Environment of Earth

March 11, 2008


Filed under: Atmospheric chemistry,Environment — gargpk @ 2:24 pm
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The atmosphere is made up of a large number of gaseous constituents and in the atmosphere, a large variety of chemical reactions are constantly going on amongst its constituents. The products of one reaction are reactants for other reactions i.e. chemical reaction constitute many of the sources and sinks of gases in the atmosphere. Since the study of atmospheric chemistry mainly attempts to understand the chemical kinetics of atmosphere, a study of the gas-phase reaction rates becomes very important.

Gas-phase reaction rates

For understanding gas-phase reaction rates, following simple reaction may be considered:

2NO + O2 ——-> 2NO2

The reaction rate (R) is given by:

R = – {d[NO]/dt} = – {d[O2]/2dt} = d[NO2]/dt

In atmospheric context, concentrations are usually expressed as number of molecules of gas per cubic centimeter (cm-3). Therefore, the rate of above reaction may be expressed in the form:

R = k[O2]1[NO]2

Where, k = reaction rate constant; exponents denote the order of reaction.

This reaction is second-order with respect to nitric oxide, first-order with respect to oxygen and third-order overall i.e. sum of individual orders. As the units of reaction rate are given in concentration per unit time (cm-3 s-1 ), the rate constant for the above reaction will have the units of cm-6 s-1.

The consideration of reactions of various orders shows that:

(a) Rate constant for first-order reaction will have the units of s-1.

R = – {d[A]/dt} = k[A]t

or, d[A]/[A]t = – k dt

On integration, it gives:

ln[A]t = – kt + c

c is concentration of A at t = 0 and is expressed as [A]0. Thus:

ln{[A]t/[A]0} = – kt

or, [A]t = [A]0 e-kt

Similar expression for reactions of other orders may be given:

Order Differential form Concentration relationship Units

0 – d[A]/dt = k [A]t = [A]0 – kt cm-3 s-1

0.5 – d[A]/dt = k[A]0.5 [A]t = {[A]00.5 – kt/2}2 cm1.5 s-1

2 – d[A]/dt = k[A]2 [A]t = [A]0/{kt[A]0 + 1} cm3 s-1

(b) In case of second-order reaction, if concentration of one reactant is significantly higher than that of other reactant then the reaction can generally be treated as a first-order reaction with respect to reactant at low concentration. For example, reaction of ozone with nitric acid in atmosphere is a second-order reaction:

NO + O3 —–> NO2 + O2

and – {d[NO]/dt} = k”[NO][O3]

k” is the second-order rate constant and for this reaction, its value is 1.8 x 10-14 cm3 s-1 at 300 K. Concentrations of NO and O3 may be assumed as being 0.1 ppb and 15 ppb respectively that are the typical concentrations for lower atmosphere. These concentration units have to be made compatible with those of rate constant. At atmospheric pressure of 1.0, one cm3 of gas contains 2.7 x 1019 (Loachmidt number). This number has to be multiplied by partial pressure of the gas to give appropriate units. Thus, the concentrations of NO and O3 come to be 2.7 x 109 cm-3 and 3.9 x 1011 cm-3 respectively. Using these concentrations, the rate of reaction comes to be 1.8 x 107 cm-3 s-1. This rate is quite high compared with concentration of NO and so the concentration of NO will decline quite rapidly in a closed system. On the other hand, concentration of ozone is very much greater than that of NO and will remain relatively constant. Therefore, the concentration of ozone may be incorporated as a constant within the rate constant and the rate expression can be given as:

– {d[NO]/dt} = k’[NO]

where, k’ is first-order constant given by k”[O3]. The value of this pseudo-first-order rate constant is 0.007 s-1 at the concentration of ozone being considered.

Apparent reduction in the reaction order of a system may occur in other ways also. All that is required for it is that the concentration of one reactant should remain constant. If a reactant is being catalyst or continuously replaced in the system, this might be the condition.

(C) When reactions can be reduced to first-order systems, their study becomes convenient. In such systems, concentration of the reactant will halve over a constant time period regardless of its initial concentration. This period of time is termed half-life (t0.5) and is related to the first-order constant by the expression:

t0.5 = ln(2)/k’ = 0.693/k’

In second-order systems, half-life is dependent on the second-order reactant and relationship between t0.5 and second-order rate constant is given by the expression:

t0.5 = 1/k [A]0

For reaction orders greater than unity, higher the concentration, shorter is the half-life.

Atmosphere as steady-state system

From the above short discussion of chemical kinetics, great stability of the atmosphere would not be expected since it is not in a state of chemical equilibrium. However, the natural atmosphere appears to be quite stable. This apparent stability of atmosphere is because it is in steady state in which various chemical species are continuously being added and removed. The steady-state situation of the atmosphere is maintained by relative constancy of the input and output i.e. by the fact that the rates of the addition/production and removal/destruction of chemical species are equal.

For describing steady-state systems, the residence time or mean lifetime of a chemical species is a useful parameter. Residence time () is easily obtained from the first-order rate constants.

The flux of a material from a system is given by:

Fo = A/

Considering the flux as material lost in a given volume through a chemical reaction of first-order and writing the amount as concentration:

Fo = – {d[A]/dt}


d[A]/[A]t = -kdt


d[A]/dt = -k[A]t


Fo = [A]/ = k[A]


 = 1/k’

Concentration [A] is assumed to remain constant due to continual input of material. This is true of a steady-state system. Thus, the residence times are useful for describing steady-state situations. With steady-state assumption for the oxidation of nitric oxide by ozone discussed above, the calculated value of pseudo-first-order rate constant is 0.007 s-1. This value indicates that nitric oxide has a residence time of about 150 s in the atmosphere.

Various chemical species are continuously added into the atmosphere from the lithosphere or biosphere and are destroyed through complex reaction systems in the atmosphere. Thus atmosphere is a steady-state system and quite complex reaction systems in the atmosphere (e.g. Methane cycle discussed later) can be easily understood if they are assumed to be in the steady state.


Water is present in the atmosphere as suspended droplets and atmospheric gases are dissolved in this atmospheric water. The dissolution of atmospheric trace gases into suspended droplets is one of the most important controls on rainfall chemistry.

Henry’s Law describes the solubility of gases in water and states that at equilibrium the partial pressure of a gas above a solution of the gas is proportional to the concentration of the gas in the solution. However, in much of the atmospheric chemistry, it is useful to imagine the relationship between the gaseous and liquid phase concentrations in terms of a equilibrium of the type:

A(g) = A(aq)

Where, A(g) and A(aq) represent the concentrations of substance A in gaseous and aqueous phases respectively.

By writing the Henry’s Law constant (KH) as the equilibrium constant for this reaction and using pressure to describe the concentration of A in the gaseous phase:

KH = [A(aq)]/pA

If units of pressure and concentration are taken as atm and mol l-1 respectively, Henry’s Law constant will have the units of mol l-1 atm-1. It is clear that larger the value of KH, more soluble the gas will be. Therefore, H2O2 is highly soluble and its large amounts can dissolve in the clouds and rainwater droplets. KH values for some atmospheric trace gases are given in the Table-1.

Table-1. KH values for some atmospheric trace gases at 288 K


Gas                                       KH (mol/l/ atm)


Hydrogen peroxide            2 x 10^5

Dieldrin                                   5800

Lindane                                   2230

Ammonia                                     90

Aldrin                                            85

DDT                                                 28

Sulfur dioxide                              5.4

Formaldehyde                             1.7

Mercury                                            0.093

Carbon dioxide                              0.045

Acetylene                                        0.05

Nitrous oxide                                 0.034

Ozone                                                0.02

Nitric acid                                      0.0023

Methane                                           0.0017

Oxygen                                             0.0015

Nitrogen                                          0.001

Carbon monoxide                       0.001


Henry’s Law constant accounts only for simple dissolution of gases and not for the condition where there is hydrolysis after dissolution. For example, formaldehyde dissolves in water and subsequently hydrolysis to methylene glycol according to the following equations:

HCHO(g) ===== HCHO(aq)

HCHO(aq) + H2O ====== H2C(OH)2(aq)

Thus, the apparent solubility of formaldehyde in water is greater than that expected from Henry’s Law constant. The total amount of formaldehyde dissolved (T(HCHO) in solution will be:

T(HCHO) = [HCHO(aq) + [H2C(OH)2(aq)]

The concentration of methylene glycol will be related to the aqueous formaldehyde by the Laws of mass action:

K = [H2C(OH)2(aq)]/[HCHO(aq)]

Where, K is the equilibrium constant fro hydrolysis reaction. This gives:

T(HCHO) = [HCHO(aq)]+ K[HCHO(aq)]

Since [HCHO(aq)] is known from Henry’s Law, above equation may be written as:


In case of formaldehyde, K is about 2000, i.e. gas is readily hydrolyzed by water so most of it will be found in aqueous solution as methylene glycol rather than as formaldehyde. This makes formaldehyde rather more soluble. The KH is about 1.7 mol l-1 atm-1. At equilibrium with atmosphere pHCHO at 10-9 atm, total concentration of the formaldehyde derived carbon would be predicted to be about 3.4 x 10-6 mol l-1.

Dissolution and hydrolysis of formaldehyde is a rather simple case. Many other atmospheric gases such as carbon dioxide, sulfur dioxide and ammonia undergo more complex hydration reactions and the pH of rainwater is significantly influenced by sets of these hydration reactions.

(I) CO2 + H20 ==== H2CO3(aq)

H2CO3(aq) ==== H+(aq) + HCO3-(aq)

HCO3-(aq) ==== H+(aq) + CO32- (aq)

(ii) SO2(g) + H2O ==== H2SO3(aq)

H2SO3(aq) ==== H+(aq) + HSO3-(aq)

HSO3(aq) ==== H+(aq) + SO32- (aq)

(iii) NH3 +H2O ==== NH4OH(aq)

NH4OH(aq) ==== NH4+(aq) + OH-(aq)

Henry’s Law constant and equilibrium constants for these reactions i.e. KH, K’ and K” respectively are given in the following Table- 2.

Table-2. KH, K’ and K” values for some important atmospheric gases.


Gas KH (mol/l/atm) K’ (mol/l) K” (mol/l)


Carbon dioxide 0.045 3.8 x 10^7 3.7 x 10^11

Sulfur dioxide 5.4 2.7x 10^2 2.7 x 10^7

Ammonia 90 1.6 x 10-5 –

Though sulfur dioxide is a soluble gas, it does not all dissolve in liquid phase of a system of cloud-droplets suspended in air. The ratio of the volume of water and the volume of air is quite low usually being less than 10-6. Therefore, most of the mass of sulfur dioxide remains in gaseous phase in a cloud. Among common atmospheric trace gases, hydrogen peroxide is probably soluble enough to partition predominantly into liquid phase. Under the acidic conditions, ammonia may also be found predominantly in the liquid phase.

Transfer of gases to liquids

When a gas is in high concentration in the atmosphere and at very low concentration in water droplet, there occurs a flux of the gas to the water. The flux of gases across an air-liquid boundary is usually described in terms of a two-film model. This model assumes that there are thin boundary layers on either side of the gas-liquid interface and transfer through these layers is governed by diffusion. As diffusion is a slow process compared with turbulent transport, transfer through the still boundary layer limits the flux of gases to water bodies. In gas-liquid transfer, the total resistance to transport will be the sum of the individual resistance of the two layers. For most of the important atmospheric gases (except perhaps formaldehyde), one of these two resistance is greater than the other resistance. Therefore, the transport of gases across gas-liquid interface may be of two types: gas phase controlled transport and liquid phase controlled transport.

(I) Gas-phase controlled transport: Such transport generally occurs in case of highly soluble gases. The flux (F) of a highly soluble gas across a boundary layer at the air-water interface is given by:

F = k c

where, k = exchange constant having units of ms-1 and c = difference in concentration of gas between the bulk gaseous concentration and the gaseous concentration at the liquid surface. The exchange constant is the reciprocal of the resistance ® and can be obtained from the diffusion coefficient of the gas (D) in the boundary layer (z) i.e.

k = 1/r = D/z

As resistance can be summed, exchange constants must be summed as reciprocals i.e. 1/K = 1/k1 + 1/k2 + …., where K = exchange constant of whole system of boundary layers.

(ii) Liquid-phase controlled transport: Such transport occurs mainly in case of less soluble gases. In such case, c represents the difference between concentration in the liquid at the surface and the concentration in the bulk of liquid. Transfer into liquids can be fast in another way also. Rapid reaction of gas in the liquid phase lowers the liquid-phase boundary layer and transfer becomes gas-phase controlled. Sulfur dioxide is such a gas that is rapidly hydrolyzed to bisulfite and sulfite ions in water so that its dissolution in the water is under gas-phase control and is quite rapid (about 0.5 to 1.0 cm/s). This hydrolysis is only significant for sulfur dioxide at pH>3.0 and so in very acidic solutions, the dissolution will remain under liquid-phase control.


The dissolution of gases in water droplets suspended in the air can be considered as a two step process:

(a) Transfer of gas from bulk atmosphere to the surface of droplet

(b) Mixing of gas within the droplet

Gas-phase transport processes are usually fast so the rate of transfer is limited by the mixing within the droplet. If gas is quite rapidly transferred through the gas-phase boundary layer then the surface of droplet can reach equilibrium with the gaseous phase. More gas can dissolve in the droplet after dissolved gas is mixed inward from the surface of droplet towards its center. In the droplet falling through atmosphere, mixing may also occur through convective stirring. If droplet is stagnant, mixing occurs only through the slower diffusive processes. Diffusive transport within a sphere is given by:

Mt/M = 1 – (6/)  1/n2 exp(-Dn22t2/r2)


where Mt and M are masses of the substance at time t and at equilibrium respectively; n is an integer, r is radius of the sphere.

Typical value for the radius of suspended droplets in air is about 50m while value of diffusion coefficient typical for many dissolved gases is about 10-9 m2s-1. These values suggest that 50% saturation of droplet is achieved in 0.3 second indicating that equilibration of water droplets with atmospheric gases is quite rapid. Large suspended drops may take longer to equilibrate but their equilibration can be quite rapid if they are falling and are being stirred by the air flow on their surface.


The dissolution of soluble atmospheric trace gases into droplet results in much increased concentration of those gases in the small volume of water droplet. This increased concentration of gases allows many new opportunities for chemical reactions inside the droplet despite the fact that only very soluble gases are found partitioned predominantly into the droplet phase. Important such reactions are discussed below.

1. Oxidation of sulfur dioxide: This has been the most frequently studied reaction in aqueous atmospheric droplets due to its importance in acid rain problem. Oxidation of sulfur dioxide by oxygen is very slow in absence of catalysts. Only through the presence of catalysts, such as iron or manganese, such oxidation can be fast enough to be important in atmospheric droplets. At typical acidity (i.e. about pH 5.0) of atmospheric aerosols, sulfur dioxide will be present mainly as bisulfate ion (HSO3-) so the oxidation reaction may be given as:

HSO3-(aq) + 0.5O2(aq) –Fe, Mn———–> SO42- + H+(aq)

Despite many attempts, it is still not clear whether iron or manganese present in rainwater is principally responsible for oxidation of sulfur dioxide.

In remote areas these catalysts, though abundant from crustal sources, may not be in a form that is soluble enough to promote the oxidation reaction. In such areas, H2O2 and O3 may be the oxidants though they are present in atmosphere at quite low concentrations:

HSO3-(aq) + H2O2(aq) ——–> SO42- (aq) + H2O + H+(aq)


HSO3-(aq) + O3 ——–> SO42- (aq) + O2 + H+(aq)

The product of these reactions is sulfuric acid that being much stronger acid than sulfurous acid is responsible for appearance of proton on the right side of above reactions. As oxidation proceeds, droplet becomes more acidic than it was owing to sulfurous acid alone: first due to production of sulfuric acid and secondly from dissolution of more sulfur dioxide to replace the oxidized gas. The solubility of sulfur dioxide is lowered with the increase in acidity but in case of catalyzed reaction, the actual rate of oxidation slows down too. Thus, the oxidation reaction can rapidly come to a standstill. However, many workers consider H2O2 to be a very effective oxidizing agent of SO2 in atmosphere because the rate of oxidation by it actually increases under acidic conditions. So the oxidation rate is not slowed down as the reaction proceeds and more H2SO4 is produced.

2. Oxidation of nitrogen oxides: Nitrogen oxides may be oxidized in droplets to a lesser extent to form nitric acid. This system has not been studied in much detail. There is also the possibility of the dissolution and reduction of two nitrogen oxides:

NO2(aq) + NO3(aq) + H2O ———-> 2H+(aq) + 2No3- (aq)

3. Reaction with chloride ion: The production of sulfuric or nitric acid may result in an important subsequent reaction if chloride ion is present in high concentration in the atmosphere. Hydrogen chloride is more volatile than other strong acids found in aerosol droplets, so it may be lost from the droplet according to the reaction:

H+(aq) + NaCl(s) ———> HCl(g) + Na(aq)

Study of the Na:Cl ratio in maritime aerosols has provided evidence of the occurrence of this reaction. However, quantitative description of the reaction is rather difficult because it probably occurs in droplets which have nearly evaporated to dryness. These aerosols will have a very high salt concentration. Under such situation, behaviour of sodium departs from ideal condition meaning that thermodynamic predictions made using equilibrium constants obtained from low concentrations may be wrong i.e. the solubility of gases in saline droplets may not follow Henry’s Law.

4. Dissolution and hydrolysis of carbonyl sulphide: This reaction can be important in remote areas in generating sulfuric acid. Initial Hydrolysis step would be:

OCS(aq) + H2O ———> CO2(aq) + H2S(aq)

This would be followed by the oxidation of hydrogen sulphide through to sulfuric acid.

5. Reactions of photochemically generated species: Hydroxy and hydroperoxy radicals are produced photochemically in the atmosphere and these radicals are scavenged by the cloud droplets. These radicals can then promote various important reactions in the droplets. Such reactions have been discussed later.


Most of the particulate material suspended in the atmosphere has very small size and so has a very large surface area per unit mass (around 1 million square meter per gram). Such large surface area offers considerable opportunity for the absorption of molecules from the gas phase. This is particu­larly true if these molecules have a low volatility. A sub­stance having vapor pressure less than 10-6 Pa at ambient temperature will largely be adsorbed on the aerosol particles. Therefore, metals volatilized through volcanic or biological processes will probably end up at­tached to aerosols. The likelihood of surface reactions also increased by the large surface to volume ratio of aero­sols. Generally, two types of reactions occur on aerosol: thermal reactions and photochemical reactions.

Thermal reactions: For describing thermal reactions on aerosol surfaces, following two surfaces have been common models of atmospheric aerosols:

(i) Sulfuric acid surface: Sulfuric acid is a liquid surface but acid covers the surface of many atmospheric aerosol particles so this is a good model. The effectiveness of sulfuric acid surfaces as sink has been investigated for a number of atmospheric trace gases. The effectiveness of surface may be measured in terms of the probability of reactions occurring on collision of the molecules of the gas with the surface. Such probabilities for some major atmospheric trace gases are given in Table- 3.

Table- 3. Probabilities of reactions on collision of gas molecules with surface.


Molecule Probability


Water vapor 2 x 10^3

Ammonia >1 x 10^3

Hydrogen peroxide 7.8 x 10^4

Nitric acid 2.4 x 10^4


For species like nitric acid or hydrogen peroxide, the absorption of the gas by sulfuric acid surfaces could be a sink of atmospheric gases as much important as the photolysis.

(ii) Graphite carbon surface: Absorption of gases by graphite carbon is well known. A gas like sulfur dioxide is readily absorbed and presumably oxidized on the surface. However, aerosol surface soon becomes saturated or poisoned. Absorption of gas molecules can not occur further unless there is some mechanism for ‘cleaning’ the surface. Thus it is diffi­cult to visualize the mechanism of the removal of large amounts of a gas like sulfur dioxide from atmosphere by such a heterogeneous solid phase process.

Photochemical reactions: In addition to possibility of ther­mal reactions on particle surface subsequent to the absorp­tion of the gas molecules, photochemical reactions are also possible. For example,

2CO + O2 —————-> 2CO2
TiO2, ZnO

2N2 + 6H2O ————-> 4NH3 + 3O2

The importance of these reactions in the atmosphere is not known. However, it is known that photo-assisted reactions on titanium oxide or zinc oxide desert sands lead to production of ammonia. It has been postulated that such reactions were the source of ammonia in the early atmosphere of Earth.


In case of carbon dioxide and sulfur dioxide gases, the K” is very much smaller than K’ and may be neglected at acidic pH values. Thus the pH of a droplet of water in equilibrium with atmospheric carbon dioxide can be determined by combining two equilibrium constant equations, one governing the dissolution and the other the first step in the dissociation. The concentration of HCO3- will be:

[HCO3-] = KH K’ pCO2 / [H+]

If dissociation of carbon dioxide is the only source of hydrogen ions in the system, then [HCO3-] = [H+] so that:

[HCO3-] = [H+] = (KHK’ pCO2)0.5

By substituting appropriate values of equilibrium constants and use of a carbon dioxide partial pressure of 3.4 x 10-4 atm, a hydrogen ion concentration of 2.3 x 10-6 mol l-1 or a pH of 5.6 will be obtained. In remote regions, pH of pure rainwater may be close to this value and been assumed to be the pH of normal rainwater. However, trace amounts of other compounds can affect the rainwater pH. For example, sulfur dioxide concentration of having partial pressure of 5 x 10-9 atm in air will give an equation analogous to the one used for CO2:

[HSO3-] = (KH K’ pSO2)0.5

Use of appropriate values will give a rainwater pH of 4.6. Thus even at low concentrations sulfur dioxide has profound effect on pH despite carbon dioxide being present in much higher concentration. High solubility of sulfur dioxide gas and high dissociation constants for it make it more effective at acidifying water droplets than carbon dioxide. Its oxidation to sulfuric acid is comparatively easy and yields a further proton. Therefore, this gas has even more dramatic effect on rainwater pH.

Equations describing the dissolution of an acidic gas such as sulfur dioxide show that presence of acids already in solution will depress the solubility of the acidic gas while alkalis will enhance it. Ammonia is common alkaline gas in atmosphere and it will neutralize the dissolved acids, particularly the sulfuric acid. This means that ammonium salts, particularly ammonium sulphate, that are present in the atmosphere also effect the rainwater pH. In a system of water in equilibrium with CO2, SO2 and NH3 gases at the same time with pCO2 = 3.4 x 10-4 atm, pSO2 = 5 x 10-9 atm and pNH3 = 10-9 atm, pH can be calculated as for a single gas. But this calculation is a little more complex because the solution is not particularly acidic and second dissociation constant of sulfurous acid becomes important. The value of pH obtained is about 5.8 showing that ammonia even at very low concentrations has an effect on pH.


Many key reactions in the atmosphere are photochemical reactions which are initiated by absorption of a photon of light. Such reactions can be written as if they were normal chemical reactions by substituting photon (hv) as one of the reactants:

NO2 + hv —— NO + O

The rate constant of such a reaction is given as:

– {d[NO2]/dt} = k” [hv][NO2]

This expression is not very useful because the second-order rate constant k” would probably vary dramatically with the energy of photon involved in the reaction. However, by taking a content flux of photon with respect to wavelength and incorporating it into a pseudo first-order rate constant, the rate expression becomes:

d[NO2]/dt = J[NO2]

J is special first-order constant that embraces the absorption coefficient of the reactant, quantum efficiency of the reaction in question and the solar spectrum and intensity at the altitude and latitude under consideration. Estimates of J for many atmospheric trace gases can be made with a little information on the photochemistry. For example, a typical mid-latitude mid-day value of JNO2 , for the photodissociation of nitrogen dioxide is 5 x 10-3 s-1 which suggests a residence time of 200 s.

Many photochemical are important in the atmosphere as they yield atoms or free radicals and these species are greatly more reactive than the molecular species found in the air. For example, photodissociation of NO2 yields atomic oxygen which can subsequently lead to the formation of ozone:

O + O2 + M ——-> O3 + M

where M is a third body i.e. a molecule such as molecular nitrogen which carries off the excess energy that might disrupt the ozone molecule. The ozone thus produced might further be photodissociated:

O3 + hv ——–> O(3P) or O(1D) + O2

If wavelength of photon is less than 315 nm, the oxygen atom is produced in excited 1D state, otherwise in the 3P ground state. The ground state oxygen will probable recombine with a molecule of oxygen to for ozone again i.e. no net reaction would occur. The excited oxygen atom may be collisionally de-excited to ground state, or more importantly, may react with water molecule providing a source of hydroxyl radical (OH):

O(1D) + H2O ——–> 2OH

Hydroxyl and hydroperoxy radicals in atmosphere

Hydroxyl radical (OH) produced by reaction of excited oxygen atom (formed by photodissociation of atmospheric ozone) with water as described above is probably the most important radical in the chemistry of troposphere. A number of reactions in the troposphere involving hydroxyl radical can produce hydrogen atom or hydroperoxy radical:

OH + CO —–> CO2 + H

OH + O ——> O2 + H

OH + O + M ——> M + HO2

OH + O3 ——> O2 + HO2

H + O2 + M ——> M + HO2

Very quickly a range of radicals and atoms can be generated. These highly reactive species are basic to the gas-phase chemistry of atmosphere. Due to their high reactivity, these species are naturally found in very low concentrations in the atmosphere. Their typical background concentrations are:

Hydroxyl radical – 7 x 105 cm-3; Hydroperoxy radical – 2 x 107 cm-3

High reactivity of these radicals is indicated by their short residence times. The residence time of OH radical is less than 1 s while that of HO2 radical is perhaps 1 minute.

Reactions of hydroxyl and hydroperoxy radicals with atmospheric trace gases

Reactions of many trace gases found in the atmosphere with the hydroxyl radical exert a profound effect on the composition of atmosphere. Reactions with some of the atmospheric trace gases are discussed below.


Biologically produced sulfur gases are emitted into the atmosphere mainly as sulphides such as dimethyl sulphide, hydrogen sulphide and carbon disulphide. All these react with hydroxyl radical in the atmosphere.

(a) Dimethyl sulphide: This is the major sulphide emission into the atmosphere which reacts with OH radical as follows:

CH3SCH3 + OH —–> CH3SOH + CH3

O3 + CH3SOH —–> CH3SO3H

The product of these reactions is methyl sulphonic acid (CH3SO3H) and most of it persists in the ambient atmosphere though a relatively small amount may be oxidized through sulfur dioxide.

(b) Hydrogen sulphide: This gas in atmosphere is also attacked by OH radical as follows:

H2S + OH ——-> HS + H2O

The resulting bisulphide radical (HS) is oxidized through SO2 in a number of subsequent reactions. The SO2 can also be oxidized by OH and HO2 radicals:

SO2 + OH + M —–> HSO3 + M

SO2 + HO2 —–> SO3 + OH

Bisulfite radical (HSO3) and SO3 react with OH and water respectively to yield sulfuric acid which is the ultimate product of oxidation of atmospheric sulfur.

HSO3 + OH ———> H2SO4

SO3 + H2O ———-> H2SO4

(c) Carbon disulphide (CS2): This has been experimentally shown to be oxidized by OH radical yielding equal proportions of carbonyl sulphide and sulfur dioxide as final products. However, in atmosphere CS2 may not react with OH radical principally and its reactions with oxygen atoms may be more important.


Most of the atmospheric ammonia is removed through dissolution in liquid water in the atmosphere. However, ammonia is also attacked by OH radical though this reaction accounts for only a few percent of the ammonia removed from Earth’s atmosphere:

NH3 + OH ——> NH2 + H2O

Various subsequent reactions are possible:

NH2 + O ——> HNO + H

HNO + O2 —–> NO + HO2

The NO can be oxidized to NO2 which subsequently may react with OH radical to yield HNO3 and this is effectively removed from atmosphere through dissolution in rainwater.

NO2 + OH ——> HNO3


Hydroxyl radical on reaction with carbon monoxide yields carbon dioxide and hydrogen radical.

CO + OH ———-> CO2 + H


Formaldehyde found in trace quantities and formed in various atmospheric reactions is oxidized by OH radical in the following manner:

HCHO + OH ——–> HCO + H2O


Methane is naturally emitted from earth’s surface. In the atmosphere, methane is oxidized by OH radical yielding methyl radical and water:

CH4 + OH ——–> CH3 + H2O

CH3 undergoes following reactions in the methane cycle in the atmosphere yielding CH3O2.

CH3 + O2 + M ——-> CH3O2 + M

CH3O2 reacts with hydroperoxy radical in the following manner:

CH3O2 + HO2 ——–> CH3COOH + O2


In the presence of some suitable molecular species (M), hydroxyl radicals may react with each other to for hydrogen peroxide:

OH + OH + M —— H2O2 + M

Hydroperoxy radical may be a more efficient route for the formation of hydrogen peroxide:

HO2 + HO2 + M —— H2O2 + M

H2O2 is highly water-soluble and a strong oxidizing agent so it probably plays an important role in oxidation processes within water droplets in the atmosphere.


Presence of water and enough light in the clouds may result in the formation of hydroxy and hydroperoxy radicals there. These radicals shall be scavanged by cloud droplets and then could promote a variety of reactions in the droplets.


(a) Oxidation of inorganic species such as ammonia:

NH3(aq) + OH(aq) —— NH2(aq) + H2O

NH2(aq) + O2(aq) —– NH2O2(aq)

NH2O2(aq) + OH(aq) —— HNO2(aq) + H2O

(b) Oxidation of nitrogen oxides:

NO-2(aq) + OH(aq) —- NO2(aq) + OH-(aq)

NO(aq) + OH(aq) —— HNO2(aq)

NO2(aq) + OH(aq) —— HNO3(aq)

(c) Oxidation of sulfur compounds:

H2S(aq) + OH(aq) —— HS(aq) + H2O


Scavenged HO2 radicals have a longer lifetime than OH radicals in the water droplets, so they may be at much higher concentrations there and, therefore, could be important in reactions such as:

(a) Oxidation of SO32-

HO2(aq) + SO2-3(aq) —— SO2-4(aq) + OH(aq)

(b) Generation of Hydrogen peroxide:

HO2(aq) —— H+(aq) + O-2(aq)

O-2(aq) + HO2(aq) —– HO-2(aq) + O2(aq)

H+(aq) + HO-2(aq) —– H2O2(aq)


The possibility of radical chemistry opens up a whole range of organic reaction chemistry also, in particular the oxidation of naturally occurring alcohols and aldehydes e.g.

CH3OH(aq) –oxidant—- HCHO(aq) + H2O

HCHO(aq) + H2O ——- H2C(OH)2(aq)

H2C(OH)2(aq) + OH(aq) ——- HC(OH)2(aq) + H2O

HC(OH)2(aq) + O2(aq) ——- HO2(aq) + HCOOH(aq)

Formic acid, acetic acid and oxalic acid have been detected in the rainwater and point to the possibility of detection of a wide range of dissolved organic substances. These may indicate a complex radical-initiated chemistry that has an important effect on the acidification of rainwater.


Ionosphere is the conducting layer at an altitude of about 80 km and above. This zone of atmosphere was initially probed by radio-waves from ground and later by radio-sounders carried by rockets or direct measurements of gaseous components. Salient features of the chemistry of ionosphere are discussed below.

1. Ionosphere can be differentiated into various layers which represent zones of different electron densities. As a whole, ionosphere is electrically neutral since it also has positive ions like O2+, O+ and NO+. The positive ion chemistry is highly distinctive for various layers of ionosphere.

2. Ionosphere structure shows diurnal and long-term changes. Most important long-term changes correspond to solar sunspot cycle. Changes affect reflection of radio-waves and also alter the concentrations of various species in upper atmosphere.

3. Electrons in ionosphere are produced by photo-ionization. Above the altitude 100 km, this photo-ionization is brought about largely by extreme ultra-violet radiation. At lower altitudes, Lyman-A radiation is important. Some contribution to photo-ionization at somewhat lower altitudes is also made by cosmic rays. However, due to magnetic shielding of Earth, cosmic radiation is only important at fairly high latitudes. Night-time ionization is attributed to a downward flux of protons and radiation from excited species in the upper atmosphere (i.e. UV night glow).

4. In D region of ionosphere, electrons are produced principally by photo-ionization of nitrous oxide because it has lowest ionization potential among dominant species in the atmosphere. However, NO+ is not the most abundant positively charged species in the upper atmosphere. At altitude about 80 km, principal ion is a water cluster or hydrated proton i.e. H+(H2O)2. The charge initially carried by NO+ is transferred to water via an O2+ intermediate.

5. Production of electrons and ions is balanced by loss processes in a quasi-steady-state ionosphere. Loss processes usually involve reduction of photo-electron to thermal energies followed by ion-electron recombination or electron attachment. Typical processes are:

NO+ + e- ——> N + O (dissociative recombination)

O+ + e- ——–> O + hv (radiative recombination)

O2 + O2 + e- ——> O2- + O2 (three-body attachment)

6. In F-layer of ionosphere, positive charge is largely carried by O+ while at lower levels, it is more likely to be present on NO+, O2+ and lower down in atmosphere, on hydrated proton.

7. Though hundreds of reactions are used in descriptions, positive ion chemistry is still poorly understood. D-region of ionosphere is particularly complex because of the presence of an extensive array of negatively charged poly-molecular hydrates of water.

8. E-region of ionosphere is interesting because it sometimes shows thin sporadic layers that appear to be derived from metal-ion chemistry in mid-latitudes. Intensities of these layers show significant increases in response to meteor showers so it is possible that metal ions have extraterrestrial origin. Typical reactions are:

Mg + hv ——-> Mg+ + e-

Mg+ + O2 + M ——-> MgO2+ + M

Mg+ + O3 ——–> MgO+ + O2

The first reaction produces electrons but subsequently they react with charged metal and metal oxide species.

9. In the ionosphere, O+ ions are normally removed through reaction with oxygen and nitrogen:

O+ + O2 ——> O2+ + O

O+ + N2 ——> N2+ + O

But reactions involving hydrogen or water are about 1000 times faster. This leads to considerable reduction in concentration of electrons through following reactions:

O+ + H2O —–> H2O+ + O

O+ + H2 ——> OH+ + H

followed by:

e- + H2O+ —–> H2 + O

e- + H2O+ ——> OH + H

e- + OH+ ——-> O + H

10. Human activities can also affect the ionosphere chemistry. For example, at first launch of Skylab, a large booster operated in upper portion of ionosphere (at altitude 190 km). During the portion of flight through ionosphere, some 1.2 x 1031 molecules of water and hydrogen were released were released due to which electron densities were lowered over a radius of 1000 km around the flight path of the rocket thus creating an electron-hole.


Methane is emitted from the earth’s surface mainly due the activity of methanogenic bacteria. Mean rate of its emission is 2 x 1011 cm-2 s-1. It undergoes a complex series of reactions which together constitute the methane cycle in the atmosphere. The cycle may be divided into three main parts:

1. Oxidation of methane and formation of formaldehyde;

2. Oxidation (removal) of formaldehyde and formation of carbon monoxide;

3. Oxidation (removal) of carbon monoxide and formation of carbon dioxide.

Oxidation in various reactions of these three parts of methane cycle in achieved by reaction with OH, O2, or by photochemical oxidation. Reduction at places in the cycle is achieved by reaction with NO and HO2.

Oxidation of methane and formation of formaldehyde

Methane in the atmosphere is first attacked by hydroxyl radical yielding methyl radical and water. Methyl radical, through various oxidation and reduction reactions in which methyl peroxide (CH3O2), methyl hydroperoxy (CH3OOH), methyl oxide (CH3O) are formed, finally yields formaldehyde. In the sequence of these reactions, OH and HO2 radicals used are again formed. The reactions involved in this part of methane cycle are:

1. CH4 + OH ——–> CH3 + H2O –

2. CH3 + O2 + M ——> CH3O2 + M

(M is some molecule acting catalytically and carrying off the excess energy of reaction)

3. CH3O2 + NO ——–> CH3O + NO2

3A. CH3O2 + HO2 ——–> CH3OOH + O2

3B. CH3OOH + hv ———–> CH3O + OH

4. CH30 + O2 ——-> HCHO + HO2

Oxidation of formaldehyde and formation of carbon monoxide

Formaldehyde formed ultimately in the above part of methane cycle is removed by photochemical or chemical oxidation reactions in this second part of methane cycle. A very small part of formaldehyde may be removed from atmosphere through dissolution in rainwater. Ultimate product of chemical removal of formaldehyde is carbon monoxide.

5. HCHO + OH ——–> HCO + H2O

5A. HCHO + hv ——–> HCO + H

6. HCHO + hv ———> CO + H2

7. HCO + O2 ———-> CO + HO2

Oxidation of carbon monoxide and formation of carbon dioxide

Carbon monoxide formed in the second part of methane cycle is finally oxidized by reaction with hydroxyl radical to yield carbon dioxide and hydrogen atom.

8. CO + OH ——–> CO2 + H

The notion of continuity helps in understanding the transfer of material along various reaction pathways in the complex set of reactions of methane cycle given above. Formaldehyde formed by oxidation of methane in atmosphere is removed by four possible processes (reaction numbers 5, 5A, 6 and rainout). The notion of continuity requires that the sum of fluxes through these four pathways is equal to the production rate. From the reactions given above the destruction of formaldehyde can be equated with the production of methane at the surface of Earth or with destruction of methane in the atmosphere. Thus may be written as:

– {d[CH4}/dt} = – {d[HCHO]/dt}

Since washing out with rain (rainout) is very insignificant, equation can be rewritten as:

k1 [OH][CH4] = k5 [HCHO][OH] + J5A [HCHO] + J6 [HCHO]

(subscript numbers refer to reaction numbers given above)

The atmospheric concentration of methane is 1.6 ppm or 4.2 x 1013 cm-3 and of hydroxyl radicals is about 7 x 105 cm-3. Taking the rate constant k1 = 8 x 10-15 cm3 s-1, the destruction rate of methane can be estimated as 2.3 x 105 cm-3 s-1. Furthermore, the above equation can be rearranged as:

[HCHO] = k1 [OH][CH4] / {k5 [OH] + J5A + J6}

This gives the estimate of formaldehyde as 4.3 x 109cm-3 (where k5 = 1.3 x 10-11 cm3 s-1 and J5+J6 = about 4.5 x 10-5 s-1). This estimate of the concentration is a little low but not too far from the typical value of 1010 cm-3 that is observed in the atmosphere.

Notion of continuity can be applied to the formation of carbon monoxide from the oxidation of methane. Carbon monoxide in atmosphere may arise from various sources but the magnitude of natural sources of production of CO can be easily assessed. If small loss of formaldehyde and possibly of methyl hydroxyperoxide (CH3OOH) due to rainout is neglected then CO should be formed at the same rate as methane is released into the atmosphere i.e. at

2 x 1011 cm-2 s-1 or about 0.7 x 1015 g (C) a-1. This is larger than the amount which arises from human activities (0.3 x 1015 g (C) a-1).


The ozone present in troposphere and stratosphere together constitute the total atmospheric ozone. The atmospheric ozone has important impact on the global climate system. The production and loss of ozone in both troposphere and stratosphere are strongly linked to atmospheric chemistry at both levels. Both areas of ozone are also influenced by four major processes that basically dominate the biogeochemical cycles in atmosphere:

1. Emissions from natural and anthropogenic sources

2. Chemical transformations and reactions

3. Atmospheric transport through circulation

4. Removal mechanisms

Ozone chemistry of troposphere

The troposheric ozone concentrations make up only 13% of total ozone in atmosphere yet ozone of this zone has major impact on climatic change through its effect on global warming. Natural background ozone concentrations can only be found in atmospheres of rural and remote areas while over urban centres, unnatural ozone concentrations are created by anthropogenic emissions of various substances that have profound effect on ozone chemistry. In unperturbed troposphere, the formation and destruction of ozone are part of a dynamic balance controlled mainly through ozone sources from marine and terrestrial biospheres and sinks atmospheric photochemistry and surface depositions. Anthropogenic emissions entering this system change the balance both spatially and temporally and such changes can be transferred globally by atmospheric transport mechanisms.

Naturally the tropospheric ozone is a secondary constituent originating from two main sources:

1. In upper troposphere, major source is transport of ozone from stratosphere

2. In middle and lower troposphere, photochemical mechanisms of ozone production

The concentration of ozone at any level in troposphere is determined mainly by photochemical mechanisms of its formation and destruction. Photochemistry dominates the ozone cycle particularly in middle and lower troposphere atleast for three reasons:

(a) Presently calculated rate of loss of ozone are about four times higher than the rate which would have occurred if tropospheric ozone originated completely in stratosphere

(b) Measured increase in ozone over urban areas can only be photochemical in origin

(c) Larger concentration of ozone in Northern Hemisphere than in Southern Hemisphere despite larger land surface sink can only be attributed to atmospheric photochemical reactions.

Recent estimates show that maximum ozone produced per year in troposphere is about 6.5 x 1011 molecules cm-1 s-1. Higher concentrations of ozone occur in mid-latitudes of Northern Hemisphere because of higher number of precurssor sources there. Minimum ozone concentrations occur in equatorial regions around 100 S caused partially by stronger photochemical destruction in the tropics and partially by background ocean conditions in Southern Hemisphere. Average ozone concentrations in free troposphere are 39 ppbv in Northern Hemisphere and 24 ppbv in Southern Hemisphere. Representative latitudinal ozone concentrations in free troposphere are 30-40 ppbv in 30-600 S, 20 ppbv in 0-300 S, 20-30 ppbv in 0-200 N and 30-50 ppbv in 20-600 N.

Vertical distribution of ozone differs between hemispheres and with distribution of important chemical precursors, particularly CO. In Northern Hemisphere, on average the ozone concentration increases slightly with altitude and boundary layer ozone concentrations are about 1.1 to 1.5 lower than the free troposphere. In Southern Hemisphere, there is little variation in ozone concentration with altitude and ozone in boundary layer does not decrease significantly compared to free troposphere.

The formation and destruction of ozone in troposphere depends heavily on the OH radical concentration and associated reaction efficiency. The process is initiated by photodissociation of ozone by sunlight and the formation of OH from water and oxygen:

O3 + hv (300-330 nm) ——-> O2 + O(1D)

O(1D) + H2O ——> OH + OH ( R = 2.3 x 10-10)

OH formation depends on water in troposphere. As a rough estimate, H2O0.5-1.0 approximates OH concentrations. Since Oh is a highly reactive radical, it is very short-lived in troposphere. Its concentrations sow diurnal variations, particularly in higher latitudes linked to solar-energy variations. At night, OH concentrations are supposed to fall by two orders of magnitude as compared to daytime with minimum concentrations about 105 molecules cm-3 and maximum concentrations near mid-noon about 107 molecules cm-3.

In the troposphere, apart from OH radical other critical species for basic gas-phase reactions are nitrogen oxides (NO, NO2, Nox), free hydrogen/oxygen radicals (OH, HO2), methane and non-methane hydrocarbons (designated by general term RO2 and carbon monoxide. These processes are strongly linked to one another and depend heavily in the concentrations of the relevant molecules in the atmosphere. These reactions in troposhpere have been described in detail in the discussion of photochemical smog problem. However, important features of main molecules affecting tropospheric ozone may be summarized as following:

1. Nitrogen oxides: Nitrogen gases help control OH concentrations in troposphere and concentrations of NO and NO2 are needed to form ozone. Since both molecules are active in ozone process, they are described by their conserved quantity, NOx. The rate of ozone production in troposphere seems to be controlled by NOx concentration. NOx acts as catalyst to photochemical reaction processes and provides the environment which allows further ozone formation or loss reactions in various chains. For example, in NOx-poor environment, oxidation of one methane molecule to carbon dioxide via CO results in net loss of about 3.5 H atoms and 1.7 ozone molecules. In NOx-rich environments, the same process will create about 0.5 H atoms and 3.7 ozone molecules. The transfer point between ozone loss and ozone production seems to an NO concentration of about 30.0 pptv. The efficiency of NOx in ozone-formation processes decreases with increasing NOx concentration. However, in terms of total production of ozone, this inefficiency is overcome in the atmosphere with higher NOx.

Main sink of NOx in atmosphere is conversion to nitric acid by OH. This sink acts within a time frame of 1 to 2 days and nitric acid during this time is either washed out of atmosphere or is removed by surface deposition. Another mechanism associated with lifetime of NO2 is the day-night cycle of its release and capture associated with N2O5. During night, NO2 and nitrate radical (NO3) combine in presence of some catalyst to form N2O5 which acts as a strong reservoir. During daytime, sunlight reverses the process and NO2 is released.

Associated with NOx and its impact on ozone are RO2 reactions which can lead to a wide variety of complex non-methane hydrocarbon reactions. Most well known byproduct of this process is PAN (peroxyacetyl nitrate) which acts as a reservoir for NOx in clean marine air. Its free tropospheric values tend to be in the 25-35 pptv range.

2. Carbon monoxide: In natural atmosphere, CO is created as byproduct of reactions sequence of oxidation of methane during photodissociation of HCHO. There is strong correlation between concentrations of CO and methane in troposphere. Average concentrations of CO are on the order of 30-200 ppbv and its lifetimes are relatively short (about 1-2 months) due mainly to reactions with OH. CO and ozone show positive relationship in areas of higher NOx where ozone is being created photochemically. However, in areas of ozone destruction, where NOx concentrations are less than 0.01 ppbv, CO concentrations are independent of ozone.

3. Methane and Non-methane hydrocarbons (MHC & NMHC): Methane is the most important and most abundant atmospheric hydrocarbon. Its lifetime in troposphere is about 5-10 years. Major sink of methane is its reaction with OH leading to the formation of ozone. Another sink is its gradual transfer to stratosphere through exchange processes across tropopause. Methane then acts as an important factor in strotospheric chemistry.

Non-methane hydrocarbons (NMHC) in atmosphere may also contribute to the formation of ozone. However, which species of NMHCs are important and in what amounts is yet not well established.

Major molecules associated with tropospheric ozone chemistry and their energy requirements are listed in Table- 4.

Table- 4. Energy requirements of some major molecules associated with tropospheric ozone chemistry.

Chemical species Enthalpy of formation Free-energy of reaction*

O(3P) 59.6 55.4

O(1P) 104.8

O3 34.1 39.0

OH 9.3 8.2

HO2 ~3.4 4.4

H2O2 -32.6 -25.2

H2O -57.8 -54.6

N2O 21.6 20.7

NO2 7.9 12.3

CH4 -17.9 -12.1

CO -26.4 -32.8

+RO2 var. var.

* energy needed to create or destroy chemical bonds. Positive numbers indicate energy must be added to create formation reaction.

+ Complex organic peroxy radicals

Ozone chemistry of stratosphere

Most of the ozone in the atmosphere forms the Ozone layer in the stratosphere at altitudes between 10 and 40 km (100 to 0.1 mb pressure altitude) depending on latitude, just above the tropopause. This layer is crucial for life because only ozone absorbs UV-B radiation between 280-320 nm. UV-A rays between 320 and 400 nm are not affected by ozone while UV-C rays between 200 and 280 nm are absorbed by other atmospheric constituents also beside ozone.

Stratospheric ozone distributions are strongly dependent on stratospheric circulation patterns, varying according to latitude, seasons, short-term meteorological changes and the photochemical processes of formation and destruction. Major driving forces are availability of sunlight and thus of UV radiation and in upper stratosphere (above pressure altitude of 5 mb) the latitudinal temperature gradient which assists ozone transport. The ozone content of stratosphere is highly dynamic and variable. Its concentrations peak around the altitude of 30 km in tropics and around 15 to 20 km in polar regions.

Though hundreds of reactions are known to be involved in the ozone chemistry of stratosphere, only a few can be described properly. The ozone chemistry basically involves two types of reactions: those involved with ozone formation and those involved with ozone destruction. These two types of reactions are important because relationship between stratospheric ozone and climate has been studied particularly in association with ozone depletion and ultra-violet radiation. Another important feature is that above tropopause, liquid water does not play significant role and stratospheric ozone chemistry here is dominated by photochemical reactions.

1. Ozone formation: This itself is a photochemical process involving UV radiation of wavelength less than 242 nm. Though photodissociation of oxygen by UV radiation at less than 175 nm may yield an oxygen atom in excited state i.e. O(1D), such photodissociation is important only in the upper stratosphere because such short wavelength can not penetrate lower into stratosphere.

Thus in upper stratosphere reaction may be:

O2 + hv ( O(3P) + O(1D)

Oxygen atom in excited state on collision with some diatomic molecule (M2) yields oxygen atom in ground state i.e. O(3P):


O(1D) + M2 —————–> O(3P) + M2


while in lower stratosphere reaction is:

O2 + hv (175-242 nm) ———> O(3P) + O(3P)

The oxygen atoms in ground state react with diatomic oxygen molecules to form ozone:

O(3P) + O2 ———> O3

2. Ozone destruction: This involves those reactions which balance the photochemical formation of ozone in stratosphere:

O3 + hv ——> O2 + O(1D)

O3 + O ——-> 2O2

Another additional reaction for removal for oxygen atoms is:

O + O + M ——–> O2 + M

Many analogous reactions involving H, N and Cl radicals also occur in stratosphere:

OH + O3 ———> O2 + HO2

HO2 + O ———> OH + O2

NO + O3 ——–> O2 + NO2

NO2 + O ——–> NO + O2

O3 + Cl ——> O2 + ClO

ClO + O ——> O2 + Cl

All the above pairs of reactions are summed as:

O3 + O —–> 2O2

i.e. each pair of reactions involves destruction of ozone and atomic oxygen while restoring the OH, NO or Cl radical.


N2O which is relatively stable in the troposphere, usually moves into stratosphere and undergoes following photochemical reactions:

N2O + hv ( N2 + O(1D)

N2O + O(1D) ———-> N2 + O2

N2O + O(1D) ———-> 2NO

NO or NO2 which may also move from troposhpere into stratosphere or are produced in the stratosphere, undergo following reactions there:

NO + O + M ——> NO2 + M

NO2 + OH + M —–> HNO3 + M

HNO3 + hv ( OH + NO2

Ozone layer in stratosphere absorbs sufficient amount of UV radiation so that at tropopause HNO3 has photochemical lifetime of about 10 days. This time is long enough for much of it to cross the tropopause and come down with rainfall thus being removed from stratosphere.


Chief natural source of chlorine in atmosphere is probably methyl chloride from marine algae but it accounts for only 25% of the chlorine currently being transported across tropopause into the stratosphere. Other natural sources adding minor amounts are HCl acid from volcanoes and chlorine from sea sprays. In the past few decades, chlorofluorocarbons (mostly CFCl3 i.e. Feron-11) and CF2Cl2 i.e. Feron-12) added to the atmosphere by human activities have become chief source of stratospheric chlorine. Ammonium perchlorate-aluminium solid rocket propellents are another anthropogenic source of atmospheric chlorine. These compounds absorb UV radiation in the range of 190 to 220 nm resulting in their photodissociation:

CH3Cl + hv —–> CH3 + Cl

CFCl3 + hv ——> CFCl2 + Cl

CF2Cl2 + hv —–> CF2Cl + Cl

Free Cl atoms in stratosphere may undergo various reaction cycles:

1. Reaction with ozone: Free chlorine atoms in stratosphere react with ozone in catalytic manner and cause depletion of ozone:

Cl + O3 —–> ClO + O2

ClO produced may react with nitrogen compounds:

ClO + NO —–> Cl + NO2

ClO + NO2 + M ——–> ClNO3 + M

ClNO3 may be decomposed by UV radiation or by reaction with atomic oxygen:

ClNO3 + hv —–> ClO + NO2

ClNO3 + O —-> O2 + ClO + NO

Reactions of ClO with NO or NO2 are important because they effectively remove N- and Cl- species involved in ozone destroying cycles.

2. Reaction with methane and hydrogen: Free Cl may also react with CH4 or H2:

Cl + CH4 —–> HCl + CH3

Cl + H2 ——> HCl + H

Some of the HCl may react with OH radical in stratosphere:

OH + HCl —–> H2O + Cl

However, most of the HCl moves down to tropopause and is removed with rainfall as HCl acid.


In atmospheres of urban centres, under conditions of relatively low humidity, plenty of sunshine, a large amount of exhaust emissions from motor vehicles and moderate to low wind speeds, photochemical processes lead to a secondary pollution situation commonly termed “photochemical smog’. A large number of compounds and reactions have been characterized in the urban air where such smog situation occurs. The chemistry of this photochemical smog condition is extremely complex. Major photochemical processes associated with this condition have been discussed below.

1. Nitrogen oxide pseudo-equilibrium

The oxides of nitrogen, particularly NO and NO2 are at the root of photochemical smog problem. Oxidation of atmospheric nitrogen during high temperature combustion processes (particularly in motor vehicles) results in formation of NO which is further oxidized to NO2:

(a) O + N2 —-> NO + N

N + O2 —-> NO + O


N2 + O2 —-> 2NO

(b) 2NO + O2 ——> 2NO2

R1 = k1[NO]2[O2] where R1 and R2 are reaction rates and k1 and k2 are the rate

and, constants

NO + O3 —-> NO2 + O2

R2 = k2[NO][ O3]

The reaction of NO with oxygen at the concentrations found even in the polluted air is very slow, therefore, NO2 is mainly produced by oxidation of NO by ozone. In the polluted air, typical early morning concentrations of ozone and NO are 40 ppb and 80 ppb respectively. Values for k1 and k2 are 1.93 x 10-38 cm6s-1 and 1.8 x 10-14 cm3s-1 for ozone and NO respectively. From these values the calculation shows that R1 = 4.6 x 10-5 cm-3 s-1 and R2 = 3.8 x 1010 cm-3 s-1. This confirms the far greater importance of the oxidation of No by ozone.

The NO2 produced in this way can be photodissociated back to NO. Thus a sequence of reactions describing its destruction and regeneration can be given:

NO2 + hv ( O(3P) + NO

O(3P) + O2 + M ——> O3 + M

O3 + hv (300-330 nm) ——> O2 + O(1D)

O3 + HO2 ——> 2O2 + OH ( R = 1.1 x 10-14)

O3 + NO —–> NO2 + O2 (R = 2.3 x 10-12)

In a volume of air in steady-state where production and destruction rates of NO2 are equal and where oxidation of NO by oxygen is assumed to be unimportant, the reaction rate may be written as:

k2[NO][O3] = J[NO2]

where J is effective first-order rate constant for photodissociation. The equation may be rearranged as:

J/k2 = [NO][O3]/[NO2]

where the term on right-hand side may be ignored as a pseudo-equilibrium constant relating the partial pressures of NO, NO2 and O3. The value of J will varies with change in intensity of sunlight throughout the day. However, measurements have shown that overall the equality implied in this equation holds in the polluted atmosphere. During first half of the day radiation intensity increases which means J will increase and during this period increasing amounts of ozone and NO would be expected. Since both these are produced by destruction of NO2, the amount of ozone should approximately equal the amount of NO.

Measurements from polluted atmospheres show that neither of the above predictions are borne out. The level of NO rises in the early morning but level of ozone rises much more slowly. Further, the fact that the level of NO2 falls by mid-day is even more in contrast to the theoretical prediction. A possible explanation for these observations is that the observed rises and falls in the concentrations of pollutants are merely functions of the pattern of generation and dispersion in the atmosphere.

2. Role of organic molecules in smog

Under constant illumination the rise in the level of ozone indicates a decreasing NO:NO2 ratio in the pseudo-equilibrium. For the latter to happen, another source of oxidant is needed because above described sequence of reactions does not result in any overall production of ozone. However, ozone production in polluted atmosphere may be explained by following scheme.

As in unpolluted atmosphere, oxidation in polluted atmosphere also occurs through reactions in which hydroxyl radical plays a key role. The hydroxyl radical attacks a veriety of pollutants in the urban air resulting in formation of free radicals like methyl radical (CH3), acetyl (CH3CO) and atomic hydrogen (H) which may become involved in subsequent reactions which oxidize to NO to NO2 and regenerate hydroxyl radical at the same time.

(a) Alkanes in smog: Presence of alkanes such as methane in polluted air provides a way in which NO can be oxidized to NO2 without consuming ozone. For example, methane may be oxidized by OH radical to produce methyl radical which further undergoes a series of reactions:

CH4 + OH —-> H2O + CH3 (R = 2.4 x 10-12)

CH3 + O2 +M —-> CH3O2 + M

CH3O2 + NO —-> CH3O + NO2 (R = 7.0 x 10-12)

CH3O + O2 —–> HCHO + HO2 (R = 5.0 x 10-13)

HCHO = hv ( 2 H + CO

CO + OH —–> CO2 + H (R = 1.35 x 10-13)

HO2 + NO —–> NO2 + OH (R = 4.3 x 10-12)

HO2 radical can also react photochemically or with ozone, atomic hydrogen or atomic oxygen to regenerate OH radical. HCHO can photodissociate into atomic hydrogen or react with oxygen to give the HO2 radical and CO.

The above reactions can be summed up and show the importance of methane in generating NO2 in photochemical smog:

CH4 + 2 O2 + 2 NO —-> H2O + HCHO + 2 NO2

This indicates net oxidation of NO in a manner that has not used ozone, therefore, it is different from pseudo-equilibrium situation.

(b) Aldehydes in smog: Aldehydes also provide effective ways of oxidizing NO to NO2. For example, acetaldehyde is attacked by OH radical producing acetyl radical which undergoes following subsequent reactions:

CH3CO + O2 —–> CH3COO2

CH3COO2 + NO —–> NO2 + CH3CO2

CH3CO2 ——> CH3 + CO2

Methyl radical produced is oxidized as described above. There are analogous reactions for higher aldehydes.

(C) Atomic hydrogen in smog: The atomic hydrogen produced by attack of OH on CO or photodissociation of HCHO can react with HO2 radical to produce two OH radicals that can initiate further attack on organic compounds in air.

OH + CO —–> CO2 + H

H + HO2 —–> OH + OH

Atomic hydrogen can also form HO2 radical which can oxidize NO to NO2:

H + O2 + M —–> HO2 + M

HO2 + NO —–> NO2 + OH

In general, hydrocarbons present in the polluted urban air promote the oxidation of NO to NO2 by reactions of the types described above. The NO2 is subsequently photolysed to produce NO for reoxidation and increasing amount of ozone.

NO2 + hv ( O(3P) + NO

O(3P) + O2 + M —–> O3 + M

Though there are losses in the above described scheme, the built-up of ozone throughout the day can thus be well explained.

3. Other products in photochemical smog

A number of other features of photochemical smog can also be explained by photochemical mechanism described above.

1. Formation of PAN: Peroxyacetylnitrate (CH3COO2NO2) or PAN is a major eye-irritant found characteristically in photochemical smog. The peroxyacetyl radical (CH3COO2) produced by attack of acetyl radical on oxygen can combine with NO2 to form PAN:

CH3COO2 + NO2 ——> CH3COO2NO2

PAN is the principal member of a group of rather similar nitrated compounds which includes higher peroxyalkyl compounds such as peroxypropionyl nitrate which has also been detected in low concentrations in photochemical smog. There is also much current interest in the natural production of compounds like PAN.

2. Formation of N2O5 : NO2 is oxidized by ozone to NO3 which subsequently reacts with NO2 to form N2O5. The NO3 may also react with NO to produce more NO2.

NO2 + O3 —-> NO3 + O2

NO3 + NO2 ——> N2O5

NO3 + NO ——-> 2NO2

3. Formation of nitric acid: The OH radical formed in the smog reacts with NO to form HNO2 and with NO2 to produce HNO3. There may be reaction between NO and NO2 to form HNO2.

NO + OH —-> HNO2

NO2 + OH + M—–> HNO3 + M

NO + NO2 + H2O ——> 2HNO2

HNO2 undergoes photodissociation to produce NO and provide a source of OH radicals.

HNO2 + hv ( NO + OH

4. Formation of hydrogen peroxide:
Formaldehyde in polluted air is an important source of atomic hydrogen and hence OH and HO2 radicals:

HCHO + hv ( 2H + CO

H + O2 + M ——> HO2 + M

HO2 + NO —–> NO2 + OH

The OH and HO2 radicals may produce H2O2:

OH + OH + M —–> H2O2 + M

HO2 + HO2 ——–> H2O2 + O2 (R = 3.8 x 10-14)

5. Oxidation of sulfur dioxide: Sulfur dioxide can be oxidized under photochemical conditions but the S-O bond is very strong. So the sulfur dioxide can not undergo photodissociation as in the familiar case of NO2. The oxidation of SO2 involves OH radical:

OH + SO2 ——> HSO3

HSO3 + O2 —–> HSO5 or,

HSO5 ——-> HO2 + SO3 HSO3 + O2 ——> HO2 + SO3

SO3 + H2O —–> H2SO4

There is increasing evidence that the two middle reactions occur as a single reaction.

4. Degradation of larger organic molecules

Larger organic molecules (other than methane and acetaldehyde) are also split up in photochemical smog.

(a) Alkanes: Degradation of large alkane molecules (e.g. butane) starts with attack by OH radical:

OH + CH3CH2CH2CH3 ——-> H2O + CH3CH2CH2CH2

O2 + CH3CH2CH2CH2 ——-> CH3CH2CH2CH2O2


CH3CH2CH2CH2O + O2 ——-> CH3CH2CH2CHO + HO2

CH3CH2CH2CHO + hv ——–> CH3CH2CH2 + HCHO

(b) Alkenes: Large alkane molecules may be degraded by being attacked by ozone, atomic oxygen (O(3P) or OH radical. Attack by OH radical predominates in polluted atmosphere. A typical reaction scheme may be illustrated using butane as example:




The process goes on and on.

The above reaction schemes show degradation of larger organic molecules into smaller ones resulting in greater predominance of low molecular weight compounds in typical urban atmosphere with exhaust fumes of automobile.

5. Heterogeneous reactions in photochemical smog

Gas-phase photochemical reactions may lead to formation of aerosols in polluted urban atmosphere and these give rise to visual obscurity associated with smog condition. High opacity of smog gives an exaggerated impression of the amount of particulate material present yet it is estimated that as little as 5% of pollutants present in photochemical smog could be converted into suspended particulate materials. Various heterogeneous reactions could occur on the surface of these particles or in cloud or rain droplets associated with smog. The material forming condensed phase of smog may consist of both inorganic and organic substances.

(i) Inorganic substances: These include metal oxides and the salts of acids produced within urban air. The acids (particularly sulfuric and nitric acids) are usually present in association with solid particles or more probably as droplets due to their high affinity for water. The latter can react rapidly with atmospheric ammonia. The ammonium sulphate and ammonium nitrate produced are important aerosols that are main causes for the reduction of visibility that accompanies photochemical smog.

(ii) Organic solids: Relatively little is known of the reaction pathways that produce organic particulate materials in the polluted urban air. Nitrogen has been detected in rather unusual reduced oxidation states on particles in photochemical smog. This nitrogen is thought to be present as nitriles, amines or amides bound onto the surface of soot particles. By denoting the soot surface as S, the process may be written as:

S-OH + NH3 —-> S-ONH4

(a phenolic hydroxy ammonium complex)

S-ONH4 —–> S-ONH2 + H2O (at higher temperature)



S-COONH4 —–> S-COONH2 + H2O (at higher temperature)

S-COONH2 ——> S-CN + H2O

Most thoroughly studied heterogeneous reaction in the atmosphere Is the oxidation of sulfur dioxide in atmospheric liquid droplets by the ozone, hydrogen peroxide or oxygen in the presence of a transition metal ion catalyst. This oxidation reaction has been discussed earlier and may proceed much faster in polluted urban atmospheres than in unpolluted atmospheres because the concentrations of oxidants (H2O2 or O3) and metal ion catalysts may be much higher. Metal ions may, in particular, be leached from particulates that are added into the air through anthropogenic activities. Leaching of metals from ash may be particularly significant in their surface concentrations being enriched. High amounts of soluble metal ions have been observed in association with fly ashes from the combustion of refuge derived fuels. A further mechanism for increasing the rate of oxidation involves dissolution of materials such as calcium oxide which are present in high concentrations in coal fly ash making the droplet alkaline:

CaO + H2O —–> Ca2+ + 2OH-

This allows dissolution of larger amounts of sulfur dioxide and thus increases the rate of catalytic oxidation. Alternatively, dissolution of ammonia from a polluted atmosphere will also increase the pH and enhance both the dissolution and oxidation of sulfur dioxide.

Oxidation of sulfur dioxide may also occur via absorption of gas onto solid surfaces followed by subsequent oxidation. However, the surface area of particulate material even in polluted atmosphere is quite small and, therefore, such mechanism requires some method of ‘cleaning’ the surface in order to make oxidation process significant. If particulates are wet, this mechanism may be effective since water would ‘clean’ the surface of particulate material.

In the atmosphere, changes in the size and/or composition of particles also occur. These include leaching of particulate material by water, oxidation or reduction of particles. Zinc vapour from copper smelters condenses to form highly angular and crystalline zinc oxide crystals in the atmosphere. These are gradually degraded, then rounded and now acquire a carbonaceous coating. Slowly zinc oxide core decomposes and particle ends up as a carbonaceous pseudomorph with little or no zinc. Possibly, carbonaceous particles are formed by reduction of zinc oxide following deposition of hydrocarbons onto the surface of the particle.


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Filed under: terrestrial vegetation — gargpk @ 1:47 pm
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A phytogeographical region is defined as an area of uniform climatic conditions and having a distinctly recognisable type of vegetation. According to D. Chattarjee (1962), India can be divided into nine phytogeographical regions.

  1. Western Himalayas

This region comprises north and south Kashmir, part of Punjab and Kumaon region of Uttaranchal. Average annual rainfall in the region is 100-200 cm. The region is wet in outer southern ranges and slightly dry in the inner areas. At high altitudes, snowfall occurs during winters. The region is subdivided into three zones.

  1. Submontane (lower, tropical and subtropical) zone: This zone includes outer Himalayas i.e. regions of Siwalik Hills and adjoining areas from 300 to 1500 m altitude. Average annual rainfall of the zone is around 100 cm. The vegetation consists of subtropical dry evergreen, subtropical pine and tropical moist deciduous forests.

  2. Temperate (montane) zone: This zone extends in the western Himalayas between the altitudes 1500 and 3500 m. The climate is wet between the altitudes 1500 and 1800 m and is drier at higher altitude. The vegetation consists of wet forests, Himalayan moist and Himalayan dry temperate forests.

  3. Alpine zone: This zone extends between 3500 m and 5000 m altitudes. The rainfall is very scanty and climate is very cool and dry. The vegetation consists of alpine forests.

  1. Eastern Himalayas

This region extends in the Himalyas from east of Nepal up to Arunachal. The climate is warmer and wetter than in western Himalayas. Tree line and snow line are higher by about 300 m than in the western Himalayas. The tropical temperature and rainfall conditions result in vegetation of the region having greater general species diversity, greater variety of oaks but lesser variety of conifers than in the western Himalayas. This region is also divided into three zones.

  1. Submontane (lower, tropical and subtropical) zone: This zone extends from the foothills up to the 1850 m altitude. The climate is nearly tropical and subtropical. The vegetation consists of subtropical broad-leaved forests, pine forests and wet temperate forests.

  2. Temperate (montane) zone: The zone extends from 1850 m to 4000 m altitude, about 500 m higher than in the western Himalayas. The vegetation consists of typical temperate forests with oaks and Rhododendron at lower and conifers at higher altitudes.

  3. Alpine zone: This zone extends from 4000-5000 m altitude. The climate is very cool and dry. The vegetation consists of alpine forests.

  1. Indus plain

This region comprises a part of Punjab, Delhi, Rajasthan, a part of Gujrat and Cutch. The climate has very dry and hot summers alternating with dry and cold winters. The annual rainfall is generally less than 70 cm and may be 10-15 cm in some areas. Most of the region is desert today though it had dense forests about 2000 years ago that were destroyed due to biotic factors particularly extensive cattle grazing. The vegetation today consists of tropical thorn forests and grasslands in some areas.

  1. Gangetic plain

This region covers part of Delhi, Uttar Pradesh, Bihar, West Bengal and part of Orissa. Average annual rainfall ranges from 50 cm to 150 cm from east to west. The vegetation consists of tropical moist deciduous forests, dry deciduous forests, thorn forests and mangrove forests.

  1. Assam

The region covers most of the Assam. The climate is characterized by very high temperature and rainfall. The vegetation consists of tropical evergreen and wet temperate forests in the lower plains while hilly tracts up to 1700 m altitude have subtropical pine forests.

  1. Central India

This region comprises part of Orissa, Madhya Pradesh, Vindhyan region and Gujrat. The areas are mostly hilly with some places at 500-700 m altitude. The average annual rainfall is 100-170 cm. Biotic disturbances are very common in this region resulting in degradation of forests into thorny forests in the open area. The vegetation consists of tropical moist deciduous forests, chiefly Sal forests in areas of annual rainfall above 150 cm and mixed deciduous forest in areas of 125-150 cm annual rainfall. Tropical thorn forests are found in the areas of annual rainfall below 125 cm.

  1. Western coast of Malabar

This is a small region extending from Gujrat to Kanyakumari along Western Ghats. The climate is warm humid having annual rainfall over 400 cm. The climate is tropical on the coasts and temperate in the hills. The vegetation consists of tropical wet evergreen, moist evergreen and moist deciduous forests. Wet temperate forests (Sholas) are present in Nilgiri while mangrove forests are found in the saline swamps on the coasts.

  1. Deccan

The region comprises southern Peninsular India from southern Madhya Pradesh up to Kanyakumari excluding the Western Ghats. The average annual rainfall in the region is about 100 cm. The vegetation consists of tropical dry evergreen, dry deciduous and swamp forests.

  1. Andman and Nicobar

This region includes Andman and Nicobar Islands. The climate of the region is warm and humid with very high temperature and annual rainfall. The vegetation consists of littoral mangrove, evergreen, semi-evergreen and deciduous forests.


Filed under: terrestrial vegetation — gargpk @ 1:45 pm
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India is situated at tropical latitudes and has diverse temperature and rainfall regimes. The overall climate of India is suitable for the growth of forests. The climax formations of Indian subcontinent have been altered much due to human activities in the last few thousand years. However, the remaining vegetation shows that the natural vegetation of India primarily consists of forests. The grasslands found in the region are not natural plant formations but have originated secondarily due to destruction of natural forests in some places. Therefore, these represent various stages of seral (successional) development due to the influence of a variety of biotic influences.

Source : Forest Survey of India, Dehradun. State of forest report 2001. Dehradun, FSI, 2002. 12p.


The most important factors influencing the physiognomy, species composition, phenology etc. of Indian forests are temperature, rainfall, local edaphic and biotic factors. These factors have been used in the classification of Indian forests. Most detailed classification of Indian forests is by Champion and Seth (1967) in which 16 major types of forests have been recognized. These 16 major types can be grouped into 5 major categories viz. moist tropical, dry tropical, montane sub-tropical, temperate and alpine forests.

Natural Vegetations in India

See also: http://www.envfor.nic.in/fsi/sfr99/misc/ifcmap.html


These forests are found in the areas of quite high temperature and rainfall. The forests are dense, multi-layered and have many types of trees, shrubs and lians. These forests are further categorized into 4 types depending on the degree of wetness in the area and the dominant life form in the forest.

(1)  Tropical moist evergreen forests

These are climatic climax forests found commonly in areas having annual rainfall above 250 cm and temperature 25-30oC. These forests are chiefly distributed on the western face of Western Ghats, Assam, Cachar, parts of West Bengal, northern Canara, Annamalai Hills and Coorg in Meysore and Andman Islands.

The characteristic feature of these forests is dense growth of very tall trees having height of more than 45 m. Climbers, lians, epiphytes and shrubs are abundant but herbs and grasses are rare in these forests. The carpet layer of herbs and grasses can not grow because very dense layer of leaf canopy of trees does not allow enough light to reach to the ground.

Dominant trees in forests of west coast are Dipterocarpus indica, Palaquim and Cellenia while in forests of Assam Diptercarpus macrocarpus, D. turbinatus, Shorea assamica, Mesua ferrea and Kayea are the dominant trees.

Common subdominants in these forests are Mangifera, Eugenia, Myristica, Pterospermum, Polyalthia, Elaeocarpus, Schlechera, Artocarpus, Memeocylon, Poeciloneuron, Cinnamomum, Diospyros, Sapindus, Vitex, Holigarna, Alstonia, Hardwickia, Spondias, Dendrocalamus, Calamus, Bombax, Veteria, Calophyllum, Pandanus, Cedrela, Tetrameles, Strobilanthes, Emblica, Michelia, Ixora, Hopea, Lagerstroemia, several species of ferns and orchids.

See also: http://www.indianetzone.com/2/forests.htm

(2) Tropical moist semi-evergreen forests

These are also climatic climax forests found commonly in areas of annual rainfall 200-250 cm and temperature 25-32oC.These forests are chiefly distributed along the Western Ghats, in upper parts of Assam and Orissa and in Andman Islands. These forests are more developed in the northern India than in southern India.

Characteristic feature of these forests is dense growth of evergreen trees intermixed with deciduous trees that shed their leaves for very brief period of relative dryness. Average height of trees in these forests is 25-35 m and shrubs are common. Forests have rich carpet layer of herbs, grasses ferns and orchids.

Dominant trees in these forests are Dipterocarpus alatus, Hopea, Terminalia and Salmalia in Andman Island; Artocarpus, Micheliaand Mangifera in Orissa; Schima wallichii, Bauhinia, Phobe and Ammora in Assam.

Common subdominants in these forests are Mylia, Schleichera, Bambusa, Ixora, Calamus, Sterculia, Webera, Strobilanthus, Cedrela, Shorea, Actinodaphne, Garcinia, Lagerstroemia, Mallotus, Vernonia, Dendrocalamus, Pelvetta, Elattaria, Pothos, Vitis, Garuga, Albizzia and Dellenia. Common herbs and grasses in the ground (carpet) layer are Inula, Andropogon, Crotolaria, Imperata, Leca, Desmodium, Fambosa and Woodfordia.

(3) Tropical moist deciduous forests

These forests are found in the area having temperature of 25-30oC and quite high annual rainfall of 150-200 cm spread over most of the year but periods of rain alternating with very short periods of dryness. In several areas, the forests have been converted into open savannahs due to intensive biotic factors. These forests are chiefly distributed in a narrow belt along Himalayan foothills, on the eastern side of Western Ghats, Chota Nagpur, Khasi hills, in moist areas of Kerala, Karnataka, sothern Madhya Pradesh, parts of northern Uttar Pradesh, Bihar and West Bengal.

Chief characteristic of these forests is dominance of deciduous trees that remain leafless for one or two months only along with lower story of smaller trees and evergreen shrubs.

Dominant trees of these forests in north India are Tectona grandis, Shorea robusta, Salmella, and Dalbergia while in south India only Tectona grandis and Shorea sp. are dominant.

Common subdominants in the forests are Cedrela, Albizzia, Terminalia, Adina, Melia, Sterculia, Grewia, Gariya, Lagerstroemia, Cordia, Pongamia, Bambusa, Dendrocalamus, Chloris, Mallotus, Anogeissus, heteropogon, Cymbopogon and Andropogon.

See also:


(4) Littoral and swamp forests

These forests are found in wet marshy areas, in river deltas, in saline or other swampy areas and along the sea coasts. They are chiefly distributed in deltas of large rivers on the eastern coast and in pockets on the western coast (Tidal forests), in saline swamps of Sundarban in West Bengal, coastal areas of Andhra and Orissa (Mangrove forests) and in less saline or non-saline swampy pockets throughout the India.

Chief characteristic of these forests is dominance of halophytic evergreen plants of varying height with varying density of plants in different area.

Dominant plants of tidal and mangrove forests are Rhizophora, Bruguiera, Ceriops, Horitora, Avicennia, Nipa, Sonneratia and Acanthus. In less saline swamps, dominant plants are Ipomea, Phoenix, Phragmitis, Casuarina, Manilkara and Calophyllum. In other swamps, the dominant plants are Barringtonia, Syzygium, Myristica, Bischofia, Trowia, Lagerstroemia, Sophora, Pandanus, Entada and Premna.

See also: http://www.indianetzone.com/39/indian_tidal_or_mangrove_forests.htm


These forests are found in the areas where wet season is followed by a relatively long period of dryness during which trees remain leafless. These forests are dominated by smaller trees and shrubs and have abundance of shrubs or sometimes grasses. This category includes three types of forests.

(1) Tropical dry deciduous forests

These forests are found in areas having temperature of 25-32oC and annual rainfall of 75-125 cm along with a dry season of about six months. Distribution of these forests in northern India is in areas of Punjab, Haryana, Uttar Pradesh, Bihar and Orissa. In the southern and central India, these forests are distributed in dry areas of Maharashtra, Tamilnadu, Karnataka and Madhya Pradesh.

Chief characteristic feature of the forests is open canopy of small (10-15 m high) trees and abundance of shrubs.

Dominant species of the forests in north India are Shorea robusta, anogeissus, Terminalia, Buchnnania, Somocarpus, Carissa, Emblica, Madhuca, Acacia, Aegle, Diospyros, Bauhinia, Eugenia, Zyzyphus, Lannea, Sterculia, Dendrocalamus, Salmelia, Adina, Grewia, Adathoda and Helicteres. In south India, dominant plants are Tectona grandis, Dalbergia, Kydia, Terminalia, Pterospermum, Dillenia, Acacia, Diospyros, Anogeissus, Boswellia, Bauhinia, Chloroxylon, Hardwickia, Soymida, Gymnosporia, Zyzyphus, Dendrocalamus and Holorrhena.

Subdominant species in these forests are Bambusa, Lantana and grasses like Panicum, Andropogon and Heteropogon.

See also: http://www.indianetzone.com/39/indian_dry_deciduous_forests.htm

(2) Tropical thorn forests

These forests are found in the areas of high temperature of 27-30oC and very low annual rainfall of 20-60 cm with long periods of dryness. These forests are distributed in western Rajasthan, parts of Maharashtra, Madhya Pradesh and Tamilnadu.

Chief charateristic of such forests is sparse distribution of small (8-10 m high) mostly thorny trees with shrubs being more common than trees. The plants in these forests remain leafless for most of the year. They develop leaves only during the brief rainy season when grasses and herbs also become abundant.

Dominant plants in these forests are Acacia nilotica, A. leucophloea, A. senegal, Prosopis spicigera, P. juliflora, Albizzia and Capparis.

Common subdominant plants are Zyzyphus, Anogeissus, Erythroxylon, Euphorbia, Cordia, Randia, Balanites, Salvadora, Gymnosporis, Leptadenia, Suaeda, Grewia, Gymnoma, Asparagus, Butea, Calotropis, Adathoda, Madhuca, Salmelia, Crotolaria, Tephrosia and Indigophera.

(3) Tropical dry evergreen forests

These forests are found in the areas of relatively high temperature and small rainfall available only during summers. The forests are distributed in some parts of Tamilnadu and Karnataka.

Chief characteristic features of the forests are dense distribution of mixed small evergreen and deciduous trees of 10-15 m height, absence of bamboos and abundance of grasses.

Dominant plants in the forests are Memecylon, Maba, Pavetta, Foronia, Terminalia, Ixora, Sterculia, Mesua and Schleichora.


These forests occur in the areas where climate is cooler than tropical but warmer than temperate areas i.e. on the hills between the altitudes of 1000 m and 2000 m. The forests are dominated by semi-xerophytic evergreen plants. This category includes three types of forests.

(1) Sub-tropical broad-leaved hill forests

These forests occur in relatively moist areas at lower altitudes on mountain ranges. Their chief distribution is in eastern Himalayas of West Bengal and Assam , hills of Khasi, Nilgiri and Mahabaleshwar.

Chief characteristic feature of the forests is dense growth of evergreen browd-leaved trees with abundant growth of climbers and epiphytic ferns and orchids.

Dominant trees in the forests of north are Quercus, Schima and Castanopsis with some temperate species. In the southern areas, dominants are Eugenia and members of family Lauraceae.

Common co-dominants and subdominants in the eastern Himalayas are Dalbergia sissoo, Acacia, catachu, Sterospermum, Cedrela toona, Bauhinia, Anthocephalus cadamba, Lagerstroemia parviflora, Albizzia procera, Salmella, Artocarpus chaplasha and Dendrocalamus. In the western Himalayas, codominants and subdominants are Shorea robusta, Dalbergia sissoo, Cedrela toona, Ficus glomerulata, Eugenia jambolina, Acacia catachu, Butea monosperma, Carissa and Zizyphus. Other common plants in these forests of both north and south India are Actinodaphne, Randia, Glochidion, Terminalia, Olea, Eleagnus, Murraya, Atylosia, Ficus, Pittosporum, Saccopetalum, Carreya, Alnus, Betula, Phobe, Cedrela, Garcinia and Polulus. In the south, Mangifera and Canthium and climers like Piper trichostachyon, Gnetum scandens and Smilax macrophylla are also common.

(2) Sub-tropical dry evergreen forests

These forests occur in areas having quite low temperature and rainfall. The forests are distributed in the lower altitudes of eastern and western Himalayas.

Chief characteristic feature of the forests is presence of thorny xerophytes and small-leaved evergreen plants.

Dominant plants in the forests are Acacia modesta, Dodonea viscosa and Olea cuspidata.

(3) Sub-tropical pine forests

These forests occur at middle altitudes between 1500-2000 m in Himalayas. They are distributed in western Himalayas from Kashmir to Uttar Pradesh. In eastern Himalayas, the forests occur in Khasi Jayantia Hills of Assam.

Chief characteristics of the forests in open formations of pine trees.

Dominant trees in the forests are P. roxburghii and Pinus khasiana.


These forests are found in the areas having quite low temperature along with comparatively high humidity than the comparable areas of higher latitudes. The cause of high humidity is greater rainfall in Himalayas except in parts of Uttar Pradesh, Punjab, Himachal Pradesh and Kashmir where humidity is lower. The forests occur mainly in the Himalayas at altitudes 2000-4000 m. The forests are generally dominated by tall conifers or angiospermic evergreen trees with abundance of epiphytic mosses, lichens and ferns. The category includes three types of forests.

See also http://www.indianetzone.com/2/temperate_deciduous_forests.htm

(1) Wet temperate forests

These forests are found at altitudes of 1800-3000 m in the cooler and humid mountains. They are distributed in the eastern Himalayas from eastern Nepal to Assam, in the western Himalayas from Kashmir to western Nepal and in Nilgiri Hills of south Indian.

Chief characteristic feature of the forests in the Himalayas is dense formation of evergreen, semievergreen broad-leaved and coniferous trees of up to 25 m height. In south India, these forests are termed Shola forests and mostly have 15-20 m high broad-leaved trees with dense leaf canopy, abundant epiphytic flora and rich herbaceous undergrowth.

Dominant trees in the forests of western Himalayas are angiosperms like Quercus, Betula, Acer, Ulmus, Populus, Corylus, Caprinus etc. and conifers like Abies, Picea, Cedrus etc. In eastern Himalayas, dominants are Quercus, Acer, Prunus, Ulmus, Eurya, Machilus, Symplocos, Mahonia, Begonia, Michelia, Thunbergia, Rhododendron, Arundinaria, Bucklandia, Pittosporum, Loranthus, Tsug and, Abies. In the Nilgiri Hills, the dominants are Rhododendron nilagiricum, Hopea, Balanocarpus, Artocarpus, Artocarpus, Elaeocarpus, Pterocarpus, Hardwickia, Myristica, Cordonia, Salmalia, Mucuna and Dioscorea. In all the areas, the undergrowth is formed by members of Asteraceae, Rubiaceae, Acanthaceae and Fabaceae.

(2) Himalayan moist temperate forests

These forests are found at 1700-3500 m altitude in eastern and western Himalayas. These occur in areas having annual rainfall above 100 cm but relatively less than that in areas of wet temperate forests.

Chief characteristic feature of the forests is presence of tall (up to 45 m high) conifers, oaks or their mixture along with thin partly deciduous undergrowth.

Dominant trees in the eastern Himalayas are Tsuga dumosa, Quercus lineata, Picea spinulosa, Abies densa and Quercus pachyphylla. In the western Himalayas, dominants in lower zones are Quercus incana, . dialata, Cedrus deodara, Pinus wallichiana, Picea smithiana, Abies pindrew, Cotoneaster, Berberis and Spire while in the higher zones the dominants are Quercus semicarpifolia and Abies pindrew.

(3) Himalayan dry temperate forests

These forests occur in the regions of Himalayas having very low rainfall. They are distributed in both eastern and western Himalayas.

Chief characteristic feature of the forests is dominance of evergreen oaks and conifers. Undergrowth is formed by scrubs.

Dominant trees in the forests of comparatively drier western Himalayas are Pinus gerardiana and Quercus ilex. In the comparatively wetter western Himalayan region, the dominants are Abies, Picea, Larix griffithia and Juniperus wallichiana.

Subdominant plants in these forests are Daphne, Artemesia, Fraxinus, Alnus, Cannabis and Plectranthus.

See also: http://www.indianetzone.com/11/himalayan_subtropical_pine_forests.htm


These forests are found in the regions of Himalayas having extremely low temperature and humidity. The forests are dominated by perennial and annual herbs and grasses though some trees may also be present in areas of relatively high humidity. Abundant lichen flora is characteristic feature of these forests. This category includes three types of forests.

(1) Sub-alpine forests

These forests are found in open strands throughout the Himalayas between the altitude 3500 m and the tree tine.

Chief characteristic feature of the forests is presence of some evergreen conifers and broad-leaved trees along with prominent shrub layer.

Dominant trees in the forests are Abies spectabilis, Rhododendron and Betula. Prominent shrubs in the forests are Cotoneaster, Rosa, Smilax, Lonicera and Strobilanthus.

(2) Moist alpine scrub forests

These forests are found in the Himalayas above the tree line up to 5500 m altitude in somewhat moist areas.

Chief characteristic feature of the forests is dominance of dwarf, evergreen shrubby conifers and broad-leaved trees along with prominent shrub layer under them.

Dominant trees in the forests are Juniperus and Rhododendron while prominent shrubs are Thalictrum, Lonicera, Saxifraga, Arenaria, Bergia, Sedum and Primula.

(3) Dry alpine forests

These forests are found in comparatively more dry areas of Himalayas upto 5500 m altitude.

Chief characteristic feature of the forests is open formation of xerophytic scrubs with many herbs and grasses.

Dominant plants in the forests are Juniperus, Caragana, Eurctia, Salix and Myricaria.


The grasslands of India are not of primary origin. These have originated secondarily in many areas due to destruction of natural forests by biotic interference, particularly due to excessive grazing and land clearing for agriculture. These grasslands are maintained in various seral (successional) stages by a variety of biotic factors.

According to the dryness of the area, the Indian grasslands may be categorized into three types.

  1. Xerophilous grasslands: These are found in semi-desert areas of north and west India.

  2. Mesophilous grasslands (Savannahas): These are found in areas of Uttar Pradesh having moist deciduous forests.

  3. Hygrophilous grasslands (Wet savannahas): These are found in wet regions of India.

Whyte et al. (1954) classified Indian grasslands on the basis of dominant grass species into eight major grass associations.

(1) Sehima-Dichanthium association

These grasslands develop on black soil. They are found in some areas of Maharashtra, Madhya Pradesh, Uttar Pradesh, south western Uttar Pradesh, Tamilnadu and Karnataka.

Dominant grass species in the grasslands are Sehima sulcatum, S. nervosum, Dichanthium annulatum, Chrysopogon montanus and Themeda quadrivalvia.

(2) Dichanthium-Cenchrus association

These grasslands develop on sandy-loam soils. They are found in Plains of Punjab, Haryana, Delhi, Rajasthan, Saurashtra, eastern Uttar Pradesh, Bihar, West Bengal, eastern Madhya Pradesh, coastal Maharashtra and Tamilnadu.

Dominant species in these grasslands are Dichanthium annulatum and Cenchrus ciliaris.

(3) Phragmitis-Saccharum association

These grasslands develop in marshy areas. They are found in terai regions of northern Uttar Pradesh, Bihar, West Bengal, Sundarban region of Bengal, Tamilnadu, and Kaveri delta.

Dominant species in these grasslands are Phragmitis karka,, Saccharum spontaneum, Imeerata cylindrica and Bothriochlo pertusa.

(4) Cymbopogon type

These grasslands develop on low hills. They are found in Eastern Ghats, Vidhyas, Satpura, Aravali and Chota Nagpur.

Dominant species in the grasslands is Cymbopogon.

(5) Arundinella type

These grasslands develop on high hills. They are found in Western Ghats, Nilgiris and lower Himalayas from Assam to Kashmir.

Dominant species in the grasslands are Arundinella nepalensis, A. setosa and Themeda anthera.

(6) Bothriochloa type

These grasslands develop on paddy tracts in areas of heavey rainfall in Lonavala tract of Maharashtra.

Dominant species in the grasslands is Bothriochloa odorata.

(7) Deyeuxia-Arundinella association

These grasslands develop in temperate areas of upper Himalayas between 2100-3500 m altitudes.

Dominant species in the grasslands are Deyeuxia, Arundinella, Brachypodium, Bromus and Festuca.

(8) Deschampsia-Deyeuxia association

These grasslands develop in temperate to alpine regions having thin soil cover over rocky substratum. They are found in Kashmir and in Himalayas above 2600 m altitude.

Dominant species in the grasslands are Deyeuxia, Deschampsia, Poa, Stipa, Glycera and Festuca.


Filed under: Soil — gargpk @ 1:28 pm
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Indian soils have been classified into eight major types according to their origin and physico-chemical properties.


These soils cover an area of about 1,500,000 square kilometers. The soils develop from the weathering material transported by rivers from their catchment areas and deposited in their basins during floods. Chemical nature and fineness of inorganic material depends on the type of rock in the catchment area and the degree of weathering of the rocks. Geologically these soils are of two types.

  • New alluvium (Khadar): These are soils that have been deposited more recently and are constantly replenished by deposition or more transported material during regular floods. The soils are found in low lands and deltas. The soils are generally sandy and light coloured having lesser proportion of soil aggregates (kankars).

  • Old alluvium (Bangar or Bhangar): These soils were deposited at an earlier time and have not naturally been disturbed much. These soils are found in upland areas of river deltas.They are more differentiated, more clayey, generally dark coloured and full of soil aggregates (kankars) than the new alluvium soils. Occasional pebble-beds may be present and hard pans may be formed sometimes at certain depths due to binding of soil grains by infiltrating silica or calcareous matter.

Alluvial soils may be broadly divided into coastal/deltaic alluvium and inland alluvium according to their location.

  1. Coastal/deltaic alluvium: These soils are distributed in the deltas of major rivers e.g. Ganges, Brahmputra, Krishna, Kaveri, Narmada, Tapti, Godavari, Mahanadi etc. and on the coasts of Orissa, Andhra, Tamilnadu, Kerala, Karnataka and Maharashtra. These are new alluvial soils and vegetation provides humus in them. New inorganic material is deposited regularly during floods. The soils have dark colour (in southern India), loamy texture (in Sundarban delta), high humus, high water holding capacity, high base exchange capacity, neutral to saline pH, humid or swampy nature and rich in potassium in most of the areas. The soils are suitable for crops that require humid soils e.g. rice, jute and coconut.

  2. Inland alluvium: These soils are found in the river basins of northern and southern India. They include new as well as old alluvium. Characters of the soil vary in different regions due to the type of parent rock material, climate and type of plant cover.

  1. Brahmputra alluvium: This is old alluvium distributed in the uplands of Brahmputra valley. The soil is acidic, sandy, rich in total potash and phosphorus. They have moderate amount of organic matter and nitrogen. New alluvium is found along the banks of the river. It is less acidic and neutral or even alkaline in pH.

  2. Indogangetic alluvium: These soils are distributed in large plains of the rivers Indus and Ganges in Punjab, Uttar Pradesh, Bihar and West Bengal.

  1. Punjab alluvium: This soil is sandy or loamy in texture with soil crust of varying depth. The soil is undifferentiated and has soil aggregates (kankars) in lower layers. The soil has high soluble salts and high base-exchange capacity. Phosphorus and potash content of the soil is adequate while nitrogen and organic matter contents are deficient. The soil has alkaline pH.

  2. Allluvium of Uttar Pradesh: These soils are broadly categorized into:

  1. Light coloured alluvium: This soil is found in the plains of west and north-western Uttar Pradesh.

  2. Intermediate textured alluvium: It is distributed in the central basins of rivers Ganges and Jamuna.

  3. Dark coloured alluvium: The soil is found in the eastern Uttar Pradesh.

  4. Calcareous alluvium: This soil develops on calcareous parent rock material and is found in north eastern Uttar Pradesh. These soils have varying amounts of soluble salts and base-exchange capacity in different regions. The pH of the soil is neutral to alkaline.

  1. Bihar alluvium: It is categorized into:

  1. Northern Bihar alluvium: These soils are found in the area north of river Ganges up to Himalayas in Bihar. The area includes a calcareous belt as a triangle in the west and broken inundated areas that remain flooded for different periods of the year in the middle of the region. The soils are sandy loam to clayey loam in texture, neutral to alkaline in pH rich in lime and potash and poor in phosphorus.

  2. Southern Bihar alluvium: These soils are found in the area between river Ganges and the southern hilly region in Bihar. The area includes a depressed area in the middle that remains flooded during monsoon. The soils have texture similar to northern Bihar alluvium and almost neutral pH. In these soils, potash and phosphate content is higher while lime content is lower than in northern Bihar alluvium.

  1. West Bengal alluvium: These soils are categorized into:

  1. Old alluvium of West Bengal: These soils are distributed in Rath region of Murshidabad and Bankura areas. The soils are clayey in texture.

  2. New alluvium of West Bengal: This soil is distributed in rest of the plains of West Bengal. The soil is sandy loam to clayey in texture with clay pan at certain horizons. This soil is more fertile than the old alluvium.

  1. Gujrat alluvium: These soils are locally known as goradu and distributed in north Gujrat tract, Ahmedabad and Kaira districts. The soils are categorized into:

  1. Old alluvium (Gorat soils): These are found in Baroda region. They are clayey in texture, brown in colour and have soil aggregates (kankars).

  2. New alluvium (Bhata soils): These are fairly deep soils and mainly of secondary deposition. The soils are fairly rich in phosphorus and potash but poor in nitrogen and organic matter.

  1. Red and yellow alluvium: The soils are found in the basin of river Mahanadi in Madhya Pradesh distributed in Balaghat, Durg, Raipur and Bilaspur areas. These soils are grouped into:

  1. Bhata soils: These are gravely sandy in texture, reddish brown in colour and found in uplands of barren wastelands.

  2. Matasi soils: These are loamy to clayey in texture, yellow in colour and found in upland areas. These soils are suitable for paddy cultivation.

  3. Dorsa soils: These soils have texture like Matasi soils but a darker colour than those. The soils are found on the slopes. These soils are good for paddy cultivation.

  4. Kanhar soils: These are darker and heavier soils found in lowlands. These are good for paddy and suitable for wheat cultivation.

  1. Orissa alluvium: These soils are sandy and of finer texture having sufficient potash but poor phosphate content. According to the location, these soils are grouped into:

  1. At soils: These are found in dry upland areas.

  2. Mal soils and Berna soils: These are found on intermediate slopes.

  3. Behal soils: These are the heaviest soils found in lowland area.

Alluvium soils are generally suitable for a variety of crops like wheat, rice, millets, pulses, maize, sugarcane, rubber, jute, vegetables etc. depending on the climate, texture and other soil characteristics. The nitrogen deficiency in the soil is the main limiting factor in crop production in all the alluvial soils except the calcareous and saline alkali alluvial soils. In these later soils, high salinity and low phosphate limit the crop production. With suitable irrigation, fertilizers and proper agricultural practices, alluvial soils are generally very fertile.


These soils cover about 5,46,000 square kilometers in India. They are distributed in Bundelkhand region of Uttar Pradesh, western Madhya Pradesh, Bundi and Tonk districts of northern Rajasthan. They are also present in parts of Gujrat, Karnataka, Andhra, Maharashtra and Tamilnadu. The soils are residual soils derived from the rocks of Deccan Trap, Rajmahal Trap, ferruginous gneisses and schists under semi-arid conditions. Soils have been formed by alternate deposition and assimilation of organic matter with weathering products of the rocks. The black colour of soils is due to superficial rock iron. The soils have high content of montmorillonite mineral. These soils are categorized into three types according to the depth, colour and rock matter.

  1. Shallow black soils: These are 30-50 cm deep and derived from basalts of trap rocks.

  2. Medium black soils: These are 50-120 cm deep soils derived from a variety of rocks including basaltic trap, Dharwar schists, basic granites, gneisses, hornblende and chlorite schists.

  3. Deep black soils: These are 120-200 cm deeps soils derived from basaltic traps and contain irregular lime nodules (kankars).

Black soils are deep gray to black in colour, clayey or clay-loam in texture, fine grained with small surface grains, transparent or semi-transparent, cemented by coloured matrix of double hydrated ferrous and aluminum silicates. They are rich in oxides of iron and aluminum and carbonates of calcium and magnesium. The soils are poor in phosphate and nitrogen. The soils have high water holding capacity, high content of exchangeable bases and alkaline soil pH. The soils are sticky in wet condition and contract in dry condition resulting in heavy cracks and fissures in the soil. This phenomenon is termed self-ploughing. In Maharashtra, upland soils and soils of slopes are light-coloured, thin and poor while lowland and valley soils are deep, clayey and rich with much sub-soil lime. The soils along the Ghats are very coarse and gravely. In Madhya Pradesh, both shallow and deep black soils are found. Black soils in Karnataka are quite heavy with high amounts of salts, lime and magnesia.

Black soils are best suited for the cultivation of cotton. With proper irrigation and ploughing, the soils can also be used for cultivation of wheat, millet, sugarcane, linseed, gram and pulses. These soils require good preparation of land before irrigation. Conservation of the moisture in shallow and medium black soils by contour-bunding is useful in increasing the production of rabi crops like jowar. Owing to high amounts of salts, uncontrolled irrigation of black soils makes these soils alkaline or saline.


These soils cover about 3,50,000 square kilometers in India. The soils are distributed in major portion of Tamilnadu, Karnataka, southeastern Maharashtra, eastern Andhra, Madhya Pradesh, Orissa, Chota Nagpur, southern Assam, western West Bengal, eastern hills of Aravali in Rajasthan and southern districts (Jhansi, Banda, Hamirpur, Varanasi and Mirzapur) of Uttar Pradesh. The soils have developed on crystalline rocks by prolonged weathering in hot climate. Ferrous oxide is precipitated after evapouration of soil moisture giving the red colour. The soils are generally dark red, red-black or grey-brown in clour, rich in mineral kaolinite and neutral to acidic in pH. The soils have low base-exchange capacity, low water-holding capacity, low organic matter content, low nitrogen, low phosphate, high potash and high lime. Soil aggregates (kankars) are absent in the soils. The soils have silica and aluminum with free quartz or sand and the soil aggregation is poor. The soils are grouped into:

  1. Red loam soils: These are dark red in colour and loamy in texture.

  2. Red gravelly soils: These are dark red and gravelly in texture.

  3. Mixed red and black soils: These soils are found in areas that intergrade into black soil regions.

Cultivation of red soils depends on the irrigation. Most of soils in Tamilnadu, Karnataka and Andhra have been cultivated by tank irrigation and manuring. These soils can be used for the cultivation of wheat, rice, millet, gram and sugarcane. Groundnut, coconut and ragi are also cultivated on these soils.


These soils are found in about 2,48,000 square kilometers in India. They are distributed on the hills of peninsular India, in eastern Ghat region of Orissa, parts of Assam and central India. The soils are also found in Karnataka, Maharashtra, Malabar, parts of West Bengal, Bihar, Madhya Pradesh and Tamilnadu. The soils are residual soils developed on aluminum-rich rocks e.g. greisze, sandstone, granite and basaltic rocks in regions of high rainfall with alternating wet and dry seasons. The soils develop by the process of laterization in which silica and alkali are leached down and remaining compounds rich in oxides of iron and aluminum form the soil. The soil is compact to vesicular rock essentially composed of a mixture of hydrated oxides of aluminum and iron with small amounts of the oxides of manganese, titanium etc. The soil may be broken and transported by streams to lower levels and deposited there becoming cemented again into compact mass. Thus the higher elevation laterite is residual and lower elevation laterite is alluvial in nature.

The soils are compact, reddish or yellowish-red in colour and turn black on exposure to the sun. The texture of these soils is gravelly, loamy or clay-like loam. The soils are deficient in available exchangeable bases, essential plant nutrients, potash, nitrogen and organic matter. The nitrogen content of the soils varies from 0.03% to 0.06% and pH ranges from 4.8 to 5.5. The base-exchange capacity of these soils is generally low. The laterites of higher elevation are more acidic and comparatively richer in nutrients than the laterites of lower elevation. Occasionally the humus content of the soils may be high e.g. in Kanara region of Maharashtra where nutrients, nitrogen and potash are also high. Sometimes soils may contain iron phosphate.

The cultivation of these soils requires good management practices to check the soil erosion, application of proper fertilizers, liming and proper irrigation. Higher elevation soils grow good crops of tea, cinchona, rubber and coffee. Lower elevation soils with proper management can grow paddy, rice, sugarcane and pulses like green gram and gram. Soils generally show good response to liming and application of potash and nitrogenous fertilizers according to the type and pH of the soil. In some lowland areas, iron and manganese cause toxicity due to reduction of ferric and trivalent Mn to ferric and divalent Mn respectively. Liming can correct the situation and uptake of nitrogen and phosphorus is normalised.


  1. Forest soils: These cover an area of about 2,85,000 square kilometers in India. Soils are found in forests of Malabar, hills of Coorg, hilly districts of Assam, Darzeeling in West Bengal, sub-Himalayan tracts of Uttar Pradesh and Himachal Pradesh. These soils develop in areas under heavy forest cover. In some areas, the soils are under acidic conditions with presence of acidic humus and low base-exchange capacity. In other areas, the soils are under slightly acidic to neutral conditions where brown earths with high base-exchange capacity are developed. The chief factor in the development of these soils is abundance of organic matter derived from forest growth. These soils are mostly under forest cover but are suitable for the cultivation of crops particularly sugarcane, rice, paddy and millet.

  2. Peat and organic soils: These develop in waterlogged, swampy or marshy areas. The soils develop due to deposition of large amount of poorly decomposed organic matter on the weathered rock material. These soils are found in coastal tracts of Orissa, Sunderban in West Bengal, Almora in Uttar Pradesh, northern Bihar, southeastern Tamilnadu and Kerala. The soils are bluish, very loose due to excess of organic matter, acidic, very rich in free aluminum-iron componds of blue colour and poor in lime. The soils become waterlogged during monsoon. These soils are put under paddy cultivation after monsoon is over. These soils are generally very poor for agriculture.


These soils develop in the foothills and mountains up to the altitude of 12,000 feet. Thickness, texture and properties of the soils depend upon altitude. At higher altitudes, the soils are thin and gravelly while at lower levels, they are fine and rich. The soils generally have variable colours depending on the type of rock and the amount of organic matter. The texture of the soils is loose, gravelly, sandy, loamy or sometimes clayey. These are young and immature soils. The A-horizon of the profile is formed directly on the rock and soil particles are loosely aggregated on the soft sandy beds. Humus content of the soils is generally high. Soils are usually acidic but those developed on limestone are basic in reaction. Soils are deficient in lime but rich in minerals, iron salts and nitrogen. These soils are categorized as:

  1. Terai soils: These are found near foothills of Himalaya in Uttar Pradesh and in Simla region of Himachal Pradesh. The soils are clayey in texture with high humus content and high water-holding capacity.

  2. Tea soils: These are found in Assam, Dehradun district of Uttar Pradesh, Darzeeling district of West Bengal and slopes of Nilgiri hills. The soils are loamy in texture with low lime content.

  3. Igneous soils: These are fine soils derived from granite rocks and rich in nutrients.

  4. Soils of old rocks: These are found in Nanital district of Uttar Pradesh. The soils are rich in clay and iron.

  5. Limestone soils: The soils are found near Mussouri in Uttar Pradesh. The soils are rich in lime, loamy in texture and basic in pH.

The Terai soils are very good for the cultivation of wheat, maize and sugarcane. Tea soils are very good for plantations of tea. Limestone soils are particularly good for rice. Igneous soils and soils of old rocks are fertile soils good for various crops, fruit trees and vegetables. All these soils are generally suitable for timber plantations of sal, pine, deodar etc.


These soils cover about 1,42,000 square kilometers in India and develop in areas of very low rainfall, very high temperatures and absence of vegetation. The evapouration from the soil is very high owing to absence of plant cover and soils have almost no soil moisture for major part of the year. These soils are distributed in northwest Rajasthan, southern Punjab, Haryana, in the areas between the river Indus and Arawali hills. The soils are of geologically recent origin. The areas are covered under a mantle of brown sand that inhibits the growth of soil. Sands of Rajasthan are partly derived from the disintegration of subjacent tracks but are largely blown in from coastal regions and valley of Indus river.Some soils have high soluble salts, alkaline pH and deficiency of organic matter.

These soils are normally infertile but management of irrigation can make these soils cultivable. By gradually introducing plant cover, these soils can be stabilized along with the increase in their humus content and water-holding capacity. With availability of proper irrigation, these soils can support the cultivation of paddy, millet, oat and even wheat.


These soils develop in arid areas having low rainfall, poor drainage, high temperature and high evapouration. These soils are found scattered within all the major types of soils. They are present in the indogangetic alluvium of Punjab, Haryana, Delhi, Uttar Pradesh and Rajasthan; in deltaic coastal alluvium of West Bengal, Orissa, Andhra, Tamilnadu, Kerala and Gujrat; in deep and medium black soils of Gujrat, Madhya Pradesh, Mharashtra, Karnataka and Andhra and in desert gray brown soils of Gujarat and Rajasthan. The normal soil of the area becomes alkaline or saline owing to deposition of high amounts of the soluble salts of sodium, magnesium or calcium due to poor drainage. In summers, capillary action in the soil brings the salts to the surface where they are deposited in the form of a white or black effervescent crust.

These soils are normally infertile (Usar soils). However, the management of drainage, good irrigation and application of lime and/or gypsum combined with biological management can make these soils fit for the cultivation of a variety of crops. Alkali soils can grow wheat, barley, oat, sorghum, sugarbeat, cotton, grapes etc. Saline-alkali soils can grow paddy, oat and barley while saline soils can grow rice, sugarcane, barley and castor. These soils can be successfully managed with green manuring by Dhaincha or Barseem in rotation with cereal crop.


The soils of Uttar Pradesh can be grouped into following categories:


These include both new and old alluvium developed by transportation of weathered rock material from Himalayas by rivers Ganges, Jamuna, Sarda and Rapti and its deposition in their basins.

Alluvium of Uttaar Pradesh has sandy loam, loam or clay loam texture. It is generally poor in nitrogen and available phosphorus but rich in calcium carbonate. The colour of the soil depends on the amount of sand and organic matter in it. In the areas of poor drainage due to presence of hard pans in the sub-soil or in areas where water table has risen due to heavy irrigation, the soils have become impregnated with salts. In extreme cases, these have turned into saline-alkaline soils. Alluvial soils of Uttar Pradesh are categorized into:

  1. Alluvium of northeast: It is light sandy loam in texture, calcareous in nature with neutral to slightly alkaline pH and deficient in nitrogen. These soils are quite fertile if nitrogen and phosphate fertilizers are added.

  2. Alluvium of west: It is comparatively more sandy in texture, grayish-yellow in colour, low in nitrogen, phosphorus and organic matter, impregnated with salts, mostly neutral in pH and quite fertileif proper fertilizers and irrigation are given.

  3. Alluvium of central region: These soils having medium loam texture may be grouped as:

    1. Northern calcareous alluvium: It is rich in calcium carbonate.

    2. Southern gray-brown alluvium: It is rich on impregnated salts and gray-brown in colour.

  4. Alluvium of east: These are heavier and more clayey in texture with slightly higher available potash and phosphate than the southeastern alluvium. These soils are also gray-brown in colour and impregnated with salts. Soil pH is almost neutral.

  5. Alluvium of northeast: Soils in the districts of Gorakhpur and Deoria are developed on calcareous parent material. These have variable amounts of calcium carbonate and soluble salts. Soil pH is neutral to alkaline. The soils are quite fertile and sugarcane is the major crop on these.


These soils are found in the northern districs of Jhansi, Hamirpur, Banda, Mirzapur and some southern areas of Varanasi. The soils have patches of red soil over the black soil or vice versa. In some areas, red and black soils have become thoroughly mixed giving a brown colour. These soils may be grouped as:

  1. Black clay soils: These are found in lower Gangetic basin and are locally known as karail. These soils are thought to have developed from the parent rock material similar to that of black (regur) soils of southern India. The basaltic alluvial material from Deccan trap rocks of Bundelkhand region has been by rivers coming from that region and added to the river Ganges. From there, it has been deposited in its southern basin in the areas where conditions were favourable. These soils have high montmorillonite mineral, high water-holding capacity and high exchangeable bases like potassium, calcium and Magnesium. Calcium carbonate kankars are common in the soils. The soils become saline alkaline on uncontrolled irrigation due to the presence of high amounts of soluble salts.

  2. Red soils: These are formed by very old crystalline and metamorphic rock material. The soils are generally poor in nitrogen, phophorus, potash and humus. Compared to black soils, these soils are poor in lime, potash, phosphorus and iron oxide.In Jhansi district, these soils are of two types that are locally known as:

    1. Parwa soils: These are brownish-gray in colour owing to mixing of black and red soils. Texturre of the soil varies from good loam to clayey loam.

    2. Rakar soils: These are true red soils having low nutrients and are not fit for agriculture.

  3. Red-yellow soils: These soils are found in Mirzapur district where black soil component is very low and yellow colour is due to hydration of the ferric oxide. Soil pH is neutral and texture is loam to silty loam. Soils are well drained, rich in aluminum and moderately rich in humus.


These soils have eeveloped on lower altitudes of Himalayas in the areas that are under dense plant cover and have low temperatures with temperate climate.

  1. Forest soils: These soils are found in Himalayan region of Uttar Pradesh between the altitudes of 1,000 and 3,500 meters. The soils show strong characteristic influence of vegetation on their development. These soils are categorized as:

    1. Sub-montane soils: These soils are found in the regions having rainfall of 170-225 cm and vegetation of coniferous forests. Soil has dark brown to black colour, acidic soil pH and a thick surface layer of organic matter. Soils are similar to podosol soils but show lesser leaching.

    2. Brown hill soils: These soils are found under temperate type vegetation. The soils have been derived from Shale and Sandstone rock materials. The are brown in colour, loamy to silty loam in texture and quite rich in organic matter. Soil pH is acidic but base exchange capacity of the soil is quite high. The soils are suitable for cultivation of cereals like rice, paddy, millets etc.

  2. Peaty soils: These soils have developed in waterlogged, low-lying areas in under the plant cover in Almora district. Low oxygen content and low temperature results in poor and slow decomposition of the organic matter deposited by the forest cover resulting in accumulation of a thick layer of peat on the weathered parent rock material. The soils are bluish in colour, very loose in structure, acidic in pH, rich in blue coloured free compounds of Aluminum and Iron, poor in lime and have low base exchange capacity. The soils are not good for agricultural purposes.


These soils are found in the foothills of Himalayas. They are present from lower Kumaon-Garhwal region up to Gorakhpur district in the east extending along the Nepal border. These soils are grouped into five categories in Uttar Pradesh.

  1. Bhabar soils: These soils are found immediately below the hills and haveen formed by deposition of parent material from erosion of hills. These soils are dry and generally have food good organic matter.

  2. Terai soils: After the Bhabar soils, these soils form a zone of different width along the foothills and Nepal border. The soils are sandy or silty loam in texture with moderate clay and high humus content. Generally these soils are waterlogged and have very rich vegetation. In the areas of good drainage, these soils form very fertile lands for various cash crops particularly sugarcane.

  3. Plain mountaineous soils: After the zone of Terai soils, these soils are found in comparatively more level areas. These soils have clayey texture and are rich in humus. The base-exchange capacity of the soils is quite high and soils are very fertile.

  4. Tea soils: These soils are found on the slopes in the hills of Dehradun district. The soil pH is acidic and they are very good soils for plantations of tea.

  5. Limestone soils and soils of old rocks: These soils have developed on the limestone rock material and show characteristic features.

    1. Limestone soils: These are found near Mussouri. These have developed on the limestone rock material. The soils have very high lime content, basic pH and loamy texture.

    2. Soils of old rocks: These are found in some areas of Nainital district. They have developed on very old rocks in very long geological time and are very rich in clay along with iron contents.


These soils are found in higher altitudes of Himalayan regions i.e. Kumaon and Garhwal regions of Uttar Pradesh. The soils extend in regions beyond the tree line up to the region of snow covered rocks. The soils are very thin, undifferentiated and have very slow decomposition of organic matter due to cold temperatures. These soils are commonly grouped as:

  1. Mountain meadow soils: These soils are found in the hills beyond tree line in the areas of comparatively more rainfall but quite low temperature. The soils have been derived from Shale and Sandstone parent materials and are gravely in texture with surface layer of undecomposed or partially decomposed organic matter. Soils are covered mostly with herbaceous vegetation or grass cover.

  2. Skeletal soils: These soils are found at higher altitudes in the hills where both temperature and rainfall are very low. Due to very slow weathering, the soils are very thin, usually 7-15 cm in thickness and undifferentiated. The soils are pale brown to dark brown in colour and sandy loam to loam in texture. The soils are covered mostly by xerophytic and sclerophyllous vegetation.


These soils are found in patches all over the state, particularly in the districts of Kanpur, Hardoi, Lucknow, Unnao, RaiBareilly, Azamgarh and Mirzapur. Such soils have also developed secondarily in area of heavy irrigation and high summer temperatures. In these areas, water table rises that brings the soluble salts of calcium, magnesium and sodium to the surface of soil. The saline and alkaline soils are infertile termed Usar. However, these soils can be reclaimed by proper soil management practices and can be cultivated.


In the global environment, a vast number of elements exist in a variety of chemical species and are continually transformed from one species to another. These transformations from one chemical species to another involve cycling of these elements or chemical species amongst different components of the environment i.e. amongst atmosphere, lithosphere, hydrosphere and biosphere. These cycles of elements involving different components of the environment are, therefore, considered as biogeochemical cycles. These biogeochemical cycles are highly complex and interact strongly with each other and, therefore, are of fundamental importance in maintaining global environmental balance and in understanding the dynamics of environment. Further, human activities cause increase or decrease in natural amounts of chemical species in the environment or cause addition of chemical species not found in nature resulting in disturbance in natural biogeochemical cycles. Such disturbances constitute environmental pollution, which has profound impact on the stability of global environment. The biogeochemical cycles of element carbon, nitrogen and sulfur are most important from the point of view of global environmental balance. Therefore, important features of the cycles of these three elements have been discussed below.


The carbon cycle is mainly associated with living matter, although inorganic carbon provides important segments to complete the cycle. The cycling of carbon is strongly controlled by its storage in natural reservoirs. The time period of such storage may range from millennia in rocks, through decades in deep ocean layers to seasons in active biota. Relevant time periods of such storage suggested by Warneck (1988) are:

1. Geological activity involving rocks: 2,400 to 30,000 years

2. Soil humus: 200 years

3. Long-term biosphere storage: 75 years

4. Short-term biosphere storage: 15 years

5. Ocean mixed layers: 4 to 10 years

Estimates of mass content of carbon in various global reservoirs are given in the Table-1.

Carbon in oceans

Major storage of carbon in oceans occurs in the intermediate and deep water below the thermocline. The deep layers of oceans have a very slow mixing period and carbon remains in situ for atleast 20 years in these layers. Far above in oceans, in the mixed layer, which provides the main medium of interchange with the atmosphere, carbon storage is about 1.5 orders of magnitude lower. Ninety percent of the carbon in the oceans is stored as bicarbonate (CO32-) and about 9% as carbonate (CO3). About 3% of carbon is present in organic matter in environment.

The mixing layer in oceans, broadly the layer above the thermocline, is assumed to be at depth of 75 meters. The average concentration of carbon dioxide in the oceanic surface layer (above the mixing layer) is 2.05 mmol m-3. This concentration rises rapidly with depth to about 2.29 mmol m-3 at the depth of about one-kilometer and remains fairly constant thereafter. The average oceanic carbon dioxide concentration is calculated to be about 2.25 mmol m-3. Since colder ocean water is able to hold more carbon dioxide, variations in its concentration occur with temperature of ocean water. The mass of carbon dioxide in the mixed layer is about the same as that in the atmosphere, with a total exchange between the two estimated to occur over a period of about seven years.

Table-1. Mass content of carbon in global reservoirs



in Pg (1015 g)


1. Total dissolved CO2


2. Dissolved CO2 in mixed layer (75 m depth)


3. Living biomass carbon


4. Dissolved organic carbon



1. Continental and shelf carbonates

270 x 105

2. Carbonates in oceans

230 x 105

3. Continental & shelf organic carbon

100 x 105

4. Organic carbon in oceans

200 x 104


1. Terrestrial biomass


2. Soil organic


3. Oceanic organic


ATMOSPHERE (mostly as CO2)

1. Pre-industrial estimate (290 ppmv)


2. present estimate (350 ppmv)


Organic carbon in oceans comes from precipitated remains of living organisms. About 80% of the precipitated material may be redissolved in the deep ocean layers. Dissolved organic carbon content of ocean waters is roughly estimated to be about 0.7 g m-3. Rest of the carbon in the ocean is particulate, mainly as calcium carbonate and this portion of oceanic carbon has a concentration of about 20 mg m-3. Living organisms contribute a total of only 3 Pg to the oceanic carbon storage.

Carbon in sediments and rocks

Carbon makes up only 0.032% of the Earth’s crust by mass. In terrestrial rocks, it is dissolved by rains or surface water over long periods of time and is carried by the surface runoff water to be deposited on the continental shelf sediments. In deeper oceans, deposits from organisms are built up on the ocean floor over millennia. Exchange of carbon from these locations occurs over thousands of years and is associated with activity of Earth’s crust. About two third of this carbon is inorganic carbon and rest is organic in form. Only about 1% of carbon in the form of oil and coal present in Earth’s crust can be used economically. It is estimated that if all the carbon stored in sediments is released suddenly, the atmospheric pressure will rise by 38 bars and the Earth’s atmosphere will become similar to that of planet Venus.

Carbon in biosphere

In the biosphere carbon is exchanged through:

  1. Photosynthetic activity of photosynthetic living organisms, mainly the green plants

  2. Release of carbon on decay and decomposition of dead living organisms

  3. Respiratory activity of all the aerobic living organisms including both plants and animals

  4. Release of carbon from soil humus

The mass of carbon is about three times higher than in living biosphere. The biospheric exchange processes are relatively inactive and the carbon storage may occur for 200 years. Long-lived species, particularly the plants store about 75% of the carbon present in the living biota. The major impacts on global carbon content present in the active biosphere occur in the forests, which store over 80% of the world’s biomass. Though estimates are uncertain because global distribution of different ecosystems is not known accurately, it is quite clear that tropical rain-forests, boreal forests and temperate forests are the most important ecosystems regarding storage and exchange of carbon.

Carbon in atmosphere

Exchange of carbon with the atmosphere occurs mainly through the biosphere with oceanic mixed layer being an important secondary source. Most important atmospheric form of carbon is CO2 gas and global estimates of its exchange between atmosphere and biosphere are:

1. Assimilation of CO2 into plants: 113 Pg Y-1

2. Re-release into atmosphere from:

  1. Respiration of living organisms: 55 Pg Y-1

  2. Microbial decay: 42 Pg Y-1

  3. Soil humus: 10 Pg Y-1

  4. Forest fires and agricultural burning: 1 Pg Y-1

3. Herbivore consumption: 5 Pg Y-1

In addition to CO2, other minor gases in the carbon chain are carbon monoxide (CO), methane (CH4) and non-methane hydrocarbons (NMHCs e.g. HCHO). Carbon dioxide gas is relatively inert while others are quite active in global atmospheric chemistry. Important features of atmospheric carbon species are discussed below.

1. Carbon dioxide: Though CO2 is a minor gas in the atmosphere in comparison with oxygen and nitrogen, it has major impact on global heat balance because of its high capacity of absorbing infra-red radiation. Continuously rising concentration of atmospheric CO2 due to various human activities, particularly the fossil-fuel burning, is major factor in global greenhouse warming. Anthropogenic carbon contributes about 3% of annual carbon loading. Further, its importance in relation to biosphere is supreme since it is required for photosynthesis and existence of biosphere depends on photosynthesis.

2. Carbon monoxide: About 90% of CO originates during photochemical production of methane in atmosphere. Some CO is produced during biomass burning and some during atmospheric oxidation of organic gases that are emitted from vegetation. Highest concentrations of CO are found in middle and high latitudes of Northern Hemisphere, which may reach 150 – 200 ppbv. The concentrations of atmospheric CO show a definite seasonal rhythm and are higher in summers than in winters. In Southern Hemisphere, CO concentrations are lower than in Northern Hemisphere by a factor of upto three. CO is removed from the atmosphere mainly by being oxidized to CO2.

3. Methane: This is a trace gas in atmosphere and is released mainly from rice paddies, wetland areas, enteric fermentation from animals and biomass burning. It has a uniform latitudinal distribution with an average concentration of about 1.6 ppmv. Major sinks of methane are temperate and tropical soils and oxidation to carbon monoxide.

4. NMHCs: This group includes a complex set of hydrocarbons with highly varying characteristics. Most of these are chemically active and have short lifetimes. The usual concentrations in the atmosphere are only few ppbv with localized peaks occurring near the sources. These compounds are removed from atmosphere usually by atmospheric photochemical reactions.

5. Particulate organic carbon (POCs): These complex mixtures of hydrocarbons, alcohols, esters and organics in particulate form. These are usually produced from secondary reactions (gas to particle conversions) and are important in cloud and precipitation processes. The concentrations of POCs in marine air may be around 0.1 to 0.5 g m-3 and in background continental air may be around 1.0 g m-3. In general, the composition of POCs has about 60% neutral compounds, 30% acids and 10% bases.

6. Elemental carbon: This comes into the atmosphere exclusively form biomass and fossil-fuel combustion. Its typical atmospheric concentration over continents is 0.02 g m-3. It is present as fine black powder and can be used as excellent tracer substance for studying long-range transport phenomena in atmosphere.

In addition to above forms, carbon is also present in the atmosphere as carbonyl sulfide, carbon disulfide and dimethyl sulfide. These compounds are important in sulfur-loading of atmosphere and have been discussed with atmospheric sulfur.

Table-2: Indicative characteristics of primary carbon compounds in atmosphere.


Major sources


(Tg Y-1)








Oceans, biosphere,

fossil fuels

7.6 x 104

350 ppmv

380 ppmv

5 years



Biomass burning,




<50 ppbv

150-200 ppbv




to CO2


Animals, wetlands,

decay of vegetation


1650 pptv

>1800 pptv

10 years

Oxidation to

CO, soils


Vegetation, human



few ppbv






Secondary atmospheric



0.1 g m-3

>2.0 g m-3

1 week

Wet and dry




Biomass burning


0.2 g m-3

>1.0 g m-3

1 week

Wet and dry



Nitrogen is primarily exchanged between atmosphere, biosphere and soil. Following Table-3 shows the estimated total stored in the atmosphere and surface locations on a global scale.

Nitrogen in hydrosphere

In comparison to biosphere or atmosphere, very little nitrogen is present in oceans and continental surface waters. Over 95% of nitrogen stored in oceans is present in inactive molecular form. Only nitrate (about 2.5% of total oceanic nitrogen) and organic matter (about 1.5% of total oceanic nitrogen) have some active role. Oceanic nitrogen comes through river runoff from continents and wet and dry deposition from atmosphere. Its loss occurs through deposition to sediments in the bottom of oceans and through release to atmosphere in areas of biological activity. Nitrogen content in ocean water can vary spatially; for example, ammonia in surface oceanic waters varies between 0.05 to 2.0 mmol m-3 with smallest concentrations in the open oceans where biological activity is lowest. The amount of nitrogen released from oceans to the atmosphere (about 0.5 Tg Y-1) is quite low in comparison to that from other sources.

Table-3. Nitrogen storage in various

components of global environment.


Nitrogen storage

in Tg (1012 g)


2 to 6 x 106


85 x 103



10 x 103


3.8 x 103

Surface litter

1.5 x 103

marine biomass




Human beings


Nitrogen in rocks

The amount of nitrogen stored in lithosphere is much greater than the amounts stored in all other locations combined together. In lithosphere, most of the nitrogen is stored in primary igneous rocks and thus is not available to ecosystem. Weathering and other natural processes release only a very small fraction (<<1%) of this stored nitrogen into global ecosystem.

Nitrogen in soil and biosphere

Major active zone of nitrogen use and transfer occurs in the soil and biosphere on continents with very minor activity occurring in aquatic ecosystems. Inactive N2 of atmosphere is converted to form available to ecosystem through the process of nitrogen fixation, which mainly involves bacterial activities (though some nitrogen fixation also occurs during atmospheric lightening). Fixed nitrogen is made available first to plants in the ecosystem through mineralization to ammonia or through oxidation of reduced ammonia to nitrate (NO3). This process termed nitrification occurs under aerobic conditions. The oxidized nitrogen in soil is returned to atmosphere through the process termed denitrification under anaerobic conditions.

Nitrogen content of soil determines the nitrogen availability to biosphere and various soil types differ in their nitrogen content. Most of the soils contain about 0.05% to 0.2% nitrogen by weight though richest organic soils may contain upto 0.5% of total mass as nitrogen. During rains, some of the soil nitrogen is leached by runoff or infiltration and reaches groundwater or river water to be transported elsewhere.

Nitrogen entering the plants mainly as nitrate or ammonium is assimilated there into a variety of organic nitrogenous compounds, mainly the proteins and amino acids which are passed on from plants to animals as food. Nitrogen then traverses to different trophic levels in the ecosystem as different animals eat each other. Finally, nitrogen is returned back to soil or atmosphere from the biosphere after death and decay of plants and animals. In the ecosystem, aerobic processes form NO2 also while anaerobic processes produce NO, N2O and N2. Most of these products is released to atmosphere.

All the processes and pathways involved with nitrogen cycle depend on the environmental conditions such as soil pH, water content, soil type etc. Temperature is crucial factor in nitrogen cycle because biological activity is highly sensitive to temperature.

Though nitrogen fixation is the natural source of biospheric nitrogen, nitrogen fertilizers added to soil and surface deposition of nitrogenous materials that are emitted into atmosphere by human activities have also become important inputs to biospheric nitrogen.

Nitrogen in atmosphere

Nitrogenous species important in global nitrogen cycle found in atmosphere are:

1. Molecular nitrogen: The N2 gas constitutes about 79% of air by volume and it provides the main source of nitrogen to biosphere through nitrogen fixation as discussed above.

2. Ammonia and ammonium: Ammonia is very important component of nitrogen cycle as it is the only water-soluble gaseous nitrogen species. It can directly act as plant nutrient being converted to ammonium (NH4+) which forms the atmospheric nitrogen aerosol component. About 54 Tg nitrogen is emitted to atmosphere per year and ammonia released from animal urea makes up about half of this. Nitrogen inputs through biomass burning depend on the nitrogen content of the biomass which differs in different ecosystems. Average nitrogen content of tropical forest wood is 0.45%, of tropical litter is 0.85%, of coniferous and deciduous forest wood is 0.32%, of fuel wood is 0.2% and of tropical grasses is 0.2% to 0.6%. Other minor sources include coal combustion, human excreta and fertilizers.

It is difficult to establish the global representative concentrations of ammonia and ammonium. Ammonia concentration is lowest over remote oceans (about 0.1 ppbv); while in continental background air it is 6-10 ppbv. The ammonia concentrations are higher in summers than in winters and during daytime than in night due to higher temperatures influencing the activities of soil-based microbial sources. The lifetime of ammonia is only about 6 days and so it is rapidly converted to ammonium, which is the major component of two most prevalent atmospheric aerosols, ammonium sulfate and ammonium nitrate. Concentrations of both these aerosols and the gas decrease exponentially with altitude. Major sink of these aerosols is wet and dry deposition that removes about 49 Tg of nitrogen per year from atmosphere.

3. Nitrous oxides: Apart from N2, nitrous oxide (N2O) is the other inert gas in the atmosphere. Its lifetime is about 179 years and its major sink is photochemical reactions in stratosphere. It is also a greenhouse gas. Major sources of N2O emission are soil and oceans through microbial processes. Highest concentrations of the gas over oceans occur in areas where strong upwelling brings deep-water nutrients to the surface waters. Emissions due to human activities are adding about 8% of the natural input. N2O emissions increase with higher temperature and moisture and, therefore, reach a daily maximum around noon and seasonal maximum in summers. Emissions can be greatly increased on a local scale by irrigation practices. The gas shows very little variation in global distribution due to its long lifetime and major natural sources. Depending on the photochemical activity, the concentration of gas decreases slightly with altitude in the troposphere.

4. Nitrogen oxide species: NO and NO2 are major part of a series of highly active primary and secondary compounds (including HCN and N2O5). Primary emission occurs mainly of NO which is rapidly converted to NO2, which thus becomes dominant in the atmosphere. Both these are quite short-lived species and are rapidly oxidized to nitrate aerosol or sulfuric acid. Both the gases are crucial in tropospheric and stratospheric ozone chemistry and in the chemistry of photochemical smog.

NO and NO2 are strongly influenced by anthropogenic emissions. Over 60% of nitrogen oxides come from combustion of fossil fuels and biomass. The amount of gases released from fossil-fuel combustion depends on the temperature of combustion process and nitrogen content of the fuel. Nitrogen content of coal is 1-2%, of crude oil is <1% and of natural gas is 5-10%. Concentrations of nitrogen oxides show high spatial variability during their short lifetime indicating that local and regional sources are highly important to their global budget. Natural sources of these oxides are soil and thermal dissociation of atmospheric N2 during lightening. Global emission of nitrogen oxides is about 50 Tg Y-1, which forms about 33% of total nitrogen, input into the atmosphere. About 43 Tg nitrogen is removed from atmosphere per year. This removal involves almost entirely the wet and dry deposition with a very small quantity lost to photochemical reactions. Concentrations of nitrogen oxides in clean ocean air in the troposphere are <100 pptv. Concentrations in rural air over the continents are 200-300 ppbv and in air influenced by human activities may be >10 ppbv reaching upto 500 ppbv in urban air. Highest concentrations are found in Northern Hemisphere around 400 N latitude where major anthropogenic sources of these oxides are located. Concentrations rapidly decrease with altitude to a background value of 10 pptv in the upper troposphere. Higher concentrations occur in winters, particularly in the mid-latitude areas under urban influence since temperature inversions are more prevalent and photochemical activity is at a minimum.

Table-4. Indicative characteristics of major atmospheric nitrogen compounds.


Major sources



(Tg Y-1)

Background concentration







soils, biomass burning


0.1 ppbv

>6.0 ppbv

6 days

Conversion to NH4


Conversion from NH3


0.05 g m-3

>1.5 g m-3

5 days

Wet & dry deposition


Secondarily from NOx


0.5 g m-3

>10.0 g m-3

5 days

Wet & dry deposition




310 ppbv

170 days

Strato-spheric photo-chemistry


Fossil fuels, lightening, biomass burning, intercons-versions


<100 pptv

100 pptv

<2 days

Oxidation to HNO3 & NO3, photo-dissociation

Sulfur cycle

Most of the sulfur on Earth is stored in oceans (about 1.3 x 106 Pg), sedimentary rocks (about 2.7 x 106 Pg) and evaporites (about 5 x 106 Pg). Very small percentage reaches the surface and is exchanged with atmosphere. Accuracy of the natural emissions of sulfur is about 50% only.

Sulfur in lithosphere

Sulfur is 13th most abundant element in Earth’s crust (0.1%) and 9th most abundant in sediments. Sulfur content of rocks varies considerably e.g. sedimentary rocks have about 0.38% while igneous rocks have only 0.032%. Sulfur in lithosphere is mobilized by slow weathering of rock material. Dissolved in runoff, it moves with river-water and is deposited in continental shield sediments in oceans. Eventually on geological time-scale, this uplifts to surface again thus completing the geological part of the sulfur cycle.

Sulfur in hydrosphere

Main storage of sulfur in oceans is through dissolved sulfate, averaging about 2.7 g per kg. Most volatile sulfur compound in sea water is dimethyl sulfide (DMS; (CH3)2S) which is produced by algal and bacterial decay. Its concentration in sea water is about 100 x 10-9 L-1, highest concentrations being in coastal marshes and wetlands.

Sulfur is second most abundant compound in rivers with concentrations fluctuating highly with seasons and frequency of drought, flood and normal flow. Rivers transport about 100 Tg of sulfur per year to the oceans. The storage of main sulfur mass in oceans, sedimentary and evaporite rocks establishes the base for sulfur cycle.

Sulfur in soil and biosphere

Sulfur is major essential nutrient in the biosphere and is concentrated mainly in soil from where it enters biosphere through plant uptake. From soil, sulfur is also removed in solution to groundwater and by chemical volatilization. Its main sources are deposition from atmosphere, weathering of rocks, release from decay of organic matter and anthropogenic fertilizer, pesticides and irrigation water. In soil, it is present mainly in oxidized state (e.g. SO4) with concentrations varying according to the amount of organic matter in soil. Rich organic soils may have upto 0.5% sulfur by dry weight.

Sulfur in soil may be in bound or unbound form, as organic or inorganic compounds, organic sulfur being most prevalent. Plants take up sulfur from the soil mainly as sulfate and it is passed on with the food chain in the biosphere. It leaves biosphere on death of living organisms when aerobic decay and decomposition brings back sulfate in the soil. Finally, anaerobic decomposition in soil releases part of organic sulfur as H2S, DMS and other organic compounds into the atmosphere. About 7 Tg of sulfur per year is released from global soils, with considerable latitudinal variation. The release of sulfur is dependent upon warmer temperatures.

Sulfur in atmosphere

Several sulfur compounds are released into the atmosphere due to interaction of processes between Earth’s surface and the atmosphere. Of these, most important six compounds are discussed below.

1. Carbonyl sulfide (COS): It is the most abundant sulfur species in atmosphere and in nature is mainly produced by decomposition processes in soil, marshes and wetlands along ocean coasts and areas of ocean upwelling that are rich in nutrients. Anthropogenic combustion processes produce less than 25% of COS. Its average concentration of about 500 pptv shows enough uniformity throughout latitudes and altitudes to suggest a long lifetime and no rapid sinks of this compound. A lifetime of 44 years is suggested with only sink being stratospheric photolysis and slow photochemical reactions in troposphere. Ocean may act both as source and sink. About 80% of total atmospheric sulfur is COS, but it is relatively inert and does not add much to atmospheric sulfur pollution problem.

2. Carbon disulfide (CS2 ):It is far more reactive than COS and has similar sources though on a smaller scale. It has lifetime of 12 days only and its major sink is photochemical reactions. As a result, CS2 shows greater spatial variation across the globe, ranging from 15 pptv in clean air to 190 pptv in polluted air. Its concentration decreases rapidly with altitude. The most important source of the compound is microbial processes in warm tropical soils. Major secondary sources are marshes and wetlands along sea coasts. Small anthropogenic inputs are from fossil fuel combustion.

3. Dimethyl sulfide (DMS): It is released from oceans in much greater amounts than COS or CS2 and has extremely small lifetime and is very rapidly oxidized to sulfur dioxide or is redeposited to oceans. In the sulfur cycle, most of natural gas released from oceans is DMS. Its concentrations are high during night, particularly in areas under some influence from continental sources.

4. Hydrogen sulfide (H2 S): It is mainly produced in nature during anaerobic decay in soils, wetlands, salt marshes and other areas of stagnant water with maximum concentrations occurring over tropical forests. This highly reactive is removed by reaction with hydroxyl radical (OH) and COS. Its highest concentrations occur at night and in early morning when photochemical activity is at a minimum.

4. Sulfur dioxide (SO2 ): Its natural source is oxidation of H2S and major anthropogenic source is combustion of fossil fuels. Its atmospheric concentrations are most influenced by anthropogenic emissions. In some industrialized areas such as eastern North America, over 90% of SO2 is from anthropogenic sources. Normally about half of global SO2 originates from natural sources. The lifetime of the gas is 2-4 days indicating that loss due to photochemical conversion to sulfate is quite important. Rest of the gas (about 45%) is removed from atmosphere by wet and dry deposition.

5. Sulfate aerosol: Sulfate aerosol particles originate from sea spray that is the largest natural source of sulfur to the atmosphere. Only 3 TG per year of sulfate is added to atmosphere from anthropogenic sources directly but much greater amounts are formed through secondary reactions from various sulfur species in atmosphere. Most of the salt spray sulfate falls back to oceans but some is carried over the continents to be included in deposition processes there.

Table-5. Indicative characteristics of major tropospheric sulfur compounds.


Major sources



(Tg Y-1)

Background concentration







coastal marshes, biomass burning


500 pptv




slow photoche-mistry, stratosphere, oceans





15-30 pptv





Photoche-mical production of SO2



algal deposition


<10 pptv





Oceans, oxidation to SO2


Bacterial reduction, soils,



30-100 pptv








sources, volcanoes, oxidation

of H2S


24-90 pptv

>5 ppbv



Wet & dry deposition


Sea-sprays, oxidation

of SO2


0.1 g m-3

>2.5 g m-3



Wet & dry deposition


Filed under: Environment — gargpk @ 1:21 pm
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The atmosphere of Earth comprises air envelop over the Earth’s surface extending to several kilometers into the space. Air is a mixture of gases and is held to the Earth by gravita­tional attraction. The density of atmosphere is maximum at sea level and decreases rapidly upward.

The development of atmosphere is closely related to the geological and geochemical processes and to the activities of living organisms. Initially the Earth did not possess atmosphere and it was created later in the course of Earth’ evolution. Important components of atmosphere i.e. nitrogen, carbon dioxide and water vapor arose in course of volcanic activities which brought these from the depths of lithosphere. Atmospheric oxygen was added later in Earth’s history as a result of photosynthetic activity of living organisms. In its turn, atmosphere has exerted a great influence on the evolution of lithosphere. Throughout Earth’s history, atmosphere has played important role in physical and chemical erosion of rocks. Winds, atmospheric precipitation and changes in atmospher­ic temperature and other atmospheric factors have been major factors in physical erosion of rocks while atmospheric oxygen and moisture have been extremely important in chemical erosion of rocks.

Evolution of atmosphere has played significant role in the evolution of hydrosphere since water balance of water bodies is influenced directly by precipitation and evaporation processes governed by atmospheric characteristics. On the other hand atmos­pheric processes have also been influenced by the state of hydro­sphere, and especially by the state of the oceans. Generally, the evolution of atmosphere and hydrosphere represents a single process.

Atmospheric factors have also played important role in influencing the evolution of biosphere. In the beginning the atmosphere had no oxygen and only anaerobic organisms could evolve. These organisms evolved the mechanisms by which they could split up water and release oxygen. As the oxygen accumulated into the atmosphere, the earlier reducing atmosphere turned into an oxidizing one. Only after sufficient accumulation of oxygen in atmosphere, aerobic organisms could evolve. Thus evolu­tion of complex living organisms is linked to increase in oxygen content of atmosphere which lead to the development of aerobic processes needed for the energetics of higher organisms. Carbon dioxide content of atmosphere is most important factor in activi­ties of autotrophic plants. Changes in its content during the course of evolution of atmosphere have exerted great influence on the structure and Earth’s plant cover. Further, the nature and characteristics of biotic communities depends on climatic condi­tions which are mainly governed by atmospheric factors.

Thus the atmosphere is very important component of the global environment and the important atmospheric features with a view to the study of global environment are:

1. Atmospheric stratification

2. Atmospheric gases and aerosols

3. Air pressure, winds and global air circulation

4. Atmospheric moisture and precipitation


The Earth’s atmosphere shows quite well defined layers one above the other. These layers are defined mainly on the basis of temperature. Broadly, the pattern consists of three relatively warm layers near the surface between 50 and 60 km and above 120 km separated by two cold layers between 10 and 30 km and about 80 km. Major layers in the atmosphere are as discussed below.

1. Troposphere

It is the lowermost layer of atmosphere where weather phe­nomena and atmospheric turbulence are most marked. This layer contains about 75% of total molecular or gaseous mass of atmos­phere and virtually all the water vapor and aerosols. Temperature in this layer decreases with altitude at a mean rate of about 6.5o C per kilometer. At the top the troposphere is limited and separated from the next higher layer by a layer called tropo­pause.

Tropopause is a temperature inversion level i.e. it is a layer where a layer of relatively warm air is present above a layer of colder air. This inversion level acts as a lid over most of the top of troposphere. As a result troposphere is largely self-contained and convection in it is effectively limited. The altitude at which tropopause is present is not constant but seems to be correlated with sea-level temperature and pressure which are in turn related to factors of latitude, season and daily changes in surface temperature. The altitude of tropopause varies from about 16 km at equator where heating and vertical turbulence are greatest to only about 8 km at poles. Thus troposphere ex­tends from ground surface to the altitude of 8-10 km at high latitudes to 16-18 km in equatorial zone. Physical processes within the troposphere determine the changes in the weather and exert a major influence on the climat­ic conditions in different regions of our planet. These processes include the absorption of solar radiation; the formation of fluxes of long-wave radiation, which is dissipated into outer space (and which changes in the higher layers of the air); and the water exchange that is associated with the formation of clouds and with precipitation

2. Stratosphere

This layer extends from tropopause to the altitude of about 50 km and contains most of the atmospheric ozone. Peak density of ozone in stratosphere is approximately at 22 km altitude. The ozone in this layer is responsible for the absorption of ultra-violet wave-lengths of solar radiation. Due to very low density of air such absorption results in large increase in the tempera­ture within this layer. The stratospheric temperature fairly generally rises with height in summers while the thermal struc­ture of this layer in winters is more complex. Thus marked sea­sonal changes of temperature affect the stratosphere. At the top stratosphere is limited by stratopause which has highest tempera­ture which may exceed 0o C.

3. Mesosphere

This zone of atmosphere extends from stratopause upward to about 80 km altitude. The temperature in this layer again de­creases from about 0o C at stratopause level to average of -90o C at the top of mesosphere. Air pressure in this layer is very low, decreasing from about 1 mb at 50 km altitude to 0.01 mb at 90 km altitude. Temperature again begins to increase above 80 km alti­tude. The temperature inversion level at 80 km altitude i.e. at top of mesosphere is termed mesopause.

4. Thermosphere

Above the mesopause, the thermosphere is the zone of ex­tremely low atmospheric density. This layer extends from 80 km to 100 km altitude. Molecular and atomic nitrogen and oxygen are the main constituent of this zone. The temperature in thermosphere increases with altitude owing to absorption of extreme ultra-violet radiation of 0.125-0.205 um wave-length by molecular and atomic oxygen.

5. Ionosphere

This layer is the region of high electron density extending between altitudes of 100 km and 300 km. Above 100 km altitude, this atmospheric zone is increasingly affected by cosmic radia­tion, solar X-rays and ultra-violet radiation. These radiations cause ionization of oxygen atoms and nitrogen molecules separat­ing the electrons from them. The frequency of ionized particles continues to increase upward alongwith the increase of tempera­ture in ionosphere.

6. Exosphere and Magnetosphere

Above the ionosphere, ions of oxygen, hydrogen and helium form the tenuous atmosphere generally called exosphere. In this zone the natural gas laws cease to be valid. Since natural atoms of hydrogen and helium are low molecular weight atoms, the atoms of these two elements escape from this zone into outer space. The escaped hydrogen is continuously replenished by breakdown of water and methane molecules near the mesopause. The escaped helium is similarly replenished through its formation by the action of cosmic radiation on nitrogen and from slow breakdown of radioactive elements in Earth’s crust.

The magnetosphere is the zone containing only plasma of electrically conducting gases. Charged particles are concentrated in two bands at altitudes of about 3000 km and 16,000 km. These zones form the Van Allen radiation belts or plasmosphere. The behavior of plasma particles in magnetosphere and the ‘precipita­tion’ of high energy plasma particles into Earth’s atmosphere produces ionization in lower atmospheric layers.

The stratification of atmosphere into distinct layers having specific structural and functional properties is very important feature of atmosphere since it governs the heat budget of Earth-atmosphere system, air motion and meteorological conditions. All these together result in weather phenomena in the troposphere which in turn govern the spatial and temporal pattern of climates i.e. the distribution of different climatic regimes in different geographical regions and their seasonal variations on Earth.


1. Atmospheric gases

The air forming the atmosphere is a colourless, odorless mixture of gases. Normally air consists largely of nitrogen (78%) and oxygen (21%) by volume. Remaining 1 % of the volume of air is made of small quantities of several other gases among which ex­tremely important gas is carbon dioxide (0.03%) because it can absorb heat and thus has primary role in maintaining temperature of Earth-atmosphere system. In the zone of atmosphere extending from ground surface upto the altitude of about 50 km, all the gaseous constituents of air are perfectly diffused among one another so as to give the air definite physical qualities just as if it were a single gas.

The gaseous composition of atmosphere is characterized both by permanent and variable components. Apart from carbon dioxide, another extremely important variable component of atmosphere is water vapour. It is a colourless, odorless gaseous form of water which mixes perfectly with other gases of air. Most of the atmos­pheric water vapor is concentrated in the troposphere zone. Changes in water vapor content of atmosphere over space and time are determined by interaction between evaporation, condensation and horizontal movement of water in atmosphere. The degree to which water vapor is present in the atmosphere is designated as the humidity and is of tremendous importance in weather phenome­na. Condensation of atmospheric water vapor results in formation of clouds and fog while excessive condensation results in rain, storm, hail or sleet collectively termed precipitation. The atmospheric water vapor like carbon dioxide can absorb heat and, therefore, like carbon dioxide is extremely important in ‘green-house effect’ of atmosphere.

Ozone gas is also very important constituent of atmosphere because of its ability to absorb ultra-violet radiation. Most of the atmospheric ozone is concentrated in stratosphere. It is formed from oxygen in the atmosphere.

2. Atmospheric aerosols

Apart from gases atmosphere contains many types of extremely small and light matter suspended in it generally included under the term aerosol. The word aerosol includes a wide range of material that remains suspended for a period of time in the atmosphere and usually refers to small solid and liquid matter. Solid aerosols are usually defined as particles or particulates and are distinct from dust which includes large pieces of solid material (>0 m in diameter) which settle out of atmosphere due to gravitation after short period of suspension. While effects of dust are limited locally, smaller aerosols can be transported to long distances and affect air quality and climate on regional and global scales. Aerosols originate from two main sources and are accordingly termed primary aerosols or secondary aerosols.

(i) Primary aerosols: These include matter that has been swept into the atmosphere from the surface of Earth such as dry desert plains, lake beds and beaches, volcanic eruptions, forest fires, ocean surfaces, disintegration of meteors in atmosphere, biological sources (e.g. bacteria, pollen and (fungi) etc. About 90 percent of these aerosols are found in troposphere while they are also found in upper layers of atmosphere also. Primary aerosols of size 2.0-20.0 m are defined as coarse aerosols while those 2.0 m in diameter are defined as fine aerosols.

(ii) Secondary aerosols: These aerosols are formed after various types of chemical conversion processes in atmosphere which involve gases, other aerosols and atmospheric contents particularly the water vapor. Very little is know about the details of the chemistry of trace gases to aerosols. These aerosols are almost always less than 2.0 m in size at the time of their initial formation when they are at nucleation mode (<0.1 m) but grow rapidly to accumulation mode (upto 2.0 m). General age of a layer of these aerosols can be determined by the relative amount of nucleation versus accumulation sizes. The smaller aerosols coagulate rapidly and aerosols larger than accumulation mode are efficiently removed from atmosphere by wet and dry processes and depos­ited onto the Earth’s surface.

a) Sulfate aerosols: A large fractions of aerosols are sulfate aerosols. In the nucleation stage, liquid droplet of sulfuric acid grows rapidly to accumulation size and eventually forms a stable non-reactive particle containing sulfate. Most often eventual result is ammonium sulfate in ages aerosols or ammonium bisulphate. Typical concentrations of sulfate aerosols are :

Remote background area – 1-2 g/cubic meter

Non-urban continental areas – <10 g/cubic meter

Urban areas under anthropogenic influence – >10 g/cubic meter

b) Nitrate aerosols: Nitrate is another important component of aerosols and mainly comes from oxidation of nitrogen gas. Most common compound in fine aerosol range is ammonium nitrate. It is not as stable as ammonium sulfate and its concentration is con­trolled by the relative abundance of ammonium, nitrate, sulfate and the level of atmospheric temperature. Nitrate also exists in coarse aerosols as a reactive interchange between crustal ele­ments over the continents or sea salt (ammonium nitrate) over the ocean.

c) Other aerosols: Most other aerosols can be further classified into size components with their areas of impacts as given in the subsequnet Table-1.

Optical effects of aerosol particles

High concentration of particulate material in the atmosphere is responsible for the visible hazes. Suspended material can cause a range of rather unusual atmospheric phenomena such asblue moons, green suns and green flashes or arcs about the sun or moon.

The distances between aerosol particles are generally greater than 10-100 particle radii and with such distances, scattering of light by particles is incoherent. Therefore, optical effects due to atmospheric aerosol particles are explained by light scattering.

Table-1: Properties of miscellaneous aerosol particles present in atmosphere.


Class Size range m) Impact area


(i) Aerosol size

Aitken 0.005-0.1 Air electricity

Large 0.1-1.0 Suspended particulate

Giant 1.0-15.0 Suspended particulate

Dust >15.0 Gravitational fallout

(ii) Aerosol type

Small ions <0.001 Air electricity

Large ions 0.005-0.5 Atmospheric chemistry

Haze 0.08-2.0 Visibility, human respiratory problems

Mist & fog 1.0-20.0 Visibility, atmospheric chemistry

Cloud condensation

nuclei 0.05-5.0 Cloud processes

Main aerosol 0.5-5.0 Visibility,atmospheric mass chemistry, cloud processes, human respiratory problems


Reyleigh Law for unpolarized light applicable only to particles of radius <0.03 m implies that scattered intensity will be proportional to r6/4 where r = radius of particle and = wavelength of light. Blue colour of scattered light from sky is explained in terms of effective scattering at shorter wave-lengths as the scattered intensity is inverse function of wave-wavelength. Red colour of setting Sun is because light passes over a very long path through atmosphere and most of its blue region of spectrum is lost due to scattering. Spec­tacular sunsets after volcanic eruptions or bush-fires arise due to higher than normal concentrations of very fine particulate material in the atmosphere after such eruptions.


The atmosphere above the earth is not a static body of air. The air masses have different and definite patterns of movement and all such movements ultimately result in global air circula­tion which is most important phenomena from the point of view of climatic conditions in different regions, weather conditions and distribution of pollution. The global movement of air is inti­mately associated with changes in the air pressures at different places. Therefore, the present chapter deals with concepts asso­ciated with air pressure, winds and global air circulation.


The air being in gaseous state, is readily compressible and if we consider a vertical column of air, the air nearer the ground level is compressed more due to the greater weight of air mass above it. This greater compression of air nearer to ground surface results in greater air density. The higher density of air results in increased expandability i.e. greater air pressure. Thus we find a vertical distribution of air pressure; the air pressure is greatest at sea level and gradually decreases with height. This vertical distribution of air pressure is important in many aspects of atmospheric science. In addition of height, air pressure is also affected by temperature. Increase in temper­ature results in increase in air pressure. Since different re­gions of earth receive different amounts of solar radiation daily during different times and also yearly during different months, the air is heated to different levels in different regions. This results in different air pressures in different regions i.e. horizontal distribution of air pressure over the globe. This horizontal distribution of air pressure is particularly important from the point of view of the origin, direction and velocity of winds.

The air pressure is measured by barometer which may be mer­cury barometer or aneroid barometer. The unit representing air pressure is either inches or millimeter of mercury or milibar (mb) which is a unit of pressure equivalent to a force of 100 dynes per square centimeter. The standard air pressure at sea level is 29.92 inches = 760 mm of mercury or 1013 milibar (mb).

Isobaric maps

Pressure conditions can be shown on map by means of isobars, which are lines connecting all the places that have same baromet­ric pressure. On the daily weather map, which shows conditions for a specific time only, the isobars are essential in showing the location of moving centers of high or low air pressures. On climatic maps the isobars show average air pressures which have been computed from the data accumulated over the years. First attention shall be paid to the average world conditions of air pressures.

World pressure belts

Major air pressure belts found on earth’s globe are:

(a) Equatorial trough: In the general vicinity of equator, there is a broad zone of somewhat lower than normal pressure (1013 and 1009 mb) which is known as equatorial trough.

(b) Subtropical high-pressure belts: On the north and south of this equatorial trough, there are subtropical belts of high air pressures. These belts are centered on about latitudes 30 degrees North and South. In the Southern hemisphere, this belt is clearly defined. In the North­ern hemisphere, this belt is broken into two oceanic centers or cells, one over the eastern Pacific and other over the eastern North Atlantic. High pressure at these latitudes is the result of convergence of air at higher levels and is accompanied by a general subsidence of the air.

(c) Sub-polar low-pressure belts: Extending from latitudes 45 degrees North and South to ice-covered North and South polar centers respectively are two broad belts of low pressure. In Southern hemisphere, there is a well developed subpolar low-pressure belt extending over the continuous expanse of southern ocean. The low pressure in these high latitudes in both the hemispheres is the result of numerous intense storms, each of which is a moving low-air pressure cen­ter. The pressure belts shift seasonally through several degrees of latitude alongwith the isotherm belts accompanying them. These seasonal shifts are important in explaining the world climates.

Northern hemisphere pressure centers

In the Northern hemisphere, the belted arrangement typical of Southern hemisphere is absent. This is due to the powerful influence that the vast land masses of Northern America and Asia separated by North Atlantic and North Pacific oceans exert over the pressure conditions in the Northern hemisphere.

Land areas develop high-pressure centers at the same time when winter temperatures fall far below those of adjacent oceans.Land areas develop low-pressure centers in summers when land surface temperatures rise sharply above temperatures over the adjoining oceans. Ocean areas show centers of pressures opposite to those on the lands, as seen in the January and July isobaric maps. In winters, the pressure contrasts as well as the thermal contrasts are greater. Over north central Asia, there develops Siberian high with pressure average exceeding 1036 mb. Over the central North America, there develops a clearly defined but much less intense center of high pressure, called the Canadian high.

Over the oceans, there are found Aleutian low and Icelandic low, named after the localities over which they are centered. These two low-pressure areas have much cloudy, stormy weather in win­ter, whereas the continental highs characteristically have a large proportion of clear, dry days.

In summer, pressure conditions are exactly opposite of winter conditions. Asia and North America develop lows, but the low in Asia is more intense. It is centered in southern Asia where

it is fused with the equatorial low-pressure belt. There are two well developed cells of the subtropical belt of high pressure over the Atlantic and Pacific oceans. These high pressure cells are shifted northward of their winter position and are considera­bly expanded. These cells are termed Bermuda high and Hawaiian high respectively.


Wind is simply defined as air in motion. Local winds are produced on a local scale by processes of heating and cooling of lower air. Following two categories of local winds may be recog­nized.

(i) Katabatic winds: The first category includes local winds in hilly or moun­tainous regions, where on clear and clam nights, heat is rapidly lost by ground radiation. This produces a layer of cold, dense air close to ground. A component of the force of gravity, acting in the downslope direction, causes this cold air to move down the mountain sides, pouring like a liquid into ravines and thence down the grade of the larger valley floors. Mountain breezes of this origin are of a variety termed katabatic winds. Particular­ly strong, persistent katabatic winds are felt on the great ice caps of Greenland and Antarctica where the lower air layer becomes intensely chilled. Certain occurrences of severe blizzards in these regions are katabatic winds.

(ii) Convection winds: In the second category are included land and sea breezes, which affect only a coastal belt a few km in width. Heated during the day by ground radiation, the air over land becomes lighter and rises to higher elevations. Somewhat cooler air over the adjoining water then flows land-ward to replace the rising warmer air creating a pleasant sea breeze. At night, rapid cooling of the land results in cooler, denser air which descends and spreads seaward to create a land breeze. These daily alternations of air flow are parts of simple convection systems in which flow of air takes a circular pattern in vertical cross section. Land and sea breezes are limited to periods of generally warm, clear weather when regional wind flows is weak, but they form an important element of the summer climate along coasts.

Irrespective of whether there are pressure centers or belts, a pressure gradient always exists, running from higher to lower pressure. If isobars are closely placed, it indicates that the pressure gradient is strong and pressure changes occur rapidly within a short horizontal distance. Widely placed isobars indicate a weak pressure gradient. Most of the widespread and per­sistent winds of the earth are air movements set up in response to pressure differences. The pressure gradient force acts in the direction of pressure gradient and tends to start the air flow from higher to lower air pressure. Strong pressur gradients cause strong winds and vice versa. Calm exists in the centers of high pressures.

Coriolis force and geostrophic winds

If the earth did not rotate upon its axis, winds would follow the direction of pressure gradient. However, the rotation of earth upon its axis produces another force, the Coriolis force which tends to turn the flow of air. The direction of action of Coriolis force is stated in the Ferrels’s Law‚ which states that any object or fluid moving horizontally in the Northern hemi­sphere tends to be deflected to the right of its path of motion, regardless of the compass direction of the path. In the Southern hemisphere, similar deflection occurs towards the left of the path of motion. The Coriolis force is absent at the equator but increases progressively poleward. It should be noted especially that the compass direction is not of any consequence. If we face down the direction of motion, turning will always be towards the right hand in Northern hemisphere. Since the deflective force is very weak, it is normally apparent only in freely moving fluids such as air or water. Ocean currents patterns are, to some ex­tent, governed by it, and streams occasionally will show a tend­ency to undercut their right-hand banks in hemisphere. Driftwood floating in rivers at high latitudes in Northern hemi­sphere, concentrates along the right-hand edge of the stream.

Applying these principles to the relation of winds to pressure, the gradient force (acting in the direction of the pressure gradient) and the Coriolis force (acting to the right of the path of flow) reach a balance or equilibrium only when the wind has been turned to the point that it flows in the direction at right angles to the pressure gradient i.e. parallel with the isobars. The ideal wind in this state of balance with respect to the forces, is termed the geostrophic wind for cases in which the isobars are straight. In general, air flow at high altitudes parallels the isobars. The rule for the relation of wind to air pressure in the Northern hemisphere states that: Standing with back to the wind, the low pressure will be found on the left-hand side and high pressure on the right-hand side.

Between the ground level and altitude of about 2000-3000 ft., still another force modifies the direction of wind. This force is the friction of air with ground surface. This force acts in such a way as to counteract, in part, the Coriolis force and to prevent the wind from being deflected until parallel with isobars. Instead, the wind blows obliquely across the isobars, the angle being from 20 to 45 degrees.


The wind systems present on the earth’s surfaces may be categorized as following:

(1) Doldrums: In the equatorial trough of low pressure, intense solar heating causes the moist air to break into great convection columns, so that there is a general rise of air. This region, lying roughly between 5 degrees N and 5 degrees S latitudes was long known as the equatorial belt of variable winds and calms or the doldrums. There are no prevailing surface winds here, but a fair distribution of directions around the compass. Calms prevail as much as a third of the time. Violent thunderstorms with strong squall winds are common. Since this zone is located on a belt of low pressure, it has no strong pressure gradients to induce persistent flow of wind.

(2) Trade wind belts: In the north and south of the doldrums are the trade wind belts. These roughly cover the two zones lying between latitudes 5 degrees and 30 degrees N and S. These winds are the result of a pressure gradient from the subtropical belt of high pressure to the equatorial trough of low pressure. In the Northern hemi­sphere, air moving towards equator is deflected by the earth’s rotation to flow southwestward. Thus the prevailing wind is from the northeast and the winds are termed northeast trade winds. In the Southern hemisphere, deflection of moving air towards left causes the southeast trades. Trade winds have a high degree of steadiness and directional persistence. Most winds come from one quarter of the compass.

The systems of doldrums and trades shifts seasonally north and south, through several degrees of latitudes alongwith the pressure belts that cause them. Because of the large land areas of northern hemisphere, there is a tendency for these belts to be shifted farther north in summer (July) than they are shifted south in winter (January). The trades are best developed over Atlantic and Pacific oceans, but are upset in the Indian Ocean region due to proximity of the great Asian land mass.

(3) Winds of horse latitudes: Regions between latitudes 30 and 40 degrees in both hemi­spheres have long been called the subtropical belts of variable winds and clams or the horse latitudes. These coincide with subtropical high-pressure belts. However, these are not continu­ous belts and high-pressure areas are concentrated into distinct centers or cells located over the oceans. The apparent outward spiraling movement of air is directed equatorward into the east­erly trade wind system; poleward into the westerly trade wind system. The cells of high pressure are most strongly developed in the summer (January in Southern and July in Northern hemisphere). There is also a latitudinal shifting following the sun’s declina­tion. This amounts to less than 5 degrees in Southern hemisphere, but it is about 8 degrees for the strong Hawaiian high located in the north eastern Pacific.

Winds in these regions are distributed around a considerable range of compass directions. Calms prevail upto quarter of the time. The cells of high pressure have generally fair, clear weather, with a strong tendency to dryness. Most of the world’s great deserts lie in this zone and in the adjacent trade-wind belt. An explanation of the dry, clear weather lies in the fact that the high pressure cells are centers of descending air, settling from higher levels of the atmosphere and spreading out near the earth’s surface and the descending air becomes increas­ingly dry.

(4) Westerlies: Between the latitudes 35 and 60 degrees, both N and S, is the belt of westerlies or the prevailing westerly winds. Moving from the subtropical high-pressure centers towards the subpolar lows, these surface winds blow from a southwesterly quarter in the Northern hemisphere and from a northwesterly quarter in Southern hemisphere. This generalization is somewhat misleading because winds from polar direction are frequent and strong. More accurately, winds within the westerly wind belts blow from any direction of the compass but the westerly components are defi­nitely predominant. In these belts, storm winds are common cloudy days with continued precipitation are frequent. Weather is highly changeable.

In Northern hemisphere, land masses cause considerable disruption of the westerly wind belt but in Southern hemisphere, there is an almost unbroken belt of ocean between the latitudes 40 and 60 degrees S. Therefore, in Southern hemisphere the west­erlies gain great strength and persistence.

(5) Polar easterlies: The characteristic wind systems of the arctic and antarctic latitudes is described as polar easterlies. In the Antarctic, where an ice-capped mass rests squarely upon the south pole and is surrounded by a vast oceanic expanse, polar easterlies show an outward spiraling flow. Deflected to the left in Southern hemi­sphere, the radial winds would spiral counterclockwise, producing a system of southeasterly winds.


In the Northern hemisphere, continents of Asia and North America exert powerful control upon the conditions of atmospheric temperature and pressure. Since pressure conditions control winds, these areas obviously develop wind systems that are rela­tively independent of the belted system of earth’s surface winds which is very developed in the Southern hemisphere. These inde­pendent wind systems are termed monsoon winds.

In summer, southern Asia develops a center of low pressure, into which there is a considerable flow of air. This may be a heat low (thermal low)“ limited to the lower levels of atmosphere. Warm, humid air from the Indian ocean and southwestern Pacific moves northward and northwestward into Asia, passing over India, Indochina and China. This air flow is summer monsoon winds which is accompanied by heavy rainfall in southeast Asia.

In winter, Asia is dominated by a strong center of high pressure, from which there is an outward flow of air reversing that of the summer monsoon. This flow is the winter monsoon winds which blows southward and southeastward toward the equatorial oceans and brings dry, clear weather for a period of several months.

The North America is smaller in extent as compared to Asia, and so it does not have such remarkable extremes of monsoon winds as is experienced by southeast Asia. Nonetheless, North America also experiences an alternation of temperature and pressure conditions between winter and summer. In summer, there is a prevailing tendency for air originating in the Gulf of Mexico to move northward across central and eastern part of U.S.A. In winter, there is a prevailing tendency for air to move southward from sources in Canada.

The continent of Australia also shows a monsoon effect, but being situated south of equator, it exhibits conditions reverse to those in Asia.


The surface wind systems described above represent only a shallow basal air layer of a few thousand feet thickness, whereas the troposphere is five to twelve miles thick. Since 1945 much knowledge has been gained about the nature of air flow at higher levels in troposphere and weather maps of upper air conditions have been drawn. It has been found that high above, there are slowly moving high- and low-pressure systems but these are gener­ally simple in pattern with smoothly curved isobars. Winds, which may be extremely strong and follow the isobars closely, move counterclockwise around the lows (Northern hemisphere), but clockwise around the highs. In general or average pattern of upper air flow, two systems dominate:

(i) Westerlies: This is the system of winds blowing in a com­plete circuit above the earth from about latitude 20 degree almost to the poles in both hemispheres. At high latitudes these westerlies constitute a circumpolar whirl, coinciding with a great polar low pressure center. Towards low latitudes the pressure rises steadily at a given altitude, to form two high-pres­sure ridges at latitudes 15 to 20 degrees N and S. These are the high altitude parts of the subtropical highs, but are shifted somewhat equatorward. In the high-pressure zones, wind velocities are low, just as in the horse latitudes at sea level.

(ii) Equatorial easterlies: This second major global air circu­lation system is between the high-pressure ridges where there is a trough of weak low-pressure, in which the winds are easterly. At lower elevation their influence spreads into somewhat higher latitudes as the trade winds.


The upper-air westerlies tend to form somewhat serpentine and meandering paths, giving rise to slowly moving upper air waves‚ in which the winds are turned first equatorward, and then poleward. At altitudes of 30,000 to 40,000 feet, associated with the development of such upper air waves, are narrow zones in which wind streams attain velocities upto 200 to 250 miles per hour. This phenomenon is termed the jet stream and it consist of pulse-like movements of air following a broadly curving track. In cross section, the jet may be likened to a stream of water moving through a hose, the center line of highest velocity being surrounded by concentric zones of less rapidly moving fluid.

Most important function of upper air waves is that by means of these, the warm air of tropics is carried far north at the same time that cold air of polar regions is brought equatorward. In this way the horizontal mixing i.e. advection develops on a gigantic scale and serves to provide heat exchange between re­gions of high and low insolation.


The water evaporating from the surface of water bodies and also transpiring from the plants is held in the atmosphere as atmospheric moisture. The amount of water vapor held in the atmosphere at a given time varies widely from place to place. It ranges from virtually nothing in cold, dry air of arctic regions in winter to as much as 4 to 5 percent of the volume of atmos­phere in humid, hot tropical areas. The atmospheric water vapor returns to Earth’s surface in the form of precipitation which includes rain, snow, sleet and hail. This cycle of water from Earth’s surface to atmosphere and back to Earth constitutes very important part of global hydrological cycle. Further, this part of hydrological cycle is most important in creating and maintain­ing particular climatic conditions in different areas of Earth.

Important concepts related to atmospheric moisture and precipita­tion are discussed below.

Atmospheric humidity

The term humidity refers to the quantity of water vapor present in the air. For a given temperature, there is a definite limit to the quantity of moisture that can be held by the air. This limit is called saturation point. The actual total amount of water vapor present at a given temperature in a given volume of air is termed absolute humidity. The quantity of water vapor that can be held in a given volume of air increases with temperature i.e. absolute humidity is directly proportional to temperature.

The proportion of water vapor present relative to the maxi­mum quantity of water vapor that can be held at a given tempera­ture expressed as percentage is termed relative humidity at that temperature. At saturation point, relative humidity is 100%. Change in relative humidity can be caused in two ways:

(i) Addition of water vapor: When evaporation or transpiration adds water vapor to atmosphere, relative humidity increases. However, it is a slow process requiring that the water vapour diffuse upward through the air.

(ii) Decrease in temperature: Relative humidity can increase even without addition of water vapor to the air through a decrease in temperature because capacity of air to hold water vapor increases with decrease in temperature.

Dew point is that critical temperature at which air is fully saturated with the amount of water vapor present in it. If tem­perature falls below this point, condensation of atmospheric water vapor normally occurs.

When air rises or sinks in elevation, it undergoes changes of volume i.e. volume of air increases when it rise to higher elevation due to fall in atmospheric pressure and decreases when air sinks to lower elevation due to rise in atmospheric pressure. Due to this change in air volume with change in elevation, the value of absolute humidity can not remain a constant figure for the same body of air. Therefore, meteorology makes use of specif­ic humidity which is the ratio of weight of water vapor to weight of moist air (including water vapor). When a given air parcel rises or sinks in elevation without gain or loss of water vapor, the specific humidity remains constant. Specific humidity is often used to describe the moisture characteristics of a large mass of air. Its value ranges from 0.2 gm/kg for extremely cold, dry air over arctic regions in winter to 18-20 gm/kg for extreme­ly warm, moist air of tropical regions.

Another index of moisture used in meteorology is mixing ratio which is the weight of water vapor to weight of water vapor of dry air (excluding water vapor) stated in units of grams per kilogram. Mixing ratio commonly differs very little in actual numerical value of specific humidity.


When large masses of air are experiencing steady drop in temperature below dew point, condensation of water vapor occurs very rapidly within the clouds in atmosphere and precipitation occurs. This condition can not be brought about by simple cooling of air through loss of heat by radiation during night. Instead, rise of air to higher elevation is necessary. When air rises to higher elevation, its temperature drops even without loss of heat energy to outside. Because of drop of atmospheric pressure and increase in volume, air molecules strike each other less fre­quently and this imparts lower sensible temperature to air. In absence of condensation, the rate of drop of temperature is termed dry adiabatic rate and is about 5.5 degrees F per 1000 feet of vertical elevation. If water vapor in the air is condens­ing, latent heat is liberated which counteracts the temperature loss. The modified rate of adiabatic temperature loss in such condition, termed wet or saturation adiabatic rate becomes slightly lower to about 3 degree F per 1000 feet.

Water vapor does not necessarily condense in clean air even when the vapor pressure of water is many times greater than that required to form liquid water. The reason of this supersaturation condition of clean air is that the equilibrium vapor pressure over small droplets is much greater than that over plane surfaces (p ). Relation between radius of water vapor droplet and partial pressure of water vapor (pr ) above it is given by:

loge (pr /p ) = 2 M / L RTr

where, = surface tension of water (N/m); L = density of water (kg/cubic meter); M= molecular weight of water; R = universal gas constant (J/K/mol); T= absolute temperature in K; r = radius of droplet (m).

Presence of salts in water has pronounced effect on equilib­rium of relative humidity with a small water droplet .

Very small droplet of pure water requires that air be supersatu­rated for condensation to occur. Presence of even a small amount of salt lowers the water vapor pressure considerably and thus salt acts as condensation nucleus lowering the equilibrium rela­tive humidity remarkably. Removal of water vapor by condensation leads to fall in vapor pressure to below saturation level. In the cloud, the condition of supersaturation is maintained due to cooling of air rising to higher elevation. With cooling the relative humidity of air increases i,e, colder air becomes saturated at a lower water vapor content than warmer air, so the condition of supersatura­tion is maintained. Very high in a cloud, temperature may drop to below freezing point of water and so ice, snow, rain and hail can form.

Prior to actual precipitation clouds are formed in the atmosphere. These consist of tiny droplets of water 20-60 microns in diameter or minute crystals of ice. These are sustained in the atmosphere by the slightest upward movement of air. For the formation of cloud droplets, it is necessary that microscopic dust particles serve as condensation nuclei. Hydroscopic particu­late aerosols found in atmosphere serve as condensation nuclei. Precipitation occurs when condensation is occurring very rapidly in the clouds.


Formation of clouds and precipitation occurs only when large air masses rise to higher elevation. This rise of air masses occurs in three ways and accordingly precipitation is of three types:

(i) Convectional precipitation: Such precipitation results from a convectional cell which is an updraft of warmer air rising up because it is lighter than surrounding colder air. When bare land surfaces rapidly heated, it transmits radiant heat to overlying air. Air over warmer land is heated more than adjacent air and begins to rise in a tall column called thermal. As rising air cools adiabatically, it eventually reaches same temperature as the surrounding air and comes to rest. However, before coming to rest it may be cooled below the dew point and immediately conden­sation begins. The rising air column appears as cumulus cloud whose flat base shows the critical level above which condensation is occurring. Bulging ‘cauliflower’ top of this cloud represents the top of rising warm air column pushing into higher levels of atmosphere. If this convection column continues to develop, the cloud may grow into a cumulonimbus cloud mass from which heavy rainfall will occur. In most of the natural conditions, the unequal heating of ground serves only as a trigger effect to release a spontaneous updraft of air mass. Later it rises due to heating by release of latent heat from condensing water vapor. This heating causes air to continuously rise upward even during condensation. Such air is described as unstable air.

(ii) Orographic precipitation: This type of precipitation is related to mountains. Prevailing winds or other moving air masses may be forced to move over mountain ranges in some areas. As the air rises on the windward side of the range, it is cooled adia­batically. If cooling is sufficient, precipitation results on the windward side of range. Much orographic rainfall is actually of convectional type, in that it takes the form of heavy convection­al showers and storms. Storms are induced, however, by the forced ascent of unstable air as it passes over the mountain barrier.

(iii) Cyclonic precipitation: This type of precipitation occurs when air converges in cyclonic storms or eastward-moving centers of low pressure and is forced to rise resulting in cooling and condensation. Much of the precipitation in middle and high lati­tudes is of such type.

Global warming data tables

Filed under: Global warming — gargpk @ 1:15 pm

Table 1: Trace gas characteristics and their estimated future impacts on atmospheric warming.









Concentration (ppmv)


1985 A.D.

2030 A.D.



















Upto 25% less than now


Upto 12.5% more than now

Main IR absorption wavelength (m)







4.5, 7.6








Observed (%) increase 1975-1985








Approximate lifetime in atmosphere in years








IR trapping (W m-2)

















Estimated temperate increase (oK)








Table 2. Effects of differences in potential climatic changes from doubling of carbon dioxide concentrations in three global GCMs compared to present-day control.

(Data: Schlesinger and Mitchell, 1987)



Model NCAR

Model GISS

Model DFDL


Latitude average (0C)

Northern Hemisphere

Southern Hemisphere

Geographical Location

Northern Hemisphere

Southern Hemisphere


















Northern Hemisphere (0N)

Southern Hemisphere (0S)









Tropospheric min.

Maritime warming (0C)





Stratospheric max.

cooling (0C)





Precipitation, 300N-300S


More to S

Less to N

More to N

Less to S

More to S

Less to N

Indian summer monsoon strength

Tropical rain rate increase (mm/day)







Much more








Drier except N


Wetter except N


Dry in E wetter in W

DJF = December, January, February (Winter in Northern Hemisphere); JJA = June, July, August (Winter in Southern Hemisphere)

Table 3. Estimated average growth and yield changes in C3 species assuming a doubling of CO2 concentration. (SCOPE, 1986)



Immature crop

Mature crop

No. of records

% increase in biomass

No. of records

% increase in biomass

Fiber crop

Fruit crop

Grain crop

Leaf crop


Root crops

C3 weeds



Cucumber, eggplant, okra, pepper, tomato

Barley, rice, wheat, sunflower

Cabbage, white clover, fescue, lettuce, chard

Pea, bean, soybean

Sugar beet, radish

Jimson weed, pigweed, ragweed, sicklepod, other weeds


































Mean of all C3 plants





* = 95% confidence limits.

Note: C3 plants have the highest rate of photosynthesis and can assimilate carbon dioxide better than lower plants.

See site


Filed under: Environment — gargpk @ 1:02 pm
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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

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


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

  1. 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.

  2. Indirect influence: The indirect influence of local evaporation on totalprecipitation derives from the linkage between the volume of precipitation and the relative humidity atmosphere.

Table-4: Atmospheric water exchange over continents.


r (sq. km /yr)

ra (sq. km /yr)

rl (sq. km /yr)

















N. America





S. America










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.

Table-5: Precipitation, evaporation and run-off on continents & oceans.


Precipitation (cm/year)

Evaporation (cm/year)

Run-off (cm/year)













N. America




S. America








All land
















World ocean




(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.

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 R when R/Lr 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 = (R/Lr)

The form of the function for R/Lr 0 and for R/r 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.

  1. 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. 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. 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.


Filed under: Environment — gargpk @ 12:59 pm
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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 (Fig. 1):

(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:

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

(ii) Energy expenditure associated with friction of air currents, ocean waves produced by winds and ocean tides

Figure-1. Components of the energy balance of Earth-Atmosphere-Hydro-Lithosphere system.

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

(iv) Energy expenditure on photosynthesis

(v) 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 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.

Table-1. Average values of Earth’s surface energy balance components at various latitudes (kcal/sq. cm/year)

Latit-sude (in degr-ees)























70-60 N












60-50 N
























50-40 N
























40-30 N
























30-20 N
























20-10 N
























10-0 N
























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

Table 2. Mean latitudinal values of energy balance components of Earth-atmosphere system for six-month periods (kcal/sq. cm/year) April-September

Latitude (in degrees)






80-90 N












70-80 N












60-70 N












50-60 N












40-50 N












30-40 N












20-30 N












10-20 N












0-10 N












October- March

80-90 N












70-80 N












60-70 N












50-60 N












40-50 N












30-40 N












20-30 N












10-20 N












0-10 N












Table-3. Energy balance of continents and oceans (kcal/sq. cm/year)













































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.

March 9, 2008


Filed under: Climate,Environment — gargpk @ 3:00 pm
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The global climate is the result of innumerable interactions occurring amongst components of the global environment, chiefly amongst hydrosphere (particularly the oceans) and the atmosphere. These interactions result in particular meteorological conditions and spatial average fields of these conditions produce climatic regimes. The purpose of climate theory is to identify the average distribution of meteorological elements in space and time as well as their responsiveness to external factors. Numerical models of climate are devised for this purpose, which make it possible to calculate average fields of meteorological elements. With the development of high-power computers, it has also become possible to design numerical models that can reproduce non-average fields of meteorological elements and calculate such fields for prolonged time intervals thus allowing calculation of average fields describing climatic conditions. Various theoretical/numerical models have been developed in the past half a century. These allow description and/or study of present climate as well as changes in climatic conditions produced by natural or anthropogenic activities. For example, numerical model of climate given by Manabe and Bryan (1969) shows the influence of the circulation of the ocean waters on climatic conditions. The model given by Holloway and Manabe (1971) shows the distribution of basic components of heat and water balance at the Earth’s surface similar to that actually present on Earth. As has been pointed above, the numerical models of climate can be employed for the study of current climate as well as climatic changes. However, the models used to study climatic change have to meet more rigorous requirements than the models used to study current climatic regime. The important requirements for such models are:

  1. Model should not include empirical data concerning the distribution of individual elements of climate, particularly those that change substantially during the changes in climate.

  2. Model must recognize realistically all types of inflow of heat that influence the temperature field appreciably; in particular, the law of conservation of energy.

  3. Model must include major feedback relations among various elements of climate.

The third requirement is most important and has been discussed below.

Feedback Relations Among Elements Of Climate

Feedback relations among climatic elements are highly complex and interrelated. Such relations include both negative and positive feedback relations.

Negative feedback relationships

These relationships reduce the anomalies of meteorological elements and contribute to the approximation of the values of these elements towards their climatic normals. Thus, the climatic stability is maintained by negative feedback relationships among climatic elements. Major such negative feedback relationships are:

1. Long-wave radiation and temperature at Earth’s surface: Intensity of long-wave radiation increases with increase in temperature at Earth’s surface. This produces a greater expenditure of heat energy, which in turn inhibits further increase in temperature.

2. Heat transfer in atmosphere and air temperature gradient: The usual flow of heat into the atmosphere from a zone of higher temperature produces a smoothing process that eliminates the differences in temperature distribution.

Positive feedback relationships

Such relationships among climatic elements play major role in climatic change since these increase the anomalies in meteorological elements. Thus, positive feedback relationships reduce the climatic stability. Major such relationships are:

  1. Absolute air humidity and air temperature: Absolute air humidity increases with rise in air temperature. With rise in temperature, evaporation increases leading to comparative constancy of relative humidity in most climatic zones (except in dry continental regions). Increase in absolute air humidity decreases long-wave radiation. Thus, the increase in absolute air humidity with rise in temperature partly compensates the increased long-wave radiation attributable to increased temperature. Manabe and Wetherald (1967) showed that influence of the change in solar constant on air temperature at Earth’s surface is almost twice in the condition of constant relative humidity as compared to the condition when absolute humidity is stable. This particular feedback relationship is important in numerical models of thermal regimes employed in studies of climatic changes.

  2. Snow and ice cover and albedo of Earth’s surface: Positive feedback relationship between snow and ice covers on the albedo of Earth’s surface plays a very great influence on the patterns of changes in atmospheric thermal regimes. Ice or snow covers have high albedo and so reduce air temperature above them and climatic changes are intensified by formation and melting of ice. Available data shows that during summer months, albedo over ice cover in Central Arctic is about 0.7 while in Antarctica, it is about 0.8 to 0.85. In regions free of ice and snow, albedo of Earth’s surface does not exceed 0.15. This indicates that other conditions being same, snow and ice covers reduce the radiation absorbed by Earth’s surface by several times.

  3. Snow and ice covers and Earth-atmosphere system: The data shows that albedo of Earth-atmosphere system during summers in Central Arctic region is 0.55 and in Antarctic region is about 0.6. This is approximately twice the value of the estimated albedo for the planet as a whole which is about 0.33. Such large differences in values of albedo must exert considerable influence on the atmospheric thermal regime.

  4. Air temperature and Earth’s albedo: Ice and snow covers are created at the Earth’s surface due to reduction in air temperature. These covers cause a sharp decline in the absorbed radiation. This contributes to a further reduction in Earth’s temperature and consequently further increases the area under snow and ice. The reverse process may be effected by increases in temperature, which results in melting of ice and snow. Budyko (1968) showed that inclusion of this feedback relationship into a numerical model of atmospheric thermal regime invariably always exerts a very substantial influence on the distribution of air temperature at Earth’s surface. This influence can be shown by a simple example, which shows how average global temperature will change if Earth’s surface becomes fully covered by snow and ice and clouds in the atmosphere are absent. In such a condition, Earth’s albedo will change from its present value of 0.33 to the value for dry snow cover i.e. 0.8. Thi
    s increase in albedo will reduce the air temperature. The absence of clouds will further reduce the temperature of lower layers of atmosphere near Earth’s surface. In the present times, the average temperature in lower layers of atmosphere rises substantially almost everywhere at the Earth’s surface. This is due to greenhouse effect associated with absorption of long-wave radiation by water vapour and carbon dioxide present in the atmosphere. However, at very low temperature resulting in the formation of snow and ice covers at Earth’s surface, green-house effect would become insignificant and dense clouds which perceptibly change the radiation flows are not formed. Under such conditions, atmosphere will become more or less transparent to both short-wave and long-wave radiation. The average temperature of the Earth’s surface for such transparent atmosphere is determined by the formula

4/So(1- s) /4

where So = Solar constant; s = albedo of Earth’s surface; = Stefen’s constant.

The above formula shows that Earth’s average surface temperature at albedo value of 0.8 will be _ 87o C (186o K). Thus, the Earth’s average temperature will decline by approximately 100o C from its present value of 15o C if ice or snow were to cover the entire Earth even for a short period. Thus, the enormous influence of snow cover on the thermal regime of Earth becomes quite clear.

A number of studies have attempted to calculate the influence of sea polar ice on Arctic thermal regime. These studies are based on the available data of thermal balance in central regions of Arctic Ocean and approximate values of the proportions derived from a semi-empirical theory of climate. It has been established that polar ice reduces average air temperature in the Central Arctic during summer months by several degrees and by about 20o C in winters. It has been concluded from these studies that the Arctic Ocean could be free of ice in the present age. However, this state would be extremely unstable and it could develop an ice cover as a result of a relatively small change in climate.

Since a permanent ice cover exerts a substantial influence on the atmospheric thermal regime even when it covers only a small part of the Earth’s surface, this must be taken into account in studies of climatic changes.

Present-Day Climate

The climatic conditions of present century have been established on the basis of meteorological data collected from worldwide network of climatic stations. The data shows that elements of meteorological regime change perceptibly over time. These changes are both periodic (daily or yearly) oscillations as well as non-periodic oscillations of different time intervals.

Short interval non-periodic changes (of days or months) occurring in meteorological regime determine the oscillations in weather. These changes are not spatially homogeneous and are largely explained by the instability of atmospheric circulation. For longer time intervals (of several years), irregular oscillations of individual elements of meteorological regime occur along with long-term changes that are similar over large territories. Such changes characterize fluctuations in climate.

Since climatic fluctuations at present time are relatively modest, average values of meteorological elements over a period of several decades could be used in order to describe the climatic features of the present age. The use of such average values makes it possible to exclude the influence of unstable atmospheric circulation pattern. Following Table shows average temperature in January and July and also the average yearly atmospheric precipitation at various latitudes.

Data show that difference between average temperatures at Earth’s surface at various latitudes is almost 700 C. The Earth’s surface temperature is maximum at Equator and lowest at South Pole. Earth’s spherical shape exerts a substantial influence on the distribution of these temperatures by producing variations in the total radiation reaching the upper boundary of atmosphere. Further, permanent ice covers are found at high latitudes where air temperature does not rise above freezing point almost throughout the year. Apart from substantial changes in meridonial direction, average air temperature at Earth’s surface also changes substantially in most latitudinal zones at various longitudes. This is largely explained by the distribution of continents and oceans.

The influence of ocean’s thermal regime extends to a large part of the surface of continents on which maritime climate exists. This influence is characterized by relatively modest yearly oscillations in air temperature at middle and high latitudes. The amplitude of yearly temperature fluctuations increases sharply in those extratropical continental regions where influence of oceanic thermal regime is less pronounced, characterizing the continental climate.

The distribution of average latitudinal precipitation values produces a pattern in which the principal maximum value occurs in the equatorial zone, total precipitation declines at subtropical latitudes, two secondary maxima lie at middle latitudes and precipitation declines in polar latitudes. Changes in average precipitation at different latitudes are explained by the distribution of average air temperatures and by specific characteristics of atmospheric circulation. Other conditions being equal, total precipitation increases with temperature because it increases the volume of atmospheric water vapour. Vertical air currents that carry water vapour through condensation level producing clouds also have important role in precipitation.

The atmosphere’s overall circulation is closely associated with geographical distribution of stable pressure systems. Particularly important such systems are:

  1. low-pressure belt near the equator,

  2. high-pressure region at high tropical and subtropical latitudes

  3. the region of frequent cyclone-formations at middle latitudes.

Downward movements of air within high-pressure zones substantially reduce the precipitation. On the other hand, pronounced upward air movements increase the precipitation at equatorial latitudes and in several regions at middle latitudes.

Largest desert areas on Earth having negligible precipitation are found in the subtropical high-pressure zone. Total precipitation also declines in continental regions at middle latitudes, which are distant from ocea
ns because very small quantity of water vapour carried by air currents from oceans reaches these regions.

Thus in the continents, zones of humid climates are largely located at equatorial latitudes and in regions of maritime climate, at middle and high latitudes. Similarly, at high tropical and subtropical latitudes and in regions of continental climates, conditions of insufficient moisture prevail.

Table : Present-day latitude-wise air temperatures and precipitations

Latitude      January temperature (0C)        July temperature (0C)           Precipitation(cm/year)

90-80 N/S             -31/-13                                      1/-42                                        19/11

80-70 N/S             -25/-8                                        2/-30                                       26/25

70-60 N/S             -22/0                                       12/-12                                        52/67

60-50 N/S            -10/5                                        14/1                                           80/101

50-40 N/S             -1/12                                       20/8                                          75/108

40-30 N/S            -11/20                                     26/14                                         77/103

30-20 N/S             19/25                                     28/18                                         73/91

20-10 N/S             25/26                                     28/24                                      114/122

10-0 N/S               27/27                                     27/26                                       201/150

Stability Of Climate

Budyko (1968) first used a semi-empirical model of atmospheric thermal regime to study the single-valued character and stability of current global climatic regime. The study using distribution of average air temperatures among different latitudes showed that the present climate is not the only possible one for existing climate-forming factors. Aside from the existing climate, current external conditions may produce a climate corresponding to a ‘white Earth’ as well as other variants of climate. Relatively small changes in external climate-forming factors may greatly alter the existing climate. Numerous subsequent studies employing similar models of climate have confirmed this conclusion.

The study of Wetherald and Manabe (1975) about the stability of existing climate is especially interesting. Unlike studies using semi-empirical models of t
he distribution of average air temperatures among different latitudes, their study applied a three-dimensional model of a general theory of climate that includes a detailed consideration of dynamic processes occurring in the atmosphere. The model takes into account the influence of state transformations of water on thermal regime including the feedback relationships between the snow cover, polar ice and air temperature. The study showed that if solar constant increases by more than 2%, the average yearly air temperature shall increase by approximately 2o C at low latitudes, more at higher latitudes and by about 10o C at 80o N. The ice cover on Earth shall be reduced by such rise in average yearly air temperature. The conclusions of this study are similar to those obtained by using semi-empirical models of thermal regime of atmosphere. However, the calculations from such models show slightly larger changes in average yearly air temperature following increases in the solar constant by two percent i.e. 3o to 4o C at low latitudes and 12 to 14o C at 80o N.

Thus, it may be concluded that the contemporary climate of Earth is neither unambiguous nor highly stable. It is highly sensitive to small changes in inflows of heat arriving at the upper boundary of atmosphere.


Budyko, M.I. (1968) On the origin of Ice Ages. Meteorologiya I gidrologiya No. 11.

Holloway, J.L. & Manabe, S. (1971) A Global General Circulation Model with Hydrology and Mountains. Monthly Weather Review Vol. 99, No. 5

Manabe, S. & Bryan, K. (1969) Climate Calculation with a Combined Ocean- Atmosphere Model. Journal of the Atmospheric Sciences. Vol. 26, No. 4.

Manabe, S. & Wetherald, R.T. (1967) Thermal Equilibrium of the Atmosphere with a Given Distribution of Relative Humidity. Journal of the Atmospheric Sciences Vol. 24, No. 3.

Wetherald, R.T. & Manabe, S. (1975) The Effect of Changing the Solar Constant on the Climate of a General Circulation Model. Journal of the Atmospheric Sciences Vol. 32, No. 11.

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