Environment of Earth

March 11, 2008

ATMOSPHERE OF EARTH

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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

ATMOSPHERIC STRATIFICATION

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.

ATMOSPHERIC COMPOSITION

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.

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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.

AIR PRESSURE, WINDS AND GLOBAL AIR CIRCULATION

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.

AIR PRESSURE

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.

WINDS AND WIND SYSTEMS

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.

MONSOON WINDS OF ASIA AND NORTH AMERICA

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.

GLOBAL CIRCULATION SYSTEMS

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.

JET STREAM

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.

ATMOSPHERIC MOISTURE

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.

CLOUD PHYSICS

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.

TYPES OF PRECIPITATION

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.

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