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Chapter XIII. Water in the Local Air

Chapter XIII. Water in the Local Air

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346



XIII.



WATER I N THE LOCAL AIR



vapor concentration. The profile of humidity decreases upward in an

exponential decay curve log eh = log e , - hi6.5, in which e represents

vapor pressure at sea level (subscript 0) and at any height h (in

kilometers) (Hann, 1897, p . 279). The removal mechanisms operate

throughout the troposphere but with decreasing effect a t the higher

altitudes where they find less vapor to condense on the average.

Atmospheric vapor concentration is greatest in the local air that bathes

the ecosystems on the earth's surface.

Table I shows values of vapor pressure (in millibars) at typical wet

surfaces in nature, and the sea-level specific humidity of moist air in

contact with these surfaces (as grams of vapor per kilogram of air). The

dryness of cold high-latitude surfaces and of the air above them form

one extreme. The other extreme is found in the high humidities of

equatorial oceans or rainforest and the air in contact with them, in

which vapor concentration is more than two orders of magnitude

greater than in the high latitudes, as the surfaces approach the 32-34°C

temperature limit discussed in Chapter X.



Regimes

Diurnal Regimes The diurnal regime of vapor pressure in the air

near the evaporating surfaces depicts the response of a stor'ige

function in the local air, which buffers moisture input from evapotranspiration during dayIight hours to a varying outflow, the upward

mixing of vapor. Over many land surfaces nocturnal inversions that



TABLE I



V a p o r Pressure a n d Specific H u m i d i t y a t Selected Surfaces"

Corresponding humidities

Typical

surface

Vapor

temperature pressure Specific humidity

("C)

(mb)

(g kg-7



Surface type

Snow surface of ice cap

Mid-latitude snow field i n winter

Melting snow

Mid-latitude water or vegetation

Low-latitude ocean or forest

Under intense radiation

-



-



-



-



-40

20

0

+ 15

+25

33

-



+

-



' Source: Smithsonian Tables (1966).



0.13

1

6

17

32

50



0. I

1

4

11

20

32



WATER VAPOR I N THE LOCAL A I R



347



persist into the forenoon hours limit upward mixing and a maximum

in the daily vapor pressure cycle occurs in the late morning.

This often is followed by a minimum at the time when upward

mixing of vapor out of the local air is most vigorous. A second

maximum may follow, when turbulent mixing diminishes in the

evening. The primary minimum, at night, indicates the lack of input

from the surface. Sometimes, in fact, a net downward flow of vapor

takes place out of the air to the surface, and the vapor pressure in the

local air declines still more.

The vapor concentration in the local air over water surfaces, which

generate a fairly constant upward flow of vapor, exhibits a different

diurnal regime. This regime is primarily influenced by the variations

in upward mixing of vapor out of the local air into the free atmosphere. The maximum vapor pressure therefore comes at the time of

weakest vertical mixing, near sunrise, and the lowest at the time of

strongest mixing.



Annual Regime The dependence of atmospheric vapor pressure on

the rate of evaporation from the underlying surface influences its

variation through the year. Over Lake Ontario, for example, it varies

from 4.5 in late winter to 20 mb in late summer. The effect of the

underlying surface is further demonstrated by the fact that during the

months of active evaporation from the lake, the atmospheric vapor

pressure over i t is substantiallv greater than over the land adjacent

(Richards and Fortin, 1962). The ratio of l a k e - b l a n d values shown

over the annual cycle in Fig. XIII-1 is largest-about

1.3-in midwin-



MONTHS



Fig. XIII-I. Annual regime of the ratio between vapor pressure over Lake Ontario to

that over adjacent land surfaces (Richards and Fortin, 1962).



348



XIII.



W A T E R IN T H E LOCAL A I R



ter, when evaporation is weak from the frozen land but active from the

lake. The ratio during May and June, on the other hand, indicates the

coldness of the lake surface relative to the surrounding lands; vapor

pressure is lower over the lake than over the land.

Annual regimes of vapor pressure at four places in Australia are

given in Table 11. Between winter and summer at Sydney, vapor

pressure doubles, a response more to the change in evaporation from

land than to the smaller change in sea-surface temperature and

evaporation. The author’s experience with Sydney humidity in February and March, however, confirms the high vapor pressures then,

which are associated with the continued warmth of the near-shore

waters. Days of easterly (on-shore) winds are especially sultry.

The change between winter and summer at Darwin is large, 13 mb.

The whole annual cycle is found at a higher level than at the other

stations; this indicates the warmth of the Timor Sea. The large change

through the year indicates the effect of a small change in a high seasurface temperature as well as the occurrence of dry air in winter (June

and July) moving from the interior of the continent. In the Northern

Hemisphere continents, such movement of dry air also occurs in

winter, but it is much drier than in Australia.



The World Pattern

The air over the cold snow surfaces of the northern continents in

winter is particularly dry. Over great areas, vapor pressure is less than

2 mb. These are source regions of very dry air, which moves south and

dominates many winter days in the midwestern United States.

The contrasting surfaces of moist warm areas of the equatorial

latitudes average higher than 25 m b in vapor pressure. Latitudinal

means of vapor pressure over the land are shown in Table 111, which

TABLE I1



V a p o r Pressure in Australiaa

Location

Darwin, 12s

Brisbane, 27s

Sydney, 34s

Hobart, 435



J

31

21

18

9



F

31

22

19

12



M

31

20

18

11



A

27

17

15

10



M

22

14

12

9



J



J

18

12

10

8



A



S



O



N



D



18 21 25 28

11 11 13 15

10 10 11 13

9

8

8

8



Units: mb. Source: Commonwealth of Australia Yearbook (1965).



29

18

15

9



M

30

20

17

10



e



a

26.0

16.4

13.9

9.3



n



349



WATER V A P O R I N THE LOCAL AIR



TABLE I11



V a p o r Pressure



in



the Air near the Surface o f



the Continents'

Latitude

90N

60

40

30

20

1G

Equator

10s

20

30

40

60

90



Jan




2

5

8

14

22

28

28

22

16

14



4



July

3

12

18

20

23

28



28

25

15

10

8

1



'' Units: mb. Source: Kessler (1968, p . 14).

shows that the zone of high humidity extends from 20s to 1ON

latitude.

Figure XIII-2, the world pattern in July (Landsberg, 1964), displays

large areas where vapor pressure exceeds 30 mb. The Caribbean Sea,

the Gulf of Mexico, and their coastal lands are well known sources of



Fig. XIII-2.



Average vapor pressure (mb) in July (Landsberg, 1964)



350



XIII.



WATER I N THE LOCAL AIR



humid air streams. Days of sweltering humidity in central North

America occur during southerly airflow, which on its traverse across

the evaporating fields and forests of the lower Mississippi Valley gains

even more moisture.

It is significant that, when we look for the sources of the rain and

snow that fall on central and eastern North America, we see over the

western Pacific Ocean a vapor pressure that lies between 12 and 15 mb

during most of the year. Over the Arctic Ocean we see a vapor

pressure lower than 10 mb. In contrast, over the Gulf of Mexico we see

a vapor pressure of 20 mb in winter and 30 nib in summer. Which

body of water is most likely to provide moisture for the central plains

of the continent?



Humidity As a Component of the Environment

Vapor content is affected by the source of the airstream, but this

quality changes as the air moves over drier or moister land. Characteristics of the boundary layer and even more of the local air reflect

conditions at the directly underlying surface. Movement of the local

air is braked over rough ecosystem canopies and its warmth, C 0 2 , and

water-vapor content are modified by the fluxes of sensible heat, CO.,,

and vapor from the ecosystems that make up its porous lower

boundary.

Over an ecosystem surface having "infinite" extent, as was hyyothesized in the definition of potential evapotranspiration in an earlier

chapter, the partnership between ecosystem and local air tends to be

dominated by the ecosystem. The rhythms of warmth, moisture, and

so on, in the local air are forced by the rhythms of the corresponding

fluxes at the underlying surface.

In reality, this hypothetical situation is rare. Most terrestrial ecosystems are limited in size, often covering an area of only a few hectares,

100 m or so across. The air that fills their foliage volumes is still

dominated by exchanges of water and energy at the leaf surfaces, as

evidenced by the vertical profiles of humidity or temperature within a

forest stand; but above these systems the main body of the local air

moves on across the countryside, floating above a mosaic of many

contrasting ecosystems and carrying, for example, the moisture it

acquires from one system over a neighboring system.

For this reason, the expressions for evapotranspiration discussed

previously include terms for atmospheric humidity. The Bowen ratio

can be approximated by measuring the differences in temperature and

moisture between an ecosystem and the overlying air. The gradient of



WATER V A P O R I N THE LOCAL AIR



351



moisture is important in the evaporation process and depends on the

concentration of vapor within the storage zone represented by the

local air.

One effect of high specific humidity is, by reducing transpiration, to

slow the flow of nutrients brought into a plant in the stream of water

moving from soil to leaves. This decrease in nutrient intake might well

account for the stunted growth of trees in perennially cloudy, foggy

zones of some mountains (Odum, 1971, p. 376).

The local air, not being precisely defined as to thickness, forms a

reservoir of indefinite size for water vapor. We can obtain a rough

idea of its storage capacity from the following: Assume it to be 1 km

deep; it then contains 1300 kg of air in a column of 1-mz crosssectional area. Over an ice cap, at a specific humidity of 0.1 g kg-' (see

Table I), this column contains about 130 g of vapor m-' of surface.

Over an equatorial forest, at a specific humidity of 20 g kg-' (Table I

again), it contains 26 kg of vapor. Obviously, the storage of water in

this zone can vary tremendously. Suppose we take a typical midlatitude instance of 10 kg of vapor stored m-' of area. This value, 10 kg

m-2

. of the same general size as the mass of water delivered to the

, is

surface in a day of moderate rain (10 kg m-' day-'), or the amount

evaporated from a corn field in two or three days of hot summer

weather.

Conditions of high moisture concentration slow down the further

transfer of water from the ecosystems into the air, as quantitatively

shown in the various evaporation formulas. Dry air, on the other

hand, accepts vapor from below avidly, and brings about high

evaporation rates best exemplified where desert air invades an oasis.

The cool air we feel as we drive from the desert onto a road between

irrigated alfalfa fields is a consequence of the fact that available energy

at the alfalfa system is channeled almost entirely into evaporation,

leaving little or none to warm the air. The moisture stored or not

stored in the local air thus is an important factor in the environment of

an ecosystem, affecting its rate of evaporation and indirectly its

openness to assimilate CO, from the air.

Gentilli (1955) investigated the generally accepted correspondence of

dew-point temperature (a common measure of vapor pressure or

specific humidity) with the minimum air temperature reached during

the preceding night (a common forecast variable). The correspondence

in the Plains States and eastward is good, especially in seasons when

nights are warmer than about -10°C. It demonstrates the important

role played by water in the atmosphere. Water vapor effectively

transfers atmospheric energy to the ground by radiation, and main-



352



XIII.



W A T E R I N THE LOCAL A I R



tains ecosystems in an equable environment during the stress period

of the night. This energetic bond from vapor in the local air to

moisture at the surface means that vapor concentration influences

snow melting, freeze-thaw cycles, evaporation, and dew formation.

This income of 30-50 W m-’ in long-wave flux density is particularly

important in the small energy budgets of night hours, when the

warmth it brings accelerates plant respiration and reduces net photosynthetic productivity of ecosystems, and adds to human heat stress in

urban systems.



Humidity in the Environment of M a n Moisture in the local air also

affects man when i t reaches high values. The isarithm of 20 mb in Fig.

XIII-2 is of interest because this value is generally accepted a s an index

of sultriness in the human environment. In Asia this line takes in most

of China and Japan in July; it covers the eastern Mediterranean and its

southern shores. In North America it takes in the eastern part of the

continent to a latitude of 40N. The zonal averages of Table I11 show

that the 20-mb area stretches from 15s to 30N latitude and encompasses a major fraction of the world.

Air conditioning, the American term for artificial cooling and drying

of summer air, has become one of the major consumers of electric

power in this country. A rough indication of the desire for it is given

by a combined index* of air temperature and dew point. Dew point i s

uniquely related to vapor pressure and is weighted about a third as

heavily as air temperature in the formu1ation.t Table IV shows the

frequency distribution of the units of this index at Baltimore in

midsummer, when evapotranspiration in eastern North America i s

most active. At index values above 75 units about half the people are

uncomfortable. These values occur 0.82 of the time in the afternoon

and even at night Q.09 of the time. At index values above 80 units

nearly everyone is uncomfortable; such values occur 0.50 of the

afternoon hours at Baltimore, in the open air. In the city, and

especially in buildings, the frequency of occurrence is larger.

The physiological effects of moisture in the atmosphere are not well

understood, beyond the obvious fact that high vapor pressure suppresses evaporative cooling of the heat-stressed human body. One

panel (Sargent et al., l967) also notes that ”there is need to know the

* At first called the ”discomfort index” (Thorn, 1956), this was later euphemized to

”temperaturehumidity index.”

t A similar index, called “effective temperature” (Landsberg, 1969, p. 54) is ‘tffectetl

about the same by an increase of 6 m b in vapor pressure alone (= 4 g kg ‘ increase in

specific humidity) as by a 6°C rise in air temperature alone.



353



V I S I B L E FORMS OF WATER I N THE L O C A L AIR



TABLE IV

Frequency Distribution of the Discomfort index (or

Temperature-Humidity Index) at Baltimore in July"



Scale units

<60

60-64

6569

7&74

7579

8G84



8589



Night

0000-0500

5

22

21

43



Afternoon

1200-1700

0

0



1



0

0



17

32

38

12



100



100



9



'' Units: percent of occurrences. Source: Thorn

(1956).



chronic effects of exposure to very low humidity, for man now spends

so much of his life in artificial atmospheres that may be exceedingly

dry." As midwesterners know, heating polar air at a vapor pressure of

only 2 4 mb without humidifying it results in very low relative

humidity in the home and severe respiratory problems.



VISIBLE FORMS OF WATER IN THE LOCAL AIR



Most atmospheric water is in the vapor state, but sometimes a

myriad of tiny droplets or ice crystals also are present. Although these

total only a small mass of water substance, they are conspicuous in the

landscape and form an important part of the environment of ecosystems at the surface.



Amounts and Significance

A typical figure for the liquid-water content of a cloud is 0.5 g m-'3.

This means only 0.5 g of water per 1300 g of air, or, in terms of specific

humidity, only 0.4 g kg-l. Compare this figure with ordinary vaporcontent values of 10 g kg-' in the middle latitudes (Table I), or more

than 20 g kg-l in the low latitudes. Yet clouds and fog are prominent

features of the environment of ecosystems.



354



XIII.



WATER I N THE LOCAL AIR



The significance of the small amounts of condensed water in clouds

or fogs stems from the finely divided state. Each tiny crystal or droplet

of the size of a raindrop) refracts and reflects light. A cloud of

droplets reduces vision through the atmosphere to almost zero, cuts

off the direct beam of the sun, and becomes a source of diffused shortwave radiation and of long-wave radiation to the ecosystems below. A

cloud deck passing across the sky causes a rapid change in the flow of

radiant energy to an ecosystem. The total input of energy changes,

usually decreasing, and its spectral composition shifts to favor both

the chemically and biologically effective blue wavelengths and far

infrared wavelengths that have chiefly a heating effect. At the same

time, the change from a direct beam to a diffused source for the

shortwave radiation brings about deeper penetration of light into the

ecosystem, benefiting photosynthesis. Water droplets and ice particles

in the local air thus have important effects on the functioning of

ecosystems, affecting their water balances.



Radiation Fog

One mode of formation of visible water particles in the local air

results from condensation in place of vapor already present in the air.

On nights of weak vertical mixing, radiation cooling is concentrated in

a shallow layer of air, and if initial humidity were high, some of the

vapor in this layer would condense as radiation fog. Radiative cooling

of the top surface of the fog now goes on more rapidly after droplets

have formed, and supports further condensation of vapor, strengthening the development of the fog.

Radiation fog has many hours in which to develop during winter

nights. It continues well into the next morning, perhaps even lasting

through the day if it is thick enough to prevent solar heating of the

ground.

Cold air that has drained into topographic depressions is often the

medium in which vapor easily condenses into fog. Inversions above

the cold air reduce upward mixing of the droplets and confine vapor,

cold air, and fog droplets with their associated pollutants within the

basins (see Plate 24). Ecosystems in such sites are subjected to

prolonged attacks by such pollutants as sulfur dioxide, often in the

form of droplets of dilute sulfuric acid. They are experiencing ”one of

the most common causes of the accumulation of pollutants to obnoxious concentrations for long periods” (Scorer, 1968, p. 31) that human

activity produces.

In periods of weak general movement in the atmosphere, such fog



V I S I B L E F O R M S OF W A T E R I N T H E L O C A L A I R



355



Plate 24. Patches of radiation fog in Morioka, northern Japan, in winter. Some

downslope movement also is taking place (December 1966).



tends to perpetuate itself because it screens the wet, cold soil from

solar heating. The author recalls a winter of frequent radiation fog at

airports on the Colorado Plateau, a region in which fog is seldom

forecast. Early winter rains saturated the soil and were followed by a

long period of stagnant air circulation, in which water repeatedly

circulated between the wet soil and the foggy local air. Closed airports

disrupted the short-hop air operations of the day (1943-1944), trapping

almost all of one airline’s planes at Albuquerque. The lifting of the fog

was forecast correctly after the high level of soil moisture was taken

into account.

The tectonic valley of the Rhine between the Vosges and the Black

Forest often fills with cold, damp air in periods of anticyclonic

stagnation. One such period in 1972 and 1973 lasted for 32 days

continuously (Weischet, 1973). The sun can barely be discerned,

visibility at ground level is low, and heavy frost forms on trees.

Meanwhile, above the fog clear skies reign. Skiers go up into the Black

Forest, only 0.5 km higher in altitude, and enjoy solar warmth, an air

temperature around 5°C higher, and a distant view of the Alps. Indeed

an abrupt transition to be caused by a thimbleful of water! The

contrast increases as pollutants continue to accumulate in the valley

fog.



356



XIII.



WATER I N THE LOCAL AIR



Plate 25. Visible-channel scanning by DMSP satellite 7529 R showing radiation fog

filling the lowlands of Puget Sound, the Willamette Valley, and the Columbia Basin on

18 December 1975. Two thousand holiday travelers were immobilized for several days

at the Seattle-Tacoma airport alone, and a like number waited at other airports to get

into Seattle. Meanwhile, AMTRAK trains arrived on schedule.



The humid conditions here can be expected to worsen still more

after new power-plant cooling towers are built and further humidify

the hapless valley’s air. The wet plumes from these gigantic evaporating devices* represent an impact on the local air that needs study in

many parts of the industrial world.

The valleys of the western US under winter anticyclones often fill

with long-lived bodies of radiation fog. Plate 25 is a satellite photograph taken at a time of anticyclonic dominance in December 1975.

The upper, visible surfaces of the fog bodies are uniformly white but

the ecosystems in the valleys-and

people, to+--live in a gray,

windless, clammy world of weak light and slowed biological activity.

The amount of liquid water that shows so clearly in these photographs from space is very small. If we take the fog bodies as being 200

m deep and having a liquid-water content of 0.5 g mP3, we compute a

* The source strength of such a system approaches



lo5 tons of



vapor per day



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