Download Adiabatic Heating and Cooling

153
C O N D E N S AT I O N
Meters
Feet
55,000
Cirrus
15,000
Cumulonimbus
40,000
12,000
Cirrocumulus
9000
25,000
Cirrostratus
6000
Altocumulus
Altostratus
Cumulus
3000
10,000
Stratocumulus
Smog
(Sea level) 0
● FIGURE
Stratus
Nimbostratus
Fog
0
6.10
Cloud classification scheme. Clouds are named based on their height and their form.
Observe this figure and Figure 6.11; what cloud type is present in your area today?
anywhere from 500 to 12,000 meters (1650–39,600 ft) above sea
level. From this base, they pile up into great rounded structures, often
with tops like cauliflowers.The cumulus cloud is the visible evidence
of an unstable atmosphere; its base is the point where condensation
has begun in a column of air as it moves upward.
Examine Figures 6.10 and 6.11 to familiarize yourself with
the basic cloud types and their names. Keep in mind that some
cloud shapes exist in all three levels—for example, stratocumulus
(strato = low level + cumulus = a rounded shape), altocumulus, and
cirrocumulus. These three share the similar rounded or cauliflower
appearance of cumulus clouds, which can exist at all three levels.
You may notice that altostratus (alto = middle level + stratus = layered shape) and cirrostratus have two-part names, but low-level layered clouds are called stratus only. Lastly, thin, stringy cirrus clouds
are found only as high-level clouds, so the term cirro (meaning
high-level cloud) is not necessary here.
Other terms used in describing clouds are nimbo or
nimbus, meaning precipitation (rain is falling). Thus, the nimbostratus cloud may bring a long-lasting drizzle, and the cumulonimbus
is the thunderstorm cloud. This latter cloud has a flat top, called
an anvil head, as well as a relatively flat base, and it becomes darker
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as it grows higher and thicker and thus blocks the incoming sunlight. The cumulonimbus is the source of many atmospheric concerns including high-speed winds, torrential rain, flash flooding,
thunder, lightning, hail, and possibly tornadoes. This type of cloud
can develop in several different ways as we will soon discuss.
Adiabatic Heating and Cooling The cooling process
that leads to cloud formation is quite different from that associated with the other condensation forms that we have already examined. The cooling process that produces fog, frost, and dew is
either radiation or advection. On the other hand, clouds usually
develop from a cooling process that results when a parcel of air on
Earth’s surface is lifted into the atmosphere.
The rising parcel of air will expand as it encounters decreasing atmospheric pressure with height. This expansion allows the
air molecules to spread out, which causes the parcel’s temperature
to decrease. This is known as adiabatic cooling and occurs at
the constant lapse rate of approximately 10°C per 1000 meters
(5.6°F/1000 ft). By the same token, air descending through the
atmosphere is compressed by the increasing pressure and undergoes adiabatic heating of the same magnitude.
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© Steve McCutcheon/ Visuals Unlimited
C H A P T E R 6 • M O I S T U R E , C O N D E N S AT I O N , A N D P R E C I P I TAT I O N
© C. Donald Ahrens
154
Cirrostratus
© Mark A. Schneider/ Visuals Unlimited
© Mark A. Schneider/ Visuals Unlimited
Cirrocumulus
Altostratus
M. Trapasso
© Ralph F. Kresge/ NOAA
Altocumulus
Stratocumulus
● FIGURE
Stratus
6.11
Types of clouds.
However, the rising and cooling parcel of air will eventually reach its dew point—the temperature at which water vapor begins to condense out, forming cloud droplets. From this
point on, the adiabatic cooling of the rising parcel will decrease
as latent energy released by the condensation process is added
to the air. To differentiate between these two adiabatic cooling
rates, we refer to the precondensation rate (10°C/1000 m) as the
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dry adiabatic lapse rate and the lower, postcondensation rate
as the wet adiabatic lapse rate. The latter rate averages 5°C per
1000 meters (3.2°F/1000 ft) but varies according to the amount
of water vapor that condenses out of the air.
A rising air parcel will cool at one of these two adiabatic
rates. Which rate is in operation depends on whether condensation is (wet adiabatic rate) or is not (dry adiabatic rate) occurring.
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NOAA/NWS
© John Cunningham/Visuals Unlimited
C O N D E N S AT I O N
Cumulonimbus
M. Trapasso
©Martin Miller/ Visuals Unlimited
Cirrus
Nimbostratus
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Nor
m
(6.5 al lap
s
°C/
100 e rate
0m
)
6000
4000
Dry a
diaba
2000
0
0°
tic ra
te (10
.0°C/
10°
Temperature of air at 1000 m
8000
Temperature of air at 2000 m
10,000
Temperature of rising
parcel of air at 2000 m
On the other hand, the warming temperatures of descending air allow it to hold greater quantities of water
vapor. In other words, as the air temperature rises farther above the dew point, condensation will not occur,
so the heat of condensation will not affect the rate of
rise in temperature. Thus, the temperature of air that is
descending and being compressed always increases at
the dry adiabatic rate.
It is important to note that adiabatic temperature
changes are the result of changes in volume and do
not involve the addition or subtraction of heat from
external sources.
It is also extremely important to differentiate between the environmental lapse rate and adiabatic lapse rates.
In Chapter 4, we found that in general the temperature
of our atmosphere decreases with increasing height above
Earth’s surface; this is known as the environmental lapse
rate, or the normal lapse rate. Although it averages 6.5°C
per 1000 meters (3.6°F/1000 ft), this rate is quite variable
and must be measured through the use of meteorological
instruments sent aloft. Whereas the environmental lapse
rate reflects nothing more than the vertical temperature
structure of the atmosphere, the adiabatic lapse rates are
concerned with temperature changes as a parcel of air
moves through the atmospheric layers ( ● Fig. 6.12).
Temperature of rising parcel of
air at 1000 m
6.11 (continued)
Altitude (m)
● FIGURE
Cumulus
1000
m)
20°
30°
Temperature (°C)
● FIGURE
6.12
Comparison of the dry adiabatic lapse rate and the environmental lapse rate. The environmental lapse rate is the average vertical change in temperature. Air displaced upward
will cool (at the dry adiabatic rate) because of expansion.
In this example, using the environmental lapse rate, what is the temperature of the
layer of air at 2000 meters?
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Stability and Instability Although
4000
4000
Height (m)
Height (m)
Stable
Unstable
adiabatic cooling results in the development
of clouds, the various forms of clouds are reDry
lated to differing degrees of vertical air move3000
3000
adiabatic
ment. Some clouds are associated with rapidly
Lapse
rate
rising, buoyant air, whereas other forms result
rate
when air resists vertical movement.
2000
2000
An air parcel will rise of its own accord
as long as it is warmer than the surrounding
layer of air. When it reaches a layer of the
1000
1000
atmosphere that is the same temperature as
Dry
Lapse
itself, it will stop rising. Thus, an air parcel
adiabatic
rate
rate
warmer than the surrounding atmospheric
air will rise and is said to be unstable. On
0
10
20
30
0
10
20
30
the other hand, an air parcel that is colder
Temperature (°C)
than the surrounding atmospheric air will
● FIGURE 6.13
resist any upward movement and will likely
Relationship between lapse rates and air mass stability. When air is forced to rise, it cools adiabatisink to lower levels. Then the air is said to
cally. Whether it continues to rise or resists vertical motion depends on whether adiabatic cooling is
be stable.
less rapid or more rapid than the prevailing vertical temperature lapse rate. If the adiabatic cooling
Determining the stability or instability
rate exceeds the lapse rate, the lifted air will be colder than its surroundings and will tend to sink
of an air parcel involves nothing more than
when the lifting force is removed. If the adiabatic cooling rate is less than the lapse rate, the lifted
asking the question, If an air parcel were lifted
air will be warmer than its surroundings and will be buoyant, continuing to rise even after the
to a specific elevation (cooling at an adiabatic
original lifting force is removed.
lapse rate), would it be warmer, colder, or the
In these examples, what would be the temperature of the lifted air if it rose to 2000 meters?
same temperature as the atmospheric air (determined by the environmental lapse rate at
that time) at that same elevation?
that aloft, and the environmental lapse rate will be low, thus enIf the air parcel is warmer than the atmospheric air at the selected
hancing stability. With the rapid heating of the surface on a hot
elevation, then the parcel would be unstable and would continue to
summer day, there will be a very steep environmental lapse rate
rise, because warmer air is less dense and therefore buoyant.Thus, under
because the air near the surface is so much warmer than that above,
conditions of instability, the environmental lapse rate must be greater
and instability will be enhanced.
than the adiabatic lapse rate in operation. For example, if the environPressure zones can also be related to atmospheric stability.
mental lapse rate is 12°C per 1000 meters and the ground temperature
In areas of high pressure, stability is maintained by the slow subis 30°C, then the atmospheric air temperature at 2000 meters would
siding air from aloft. In low pressure regions, on the other hand,
be 6°C. On the other hand, an air parcel (assuming that no condensainstability is promoted by the tendency for air to converge and
tion occurs) lifted to 2000 meters would have a temperature of 10°C.
then rise.
Because the air parcel is warmer than the atmospheric air around it, it
is unstable and will continue to rise ( ● Fig. 6.13).
Now let’s assume that it is another day and all the conditions
are the same, except that measurements indicate the environmental
lapse rate on this day is 2°C per 1000 meters. Consequently,
Condensed droplets within cloud formations stay in the air and do
although our air parcel if lifted to 2000 meters would still have
not fall to Earth because of their tiny size (0.02 mm, or less than
a temperature of 10°C, the temperature of the atmosphere at
1000th of an inch), their general buoyancy, and the upward move2000 meters would now be 26°C. Thus, the air parcel would be
ment of the air within the cloud. These droplets of condensation
colder and would sink back toward Earth as a result of its greater
are so minute that they are kept floating in the cloud formation;
density (see again Fig. 6.13). As you can see, under conditions of
their mass and the consequent pull of gravity are insufficient to
stability, the environmental lapse rate is less than the adiabatic
overcome the buoyant effects of air and the vertical currents, or
lapse rate in operation. If an air parcel, upon being lifted to a
updrafts, within the clouds. ● Figure 6.14 shows the relative sizes
specific elevation, has the same temperature as the atmospheric
of a condensation nucleus, a cloud droplet, and a raindrop. It takes
air surrounding it, it is neither stable nor unstable. Instead, it is
about a million cloud droplets to form one raindrop.
considered neutral; it will neither rise nor sink but will remain at
Precipitation occurs when the droplets of water, ice, or frozen
that elevation.
water vapor grow and develop masses too great to be held aloft.
Whether an air parcel will be stable or unstable is related to
They then fall to Earth as rain, snow, sleet, or hail. The form that
the amount of cooling and heating of air at Earth’s surface. With
precipitation takes depends largely on the method of formation
cooling of the air through radiation and conduction on a cool, clear
and the temperature during formation. Among the many theories
night, air near the surface will be relatively close in temperature to
Precipitation Processes
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P R E C I P I TAT I O N P R O C E S S E S
Typical raindrop
2 mm
Large cloud
droplet
Small cloud
droplets
Typical
cloud droplet
0.02 mm
•
Condensation
nucleus
0.0002 mm
(a)
● FIGURE
● FIGURE
6.14
The relative sizes of raindrops, cloud droplets, and condensation
nuclei.
If the diameter of a raindrop is 100 times larger than a cloud
droplet, why does it take a million cloud droplets to produce
one raindrop?
that try to explain the formation of precipitation, the collision–
coalescence process for warm clouds in low latitudes and the
Bergeron (or ice crystal) Process for cold clouds at higher
latitudes are the most widely accepted.
Precipitation in the lower latitudes of the tropics and in
warm clouds is likely to form by the collision–coalescence process. The collision–coalescence process is one in which the name
itself describes the process. By nature, water is quite cohesive (able
to stick to itself). When water droplets are colliding in the circulation of the cloud, they tend to coalesce (or grow together). This
is especially true as the water droplets begin to fall toward the
ground. In falling, the larger droplets overtake the smaller, more
buoyant droplets and capture them to form even larger raindrops.
The mass of these growing raindrops eventually overcomes the
updrafts of the cloud and fall to Earth, under the pull of gravity.
This process occurs in the warm section of clouds where all the
moisture exists as liquid water ( ● Fig. 6.15).
At higher latitudes, storm clouds can possess three distinctive layers. The lowermost is a warm layer of liquid water. Here
the temperatures are above the freezing point of 0°C (32°F). Above
this is the second layer composed of some ice crystals but mainly
supercooled water (liquid water that exists at a temperature below 0°C). In the uppermost layer of these tall clouds, when temperatures are lower than or equal to –40°C (–40°F), ice crystals will
dominate ( ● Fig. 6.16). It is in relation to these layered clouds
that Scandinavian meteorologist Tor Bergeron presented a more
complex explanation.
The Bergeron (or ice crystal) Process begins at great heights
in the ice crystal and supercooled water layers of the clouds.
Here, the supercooled water has a tendency to freeze on any
available surface. (It is for this reason that aircraft flying through
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Small droplets
captured in
wake
(b)
6.15
Collision and coalescence. (a) In a warm cloud consisting of small cloud
droplets of uniform size, the droplets are less likely to collide because
they are falling very slowly and at about the same speed. (b) In a cloud
of different-sized droplets, some droplets fall more rapidly and can overtake and capture some of the smaller droplets.
Why do these tiny droplets fall at different speeds?
middle- to high-latitude thunderstorms run the risk of severe icing and invite disaster.) The ice crystals mixed in with the supercooled water in the highest layers of the clouds can become freezing nuclei and form the centers of growing ice crystals. (Essentially,
this is the process that can also create snow.) As the supercooled
water continues to freeze onto these frozen nuclei, their masses
grow until gravity begins to pull them toward Earth. As this frozen precipitation enters the lower layer of the clouds, the abovefreezing temperatures there melt the ice crystals into liquid rain
before they hit the ground. Therefore, according to Bergeron, rain
in these clouds begins as frozen precipitation and melts into a liquid before reaching Earth.
As the melted precipitation falls through the lower, warmer
section of the cloud, the collision–coalescence process may take
over and cause the raindrops to grow even larger as they descend
toward the surface.
Major Forms of Precipitation
Rain, consisting of droplets of liquid water, is by far the most
common form of precipitation. Raindrops vary in size but are
generally about 2–5 millimeters (approximately 0.1–0.25 in.) in
diameter (see again Fig. 6.14). As we all know, rain can come in
many ways: as a brief afternoon shower, a steady rainfall, or the
deluge of a tropical rainstorm. When the temperature of an air
mass is only slightly below the dew point, the raindrops may be
very small (about 0.5 mm or less in diameter) and close together.
The result is a fine mist called drizzle. Drizzle is so light that it is
greatly affected by the direction of air currents and the variability
of winds. Consequently, drizzle seldom falls vertically.
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NOAA
stones have been known to kill animals and
humans.
Hail forms when ice crystals are lifted by
strong updrafts in a cumulonimbus (thunderIce crystals dominant
7600 m
(-40°C)
storm) cloud. Then, as these ice crystals circu(25,000 ft)
late around the storm cloud, supercooled water droplets attach themselves and are frozen
as a layer. Sometimes these pellets are lifted
up into the cold layer of air and then dropped
5500m
again and again. The resulting hailstone, made
Mixed ice and water
(18,000 ft)
(-20°C)
up of concentric layers of ice, has a frosty,
opaque appearance when it finally breaks out
of the strong updrafts of the cloud formation
and falls to Earth. The larger the hailstone, the
Freezing level (0°C)
more times it is cycled through the freezing
process and accumulated additional frozen
Liquid water only
layers.
On occasion, a raindrop can form and
1000 m
have
a temperature below 0°C (32°F). This
(3000 ft)
will occur when there is a shallow layer of
below-freezing temperatures all the way to
the ground so that the liquid rain can reach
a supercooled state. These supercooled droplets
will freeze the instant they fall onto a surface
● FIGURE 6.16
that is also at a below-freezing temperature.
The distribution of water, supercooled water, and ice crystals in a high-latitude storm cloud
The resulting icy covering on trees, plants,
according to the Bergeron Process theory.
and telephone and power lines is known as
What is the difference between water and supercooled water?
freezing rain (or glaze). People usually call
the rain and its blanket of ice an “ice storm”
( ● Fig. 6.18). Because of the weight of ice,
Snow is the second most common form of precipitation.
glazing can break off large branches of trees, bringing down teleWhen water vapor is frozen directly into a solid without first
phone and power lines. It can also make roads practically impasspassing through a stage as liquid water (or sublimation), it forms
able. A small counterbalance against the negative effects of glazing
minute ice crystals around the freezing nuclei (of the Bergeron
is the beauty of the natural landscape after an ice storm. Sunlight
Process). These crystals characteristically appear as six-sided, symcatches on the ice, reflecting and making a diamond-like surface
metric shapes. Combinations of these ice-crystal shapes make up
covering the most ordinary weeds and tree branches.
the intricate patterns of snowflakes. Snow will reach the ground
if the entire cloud and the air beneath the cloud maintain
below-freezing temperatures.
● FIGURE 6.17
Sleet is frozen rain, formed when rain, in falling to
Hailstones can be the size of golf balls, or even larger.
Earth, passes through a relatively thick layer of cold air near
What gives them their spherical appearance?
the surface and freezes. The result is the creation of small,
solid particles of clear or milky ice. In English-speaking
countries outside the United States, sleet refers not to this
phenomenon of frozen rain but rather to a mixture of rain
and snow.
Hail is a less common form of precipitation than
the three just described. It occurs most often during the
spring and summer months and is the result of thunderstorm activity. Hail appears as rounded lumps of ice,
called hailstones, which can vary in size from 5 millimeters (0.2 in.) in diameter and up to sizes larger than
a baseball ( ● Fig. 6.17). The world record is a hailstone
30 centimeters (12 in.) in diameter that fell in Australia.
Hailstones dropping from the sky can be highly destructive to crops and other vegetation, as well as to cars and
buildings. Though primarily a property destroyer, hail-
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P R E C I P I TAT I O N P R O C E S S E S
FEMA Photo/Michael Raphael
●
● FIGURE
6.18
An ice storm can cover a city with a dangerous glazing of ice.
Why are power failures a common occurrence with ice storms?
Factors Necessary for Precipitation
Three factors are necessary for the formation of any type of
precipitation on Earth. The first is the presence of moist air on
the surface. This air obviously represents the source of moisture
(for the precipitation) and energy (in the form of latent heat of
condensation). Second are the condensation nuclei around which
the water vapor can condense, discussed earlier in this chapter.
Third is a mechanism of uplift. These uplift mechanisms are responsible for forcing the air higher into the atmosphere so that it
can cool down (by the dry adiabatic rate) to the dew point. These
uplift mechanisms are vital to the process of precipitation.
A parcel of air can be forced to rise in four major ways. All
the precipitation that falls anywhere on Earth can be traced back
to one of these four uplift mechanisms ( ● Fig. 6.19):
■
■
■
■
Convectional precipitation results from the displacement of
warm air upward in a convectional system.
Frontal precipitation takes place when a warm air mass rises
after encountering a colder, denser air mass.
Cyclonic (or convergence) precipitation occurs when air
converges upon and is lifted up into a low pressure system.
Orographic precipitation results when a moving air mass
encounters a land barrier, usually a mountain, and must rise
above it in order to pass.
Convectional Precipitation The simple explanation
of convection is that when air is heated near the surface it expands, becomes lighter, and rises. It is then displaced by the cooler,
denser air around it to complete the convection cycle. The important factor in convection for our discussion of precipitation is
that the heated air rises and thus fulfills the one essential criterion
for significant condensation and, ultimately, precipitation.
To enlarge our understanding of convectional precipitation,
let’s apply what we have learned about instability and stability.
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Figure 6.20 illustrates two different cases in which air
rises due to convection. In both, the lapse rate in the free
atmosphere is the same; it is especially high during the
first few thousand meters but slows after that (as on a hot
summer day).
In the first case (Fig. 6.20a), the air parcel is not very
humid, and thus the dry adiabatic rate applies throughout
its ascent. By the time the air reaches 3000 meters (9900 ft),
its temperature and density are the same as those of the surrounding atmospheric air. At this point, convectional lifting
stops.
In the second case (Fig. 6.20b), we have introduced
the latent heat of condensation. Here again, the unsaturated rising column of air cools at the dry adiabatic rate
of 10°C per 1000 meters (5.6°F/1000 ft) for the first
1000 meters (3300 ft). However, because the air parcel is
humid, the rising air column soon reaches the dew point,
condensation takes place, and cumulus clouds begin to
form. As condensation occurs, the heat locked up in the
water vapor is released and heats the moving parcel of air,
retarding the adiabatic rate of cooling so that the rising
air is now cooling at the wet adiabatic rate (5°C/1000
meters). Hence, the temperature of the rising air parcel
remains warmer than that of the atmospheric layer it is passing
through, and the air parcel will continue to rise on its own. In
this case, which incorporates the latent heat of condensation,
we have massive condensation, towering cumulus clouds, and a
thunderstorm potential.
Convectional precipitation is most common in the humid
equatorial and tropical areas that receive much of the sun’s energy and in summer in the middle latitudes. Though differential
heating of land surfaces plays an important role in convectional
precipitation, it is not the sole factor. Other factors, such as surface topography and atmospheric dynamics associated with the
upper air winds, may provide the initial upward lift for air that is
potentially unstable. Once condensation begins in a convectional
column, additional energy is available from the latent heat of condensation for further lifting.
This convectional lifting can result in the heavy precipitation,
thunder, lightning, and tornadoes of spring and summer afternoon
thunderstorms. When the convectional currents are strong in the
characteristic cumulonimbus clouds, hail can result.
Frontal Precipitation The zones of contact between
relatively warm and relatively cold bodies of air are known
as fronts. When two large bodies of air that differ in density, humidity, and temperature meet, the warmer one is lifted
above the colder. When this happens, the major criterion for
large-scale condensation and precipitation is once again met.
Frontal precipitation thus occurs as the moisture-laden warm
air rises above the front caused by contact with the cold air.
Continuous frontal precipitation has caused some devastating
floods through time.
To fully understand fronts, we must examine what causes unlike bodies of air to come together and what happens when they
do.This will be discussed in Chapter 7, where we will take a more
detailed look at frontal disturbances and precipitation.
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Warm air
Convectional
Cyclonic (Convergence)
Rain
shadow
Warm air
Cold air
Front
Frontal
● FIGURE
Orographic
6.19
The principal cause of precipitation is upward movement of moist air resulting from convectional,
frontal, cyclonic, or orographic lifting.
What kind of air movement is common to all four diagrams?
Cyclonic (Convergence) Precipitation The third
mechanism, the cyclonic (also known as convergence), was first
introduced in Chapter 5 (see again Fig. 5.4). When air enters a
low pressure system, or cyclone, it does so (in a counterclockwise
fashion in the Northern Hemisphere) from all directions. When
air converges on a low pressure system, it has little option but to
rise. Therefore, clouds and possible precipitation are common
around the center of a cyclone.
Orographic Precipitation As was the case with convectional rainfall, orographic rainfall has a simple definition and a
somewhat more complex explanation. When land barriers—such
as mountain ranges, hilly regions, or even the escarpments (steep
edges) of plateaus or tablelands—lie in the path of prevailing
winds, large portions of the atmosphere are forced to rise above
these barriers. This fills the one main criterion for significant precipitation—that large masses of air are cooled by ascent and expansion until large-scale condensation takes place. The resultant
precipitation is termed orographic (from Greek: oros, mountains).
As long as the air parcel rising up the mountainside remains stable
(cooling at a greater rate than the environmental lapse rate), any
resulting cloud cover will be a type of stratus cloud. However, the
situation can be complicated by the same circumstances illustrated
in Figure 6.20b. A potentially unstable air parcel may need only
the initial lift provided by the orographic barrier to set it in motion. In this case, it will continue to rise of its own accord (no
longer forced) as it seeks air of its own temperature and density.
Once the land barrier provides the initial thrust, it has performed
its function as a lifting mechanism.
Because the air deposits most of its moisture on the windward side of a mountain, there will normally be a great deal less
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precipitation on the leeward side; on this side, the air will be
much drier and the dew point consequently much lower. Also, as
air descends the leeward slope, its temperature warms (at the dry
adiabatic rate), and condensation ceases. The leeward side of the
mountain is thus said to be in the rain shadow ( ● Fig. 6.21a). Just
as being in the shade, or in shadow, means that you are not receiving any direct sun, so being in the rain shadow means that you do
not receive much rain. If you live near a mountain range, you can
see the effects of orographic precipitation and the rain shadow in
the pattern of vegetation (Figs. 6.21b and c). The windward side
of the mountains (say, the Sierra Nevada in California) will be
heavily forested and thick with vegetation. The opposite slopes in
the rain shadow will usually be drier and the cover of vegetation
sparser.
Distribution of Precipitation
The precipitation a region receives can be described in different
ways. We can look at average annual precipitation to get an overall
picture of the amount of moisture that a region gets during a year.
We can also look at its number of raindays—days on which 1.0
millimeter (0.01 in.) or more of rain is received during a 24-hour
period. Less than this amount is known as a trace of rain. If we
divide the number of raindays in a month or year by the total
number of days in that period, the resulting figure represents the
probability of rain. Such a measure is important to farmers and
to ski or summer resort owners whose incomes may depend on
precipitation or the lack of it.
We can also look at the average monthly precipitation. This
provides a picture of the seasonal variations in precipitation
( ● Fig. 6.22). For instance, in describing the climate of the west
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P R E C I P I TAT I O N P R O C E S S E S
−18°
−15°
−12°
5500
−6°
4500
0°
0°
3500
3000
10°
Dry
adiabatic
rate
(10°C/
1000 m)
8°
20°
18°
24°
30°
Temp (°C)
(a)
● FIGURE
2000
1000
−3°
10°
500
Wet
adiabatic
rate
(5°C/1000 m)
4°
15°
13°
18°
Condensation
begins
20°
Dry adiabatic
rate
(10°C/1000 m)
24°
30°
30°
Temp (°C)
5°
8°
1500
13°
−9°
0°
2500
4°
0°
−6°
4000
−3°
−15°
−12°
5000
−9°
−5°
−18°
6000
Altitude
Temp (°C)
(m)
Existing lapse rate
30°
Temp (°C)
(b)
6.20
Effect of humidity on air mass stability. (a) Warm, dry air rises and cools at the dry adiabatic rate, soon becoming the same temperature as the surrounding air, at which point convectional uplift terminates. Because the
rising dry air did not cool to its dew point temperature by the time that convectional lifting ended, no cloud
formed. (b) Rising warm, moist air soon cools to its dew point temperature. The upward-moving air subsequently cools at the wet adiabatic rate, which keeps the air warmer than the surrounding atmosphere so that
the uplift continues. Only when all moisture is removed by condensation will the air cool rapidly enough at
the dry adiabatic rate to become stable.
What would be necessary for the cloud in (b) to stop its upward growth at 4500 meters?
coast of California, average annual precipitation would not give
the full story because this figure would not show the distinct wet
and dry seasons that characterize this region.
Horizontal Distribution of Precipitation
● Figure 6.23 shows average annual precipitation for the world’s continents. We can see that there is great variability in the distribution of precipitation over Earth’s surface. Although there is a
zonal distribution of precipitation related to latitude, this distribution is obviously not the only factor involved in the amount
of precipitation an area receives.
The likelihood and amount of precipitation are based on two
factors. First, precipitation depends on the degree of lifting that
occurs in air of a particular region. This lifting, as we have already
55061_06_Ch06_p140_169 pp3.indd 161
seen, may be due to the collision of different air masses (frontal), to
the convergence of air into a low pressure system (cyclonic or convergence), to differential heating of Earth’s surface (convection), to
the lifting that results when an air mass encounters a rise in Earth’s
surface (orographic), or to a combination of these processes. The
second factor affecting the likelihood of precipitation depends on
the internal characteristics of the air itself, including its degree of
instability, its temperature, and its humidity.
Because higher temperatures, as we have seen, allow air masses
to hold greater amounts of water vapor and because, conversely, cold
air masses can hold less water vapor, we can expect a general decrease of precipitation from the equator to the poles that is related to
the unequal zonal distribution of incoming solar energy discussed in
Chapter 3.
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162
C H A P T E R 6 • M O I S T U R E , C O N D E N S AT I O N , A N D P R E C I P I TAT I O N
However, if we look again at Figure 6.23, we see a great
deal of variability in average annual precipitation beyond the
general pattern of a decrease with increased latitude. In the
following discussion, we examine some of these variations and
give the reasons for them. We also apply what we have already
Orographic
Precipitation
Clouds
2500 m
Rate of cooling
after condensation
5.0°C/1000 m
Condensation
level
1500 m
Rate of cooling
10.0°C/1000 m
500 m
Uplift
19°C
0
(a)
4.0°C
Windward
slope
9°C
learned about temperature, pressure systems, wind belts, and
precipitation.
Distribution within Latitudinal Zones The
equatorial zone is generally an area of high precipitation—more
than 200 centimeters (79 in.) annually—largely
due to the zone’s high temperatures, high
humidity, and the instability of its air. High
Rain
temperatures and instability lead to a general
Shadow
pattern of rising air, which in turn allows for
precipitation. This tendency is strongly reinRate of warming
forced by the convergence of the trades as they
10.0°C/1000 m
move toward the equator from opposite hemispheres. In fact, the intertropical convergence
Leeward
zone is one of the two great zones where air
slope
masses converge. (The other is along the polar
14°C
front within the westerlies.)
In general, the air of the trade wind zones
is stable compared with the instability of the
equatorial zone. Under the control of these
24°C
steady winds, there is little in the way of atmospheric disturbances to lead to convergent
or convectional lifting. However, because the
trade winds are basically easterly, when they
Station:
Latitude:
San Francisco
38°N
Average annual prec.:
12.8°C (55°F)
Mean annual temp.:
R. Gabler
°F
100
80
(b)
Longitude:
55 cm (21.7 in.)
Range:
122°W
7.2°C (13°F)
°C
Cm
30
70
20
60
In.
30
25
60
10
50
40
20
0
40
20
15
−10
0
R. Gabler
−20
−40
(c)
● FIGURE
30
−20
10
20
−30
5
10
−40
6.21
Orographic precipitation and the rain-shadow effect.
(a) Orographic uplift over the windward (western) slope of
the Sierras produces condensation, cloud formation, and
precipitation, resulting in (b) dense stands of forest. (c) Semiarid or rain-shadow conditions occur on the leeward (eastern)
slope of the Sierras.
Can you identify a mountain range in Eurasia in which the
leeward side of that range is in the rain shadow?
55061_06_Ch06_p140_169 pp3.indd 162
J
● FIGURE
F
M
A
M
J
J
A
S
O
N
D
6.22
Average monthly precipitation in San Francisco, California, is represented by colored
bars along the bottom of the graph. A graph of monthly precipitation figures like this
one gives a much more accurate picture than the annual precipitation total, which
does not tell us that nearly all the precipitation occurs in only half of the year.
How would this rainfall pattern affect agriculture?
6/5/08 10:22:34 PM
163
P R E C I P I TAT I O N P R O C E S S E S
GEOGRAPHY’S PHYSICAL SCIENCE PERSPECTIVE
The Lifting Condensation Level (LCL)
W
reached. Any additional lifting and clouds
will form and build upward. Therefore, the
height at which clouds form from lifting is
called the lifting condensation level (LCL)
and can be estimated by the equation:
LCL (in meters) = 125 meters × (Celsius
temperature – Celsius dew point)
For example, if the surface temperature
is 7.2°C (45°F) and the dew point temperature is 4.4°C (40°F), then the LCL is
estimated at 350 meters (1148 ft) above
the surface.
Caution: Keep in mind that different
layers of clouds may exist at the same
time. Low, middle, and high clouds as defined in this chapter may all appear on the
same afternoon. These clouds may have
formed in other regions and be only passing overhead. The formula presented here
is best used with the lowest level of cloud
cover that appears overhead.
M. Trapasso
hen you look at clouds in our
atmosphere, it is often quite
easy to see their relatively flat
bases. Cloud tops may appear quite irregular, but cloud bases are often flat. Even if
the cloud bases do not seem flat, it will be
obvious that the clouds you see all seem
to be formed at the same level above the
surface. This level represents the altitude to
which the air must be lifted (and cooled at
the dry adiabatic rate) before saturation is
The stratocumulus clouds (bottom layer) show the lifting condensation level (LCL).
move onshore along east coasts or islands with high elevations,
they bring moisture from the oceans with them. Thus, within the
trade wind belt, continental east coasts tend to be wetter than
continental west coasts.
In fact, where the air of the equatorial and trade wind
regions—with its high temperatures and vast amounts of moisture—moves onshore from the ocean and meets a landform
barrier, record rainfalls can be measured. The windward slope
of Mount Waialeale on Kauai, Hawaii, at approximately 22°N
latitude, holds the world’s record for greatest average annual
rainfall—1168 centimeters (460 in.).
Moving poleward from the trade wind belts, we enter the
zones of subtropical high pressure where the air is subsiding. As
55061_06_Ch06_p140_169 pp3.indd 163
it sinks lower, it is warmed adiabatically, increasing its moistureholding capacity and consequently reducing the amount of precipitation in this area. In fact, if we look at Figure 6.23, which
shows average annual precipitation on a latitudinal basis, we can
see a dip in precipitation level corresponding to the latitude of
the subtropical high pressure cells. These areas of subtropical high
pressure are in fact where we find most of the great deserts of the
world: in northern and southern Africa, Arabia, North America,
and Australia. The exceptions to this subtropical aridity occur
along the eastern sides of the landmasses where, as we have already noted, the subtropical high pressure cells are weak and wind
direction is often onshore. This exception is especially true of regions affected by the monsoons.
6/5/08 10:22:38 PM