INFLUENCE OF AEROSOL CLOUD HEIGHT ON THE CHANGE IN THE ATMOSPHERIC RADIATION BALANCE DUE TO AEROSOLS RUTH A. RECK Research
Laboratories,
General
Motors
Corporation,
Warren,
Michigan
48090, U.S.A.
(First receivrd 2 April 1974 and injifina/,form 22 July 1974)
Abstract-A radiative-convective atmospheric model for 35”N latitude in April is used to demonstrate the influence of aerosol cloud height on the change in the atmospheric radiation balance and temperature profile due to aerosols. Three different aerosol optical densities (0.065. 0.325 and 0.650) are considered as well as three mean aerosol cloud heights, corresponding to pressures of 958, 418 and 41.5 mb. In all cases considered. the surface temperature difference due to aerosols ranges between - 3 and + 0.4”K, with small magnitudes corresponding either to high aerosol clouds or to high surface albedo. The effect of surface albedo w, is included by considering w, values of 0.07, 0.3 and 0.6. Aerosols at all heights produce two opposing etfects in the atmosphere: (1) they reduce the magnitude of the total radiative cooling of the atmosphere, resulting in an effective “heating” at the earth’s surface, and (2) they increase the “reflectivity” of the earth-atmosphere system to solar radiation and thus reduce the amount of solar energy absorbed resulting in an effective “cooling” at the earth’s surface. Effect (1) apparently results from a “trapping” of the radiation by the aerosol and the absorption of this radiation by 03. HZ0 and CO,. A low-lying aerosol with an optical density ten times the present background value. when situated oker a surface such as water (0, = 0.07) causes a I .5 per cent decrease in the radiative cooling of the atmosphere relative to the present cooling rate. (The effect due to a low-lying aerosol having the background optical density T = 0.065 was found to be negligible.) Accompanying the’ I .5 per cent “heating” effect was a 3.8 per cent decrease in solar flux absorbed in the earth-atmosphere system. The combined heating and cooling produces a surface temperature difference of -3’K. For aerosols near the surface and over a reflecting surface such as snow (w, = 0.6) an aerosol optical density ten times the present background value produces a _ 0.7 per cent decrease in the present value of the radiative cooling of the atmosphere. Accompanying this “heating” effect is a negligible decrease in the solar radiation flux absorbed in the earth-atmosphere system. The combined effect in this case results in a surface temperature increase of 0.4-K.
INTRODUCTION
In a previous paper (Reck, 1974) a radiative-convective atmospheric model was used to demonstrate the role that the global surface shortwave albedo plays in determining the effect of a low-lying aerosol cloud upon the atmospheric radiation balance. In those calculations both heating and cooling was obtained near the surface of the earth, depending upon the magnitude of the surface albedo. In addition, the sign of the temperature effect at the earth’s surface was found to be independent of the optical density of the aerosol cloud. An analytic solution of the radiative-transfer equation in the two-stream approximation confirms both of these numerical results (Chjlek and Coakley, 1973). Calculations for increasing values of the aerosol optical density demonstrated a leveling of the surface temperature effect. In the present calculations the aerosol cloud type is unchanged from that previously considered, but we now use three aerosol attitudes (defined in terms of mean pressure) while holding the aerosol cloud optical density fixed. (Here fixed optical density is equivalent 89
to constant total number of particles in each of the throo clouds.) The thickness il % 01 each aerosol cloud is dotcrminrd in the model (Manabc and Stricklcr, 19h.T: Manabc and W~therald. 1967: Reck, 1974) bq the mean distance bctweun the two adjacent ~itrnos~~h~ri~ layers. Thus aerosol cloud thickness increases with height: this is physically ruasonablc since aerosol number density decreases \vith dccrcasing pressure (Hoffman ~1 trl.. 197.7). As a result, for fixed optical density 5 ( =I CTA %) the mean extinction coeficicnt c dccrcases with aerosol cloud height. The calculations were perfhrmed for acrosois havin g an opricai dcnsitv 01‘0.065. ;I ~aluc within the range quoted for the present global-uvcrage background f Porch 01 ~1.. 1~70) and also for optical densities 5 and IO times that value. At the earth’s surfi~e thcsc could corrcspond to a dust particle density of about 1SO. 1000 and 3000 ~(g m ‘, Only direct radiative effects are consider4 although the? map not bc the only ones 01‘ importance. For example. aerosols may also lead to seeding of water clouds (SMIC’. 197 I J with secondary changes in the global alhcdo as ~41. In addition to the i&t that the aerosol number density decrcabcs 1%ith altitude. the number density and spatial and silt distribution is highly dynamic and strongly dcpcndent on meteorological conditions. A complete model of the cfEct of aerosols on the atmospheric oncrgy balance would include all these factors. As ;I step toivard this complctcx model. the inlluencc of a given number of particles when placed at dif%zrent heights in the atmosphere is of interest. The purpose of the present work is to characterize the cncrgq e%ccts of ;I given aerosol cloud with respect to its mean height. Howcvcr. caution must be cxcrcised in comparing these calculations (for a single type ofaorosol cloud) with height effects in the real atmosphcrc, It is well known experimentally that the aerosols in the troposphero may bc of man! types (Drayson P! ul.. 1972) (e.g. sulfa&s. nitrates and ~i~~~ro~~~ri~(~~is~ each having diffcrcnt optical properties. while the stratospheric aerosols (Drayson (jr rii.. 1972) arit predominato14 sulf:~tes. The atmospheric model \ve have LISC~ in thcsc calculations has been previousI) dcscribed and the reader should refer to that original work (Reck. 1973) for details. As before WC have chosen to 11s~ values of the global parameters for 35 Y latitude in April \vith 3 ]a~cr_s &water clouds in a Manahe W~ther~~id-t~p~ (MW) (M~~~l~~beand Wetherald. 1967) atmospheric model. Again absolute humidit> is h&i iixcd and aerosols arc introduced in the two-stream approximation. In the next section various calculated yuantitics from the acro4~~l-t’rcc radiative convcctivc model at 35 ‘N latitude (April) will bc compared with csperimentally determined values for the average atmosphere as it exists toda,. This comparison is important to contit-m the validity of both the MW model and the parameter values used. and also to tiw thi: magnitudes of the global energy tlux quantities with present aerosols to that the relativc importance of chanpcs in them due to increasing amounts of aerosols can be determined.
The results of the MW model at 35 N iatitudc (April) arc compared with various available experimental values in Table I. Other extensive comparisons of the MW model fog other latitudes have also been made in the past (Manabc and Strickler. 1964; Manabe and Wet herald, 1967).
Influence Table
I. Comparison
of calculated
of aerosol
cloud height
and experimental thermal-radiative (no aerosols)
Nonradiative the global
flux from surface
r(mean earth-atmosphere short-wave albedo) T,(mean sea-level surface temperature) Solar radiation absorbed at the surface Solar radiation absorbed in the atmosphere
parameters
for 35” north latitude
(April)
Calculation (April 35” N lat) using M-W Atm. Model
Experimental r(mean surface shortwave albedo) Net solar flux at the top of the atmosphere
91
0. I (Sellers, 1965) 0.34 cal cm-’ min- ’ (Von der Haar and Suomi, 1971) 0.14 cal cm-* min-I (Budyko et al., 1962) (Sellers. 1965) 0.3 (Von der Haar and Suomi, 1971) 294.15”K (Valley, 1964) 0.25 cal cm-’ min- ’ (Reck, 1972) 0.09 cal cm-’ mini I (Sellers, 1965)
0.07 (assumed) 0.351 cal cm-’
mini’
0.138 cal cm-’
min-’
0.298
294.9”K 0.254 cal cm- * min0.097 cal cm-*
1
min-’
The annual value of surface albedo as given by Sellers (1965) for (3@40”) N latitude is T 0.1 and the value 0.07 in April (more nearly that for ice-free water) is assumed. The calculated value of solar flux at the top of the atmosphere is within 3 per cent of the value determined by satellite measurements (Von der Haar and Suomi, 1971). Of this total radiation flux density, 0.34 cal cmP2 min- I, the model predicts 0.254 cal cm- ’ min- 1 reaches the surface (compared with 0.25 cal cmm2 min- ’ determined from a vertical resolution of the experimental values (Reck, 1972)). Furthermore, 0.097 cal cme2 min- 1 of the solar radiation is absorbed in the atmosphere compared with 0.09 cal cm-’ min- ’ from the resolution of experimental quantities (Sellers, 1965; Reck, 1972). The calculated asymptotic mean sea-level steady-state temperature near the surface is 294.9” K established after _ 300,8-h iterations using the MW model. (Note that the number of iterations is strongly a function of the initial temperatures chosen; 300 iterations are required for an initial isothermal 300°K atmosphere.) This is to be compared with the value 294.15”K which is the mean of the values for January (287.15”K) and July (301.15”K) of the standard subtropical atmosphere at 30”N latitude at 1000 mb pressure (Valley, 1965). The convective correction introduces a nonradiative flux of 0.138 cal cmb2 mini ’ from the surface to the atmosphere compared with estimates by Budyko et al. (1962) and Sellers (1965)of -0.14calcm-2min-1. The calculated mean shortwave earth-atmosphere albedo at 35”N latitude (April) from this model is 0.298 compared with a satellite value (Van der Haar and Suomi, 1971) of 0.3. The fact that these derived quantities are all within the uncertainty limits of the experimentally determined values justifies the parameters chosen for 35”N latitude in April. DESCRIPTION
OF
THF
PRESENT
CALCULATIONS
asymptotic steady-state values of surface temperature, radiative cooling, and earthatmospheric albedos for each of three values of the aerosol optical density, 0.065 (present The
background ~due), N.335 and 0.650. for mean aerosol cloud positions in the vertical of 95s. 41 X and 41.5 mb have been calculated. The lowest position (95X mb) corresponds to the previously studied (Reek, 1’974) low-lying aerosol cloud. The intermediate position (41 X mb) is between the uppermost water cloud layer situated at 336 mb and two loucr-lying water cloud layers at X81 and 935 mb. Values of the water cloud optical paramctcrs arc given in Table 2; values of the aerosol hackscattcr and absorption in the visible. and transmission and backscatter in the infrared arc given in Table 3. These nine calculations have been repeated for cuch of three valut~ of the surface albcdo. 0.07. 0.30 and 0.60. In these calculations the criterion for reaching a steady-state tcmperature profile was that the tcmperaturc at every level should not change by more than 0.003 0.006 K da).- ’ In some cases for the high level aerosol cloud the 8-h integration time-step was shortened due to instability problems. This difficulty has been discussed earlier
Influence
of aerosol
93
cloud height
(Manabe and Strickler, 1964). The results of these 27 different calculations will now be presented. AEROSOL COOLING,
EFFECTS ON TEMPERATURE, AND MEAN GARTH-ATMOSPHERE
RADIATIVE ALBEDG
As shown in Fig. 1, for an aerosol cloud over a surface having an albedo of 0:07; a decrease in surface temperature (i.e. a negative temperature difference) was obtained for all heights and optical densities of the aerosol cloud. For each cloud height the surface temperature difference increased as the optical density increased. This temperature effect, however, decreased as the height of the aerosol cloud increased. Aerosol clouds at the middle and high altitudes having an optical density equal to the present background value (r = 0.065) show very little difference in their effects on the surface temperature (both AC < /0.2j”K). However, for larger T-values this is no longer true. In Fig. 2 the corresponding differences in the net cooling of the atmosphere are shown. The magnitudes of these quantities (- 10s4 cal cm- ’ min- ‘) are small compared to the present value of 0.14cal cmW2 mini 1 for the nonradiative flux shown in Table 1. The largest change in cooling rate shown in Fig. 2 (for a lowlying aerosol cloud with an optical density 10 times the present background level) would lead to a decrease of only 1.5 per cent in the magnitude of the radiative cooling of the atmosphere. The mean calculated earth-atmosphere albedos for the various cloud heights and optical densities are shown in Fig. 3. The dotted line corresponds to the albedo without aerosols. The maximum increase of the albedo from the present vaIue is - 0.026 or IO per cent. This represents a decrease in the amount of solar energy absorbed by the earth-atmosphere system of - 0.013 cal cm-’ min- ‘, or 3.8 per cent of the amount absorbed without the aerosols. IO
ld
L
-4.0
-3.0
-2.0
-1.0
0
1.0
ATsLK) Fig. 1. Surface temperature differences (^K) (temperature with aerosols minus temperature without aerosols) for various man aerosol cloud heights (mb) over a surface having a mean shortwave albedo, w,, of 0.07. Each curve is for a different aerosol optical density. Note: These results show surfice temperature differences for nine separate calculations and should not be confused with the usual atmospheric vertical temperature profiles.
w Coohng
(Effective
decreastng
_
*heating”)
Figure 4 shows the difference in surface temperature for an aerosol cloud over a surf&e having an albedo of 0.3. These curves have roughly the same shape as those in Fig. 1 cxoept that all are shifted toward more positive values of the abscissa. In fact the highest aerosol cloud produces a positive AT,, since the magnitude of the “heating” effect has become greater than the magnitude of the decrease in absorbed solar radiation tlux. In Fig. 5 the change in net atmospheric cooling for ~1)~= 0.3 is similar to that in Fig. 2 for CO,? = 0.07. Again there is a shift to so~ne~ll~lt smaller- values of the abscissa. The mean earth-atmosphere albedos for (r), = 0.3 ;trc shown in Fig. 6. Again. the dotted line refers to the no-aerosol value. Here the maximum range ofalbedo diffcrcnce is .. 0.013 at the surface. For the high and middle aerosoi cloud the alhedo increabcs :15 T inu-c:ux \vfliic fix ;I low-lying cloud the greatest effect in Fig. 6 is for z = 0.065. Lvith a decreasing effect as T increases. It is important to note that in previous calculations I Reck, 1974) a reversal 01 the trend in the A?; behavior was observed at W, : 0.3 when the cloud optical densit) exceeded 3.0. In Fig. 6. an additional point(r) for z = 0.02 confirms the fact that the alhedo change approaches zero as ?--+O. In an effort to understand the reversal in behavior shown in Fig. 6. it is neccssar! to realize that when (0, = 0.07 most of the solar radiation getting through the lowest water cloud is absorbed. In this case the atmosphere -earth albedo is mainly determined by water
Influence
of aerosol
cloud height
WithoutAerosols
I o* = 0.07
II I I I I I I I I
I a (Mean Albedol
Fig. 3. Mean shortwave earth-atmosphere albedo given for various mean aerosol cloud heights (mb) over a surface having a mean shortwave albedo of 0.07. Each curve is for a different aerosol optical density. The dashed line refers to the mean earth-atmosphere albedo without aerosols. Note: These results show mean shortwave earth-atmosphere albedos for nine separate calculations and should not be confused with the usual atmospheric vertical profiles.
as = 0.30
-2.0
-1.0
0
1.0
2.0
AT, (K) Fig. 4. Surface temperature differences (‘K) (temperature with aerosols minus temperature without aerosols) for various mean aerosol cloud heights (mb) over a surface having a mean shortwave albedo, CO,,of 0.3. Each curve is for a diRerent aerosol optical density. Note: These results show surjuc~ temperature differences for nine separate calculations and should not be confused with the usual atmospheric vertical temperature profiles.
Cooli:lg
decreuslng
(Effective
‘heating')
I
I Fig. 5. Differences in RC. the integrated radiative cooling rate of the atmosphere in cal cm ’ min ’ (RC with aerosols minus RC Gthout aerosols). for various mean aerosol cloud heights (mh) over a surface having a mean shortwave albedo. CO,.of O.?. Each curve is for a different aerosol optical donsit>. .\‘,)rr: These results show integrated radiative ditkences for nine separate calculations and should not bc confused with the usual atmosphertc vertical profiles. 1
I
; WithoutAerosols
0.41
0.43
0.42 MEAN
0.44
ALBEDO
Fig. 6. Mean shortw-avc earth atmosphcrc alhedo ptvcn for various mean aerosol cloud heights (mh) over a surface having a mean shortwavc albedo of Cf.?. Each curve is for a different optical density, The dashed line refers to the mean earth-atmosphere alhedo without aerosols. The addttional point (s) for T = 0.02 confirms the fact that the al&do change approaches zero as r--d 0. .Yote: These results show mean shortwave earth-atmosphere albedos for ten separate calculations and should not be confused with the usual atmospheric vertical profiles.
Influence
of aerosol
cloud height
97
cloud backscatter. However when o, = 0.30 a larger amount of the albedo (0.418-0.298) is caused by the surface reflection. In this latter case a low-lying aerosol cloud of small optical density (T = 0.02) increased the atmosphere-earth albedo. But as the optical density of the aerosol cloud increases it not only increases the backscatter of solar radiation upwards, it also scatters back toward the earth the solar radiation reflected from the earth’s surface. These two competing effects cause the mean albedo to be less for an aerosol cloud with r = 0.02 and 0.325 than for one with r = 0.065. When the aerosol cloud is above the lowest water cloud the reflected solar radiation is reduced in amount by the water cloud and the effect does not occur for the parameters used. As shown in Fig. 7, for an aerosol cloud over a surface having an albedo of 0.6 the surface temperature effect is small, and the behavior is nonuniform. For r = 0.065 there is a small (0.02” K) heating effect at the surface which increases linearly with cloud height. For z = 0.325, AT, t 0 for the low or middle height clouds whereas AT, 5 0.4’ K for the high aerosol cloud. For T = 0.65, both the low and high-lying aerosols give positive AT, while the middle aerosol cloud shows a slight cooling effect. The change in the integrated radiative cooling of the atmosphere for o, = 0.6 is shown in Fig. 8. Here there is very little effect for r = 0.065, and essentially no effect for all values of T for the high aerosol cloud. The mean albedo for o, = 0.6 is 0.574 and shows no significant change with the addition of an aerosol cloud. Therefore no curve is shown.
aIs =
0.60
0
1.0
Fig. 7. Surface temperature differences (“K) (temperature with aerosols minus temperature without aerosols) for various mean aerosol cloud heights (mb) over a surface having a mean shortwave albedo, o, of 0.6. Each curve is for a different optical density. A’ote: These results show surface temperature differences for nine separate calculations and should not be confused with the usual atmospheric vertical temperature profiles.
L
c
p
i
c
Influence
of aerosol
99
cloud height
The results presented in this work are limited by the radiative-convective model used in that they do not include any of the additional dynamic or hydrologic effects that would be important in a “real” atmosphere. Ack,low/edgrmr/lts-The author wishes to express gratitude to S. Manabe and R. Wetherald Fluid Dynamics Laboratory of Princeton University for a copy of their radiative-convective
at the Geophysical computer code.
REFERENCES Budyko M. I., Yefimova N. A., Zubenok L. I. and Strokin L. A. (1962) Sov. Geogruk 3, 3. Chjilek P. and Coakley J. A. Jr. (1973) Aerosols und Climate. NCAR, June. Drayson S. R., Bartman F. L., Kuhn W. R. and Tallamraju R. (1972) Satellite measurement of stratospheric pollutants and minor constituents by solar occulation: A preliminary report. NOAA Grant NG-10-72. Hofmdnn D. J., Rosen J. M. and Pepin T. J. (1973) Satellite measurement of stratospheric pollutants and minor constituents by solar occulation: A preliminary report. CIAP. ONR No. NOOO14-70-A-0266-0005, Report GM-9. May. London J. and Sasamori T. (1971) Space Resvurch XI. Akadenusche. Berlin. Manabe S. and Strickler R. F. (1964) J. Acrnos. Sci. 21, 361. Manabe S. and Wetherald R. T. (1967) J. Atmos. Sci. 24, 241. Porch W. M., Charlson R. J. and Radke L. F. (1970) Sciericr 170, 315. Reck R. A. (1972) Vertical resolution of the earth-atmosphere annual energy balance. GMR-1313, General Motors Corp. Reck R. A. (1974) Influence of surface albedo on the change in the atmospheric radiation balance due to aerosols. Atmospheric Environment 8, 3233333. Sellers W. D. (1965) Physical Climatology. The University of Chicago Press. Chicago. SMIC (Study of Man’s Impact on Climate) (1971) Inadwrterlr CIimatr Modification, p. 188. M.I.T. Press, Cambridge. Valley S. L. (Editor) (1965) Handhook qfGrophy.sics and Space Ewiromwrlts. McGraw-Hill, New York. Von der Haar T. H. and Suomi V. E. (1971) J. Atrnos. Sci. 28, 306.
Laboratories,
General
Motors
Corporation,
Warren,
Michigan
48090, U.S.A.
(First receivrd 2 April 1974 and injifina/,form 22 July 1974)
Abstract-A radiative-convective atmospheric model for 35”N latitude in April is used to demonstrate the influence of aerosol cloud height on the change in the atmospheric radiation balance and temperature profile due to aerosols. Three different aerosol optical densities (0.065. 0.325 and 0.650) are considered as well as three mean aerosol cloud heights, corresponding to pressures of 958, 418 and 41.5 mb. In all cases considered. the surface temperature difference due to aerosols ranges between - 3 and + 0.4”K, with small magnitudes corresponding either to high aerosol clouds or to high surface albedo. The effect of surface albedo w, is included by considering w, values of 0.07, 0.3 and 0.6. Aerosols at all heights produce two opposing etfects in the atmosphere: (1) they reduce the magnitude of the total radiative cooling of the atmosphere, resulting in an effective “heating” at the earth’s surface, and (2) they increase the “reflectivity” of the earth-atmosphere system to solar radiation and thus reduce the amount of solar energy absorbed resulting in an effective “cooling” at the earth’s surface. Effect (1) apparently results from a “trapping” of the radiation by the aerosol and the absorption of this radiation by 03. HZ0 and CO,. A low-lying aerosol with an optical density ten times the present background value. when situated oker a surface such as water (0, = 0.07) causes a I .5 per cent decrease in the radiative cooling of the atmosphere relative to the present cooling rate. (The effect due to a low-lying aerosol having the background optical density T = 0.065 was found to be negligible.) Accompanying the’ I .5 per cent “heating” effect was a 3.8 per cent decrease in solar flux absorbed in the earth-atmosphere system. The combined heating and cooling produces a surface temperature difference of -3’K. For aerosols near the surface and over a reflecting surface such as snow (w, = 0.6) an aerosol optical density ten times the present background value produces a _ 0.7 per cent decrease in the present value of the radiative cooling of the atmosphere. Accompanying this “heating” effect is a negligible decrease in the solar radiation flux absorbed in the earth-atmosphere system. The combined effect in this case results in a surface temperature increase of 0.4-K.
INTRODUCTION
In a previous paper (Reck, 1974) a radiative-convective atmospheric model was used to demonstrate the role that the global surface shortwave albedo plays in determining the effect of a low-lying aerosol cloud upon the atmospheric radiation balance. In those calculations both heating and cooling was obtained near the surface of the earth, depending upon the magnitude of the surface albedo. In addition, the sign of the temperature effect at the earth’s surface was found to be independent of the optical density of the aerosol cloud. An analytic solution of the radiative-transfer equation in the two-stream approximation confirms both of these numerical results (Chjlek and Coakley, 1973). Calculations for increasing values of the aerosol optical density demonstrated a leveling of the surface temperature effect. In the present calculations the aerosol cloud type is unchanged from that previously considered, but we now use three aerosol attitudes (defined in terms of mean pressure) while holding the aerosol cloud optical density fixed. (Here fixed optical density is equivalent 89
to constant total number of particles in each of the throo clouds.) The thickness il % 01 each aerosol cloud is dotcrminrd in the model (Manabc and Stricklcr, 19h.T: Manabc and W~therald. 1967: Reck, 1974) bq the mean distance bctweun the two adjacent ~itrnos~~h~ri~ layers. Thus aerosol cloud thickness increases with height: this is physically ruasonablc since aerosol number density decreases \vith dccrcasing pressure (Hoffman ~1 trl.. 197.7). As a result, for fixed optical density 5 ( =I CTA %) the mean extinction coeficicnt c dccrcases with aerosol cloud height. The calculations were perfhrmed for acrosois havin g an opricai dcnsitv 01‘0.065. ;I ~aluc within the range quoted for the present global-uvcrage background f Porch 01 ~1.. 1~70) and also for optical densities 5 and IO times that value. At the earth’s surfi~e thcsc could corrcspond to a dust particle density of about 1SO. 1000 and 3000 ~(g m ‘, Only direct radiative effects are consider4 although the? map not bc the only ones 01‘ importance. For example. aerosols may also lead to seeding of water clouds (SMIC’. 197 I J with secondary changes in the global alhcdo as ~41. In addition to the i&t that the aerosol number density decrcabcs 1%ith altitude. the number density and spatial and silt distribution is highly dynamic and strongly dcpcndent on meteorological conditions. A complete model of the cfEct of aerosols on the atmospheric oncrgy balance would include all these factors. As ;I step toivard this complctcx model. the inlluencc of a given number of particles when placed at dif%zrent heights in the atmosphere is of interest. The purpose of the present work is to characterize the cncrgq e%ccts of ;I given aerosol cloud with respect to its mean height. Howcvcr. caution must be cxcrcised in comparing these calculations (for a single type ofaorosol cloud) with height effects in the real atmosphcrc, It is well known experimentally that the aerosols in the troposphero may bc of man! types (Drayson P! ul.. 1972) (e.g. sulfa&s. nitrates and ~i~~~ro~~~ri~(~~is~ each having diffcrcnt optical properties. while the stratospheric aerosols (Drayson (jr rii.. 1972) arit predominato14 sulf:~tes. The atmospheric model \ve have LISC~ in thcsc calculations has been previousI) dcscribed and the reader should refer to that original work (Reck. 1973) for details. As before WC have chosen to 11s~ values of the global parameters for 35 Y latitude in April \vith 3 ]a~cr_s &water clouds in a Manahe W~ther~~id-t~p~ (MW) (M~~~l~~beand Wetherald. 1967) atmospheric model. Again absolute humidit> is h&i iixcd and aerosols arc introduced in the two-stream approximation. In the next section various calculated yuantitics from the acro4~~l-t’rcc radiative convcctivc model at 35 ‘N latitude (April) will bc compared with csperimentally determined values for the average atmosphere as it exists toda,. This comparison is important to contit-m the validity of both the MW model and the parameter values used. and also to tiw thi: magnitudes of the global energy tlux quantities with present aerosols to that the relativc importance of chanpcs in them due to increasing amounts of aerosols can be determined.
The results of the MW model at 35 N iatitudc (April) arc compared with various available experimental values in Table I. Other extensive comparisons of the MW model fog other latitudes have also been made in the past (Manabc and Strickler. 1964; Manabe and Wet herald, 1967).
Influence Table
I. Comparison
of calculated
of aerosol
cloud height
and experimental thermal-radiative (no aerosols)
Nonradiative the global
flux from surface
r(mean earth-atmosphere short-wave albedo) T,(mean sea-level surface temperature) Solar radiation absorbed at the surface Solar radiation absorbed in the atmosphere
parameters
for 35” north latitude
(April)
Calculation (April 35” N lat) using M-W Atm. Model
Experimental r(mean surface shortwave albedo) Net solar flux at the top of the atmosphere
91
0. I (Sellers, 1965) 0.34 cal cm-’ min- ’ (Von der Haar and Suomi, 1971) 0.14 cal cm-* min-I (Budyko et al., 1962) (Sellers. 1965) 0.3 (Von der Haar and Suomi, 1971) 294.15”K (Valley, 1964) 0.25 cal cm-’ min- ’ (Reck, 1972) 0.09 cal cm-’ mini I (Sellers, 1965)
0.07 (assumed) 0.351 cal cm-’
mini’
0.138 cal cm-’
min-’
0.298
294.9”K 0.254 cal cm- * min0.097 cal cm-*
1
min-’
The annual value of surface albedo as given by Sellers (1965) for (3@40”) N latitude is T 0.1 and the value 0.07 in April (more nearly that for ice-free water) is assumed. The calculated value of solar flux at the top of the atmosphere is within 3 per cent of the value determined by satellite measurements (Von der Haar and Suomi, 1971). Of this total radiation flux density, 0.34 cal cmP2 min- I, the model predicts 0.254 cal cm- ’ min- 1 reaches the surface (compared with 0.25 cal cmm2 min- ’ determined from a vertical resolution of the experimental values (Reck, 1972)). Furthermore, 0.097 cal cme2 min- 1 of the solar radiation is absorbed in the atmosphere compared with 0.09 cal cm-’ min- ’ from the resolution of experimental quantities (Sellers, 1965; Reck, 1972). The calculated asymptotic mean sea-level steady-state temperature near the surface is 294.9” K established after _ 300,8-h iterations using the MW model. (Note that the number of iterations is strongly a function of the initial temperatures chosen; 300 iterations are required for an initial isothermal 300°K atmosphere.) This is to be compared with the value 294.15”K which is the mean of the values for January (287.15”K) and July (301.15”K) of the standard subtropical atmosphere at 30”N latitude at 1000 mb pressure (Valley, 1965). The convective correction introduces a nonradiative flux of 0.138 cal cmb2 mini ’ from the surface to the atmosphere compared with estimates by Budyko et al. (1962) and Sellers (1965)of -0.14calcm-2min-1. The calculated mean shortwave earth-atmosphere albedo at 35”N latitude (April) from this model is 0.298 compared with a satellite value (Van der Haar and Suomi, 1971) of 0.3. The fact that these derived quantities are all within the uncertainty limits of the experimentally determined values justifies the parameters chosen for 35”N latitude in April. DESCRIPTION
OF
THF
PRESENT
CALCULATIONS
asymptotic steady-state values of surface temperature, radiative cooling, and earthatmospheric albedos for each of three values of the aerosol optical density, 0.065 (present The
background ~due), N.335 and 0.650. for mean aerosol cloud positions in the vertical of 95s. 41 X and 41.5 mb have been calculated. The lowest position (95X mb) corresponds to the previously studied (Reek, 1’974) low-lying aerosol cloud. The intermediate position (41 X mb) is between the uppermost water cloud layer situated at 336 mb and two loucr-lying water cloud layers at X81 and 935 mb. Values of the water cloud optical paramctcrs arc given in Table 2; values of the aerosol hackscattcr and absorption in the visible. and transmission and backscatter in the infrared arc given in Table 3. These nine calculations have been repeated for cuch of three valut~ of the surface albcdo. 0.07. 0.30 and 0.60. In these calculations the criterion for reaching a steady-state tcmperature profile was that the tcmperaturc at every level should not change by more than 0.003 0.006 K da).- ’ In some cases for the high level aerosol cloud the 8-h integration time-step was shortened due to instability problems. This difficulty has been discussed earlier
Influence
of aerosol
93
cloud height
(Manabe and Strickler, 1964). The results of these 27 different calculations will now be presented. AEROSOL COOLING,
EFFECTS ON TEMPERATURE, AND MEAN GARTH-ATMOSPHERE
RADIATIVE ALBEDG
As shown in Fig. 1, for an aerosol cloud over a surface having an albedo of 0:07; a decrease in surface temperature (i.e. a negative temperature difference) was obtained for all heights and optical densities of the aerosol cloud. For each cloud height the surface temperature difference increased as the optical density increased. This temperature effect, however, decreased as the height of the aerosol cloud increased. Aerosol clouds at the middle and high altitudes having an optical density equal to the present background value (r = 0.065) show very little difference in their effects on the surface temperature (both AC < /0.2j”K). However, for larger T-values this is no longer true. In Fig. 2 the corresponding differences in the net cooling of the atmosphere are shown. The magnitudes of these quantities (- 10s4 cal cm- ’ min- ‘) are small compared to the present value of 0.14cal cmW2 mini 1 for the nonradiative flux shown in Table 1. The largest change in cooling rate shown in Fig. 2 (for a lowlying aerosol cloud with an optical density 10 times the present background level) would lead to a decrease of only 1.5 per cent in the magnitude of the radiative cooling of the atmosphere. The mean calculated earth-atmosphere albedos for the various cloud heights and optical densities are shown in Fig. 3. The dotted line corresponds to the albedo without aerosols. The maximum increase of the albedo from the present vaIue is - 0.026 or IO per cent. This represents a decrease in the amount of solar energy absorbed by the earth-atmosphere system of - 0.013 cal cm-’ min- ‘, or 3.8 per cent of the amount absorbed without the aerosols. IO
ld
L
-4.0
-3.0
-2.0
-1.0
0
1.0
ATsLK) Fig. 1. Surface temperature differences (^K) (temperature with aerosols minus temperature without aerosols) for various man aerosol cloud heights (mb) over a surface having a mean shortwave albedo, w,, of 0.07. Each curve is for a different aerosol optical density. Note: These results show surfice temperature differences for nine separate calculations and should not be confused with the usual atmospheric vertical temperature profiles.
w Coohng
(Effective
decreastng
_
*heating”)
Figure 4 shows the difference in surface temperature for an aerosol cloud over a surf&e having an albedo of 0.3. These curves have roughly the same shape as those in Fig. 1 cxoept that all are shifted toward more positive values of the abscissa. In fact the highest aerosol cloud produces a positive AT,, since the magnitude of the “heating” effect has become greater than the magnitude of the decrease in absorbed solar radiation tlux. In Fig. 5 the change in net atmospheric cooling for ~1)~= 0.3 is similar to that in Fig. 2 for CO,? = 0.07. Again there is a shift to so~ne~ll~lt smaller- values of the abscissa. The mean earth-atmosphere albedos for (r), = 0.3 ;trc shown in Fig. 6. Again. the dotted line refers to the no-aerosol value. Here the maximum range ofalbedo diffcrcnce is .. 0.013 at the surface. For the high and middle aerosoi cloud the alhedo increabcs :15 T inu-c:ux \vfliic fix ;I low-lying cloud the greatest effect in Fig. 6 is for z = 0.065. Lvith a decreasing effect as T increases. It is important to note that in previous calculations I Reck, 1974) a reversal 01 the trend in the A?; behavior was observed at W, : 0.3 when the cloud optical densit) exceeded 3.0. In Fig. 6. an additional point(r) for z = 0.02 confirms the fact that the alhedo change approaches zero as ?--+O. In an effort to understand the reversal in behavior shown in Fig. 6. it is neccssar! to realize that when (0, = 0.07 most of the solar radiation getting through the lowest water cloud is absorbed. In this case the atmosphere -earth albedo is mainly determined by water
Influence
of aerosol
cloud height
WithoutAerosols
I o* = 0.07
II I I I I I I I I
I a (Mean Albedol
Fig. 3. Mean shortwave earth-atmosphere albedo given for various mean aerosol cloud heights (mb) over a surface having a mean shortwave albedo of 0.07. Each curve is for a different aerosol optical density. The dashed line refers to the mean earth-atmosphere albedo without aerosols. Note: These results show mean shortwave earth-atmosphere albedos for nine separate calculations and should not be confused with the usual atmospheric vertical profiles.
as = 0.30
-2.0
-1.0
0
1.0
2.0
AT, (K) Fig. 4. Surface temperature differences (‘K) (temperature with aerosols minus temperature without aerosols) for various mean aerosol cloud heights (mb) over a surface having a mean shortwave albedo, CO,,of 0.3. Each curve is for a diRerent aerosol optical density. Note: These results show surjuc~ temperature differences for nine separate calculations and should not be confused with the usual atmospheric vertical temperature profiles.
Cooli:lg
decreuslng
(Effective
‘heating')
I
I Fig. 5. Differences in RC. the integrated radiative cooling rate of the atmosphere in cal cm ’ min ’ (RC with aerosols minus RC Gthout aerosols). for various mean aerosol cloud heights (mh) over a surface having a mean shortwave albedo. CO,.of O.?. Each curve is for a different aerosol optical donsit>. .\‘,)rr: These results show integrated radiative ditkences for nine separate calculations and should not bc confused with the usual atmosphertc vertical profiles. 1
I
; WithoutAerosols
0.41
0.43
0.42 MEAN
0.44
ALBEDO
Fig. 6. Mean shortw-avc earth atmosphcrc alhedo ptvcn for various mean aerosol cloud heights (mh) over a surface having a mean shortwavc albedo of Cf.?. Each curve is for a different optical density, The dashed line refers to the mean earth-atmosphere alhedo without aerosols. The addttional point (s) for T = 0.02 confirms the fact that the al&do change approaches zero as r--d 0. .Yote: These results show mean shortwave earth-atmosphere albedos for ten separate calculations and should not be confused with the usual atmospheric vertical profiles.
Influence
of aerosol
cloud height
97
cloud backscatter. However when o, = 0.30 a larger amount of the albedo (0.418-0.298) is caused by the surface reflection. In this latter case a low-lying aerosol cloud of small optical density (T = 0.02) increased the atmosphere-earth albedo. But as the optical density of the aerosol cloud increases it not only increases the backscatter of solar radiation upwards, it also scatters back toward the earth the solar radiation reflected from the earth’s surface. These two competing effects cause the mean albedo to be less for an aerosol cloud with r = 0.02 and 0.325 than for one with r = 0.065. When the aerosol cloud is above the lowest water cloud the reflected solar radiation is reduced in amount by the water cloud and the effect does not occur for the parameters used. As shown in Fig. 7, for an aerosol cloud over a surface having an albedo of 0.6 the surface temperature effect is small, and the behavior is nonuniform. For r = 0.065 there is a small (0.02” K) heating effect at the surface which increases linearly with cloud height. For z = 0.325, AT, t 0 for the low or middle height clouds whereas AT, 5 0.4’ K for the high aerosol cloud. For T = 0.65, both the low and high-lying aerosols give positive AT, while the middle aerosol cloud shows a slight cooling effect. The change in the integrated radiative cooling of the atmosphere for o, = 0.6 is shown in Fig. 8. Here there is very little effect for r = 0.065, and essentially no effect for all values of T for the high aerosol cloud. The mean albedo for o, = 0.6 is 0.574 and shows no significant change with the addition of an aerosol cloud. Therefore no curve is shown.
aIs =
0.60
0
1.0
Fig. 7. Surface temperature differences (“K) (temperature with aerosols minus temperature without aerosols) for various mean aerosol cloud heights (mb) over a surface having a mean shortwave albedo, o, of 0.6. Each curve is for a different optical density. A’ote: These results show surface temperature differences for nine separate calculations and should not be confused with the usual atmospheric vertical temperature profiles.
L
c
p
i
c
Influence
of aerosol
99
cloud height
The results presented in this work are limited by the radiative-convective model used in that they do not include any of the additional dynamic or hydrologic effects that would be important in a “real” atmosphere. Ack,low/edgrmr/lts-The author wishes to express gratitude to S. Manabe and R. Wetherald Fluid Dynamics Laboratory of Princeton University for a copy of their radiative-convective
at the Geophysical computer code.
REFERENCES Budyko M. I., Yefimova N. A., Zubenok L. I. and Strokin L. A. (1962) Sov. Geogruk 3, 3. Chjilek P. and Coakley J. A. Jr. (1973) Aerosols und Climate. NCAR, June. Drayson S. R., Bartman F. L., Kuhn W. R. and Tallamraju R. (1972) Satellite measurement of stratospheric pollutants and minor constituents by solar occulation: A preliminary report. NOAA Grant NG-10-72. Hofmdnn D. J., Rosen J. M. and Pepin T. J. (1973) Satellite measurement of stratospheric pollutants and minor constituents by solar occulation: A preliminary report. CIAP. ONR No. NOOO14-70-A-0266-0005, Report GM-9. May. London J. and Sasamori T. (1971) Space Resvurch XI. Akadenusche. Berlin. Manabe S. and Strickler R. F. (1964) J. Acrnos. Sci. 21, 361. Manabe S. and Wetherald R. T. (1967) J. Atmos. Sci. 24, 241. Porch W. M., Charlson R. J. and Radke L. F. (1970) Sciericr 170, 315. Reck R. A. (1972) Vertical resolution of the earth-atmosphere annual energy balance. GMR-1313, General Motors Corp. Reck R. A. (1974) Influence of surface albedo on the change in the atmospheric radiation balance due to aerosols. Atmospheric Environment 8, 3233333. Sellers W. D. (1965) Physical Climatology. The University of Chicago Press. Chicago. SMIC (Study of Man’s Impact on Climate) (1971) Inadwrterlr CIimatr Modification, p. 188. M.I.T. Press, Cambridge. Valley S. L. (Editor) (1965) Handhook qfGrophy.sics and Space Ewiromwrlts. McGraw-Hill, New York. Von der Haar T. H. and Suomi V. E. (1971) J. Atrnos. Sci. 28, 306.
