In the ecosystem, autotrophic organisms (chiefly the green plants) use the energy of solar radiation to produce organic matter, which is used by all the living organisms, including autotrophic organisms themselves, in running their life activities. The organic matter is used by the living organisms through their respiratory activity. Out the total organic matter produced by photosynthetic activity of autotrophic organisms, a certain portion is consumed by these organisms themselves and the remaining organic matter is available in the ecosystem as the net organic matter production of autotrophic organisms. A relatively very small part of this net organic matter production in the ecosystem is directly transformed into mineral substances. This takes place without the participation of any other living organisms through the processes such as forest and prairie fires during which organic matter is transformed into carbon dioxide, water vapor and certain mineral compounds. Further, a still smaller portion of organic matter is deposited in the upper layers of lithosphere and at the bottom of water bodies in the form of coal, peat and other organic compounds. The remaining organic matter is now passed on to heterotrophic organisms in the ecosystem through various food chains. All the living organisms of a particular type in the ecosystem that receive organic matter as food in a particular manner constitute a trophic level. The organic matter received by a trophic level undergoes three fates:

1. A portion is consumed by that that trophic level itself though respiration in that trophic level

2. A certain other portion is passed on to next higher trophic level as organic food and

3. Remaining organic matter is stored in the trophic level as increase in the biomass of that trophic level (i.e. increase in the number of organisms of that trophic level).

From the point of view of ecosystem energetics, the organic matter that is received, passed on to next higher trophic level or stored by a trophic level represents the amount of energy received, passed on or stored by that trophic level. It is obvious that in a dynamically stable ecosystem, there can not be any storage of energy (i.e. organic matter) in any of its trophic levels. Therefore, in the dynamically stable global ecosystem, a very small portion of the net production of organic matter by autotrophs is stored in the abiotic components of the environment (i.e. lithosphere and hydrosphere) while major portion is consumed by heterotrophic organisms through their respiration.

The consumption of organic matter in a trophic level (including autotrophic organisms themselves) through the respiration in that level represents the loss of energy in that trophic level. It is a feature of global ecosystem that the flow of energy (represented by flow of organic matter as food) between trophic levels is associated with large losses of energy at each trophic level. The ratio of the amount of energy passed on from a trophic level to its next higher trophic level (_n) and the amount of energy received by that trophic level from its previous trophic level (_n1) is termed ecological efficiency () of that trophic level i.e.

Ecological efficiency () = _n/_n-1

The ecological efficiency of trophic levels, in general, is estimated to range between 10-20%. Such small general value of ecological efficiency indicates that biomass in each successively higher trophic level in the ecosystem is bound to be substantially reduced. Since ecological efficiency of a trophic level depends on the respiration of that level, smaller the value of ecological efficiency of a trophic level, greater is the consumption of organic matter through respiration (i.e. loss of energy) at that trophic level. As a result, there is greater reduction of biomass in that trophic level and in the next higher trophic level.

Nature of organisms and transformation of organic matter

Since intensity of metabolism per unit mass of a live organism usually increases with decrease in the size of organism, the biomass present at a specific trophic level in the food chain depends on the size of organisms of that trophic level. One of the causes of this relationship is that the metabolism depends substantially on the ratio of the rate of diffusion of gases through the surface and the mass of organism. This ratio increases as the size of organism decreases. Thus the rate of metabolism of a given unit weight of microorganisms is many times greater than that of macro-organisms. Further, metabolism also depends on the nature of physiological processes within the tissues of organisms. In wood of plants, the metabolism is usually much slower than in vertebrate tissue of similar size. These general principals largely determine the total biomass of various types of organisms in the global ecosystem.

The largest proportion of forests in the overall biomass of living organisms is due to the fact that autotrophic trees are located at the first link in the food chains and also due to the large size of individual trees. Together with specific properties of the wood, this feature substantially reduces the rate of metabolism per unit biomass in forests. Though the productivity of ocean phytoplankton is comparable with forests, small size of individual plankton organisms intensifies their metabolism per unit weight so much that the total mass of plankton on Earth is negligible in comparison with that of forests.

About 95% of the total biomass on Earth belongs to plants and rest to the animals. Biomass of aquatic organisms is substantially less than that of terrestrial organisms. Therefore, the distribution global biomass is largely determined by the distribution of terrestrial plant cover i.e. by the forest cover on continents. Considering that total biomass on Earth (global biomass) is approximately 3×10¹²tonnes and total productivity of plants on continents is approximately 140×10⁹tonnes, the time period of one cycle of organic matter for the plants on Earth comes to be approximately 20 years. This average figure relates to forests that constitute major portion of the biomass of plants on Earth. In other natural zones on continents, the duration of one cycle of plant organic matter is much shorter. The duration of this cycle in the oceans having phytoplankton is still shorter and appears to be only a few days.

The total biomass of animals is assumed to be approximately 10¹¹ tonnes. Assuming that the animals assimilate about 10% of the total productivity of plants, the average duration of one cycle of animal organic matter comes to be several years. However, the actual length of life of one generation varies widely in animal kingdom and the nature of the distribution of biomass among different animal groups is still not much clear.

Invertebrates are the largest components of animal biomass and among them, most important are organisms living in soil. The zoological mass of large animals per unit area on Earth is relatively quite low. Calculations of Huxley (1962) show that while in African savannahs, the biomass of large wild animals may be 15-25 tonnes/km², this figure is only about 1.0 ton/km² in middle latitudes, 0.8 ton/km² in tundra and 0.35 ton/km² in semi-desert areas.

Man occupies topmost position in the food chain on the Earth and consumes both the primary production of autotrophic plants and the biomass produced by many herbivorous and carnivorous animals. For the present size of human population of over 4.0 billion, its biomass is approximately 0.2×10⁹ tonnes. Assuming that each human being expends on average about 2.5×10³ kcal of energy per day, the total energy consumption of human population comes to be about 1.8×10¹⁵ kcal/year. Thus, the human population consumes about 0.2% of the total production of Earth’s organic world.

Average values of the productivity of natural plant covers of Earth have been derived by using various theoretical and numerical models and data from a variety of studies including empirical determinations of productivity in individual biogeographical zones.

Terrestrial plant covers

Yefimova (1979) has made use of quite precise relationships between productivity of natural plant cover and meteorological factors in calculating values of the productivity and the coefficient of utilisation of photosynthetically active radiation for each continent. Results of her calculations are given in the Table-1. The data shows that the average productivity per unit area for the five continents of Earth does not differ very much. In each of these continents, magnitude of productivity over large part of continental territory is greatly limited by insufficient moisture or heat. The continent of South America is exception to this general condition since climatic conditions over large part of its territory are favourable for plant life.

Table 1: Productivity and coefficients of utilization of photosynthetically active radiation in various continents of Earth. (Yefimova 1979)

In Australia and Africa, coefficients of utilization of photosynthetically active radiation are lower than average. This can be attributed to insufficient moisture over large parts of these continents, which inhibits the complete utilization of available radiation by plant covers.

In Europe and South America, most favorable conditions for the development of plant life are found. In Europe, located at higher latitudes and exposed to less solar radiation, its utilization is relatively greater.

Smil (1985) gave estimates of the productivity and storage of biomass in major biomes of the Earth. These estimates are given in Table-2. Data in this table shows that there is not much difference in the area occupied by different types of ecosystems except wetlands that occupy smallest area on the Earth. However, productivity is highest in cultivated lands where one ton of biomass is produced per one ton of phytomass, followed by tropical and temperate grasslands where 0.5 ton of biomass is produced by each ton of phytomass. Next in productivity are tundra, deserts-semi-deserts and wetlands where 0.2 tonnes of biomass is produced per hectare from one ton of phytomass per hectare. These areas are followed by wetlands and shrub-lands where productivity is 0.13 tonnes per hectare. Tropical, temperate and boreal forest, though occupy almost same area on Earth, produce 0.067, 0.04 and 0.02 ton of biomass per tone of phytomass per hectare respectively. Despite these facts, most important on Earth are tropical, temperate and boreal forests that have the highest concentration of biomass on Earth (totaling about 750 tonnes per hectare). These ecosystems also have the highest total storage of biomass on Earth totaling about 850 x 10⁹ tonnes. Further, it may be noted that contribution to total biomass production is equal for tropical rainforests and tropical grasslands (20 x 10⁹ t/yr), followed by boreal forests and tropical grasslands (15 x 10⁹ t/yr) and temperate forests, woodlands-shrub-lands and temperate grasslands (10 x 10⁹ t/yr). Tundra and deserts have quite high average of net biomass production per unit area and also quite high weight of phytomass per unit area. Despite this they contribute very little to total global biomass production (1.0 – 2.0 x10⁹ t/yr). However, if total biomass storage in different types of ecosystems on Earth is considered, tropical rainforests, temperate forests and boreal forests are the most important storehouses of organic matter on Earth having 850×10⁹ tonnes of biomass. Woodland and shrub-lands having 75×109 tonnes and then tropical and temperate grasslands having 60×109 tonnes of biomass storage follow these.

Table 2: Area, productivity and storage of major global ecosystems. (Smil, 1985)

Aquatic plant covers

There is much less data about productivity of autotrophic plant covers in water bodies as compared to that about terrestrial plant covers. However, the available data indicates that the seas and oceans have the greatest volume of organic matter produced by phytoplankton located in the 30-40 meters deep layer of hydrosphere. At greater depths, quantity of solar radiation is insufficient for active development of photosynthesis.

In general, the productivity of shelf zones is substantially less than open ocean. It may attain maximum values in small bodies of water possessing large quantities of minerals required by the plants. The overall value of productivity for the oceans is estimated to be about 55 billion tonnes per year i.e. approximately 15 centner per hectare. This last figure is less than 1/6^th of the average productivity per unit area on continents.

Thus the estimates show that the yearly volume of productivity for the Earth as a whole is approximately 200 billion tonnes i.e. about 40 calories per hectare. This corresponds to an energy expenditure of approximately 0.15 kcal/cm² per year. This is about 0.1% of the solar radiation reaching the Earth’s surface.

In the environment, organic matter is synthesized in its biotic component i.e. biosphere. Autotrophic organisms are the only organisms that can synthesize organic matter using solar radiation and mineral matter taken from atmosphere, hydrosphere and edaphosphere. Autotrophs synthesize organic matter either by photosynthesis or by chemosynthesis. While chemosynthesis is important for cycling of nitrogen and certain other processes in the environment, photosynthesis is the major process responsible for formation of organic matter in the environment. Autotrophic green plants, particularly land plants are most important from the point of view of photosynthetic production of organic matter. In the photosynthesis, carbon dioxide and water are used and a certain portion of short-wave solar radiation is absorbed and expended within the plant cover. In considering the role of plant cover of Earth in the global energy and water balance, it is necessary to consider the amount of solar radiation and water utilised by plants in production of biomass i.e. in photosynthesis. For this, the quantities that are calculated and studied are efficiency of photosynthesis and productivity of transpiration.

1. Efficiency of photosynthesis: It is the ratio of energy expenditure on the synthesis of biomass to the total quantity of solar energy absorbed by plant cover in an area. Many experimental studies have shown that this efficiency of photosynthesis is very modest and under normal conditions, usually does not exceed 0.1 to 1.0 percent. However, under very favorable conditions, it may increase to several percent.

2. Productivity of transpiration: It is the ratio of the amount of biomass produced to the quantity of water transpired by photosynthesizing plant cover. This productivity of transpiration usually ranges from 0.5 to 0.1 percent which indicates that photosynthesizing plants use very little water and abundant transpiration in them merely circulates the water in the environment.

Thus general low values of both the above quantities indicate that under natural conditions, plant cover assimilates only a negligible part of available energy and water resources i.e. there is substantial limitation on the use of natural resources in production of biomass in the environment. It is important to establish the causes of this limitation for the study of the relationship of productivity of plant cover to climatic factors. Experimental studies of maximum possible efficiency of photosynthesis in controlled environmental conditions when carbon dioxide of the atmosphere is fully utilized indicate that under such conditions, plants can assimilate 5% or more of the solar energy received and the productivity of transpiration also increases manifold. However, in natural conditions maximum possible photosynthesis and, therefore, the production of biomass is greatly limited by various factors other than the availability of resources.

PHOTOSYNTHESIS WITHIN TERRESTRIAL PLANT COVER

In nature, most of the photosynthesis takes place within the terrestrial plant cover in which different meteorological conditions exist at different levels. The efficiencies of photosynthesis at various levels within the plant cover are not same and are determined by particular microclimatic (meteorological) conditions prevailing at different levels. The microclimatic effects of a forest cover are explained in terms of:

1. Plant coverage characteristics: These characteristics depend upon:

1. 1. Density of dominant forms in the forest covers.

   2. Distribution of different forms in the forest covers.

2. Stratification characteristics of plant cover: These characteristics depend upon:

   1. Total vertical height of plant cover.

   2. Number of vertical strata in the plant cover.

   3. Morphological characteristics of each strata in the plant cover which are determined by branching pattern of plants, evergreen or deciduous nature of foliage, size, density, texture and orientation of leaves.

Importance of the above features can be judged from comparison of tropical and temperate forest plant covers. In tropical forests, average height of tall trees is 46-55 metres, species diversity is 40-100 species per hectare, stratification is strong with 4-5 strata, undergrowth is dense commonly with two upper foliage strata and lower strata being denser. In temperate forests, average height of tall trees is about 30 metres, species diversity is less than 20 species per hectare, stratification is poor with usually 2-3 strata which are almost continuous from low shrubs to top of trees.

In the study of photosynthesis at various levels within plant cover, averaging of the values of meteorological elements at one level along the horizontal line is appropriate and it makes it possible to exclude the influence of individual plants on the meteorological regime. By applying such averaging techniques, following conclusions have been established:

1. Microclimate within the plant cover may be represented by a series of vertically varying profiles of meteorological elements, particularly of solar radiation, water vapor pressure, air temperature, carbon dioxide concentration and wind speed.

2. The profiles of meteorological elements show diurnal and seasonal variations.

3. The average vertical flow of short wave and long-wave radiation, heat and water vapor within the layer of plant covers and the momentum of the system depend substantially on height.

Microclimatic profiles within plant cover

1. Solar radiation: Plant cover significantly changes the pattern of incoming and outgoing radiation. Short-wave reflectivity of area depends somewhat on the density and characteristics of the plant cover. The albedo of areas having coniferous forests is about 8-14 while that having deciduous forests is about 12-18. Albedo of semiarid savannas and woodlands is much higher.

Large amount of solar radiation is trapped within the foliage canopy e.g. Fagus sylvetica forest traps about 80% of incoming radiation in the top strata of canopy and less than 5% reaches the ground. Such trapping is more pronounced on sunny days.

The foliage canopy absorbs more short-wave radiation than long-wave infrared radiation e.g. in tropical forests of Nigeria, only 7.6% radiation of <0.5 m reaches the ground while 45.3% of radiation of >0.6 m reaches the forest floor.

Effect of the age of plant cover on the penetration of light into the plant cover can be judged by the observation that in Pinus sylvestris forest in Germany, percentage of light reaching the ground floor is about 50% at 1.3 year, only 7.0% at 20 years and again 35% at 130 years.

Penetration of solar radiation within the plant cover generally obeys Bougner-Lambert Law:

I = I_o e^-KL

Where, I = radiation intensity on a horizontal plane within the plant cover; I_o = radiation intensity on a horizontal plane above the plant cover; L = leaf area index; K = Extinction coefficient.

Extinction coefficient (K) is constant for a given species and is related to:

• Amount and type of leaf chlorophyll.
• Canopy architecture and
• Reflectivity of leaves. Its value lies between 0.3 and 0.5 for grass-type plant cover and approaches 1.0 for nearly horizontal leaves. Value of K shows inverse relationship to chlorophyll content and reflectivity of leaves.

In general, light penetration into plant cover depends upon the type of plants (particularly trees), spacing of plants, age of plants, crown density, height of plants (particularly trees) and time of year. Percent light reaching the forest floor in some types of forests is given below:

• Birch-beech forest 50-75%
• Pine forest 20-40%
• Spruce-fir forest 10-25%
• Tropical forest 0.1-0.01%

In deciduous forests, light penetration increases during leafless conditions.

Thus the intensity of solar radiation decreases exponentially from top of plant cover towards Earth’s surface due to absorption and radiation scattering by the surface of plants. The resulting radiation balance, therefore, also decreases in the same direction due to screening effect of plants.

1. Air temperature: During day time, heating of foliage canopy causes a convectional transfer of sensible heat and so air temperature within upper canopy may be higher than above the canopy or below. At night, the relationship is reversed as upper canopy layer of air is cooled by contact which are both losing heat by radiation and also transpiring slowly.

Modification of thermal environment is due to shelter from sun, blanketing at night, heat loss by evotranspiration, reduction in wind speed and obstruction to vertical airflow.

Blanketing causes lower maximum and higher minimum temperatures and causes lower mean monthly temperatures in tropical and temperate forests.

At sea level, mean monthly differences in air temperature in temperate forest may reach 2.2^OC in summer but only 0.1^OC in winters. In hot summers, this difference can be more than 2.8^OC.

In forests, which do not transpire greatly in summers e.g. forteto oak maquis of Mediterranean area, day temperatures in woods may cause mean monthly temperatures to be higher than in open.

Altitude in the same climatic zone may affect the degree of temperature decrease in temperate forests. At 1000 meters altitude, lowering of temperatures may be twice that at sea level.

Vertical stratification in plant cover modifies the thermal profile within it in complex ways. In tropical forests, dense foliage canopy heats up greatly during daytime and cools rapidly during night. It shows a much greater diurnal temperature range in denser canopy than in the lower strata. Whereas daily temperatures of second story are intermediate between those of the tree tops and undergrowth, the nocturnal minima are higher than either tree tops or undergrowth because the second story is insulated by trapped air both below and above.

1. Saturation water vapour pressure: This profile within plant cover shows close correspondence with temperature profile both during day and night. Forest temperatures differ strikingly from those in open and the forest water vapour pressures were found to be higher within an oak stand than outside it for every month except December.

At night, actual water vapor pressure almost reaches saturation as the air and canopies are cooled by radiation and convection. Some water vapor is transferred through transpiration from the canopy. During day, upper canopy is air heated by convection and water vapor pressure curve shifts much from saturation curve. Deficit between the two increases the downward and at quite lower level, actual water vapor curve inflects. Towards the bottom of canopy, it reapproaches saturation curve due to transpiration coupled with low air movement and low temperature towards base of plant cover.

The flow of water vapor within the plant cover increases with height because of the influence of transpiration by plants and the momentum of the system declines downwards from the plant cover’s upper boundary as a result of the inhibiting effect of plants on the movement of air. This effect is associated with the reduction in turbulent exchange within the layer of plant cover compared with higher air layers. The coefficient of turbulent exchange within the layer also declines towards Earth’s surface.

Humidity conditions within the plant cover are very much different from those outside it due to evotranspiration characteristics of the cover. It generally depends upon the type of plant cover, density of plant cover, structure of vertical stratification and temperature effects. Time of the day and season also affect evotranspiration and, therefore, humidity within the plant cover.

Evotranspiration generally increases with density of vegetation and within the plant cover, relative humidity may be 3-10% higher than outside. This effect is more pronounced in summers.

Rainforests have high transpiration and so have high humidity inside their plant cover. Mean annual relative humidity excess is reported to be 9.4% in beech, 8.6% in Pinus abies forest, 7.9% in larch forest and 3.9% in Pinus sylvestris forests.

In tropical forests, night exhibits complete saturation while in daytime, the humidity decreases with height.

1. Carbon dioxide: The profile of carbon dioxide concentration within plant cover shows much diurnal variation due to photosynthetic uptake of carbon dioxide during daytime and respiratory addition of this gas during night. Carbon dioxide concentration in soil is very low and its use by plants is spatially and temporally very inhomogeneous.

During daytime, CO₂ concentration decreases from upper canopy towards ground. It reaches a minimum point near middle of canopy. Below this point, CO₂ concentration rapidly increases towards ground and becomes equal to CO₂ concentration outside the canopy at a point that roughly corresponds to compensation light intensity. It reaches fairly high level at soil surface. This profile is due to photosynthetic depletion of carbon dioxide in upper canopy, equilibrium corresponding to compensation point lower in the canopy and respiratory addition of carbon dioxide from lowest shaded leaves and soil microorganisms.

In the night, concentration of CO₂ gradually increases towards ground level due to its respiratory addition.

1. Wind velocity: The profile of wind speed shows no strong change in day and night but overall wind speed is higher during daytime due to convectional effects. Wind profile within the canopy develops due to steady state boundary layer flow. The profile is logarithmic above the canopy and becomes exponential within the canopy. The zero plane displacement (D) depends on the height of plants. The roughness height (z_o) is a measure of community roughness and it is effectively the thickness of a laminar sublayer through which individual elements project. Value of z_o is related to height variation and spacing of individual elements which in the plant cover are plants. In extrapolation of logarithmic curve downwards, the zero velocity intercept is found to lie at the height D+z_o. If the canopy were rigid, it would have a constant value but variation of surface roughness depends on leaf flutter, movement of branches and leaf streamlining. These variations cause variations in value of z_o with wind speed. Surface frictional characteristics are entirely specified by D and z_o. The wind profile is major factor in establishment of profiles of saturation water pressure and carbon dioxide within the plant cover.

Lateral air movement is generally lesser within the plant cover than outside it. Even large variations in outside wind velocities do not affect airflow inside forest cover. Vertical stratification structure, leaf canopy architecture, density of stand and season have marked influence on wind velocity profile within a plant cover. For this reason, reduction in wind velocity within the forest cover is different in temperate and tropical forests. Reduction in wind speed from outer edge towards deep inside a forest is greater in tropical rainforests. In temperate European forests, wind velocity at outer edge of forest is reduced to 60-80%, 50% and 7% at points 30 m, 60 m and 120 m respectively deep inside the forest. In Brazilian evergreen forest, wind velocity of 2.2 m/second at the outer edge of forest is reduced to 0.5 m/second at 100 m deep inside forest while at 1000 m inside forest the wind velocity becomes negligible. In this forest, outside storm velocity of 28 m/second was reduced to 2 m/second at 11 km deep inside the forest.

1. Flow of water vapor and momentum of system: The flow of water vapor within the layer of plant cover increases with height because of the influence of transpiration by plants. The momentum of the system decreases downwards from upper boundary of plant cover towards ground level as a result of the inhibiting effect of plants on the movement of air. This effect is associated with the reduction in the turbulent exchange within the layer of plant cover compared with higher air layers. The coefficient of turbulent exchange within the layer also decreases towards ground level.

The theory of photosynthesis within plant cover and numerical models of this process developed in recent years are based on the general idea of a transition from photosynthesis within a single leaf to photosynthesis within a layer that is homogeneous horizontally but possesses different physical conditions at various heights. Application of the theory of photosynthesis within a layer of plant cover indicates following general conclusions:

1. When assimilation process is not very sensitive to different meteorological elements, total assimilation within the plant cover strongly depends on the radiation flux for low levels of radiation. For large values of radiation, the total assimilation is independent of radiation flux and becomes dependent on other factors particularly temperature.

2. Within he plant cover, increase in total assimilation with increase in inflow of CO₂ from soil is much slower than would occur if all the inflow of CO₂ from soil were to be expended on assimilation. This is because the inflow of CO₂ from soil first encounters the leaves located in shade, which are not able to photosynthesize intensively due to insufficient radiation. The general increase in CO₂ concentration produced by its inflow from below is compensated by a reduced contribution of CO₂ from above. Thus assimilation within plant cover is influenced very little by the upward inflow of CO2 from soil and is largely influenced by flow of CO₂ coming downwards from the atmosphere.

PRODUCTIVITY OF PLANT COVER

The productivity of plant cover () is the difference between total assimilation and the expenditure of organic matter on respiration within a plant cover. Thus the productivity of a particular plant cover depends on photosynthesis and respiration in it. The leaves of plants are the major organs association with both photosynthesis and respiration. Therefore, the productivity of plant cover substantially depends on the value of the index of leaf surface (leaf index) and decreases for both very small and very large values of this index. In view of this, the value of productivity of a plant cover is calculated for an optimal value of leaf index i.e. the value of this index that corresponds to the highest value of productivity. From various studies, it has been established that the parameters and factors that affect the photosynthesis within a plant cover also influence the productivity of plant cover. Thus the productivity of plant cover is mainly determined by parameters characterizing the properties of plant cover itself and the climate.

In general, following important points can be observed in relation to productivity of plant cover in nature:

1. The structure of plants in the plant cover continuously changes throughout their life cycles and photosynthetic activity of leaves is never optimal throughout the entire vegetative period of any plant.

2. Availability of mineral nutrients in nature is always less than optimally required for maximum possible photosynthesis.

3. Under natural conditions, water regime of soil is also not constantly maintained at optimally required level.

Thus the productivity of plant cover in real natural conditions is always less than theoretically possible maximum level due to complex interactions between a variety of biological, climatic and soil factors.

Climatic factors and productivity of plant cover

In conditions of sufficient moisture, two climatic factors i.e. photosynthetically active radiation and temperature are particularly important in relation to productivity of plant cover.

The influence of radiation and temperature on productivity of plant cover is quite complex. In real natural situations, radiation is always a factor whose value is a ‘minimum’ because radiation available to leaves in lower layers of canopy is always insufficient. Therefore, increase in radiation flux always results in increased productivity of plant cover.

With increase in temperature, the productivity of plant cover increases initially. After attaining a certain maximum value that depends on the value of radiation flux, productivity begins to decrease with further increase in temperature. Thus productivity of plant cover substantially decreases above a certain threshold value of temperature which is determined by the radiation flux.

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 13^th most abundant element in Earth’s crust (0.1%) and 9^th 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-1. Indicative characteristics of major tropospheric sulfur compounds.

Nitrogen is primarily exchanged between atmosphere, biosphere and soil. Following Table-1 shows the estimated total nitrogen 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-1. Nitrogen storage in various components of global environment.

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-2. Indicative characteristics of major atmospheric nitrogen compounds.

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 (CO₃^2-) 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

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.

Soils are formed by mixture of weathered rock material with the organic matter derived from decomposition of mostly plant litter. The role of plant cover in pedogenesis and determination of the soil type of an area is clear from the fact that all zonal soil types correspond to specific types of plant covers.

Tundra soils

Rock tundra is associated with isolated patches of lichens and mosses with occasional higher plants. Tundra moor is associated with peat-forming mosses mainly Polytrichum sp., lichens like Cladonia sp. Or Cetraria sp., grasses like Carex, Eriophorum, herbs like Potentilla, Ranunculus, Gentiana, Saxiraga, Dryas octapetala, shrubs and trees like Betula nana, Salix herbacea, S. reticulata, S. arctica and heathers like Empetrum nigrum, Cassiope tetragonal.

Podsols

Iron-humus podsol develops under heathland with Ericaceous dominants while iron-podsol develops under coniferous forests. Podsolization is strongly influenced by the type of vegetation. Certain species hasten podsolization e.g. Ericaceous heath Calluna vulgaris and Erica cinerea that usually occupy cleared forests on acid brown soil. Some conifers and Fagus sylvatica and Quercus sp. In Britain and Populus trichocarpa in Alaska are strong acidifiers. Sphagnum sp., Eriophorum sp., and Molineae caerulea form blanket peat, forming peat podsol or peaty-gley.

Brown forest soils

Very high productivity of broad-leaved summer forests plays considerable part in the pedogenesis of brown forest soils by maintaining quite high activity of soil microorganisms and sol fauna, particularly earthworms.

In North Europe, clearing of broad-leaved forests present on acid brown soils that had developed on siliceous parent material resulted in the establishment of Ericaceous heath lands. This hastened podsolization in those areas. These podsols are being maintained today by burning and felling of trees along with maintenance of heath-land. In absence of such interference, Calluna-Erica heath-land is easily replaced first by bracken (Pteridium aquilinum) and then by conifers. These plants bring back the podsol soil to acid brown soil.

In the absence of normal forest, the brown forest soils can be maintained under grass cover because much organic matter is returned to the soil by extensive root system. However, inorganic fertilizers are needed even then. If land is under crop cultivation, both inorganic and organic fertilizers are needed to maintain the soil.

Red & brown soils of arid subtropics

The soils in arid subtropics formerly had luxuriant vegetation of Quercus ilex and Pinus halpensis. However, overgrazing in the areas of brown soils on limestone in Europe resulted in sparse vegetation of low trees causing conversion of brown soils (Terra fusca) to red soils (Terra rosa). Brown soils in many areas still have comparatively better sclerophyllous cover. Afforestation on red soils protects them from summer solar insolation and changes them to brown soils again.

Other red and brown soils

While red brown soils develop under subtropical dry forests, red-yellow soils develop under subtropical forests and red laterite soils under tropical forsts. Brown and grey soils are developed under deciduous forests, grey-brown soils under semi-desert or desert scrub forests and chestnut brown soils under steppe.

The plant cover and the cycling of nutrients are intricately interlinked with each other in the ecosystem. Nutrient cycling and biomass normally reach equilibrium under climax conditions. External influences, particularly destruction of plant cover may break this cycling and disrupt this equilibrium causing deterioration of the environment. For example, in North Europe, destruction of deciduous forest cover by human activities during Late Stone Age and Bronze Age caused loss of nutrients from the upper layers of soil. The deteriorated soil caused establishment of heath-lands on them. These heath-lands are presently maintaining and are being maintained by reduced nutrient cycling from the upper layers of soils only. In the tropical areas, laterite soils (latosols) have very deep crusts of weathering and are greatly leached of nutrients. However, rain forests of ancient geological ages on these soils provide rich vegetation of great biomass in which most of the ecosystem nutrients are locked up. Despite poor nutrient availability in the soils in these tropical rain forests, such forests are presently being maintained in these areas only due to highly efficient nutrient cycling from the upper layers of the soil. The huge amount of plant organic matter from the vegetation falling on the soil decomposes and provides nutrients to the plants again.

In the soil under a forest cover, the absorption of several nutrients from the deeper layers initially reduces their availability but the return of these nutrients with fall of litter again increases and maintains their availability in the upper layers of the soil. This effect is particularly marked for Mg and Ca. The efficiency of nutrient cycling is greatest in rain forests followed by deciduous forests, coniferous forests and grasslands in that order. Coniferous forests return 50-100 kg/ha/yr of ash elements while deciduous forests may return 200-270 kg/ha/yr. The return of Ca in the rain forests is 200-300 kg/ha/yr while in deciduous forests is only 150 kg/ha/yr.

The soil structure is also greatly affected by the plant cover because roots of plants have a direct influence in maintaining the rhizosphere bacteria whose capsular slimes and gums stabilize the soil crumbs. Rhizosphere zone in the soil provides nearly ideal conditions for both aggregate formation and aggregate stabilization by incorporation of bacterially synthesized macromolecules. In the grassland cover, rapid aggregate promotion is certainly due to rapid and prolific root production of these plants.

The plant cover also influences soil fauna and consequently, the soil structure. In the forest mull soils, the plant cover provides litter that promotes and maintains rich earthworm population in the soil. In these soils, earthworms create pore space through voided casts that are stabilized initially by fungal growth and later by cementation with bacterially produced polysaccharide macromolecules.

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 the Table.

Table: Probabilities of reactions on collision of gas molecules with surface.

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,

                                                    hv

2CO + O₂ —————-> 2CO₂

                                               TiO₂, ZnO

                                                 hv

2N₂ + 6H₂O ————-> 4NH₃ + 3O₂

TiO₂

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.

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.

EARTH’S SURFACE WIND SYSTEMS

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.

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  um  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.

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Class                   Size range (mm)                Impact area

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(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

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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.