganopolski et
18
Chapter 8 – Surface Climates
al. 1998; Brovkin et al. 1998; Claussen et al. 1999; Joussaume et al. 1999; de Noblet-Ducoudré et al.
2000). Clearing temperate forests for cropland cools climate by increasing albedo (Bonan 1997, 1999;
Hansen et al. 1998b; Brovkin et al. 1999; Pitman and Zhao 2000; Govindasamy et al. 2001; Zhao et al.
2001). Replacement of boreal forest with tundra cools climate by exposing the high albedo of snow (Bonan
et al. 1992; Thomas and Rowntree 1992; Chalita and Le Treut 1994; Foley et al. 1994; TEMPO 1996;
Douville and Royer 1996; Levis et al. 1999a).
8.4.3
Landscape heterogeneity and mesoscale circulations
The effect of vegetation on regional climate is most noticeable along the transitions between one
vegetation type and another, where spatially heterogeneous landscape patterns can influence boundary
layer structure and mesoscale atmospheric circulation (Cotton and Pielke 1995; Giorgi and Avissar 1997).
The influence of surface heterogeneity arises as a result of differential surface energy fluxes and
atmospheric heating. For example, dry surfaces, because they lack sufficient water for sustained
evapotranspiration, have high sensible heat flux and low latent heat flux. The overlying air tends to become
warm and dry. In contrast, wet surfaces have high latent heat fluxes; the air is cool and moist. Local
differences in albedo can also generate mesoscale circulations (Pielke et al. 1993). Indeed, at one point in
time coating large areas in coastal arid climates with asphalt was proposed as a means to induce more
rainfall (Black and Tarmy 1963; Black 1963). The contrast in surface fluxes needed to create mesoscale
circulations is particularly large in semi-arid climates where patches of irrigated croplands are interspersed
among dry native grasses. The large horizontal contrast in sensible heat flux can lead to a strong circulation
similar to a sea breeze in which surface winds blow from the cooler cropland to the warmer dry grassland
while upper winds flow in the opposite direction (Anthes 1984; Mahfouf et al. 1987; Yan and Anthes
1988; Segal et al. 1988; Avissar and Pielke 1989; Segal and Arritt 1992; Chen and Avissar 1994a,b; Seth
and Giorgi 1996).
Figure 8.25 illustrates the development of such a mesoscale circulation. In these simulations, a
mesoscale atmospheric model was run once with homogenous wet forest land surface and once with a 200km patch of dry grassland surrounded by wet forest. The difference between simulations highlights the
atmospheric changes due to the dry grassland. As a result of less latent heat and more sensible heat, the
19
Ecological Climatology
near-surface air over the dry grassland was 1-4 °C warmer and drier compared with the forest-only
simulation. The differential surface heating between the dry grasses and wet trees on either side created
two circulations that converged over the warm grassland. The circulation on the western edge was
counterclockwise, with near-surface westerly wind and easterly wind aloft. The circulation along the
eastern edge was clockwise. This was accompanied by upward movement of air over the warm, dry
grassland and descent on either side. These circulations transported moisture from the forest inward to the
grassland and then upward so that air high in the atmosphere over the grassland was moistened.
By creating gradients in surface heating, a heterogeneous mixture of vegetation can influence
precipitation (Chen and Avissar 1994b; Avissar and Liu 1996; Pielke et al. 1997). For example, the
landscape of northern Georgia is a mixture of forest and farmland (Figure 8.26, color plate). Simulations
with a mesoscale atmospheric model contrasted the effects of a homogeneous forest landscape with a
mixture of forest and farmland (Pielke et al. 1997). In these simulations, covering a spatial domain of 210
km by 210 km, forests were wet so that latent heat was high and sensible heat was low. Other vegetation
types were water-limited. By 1400 hours on the particular summer day simulated, a large cloud with
widespread convective precipitation had developed in the simulation with heterogeneous vegetation. The
homogenous vegetation had little convective activity, and clouds did not form until later in the afternoon.
Similar simulations in a region of west Texas also found convective activity to depend on the type
of vegetation specified (Pielke et al. 1997). This region is part of the High Plains dryline, which separates
warm, moist air extending northwards from the Gulf of Mexico from hot, dry air originating in Mexico and
southwestern United States (Figure 8.26, color plate). It is the favored location for thunderstorms in spring
and early summer. For the particular late spring day simulated, cloud development and convective
precipitation were more extensive and vigorous in a simulation with the observed heterogeneous mixture of
crops, trees, and short grass prairie than in a simulation with only short grass vegetation. The near-surface
atmosphere was also moister than in the short grass simulation. This is a result of differential surface
heating associated with the different vegetation types. The dry grasslands had high sensible heat and low
latent heat compared with the other vegetation types.
In Florida, consumptive use of water is draining the Everglades. Pine woodlands have been
cleared and marshes drained for agricultural and urban land (Figure 8.26, color plate). These landscape
20
Chapter 8 – Surface Climates
changes have altered the hydrology and climatology of south Florida (Pielke et al. 1999b). Simulations
with a mesoscale atmospheric model show July and August precipitation is reduced by 11% compared with
the natural landscape. This reduction in precipitation further reduces the amount of water in the Everglades
and is accompanied by greater sensible heat, less latent heat, and a warming of maximum surface air
temperature of 0.7 °C.
8.5 Hydrology
Much of the influence of land cover on surface climate arises from changes in soil water,
evapotranspiration, and the partitioning of net radiation into sensible and latent heat. Consequently, one
cannot clearly distinguish vegetation influences on climate from hydrologic influences. Both are intimate
components of the cycling of energy and water between land and atmosphere. However, two aspects of the
hydrologic cycle merit special attention.
8.5.1 Snow
The presence of snow on the ground affects surface climate in several ways. The thermal
conductivity of snow is much less than that of soil. A typical value is 0.34 W m-1 °C-1, which is one-third to
one-fifth that of mineral soil (Table 6.5). With a low thermal conductivity, less heat is transferred by
conduction. In winter, therefore, a deep snow pack on the ground acts as an insulating blanket that prevents
soil from cooling. Figure 8.27 illustrates this insulating effect of snow. Prior to the onset of snow in
November, air temperature and soil temperature are similar. As the air cools, the soil tracks air temperature
to within 1 °C. After a 10 cm snow pack covers the ground, the soil is several degrees warmer than air. By
January, the air has cooled to –10 °C but the soil is only –3 °C. Just as snow inhibits heat loss from the
underlying soil in winter, it inhibits soil warming in spring. Although the air warms by 10 °C from January
to March, the soil warms by less than 3 °C. The low thermal conductivity of snow prevents heat gain by the
soil. Once the snow is removed, the soil warms rapidly and again closely tracks air temperature.
The presence of snow on the ground is correlated with anomalously cold air temperatures (Foster
et al. 1983; Namias 1985; Leathers and Robinson 1993; Leathers et al. 1995). Figure 8.28 illustrates this
for a network of 91 stations in Northeast United States from Maine to West Virginia (Leathers et al. 1995).
21
Ecological Climatology
Long-term mean climatological data show that both maximum and minimum temperatures are several
degrees colder when snow is on the ground than when the ground is snow free. When averaged over the 6month winter season, the presence of snow is associated with 6 °C and 5 °C reductions in daily maximum
and minimum temperatures, respectively. Throughout much of the United States from the Central Plains to
the Atlantic coast, there is a strong negative correlation between snow cover and air temperature (Walsh et
al. 1982; Leathers and Robinson 1993). Snow cover in this region varies considerably from year to year
(Hughes and Robinson 1996). This suggests that unusually cold temperatures also characterize years with
anomalously high snow cover.
The correlation between snow cover and cold temperature clearly reflects variations in large-scale
atmospheric circulation. Cold air masses often produce snow in winter. However, climate model
simulations suggests the presence of snow feeds back to cool temperature. Snow reflects much more
incoming solar radiation compared with snow-free ground (Table 8.1), preventing the surface from
warming during the day. Moreover, a significant fraction of net radiation at the surface is used to melt
snow on warm days. Also, melting snow has a temperature of 0 °C, which is typically cooler than the air
above it. As a result, sensible heat is transferred from the warmer air to the colder surface. The presence of
snow is now recognized as an important condition required for accurate weather forecasts and climate
simulations (Barnett et al. 1988, 1989; Walsh and Ross 1988; Cohen and Rind 1991; Walland and
Simmonds 1997). This is especially true for Southeast Asia, where heavy snow cover in the Himalayas
reduces temperatures over land and weakens the summer monsoon.
8.5.2
Soil water
Storage of water in soils can lead to a positive feedback that increases the duration of a wet or dry
hydrologic state. Indeed, climate model studies show that hydrologic feedbacks contribute to and amplify
interannual precipitation variability over tropical and middle latitude land through the retention of
precipitation by soil and the influence of soil water on subsequent evapotranspiration (Shukla and Mintz
1982; Delworth and Manabe 1988, 1989; Koster and Suarez 1995, 1996; Koster et al. 2000). One means
by which this occurs is through recycling of precipitation (Figure 8.29). In the Amazon Basin, the average
annual precipitation is 1950 mm. More than half of this (1131 mm) is returned to the atmosphere as
22
Chapter 8 – Surface Climates
evapotranspiration. The remainder (819 mm) is lost as runoff. Three-quarters of the precipitated water
(1462 mm) comes from outside the basin. The remainder (488 mm or 25%) is obtained from local
evapotranspiration that is recycled as precipitation. Tropical deforestation interrupts this local recycling,
contributing to reduced rainfall. In the Mississippi Basin, 90% of the precipitated water is transported from
outside the region; only 10% of the annual precipitation is obtained from local recycling of
evapotranspiration.
Precipitation recycling can lead to a positive feedback in which wet soils pump more moisture
into the atmosphere, which enhances rainfall and further wets the soil. Conversely, dry soils, with their low
rates of evapotranspiration, can cause reduced rainfall. Climate model studies have routinely demonstrated
this feedback (Rind 1982; Shukla and Mintz 1982; Simmonds and Lynch 1992; Atlas et al. 1993; Dirmeyer
1994; Beljaars et al. 1996; Bonan and Stillwell-Soller 1998; Eltahir 1998; Pal and Eltahir 2001). This
feedback is difficult to demonstrate observationally, but a study of soil moisture and rainfall over a 14-year
period in Illinois found a positive correlation between initial soil water content and subsequent summer
rainfall (Findell and Eltahir 1997). High rates of evapotranspiration also tend to cool the atmosphere
because more net radiation at the surface is dissipated as latent heat than as sensible heat. Several
observational studies have found a negative correlation between summer temperature and precipitation
over much of interior United States (Madden and Williams 1978; Karl and Quayle 1981; Namias 1983;
Karl 1986; McNab 1989; Huang et al. 1996; Durre et al. 2000). The relationship is strongest in central
United States in the lower Mississippi River basin. In this region, hot summers are likely to also be dry;
wet summers are likely to be cool. Greater precipitation with wetter soils requires not only more moisture
in the atmosphere, but also the formation of clouds and enhanced convective processes. In some cases,
however, the cooler surface as a result of increased evapotranspiration may reduce atmospheric instability,
thereby reducing rather than increasing rainfall.
The severe 1988 drought in the Mississippi River basin and the extensive flooding in the same
region in 1993 may be examples of this precipitation-soil moisture-evapotranspiration feedback in which
soil moisture accentuates floods and droughts. Moisture budget analyses for these two events suggest that
although the events originated because of anomalous atmospheric circulation patterns, surface conditions
amplified these anomalies (Trenberth et al. 1988; Trenberth and Branstator 1992; Trenberth and Guillemot
23
Ecological Climatology
1996). Analyses suggest the 1993 flood was prolonged by increased evapotranspiration that increased
precipitation. This conclusion is supported by climate model studies of the July 1993 flood, which
demonstrate the importance of initial wet soils in sustaining the heavy precipitation (Beljaars et al. 1996;
Bosilovich and Sun 1999a,b; Viterbo and Betts 1999). In contrast, the 1988 drought was prolonged by
reduced evapotranspiration.
As with snow, it is now recognized that initial soil water content is an important determinant of
seasonal precipitation forecasts (Fennessy and Shukla 1999; Pielke et al. 1999a; Douville and Chauvin
2000; Koster et al. 2000; Small 2001). Furthermore, because of the importance of soil hydraulic properties
(e.g., porosity, field capacity) in determining available water and surface energy fluxes, soil texture can
have large impacts on surface energy fluxes and boundary layer development (Ek and Cuenca 1994;
Cuenca et al. 1996; Sun and Bosilovich 1996).
8.6 Microclimates and land use planning
By understanding the ecological functions that create surface climates and the specific landscape
features that alter these functions, the natural spatial variability in climate can be exploited to create
thermally efficient, water-conserving landscapes. We can always use technology to modify the prevailing
climate for human needs. Ski resorts make snow in winter to prolong the ski season. Farmers heat fruit
orchards in springtime to prevent frost damage to emerging blossoms. In arid climates, clouds are seeded
with dry ice or silver iodide to stimulate ice crystal formation and induce precipitation. Topographic,
edaphic, and biological features of the landscape create similar climate changes. The 3.5 °C difference in
air temperature between a northeast slope and a southwest slope (Figure 8.9) could be the difference
between a pleasant summer day or a hot, dry site. The typical 1 °C cooling of air with every 100 m gain in
elevation (Figure 8.10) is the difference between a long growing season at low elevations and a short
growing season in the mountains. Cold air drainage may cause a killing frost in a low-lying site while a
slightly higher location may bask in the relative warmth of a mid-slope thermal belt (Figure 8.13).
Orographic precipitation creates a stark contrast between moist and arid climates over a short distance
(Figure 8.14).
On a hot summer day, moderate temperatures and sea breezes (Figure 8.16) create
24
Chapter 8 – Surface Climates
refreshing coastal climates while inland cities may be suffering from a sweltering heat wave. One only has
to stand under a tree on a hot summer day to realize the shaded environment under the canopy is cooler
than open spaces (Figure 8.18). Windbreaks are a natural means of conserving energy by reducing heat
loss from strong winter winds (Figure 8.22). Even within a residential yard, the microclimatic differences
between a shady area and an open site baked to hot temperatures by the Sun can equal a thousand meters of
elevation or hundreds of kilometers of latitude; it may be the difference between a hot, semi-arid climate
and a cooler humid continental climate.
By understanding these climate changes, we can take advantage of natural landscape processes
and make climate work for us (e.g., Rahn 1979). Indeed, this is a recurring theme within the landscape
architecture and land use planning professions, which advocate understanding local environmental features
as part of the site planning process (Lynch and Hack 1984; Steiner 1991; Simonds 1998) and creating
designs in harmony with the environment, especially energy and water conservation (Robinette 1983,
1984; Moffat and Schiler 1994; Brown and Gillespie 1995). At the site scale, the use of vegetation to
conserve energy and to create thermally pleasant environments by blocking cold winter winds, shading hot
summer sun, and by evapotranspirational cooling is well documented (DeWalle et al. 1983; Parker 1983,
1987; Thayer et al. 1983; Enis 1984; Wagar 1984; Thayer and Maeda 1985; Heisler 1986; Wagar and
Heisler 1986; Huang et al. 1987; Hoyano 1988; Wilmers 1988; McPherson et al. 1988, 1989). In the arid
United States West, the use of native and low-water plants in urban landscapes to reduce water usage is
gaining momentum through xeriscape designs (Knopf 1991; Ellefson et al. 1992; O’Keefe 1992; Knopf et
al. 1995; Denver Water 1996; Huddleston and Hussey 1998). Precipitation can also be more effectively
managed and stored by vegetated landscapes that promote infiltration in contrast to hard surfaces that
promote runoff (Robinette 1984; Ferguson 1994, 1998; Thompson 1996).
25
Ecological Climatology
8.7 Tables
Table 8.1. Broadband albedo of various surfaces
Surface
Albedo
Natural
Fresh snow
0.80-0.95
Old snow
0.45-0.70
Desert
0.20-0.45
Glacier
0.20-0.40
Soil
0.05-0.40
Cropland
0.18-0.25
Grassland
0.16-0.26
Deciduous forest
0.15-0.20
Coniferous forest
0.05-0.15
Water
0.03-0.10
Urban
Road
0.05-0.20
Roof
0.08-0.35
Wall
0.10-0.40
Paint
White
0.50-0.90
Red, brown, green
0.20-0.35
Black
0.02-0.15
Source: Data from Oke (1987, pp. 12, 281).
26
Chapter 8 – Surface Climates
Table 8.2. Leaf orientation and reflection, transmission, and absorption of solar radiation by a leaf for
visible and near-infrared wavebands
Leaf
Vegetation
orientation
Visible
Reflected
Near-infrared
Transmitted Absorbed
Reflected
Transmitted Absorbed
Needleleaf tree
Random
0.07
0.05
0.88
0.35
0.10
0.55
Broadleaf tree
Semi-
0.10
0.05
0.85
0.45
0.25
0.30
0.11
0.07
0.82
0.58
0.25
0.17
horizontal
Grass, crop
Semivertical
Source: Data from Dorman and Sellers (1989).
27
Ecological Climatology
Table 8.3. Roughness length of various surfaces
Surface
Soil
Roughness length (m)
0.001-0.01
Grass
Short
0.003-0.01
Tall
0.04-0.10
Crop
0.04-0.20
Forest
1.0-6.0
Suburban
Low density
0.4-1.2
High density
0.8-1.8
Urban
Short building
1.5-2.5
Tall building
2.5-10
Source: Data from Oke (1987, pp. 57, 298).
28
Chapter 8 – Surface Climates
Table 8.4. Midday summer energy balance (W m-2) and temperature (°C) for bare ground and soil covered
with paper, hay, and black plastic mulch
Mulch
Bare ground
Paper
Hay
Black plastic
Net radiation, Rn
642
433
607
712
Sensible heat, H
362
349
489
635
Latent heat, λE
195
42
84
0
Soil heat, G
85
42
35
77
Temperature, Ts
40
40
51
52
Source: Data from Rosenberg et al. (1983, p. 196).
29
Ecological Climatology
Table 8.5. Surface energy budget (MJ m-2 day-1) and surface temperature (°C) for untreated and whitened
soils
Untreated soil
Whitened soil
Incoming solar radiation, S↓
27.2
27.2
Reflected solar radiation, S↑
8.2
16.3
-12.8
-9.8
Net radiation, Rn
6.2
1.1
Sensible heat, H
1.9
-2.5
Latent heat, λE
4.2
3.4
Soil heat, G
0.2
0.2
33 °C
28 °C
Net longwave radiation, L↓-L↑
Temperature, Ts
Source: Data from Stanhill (1965).
30
Chapter 8 – Surface Climates
8.8 Figure Legends
Figure 8.1. Variations in topography, soils, and land cover across a typical landscape that create mountain
and plains mesoclimates and numerous microclimates within a humid continental macroclimate.
Figure 8.2. Altitude, zenith, and incidence angles. Top: Horizontal surface. Bottom: Sloped surface.
Figure 8.3. Effect of angle of slope on diffuse radiation. The top left panel shows the sky as an inverted
bowl or hemisphere from which diffuse radiation is received. The point in the center receives diffuse
radiation from the entire hemisphere. The three right panels show a cross-section of this hemisphere for a
horizontal surface, a south-facing slope, and a wall. The gray area is the portion of the sky blocked. The
graph shows sky and terrain view factors in relation to angle of slope.
Figure 8.4. Direct beam and diffuse solar radiation on slopes. In all four panels, the Sun is due south with a
zenith angle of 27°. Direct beam radiation varies with angle of slope and aspect in relation to the incidence
angle. The portion of the sky from which diffuse radiation is received varies with angle of slope. Top left:
27° south-facing slope. Top right: 27° north-facing slope. Bottom left: 45° south-facing slope. Bottom
right: 45° north-facing slope.
Figure 8.5. Clear sky direct beam solar radiation as a function of time of day (vertical axis) and day of year
from January to December (horizontal axis). Gates (1980, pp. 96-147) describes atmospheric attenuation of
direct beam and diffuse radiation. Top: 20% southeast slope at latitude 30° N. Bottom: 42% north slope at
latitude 45° N.
Figure 8.6. Clear sky solar radiation as a function of time of day at the winter solstice (top) and summer
solstice (bottom) for a horizontal surface and north-, east-, south-, and west-facing walls at latitude 40° N.
Left: Direct beam radiation. Right: Direct beam and diffuse radiation. Diffuse radiation reflected by terrain
is neglected.
31
Ecological Climatology
Figure 8.7. Effect of slope and aspect on daily clear sky solar radiation (left) and daylength (right) for
January 15. Graphs are for north-, northeast-, east-, southeast-, south-, southwest-, west-, and northwestfacing slopes of 0% (horizontal) to 100% (45°). Diffuse radiation reflected by terrain is neglected. Top:
Latitude 30° N. Middle: Latitude 40° N. Bottom: Latitude 50° N.
Figure 8.8. As in Figure 8.7, but for July 15.
Figure 8.9. Maximum air temperature in a forest on a warm summer day in relation to slope and aspect.
Thin lines show elevation. Thick lines show temperature. Adapted from Geiger (1965, pp. 425-428).
Figure 8.10. Air temperature in relation to elevation on a 1500-m mountain. Top: Dry adiabatic cooling.
Bottom: Moist adiabatic cooling with precipitation.
Figure 8.11. Mountain and valley winds at night (top) and day (bottom).
Figure 8.12. Early morning air temperature in central Pennsylvania on a clear, cold spring night in relation
to elevation. Data from Hocevar and Martsolf (1971).
Figure 8.13. Mid-slope thermal belt on hillslopes in Austrian mountains. Top: Temperature in relation to
elevation at 0600, 1200, 1800, and 2400 hours. Middle: Frequency distribution of elevation with warmest
temperature. Bottom: Length of growing season for beech and maple trees in relation to elevation. Adapted
from Geiger (1965, pp. 435, 437, 440).
Figure 8.14. Elevation (top) and annual precipitation (bottom) along a west-to-east transect across the
United States from California to Maryland located at latitude 38.25° N. Elevation and precipitation are
from Figures 2.9 and 2.11 and have a spatial resolution of ½° (approximately 50 km).
Figure 8.15. Vertical temperature profile over the course of a year for Mirror Lake, New Hampshire.
Adapted from Likens (1985, p. 110).
Figure 8.16. Atmospheric circulation for land and sea breezes. The dashed lines are isobars of equal
pressure. Top: Daytime sea breeze. Bottom: Nighttime land breeze.
32
Chapter 8 – Surface Climates
Figure 8.17. Direct beam solar radiation in a forest based on a radiative transfer model (Sellers 1985). Left:
Canopy radiative transfer in a broadleaf forest with random leaf orientation in relation to leaf area index for
a zenith angle of 45°. The top panel shows the fraction of incident radiation absorbed by the canopy with a
soil albedo of 0.10 (visible) and 0.20 (near-infrared). The bottom panel shows canopy albedo for snowcovered ground with an albedo of 0.95 (visible) and 0.70 (near-infrared). Right: Canopy albedo in the
visible (top) and near-infrared (bottom) wavebands in relation to zenith angle for a broadleaf forest with a
leaf area index of 6 m2 m-2 and semi-horizontal, random, and semi-vertical leaf orientations.
Figure 8.18. Effect of forests on summer microclimates. Top: Vertical profile of daily average air
temperature above and within a forest canopy. Middle: Daily maximum and minimum air temperature in a
forest and with increasing distance into a clearing. Bottom: Difference in midday air temperature measured
in a clearing and in a forest in relation to clearing size. Adapted from Geiger (1965, pp. 321, 345, 352).
Figure 8.19. Shadow length in relation to tree height and solar altitude angle.
Figure 8.20. View factors in a forest clearing (Oke 1987, p. 353). Top: Sky and terrain view factors in a
clearing of diameter D surrounded by trees of height H. Bottom: Sky and terrain view factors in relation to
angle β.
Figure 8.21. Effect of clearing size D / H on nighttime radiation balance illustrated assuming L↓ = 250 W
m-2, Lv = 360 W m-2 (i.e, the radiative temperature is about 10° C), and the longwave radiation emitted by
the
ground
in
the
clearing
(L↑)
is
also
360
W
m-2.
Net
radiation
is
then
R = 250ψ
+ 360(1 − ψ
) − 360 .
n
sky
sky
Figure 8.22. Effects of trees on wind. The thick solid line represents a windbreak. Distance is measured
upwind and downwind of the windbreak as a proportion of tree height. Top: Wind speed in relation to
distance for low, medium, and high density vegetation. Bottom: Spatial pattern of wind speed around a
medium density windbreak. Adapted from Oke (1987, pp. 244, 245).
33
Ecological Climatology
Figure 8.23. Physical processes by which land affects climate and which are typically represented in the
land surface models used with climate models. Top left: Biogeophysics. Top right: Hydrology. Bottom:
Catchment hydrology and river flow.
Figure 8.24. Paired climate model simulations to examine the influence of altered vegetation cover on
climate. The atmosphere model provides temperature (T), wind (u,v), humidity (q), precipitation (P), solar
radiation (S↓) and longwave radiation (L↓) to the land model. The land model returns the surface fluxes of
latent heat (λE), sensible heat (H), momentum (τx,τy), reflected solar radiation (S↑), and emitted longwave
radiation (L↑). The coupled land-atmosphere model is integrated for many model years with a time step
typically less than 1 hour.
Figure 8.25. Mesoscale circulations created by wet and dry vegetation on a summer day. Data show the
difference (experiment minus control) between a simulation with a patch of dry grasses surrounded by wet
trees (experiment) and a control simulation with homogenous wet trees. The dry grass extends from x =
250 km to x = 450 km (thick horizontal line). Data are shown as a 2-dimensional cross-section of height
(left axis) and west-to-east distance (bottom axis). Top left: Air temperature. Top right: Humidity. Bottom
left: Horizontal zonal (west-to-east) wind. Positive values (stippling) indicate westerly wind. Negative
values (shaded) indicate easterly wind. The four large arrows show wind direction along the surface and
aloft. Bottom right: Vertical velocity. Positive values (stippling) indicate rising motion. Negative values
(shaded) indicate descending motion. The three large arrows indicate direction of motion. Adapted from
Seth and Giorgi (1996).
Figure 8.26. Landscape heterogeneity. Top left: Land cover of northern Georgia at a spatial resolution of 1
km. Atlanta is the large city near the center of the map. Top right: Land cover of the Central Plains at a
spatial resolution of ½°. Bottom: Land cover of southern Florida at a spatial resolution of 1 km. One-km
data provided by the U.S. Geological Survey Earth Resources Observation System (EROS) Data Center
(Sioux Falls, South Dakota). One-half degree data from Ramankutty and Foley (1999a).
34
Chapter 8 – Surface Climates
Figure 8.27. Influence of snow cover on soil temperature from a model of soil temperature based on heat
transfer and energy conservation (Chapter 6). The left axis shows air and soil temperature with no snow
cover (July to November), with a 10-cm snow pack (November to March), and again with no snow (March
to July). The right axis shows the depth of snow on ground.
Figure 8.28. Effect of snow cover on climatological mean daily maximum and minimum air temperature.
Data from Leathers et al. (1995).
Figure 8.29. Precipitation recycling in the Amazon Basin (top) and Mississippi Basin (bottom). Dark
arrows highlight precipitation recycling. Numbers in parentheses show fluxes as a percent of precipitation.
Data from Eltahir and Bras (1994, 1996). There is large uncertainty in these recycling estimates. Brubaker
et al. (1993) give recycling of 14-34% in these regions depending on season. Trenberth (1998, 1999) gives
higher recycling (34% for the Amazon and 21% for the Mississippi).
35