Indian Journal of Radio & Space PhysicsVol. 17, December 1988, pp. 203-219

Aerosols and Radiation Budget in the Middle Atmosphere

.// B V KRISHNA MURTHY

#' Space Physics Laboratory, Vikram Sarabhai Space Centre, Trivandrum 695 022

A brief outline of the atmospheric aerosol formation, destruction and distribution processes and size distributions ispresented. Recent experimental observations in India are discussed. Some important aspects of the radiative effects of theaerosols are presented.

1 Introduction

Atmospheric aerosols playa very important role inmany atmospheric processes some of which may havein turn an impact on climate. However, our ability toassess the role of aerosols in atmospheric processcs isgreatly handicapped by the very limited informationavailable on the properties of aerosols. The inputs toatmospheric aerosols come from natural and man­made sources. The recent thrust to aerosol studiescame from the concern about the possible effects ofman-made aerosols on atmospheric processes andclimate. However, a realistic assessment of this can­not be made because of our incomplete knowledgeabout the characteristics of natural aerosols and theirrole in various atmospheric processes.

The radiation budget of the middle atmosphere isinfluenced primarily by ozone, carbon dioxide andwater vapour ~Aerosols affect the radiation budgetthrough scattering and absorption processes. For aproper evaluation of this effect, the data required onaerosol characteristics are rather scanty. Aerosol da­ta over the Indian continent are very scarce if not vir­tually non-existent prior to nineteen eighties. A majorthrust to aerosol studies has been given in the IndianMiddle Atmosphere Programme.

2 Aerosols: Their Sources, Formation, Destructionand Distribution Processes

Aerosols are defined as suspended particles in sol­id and/or liquid phase in the atmosphere. They occurover a wide range of sizes from about 10 - 9 to 10 - 4 min diameter. The size range can again be subdividedinto three ranges based on atmospheric effects of aer­osols, The smallest particles of sizes less than- 0.1 flm are called Aitken particles and :ue import­ant in atmospheric electricity. Particles in the sizerange 0.1 flm to 1.0 flm are referred to as large parti­cles and are effective in atmospheric optics and in ra-

diation budget. These are also effective as condensa­tion nuclei. Particles of size range > 1.0 flm play animportant role in cloud physics and are effective ascondensation nuclei. The large particles (0.1 flm to1.0 fl m) are sometimes called haze particles.

Aerosol particles are either directly produced orformed by various gas-to-particle conversion pro­cesses. In either case, the 'original' aerosol particle un­dergoes a variety of chemical processes and physicalmodifications in the atmosphere. These processesand modifications result in a continually changingchemical composition and particle size distributionand hence physical properties. Particles from differ­ent sources are mixed by Brownian diffusion andcoagulation processes on a microscale and by atmos­pheric turbulence and circulation on a macroscale.The principal sources of aerosols are listed in Table 1with approximate estimates of their input levels. Thewide diversity in the different estimates is an indica- .tion of the uncertainty involved in making the esti­mates and the need for more relevant information.

Some of the characteristics of the size distributionof aerosols depend upon the various formation andtransformation processes of aerosols. Particles in thesize range below about 1 flm (transient/Aitken nuc­lei range) are produced by gas-to-particle conversionprocesses whereas particles of sizes greater than1 flm (coarse range) are formed directly by mechani­cal processes soch as wind blown dust and sea saltspray. The relationship between the particle produc­tion processes and size range is shown in Fig. 1. Pani­cles at the lower end of the size range have relativelyshort atmospheric residence times because of chemi­cal reactivity and greater physical mobility, Theseparticles will eventually be incorporated in the0.1 flm to 1.0 flm range which is generally referred toas the accumulation moue. Conversion, of particles inthe transient and a<:cumulation ranges to coarse rangeand vice versa is not a very efficient process. Thus,

1()3

INDIAN J RADIO & SPACE PHYS, VOL. ]7, DECEMBER ]988

Table1-Estimates ofParticleProduction in106 tons/year

(ii) Particles formed fromgases

Converted su]phates 200Converted nitrates 35

Converted hydrocarbons 15Sub-total 250Total of man-made 280sources

Source

Man-made

After Peterson and Junge 1

(i) Direct particleproduction

TransportationStationary fuel sourcesIndustrial processesSolid waste disposalMiscellaneousSub-total

1.89.6

12.40.45.4

29.6

After Hidyand Brook2

37-110

1102327

160197-270

WIND BLOWNDUST

SEA SPRAYVOLCANOPLANT

PARTICLES

I I I I I -.. •0.01 0.1 1.0 10 100

TRANSIENT OR ~CHANICAl

+- AITKEN NUCLEI -+ ACClNJl.RIOOl -+- GENERRIOOl _RANGE RANGE RANGE

_ FINE PARTICLES ••• COARSE PARTICLES _

Fig. I-Atmospheric aerosol surface area distribution showingthe main modes and principal source and removal mechanisms

T ...

particles in the lower two ranges are quite distinctfrom those in the coarse range.

2.1 Gas-to-Particle Conversion Processes

It is seen from Table 1 that gas-to-particle conver­sion process is a major contributor to the aerosol par­ticle production. It contributes more than half of thetotal particle production due to natural sources aswell as man-made sources. Atmospheric gases can in­teract with each other and with existing particles toform new particles or to modify the existing particlesby different processes. The main processes are (1)homogeneous homomolecular nucleation, (2) homo­geneous heteromolecular nucleation, and (3) hetero­geneous heteromolecular nucleation. The first pro­cess involves formation of new, liquid or solid ultra-

Natural

204

(i) Direct particleproduction

Sea saltWind blown dustVolcanic emissionsMeteoric debrisForest firesSub-total

(ii) Particles formed fromgases

Converted su!phatesConverted nitrates

Converted hydrocarbonsSub-totalTotal of natural sourcesGrand Total

50025025

o5

780

3356075

47012501530

109560-360

40.02-0.2

1461610

37-365600-620182-1095

208036903960

fine particles from a gas phase consisting of a singlegas species only. The second process involves forma­tion of new particles from a gas phase consisting oftwo or more gaseous species. In this, the most com­mon one of the species is water vapour. In the thirdprocess, i.e. heterogeneous, heteromolecular nuclea­tion, growth of pre-existing particles takes place dueto condensation of gaseous species.

Atmospheric trace gases undergo chemical reac­tions to produce reaction products with a low vapourpressure. As more and more of these products areformed, a state of supersaturation will be reachedwith respect to these molecules. The degree of super­saturation will determine the degree of nucleation(formation of new particles) and condensation (depo­sition on pre-existing particles). From the thermody­namic point of view, more energy is required to form anew particle (nucleation) than to enlarge the surfaceof the existing particle (condensation). The followingreactions are considered to be important for the gas­to-particle conversion: (1) sulphur dioxide reactswith hydroxyl radicals to form eventually sulphuricacid molecules, (2) non-methane hydrocarbons reactwith ozone and/or hydroxyl radicals to form alde­hydes, alcohols, carboxylic acids and dicarboxylic ac­ids, and (3) most secondary reaction products of non­methane hydrocarbons react with oxides of nitrogento form organic nitrates. The particles formed bythese mechanisms are in the fine particle size range( <.05 ,urn). Gas-to-particle conversion can also takeplace through direct reaction of gases with particleson the surface or in the interior in case of liquid parti­cles.

..

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KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

... (2)

... (1)

Gas-to-particle conversion plays a major role in theaerosol production in the stratosphere. Sulphur diox­ide is converted through various reactions to sulphateaerosols in the stratosphere. In fact it is this process ofgas-to-particle conversion which is primarily respon­sible for the Junge layer in the stratosphere. Eversince its discovery by Junge (Junge et ai. 3) the stratos­pheric aerosol layer has been a subject of great inter­est. It has been well established that the stratosphericparticles are primarily composed of liquid solutionsof sulphuric acid and water. Other substances likeammonium sulphate are also generally present(Hamill et ai.4). These sulphate particles are producedlocally in the stratosphere, mainly from precursor sul­phur-bearing vapours. Sulphur dioxide (SOz) andcarbonyl sulphide (OeS) are the most important pre­cursors. SOz, which is an important component ofvolcanic effluents, is injected directly into stratos­phere in large quantities by major volcanic eruptions.During periods of quiescent volcanic activity, the ma­jor precursor gas is carbonyl sulphide (OeS) with car­bon disulphide (eSz) as a possible secondary source(Sze and K05; erutzen6). Unlike most of the other sul­phur gases, oes is expected to have a long tropos­pheric lifetime of several years (Turco et aU). oesoriginates in biological, volcanic and industrial pro­cesses (mainly from the surface). esz is less inert thanoes but it has aeronomic significance beGause of itspossible oxidation to oes. Apart from oes and esz,SOz from surface emission is also transported to low­er stratosphere even during periods of quiescent vol­canic activity. Detailed modelling studies have beenmade of the tropospheric oes cycle and the stratos­pheric sulphur balance during volcanically quiescenttimes (Turco et ai.8•9). These studies show that oes isthe dominant sulphur source compared to esz andSOz for the stratospheric aerosol layer. The precur­sor gases from troposphere are transported to stra­tosphere through eddy diffusion and direct injection.The direct injection can occur during thundercloudpenetration, tropopause folding and vertical convec­tion especially in the tropical zone. Horizontal advec­tion subsequently distributes injected material zonal­ly and meridionally.

In the lower stratosphere, oes breaks up into sul­phur and carbon monoxide by photo decompositiondue to absorption inthe extreme ultraviolet, 200-240nm. The sulphur atoms thus produced react rapidlywith molecular oxygen to form SO which in turn is ra­pidly oxidized to SOz. SOz is further oxidized mainlythrough reactions with OH (three-body reaction) toHzS04' For detailed discussion on the sulphur chem­istry, articles by Turco et aU. 10 may be referred.

Depending on the saturation levels, the H2S04 andH20 vapours nuclease by heterogeneous heteromo-

lecular or homogeneous heteromolecular processes.It isgenerally believed that the former process ismuchmore efficient than the latter one. However, recent as­sessment in this regard (Yue and Deepakll) broughtoutthatattemperaturesaround - 7Soewhicharenotunusual in the stratosphere, the homogeneous nuc­leation mechanism can be a significant contributor tothe overall nucleation process. Yue and Deepak stud­ied the relative contribution of the two nucleation

processes at different temperatures and vapour pres­sures (of Hz SO4 and HzO )and arrived at the above re­sult.

Other processes of nucleation include ion nuclea­tion and ion-ion precondensation nucleation. Ionnucleation involves condensation of gaseous vapours(e.g.,HzSO 4 and HzO) on charged molecular clusters.The electrostatic energy of the core ion stabilizes thenucleation embryo, effectively lowering the barrier tonucleation. The ion-ion precondensation nuclei maybe formed when large cluster ions of opposite chargesrecombine. The contribution of these processes tothe total nucleation is expected to be not significantcompared to the other processes described above.

After nucleation, an aerosol continues to grow byvapour condensation. Both evaporation and attach­ment processes take place at any given instant and thenet effect of these two processes determines the rateof growth by condensation. In the troposphere,growth of aerosol particles by condensation of watervapour plays an important role. The change in theparticle size is related to the relative humidity(HanePZ)by

[ , ( )] 1/3r(aw)= ro 1+P :law

where

ro E>ryparticle radiusp' Particle density relative to that of watermw(aw) Mass of condensed watermo Dry particle massaw Water activity which is essentially the relative

humidity f corrected for curvature of theparticle surface.

tlw is given by

[ 2St Vw]aw= f exp -~.-RwTr

where

5t Surface tension on the wet particle surfaceVw Specific volume of waterRw Specific gas constant for waterT Absolute temperature in Kr Particle radius

205

INDIAN J RADIO & SPACE PHYS, VOL. 17, DECEMBER 1988

For room temperature the value of2St Vj R" T isap­proximately equal to 0.001056 ,urn,

2.2 Role of Meteoric Debris

The meteoric influx to the earth amounts to

1.6 x 107 kg per year. The meteor input to total stra­tospheric particulate mass input constitutes only afew per cent. Hunten et al.13 and Turco et al.14 madedetailed theoretical calculations on the effects of

meteoric debris on stratospheric aerosols. Turco etal,14 concluded from their model calculations thatmeteoric debris is an important natural aerosol con­stituent for particles larger than 1,um radius through­out the stratosphere and may be dominant for parti­cles smaller than 0.01 ,urn radius above 20 km. Forother sizes, the meteoric debris generally constitutesless than 10 per cent of the total aerosol mass. Theircalculations show that very few meteoric particlesreach the upper troposphere in their natural state andparticipate as cloud ice condensation nuclei. This is incontrast to the Bowen's hypothesis1) of relation be­tween precipitation and meteoric dust influxes.

2.3 Sedimentation

The fall velocity and the diffusion coefficient arebasic parameters to determine the sedimentationrate. These parameters are particle size dependent.Atsizes comparable to molecular dimensions, kinetictheory of gases is applicable and for larger particlesStokes-Cunningham formula is applicable.

For large particles, the fall velocity is given by

d Molecular diameter

ng Number density of molecules (air)YJ Viscosity of air

The slip correction factor (a) arises because thedrag force on the particle must be corrected for slip­page of gas at the particle surface. With

1;;1IIYJ=- -ng mkT

where m is the mean molecular mass, and for Ilr ~ 1and as p ~ Ps,

v= j " 2 {hp,g ... (5)mkT 9 ng

The approximate values of fall velocity for differentsizes at different height~ are given in Table 2.

2.4 Coagulation

Coagulation is controlled by the diffusion coeffi­cient and is thus important for smaller particles.Coagulation controls the smaller particle end of theaerosol size distribution whereas sedimentation con­trols the larger particle end. The rate of coagulationbetween two groups of particles of different sizeswithin a population n( r) is given by

dn(rlz)-'-= -4n(D] + Dz)(r1 +rz)dt

r ~,

... (8)

Table 2- Fall Velocities of Particles of Different Sizes

x n( r1) n( rz) d rid rz ... (6)

where D (= kT Bs) is the diffusion coefficient andn( r) is the concentration of the particles. SubstitutingforD,

dn(rl z) 2 kT[ 1 1 (1 1)]-~'-= --- -+-+113 2+2 (r1+rz)

dt 3 YJ rl rz r1 rz

x nh)n(rz)drl dr2 ... (7)

which can be written as

... (4)

... (3)v= msgBs

and the diffusion coefficient by

dsc = kT B,

where

ms Mass of the particleg Acceleration due to gravityk Eoltzmann's constant

Bs Particle mobility, defined as the velocity per un­it driving force

Bs is given by

( p) aB=l--~-

s p, 6nYJr

Velocity v(cm/s)at.r=where

p Density of airPs Density of particlea (= 1 + I3l1r) Cunningham's slip correction fac­

tor

13 1.26+ 0.40 exp (-1.10 rll)I (= 1IIi n dZ flg) Mean free path length

Altitudekrn

30201510

5,um

0.50.170.120.10

2,um

0.240,080.0550.045

1,um

0.060.020.0140.011

0.5,um

0.D15

0.0050.00350.003

206

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KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

The combination of r1 and rz results in a particle with

( 3 + 3)1/3r= rl rz

Fig. 2 (JungeI6) shows how a given model size dis­tribution of particles undergoes changes due to coag­ulation. Rapid decrease in the small particle concen­tration can be seen in this figure. This results in a dis­placement of the maximum towards larger particles.

... (13)

... (12)

or

dn(r) = Cr -(v+ I) dr

dn(r) = CIr- Vd(log r)

Fig. 2-Change in number and volume distribution with size ofcontinental aerosols due to coagulation (fromJunge 16)

>I~.•. -.•.

10-1 100

RADIUS (}Jm)

~•.Q.

...Eu

chemistry is important in understanding the effect ofwet removal processes.

The log radius-surface(s) and log radius-volume( V )distributions are written as

ds(r) = 4Jlrz dn(r)d(log r) d(log r)

2.6 Aerosol Size Distributions

The relationship between aerosol producing pro­cesses and their contribution to different aerosoi size

ranges has been discussed earlier. The size distribu­tion of aerosols is a very important information in thestudy of effect of aerosols on atmospheric processes.The aerosol size distribution can be expressed interms of number density distribution or surface areadistribution or volume distribution. The surface areaand volume distributions are very important in thestudy of aerosol chemistry.

Junge 16 found that the number density size (radiusr) distribution can be approximated by a power law inthe size range 0.04-10 flm in the case of continentalaerosols. This is written as follows

... (9)

... (11)

'" (10)

zr dr1 drdt

x V 3 3r - r1

The number of particles with radii r decreases bycoagulation with smaller and larger particles by

Iz= -Ko L~~f(~_,r)n(rl)n(r)drldrdt

So, the total change!:!..nin particles with radii ris givenby

Itl!:!..n(r)= 0 (II + Iz) dt

The largest particles which can thus participate insuch formation of particle with radios.,. are of size,r( = rl21/3. The total number of particles with radii rformed during d tis then

I, ~ K,L](;'.1r' - ..;)n(r,) n(r' - rll

2.5 Aerosol Removal Processes

There are two processes of direct removal of parti­cles from the atmosphere. These are (1) dry deposi­tion, i.e. gravitational sedimentation, discussed in thepreceding paragraphs, and (2) wet removal. Wet rem­oval is of two types, namely, rain out and washout. Inrain out the particles are in~orporated in precipita­tion nuclei during processes occurring within theclc,ud. Washout involves incorporation of a materialinto precipitation as a consequence of processes oc­curring below the cloud. These removal processes arehighly size dependent. Furthermore, the removal rateof hygroscopic and non-hygroscopic particles will bedifferent because hygroscopic aerosols will iftcreasegreatly in size with increasing relative humidity. Thewet removal processes are important mainly in thelower troposphere where cloud formations exist.However, in tropical latitudes-where the vertical ext­ent of the clouds (thunderclouds) often extends intothe stratosphere, the wet removal processes can be­come important even at altitudes of upper tropos­phere and lower stratosphere. Study of precipitation

207

INDIAN J RADIO & SPACE PHYS, VOL. 17, DECEMBER 1988

.. . (14)

and

d V(r) 4 ,dn(r)~--=- nr'--d(log r) 3 d(1og r)

The power law index (v lis typically 3 but could varyfrom 2 to 6. This type of distribution has been veryuseful as it is a simple function. However, it is general­ly found to be more representative, if the size distribu­tion (especially of tropospheric aerosols) is expressedas a sum of log normal distributions with differentmode radii. It appears'that each particle source willhave particles with their own lognormal distributions.So, with multiple sources, which is in general the case,a multi modal log normal distribution with each modespecified by a total number density, a spread parame­ter and a mode radius will represent the size distribu­tion. The multi modal log normal distribution can berepresented by

dN =I ---117 n; exp [_ [log (~nJ]d r (2 n) - r 0; In 10 201

E

~ 1~I'"'eu­..

••~E 10::::J

C

zo~~0::~Z~UZ -1o 10u~...JU~ -20::10~0.. 00() 1 0.1 1-0 10

PART I C lE RADIUS(fJm}

T .•

0:J

... (15)

where

N Particle concentration for particles of radius r

n; Integral over the individual log normal distri-butionStandard deviationIndividual mode radius

Fig. 3-Measured (from panicle impaction and simple particlelight scattering detection) size distributions of stratospheric aero­sols. The approximate height of observation above the local tropo­pause (in km) is indicated for each distribution (from Turco et

at.III).

at equatorial latitudes in view of the source of aerosolsin the inter-tropical convergence zone.

2,7 Aerosol Continuity Equation

The diffusive flux of the particles is usually derivedfrom the relationship

In Eq. ( 17), Qcan be taken to represent the rate of aer­osol production due to different processes, u( = 1/ tf'tf is the residence time lis the wet removal rate, and yisthe coagulation rate (discussed in previous sections).

where De is the eddy diffusion coefficient and ng is theair number density. The eddy diffusion coefficient isoften derived from measurements of long-lived trac­ers such as N20 and CH4 (species which are not pro­duced photochemically in the atmosphere).

The aerosol continuity equation can be written hyincluding the rates of production, destruction (remo­val) and transport. This is written as

... (17)

... (16)

an a , a-= Q- un- y-- (nv)-~ (¢II)at az az

Two modes are generally adequate to characterizemost aerosol distributions. A third mode may be ne,cessary to represent the Aitken nuclei especially nearsources of combustion particulates.

Stratospheric aerosols can be represented by asingle mode as these are dominated mainly by SO 4

aerosols. These are discussed in detail in literature] o.

Fig. 3 shows some of the observed distributions 10.

The numbers on each distribution denote the heightabove the tropopause. There is remarkable similarityin the distributions at r> 0.1 ,urn. Turco et al.17 ex­

plained this as due to the constant input rate of gase­ous sulphur to the stratosphere and to the depend­ence of aerosol removal rate on particle size. Theremoval rate varies as r hecause of an r'dependenceof particle mass and an r dependence of sedimenta­tion rate. These factors constrain the size distrihution

to a relativCly small variation. At sizes less than 0.1,urn, the small particle concentration decreases withincreasing separation from tropopause. This canprobably be attributed to increasing aerosol maturitywith altitude in the stratospherelK• The distrihutionsshown in Fig. 3 correspond to high latitudes. It wouldbe interesting to see the corresponding distributions

20~

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KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

methods and indirect methods. The direct methodsinvolve actual collection of aerosol samples ,from theatmosphere using impactors and analysing these todetermine the aerosol characteristics such as chemi­

cal composition and size distribution by laboratorymethods. The impactors are carried by balloons andthe collected samples are recovered for later analysis.Another direct method is the filter method in which

aerosols of different size ranges are collected by usingdifferent filters and are analysed for their mass dis­tribution and chemical composition.

Indirect methods use the aerosol property to scat­ter and absorb the incident light and thereby causingextinction. The scattered light signal strength or theattenuation of the light signal as it traverses a knownpath in the atmosphere is recorded. This can be ana­lysed to determine aerosol characteristics such asnumber density and size distribution. The indirectmethods can be carried out in situ using balloon orrocket platforms. Balloon-borne dustsondes211 areamong the indirect in situ experiments that have beenwidely conducted. In the balloon-borne dustsondesthe air containing aerosols is pumped into an illumi­nated cell. The light scattered by the pumped-in air isrecorded. As aerosols pass through the cell they giverise to scattered light pulses. The number of pulsesand the pulse width give information on the aerosolnumber density and the size when analysed using cali­bration data. Another indirect method using balloon­borne/rocket-borne payloads is the light scatteringmethod2!'22. In this method, the scattered sunlight byair molecules and aerosols is recorded at differentscattering angles at different altitudes. The recordedsignal strength as a function of scattering angle is ana­lysed to derive the aerosol characteristics.

The ground-based techniques for the measure­ment of aerosol characteristics employ indirect meth­ods (remote sensing). Among these are the turbiditymeasurement technique and the multi-wavelength so­lar radiometer. Turbidity measurements are carriedout by a sun photometer using wide band filtersB.These measurements can be used to obtain the Ang­strom's turbidity coefficient and the wavelength ex­ponent using power law size distribution for aerosols.As the size distribution differs significantly from pow­er law distribution, especially for lower troposphericaerosols which contribute most to the total atmos­

pheric turbidity, it is necessary to make measure­ments at a number of wavelengths using narrow bandfilters to derive the size distribution function. Multi­wavelength radiometers (MWR) use this method. InMWR, the aerosol optical depth is obtained as a func­tion of wavelength by making measurements of directsolar flux at a number of wavelengths in the visible atdifferent solar zenith angles. By adopting inversion

05

2520--.J

10 0,

60 50 40 30

NORTH LATITUDE (deg)

o90 80

wo=>

'::: 15

~«

Fig. 4-Latitudinal cross-sections of aerosols. The mixing ratiounits (indicated by the contour lines) are in particles/ mg of air. Thethick dashed band indicates the tropopause altitude (from Rosen

eta!.19).

~ 20E

Recourse has to be taken to numerical methods tosolve the continuity equation. One-dimensional mo­delling using the continuity equation has been carriedout for stratospheric aerosols and the results havebeen reviewed by Turco et al.lO The model predic­tions compare fairly well with the observations; How­ever, realistic three-dimensional modelling is neededto understand the global variation of stratosphericaerosols. Modelling of tropospheric aerosols is still inits infant stage mainly because of the plethora ofsources of different nature.

2.8 Latitudinal Structure of Aerosols

Fig. 4 shows the latitudinal distribution of aerosolparticles with radius greater than 0.15 ,um. Verticalarrows on the latitude axis indicates the location of theobservation sites. The observations have been madeusing balloon-borne particle counters. The heavydashed line indicates the tropopause height. The fi­gure shows the well-known Junge layer in the stratos­phere. The latitudinal distribution shows greater aer­osolloading in the low latitudes. It also points to rela­tive paucity of aerosol measurements at lower lati­tudes. With the advent of satellite-based observations

(SAGE II), a better global coverage of aerosol charac­teristics is now available.

3 Brief Outline of Experimental Methods forAerosol StudiesExperimental methods for the study of atmospher­

ic aerosols can be classified into two categories: direct

AEROSOL DISTRIBUTION DURING SEP-OCl1972

209

INDIAN J RADIO & SPACE PHYS, VOL. 17, DECEMBER ]988

4 Studies on Atmospheric Aerosols in IndiaIndia Meteorological Department has been con­

ducting turbidity measurements at a number of se­lected Indian stations and the results have been pu­blished and discussed in literature23• These results es­

tablish the temporal- and location-dependent var­iations of atmospheric turbidity. At many of the loc­ations, a summer maximum of turbidity is recorded.In this article no attempt is made to cover these re­sults.

. -

402.~ ;0 3:5POWER LAW INDEX

o

E..•

4O~1---...••r----"""----"'·"""---1

by Jayaraman etal.22 at Hyderabad. The profiles byJayaraman et al. were obtained by the indirect methodof light scattering which may be sensitive to a widersize range than the profiles measured by the impactorwhich is sensitive to 0.1-1.0 ,urn range with a peak at0.15 ,urn. The difference between the two profilescould be at least partly due to this. Also, the profile ofChagnon and Junge was obtained in 1960 whereasthe profiles by Jayaraman et al. correspond to eigh­ties.

One prominent feature of the impactor profile isthe stratospheric aerosol peak. This is not seen prom­inently in the profiles by Jayaraman et al. The stratos­pheric peak of aerosols is mainly a large particle peak,i.e. particles> 0.15 ,urn and the smaller particles donot exhibit this peak4. As pointed out earlier the lightscattering method gives aerosol information coveringa much wider range including smaller particleswhereas the impactor is confined to a relatively smallrange (0.1-1.0 ,urn). This may explain the differencebetween the impactor profile and profiles of Jayara­man et al. with regard to stratospheric peak.

Considering the profiles by Jayaraman et al., thetropospheric aerosol number densities at Trivan­drum are much greater than those at Hyderabad,whereas in the stratosphere no such significant differ­ences are seen. The difference could be due to a reallatitude effect or because of the profiles correspond­ing to different years. This can be resolved only by si­multaneous measurements at the two locations.

Jayaraman et al., comparing the power law indieesobtained at Hyderabad and Trivandrum (Fig. 6),found that the Trivandrum values (observed in 1980)are much smaller than the Hyderabad values (of1984), indicating that the particles found over Trivan-

•10

n{----n.10.85 Hyd .. Joyaraman et 01. -18.f.·64 Hyd

~ •...•....••'ll€'on Feb.eO TvmChagnonand -Ap( 60 Hyd

Junge2S

.. ,10 10 10

AI!AOSOl NUNBEA MN$ITY (.-1)

•10

methods24 the columnar aerosol size distributionfunction can be derived from MWR observations.

A very powerful ground-based method for thestudy of aerosol characteristics with good altituderesolution is the lidar (acronym for Light Detectionand Ranging) method. In lidar (operated in monos tat­ic mode) a high power pulsed laser is used as a trans­mitter. The laser beam isdirected into the atmosphereby suitable optics. The backscattered light from theatmosphere is received by a suitable telescope and isdetected by a photomultiplier tube whose output isfurther amplified for recording. The backscatteredsignal strength is analysed to obtain information onaerosol characteristics25 - 27.

4.1 BaIloon-borne/Rocket-borne Measurements

Alitude profiles of aerosol extinction/numberdensity have been obtained using balloon-borne androcket-borne payloads22. These payloads make useof the scattered sunlight to derive the aerosol charac­teristics. The vertical profiles22 of aerosol numberdensity obtained are shown in Fig. 5 and the derivedpower law index in Fig. 6. Chagnon and Junge2S pu­blishea a vertical profile of aerosol number densityobtained by using a balloon-borne impactor at Hy­derabad. The impactor responds to the particle sizerange 0.1-1.0 ,urn with maximum sensitivity at 0.15,urn. This profile is also shown in Fig. 5 for comparis­on. The aerosol number densities in the impactor pro­file are much smaller than the latter profiles obtained

'0

;:c

10

Fig. :;- Aerosol numhcr density profiles ohtaincd at Hyderilhad(Hyd.) and Trivandrum (Tvm.)

Fig. 6- Power law index profiles ohtained at Hyderahad andTrivandrum

210

I' I II ", '" "l'I"1 " "1 '

KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

Fig. 7- Aerosol lll!mber density profile obtained from CW lidarobservations at Trivandrum

4.3 Pulsed Ruby Lidar Observations

A pulsed ruby lidar has been in regular operation atTrivandrum since October 1986 to study the charac­teristics of the altitude profiles of aerosols. The rubylidar operates at a wavelength of 694.3 nm with laserenergy of about 1 J per pulse. The pulse width is 4 flSand it is operated at a pulse repetition rate of 2 pulses/

of turbulence layers in the lower atmosphere can leadto accretion of aerosols between these layers. In theturbulent layers, aerosols would be well mixed hori­zontally whereas in the regions between them mixingwould be poor leading to accretion of aerosols. It isknown that the height of the nocturnal atmosphericboundary layer (ABL) over land (with clear skies andweak to moderate winds) is considerably low (200 m).Typically a stable sporadic regime would prevail un­der these conditions33• T.his regime is characterizedby narrow turbulent regions under favourable condi­tions, i.e. wind shears. Specifically, in regions wherethe Richardsonnumber(Ri) is quite small ( < 0.25)in­stabilities (Kelvin-Helmholts instabilities) grow andturbulence would prevail. Acoustic radar and radio­sonde observations showed stratified turbulence in

the nighttime lower atmosphere due to windshears33,34. Parameswaran et a/.3D examined the windand temperature profiles to obtain Richardson num­ber and found conditions to be favourable for stratifi­

ed turbulence when the layered structures in aerosolswere observed. Thus aerosols can act as tracers ofstratified turbulence in the lower atmosphere.

103 104

AEROSOl. NUMBER DENSITY Icm-3)

,5,\,,,

\~I \I ,I 4\

, ,'•••••• I....... _--" .••.

~)..•.

~ ~.•.

.............. ~.

2000

18001600E 1400~ 1200

:)...5 1000c800

600400200 L102

drum are relatively bigger than those over Hydera­bad. This they attribute to the influence of SierraNegra (lat. 0 .8°S,long. 91.2°W) volcanic eruption (13Nov. 1979) on the aerosol characteristics over Triv­andrum in 1980 (observations were made in Feb.1980). However, it is difficult to see how the volcaniceruption can influence the tropospheric aerosols at aplace located - 16,000 km from the eruption after 3months .

4.2 CW Lidar Measurements

As discussed above, lidars are capable of providingrange resolved information on aerosol characteris­tics. Lidars can be operated in monostatic or bistaticmode. In the monostatic mode, high power pulsed la­sers are used as transmitters. In the bistatic mode, CWlasers are used and this mode is particularly useful inthe studies on angular dependence of aerosol scatter­ing. But the range capability of this mode of operationis limited to lower tropospheric altitudes only ( < 4km). At Trivandrum, a CW lidar and a pulsed lidarhave been operated to study aerosol characteris­tics29,3o.

In the CW lidar experiment, argon ion laser operat­ingata wavelength of 514.5 nm was used as the trans­mitter. The receiving telescope was located at a hori~zontal separation of about 430 m from the transmit­ter. Using this experimental set-up, the angular de­pendence of atmospheric scattering has been studied.A method of solution of the bistatic lidar equation hasbeen developed to obtain the aerosol characteristicssuch as size distribution parameter and refractive in­dex. The altitude profiles of aerosol number densityhave also been obtained up to an altitude of - 2 kmfrom the bistatic lidar observations30.

Monthly mean profiles of aerosol number densitycorresponding to months preceding and followingthe south west monsoon season clearly showed thewashout effect of monsoon rains3'.

The aerosol number density profiles on individualdays showed layered structures in altitude (the pro­files correspond to the early night period of 1900­2000 hrs 1ST). Some typical profiles are shown in Fig.7. The fluctuations in the aerosol number density withaltitude are well above the altitude resolution and er­rors of measurements. Aerosols in the lower tropos­phere can be expected to show normally a smoothvariation with altitude especially at a tropical stationlike Trivandrum where convective mixing would bestrong. Since the fall velocities in the size range underconsideration (0.1 flm to 10 flm) are less than 2 m/hr(Ref. 32), the aerosols lifted by convection in the day­time would remain in the lower troposphere through­out the night. The observation of structured layers isnot in accordance with this behaviour. But presence

"

'\,

2ll

INDIAN J RADIO & SPACE PHYS. VOL. 17. DECEMBER 1988

2_

-"-220

~a216

..0:\ ~:r~21.3km

~T204

II

~'].e.m

9.3k",

~H'~ __ ~ I.akm

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taining the extinction profile. The bars parallel to theX-axis indicate the total error in the extinction values

contributed by various factors. The extinction profileshows a very prominent maximum in the altituderange 3-4 km. This feature is revealed by all the pro­files obtained. While the value of maximum extinctionshowed variation from month to month, the altituderange of its occurrence remained more or less thesame. During the local summer months the maximumis quite sharp compared to that in the winter months.This maximum has been attributed to change in aero­sol size distribution due to water vapour condensa­tion leading to an increase in the number oflarger par­ticles and/or an increase in the total number of parti­cles by vertical redistribution due to lack of mixing(turbulence )which may arise due to wind velocity gra­dients. It is worthwhile to study the effect of Hadleycell circulation in the vertical distribution of aerosols

in this regard.Study of variation of monthly mean extinction

shows that the variability is least in the middle tropos­phere (Fig. 9). The month to month variations in up­per(13.8 km)and lower tropospheres (1.8kmand 4.3km) are quite different. The variation in extinction inthe lower stratosphere (21 .3 km) is different from thatin the troposphere. This is quite understandable be­cause the stratospheric aerosols are mainly of localorigin rather than due to direct transport from tropos­phere. Fig. 9 also shows the monthly mean tempera­ture variation at 21.3 km. It is seen that stratosphericaerosol extinction at 21.3 km shows variations oppo­site to those in temperature indicating an inverse rela­tionship between the two. Modelling studies on stra­tospheric sulphate aerosol (which is dominant in the

..10I()' Ilf6

AEROSOL EXTINCTION (",-1)

11

20

min. The altitude resolved backscattered signal is re­ceived and analysed by solving the lidar equation toobtain the altitude profile of aerosol extinction. In thelidar equation, two unknowns, namely, the cxtinctionand backscatter function are involved. In order to

solve the Iidar equation, an empirical relationship be­tween the two is assumed based on theoretical calcu­

lation using Mie scatteringfunctions27. The altitude ofminimum in turbidity is identified from thc lidar ob­servations which is generally around the tropopauseheight and it IS assumed that at this altitude, the extinc­tion is contributed mainly by Rayleigh scattering, i.e.contribution due to aerosols to extinction at this alti­

tude is assumed to be negligible. With this assump­tion, the lidar equation is solved to obtain the total at­mospheric extinction profile. From this, the altitudeprofile of aerosol extinction is obtained by subtract­ing the contribution due to Rayleigh scattering. Usingthis method of analysis, altitude profiles of aerosol ex­tinction have been obtained from ruby lidar observ­ations on 5-10 days in a month in the early night peri­od. The error budget has been estimated andit has been found that the total error in theaerosol extinction contributed by variousfactors is, in general, less than 15 per centand is much less at lower altitudes « 10 km).From the aerosol extinction, the aerosol numberdensity is obtained using the Mie scattering theoryand assuming a aerosol refractive index of 1.5 and apower law index (v) of 3.5 for the size distribution. Atypical profile of aerosol extinction is shown in Fig. 8(it may be noted that the height resolution of this pro­file is much less than that of CW lidar profiles shownin Fig. 7, and hence the ruby lidar profiles do not showthe fine layered structures shown by CW lidar profilesin the lower height ranges). The discontinuity in theprofile between 16.5 km and 19.5 km is because theaerosol extinction at 18 km is taken to be zero for ob-

Fig. 8-Aerosol extinction profile obtained at Trivandrum fromruby lidar observations on 2 Oct. 1986

Fig. 9- Monthly mean variation ot aerosol extinction at fixed alti­tudes. The variation of monthly mean temperature obtained from

balloonsonde observations at 21.3 km is also shown.

212

I I jJ '. j".,il~II~. L

---------

KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

Fig. 10- Annual mean profile of aerosol extinction and numberdensity obtained from ruby Iidarobservations at Trivandrum

AEROSOL NUMBER OENSITY(I;;1

107 108 109

ing the available logistics into account.The radiome­ters give the total aerosol optical depth as a function ofwavelength (in the visible and near IR) after due ac­count is taken of the Rayleigh seatter and absorptioncomponents.

Agashe and Mahajan38 and Nair et al.37 reportedmorning to evening differences in the aerosol opticaldepth on some occasions at Pune and Trivandrum.The effect of relative humidity on aerosol opticaldepths has been studied by Krishna Moorthy et al.39

using Trivandrum data. They reported a steep in­crease in optical depth at relative humidities greaterthan 70 per cent which is in agreement with theoreti­cal estimates based on growth of aerosol particleswith increase in humidity40. Nair et al. 37 reported sea­sonal v.ariation in wavelength dependence of aerosoloptical depth. Fig. 11 shows the wavelength depend­ence of aerosol optical depth for the three~easons,namely, summer, monsoon and winter months. Thepeak in optical depth at 935.0 nm is due to absorptionby atmospheric water vapour which has a strong ab­sorption band around this wavelength (measure­ments at 935 nm were made to determine the total co­

lumnar water vapour content). Nair et af.37 inferred

from the seasonal variations (Fig. 11) a monp.omOdalstructure of the columnar aerosol size distribution insummer and monsoon and a bimodal structure inwinter months. They attributed this to the presence ofdominant source of aerosols in the summer and mon­soon months in contrast to winter months.

IMAP aerosol campaigns-In the preceding sec­tions, a summary of observational results on aerosolsobtained in India is given. In the Indian Middle At­mosphere Programme, a coordinated experimentalprogramme is being carried out in a campaign modeto study the aerosol characteristics. The experimentsinclude rocket- and balloon-borne experiments (scat­tering method) and ground-based (lidar and MWR)measurements for the determination of aerosol char­

acteristics. A pre-campaign has already been carriedout in October 1986 from Trivandrum (rocket- andground-based experiments) and in April 1984 andOctober 1985 from Hyderabad (balloon-borne ex­periments). The results of this pr~-campaign indicatean altitude dependent aerosol size distribution. Theyalso indicate that the size distribution could be· quitedifferent from an inverse power law. The results havebeen reported in a scientific report41.

The main campaign has been planned to be carriedout in two phases: phase 1in premonsoon period andphase 2 in postmonsoon period. Phase 1has iilreadybeen carried out in April 1988. The experiments in­cluded (i) balloon-borne measurements of aerosols,ozone and electrical conductivity from Hyderabadand (ii) rocket-borne and ground-based experiments

-410

_ Annuol •••eon 0••.0.01 fttlftd""___ Annual ••• " aerosol nu__

_ .Ity

-7 -6 -510 10 10

AEROSOL EXTINCTION 1m-I)

-810

32

2&24! 20III

Q::216.. ....•c 12

•4

Rose et ap6 have also obtained an annual meanprofile of aerosol extinction (Fig. 10).This is the meanof sixty profiles corresponding to the period October1986 to September 1987. The bars parallel to X-axisindicate the standard deviations at the correspondingaltitudes. The mean aerosol number density profile isalso shown in the same figure. These profiles can beconsidered as reference profiles of aerosolextinction/number density at a tropical coastal sta­tion like Trivandrum.

4.4 Spectral Measurements of Total Extinction with MWR

As part of the Indian Middle Atmosphere Pro­gramme, multi-wavelength radiometers (MWRs)have been deployed at Trivandrum, Mysore, Pune,Visakhapatnam, Jodhpur and Delhi (Refs 37 and 38);these places have been selected pased on the geogra­phical and environmental conditions and also by tak-

stratosphere) formation due to homogeneous heter­omolecular and heterogeneous heteromolecularnucleation processes show that the formation pro­cesses are less effective at higher ambient tempera­tures 11 • Wang and McCormick35 from a study of asso­ciation between aerosol extinction (obtained from sa­tellite instrument SAM II) and temperature reportedboth negative and positive correlations between thetwo. The negative correlation was attributed to thestratospheric microphysical processes (formationprocesses) and the positive correlation to the hori­zontal eddy transport.

113

INDIAN J RADIO & SPACE PHYS, VOL. 17, DECEMBER 1988

....

Fig. 12-Solar radiationspectra(from Coulson42)

5.1 Effect of Aerosols on the Atmospheric Heating

Aerosols absorb energy from the solar and plane­tary radiation fields, exchange energy by collisionswith the ambient gas and emit thermal radiation. Incase any phase change is possible, they acquire or loselatent heat. The interaction with the radiation field is afunction of the size distribution and the refractwe in­

dex of the particles and the spectral density of the ra­diation. The collisional heat exchange with the am­bient gas depends upon the temperature and densityof the ambient gas and also upon the temperature of

The radiation field should be defined at each pointin space and for each wavelength for the study of in­teraction of radiaiton with the atmosphere and result­ing aeronomic and thermal effects. The analysis of theradiation field can be done from the equation of radia­tive transfer43.44.The radiative transfer equation ex­presses the energy balance in each unit volume of theatmosphere. including absorption, scattering andemission. The equation can be solved by analyticalmethods only for the most simple cases. To obtainquantitative solutions numerical methods are gener­ally used45.

The most important elements contributing to themiddle atmosphere radiation budget in terms of cool­ing and heating rates are shown in Fig. 13. The figureshows the vertical distribution of solar short wave

heating rates by 03' 02' N02, H20 and CO2 and ter­restriallong wave cooling rates by CO2, 03 and H20.The mean heating due to 03 is balanced by the corre­sponding cooling due to infrared radiation from CO2(15 ,urn), 03 (9.6 ,urn) and water vapour (18 ,urn). Nearthe stratopause the 15 ,urn band of CO2 yields a radia­tive temperature change (net cooling) about twice aslarge as the 9.6 ,urn band of 03 and about 10 timeslarger than that from the 18 ,urn band of water vapour.

• 1914

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SOLAR RADIATION

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0·80,60·40·20·0400

600 800 1000WAVElENGTH (nm)

5 Effect of Aerosols on Radiation in theMiddle AtmosphereNo detailed discussion attempted here on the much

more extensive topic of radiation budget in middle at­mosphere, which is also beyond the scope of this pres­entation. However, for the sake of completeness a fewof the important aspects of this are being presentedhere.

The solar radiation spectra at the top of the atmos­phere and at the sea level are shown in Fig. 12.The ab­sorption due to various molecular species is also indi­catt:d at the corresponding spectral regions. The dif­ference between the top of the atmosphere and the sealevel curves is due to scattering (apart from the molec­ular absorption bands which are marked).

Fig. II-Wavelength dependence of aerosol optical depth forsummer, monsoon and winter seasons in the years 1984. 1986 and

1987. The peak at 935 nmis due to water vapour absorption.

on aerosols from Trivandrum. The experiments atTrivandrum were conducted within two days of theirbeing done at Hyderabad to enable comparison.

214

! I, "

!t!,~~!UIl!,!~; ! ~:Jl.l.IIi;Ul.uIlUU,I~, j, ,I.'i.j 1,I,j,

KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

.5cx

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----IIIr /-TOTAL -t-- "..

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><2.0

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01 , , I •••• I ••.• ' I JI

I I, ••• I •••• I •• II.IJI10

!l00 !l10a:/COOlING (K/elay)

HEAnNG (K/doy)~I

~--

I.IJ IX

Fig. 13-Vertical distribution of solar short wave heating rates byD." O2, N02, H20 and CO2 and terrestrial long wave cooling rates

by CO2, D., and Hp(from London"")

Fig. 14-Real and imaginary refractive indices of aerosols as func­tions of wavelength (from Ivlev and Popova49)

Mugnai et al. used the zeroth order logarithmic dis­tribution function (Zold) suggested by Espenscheid etal.50 for the aerosol size distribution. This is given by

the particles. Thus the equilibrium conditions for agiven particle size distribution are a function of the al­titude, time of the day, season and planetary albedo.The energetic equilibrium of aerosols in the atmos­phere has been studied in detail by Fiocco et al.47 andMugnai et al.48• Their analysis is based on the equilib­rium condition for a spherical aerosol particle of radi­us r, expressed by

dN z z~= Cz exp [-(In r-ln rm) l2(ln a) ]dr

... (19)

where Pa,s is the power absorbed from solar radiationby the particle, Pa,pla the power absorbed from plane­tary radiation, p., the thermal power radiated by theparticle, and p" the power lost by the particle due tocollisions with the ambient gas. Fiocco et al.47 ana­lysed these terms in the equilibrium equation usingthe complex refractive index values for the syntheticaerosol model of Ivlev and Popova49 (shown in Fig.14) and for summer and winter conditions at 45°N us­ing the US standard atmosphere. In the particle sizeinterval 0.1 to 0.3 ,urn, daily average heating rates forsummer are 0.15-0.25 KI day and for winter these are0.07-0.1 K/day in the stratosphere.

Mugnai et al.48 carried out the analysis for particleshaving special significance to stratosphere. Their an­alysis was for the realistic case of polydisperse parti­cles instead of mono disperse particles as in the analy­sis by Fiocco et a/.47 They carried out analysis for (a)sulphuric acid droplets with no impurities and there­fore with no absorption in the visible, (b) impure sul­phuric acid droplets with a small amount of absorp­tion in the visible (imaginary refractive in­dex = 0.001), (c) impure sulphuric acid droplets withhigher absorption in the visible (imaginary refractiveindex = 0.005), and (d) volcanic ash.

where Cz is a constant whose value depends on theaerosol concentration, rm is the mode radius, and a isrelated to the width and the skewness of the distribu­tion.

From the values of p.: for each particle size, Mugnaiet al. obtained the heating rates (K/ day) from the ex­pressIOn

where T is the air temperature, x the aerosol massmixing ratio, Ps the particle density assumed(= 2.3 g/cm3 for volcanic ash particles and 1.6 g/cm3

for sulphuric acid droplets), and Cp the specific heat atconstant pressure. Aerosol mass mixing ratio xis tak­en to be constant with respect to altitude with a valueof 10 - 9. g/ g of air. The daily average heating rate forthe polydisperse aerosol given by the Zold distribu­tion is calculated from the daily average heating ratefor each particle size in the range 0.01-10 ,urn.

Fig. 15 shows the daily average heating rate as afunction of the mode radius rm for a= 1.8 and 2.2 andrefractive indices as specified for sulphuric acid parti­cles (a), (b) and (c) and volcanic ash particles (d). Theresults shown in Fig. 15 apply for equatorial atmos­phere for which the US standard atmosphere 1966

Pa.s + F..pla - Po - Pc = 0 ... (18)

dT=3xPcl4nr3 gCpdt... (20)

215

INDIAN J RADIO & SPACE PHYS, VOL. 17, DECEMBER 1988

v-=1.8

I 01

o

(b)001

~101

MODE RADIUS (~m)

Fig. 15- Daily average heating rate (K/ day) as a function of mode radius of aerosol size distribution

for equatorial conditions for a= 1.8 and 2.2 (from Mugnai etal.4~)

for 15°N is used. Fig. 15 shows that the heating ratesfor all the models (a), (b), (c) and (d) maximize aroundthe altitude of stratospheric sulphate (aerosol) layer.An increase in the heating rate from the sulphuric acidmodels (a) to (c) is quite evident (with increase of im­aginary part of refractive index) and is due to the in­creasing absorption of solar radiation by the particles.The heating rates for the volcanic ash mGdel (d) areslightly smaller than those for the sulphuric acid parti­cle model (c) but greater than tl}ose for (a) and (b).There is also a general increase of the heating ra te withthe mode radius rffi' An increase in amoves the heat­ing rate pattern to the left, towards smaller values ofrffi' Fig. 15 shows that there are large variations ofheating rate with height and aerosol composition. Theheating rates are always positive indicating that aero­sols add heat to the ambient gas by conductionthroughout the stratospheric aerosol layer region.These heating rates of aerosol when compared toother heating rates in the same region (shown in Fig.15) are quite small. However, during volcanic erup-

216

tion when stratospheric aerosol loading increases en­ormously, the aerosol heating rates can become signi­ficant. For example, the results of Cadle et al.51 indic­ate that fine ash particles from the Mt. Agung eruption(in 1963) were located during the winter of 1964/65in a broad latitude over the equatorial region with apeak mixing ratio of25 x lO-9 g/g of air at about lOoSat 20 km altitude. Thus the calculations of Mugnai etal.48 show that this would result in a heating rate of0.25 K/ day. Over and above this the heating rates dueto sulphuric acid particles (due to volcanic eruption)are to be added. Thus it is evident that during majorvolcanic eruptions, the aerosol heating rate becomesquite significant to effect the ambient thermal struc­ture. Fig. 16 shows plots of temperature deviations atthree levels (19.5,16.5 and 9.5 kIn) obtained at PortHedland (Australia) from balloon observations dueto volcanic eruption ofMt. Agung in 1963 (Ref. 52).Monthly means were calculated for five years beforethe eruption and deviations from the means werecomputed. The three month running averages of the

" I II ", " "111"1" "I '" I ;'11111'; Ii

KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

Fig. 17- Radiative convective equilibrium temperature distribu­tions of the earth's atmosphere (from Wang and Domoto54)

transfer equation. In the analytical models oh:limatechange due to aerosols, the change in the earth-at­mosphere system albedo from unperturbed (withoutaerosols) condition (A ) to perturbed (by introduc­tion of an aerosol layer in the atmosphere) condition(A ') is obtained. If A' >A, the aerosol layer will re­duce the energy of the earth-atmosphere system(causing a 'cooling') and if A' <A it will increase theenergy of the system (causing a 'heating'). The aerosollayer is characterized by the parameters a and b,where a is the fraction of incident radiation absorbedby the layer and b is the fraction of incident radiationbackscattered by the layer. Russel and Grams57showed that for an 'optically thin' aerosol layer, thecritical value Oc of 0 ( = a/b) is approximated by

Oc=(1-A)2/2A ... (21)•

The models predict that an aerosol layer witho > Oc will cause a decrease in system albedo A and alayer with 0 < Oc will cause an increase. Fig. 18 showsOcas a function of surface albedo A. The regions of in­creased and decreased energy to the earth-atmos­phere system are marked in the figure. Table 3 givesthe values of Oc for different surface types.

Harshvardhan and Cess58 gave an expression forthe surface temperature change (Ii. T.)due to a stratos­pheric aerosol layer with total optical thickness 'l' by

f.,.T;= -{1-[(1-0.37'l')/(1-0.11'l')]

-1.96 X 10-11[T]4 'l'} x 162 ... (22)

where T is the ambient temperature in the aerosollayer (in stratosphere). Although this expression isnon-linear in 'l', it will turn out to be practically linear

Ollll ~ IlAK ••••••••••

••• •• NAZI " •••••••••

000 • aIMtlIlIl

••• Q.IM --.. •.••••....

220 UO 140 no zeo 210 210 zeo 100

TEMPERAT'-"E (K)

o

1000_IOI~ 210

900

100

100

100

200

":" 3001o.Q! 4001

~

rJ-c~~~Fig. 16-Temperature changes in stratosphere due to volcanic er­

uptionoiMt.Agungin 1963 (from Newe1l52 )

5.2 Climatic Effects

The effect of aerosols on earth's climate has been

investigated by a number ofworkers55,56. The inves­tigations on this are based on analytical methods ornumerical methods involving solution of radiative

deviations are shown in Fig. 16. The figure clearlyshows a large temperature increase at the two higherlevels of 16.5 and 19.5 km (stratosphere) associatedwith theMt. Agung eruption in 1963.There isno sign­ificant effect noticeable at 9.5 km. However, an indi­cation of a decrease is seen. An increase in stratos­pheric temperature is reported due' to the EI Chichonvolcanic eruption.in 1982 (Ref. 53).

Aerosols in the stratosphere reduce the solar radia­tion flux by absorption that is incident in the tropos­phere. Thus an increase in the aerosol loading in thestratosphere after major volcanic eruption couldcause tropospheric cooling (an indication is seen inFig. 16 at 9.5 km). The extent oftropospheric coolingis proportional to the solar energy lost by the tropos­phere.

The radiative effects of the aerosols on the tropos­phere have been investigated employing numericalsolutions of radiative transfer equation. The radiativeconvective temperature distribution in the tropos­phere would be modified by the presence of aerosols,the effect being a general cooling effect in the tropos­phere. To illustrate the effect, the radiative convectivetemperature distributions as obtained by Wang andDomoto54 are shown in Fig. 17. It shows the distribu­tions for cloudless and average cloudy conditions(mean cloud albedo assumed to be 0.55 throughoutthe solar spectrum) and for clear (without aerosols),light haze and dense haze atmospheres. As seen fromthe figure the effect of aerosols is to reduce the tropos­pheric temperatures. In general, the effect of aerosolsis to decrease the solar radiation reaching the earth'ssurface and increase the global albedo of the earth-at­mosphere depending upon the absorption and scat­tering properties of aerosols.

217

INDIAN J RADIO & SPACE PHYS, VOL. 17, DECEMBER 1988

t- -

8 Turco R P,Whiuen R C, Toon 0 B, Pollack J B & Hamill P,Na­ture( GB), 283 (1980) 283.

9 Turco R P, Toon 0 B, Whitten R C, Inn E C Y & Hamill P, IGeophys Res (USA), 86 (1981) 1129.

10 Turco R P, Whitten R C & Toon 0 'B, Rev Geophys &SpacePhys( USA), 20 (1982) 233.

11 Yue G K & Deepak A,I Geophys Res( USA), 87 (1982) 3128.12 Hanel G, Adv Geophys (GB), 19 (1976) 73.13 Hunten D M, Turco R P & Toon 0 B, I Atmos Sci (USA), 37

(1980) 1342.14 Turco R P, Toon 0 B, Hamill P & Whitten R C, I Geophys Res

(USA), 86(1981) 1113.15 Bowen E G,J Meteorol( USA), 13 (1956) 142.

16 Junge C E, A ir Chemistry and RadioActivity(Academic Press,New York, USA), 1963.

17 Tur,o R P, Toon 0 B, Pollack J B, Whitten R C, Popoff I G &Hamill P, I Appl Meteorol( USA), 19 (1980) 78.

18 Oberbeck V R, Farlow N H, Ferry G V,Lem H Y & Hayes D M,Geophys Res Lett (USA), 8 (1981 ) 18.

19 RosenJ M,HoffmanDJ &PepinT J, ProcIntConfStru Compand Gen Cirofthe Upper&LowerAtmos( USA), 1 (1974)lOt.

20 Hofmann D J, Rosen J M, Pepin T J & Pinnick R G,J Atmos Sci(USA),32(1975) 1446.

21 De BaryE & Rossler F,J Geophys Res( USA), 71 (1966) 1011.

22 Jayaraman A, Subharaya B H & Acharya Y B, Phys Scr(Swed­en), 36 (1987) 358.

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INCREASED ENERGY

()O2 ()04 0·6 008 lo()

SURFACE ALBEDO A

DECREASED ENERGY

10

100

()O()l

o

..J«~ 0·1I-a:u

..JoIIIoa:ILl«

Table 3- Values of Oe for Different Surface Types57

Peterson J T & Junge CE, Man's Impact on Climate (M.I.T.Press, Cambridge, USA), 1971.

2 Hidy G W & BrockJ R, Proceedings of the Second Internation­al Congress (Academic- Press, New York, USA), 1971. 9.

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mos Sci( USA), 36 (1979) 699.

Surface Type A15,

Urban areas

0.20t.6Deserts

0.300.82Praiaries and farm lands

0.201.6Forests

0.162.2Oceans

0.085.3

Snowfields

0.700.064r

References

Fig. 18-Relationship of critical aerosol parameter 15, to surfacealbedo A for optically thin aerosol layer (from Russel and

Grams'7)

for change in optical thickness due to an increase in i .• 28

aerosol loading. This expression is obtained assum- \ 29ing sulphuric acid aerosols. The value of ~ T"from the \

above equation is calculated tobe - 0.8Kfor r= 0.03 f 30and T = 2140 K. Thus the model predicts a decrease .in surface temperature due to an aerosol layer (stra­tospheric). This means that the increased solar albedoeffect dominates over the increase in trapping or'greenhouse' effect due to a stratospheric aerosol lay­er.

218

, I " ,,, II 'I'!'I'I ""1 I.,

KRISHNA MURTHY: AEROSOLS & RADIATION BUDGET IN MIDDLE ATMOSPHERE

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219