effect of aerosol variation on radiance in the earth’s atmosphere-ocean system

Effect of Aerosol Variation on Radiance in theEarth's Atmosphere-Ocean System

Gilbert N. Plass and George W. Kattawar

The reflected and transmitted radiance is calculated for a realistic model of the atmosphere-ocean system.

Multiple scattering to all orders as well as anisotropic scattering from aerosols are taken into account by

a Monte Carlo technique. The probability for reflection or refraction at the ocean surface is calculated

for each photon. Scattering and absorption by water molecules (Rayleigh) and by hydrosols (Mie) are

taken into account within the ocean. The radiance is calculated for a normal aerosol distribution as

well as for a three and ten times normal distribution. Calculations are also made for an aerosol layer

near the earth as well as for one in the stratosphere. The upward radiance at the top of the atmosphere

depends strongly on the total number of aerosols but not on their spatial distribution. Variations in the

ozone amount also have little effect on the upward radiance. The calculations are made at the following

wavelengths: 0.7 u, 0.9 A, 1.67 M.at various levels are also discussed.

1. IntroductionThe radiation field of the earth's atmosphere is deter-

mined by the multiple scattering of photons from theaerosols and molecules within the atmosphere as well asby the interaction of the photons with the earth's sur-face. The interaction with the surface is particularlycomplicated for the majority of the earth's surface thatis covered by an ocean. Here the photons may be re-flected from the surface or refracted into the ocean.After scattering within the ocean the photons may re-emerge into the atmosphere and contribute to the radi-ance there. A realistic model for photon scatteringwithin the earth's atmosphere must include these fea-tures: (1) Rayleigh and aerosol scattering with theproportion varying with height; (2) aerosol phase func-tion with strong forward scattering; (3) aerosol numbervarying with height; (4) ozone absorption.

Calculations of the radiation field of the atmosphere-ocean system have been made by Waterman and Wes-tell,' Ivanoff and Waterman, 2 Preisendorfer, 3 and Neu-mar and Hollman.4 However, all these authors hadto make certain simplifying assumptions that includean approximate treatment of multiple scattering, theuse of a Rayleigh scattering function for the aerosolsand hydrosols instead of a more realistic function withstrong forward scattering, and the neglect of the in-

The authors are with the Physics Department, Texas A&MUniversity, College Station, Texas 77843.

Received 14 January 1972.

The radiance above and below the ocean surface as well as the flux

fluence of the radiation present in one medium uponthat in the other. An excellent review together withdata based on observations has been given by Duntley.5

With Monte Carlo techniques it is possible to calcu-late the complete radiation field in the atmosphere-ocean system with a realistic model for both the atmo-sphere and ocean. In the atmosphere, both Rayleighand aerosol scattering are included in the model, thelatter represented by a scattering function with a strongforward peak. Absorption by various atmosphericcomponents is also taken into account. Both the re-flected and refracted rays, as well as the rays thatundergo total internal reflection, are followed from theocean surface. The appropriate absorption and scat-tering by the water molecules and hydrosols of the oceanare included in the calculation. The scattering func-tion for the hydrosols as calculated from a distributionof sizes by the Mie theory agrees well with measure-ments and has a strong forward maximum. The totalradiation field is obtained as the three-dimensionalpaths of the photons are followed by the Monte Carlomethod.

The purpose of the present study is to obtain theradiance at the top and bottom of the atmosphere witha realistic model of both the atmosphere and ocean.The number of aerosols in the atmosphere as well astheir distribution with height are varied in order todetermine the eff ect of aerosol changes on the radiance.The calculations are done at three wavelengths: 0.7 ,u,0.9 a, 1.67g.

Although these calculations were performed fornormal incidence, namely Ao = 1, the principle of rec-iprocity can be used to obtain the radiance at other

1598 APPLIED OPTICS / Vol. 11, No. 7 / July 1972

'.L I Q I I flIDU I III 1WE 40 -c- RAYLEIGH

30- ~~~~~~~~AEROSOLSj_ -X X

20- o X

I0-

0_

io-6 I0-5 I-4 10-3 10-2 101- I 2

ATTENUATION COEFFICIENT (km-)Fig. 1. Attenuation length for aerosol scattering, ozone absorption, and Rayleigh scattering as a function of altitude. Five differentaerosol distributions are shown (models A-E) together with three ozone distributions (normal distribution used with models A-E and

two variations, models F and G).

solar zenith angles. Symmetry relations for the reflec-tion and transmission matrix have been rigorouslystudied by Hovenier.6 In this paper we have neglectedpolarization effects, and the reciprocity principle as-sumes the following simple form for the reflected radi-ance I:

Yu(,u¢;,olko) =- ,oI'(,uofo; ,O),

where I(O,A,k; oto) is the radiance at the observationangle specified by 4,+ due to an incoming solar beamspecified by o,0o,. The cosine of the zenith angle is ,;the azimuthal angle measured from the incident planeis 4. The reciprocity relation will also hold in the caseof coupling the atmosphere-ocean radiation fields. Thiscan best be imagined physically since the interface canbe thought of as a degenerate reflecting volume possess-ing the same symmetry as a scattering function.

I. Method of CalculationThe aerosols are represented by particles with a real

index of refraction of 1.55 and with radii distributedaccording to the haze C model proposed by Deirmend-jian.7 In this model, the number density is constantfor 0.03 < r < 0.1 , and is proportional to r 4 forr > 0.1 , where r is the aerosol radius. The singlescattering function was calculated for this distributionof spherical particles, from the Mie theory for the wave-lengths used in this calculation by the method discussedby Kattawar and Plass.' 9 The values of the singlescattering function are shown in Fig. 1 of Plass andKattawar.' 0 At each altitude a number is chosen torepresent the fraction of the photon collisions that are

with aerosols and the fraction with Rayleigh scatteringcenters.

In the ocean the photon may collide either with aRayleigh scattering center or with a hydrosol. A singlescattering function for the hydrosols was also calculatedfrom Mie theory. The particulate matter in the waterwas assumed to have n = 1.15 and n2 = 0.0001, wheren1 and n2 are the real and imaginary parts, respectively,of the index of refraction with respect to water. A sizedistribution was assumed that is constant for r < 1 and proportional to r exp(-2r) for r > 1 , where r isthe radius of the hydrosol. The modal radius for thisparticle size distribution is 3 u. The single scatteringfunction was computed from the Mie theory for par-ticles with this size distribution and index of refraction..The scattering function is highly anisotropic withstrong forward scattering. It is shown in Fig. 1 ofPlass and Kattawar.1" The agreement with measuredvalues is also discussed there.

The realistic model used to represent the earth'satmosphere, called model A in this paper, is describedby Plass and Kattawar. 12 Briefly the Rayleigh at-tenuation coefficient, ozone absorption coefficient, andthe aerosol number density as a function of height andwavelength were taken from the tables compiled byElterman 3 (earlier tables by Elterman were used byPlass and Kattawar 2 ). The total optical thickness ofthe atmosphere was calculated from the Rayleigh andaerosol attenuation coefficients and the ozone absorp-tion coefficients. The atmosphere was divided into anumber of layers, and the ratio of the Rayleigh extinc-tion to total extinction coefficient and the scattering toextinction coefficients for both the Mie and Rayleigh

July 1972 / Vol. 11, No. 7 / APPLIED OPTICS 1599

Table I. Attenuation Lengths in Ocean

X (A) 1S,R/PT,? 0S,/0T,1 3TR/PT

0.7 4.95 X 10-3 0.833 0.9140.9 1.04 X 10-4 0.333 0.9871.67 8.4 X 10-8 0.00498 0.992

scattering centers were established for each layer.The probability of ozone absorption was also calcu-lated from the ozone absorption coefficient for eachlayer. All calculations were done with the opticaldepth as the parameter.

The absorption of the ocean changes by several ordersof magnitude as the wavelength increases from 0.7 A to1.67g,. The model of the ocean requires that we specifyat each wavelength the relative contribution of Ray-leigh and Mie scattering as well as the fraction of radia-tion absorbed by each type of scattering center. Theassumed values are given in Table I for the followingratios: S,R/1T,R, S,M/3TJ, and 3T,R/3T, where sand T represent scattering the total attenuationlengths, respectively, and R and M represent the partdue to Rayleigh and iVlie scattering centers.

Five different aerosol distributions together with thenormal ozone distribution are used in these calculations.Model A uses the normal aerosol distribution as givenby Elterman" as shown by the solid line in Fig. 1.Model B corresponds to three times the normal aerosolamount and model C to ten times the normal amount.Model D has the normal aerosol distribution above 1 km(same as model A), but a heavy aerosol layer betweenthe ground and 1 km, such that the total aerosol amountis three times normal. Model E has a thick layer ofaerosols between 20 km and 25 km corresponding to tentimes the normal amount and has the normal distribu-tion (same as model A) at all other altitudes.

Two other models were used to study the effect of

ozone variations. Model F uses the same aerosoldistribution as model B and the ozone distributionindicated by open squares in Fig. 1. The total ozoneamount for this model is 0.23 cm, whereas it is 0.35 cmfor the normal ozone distribution indicated by the solidline in Fig. 1. i\Iodel G uses the same aerosol distribu-,tion as model B and the ozone distribution indicated bysolid squares in Fig. 1 (total ozone amount is 0.47 cm).These models were chosen to represent a range of valuesthat occur in the earth's atmosphere. The character-istics of the different models are summarized in TableII.

The details of the Monte Carlo calculations have beenpublished previously. 0 4 5 Briefly, the exact three-dimensional path of the photon through the atmosphereand ocean is calculated after each collision. At eachcollision in the atmosphere a choice is made betweenRayleigh scattering, aerosol scattering, and ozone ab-sorption based on the appropriate probabilities for thealtitude in question. The new scattering angle is calcu-lated from the single scattering function for the process

in question. When the photon hits the ocean surface,

the probability for reflection is calculated for the appro-

priate angle of incidence. A choice is made as towhether the photon is reflected back into the atmo-sphere or refracted into the ocean based on this prob-ability. If the photon enters the ocean, a new direc-tion is calculated from Snell's law. At each collision inthe ocean a choice is made between molecular absorp-tion or scattering (Rayleigh) and hydrosol absorptionor scattering (Mie). When the photon reaches theocean surface from within the ocean it is totally inter-nally reflected at some angles, while at other angles achoice is made between reflection and refraction backinto the atmosphere. Numerous variance reductiontechniques are used in the calculation as explained inthe Appendix of the paper by Plass and Kattawar.l

Il1. Radiance at 0.7 ,The aerosol attenuation coefficient is greater than the

Rayleigh attenuation coefficient at all altitudes less than35 km according to the 1969 Elterman aerosol distribu-tion. The ozone absorption coefficient is greater thanthe Rayleigh attenuation coefficient at all altitudes be-tween 19 km and 50 km for the normal ozone distribu-tion. The ocean is a moderately strong absorber atthese wavelengths. Approximately sixteen photonsare scattered by hydrosols for each photon scattered bya Rayleigh center within the ocean.

The upward radiance at the top of the atmospherefor the various models described in Sec. II is shown inFig. 2. The curve (crosses) for the normal aerosol andozone distributions shows a pronounced minimum nearthe zenith. These curves are normalized to a source ofunit incident flux. In all the figures in this paper, boththe incident and diffuse flux are included in the plottedvalues. The incident flux in the zenith direction isconverted into radiance by the assumption that it isrecorded by a detector with an acceptance half-angleof 9.10.

The curve for model B (open circles) gives the resultfor three times the normal aerosol amount and the

Table II. Characteristics of Aerosol-Ozone Modelsa-b

Model Aerosol number density Ozone absorption coefficient

A Based on Ref. 13 Based on Ref. 13; 0.35-cmtotal amount

B Three times model A Same as model AC Ten times model A Same as model AD Aerosol layer between Same as model A

0 km and 1 km; totalaerosol number threetimes model A

E Ten times model A for 20- Same as model A25-km region; unchangedfrom model A at otheraltitudes

F Three times model A 0.23-cm total amount; seeFig. 1

G Three times model A 0.4-cm total amount; seeFig. 1

a Rayleigh attenuation coefficient vs altitude based on Ref. 13.b Rayleigh, aerosol, and extinction ratios for ocean given in

Table I.


-0- MODEL D-U- MODEL E-0- MODEL F

I I

X-0.7/iL

-*~~~~~~~~~~~~~~~~~~-__

I I I I . I . I0.2 0.4 0.6 0.8 1.0

Fig. 2. Upward radiance at top of atmosphere as a function ofcosine of zenith angle, pu, at X = 0.7 p. Results are shown formodels A-F. The radiance is normalized to unit incident flux.The direct solar beam is included in the radiance measured by the

detector nearest the zenith (see text).

normal ozone distribution. There is still a minimumnear the zenith, but the upward radiance is greater atall angles than for model A (except for the detectorrecording the direct solar beam). The upward radi-ance is greater and relatively independent of zenithangle for model C (solid circles) with ten times normalaerosol amount and normal ozone distribution.

The next two models were used to study the influenceof variations with height of the aerosols on the upwardradiance. Model D corresponds to a strong aerosollayer in the lowest 1 km of the atmosphere, a distribu-tion corresponding to model A above 1 km, and totalaerosol amount equal to that of model B. The results(open squares) shown in Fig. 2 show no significantdifferences from the curve for model B. Thus the totalnumber of aerosols and not their distribution withheight is the most important factor in determining theupward radiance. The results for model E, which hasthe same aerosol distribution as model A except for alayer from 20 km to 25 km with ten times normal aero-sol amount, are given by the curve with solid squares.The upward radiance is slightly greater at most anglesfor this case than for model A. The increased radianceis probably due almost entirely to the fact that the totalnumber of aerosol particles is 177 greater for model Ethan for model A.

The next two models are designed to study the effectof changes in the ozone distribution on the upwardradiance. Models F and G use the ozone distributionsshown in Fig. 1 together with the aerosol distributionof model B. There is no significant variation in the

upward radiance with these variations in the ozonedistribution as is shown in Fig. 2 (diamonds); the re-sults for model G are not shown since they are virtuallyidentical to those for model F.

The downward radiance just above the surface of theocean is shown in Fig. 3 for these same model atmo-spheres. When the aerosol amount is normal, there isa minimum in the curve at pu 0.45. For greateraerosol amounts the downward radiance increases uni-formly from the horizon to the zenith. The results formodel D are again nearly identical with those for modelB, indicating that the total number of aerosols ratherthan their spatial distribution is the most importantfactor in determining the downward radiance. Theiris no dependence of the results on the ozone amounts.(The curve for model G is not plotted as it is virtuallyidentical with the curve for model F.)

IV. Radiance at 0.9 AThe aerosol attenuation coefficient is greater than the

Rayleigh attenuation coefficient at all altitudes less than42 km according to the 1968 Elterman aerosol distribu-tion. There is no ozone absorption at this wavelength.The ocean is a strong absorber at this wavelength.Approximately forty-three photons are scattered byhydrosols for each photon scattered by a Rayleighcenter within the ocean.

The calculated upward radiance is shown in Fig. 4.The value of the radiance is smaller in most cases atX = 0.9 p than at 0.7 pi. However, the curves are

3 I . I I . I I

2 -X- MODEL A-0- MODEL B-- MODEL C X-0.7.-0- MODEL D_-- MODEL E-0- MODEL F

0z

0::

3t 10lJ*ro OJ °7 :4: o or 3

0 301

- &oX2CrCE roux Fx

2x1 2 I . I I I . I I0.0 0.2 0.4 0.6 0.8 1.0

ILFig. 3. Downward radiance just above ocean surface as a func-

tion of pu at X = 0.7 pu for models A-F.


2x10'i

l0-

I - I I-X- MODEL A-0- MODEL B-0- MODEL C

'a 0.0_ . . . . . . . . . . . .

- 1lU

-r- - I I I .

r- * ~~~ *LnJn'~~~~~~~~~~~~

-X- MODEL A-0- MODEL 8-0- MODEL C-0- MODEL D-M- MODEL E X:Q9.g

I . I i I I . I

0.2 0.4 0.6 0.8 1.0

PL

Fig. 4. Upward radiance as a function of M at Xels A-E.

= 0.9 ;z for mod- I.Fig. 5. Downward radiance as a function

models A-E.of g at X = 0.9 1i for

qualitatively the same at both wavelengths and stillshow no appreciable dependence on the aerosol distribu-tion with height.

The calculated downward radiance is shown in Fig.5. The downward radiance is less at X = 0.9 p than at0.7 p for the normal aerosol distribution. The differ-ences between the two wavelengths are smaller for theother aerosol distributions.

V. Radiance at 1.67 pt

The aerosol attenuation coefficient is greater than theRayleigh attenuation at all altitudes included in thiscalculation according to the 1968 Elterman aerosoldistribution. There is no ozone absorption at thiswavelength. The ocean is a very strong absorber atthis wavelength. Approximately 500 photons arescattered by hydrosols for each photon scattered by aRayleigh center within the ocean. The influence ofabsorption within the ocean can be judged by the up-ward flux just below the ocean surface that is 1.12 X10-3, 2.95 X 10-5, and 2.44 X 10-7 at 0.7 a, 0.9 p, and1.67 pi, respectively. The ocean makes a greater con-tribution to the atmospheric radiance at 1.67 pi thanmight be assumed from these flux values, since an ap-preciable number of photons are reflected at the sur-face without entering the ocean.

The upward radiance at 1.67 pi is shown in Fig. 6.The radiance values are less here than at either of theother two wavelengths considered. Also the minimain the curves for models A and E are noticeably morepronounced. Smaller differences can also be noticed

- -

4:

0

C 0ZDl-

4xIO-

0.2 0.4 0.6 0.8 1.0

Fig. 6. Upward radiance as a function of p at X = 1.67 p formodels A-E.


l0-

w0z4:

4:

4:>0a

Il0z4:5r4:

4:

00

2x

5xlO-,3 L0.0

I I I I I:

_. leeL.r

-X- MODEL A-0- MODEL B-0- MODEL C-0- MODEL D-U- MODEL E

I I -

X-1.67,FLI . I . I . I

I -

-

In - . . . . . . . . . .

I

TV

3

2

w0gC]0

3:10'10a

2x10-2

0.0 0.2 0.4 0.6 0.8 1.0

Fig. 7. Downward radiance as a function of ,u at X = 1.67 A formodels A-E.

when the curves are compared with those at other wave-lengths.

The downward radiance is shown in Fig. 7. Theradiance values are less for models A and E than ateither of the other two wavelengths considered. On the

other hand, the changes in the curve for model C withwavelength are very small.

VI. FluxThe upward and downward flux was calculated for a

number of different detector locations in the atmosphereand ocean. The following quantities are tabulated inTable III: upward flux at the top of the atmosphere;upward and downward flux together with their ratiojust above the ocean surface; upward and downwardflux just below the ocean surface.

The upward flux at the top of the atmosphere in-creases as the aerosol amount increases (models A, B,C) at each wavelength. The upward flux is nearly thesame for models B and D with the same total aerosolamount. The upward flux is slightly greater for modelE than for model A because of a slightly greater totalaerosol amount.

The upward and downward flux just above the oceansurface is given in the fourth and fifth columns of TableIII. The upward flux just above the ocean surfaceconsists of a contribution from the downward radiationfrom both the sun and sky that is reflected at the oceansurface and from the upward radiation within the oceanthat comes back into the atmosphere. The lattercontribution is very small at 0.9 and 1.67 w. Thefraction of the downward radiation reflected at theocean surface varies with the aerosol amount, since thefraction of sky radiation compared to direct solar radia-tion increases as the aerosol amount increases. Theocean surface reflects a greater fraction of the sky radia-tion than of the direct solar radiation (assumed at thezenith), since the reflection coefficient of an air-watersurface increases as the incident angle increases awayfrom the normal. As an example, at 0.7 , the fractionof the downward flux just above the ocean surface re-

Table IlIl. Upward and Downward Fluxa

Ratio of upward DownwardUpward flux Upward flux Downward flux to downward Upward flux flux just

at top of just above just above flux just above just below below) (ii) Model atmosphere ocean surface ocean surface ocean surface ocean surface ocean surface

0.7 A 0.0650 0.0269 0.955 0.0281 1.12 x 10-3 0.9290.7 B 0.1168 0.0314 0.905 0.0347 1.23 X 10-s 0.8750.7 C 0.285 0.0346 0.740 0.0467 1.02 x 10-3 0.7060.7 D 0.1180 0.0313 0.905 0.0346 1.27 x 10-3 0.8740.7 E 0.0677 0.0269 0.948 0.0284 1.23 x 10-3 0.9230.7 F 0.1180 0.0313 0.908 0.0345 1.21 X 10-3 0.8780.7 G 0.1167 0.0313 0.903 0.0347 1.21 x 10-3 0.8720.9 A 0.0539 0.0255 0.973 0.0262 2.95 x 10-S 0.9460.9 B 0.1049 0.0304 0.929 0.0327 3.23 X 10-5 0.8980.9 C 0.266 0.0348 0.766 0.0454 1.95 X 10-s 0.7240.9 D 0.1054 0.0305 0.928 0.0329 3.68 x 10-1 0.8970.9 E 0.0581 0.0262 0.970 0.0270 3.83 X 10-1 0.9441.67 A 0.0408 0.0223 0.983 0.0227 2.4 x 10-7 0.9601.67 B 0.0831 0.0265 0.944 0.0281 2.5 x 10-7 0.9161.67 C 0.222 0.0328 0.816 0.0402 1.9 x 10-7 0.7781.67 D 0.0833 0.0266 0.944 0.0282 3.1 x 10-7 0.9171.67 E 0.0439 0.0227 0.980 0.0232 2.6 x 10-7 0.956

a All values of flux are normalized to unit incident solar flux.


I I I I I I ,_ -I

-X- MODEL A-0- MODEL B-9-MODEL C-D-MODEL D-U- MODEL E X-1.67/h

-1 1 I I I '" I .

y w . . - . . - B . B| B

. B . - .| |

I

flected by the surface is 2.7%, 3.3%, and 4.6% formodels A, B, and C, respectively; the fraction of skyto direct solar radiation is greater for model C thanfor model A because of the larger aerosol amount. Thedownward flux just above the ocean surface as shownin Table III decreases with increased aerosol amount asexpected.

The ratio of the upward to the downward flux justabove the ocean surface is tabulated in the sixth columnof Table III. This ratio is dependent on the detailsof the angular distribution of the downward radiation aspreviously discussed.

The upward and downward flux just below the oceansurface is tabulated in the seventh and eighth columnsof Table III. The upward flux depends greatly on theabsorption within the ocean; it is several orders ofmagnitude less at 0.9 u than at 0.7 ju and is down an-other two orders of magnitude at 1.67 u compared with0.9 ju. Because of the relatively small number ofcounts, these flux values are not as accurate as theothers in this table. The downward flux is some-what smaller in each case than the value just abovethe ocean surface because of losses from reflection at thesurface and small losses from absorption between thetwo detectors.

VI1. ConclusionThe upward flux at the top of the atmosphere as well

as the angular distribution of the radiation change ap-preciably as the aerosol amount increases from normalto ten times normal. At the same time the upward anddownward radiance just above the ocean surface under-goes important changes. The radiance does not changeappreciably with variations in the aerosol distribution

with height as long as the total aerosol amount remainsconstant. Similarly, changes in the ozone amountcause only small changes in the radiance at the wave-lengths considered here. Very little radiation returnsto the atmosphere from the ocean at 0.9 A and 1.67 Abecause of the high absorption of water at these wave-lengths.

These calculations were performed at the request ofGeneral Dynamics, Convair Aerospace Division, SanDiego, who specified the aerosol and ozone distribu-tions and provided support under contract NAS 1-10466.

References

1. T. H. Waterman and W. E. Westell, J. Marine Res. 15, 149(1956).

2. A. Ivanoff and T. H. Waterman, J. Marine Res. 16, 255(1958).

3. R. W. Preisendorfer, J. Marine Res. 18, 1 (1959).4. G. Neuman and R. Hollman, Int. Assoc. Phys. Oceanography

73, 72 (1960).5. S. Q. Duntley, J. Opt. Soc. Am. 53, 214 (1963).6. J. W. Hovenier, J. Atmos. Sci. 26, 488 (1969).7. D. Deirmendjian, Appl. Opt. 3, 187 (1964).8. G. W. Kattawar and G. N. Plass, Appl. Opt. 6, 1377 (1967).9. G. W. Kattawar and G. N. Plass, Appl. Opt. 6, 1549 (1967).

10. G. N. Plass and G. W. Kattawar, J. Atmos. Sci. 28, 1187(1971).

11. G. N. Plass and G. W. Kattawar, Appl. Opt. 8, 455 (1969).12. G. N. Plass and G. W. Kattawar, Appl. Opt. 9, 1122 (1970).13. L. Elterman, "UV Visible and IR attenuation for altitudes to

50 km, 1968," Report AFCRL-68-0153, Air Force CambridgeResearch Laboratories, Bedford, Mass. (1968).

14. G. N. Plass and G. W. Kattawar, Appl. Opt. 7, 415 (1968).15. G. N. Plass and G. W. Kattawar, Appl. Opt. 7, 699 (1968).

Electron Probe Analysis7th National Conference 17-21 July 1972

San Francisco

Technical sessions will include Electron Microprobe X-Ray Analysis, Techniques andInstrumentation. Principles of Electron Scattering and X-Ray Generation, QuantitativeCorrection Procedures, Soft X-Ray Emission, Computer Control and Data Reduction,Energy Dispersion Analysis, on Microprobe Analysis, Scanning Electron Microscopy,New Methods and Instrumentation in Microanalysis, and Applications. It is planned tooffer a tutorial session and/or manufacturers' forum. Anticipated attendance is esti-

mated at 400-500 and about 100 papers will be presented.

Abstracts must be submitted in English to Ted E. Lannin, General Electric Co., VallecitosRoad, Pleasanton, California 94566. Exhibitors should contact H. F. Harnsberger, ChevronResearch, Richmond, California 94802. The sponsor of the meeting is the Electron ProbeAnalysis Society of America. For general information contact David F. Kyser, Inter-

national Business Machines, Monterey & Cottle Roads, San Jose, California 94115.


effect of aerosol variation on radiance in the earth’s atmosphere-ocean system

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