Global environment

GLOBAL ENVIRONMENT


The environment of the planet as a whole may be referred as the global environment of that planet. Thus, the global environment of our planet Earth is the environment of our planet as a whole. The structure and dynamics of the global environment of Earth is the function of various properties and interactions between following components:

(a) Solar radiation: The radiation of the primary star of the planet (generally termed the sun of that planet) received by the planet. It is the essential component of all the planets.

(b) Atmosphere : The gaseous envelop covering the planet extending from the surface of the planet to some distance into the space. It is not necessarily present over every planet and may be absent. However, the planet Earth has a well developed atmosphere over it.

(c) Hydrosphere: Theoretically, this component includes all the water present in gaseous, liquid or any other form in the global environment of the planet. This component may be absent in some planets. On the Earth, this component comprises mostly the liquid water over Earth’s surface and the liquid ground water present below the ground surface. In practice, water vapor present in atmosphere is considered as part of atmosphere and water of crystallization in rocks as part of lithosphere.

(d) Lithosphere: The rocky mantle over the non-solid core of the planet extending from the planet surface to some distance towards the center of the planet. This component is always present in every planet.

(e) Biosphere: This component includes all the living organisms present in the environment of a planet.


The structure, composition, properties and interrelations of the above components together bring about particular meteorological, climatic and weather conditions, which form the unique global environment of our planet.





SOLAR RADIATION
The Sun of planet Earth emits radiation at a temperature of about 6000 degrees K. Average radiation emitted by Sun at its surface (Sun’s radiant emittance) is 73x106 W per square meter. Spectrum of this solar radiation shows distinct emission lines indicating that it comprises of radiation of different wavelengths. The intensities and thus the magnitudes of the radiation of different wavelengths are also different. The intensity of far ultra-violet radiation is very low due to absorption of radiation by outer photosphere of the Sun. Further, extreme ultra-violet and X-ray parts of solar radiation are emitted from the chromosphere and corona regions of the Sun. These regions have temperatures as high as 1 million K. The solar radiation falling at the upper boundary of Earth’s atmosphere is termed incident solar radiation. Its average magnitude over the Earth is given by (Solar constant x πr2)/4πr2, where r = radius of Earth.


The Solar constant is the irradiance on an area at right angle to solar beam and outside the Earth. Its value is 1353 W per square meter. The average incident solar radiation at upper boundary of Earth’s atmosphere is approximately 11 x 109 J/m2 /yr. Important factors that affect this solar radiation received by Earth are:

Spherical shape of Earth: Earth is a sphere and, therefore, the angle at which incoming solar radiation strikes the upper boundary of atmosphere is not same at all points. The radiation strikes Earth at right angle in the center but the angle gradually becomes more acute towards periphery. As a result, the amount of solar radiation reflected back from the upper surface of atmosphere is zero at the center and increases gradually towards periphery. Thus, the amount of radiation penetrating atmosphere and entering Earth-atmosphere system is maximum in the center and gradually decreases towards periphery.
Orbit of Earth: The orbit of Earth around Sun is not perfectly circular but is slightly elliptic. Sun occupies one focus of this elliptic orbit. The mean distance between Sun and Earth is about 150 million kilometers. However, the orbit of Earth is elliptical and so the distance changes during different times in the year. Earth is closest (about 91.5 million miles) to Sun on about January 3, at which time it is said to be in perihelion. It is at greatest distance (about 94.5 million miles) from Sun on about July 4 when it is said to be in aphelion. These differences in distance also cause some difference in amount of solar radiation received by Earth. However, the ellipticity of orbit is not the reason of seasons on Earth. An important fact related to the Earth’s orbit around Sun is that the geometry of Earth’s orbit is not constant. The orbit at present is very slightly elliptical being nearly circular but the shape of Earth’s orbit changes cyclically from almost circular to markedly elliptical and back with a periodicity of 100,000 years. The solar constant is the mean solar irradiance on an area at upper boundary of Earth’s atmosphere perpendicular to incoming solar beam, which is about 1353 W per square meter. In the present state of Earth’s orbit being nearly circular, the difference between perihelion and aphelion is about 3.5% and the difference in solar constant at these two points is about 6.66%. In the state of most elliptic orbit, difference in solar constant at perihelion and aphelion may be as large as 30%.
Inclination of Earth’s axis of rotation: The axis of rotation of Earth is not perpendicular to the plane of ecliptic i.e. the plane in which Earth’s orbit and Sun lie. Earth’s axis of rotation makes an angle of about 66.5 degrees with the plane of ecliptic and is tilted 23.5 degrees from the line perpendicular to plane of ecliptic. The Earth’s axis although always makes an angle of 66.5 degrees with plane of ecliptic, also maintains a fixed orientation with respect to stars. The Earth’s axis continues to point to the same spot in the heavens as it makes its yearly circuit around Sun. This inclination of Earth’s rotational axis alongwith its fixed orientation throughout the whole orbit around Sun causes different seasons on Earth. Between September 23 and March 21, North Pole of Earth’s axis is tilted towards the Sun and South Pole is away from the Sun. During this period, Northern Hemisphere has summers and Southern Hemisphere has winters. In this period, daylength and, therefore solar radiation received increases towards North Pole and decreases towards South Pole. From March 21 to September 23, South Pole is tilted towards Sun and North Pole is tilted away from it, North Hemisphere has winters and South Hemisphere has summers during this period. In this period, daylength and, therefore, solar radiation received increases towards South Pole and decreases towards North Pole. Maximum tilts of North Pole and South Pole towards Sun occur on June 21 and December 22 respectively and these dates are termed summer solstice and winter solstice respectively. Midway between the dates of solstices, twice the Earth’s axis is at right angle to the line drawn from Sun to Earth and neither pole is tilted towards Sun. This condition occurs on March 21 or 22 (vernal equinox) and on September 22 or 23 (autumn equinox). Two important cyclic changes related with inclination of Earth’s axis of rotation have been noted. First is the wobbling of Earth’s axis of rotation with a periodicity of 21,000 years. This causes continuous and cyclic hemispheric variation in the solar constant. Second change is cyclic variation in the angle of inclination of Earth’s axis of rotation within the range of 21.8o and 24.4o (23.45 degrees at present) with a periodicity of about 40,000 years. Therefore, the distribution of solar irradiance at Earth’s two hemispheres varies continuously with this 40,000 years cyclic periodicity.
TRANSFORMATIONS OF SOLAR RADIATION
The general fate of the incident solar radiation in Earth-atmosphere system is as below:

(a) Absorbed in stratosphere (mainly by Ozone) = 3%

(b) Absorbed in troposphere by:

(i) Carbon dioxide = 1%

(ii) Water vapor = 12%

(iii) Dust = 2%

(iv) Water droplets in clouds = 3%

(c) Reflected from clouds = 21%

(d) Scattering back into atmosphere 6%

(e) Reflected back from Earth’s surface = 4%

(f) Received at Earth’s surface as:

(i) Direct radiation = 27%

(ii) Diffused radiation via clouds or downward scattering = 21%


Solar radiation received by Earth i.e. incident solar radiation undergoes various transformations after entering the uppermost boundary of atmosphere. The solar radiation is absorbed by atmosphere, hydrosphere, lithosphere and biosphere. Some part of solar radiation absorbed in a component provides energy for the dynamic functions of that component. The remaining part of absorbed radiation is re-emitted from the component as long-wave radiation. Two important features in the study of the transformation of solar radiation are albedo and effective radiation, which are discussed below.

Albedo
The fraction of solar radiation received by a body that is returned back from it forms the albedo of that body. In general, the radiation reflected back from clouds (21%), scattered back into atmosphere (6%) and reflected back from the Earth’s surface (4%) together constitute Earth’s albedo. The average value of albedo of Earth as a whole comes to about 33% or 0.33 (represented as fraction of unity).

The albedo of Earth as a whole has two components:

(a) Albedo of Earth’s surface: It is that fraction of radiation received at the Earth’s surface, which is returned back from the surface. Its value varies depending on the extent of snow cover, vegetation and the soil characteristics. Average albedo values for the snow range from 0.7 to 0.8 and may be as low as 0.4 to 0.5 in case of wet or dirty snow. In the deserts that have light sandy soil and are without vegetation, surface albedo values are typically 0.4 to 0.5. Albedo of damp soil is usually less than that of the corresponding dry soil. In case of damp chernozem soils the albedo values may be as low as 0.05. Albedo of natural Earth surface covered with thick vegetation cover generally ranges from 0.1 to 0.25. Areas covered with coniferous forests have lower albedo than those covered with meadows.

The height of sun in the sky determines the absorption of radiation in the water bodies, mainly the oceans. When sun is relatively high, radiation reaches water surface at high angle. A large part of the incoming radiation penetrates upper layers of water body and is absorbed. When sun is low, the radiation reaches the water surface at low angles and most of it is reflected. Thus it does not penetrate much and the albedo value of water surface increases sharply at low sun. However, in case of diffused radiation, albedo of water surface is much less variable and is about 0.1.

(b) Albedo of Earth-atmosphere system: It is more complex in nature than that of Earth’ surface. Its value is largely determined by the presence-absence, nature and thickness of clouds. In the absence of clouds in the sky, albedo of Earth-atmosphere system depends largely on the albedo of Earth’s surface. If clouds are present, a large portion of solar radiation reaching atmosphere is reflected back from the upper surface of clouds and albedo value of system increases. Albedo of clouds is usually 0.4 to 0.5. In the presence of clouds, albedo of Earth-atmosphere system is usually greater than that of Earth’s surface, except where surface is covered with relatively clean snow.

Effective radiation
It is the difference between the amount of Earth’s radiation from Earth’s surface and the amount of long-wave counter-radiation from atmosphere. Most important factor connected with effective radiation is ‘green-house effect’ due to presence of atmosphere. The atmosphere has various gases viz. carbon dioxide, methane, water vapor etc. which selectively absorb long-wave radiation. Due to this Earth’s atmosphere is relatively more transparent to short-wave radiation than to long-wave radiation. Since long-wave radiation from the Earth’s surface is trapped in the atmosphere, average effective radiation from Earth’s surface as a whole is much lower than the short-wave radiation absorbed at the surface.

The effective radiation of Earth’s surface largely depends on the temperature at Earth’s surface, atmospheric humidity and clouds. Experimental data has shown that radiation of Earth’s natural surfaces is generally quite close to the radiation of black body at corresponding temperatures. Further, a significant part of long-wave radiation lost from Earth’s surface is compensated by long-wave counter-radiation from the atmosphere. This counter-radiation mainly depends on the amount of atmospheric water vapor i.e. air humidity and clouds and so these factors affect the effective radiation of Earth’s surface.

RADIATION BALANCE

In every component of the global environment, radiation is constantly coming and a portion of incoming radiation is constantly going out of it. The algebraic sum of radiation fluxes reaching and leaving a place is called its radiation balance. In the study of global environment, following three types of radiation balances are important:


(a) Radiation balance of Earth’s surface (R)

It is the algebraic sum of radiation fluxes reaching and leaving the surface of Earth. Its magnitude is equal to the difference between the amounts of direct and diffused short-wave radiation absorbed by Earth’s surface and the long-wave effective radiation and is given as:
R = Q(1 - A) - I
where, R = Radiation balance of Earth’s surface; Q = Total direct and diffused short-wave radiation reaching Earth’s surface; A = Albedo of Earth’s surface given as a fraction of unity; I = Effective radiation.

Due to ‘green-house effect’ of atmosphere, the average radiation balance of Earth’s surface is always positive. The radiation balance of Earth’s surface is linked to the radiation balance of atmosphere.


(b) Radiation balance of Earth-atmosphere system (Rs)


It is the algebraic sum of radiation fluxes reaching and leaving Earth-atmosphere system as a whole. This it is equal to the difference between radiation solar radiation reaching the upper boundary of atmosphere i.e. incoming radiation (incident solar radiation) and long-wave radiation leaving atmosphere’s upper boundary i.e. outgoing radiation. Its magnitude may be given as:

Rs = Qs(1 - As) - Is

where, Rs = Radiation balance of a vertical column extending from upper boundary of atmosphere to the Earth’s surface; Qs = incoming radiation; As = Albedo of Earth-atmosphere system and Is = outgoing radiation.

The outgoing radiation of Earth-atmosphere system includes that part of Earth’s surface radiation which passes through atmosphere unaltered and goes out of atmosphere’s upper boundary plus radiation of atmosphere itself going out of its upper boundary. The outgoing radiation is much influenced by clouds. In the absence of clouds, Earth’s surface radiation within wavelength range of 900-1200 nm plays important role. In completely cloudy conditions, the radiation from the upper surface of clouds becomes very important and this radiation depends on the temperature of that surface. The temperature at the upper surface of clouds is usually much lower than the temperature at the Earth’s surface. Therefore, clouds substantially reduce the amount of long-wave radiation into the outer space.


(c) Radiation balance of atmosphere (Ra)


It is equal to the difference between radiation balance Earth-atmosphere system and the radiation balance of Earth’s surface i.e. given as:


Ra = Rs - R

or, substituting the expressions for Rs and R, as:

Ra = Qs(1 - As) - Q(1 - A) - (Is - I)

The magnitude of the radiation balance of atmosphere is negative and equal to the absolute value of radiation balance at Earth’s surface. This negative radiation balance of atmosphere is compensated by:

(i) inflows of energy from condensation of water vapor during cloud formation and precipitation and

(ii) flow of heat from Earth’s surface associated with turbulent heat conductivity of lower atmospheric layer.

Geographical distribution of radiation balance
The radiation balance at Earth’s surface is not uniform but varies at different geographical locations. Following two factors have important effect on the variation of radiation balance over Earth’s surface.

1. Relationship between latitude and irradiance: The radiation reaching certain surface area (B) at Earth’s surface depends on the latitude and is given by Lambert’s cosine law:


QB = Qo cos z

where, Qo = Irradiance of solar beam at upper boundary of atmosphere i.e. equal to solar constant; z = solar zenith angle.

Zenith angle of direct solar beam may be defined under any condition of varying latitude, solar declination and solar time is given by:

cos z = sin φ sin δ+ cos δ cos h

where, φ= latitude; δ = angle of solar declination (angle between solar beam and equator which varies between -23.45 degrees on 22 December and +23.45 degrees on 22 June); h = hour angle of Sun (measure of time from solar noon where one hour equals to 15 degrees).

Maximum monthly range of solar irradiance shows considerable increase with latitudes and is also related to variations in photoperiod (time in hours between sunrise and sunset P). Value of h at both sunrise and sunset may be derived from solar declination and latitude of site by:

cos h = -(tan φ tan δ)

so that the photoperiod (P) is gives as:

P = 2/15 cos-1 h

2. Optical effects of atmosphere on solar irradiance: The ideal pattern of solar irradiance at the upper boundary of atmosphere is not realized perfectly at Earth’s surface because of the optical effects of atmosphere on solar beam passing through it. Following three features of the interaction between solar irradiance and the atmosphere determine solar irradiance at the Earth’s surface.

(i) Path length (m): Assuming that the thickness of atmosphere over Earth’s surface is uniform, the length of path, which solar beam traverses from upper boundary of atmosphere to the Earth’s surface i.e. the path length (m) is given as:


m = 1/cos z

The path length (m) is shortest at the place where solar beam and Earth’s surface are perpendicular to each other. The above relation is found to be correct upto zenith angle (z) of about 70 degrees (Robinson, 1966). At greater zenith angles, curvature of Earth and atmospheric refraction cause increasing overestimation of the value of path length.

(ii) Atmospheric transmittance: The radiation of solar beam passing through atmosphere may be absorbed, reflected and scattered by various gases and aerosols in the atmosphere. The effects of these phenomena on the solar irradiance are reflected in mean atmospheric transmittance (τ). If all the effects of atmosphere on the solar beam are assumed to be constant throughout the depth of atmosphere, then depletion of the irradiance of solar beam at upper boundary of Earth’s atmosphere (Qo) will be simple function of path length (also termed air mass) and mean atmospheric transmittance. Then the solar irradiance reaching Earth’s surface (Q) will be given as:

Q = Qo τm
The value of τ strongly depends on dust and pollutants present in the atmosphere and may vary from 0.4 in polluted atmosphere to 0.8 in very clear and dry atmosphere.

(iii) Cloud cover: The solar beam passing through Earth’s atmosphere is also affected by the cloud cover present in it. Various empirical relationships have been derived which relate solar irradiance reaching Earth’s surface to some measure of cloudiness. No such relation is perfect because of the variations in optical properties of different cloud types. A general relationship derived from global observations at 88 well separated meteorological stations is given as:


Q = Qo (0.803 - 0.34f - 0.458f2)

Where, f = monthly fractional cloud cover.


The variations in mean atmospheric transmittance have not been included in this relationship and the above relationship will show changes with such variations.

The salient features of the geographical distribution of radiation balance of Earth’s surface keeping in view above effects on solar irradiance at upper boundary of Earth’s atmosphere are as follows:

1. Yearly total radiation: It varies substantially within the range of less than 60 to more than 220 kcal/sq. cm/year.

(i) At high and middle latitudes, distribution of total radiation is zonal in nature while in tropical latitudes the distribution deviates from zonality.

(ii) Total radiation is greatest in high-pressure belts in Northern and Southern Hemispheres, especially in desert areas of continents. Highest total radiation is found in Northwest Africa. It is due to almost total absence of clouds in that region. However, total radiation is reduced in low latitudes near equator, in regions of monsoon climates and in certain other regions due to increased cloudiness.

(iii) Total radiation also varies seasonally i.e. in different months of the year. This may be illustrated by considering its values in June and December i.e. the months in which average height of Sun is highest and the lowest in Northern Hemisphere and vice-versa in Southern Hemisphere.

During December, zero isocline passes somewhat to the north of Arctic Circle. At latitudes above it, Sun never rises above the horizon and total radiation is zero. To the south of zero isocline, total radiation increases rapidly. Its distribution in region below Arctic Circle and above tropical latitudes is largely zonal. In low tropical latitudes, the zonal character of distribution is absent and total radiation is determined by the degree of cloudiness in different regions. Further, average zonal changes in total radiation are relatively small from low latitudes in Northern Hemisphere throughout entire Southern Hemisphere. Continuous increase in length of day towards South Pole in Southern compensates for reduction in average height of Sun and so there is negligible reduction in total radiation towards higher latitudes in Southern Hemisphere.

During June, similar situation exists with regard to distribution of total radiation. In Northern Hemisphere, total radiation shows relatively little change except for desert regions where its value is high. At middle and high latitudes of Southern Hemisphere, total radiation decreases with increasing latitudes.

2. Yearly total radiation balance: Yearly total radiation-balance at Earth’s surface is positive over entire surface of oceans and land. Negative yearly sums of radiation balance are found only in regions of permanent snow or ice cover. The yearly radiation balance shows sharp changes from land to ocean areas. As albedo values of ocean surface are lower, radiation balance of ocean areas is usually higher than that of land areas at same latitudes.

On ocean surface, distribution of radiation balance is generally zonal in character. However, some regions particularly where warm and cold currents operate, some deviations from zonality are found. At tropical latitudes, radiation balance of ocean surface shows small changes while in the middle latitudes, there is rapid reduction in corresponding balances from lower to higher latitudes. Greatest value of radiation balance of Earth’s surface is 140 kcal/sq. cm/year, which occurs in Arabian Sea.

On the land surface, changes in yearly radiation-balance values are also partly zonal in nature. However, in certain regions deviation from zonality is found due to difference in their moisture conditions. In the dry regions, radiation balance is lower in comparison with regions of sufficient or excessive moisture at same latitudes. The low radiation balance in dry regions is due to:

(i) Reflection of short-wave radiation

(ii) Higher expenditure of radiation energy on effective radiation owing to high surface temperature, low cloudiness and relatively low air humidity. Thus alongwith general reduction in radiation balance at higher latitudes, there are also found regions of further reduced radiation balance in areas of dry climate. This reduction is particularly observed in Sahara, deserts of Central Asia and in many other deserts and arid regions. In monsoon regions, yearly radiation balance of Earth’s surface is also somewhat reduced due to intense cloudiness during warm season. In humid tropical regions, highest yearly radiation balance values are found on land but even these just reach 100 kcal/sq. cm/year which is quite less than corresponding maximum value for the ocean.

ENERGY BALANCE (HEAT BALANCE)

When solar radiation enters Earth’s environment, it provides energy for maintenance and dynamic functions of different components of global environment. The continuous maintenance of particular physical and chemical states of matter in atmosphere, hydrosphere, lithosphere and biosphere requires energy provided by solar radiation. Further, various dynamic changes in these states such as air and water movements, changes in the state of water from vapor to liquid to solid and vice-versa and the activities of living organisms are found to occur. These changes are possible only through the expenditure of energy provided by solar radiation. The energy of Earth’s surface radiation balance is expended on heating of atmosphere through turbulent heat conductivity, on evaporation of water, on heat exchange with deeper layers of hydrosphere and lithosphere etc. and photosynthesis in biosphere. In general, the quantitative characteristics of all forms of transformations of solar energy on the Earth’s surface are represented in the equation of global energy (heat) balance. This equation includes the algebraic sum of flows of energy reaching and leaving the Earth’s surface. This sum is always zero according to the law of conservation of energy. The energy balance and radiation-balance at Earth’s surface are linked together.

The equations representing energy balance may be compiled for various volumes and surfaces of atmosphere, hydrosphere and lithosphere. However, in the studies of global environment, equations are often employed for an imaginary column whose upper end is at the upper boundary of atmosphere and which passes through atmosphere deep below Earth’s surface. Three equations of energy (heat) balance describing global energy balance are:

(a) Energy balance equation of Earth’s surface

(b) Energy balance equation of Earth-atmosphere system

(c) Energy balance equation of atmosphere.


(a) Energy balance equation of Earth’ surface


Major elements of this equation are :


(i) Radiation balance (R) i.e. radiation flux, which is considered positive in value when it describes inflow of energy (heat) from above to underlying Earth’s surface.

(ii) Turbulent energy (heat) flow (P) from underlying Earth’s surface to atmosphere.

(iii) Underground energy (heat) flow (A) from Earth’s surface to deeper layers of hydrosphere or lithosphere.

(iv) Energy (heat) expenditure on evaporation (or release of heat in condensation) (LE) where L is latent heat of vaporization and E is rate of evaporation.


With the above elements, energy balance equation of Earth’s surface is given as:

R = LE + P + A
The elements of energy balance not included in the above equation are:

1. Energy expenditure on melting of ice or snow on surface (or inflow of heat from freezing of water)

2. Energy expenditure associated with friction of air currents, ocean waves produced by winds and ocean tides

3. Energy (heat) flows transferred by precipitation whenever their temperature is not equal to that of underlying surface

4. Energy expenditure on photosynthesis

5. Energy (heat) inflow from oxidation of biomass.

With the addition of these elements also, comprehensive energy balance equation of Earth’s surface may be obtained.

The magnitude of underground energy (heat) flow (A) may be obtained from the energy balance equation of a vertical column whose upper base is at Earth’s surface and lower base at the depth below ground surface where heat flow is negligible (Fig. 2). Since heat flow from depths of Earth’s crust is negligible, vertical flow of heat at the lower base of column may be assumed to be zero. The equation for A is given as:

A = Fo + B

where, B represents the changes in heat content inside the column over a given period of time and Fo is the inflow of heat produced by horizontal heat exchange between the column being considered and the surrounding space of hydrosphere or lithosphere. Fo is equal to the difference between amounts of heat entering and leaving through vertical walls of column.

In lithosphere, Fo usually becomes negligible due to low heat conductivity of soil. Thus for land A = B and since over a period of whole year, upper layers of soil are neither heated nor cooled, A = B = 0.

Fo becomes large in case of water bodies having currents with a large horizontal heat conductivity determined by macroturbulence. In case of closed water bodies taken as a whole whose depth and area are large, values of A and B are close. It is because heat exchange between such bodies of water and the ground are usually negligible. However, in specific sectors of oceans, seas and lakes, magnitudes of A and B may be substantially different. The average yearly value of heat exchange of an active surface with lower is not zero but is equal to the quantity of heat received or lost due to currents and macroturbulence i.e. A = Fo.

Thus for average yearly period, energy balance equation of Earth’s surface will be:

(i) For land: R = LE + P

(ii) For ocean: R = LE + P + Fo

(iii) For deserts (where evaporation is almost zero): R = P

(iv) For global oceans as a whole (where redistribution of heat by currents is compensated and is zero): R = LE + P


(b) Energy balance equation of Earth-atmosphere system


This equation can be derived by considering the inflow and expenditure of energy in a vertical column passing downwards from the top of atmosphere to that level in hydrosphere or lithosphere at which noticeable daily or seasonal fluctuations of temperature stop (Figure-1). Energy (heat) flow through the lower base of this column is practically zero.

Energy balance equation of Earth-atmosphere system may given as:

Rs = Fs + L(E - r) + Bs

All the terms on the right-hand side of equation are assumed positive in value when they describe expenditure of energy (heat). The elements of the equations are as discussed below:

(i) Radiation balance of Earth-atmosphere system (Rs): It describes the energy (heat) exchange between the vertical column under consideration and the outer space and is equal to the difference between the amounts of total solar radiation absorbed by the entire column and the total long-wave radiation from column to outer space. It is considered positive when it describes inflow of energy (heat) into the Earth-atmosphere system.


(ii) Total horizontal heat transfer (Fs): It occurs through the sides of the column under consideration and is given as:


Fs = Fo + Fa

where, Fo = horizontal heat transfer through sides of the column in the atmosphere and Fa = horizontal heat transfer through the sides of column in the hydrosphere or lithosphere. Value of Fa is similar to that of Fo and describes the difference of inflow and expenditure of heat in the column of air resulting from atmospheric advection and macroturbulence.

(iii) Heat transfers in change of the state of water: Heat balance of column is also influenced by sources of heat (both positive and negative) that are located within the column itself. These include the inflow and expenditure of heat due to changes in state of water, especially by evaporation and condensation.

Over sufficiently homogeneous surfaces during long periods, the average difference in the magnitudes of condensation and evaporation of water drops in atmosphere is equal to the sum of precipitation (r) and the inflow of heat is equal to Lr. Corresponding component in the energy balance represents the difference between heat inflow from condensation and its expenditure in the evaporation of drops. It may differ from Lr in conditions of rugged surfaces and also in individual short periods of time.

The difference between heat expenditure on evaporation the surface of water bodies, soils and vegetation and heat inflows from condensation on these surfaces are equal to LE.

The overall influence of condensation and evaporation on the column’s energy balance may be approximated in terms L(r -E).

(iv) Changes in the heat content within the column: This change over the period being referred is represented by component Bs in the energy balance equation.


Remaining components of the balance such as heat inflow from dissipation of mechanical energy, difference between heat expenditure and inflow on melting and formation of ice, difference between heat expenditure on photosynthesis inflow from oxidation of biomass etc. are very small and may be neglected.

Consideration of different components of energy balance equation under different conditions shows that:

(i) For an average yearly period, magnitude of Bs is apparently close to zero and the equation simplifies to:


Rs = Fs + L(E - r)

(ii) For the land conditions, the equation becomes:


Rs = Fa + L(E - r)

(iii) For the entire globe, E = r over a period of one year and horizontal inflow of heat into the atmosphere and hydrosphere is apparently zero. Thus the energy balance equation of Earth-atmosphere system for the Earth as a whole simplifies to:

Rs = 0
(c) Energy balance equation of atmosphere


This equation may be obtained by either

(i) Summing up the corresponding flows of heat or

(ii) As difference between members in the heat balance equation for the Earth-atmosphere system and in that for Earth’s surface.

Assuming that atmospheric radiation balance is given by:

Ra = Rs - R

and changes in the heat content of atmosphere (Ba) are given by:

Ba = Bs - B

it can be seen that:

Ra = Fa - Lr - P + Ba

and for an average yearly period, equation is:

Ra = Fa - Lr - P


DISTRIBUTION OF ENERGY BALANCE COMPONENTS
Distribution of energy balance components of Earth’s surface
Important components of energy balance of Earth’s surface which show geographical differences in their values are heat expenditure on evaporation, turbulent heat exchange and redistribution of heat through atmospheric and oceanic currents.

1. Heat expenditure on evaporation: The magnitudes of evaporation from land surface and the oceans in the vicinity of coastlines, differ significantly. This may apparently be explained

(i) differences in the value of possible evaporation on land and on ocean and

(ii) the influence of insufficient moisture in many land areas which limits the intensity evaporation processes and of heat expenditure on evaporation.


At extratropical latitudes, absolute value of heat expenditure on evaporation generally decreases with increasing latitudes. However, major non-zonal changes on land and ocean alter this pattern. In tropical latitudes, distribution of heat expenditure on evaporation is quite complex. Compared to high-pressure regions, its value declines somewhat in the ocean regions adjoining the Equator.

In the oceans, maximum mean latitudinal heat expenditure on evaporation occurs within high-pressure belts. At 50-70 degrees where radiation balances of land and oceans are approximately same, the heat expenditure on evaporation is substantially larger for oceans. This is evidently due to large expenditure of heat brought by ocean currents. In oceans, distribution of warm and cold currents is principal cause of the non-zonal changes in heat expenditure on evaporation. All the major warm currents increase heat expenditure substantially while cold currents reduce it. This may be clearly seen in regions influenced by warm currents like Gulf stream and Kuroshio by old currents like those of Canary Islands, Bengal, California, Peru and Labrador. The yearly evaporation from ocean surface at a particular latitude may change by several time depending on the increase or decrease in water temperature brought about by the currents. In addition, non-zonal in the values of heat expenditure on evaporation and so of evaporation from oceans are also influenced by conditions of atmospheric circulation determining wind velocity and the annual humidity deficit over the oceans. The ocean surfaces have somewhat higher radiation balance than land surfaces and evaporating surfaces may additionally receive a large quantity of heat energy through redistribution of heat by ocean currents. Therefore, evaporation from ocean surface in tropical areas corresponds to a layer of water more than two meters thick.

On the land, mean latitudinal value of heat expenditure on evaporation is maximum at equator. These values change within the subtropical high-pressure belts. In both hemispheres, a certain increase in evaporation occurs with increase in latitudes though the increase is more pronounced in Northern Hemisphere. This increase is due to increased precipitation as compared with arid zones at lower latitudes. The distribution of heat expenditure on evaporation from land surface deviates from zonal pattern even more than from oceans. This is due to very great influence of climatic moisture conditions on evaporation. In regions of sufficient soil moisture found at high latitudes and in humid regions at middle and tropical altitudes, heat expenditure on evaporation and the evaporation are governed largely by balance. In regions of insufficient moisture, evaporation is reduced due to insufficient soil moisture while in desert and semi-desert areas, evaporation is almost equal to low yearly total precipitation. Highest heat expenditure on evaporation occurs in certain equatorial regions where in case of abundant moisture and large inflows of heat, it exceeds 60 kcal/sq. cm/year. This corresponds to yearly evaporation of layer of water more than one meter thick.

Further, the patterns of seasonal heat expenditure on evaporation in extratropical latitudes are different on land and oceans. On the land, this expenditure and evaporation decreases substantially during cold season and depending on moisture conditions, attains a maximum at the beginning or in middle of warm season. In contrast, evaporation from oceans usually increases in cold season due to greater difference in temperature of water and air at that time which increases difference in concentration of water vapor on the surface of water and in air. In addition, in many oceanic regions average wind velocities are greater in cold seasons and this also increases evaporation.

2. Turbulent heat exchange: The value of turbulent heat exchange is positive heat is released by Earth’s surface into air and is negative when heat is received by Earth’s surface from atmosphere during the year. Over a year, all the land surfaces except Antarctica and larger part of ocean surfaces release heat into the atmosphere.


In oceans, turbulent heat exchange gradually increases towards higher latitudes. Its magnitude is not large for greater part of ocean surfaces and usually does not exceed 10-20% of the magnitudes of principal components of energy balance equation. Large absolute values of turbulent heat flow, exceeding 30-40 kcal/cm2/year, occur in regions of powerful warm currents e.g. Gulf Stream. Here water is on average warmer than air and at higher latitudes where sea is still free from ice. Cold oceanic currents reduce temperature of water, reduce turbulent heat flow from ocean surface to the atmosphere and increase it in reverse direction.

On land, turbulent heat flow decreases towards higher latitudes. Its maximum value occurs within high-pressure belts which declines somewhat near Equator and sharply decreases at high latitudes. Magnitude of turbulent heat -exchange on continents is greatly influenced by climatic moisture conditions. In arid regions, turbulent heat flow from land surface into the atmosphere is much higher than in humid regions. Highest expenditure of turbulent heat flows on land is found in tropical deserts where it may exceed 60 kcal/sq. cm/year. In humid regions, especially in regions at middle latitudes, heat expenditure through turbulent flows is usually much lower.

The very different patterns of change in turbulent heat exchange on land and in oceans reflect differences in the mechanisms of air mass transformation on the surfaces of continents and oceans.

3. Heat redistribution through water currents: In the heat balance of oceans, inflow or expenditure of energy owing to horizontal exchanges primarily through oceanic currents is very important. A large quantity of heat is redistributed in oceans between tropical and extratropical latitudes. Both warm and cold currents play important role in redistribution of heat in oceans. Regions of increased positive values of that particular component of heat balance (reflecting outflow of heat from ocean surface to lower layers) correspond with regions of cold currents and the regions of reduced negative values correspond with warm currents. Such correspondence is observed for major warm currents e.g. Gulf Stream, Kuroshio and Southwest Pacific Stream as well as for cold currents e.g. Canary Islands, Bengal, California and Peru. Ocean currents carry away heat mainly from a zone ranging from 20 degrees N latitude to 20 degrees S latitude. Maximum of heat absorbed is slightly shifted to the north of Equator. Further, the heat is carried to higher latitudes and expended in the region of 50 degrees to 70 degrees N latitude where warm currents are especially strong.

Studies of Strokina (1963,1969) concerning changes in heat content of ocean’s upper layers over a year have shown that these changes may attain significant magnitudes which are quite comparable with changes in magnitudes of the main components of heat balance. Greatest yearly changes in heat content of ocean’s upper layers (over 25 kcal/sq. cm/year) are observed in Northwestern regions of Pacific ocean and adjoining areas.


Distribution of energy balance components of Earth-atmosphere system

Data for average yearly conditions show that relative proportions of various components of energy balance of Earth-atmosphere change perceptibly at various latitudes.

In equatorial zone, the large inflow of radiation energy is further increased by addition of a substantial inflow of heat produced by changes in state of water through condensation and evaporation. These sources of heat produce large expenditure of heat on atmospheric and oceanic advection. A relatively narrow zone adjoining Equator is and extremely important source of energy for these advection conditions.

At higher latitudes upto 30-40 degrees, a positive radiation balance that decreases with increasing latitude is accompanied by substantial expenditures of energy on water exchange. In most parts of that zone, energy of radiation balance is almost equal to heat expended on water exchange and very little heat is redistributed through air and water currents.

At latitudes above 40 degrees, a zone of negative radiation balance is found. Its absolute value increasing at higher latitudes. The negative radiation balance of that zone is compensated by inflow of heat brought by air and water currents. Proportions of those components within that balance which compensate for the deficiency of radiation energy vary at different latitudes. For the belt between 40--60 degrees, excess energy released in condensation of water is major source of heat while inflow of heat redistributed by ocean currents is also important. At higher latitudes, especially in polar regions, heat inflow from condensation is very small and influence of ocean currents is either absent (in South polar zone) or is weak due to permanent ice cover (in North polar zone). At these latitudes, redistribution of heat through atmospheric circulation is major source of heat.

The average values of various components of energy balance of Earth-atmosphere system over six-month periods at various latitudes have been studied (Table-2). These show that magnitude of radiation absorbed by Earth-atmosphere system (Qa) is not the only factor determining the magnitude of outgoing long-wave radiation at the top of atmosphere (Is). For middle and high latitudes during October to March in Northern Hemisphere and for high latitudes in Southern Hemisphere throughout the entire year, the main source of heat is heat-transfer from lower latitudes through atmospheric circulation.

Distribution of energy balance components of atmosphere
The average radiation balance of atmosphere at various latitudes changes less than other components of heat balance. The large absolute negative values for the atmospheric radiation-balance observed at all latitudes are compensated largely by inflows of heat from condensation. The role of heat from Earth’s surface through turbulent heat exchange is less important though the influence is quite perceptible.

Distribution of components of energy balance of whole Earth
Depending on the relative proportions of land and ocean areas in particular zones, mean latitudinal distribution of the components of energy balance of Earth as a whole is characterized by patterns typical of continents or by the patterns typical of oceans. Average values of energy balance components for individual continents and oceans (Table-3) show that in three continents (Europe, North America and South America) greater share of energy radiation balance is expended on evaporation. In the remaining three continents (Asia, Africa and Australia) where dry climates prevail, opposite is true.


Energy balance components of three oceans show little difference from each other. For each ocean the sum of heat expenditure on evaporation and turbulent heat exchange is close to the magnitude of radiation balance. This means that the heat exchange among different oceans resulting from currents does not exert any substantial influence on the heat balances of individual oceans.

The values of the components of energy balance for Earth a whole show that in oceans approximately 90% of the energy of radiation balance is expended on evaporation and only 10% on direct turbulent heating of atmosphere. These magnitudes are nearly same on land. For Earth as a whole, 83% of the energy of radiation balance is expended on evaporation 17% on turbulent heat exchange.

The values of the components of energy balance for the Earth as a whole are shown in FigURE-2. Overall yearly flux of solar radiation entering outer boundary of troposphere is approximately 1000 kcal/sq. cm. Due to the spherical shape of Earth, about 25% of this yearly flux (i.e. 250 kcal/sq.cm), passes through a unit surface of the upper boundary of troposphere. Assuming that Earth’s albedo (As) is 0.33, short-wave radiation absorbed by Earth represented by Qs(1-As) is approximately 167 kcal/sq. cm/year. Out of this, short-wave radiation reaching Earth’s surface is 126 kcal/sq. cm/year. Average value of albedo at Earth’s surface (A) is 0.14. This takes into account the differences in value of incoming solar radiation in various regions. Thus the amount of short-wave radiation absorbed at Earth’s surface, represented by Q(1-A), is 108 kcal/sq. cm/year and 18 kcal/sq. cm/year is reflected back from the surface. The atmosphere absorbs about 59 kcal/sq. cm/year which is substantially less than that absorbed at Earth’s surface. Since radiation balance of Earth’s surface (R) is 72 kcal/sq. cm/year, average effective radiation of Earth’s surface (I) comes to be 360 kcal/sq. cm/year. Overall value of Earth’s long-wave radiation (Is) is quite close to 167 kcal/sq. cm/year. The ratio I/Is much less than the ratio Q(1-A)/Qs(1-As). This difference shows that greenhouse effect has very large influence on the thermal processes of Earth. Due to this effect, Earth receives about 72 kcal/sq.cm/year of radiation energy. This energy is partly expended on evaporation of water (LE = 60 kcal/sq. cm/year) and partly returned to the atmosphere by turbulent heat losses (P = 12 kcal/sq. cm/year). Thus, the energy balance of atmosphere has following components:

(i) Heat inflow from absorbed short-wave radiation = 59 kcal/sq. cm/year

(ii) Heat inflow from condensation of water vapor (Lr) = 60 kcal/sq. cm/year

(iii) Heat inflow from turbulent heat losses at the Earth’s surface = 12 kcal/sq. cm/year

(iv) Heat expenditure on effective radiation into outer space (Is - I) = 131 kcal/sq. cm/year.


The last figure corresponds to the sum of first three components of energy balance.


SOLAR RADIATION AND PLANETARY TEMPERATURE

The temperature of a planet irradiated by solar radiation can be estimated by balancing the amount of radiation absorbed (Ra) against the amount of outgoing radiation (Ro). The Ra will be the product of:

(i) Solar irradiance (I)

(ii) Area of planet. Area of planet relevant for such calculations is the area of the planet as seen by incoming radiation which is given by πr2 where r = radius of the planet.

(iii) Absorbed fraction of radiation. The fraction of radiation that is absorbed is given by (1 - A where A = albedo of planet. This albedo represents the fraction of radiation that is reflected back from the planet.

Thus the energy absorbed by the planet will be:

Ra = I πρ2 (1 - A)

Intensity of outgoing radiation of a body is given by Stefan-Boltzmann Law i.e.
Io = σ T4
where σ = Stefan-Boltzmann Constant = 5.6 x 10-8 Wm-2 K-4. The total energy radiated by the planet will be the product of the intensity of outgoing radiation (Io) and the area of the whole planetary surface giving out radiation (4 π r 2). Thus the outgoing radiation (Ro) from the planet is given by:

Ro = 4 πr2 σ T4

Effective planetary temperature

Since Ra = Ro i.e. system is assumed to be in steady-state where radiation absorbed and outgoing radiation are equal, an expression for the effective planetary temperature (Te)can be obtained from the above equation and it may be given by:

Te = [I - (1 - A)/4 σ]0.25
In this expression of effective planetary temperature, effect of atmosphere has not been taken into account. For Earth, solar irradiance (I) at the top of atmosphere is about 1.4 x 103/m2/s and albedo of Earth as a whole is about 0.33. From these values, the calculated equilibrium temperature of Earth comes to be 254 K. However, the actual observed average ground level temperature of Earth is about 288 K. This higher effective temperature of Earth from the calculated value is due to the greenhouse-effect of atmosphere.

The black-body spectrum of Earth at 288 K shows that radiation from Earth is of much longer wavelength and is at much lower intensity than radiation from Sun. The absorption spectrum of Earth’s atmosphere overlaps fairly well with the solar emission spectrum. Except for a very narrow window in the absorption bands, much of the long-wave radiations from Earth correspond with the region of absorption in the atmosphere. This means that much of the incoming radiation reaches the Earth’s surface while the outgoing thermal radiation is largely absorbed by the atmosphere rather than being lost to space. Thus, the effect of atmosphere is to trap the outgoing thermal radiation. and this effect is termed green-house effect.. The thermal radiation i.e. the heat trapped by the atmosphere due to green-house effect is responsible for the effective temperature of Earth being higher than the temperature calculated without taking into account the effect of atmosphere. In general, absorption of re-emitted long-wave radiation and vertical mixing processes determine the temperature profile of the lower part of atmosphere (troposphere) which in turn determine the Earth’s temperature.

Optical depth of atmosphere and Earth’s surface temperature
The atmosphere is not transparent to the outgoing long-wave radiation and much of this radiation is absorbed in the lower part of the atmosphere, which as a result, is warmer than the upper parts. Simple radiative equilibrium models have been developed for Earth and to account for this effect, these models divide the atmosphere into layers that are just thick enough to absorb the outgoing radiation. These atmospheric layers are said to be optically thick and the atmosphere is discussed in terms of its optical depth based on the number of these atmospheric layers of different optical thickness. Earth’s atmosphere is sometimes said to have two layers while that of planet Venus has almost 70 layers which are largely due to enormous amount of CO2 in the atmosphere of Venus. The radiation equilibrium model indicates that the effective planetary temperature (Te) is thus related to ground-level planetary temperature (Tg) by the equation:

Tg4 = (1 - τ)Te4 (where τ = optical depth of atmosphere)

The optical depth of atmosphere increases with increase in atmospheric concentrations of carbon dioxide and water vapor because both these are principal atmospheric absorbers of outgoing long-wave radiation. With increasing concentrations of CO2 in lower layers of atmosphere, other such gases that are responsible for radiating heat to outer space are pushed to slightly higher and colder levels of atmosphere. The radiating gases will radiate heat less efficiently because they are colder at higher altitudes. Thus, the atmosphere becomes less efficient radiator of heat and this results in rise of atmospheric temperature. This rise in atmospheric temperature, in turn, leads to more evaporation and increase in atmospheric water vapor, which is a greenhouse gas and further increases the absorption of outgoing long-wave thermal radiation. This positive feedback results in further increase in atmospheric temperature. The model also suggests that increase in atmospheric CO2 is associated with decrease in temperatures of upper (stratospheric).

Vertical heat transport and Earth’s surface temperature
Simple models of radiation balance of atmosphere do not take into account various other processes that transport heat vertically in the atmosphere and, therefore, overestimate the surface temperature of Earth. Convection is major process of vertical heat transport and is very important in lowering the surface temperature. Convection occurs because warm air is lighter than cool air and so rises upwards carrying heat from Earth’s surface to the upper atmosphere. As warm air rises up, it expands due to fall in pressure and work done in expansion causes it to cool adiabatically. Thus,

Cv ΔT = - P ΔV (where Cv = molar heat capacity at constant volume)

Ideal gas equation PV = RT takes the differential form P dV + V dP = R dt which may be rearranged in incremental form as:

- P ΔV + R ΔΤ = V ΔP
This equation may be combined with equation Cv ΔT = - P ΔV using the fact that Cp - Cv = R, where Cp = molar heat capacity at constant pressure = 29.05 J/mol/K. This results in following equation:

Cp ΔT = (Cv + R) ΔT = - P ΔV + R ΔT = V ΔP = (RT/P) ΔP .............(a)

It can be shown that P/P = - Mmg Δz /RT where Mm = mean molecular weight of air = 0.028966 kg/mol; g = acceleration due to gravity = 9.8065 m/s/s; z = altitude. This gives:

RT/P = - Mmg Δz/ ΔP.......................................................(b)

Substitution of the above equation (b) in equation (a) gives:

Cp ΔT = - Mmg Δz

or, ΔT/ Δz = - (Mmg/Cp)

For Earth’s atmosphere, the lapse rate ( ΔT/ Δz) works out to be -9.8 K/km for dry air. However, the air is usually wet and as it rises up, it releases latent heat so the measured lapse rate is -6.5 K/km.

If atmospheric temperature falls much less slowly with height than the lapse rate (or even rises with height) then inversion conditions exist and air is very stable with respect to vertical convective mixing. Conversely, if temperature falls very rapidly with height, at a rate greater than lapse rate, then the atmosphere is unstable and convective mixing will be active.

Short-wave radiation and temperature

The discussion till now has assumed total transparency of atmosphere to incoming solar radiation. Though it is true for visible range of radiation, it is not true for ultra-violet region of the solar spectrum. Though the amount of such short-wave radiation is very small, it has important consequences for the temperature of Earth-atmosphere system.

Various ultra-violet wavelengths are absorbed in the atmosphere at different heights. AT just over 40 km, absorption of ultra-violet radiation by ozone results in considerable warming of stratosphere and in this zone, temperature rises with altitude. Average temperature of stratosphere is 250 K. Considering it to be a black-body radiator, maximum power radiation would be expected at 11.5 μm. This value is very close to absorption band of carbon dioxide which means that this gas also plays important role in stratospheric temperature. Increase in concentration of carbon dioxide in stratosphere might allow more effective radiation from stratosphere and, therefore, its cooling. This effect is quite opposite to that noted for troposphere.

Further, at the altitude of thermosphere, atmosphere is very thin. In this zone, molecules are exposed to unattenuated solar radiation of extremely short wavelength i.e. of high energy. This radiation arises from the outer region of Sun. At wavelengths below 50 nm, effective emission temperature exceeds 10,000 K. High-energy solar protons of such wavelengths are absorbed by gas molecules giving them high transitional energies i.e. high temperatures. The energies may be large enough to dissociate oxygen and nitrogen. Temperatures in thermosphere undergo wide variations depending upon the state of Sun. During solar disturbances, very much enhanced output of high-energy protons results in very high atmospheric temperatures.

Temperature in this zone may further be increased by another mechanism. The temperature is normally defined in terms of transitional energy but absorption and emission of radiation occur through vibrational and rotational changes. In upper atmosphere, the frequency of molecular collisions is relatively low and so exchange of translational, vibrational and rotational energies is infrequent. Hence the cooling of thermosphere by re-radiation is very inefficient. The temperature of thermosphere increases with height so it is also stable against convection. Heat can be lost only by very inefficient diffusion processes and as a result, thermospheric temperatures are extremely high.


WATER BALANCE


The Earth’s hydrological cycle includes exchange of water in its various forms, between the hydrosphere, atmosphere, upper layers of lithosphere and the biosphere (living organisms). The corresponding process is described in terms of the equations of a water balance defined for each of the individual component of environment.

Water balance of land surface
Water balance equation for land surface may be given as:

r = E + fw + G

where, r = precipitation; E = evaporation at Earth’s surface which is equal to the difference between evaporation and condensation at Earth’s surface; fw = surface runoff; G = flow of moisture from Earth’s surface to deeper layers.

The above equation of water balance of land surface equates to zero the algebraic sum of all types of water inflows and expenditures, in solid, liquid and gaseous states entering a horizontal sector of surface from the surrounding space a specified interval of time. This equation is generally used in a slightly modified form. If water balance for a vertical column passing through the upper layers of lithosphere and reaching the depth at which moisture exchange with deeper layer stops is taken into consideration the vertical flow of water in lithosphere (G) is given as:
G = fp + b
where, fp = run-off within the soil; b = change in water content in upper layers of lithosphere.

Further, the full run-off i.e. run-off normal (f) is given as:

f = fp + fw

where, fp = run-off within the soil; fw = surface run-off.

Taking the equations for G and f into consideration, the water balance equation for land surface becomes:

r = E + f + b

Relation between energy and water balance on land
Important indicative characteristics of hydrological regime on land are:
(a) Run-off normal (f): It is the volume of water flowing off on the average during a year from a unit of land surface in form of various horizontal flows;
(b) Coefficient of run-off (f/r): It is the ratio of run- normal to total yearly precipitation.

(c) Radiation index of dryness (R/Lr): It is the ratio of the total radiation energy received to total energy expended on precipitation.


Evaporation is one the major processes of transformation of solar energy at Earth’s surface and it very much affects the yearly run-off. Thus, run-off normal and coefficient of run-off are linked to principal components of energy balance. Following general considerations apply to relation between elements of energy and water balance on land:


(i) Average total evaporation from land surface (E) depends on the quantity of precipitation r and inflow of solar energy. Evaporation increases with increase in and radiation balance r. If soil is dry, total precipitation is caught by molecular forces on soil particles and eventually expended on evaporation. In such conditions (e.g. in deserts), run-off coefficient (f/r) approaches zero.

Since the average dryness of soil increases when inflow of radiation energy increases and amount of precipitation decreases. Thus:

f/r ---> 0 or E/r ----> 1 when R/Lr ----> ∞

(ii) With decrease in value of R/Lr, value of E/ralso declines and certain run-off appears. For sufficiently high value of total precipitation and sufficiently small value of total inflow of radiation, a state of full moistening of the upper layers of soil will be achieved. In such cases, the maximum portion of heat energy available from all sources will be expended on evaporation. The value of this expenditure may be calculated by considering the valve nature of the latent heat exchange between underlying surface and the atmosphere.

Experimental studies have shown that turbulent heat conductivity of lower layers of atmosphere depends substantially on the direction of vertical turbulent heat flow. If the direction of this flow is from Earth to atmosphere, values of turbulent heat flow become higher due to increased turbulent mixing and the values become close to the values of the main components of radiation and heat balances. If direction of turbulent heat flow is from atmosphere to Earth, inverting temperature distribution reduces exchange intensity and turbulent heat flow becomes small. Thus, turbulent heat-flow is higher in daytime than in night and in middle latitudes turbulent heat exchange in winters is low than summers. As a result of the valve effect, average turbulent heat moves from Earth’s surface to atmosphere in nearly all climatic zones of land i.e. yearly sums of turbulent heat flows are positive. Thus yearly turbulent heat flows can not produce substantial inflows of energy to underlying surface and heat expenditure on evaporation are compensated only by the radiation balance. As a result the upper boundary for LE equal to R. Therefore, in condition of full saturation of upper soil layers, it may be assumed that:
LE à R when R/Lr à 0
Linkage equation for water and energy (heat) balance: The relation between water balance and energy (heat) balance of the land may be described by the equation:


E/r = Φ (R/Lr)


The form of the function Φ for R/Lr à 0 and for R/r à ∞ is determined by respective conditions considered above in (i) and (ii).

The most important consideration related to this equation is that the magnitude of possible evaporation at a locality is determined by the radiation balance corresponding to the conditions of sufficient moisture at which vegetation can exist. The values of radiation balance at a locality differ in conditions of sufficient and insufficient moisture. The reasons for this variation may be summarized as below:

(a) In moist regions of humid climate, albedo of Earth’s surface shows little change with change in humidity. The average difference between temperatures of Earth’s surface and air is relatively small in the presence of sufficient moisture. Therefore, magnitude of evaporation may be approximately inferred from the radiation balance corresponding to the actual state of Earth’s surface.

In regions of dry climate, albedo and temperature surface changes with increasing humidity. The Earth’s surface temperature approaches the air temperature in condition of sufficient moisture. Evidently, the magnitude of evaporation has to be determined from the value of radiation balance corresponding to the albedo of a moist surface and to to a temperature at Earth’s surface that is equal to air temperature.
Comparison of R/Lr and E/r values for observed data of river basins on different continents was made was made (Fig. 18). This has allowed verification of the important considerations concerning relations between E/r and R/Lr for both small and large values of R/Lr, The ratio E/r for low values of R/r is represented as straight line OA according to condition given in (ii). Similarly, the ratio E/r for large values of R/Lr is represented as straight line AB according to condition given in (i). The experimental points representing values of E/r are calculated from water balance data by averaging E/r for specific intervals of R/Lr. The experimental points clearly show a smooth transition from line OA to line OB, which as expected, are limiting values for the relation between E/r and R/Lr.

Important features of linkage equation
(a) The figure shows that the relationship on which linking of energy and water balances is based is largely determined by two limiting conditions. One condition corresponds to the valve mechanism of turbulent heat exchange in the layer adjoining Earth while second condition corresponds obviously the small value of coefficient of run-off (f/r) in dry climates. The corresponding relation for most of the interval of changes in the parameters of the linkage equation remains close to a boundary condition. Therefore, selection of one or the other interpolation function for passing from first to second condition is not particularly significant.


Thus despite semi-empirical nature of the linkage equation it can be justified mainly with the general considerations. Further, this equation represents a supplementary relationship, which is independent of the energy balance and water balance equations.

(b) The above linkage relation may be analytically represented by employing the following equation which represents the relation between average yearly evaporation and the radiation balance:

______________________________

E = /Rr/L th Lr/R (1 - ch R/Lr + sh R/Lr)



where, th is a function for hyperbolic tangent, ch and sh are hyperbolic cosine and sine respectively.

Since run-off normal (f) is equal to difference between precipitation and evaporation (f = r - E), equations representing f and f/r in terms of above equation will be:

______________________________

f = r - /Rr/L th Lr/R (1 - ch R/Lr + sh R/Lr)

_____________________________

f/r = 1 - /R/L th Lr/R (1 - ch R/Lr + sh R/Lr)


These relations between components of energy and water balance have been verified in a number of studies.

(c) The linkage equation makes it possible to represent the relation of run-off and evaporation to total yearly precipitation and the radiation balance in a general form (Fig. 2). The regularity explains a number of empirical relations between run-off and precipitation found in various studies.

(d) The linkage equation also makes possible to express dependence of run-off on precipitation for values of radiation balance corresponding to various localities (Fig. 3). Curve A in figure represents the relation of run-off normal and yearly precipitation calculated for average conditions in European lowlands using the equation. Curve B represents the empirical relation found by H. Keller (1906) on the basis of observation at West European rivers and curve C represents the empirical relation found by D. L. Sokolovsky (1936) on the basis of observation on East European rivers. Correspondence of empirical curves B and C with calculated curve A verifies the universal nature of the linkage equation.

(e) In calculation of average empirical dependence of run-off on precipitation in data from various regions, considerable dispersion in the dependence has been observed. The linkage equation makes it possible to explain this dispersion. There is substantial variability of radiation balance in middle latitudes. This variability causes the run-off of basins with large radiation balance (in more southerly regions) to be much smaller than that of basins with low radiation balance (in more northerly regions) for equal total precipitation values. This also influences the rate of change in run-off corresponding to increases in precipitation (df/dr). The linkage equation indicates that this rate should be larger in northern basins than in southern basins. Empirical data amply confirms this prediction of linkage equation.

(f) The calculations of run-off based on linkage equation correspond with the observed data. This confirms the decisive role of climatic factors, particularly of energy factors, in total yearly run-off in large river basins with areas comparable with geographical zones. In case of small areas river run-off may change substantially as a result of local conditions of non-climatic nature.




Water balance of water bodies


The above water balance equation can also be used to calculate water balances of water bodies and also of individual sectors within such bodies. In such cases, f will describe the total redistribution of water along the horizontal plane during the time interval under consideration both within the water body itself and in layers of underlying soil (in cases where significant redistribution of moisture at such levels does occur). Similarly, b for a closed water body will also be equal to overall change in the water content both within the body of water itself and in underlying layers in those cases in which there are perceptible changes in moisture content. In many cases, b usually describes changes in levels. For an average period of one year, b is often quite negligible and the water balance equation becomes:


r = E + f


For land surfaces having no run-off, e.g. deserts, f = 0 and the water balance equation takes the form:


r = E



Water balance of atmosphere


At a particular place in the atmosphere, the water is brought by evaporation from the Earth's surface and is removed by precipitation or air currents and horizontal turbulent exchange.


The water balance equation for atmosphere is, therefore, obtained by summing all the categories of inflow and expenditure of moisture within a vertical column passing through the atmosphere and is given as:

E = r + Ca + ba


where, Ca = quantity of moisture received or lost by vertical column due to air currents and horizontal turbulent exchange; ba = change in quantity of water in the column.


Since the atmosphere can retain only very small quantities of water in all its various states, value of ba is usually much less than other quantities. Its average value over a period of one year is always close to zero and the water balance equation becomes:

E = r + Ca


The atmosphere receives a considerable amount of moisture through transpiration of plants which usually represents several tens percent of total evaporation. Therefore, transpiration may exert a considerable influence on the atmospheric water exchange and hence, on the volume of precipitation. In studying this influence principles governing atmospheric water exchange are important. These principles have been formulated in quantitative theory of water exchange discussed below.


1. Flow of water vapor brought by air currents to a territory: For a territory whose average linear scale is L, this flow may be assumed to be wu where w = atmospheric moisture content in the windward side of area and u = average velocity of air flows carrying water vapor over the territory. There will be change in w along air current’s path according to the difference between precipitation (r) and evaporation (E).


2. Flow of water vapor carried away by air currents from the territory: This flow from the territory will be:


wu - (r -E)L.


3. The overall flow of water vapor carried over the territory: This is made of two subflows:

(i) Flow of external (advective) water vapor: It is the moisture originating outside the territory by evaporative processes and brought to the territory. On the windward side of area, it is equal to wu while on leeward side (as flow leaves the area’s boundaries), it will be wu - raL where, ra precipitation produced by this advective water vapor in area. On average such flow over a territory is given by:

wu - 0.5 raL


(ii) Flow of local water vapor: It is the moisture originating within the area by local evaporation. In areas having plant cover, it consists mainly of transpiration. This local flow is equal to zero on windward side and (E - rl)L on leeward side where rl = total precipitation produced from this local water vapor in the area. On average such flow over a territory is given by:


0.5(E - rl)L


Thus, the overall flow of water vapor over the territory produced by these two subflows together is given by:


wu - 0.5(r -E)L

and,

r = ra + rl


The molecules of water vapor of external origin and local origin are mixed in the atmosphere during turbulent exchange. Therefore, the ratio of total precipitation derived from external water vapor and that derived from local water vapor is equal to the ratio of the quantity of corresponding molecules of water vapor in the atmosphere. In short, it may be assumed that:


ra/rl = [wu - 0.5 raL] / [0.5(E - rl)L]


This gives two equations:


(a) ra = r / [1 + EL/2wu];

(b) rl = r / [1 + 2wu/EL]



Coefficient of water exchange (K): This is equal to the ratio of total precipitation ® to precipitation from advective moisture (ra). It is determined from the equation for ra and given as:

K = r/ra = 1 + [EL/2wu]


The above equation shows that:

(i) The K depends on the factors determining water (vapor) balance in the atmosphere.

(ii) Relations in formulae for ra, rl and K depend on the size of territory under consideration. With increase in scale L, rl and K increase but ra decreases. For sufficiently territories, actual dependence of K on their size is not linear. For large territories, average rate of transfer of water vapor (u) somewhat declines due to curvature of air particle trajectories.



Contribution of local evaporation to total precipitation

Direct contribution: Data of water exchange over all the continents of Earth given by O. A. Drozdov and associates is given in the following table. The data shows that the principal source of water vapor for precipitation over continents is advective water vapor, which originates as oceanic evaporation. Owing to large-scale transfer of oceanic water vapor in the Earth’s atmosphere, direct contribution of local evaporation processes from the surface of continents to the total volume of precipitation is relatively modest, especially for the territories of dimensions less than 1 million square kilometers. However, influence of local evaporation on total precipitation is not limited to changes in the components of atmospheric water exchange only but includes indirect influence also.
Indirect influence: The indirect influence of local evaporation on totalprecipitation derives from the linkage between the volume of precipitation and the relative humidity atmosphere.


Studies of Drozdov (1963) have shown that volume of total precipitation may be established by following semi-empirical formula:

r = aw f(h)

where, r = total precipitation; h = average relative humidity within atmospheric layer upto 7 km altitude; w = atmospheric moisture content; a = proportionality coefficient which is equal to 1 when total precipitation of 0.1 w per day falls for 100% h.

Empirical function f(h) is generally similar for many regions of middle latitudes. The formula may be employed to establish the effect of local evaporation processes on total precipitation. The values of w and h are then determined according to the influence of local evaporation on the overall quantity of water vapor being carried over the given territory. In this connection, it becomes clear that the influence of local evaporation on total precipitation within specific limits of air humidity may be several times larger than the contribution of local evaporation to overall volume of atmospheric water vapor.

Studies of Drozdov have shown that without the large role that is often played by the indirect influence of local evaporation on precipitation, the continental regions at distance from ocean would have been transformed into deserts.

Effect of plant cover on precipitation

The precipitation in an area also depends on the plant cover which influences the quantity of local evaporation. It has been observed that when mossy swamps are drained and a forest cover develops in such area, total evaporation appears to increase. The reverse effect is observed when lowland grassy swamps are drained. It may be assumed that the replacement of various types of plant cover changes total evaporation by about 10%.

With the use of formula r = aw f(h) it is indicated that influence of evaporation on precipitation depends on the size of area in which evaporation changes. Within area of linear scale of the order of 1,000 km, the influence of fluctuations in evaporation on precipitation is negligible. However, this influence may be substantial in case of changes in evaporation over entire continents or large parts of continents.

It is possible that during Lower Palaeozoic era when plant cover existed on continents, the ability of upper layers of lithosphere to retain water was quite limited. Under such conditions, evaporation from land would be low and run-off high. Since this condition of water balance components existed over large areas of Earth, it exerted substantial influence on precipitation on continents. With appearance of plant cover on land, soil formation increased which caused increase in evaporation and decrease in run-off. Thus with increasing land area coming under plant cover, volume of atmospheric precipitation also increased. This in turn contributed to further increase in spread of plant cover within continents. Thus a positive feedback relation exists between living organisms and their environment which has contributed to spread of plant cover over much of land surface. However, this feedback relation could alter precipitation only in areas having favourable conditions of atmospheric circulation.

Water balance of whole Earth

If water balance of Earth as a whole is considered, the horizontal redistribution of water is of no importance and the water balance equation simplifies to:


r = E


The average yearly values of precipitation, evaporation and run-off for different continents and oceans is given in following Table-5. Features of water balance of Earth can be discussed on the basis of these values.


(1) For each continent, evaporation is equal to the difference between precipitation and run-off.


(2) The ratio of evaporation to run-off differs greatly on various continents. In Australia evaporation is close to precipitation. In all other continents except Africa, evaporation is less than about 66% of the sum of precipitation.


(3) The difference between evaporation from the surface of World Ocean and precipitation is equal to the river run-off from continents into the oceans.


(4) For the individual ocean, evaporation is equal to the sum of the river run-off into it and horizontal transfer of water from other oceans through global oceanic circulation processes. It is difficult to determine the magnitude of horizontal transfer through direct methods because it represents difference between two small values of inflow and outflow of water. Determination of both values is subject to significant errors. It may be simpler to estimate the exchange of water between oceans as a residual element in the water balance. The accuracy of estimates in this case is also limited.


(i) For Atlantic Ocean sum of precipitation and run-off is less than the magnitude of evaporation. It is clear that this ocean receives water from other oceans including Arctic Ocean where evaporation is substantially smaller than the sum of precipitation and river run-off.


(ii) In Indian Ocean, sum of precipitation and run-off is somewhat smaller than the magnitude of evaporation while in the Pacific Ocean, the sum of precipitation and run-off is larger than evaporation, which reflects a transfer of water into other oceans.


(5) Values of precipitation and evaporation for land and oceans indicate that for the Earth as a whole, the magnitude of precipitation over a year is 113 cm, which is equal to that of evaporation.

(6) Calculations of the values of precipitation and evaporation for various latitudinal zones show that the inflow of water vapor into the atmosphere from evaporation may be both larger and smaller in different latitudinal zones than its expenditure on precipitation. The source of water vapor in atmosphere is provided by high-pressure zones where evaporation is much more than precipitation. That surplus water is expended in zones adjacent to the Equator and, also at middle and high latitudes where precipitation is much more than evaporation.


(7) The difference between precipitation and evaporation is also equal to the difference between the inflow of water vapor into the atmosphere and the outflow resulting from the horizontal air movement. The large difference in these two in many regions indicates the importance of the transfer of atmospheric water vapor in the formation of precipitation. The influence of transfer of water vapor on the volume of precipitation has been examined in the studies of hydrological cycle in the atmosphere.


(8) The components of continental and oceanic water balance are not constant and change as a result of climatic fluctuations and other factors. Since changes in the components of water balance over a year are small in comparison with their absolute values, they may be neglected in determining the corresponding magnitudes. On the other hand, data on changes in water balance components may be very important in the studies of the evolution of hydrosphere.


(9) During glaciation periods, large masses of water used up in formation of continental ice sheets and volume of water in the World Ocean had changed. At the termination of last glaciation, approximately 20 thousand years ago, the level of World Ocean was lower by about 100 meters than it is today. Subsequently the level gradually increased and reached the present level about 5,000 years ago and has not significantly changed since then.


(10) Observations have shown that in the 20th century, level of World Ocean has increased by about 15 cm while reserves of subsoil water on land and volume of water in many lakes have declined. Though sources of water contributing to rise in water level in World Ocean are not absolutely clear, melting of ice covers, loss of groundwater and decrease of water in lakes on land are responsible for the increase in volume of oceanic water.


CLIMATIC INDICES CHARACTERIZING BIOGEOGRAPHICAL ZONES

The study of parameters of radiation and water balance of Earth clearly shows that these parameters differ with latitudes. The interactions of these parameters of radiation and water balance are responsible for particular meteorological conditions in different regions of Earth which, in turn, result in climatic zones. Further, the biosphere of Earth, particularly the plant cover over land surface is characteristically dependent on the climatic conditions of a particular region. Thus, a number of distinct biogeographical zones can be recognized over Earth’s surface on the basis of characteristic climate, soil and plant cover. Dokuchavev for the first time observed that the boundaries of biogeographical zones are largely determined by climatic factors and are particularly dependent on moisture conditions. His classical studies included estimates of the ratios of precipitation to potential evaporation for major biogeographical zones of Earth. Various studies after him examined the relationship of soil and forest types with ratios of precipitation to potential evaporation. Grigoriyev and Budyko in comprehensive studies of global environment identified the most important parameters of radiation and water balance which determine the biogeographical zonality and gave the Law of Geographical Zonality.

LAW OF GEOGRAPHICAL ZONALITY

Budyko showed that equations for mean yearly heat and water balances of land may be written in the following form:

R/Lr = E/r + P/Lr and 1 = E/r + f/r

(the members of heat balance equation are divided by Lr and those of water balance equation by r).

The linkage equation of water and heat balance may then be added to these relations:

E/r = @ (R/Lr)

These equations link four relative values of components of heat and water balances. Therefore, it is sufficient to know any one of them for determining the others. Owing to special form of the linkage equation, ratio R/Lr or P/Lr may be selected as the parameter that determines all relative values of the components of heat and water balances. The ratios E/r and f/r can not be decisive factor for the first two variables in case of dry climates when small changes in E/r or f/r produce large changes in R/Lr or P/Lr. Further, it seems more appropriate to select R/Lr as the principal parameter determining the relative values of the components of heat and water balances. This parameter may also be viewed as the ratio of potential evaporation R/L to precipitation r, or else as a ratio of the radiation balance of a moistened surface to heat expended on the evaporation of the total yearly precipitation

While the relative values of components of heat and water balances are determined by one parameter R/Lr, the determination of absolute values of these components requires two parameters namely R/Lr and R.

It has been shown by Budyko that it is possible to calculate potential evaporation from the radiation balance at the Earth’s surface. The method for calculating potential evaporation from air humidity deficit as various other empirical methods yield less accurate results.

A world map of the radiation index of dryness (R/Lr) has been constructed by Budyko (1955). A comparison of this map with geobotanical and soil maps confirms that the positions of isoclines of the index of dryness confirm well to the distribution of major climatic and biogeographical zones. It has been shown that the radiation index of dryness accords well with the boundaries of major natural climatic and biogeographical zones. In general, values of the radiation index of dryness (< 1/3) correspond to tundras, values between 1/3 and 1 to forest zone, from 1 to 2 to the steppe zone, more than 2 to semi-desert zone and more than 3 to the desert zone. Further, within particular zones at various latitudes substantial differences in conditions determining the development of natural processes can be observed. These differences derive from the fact that the energy base of natural processes may be characterized through the magnitude of the radiation balance R and it differs at various latitudes. Accordingly, while it may be possible to make use of single parameter R/L in characterizing general zonal conditions of natural processes, the characterization of absolute values of the intensity of natural processes requires the use of two parameters i.e. R/Lr and R which determine the absolute magnitudes of the elements of heat and water balances.

The relation of biogeographical conditions with the parameters R/r and R plotted along its axes and on which major biogeographical zones are divided by straight lines. A schematic form of such a graph is given in Fig.

While the use of more accurate precipitation normals found with help of currently available measuring instruments introduce certain changes in the characteristic values of moisture indices, this does not alter the general patterns of the relation of moisture indices to biogeographical zones. The close relationship of parameters R/Lr and R with biogeographical zones has been examined and named the Law of Geographical Zonality (Grigoriyev and Budyko; 1956, 1962). It was specifically noted that within each latitudinal belt there exists a definite correspondence between the boundaries of natural zones and the isoclines for particular values of R/Lr. Table- represents the overall structure. In the compilation of the table, use was made of the values of R relating to the conditions of moistened surface and of more detailed data on precipitation and plant cover in various regions. To each section in the table, characterizing gradations of moisture conditions correspond to specific values of the coefficients of river run-offs. At the same time, yearly run-off normal in each section increases as radiation balance increases. However, this does not apply to desert areas where the run-off at low latitudes is close to zero. Important conclusions of the Law of Geographical Zonality are discussed below.


1. R/Lr and botanical zones: Similar types of plant cover correspond to each section in the table. In conditions of excessive moisture, forests are prevalent at all latitudes except for areas with substantial excess of moisture i.e. R/r < 2/5. In such cases, forests are replaced by tundras at high latitudes and by swamps at lower latitudes. Since R/r < 2/5 over vast territories is found only at relatively high latitudes, swampy areas at low latitudes can not be viewed as autonomous biogeographical zones. It should be noted that the ordinate represents values of the radiation balance for actual state of Earth’s surface, which differ from the values of the radiation balance for moistened surfaces. The continuous line on the graph bounds the region of actually occurring values of R/r and R (except for mountainous regions) within whose ranges specific values of R/Lr, which are shown as vertical lines, delimit the major botanical zones. Large changes in the values of radiation balance for forest zone correspond to perceptible geobotanical changes within these zones in spite of the common features of their plant covers.

Each gradation in moistening is characterized not only by a definite type of plant cover but also by a specific value of its productivity. In the studies of Grigoriyev and Budyko (1956, 1962), it was assumed that the productivity of natural vegetation increases as the moisture conditions approach optimal ones for a given radiation balance and for given moisture conditions, productivity increases with increase in radiation balance.

2. R and R/Lr and soil zones: It has long been established that there exists a close relationship between the zonality of plant covers and the zonality of soils. Therefore, the conclusions concerning the relation of botanical zones to specific values of parameters R and R/Lr can be fully applied to soil zones as well. It can be established that as the values of parameter R/Lr increase, soil types change in the following sequence:

(i) Tundra soils;

(ii) Podzols, brown forest soils, yellow soils, red soils and laterite soils (the diversity of soil types within that group corresponds to changes in the parameter R within the broad range);

(iii) Chernozems and black soils of savannah regions;

(iv) Chestnut brown soils;

(v) Grey soils.

General relation of soil zonality to climatic indices R/Lr and R may be represented as a graph similar to the graph for botanical zones. Considering the Table-, to each section of the table, there corresponds a specific sequence of changes in soil types which is largely similar within each section. Fro example, the third section is characterized by following sequence of soil types: tundras, podzols and brown forest soils, subtropical red soils and yellow soils, tropical podzol red soils and laterites. For the fourth section, the sequence is chernozems and dark chestnut soils, black soils and brown soils, weakly leached subtropical soils, red brown tropical soils etc. To each such sequence there correspond specific values of quantitative characteristics of the process of soil formation. The region of eternal snow occupies a special place within the table. It is described by negative value of R and R/Lr due to practical absence of plant and soil covers. Similarly, the subzone of Arctic deserts has negligible values of yearly radiation balance and a high humidity.

3. R/Lr and R and zonality of hydrological regime of land: The relation of the zonality of hydrological regime of land to parameters R/Lr and R may be established on both quantitative and qualitative terms. From the linkage equation it follows that to each gradation in value of parameter R/Lr there corresponds specific gradation in the value of run-off coefficient. A consideration of the influence of energy factors makes it possible to explain the zonal changes in run-off coefficients in quantitative terms. The absolute values of total run-off are determined by two parameters, namely R/Lr and R. Graph in Fig. which is similar to that in Fig. represents the distribution of yearly total run-off and characterizes the absolute values of run-off in different biogeographical zones.

CAUSES OF PATTERNS OF GEOGRAPHICAL ZONALITY
From the above discussion, it is evident that the patterns in the Table-6 of Geographical Zones may be explained in terms of the following factors:

1. Latitudinal difference in Earth’s surface radiation balance: Owing to the spherical shape of Earth, its surface is divided into several latitudinal belts that differ in terms of radiation energy inflow to the surface of Earth.

2. Differences in moisture conditions: Within each of these belts (except for the region of eternal snow) there exist different moisture conditions ranging from excessive to highly insufficient moisture. These differences in moisture conditions within each latitudinal belt are responsible for the zonality within the belt..

Within different latitudinal belts, geographical conditions with similar moisture indices have a certain number of common characteristics, which are combined with differences resulting from different inflows of radiation energy. These common characteristics are periodically repeated as regions of two latitudinal belts in which humidity (or dryness) increases) are compared with each other.

SEASONAL CHANGES IN CLIMATIC FACTORS OF BIOGEOGRAPHICAL ZONALITY
The influence of climatic factors on biogeographical zones may be represented more clearly by considering the specific features of the climatic regimes in different seasons. In specifying the climatic conditions of particular seasons, following major types of climatic regimes may be identified:

1. Arctic climatic regime: It is characterized by the sow cover, negative air temperatures and negative values of radiation balance or else values close to zero.

2. Tundra climatic regimes: Such climatic regimes have average monthly temperatures ranging from zero to 10o C and a small positive radiation-balance.

3. Climatic regimes of forest zones: These have average monthly temperatures of more than 10o C and a positive radiation balance as well as sufficient moisture, when evaporation exceeds one half of the potential evaporation.

4. Climatic regimes of arid zones (Steppes and savannahs): These have a positive radiation balance and actual evaporation ranging from 1/10 to ½ of the potential evaporation.

5. Climatic regimes of deserts: These have positive radiation balance and an evaporation less than 1/10 of potential evaporation.


Important features of seasonal changes in climatic factors determining biogeographical zonality are:

(i) It has been shown that the type of climatic conditions in each month correspond with the type of natural zone throughout the year. However, in most biogeographical regions, several types of climatic regimes replace each other during the year.

(ii) There are substantial differences in seasonal changes of climatic regimes at various longitudes which are especially felt at middle and low latitudes where types of climatic regimes depend on moisture conditions.

(iii) In Africa and Europe, a regime of low humidity exists within a wide latitudinal belt that shifts to the north in summers and to the south in winters. This corresponds to the most favourable conditions in the subtropics in the winter and spring and at tropical latitudes in the summer. In East Asia and North America, the structure of moisture regime differs considerably from the first scheme because of substantial differences in circulation processes.

(iv) In most regions and at all longitudes there is either insufficient moisture (regimes 4 and 5) or insufficient heat (regime 1 & 2) during a major part of the year. Only in the narrow belt close to Equator, conditions corresponding to regime 3 exist throughout the year. Evidently under regimes 2 and 4 and especially under regimes 1 and 5, productivity of natural plant cover is reduced.

(v) Under conditions of insufficient heat the type of biogeographical zone is determined by climatic regime of the period in which productivity of natural plant cover is greatest even if that period is relatively short. For example, the zonal landscape of tundra is determined by conditions of the warm season, which may last no longer than 1/5th to 1/4th of the year. In such cases, climatic regime of the cold season, which extends over most of the year does not determine the landscape’s zonal character.

(vi) In regions of insufficient moisture also a regularity similar to the above one is observed. The most humid period of the year plays a determining role in establishing the type of zone, even though it may be shorter than the period with insufficient moisture. In natural zones with insufficient moisture, there may be short periods of either sufficient or excessive moisture which do not result in development of typical forest growth.


The above described regularities relating the type of biogeographical zone to seasonal changes in the climatic factors determining the zonality, in fact, complement the concept of the influence of climatic factors on biogeographical zonality described earlier in the Law of Geographical Zonality.