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How Changes in Atmospheric Properties Impact Radiative Transfer

Tracey Dorian

ABSTRACT

In a series of five different experiments, specific changes are made to the atmosphere that results in

changes in radiative flux. The five experiments include changes to baseline cloud-free tropical and

polar atmospheres, and we analyze the resulting differences in broadband fluxes, spectral fluxes, and

heating profiles. We focus on fluxes at the surface and at the top of the atmosphere, and we look at

shortwave, longwave, and shortwave and longwave combined. The experiments include increasing

water vapor amount, increasing carbon dioxide amount, decreasing ozone concentration, and adding

low clouds and high clouds to the system. After the appropriate changes to the baseline atmospheres

were made, a radiative transfer model was used to compute the associated flux differences and

heating differences which were then plotted and analyzed. The main findings of this research were

that changes in cloud coverage and cloud height seem to have a much larger impact on changes in

radiative fluxes and heating profiles than do changes in greenhouse gas concentrations. We also

find that low clouds tend to lower surface temperatures and high clouds tend to increase surface

temperatures. We also find that greenhouse gases such as water vapor, CO2, and O3 that absorb

certain wavelengths in the longwave spectrum tend to help trap heat in the atmosphere.

1. Introduction

Earths radiative energy balance is one of the most im-portant factors in understanding the fate of our climate.What determines the climate in any given location is thedifference between incoming and outgoing radiation. Thisdifference then determines whether heating or cooling oc-curs and therefore also determines if the environment is insteady state. The Earth receives solar radiation from the

Sun which then heats the surface. Thermal radiation emit-ted from the surface is then what is responsible for coolingthe Earth. Certain gases, aerosols, and clouds have an ef-fect on the amount of radiation that reaches the surfaceand the amount of radiation that eventually leaves Earthsatmosphere. Thermal infrared radiation is between 4 µmand 50 µm, and is absorbed and emitted by water vapor,carbon dioxide, ozone, and other trace gases (Petty (2006)page 58). A cloud-free atmosphere is transparent to visibleradiation, while clouds reflect visible radiation.

In this investigation, we take two baseline cases of atropical atmosphere and a polar atmosphere. For bothbaseline cases, we find the shortwave, longwave, and com-

bined upward, downward, and net broadband fluxes forboth the surface and the top of the atmosphere at 100km. Our next task was to plot upward, downward, andnet spectral fluxes for the top of the atmosphere and forthe surface. Finally, we plotted heating profiles for eachmodel atmosphere for shortwave, longwave, and combinedwavelengths. Shortwave wavelengths were designated asbands 106-129, while longwave wavelengths were desig-

nated as bands 1-105. Combined wavelengths included all129 bands. We also found the change in radiative fluxes dueto changes in the atmosphere, a process known as radia-tive forcing. For both the tropical and polar atmospheres,we increased water vapor amount, we doubled the CO2

concentration, we decreased the amount of ozone, and weadded a low liquid cloud layer and a high ice cloud layerboth with the same optical thicknesses. So our final resultsincluded an analysis of the associated changes relative tothe baseline cases in the fluxes and heating profiles.

2. Methods

In this experiment, we focus on how broadband fluxes,spectral fluxes, and heating profiles differ between two dis-tinct baseline cases and how they differ between 5 asso-ciated changes to those baseline cases. The two baselinecases are the standard tropical atmosphere and the stan-dard polar atmosphere. The standard cases both includeno clouds and standard constituent profiles (i.e., O2, WV,O3 fractional amounts set to 1); however, the tropical base-line case has a solar zenith angle set to 71◦, while the po-

lar baseline case has a solar zenith angle set to 84◦. Ourtask was to perform radiative transfer calculations using aradiative-transfer model known as Streamer. Within all of the input files (both baseline and experimental), the timeof the data is 12Z on April 2, 1999. We also chose a cloudoptical thickness of 0.6 µm and water concentration of .07g/m3 for all experiments. The idea was to change the in-put files for each of the 5 experiments and then run those

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files in Streamer. The Streamer model then performs cal-culations and creates new output files that include datafor 25 different levels from the surface to 100 km. Thedata of interest include fluxes of total downward shortwavefluxes (includes both diffuse and direct), downward long-wave fluxes, upward diffuse shortwave fluxes, upward long-wave fluxes, net fluxes, and heating rates (K/day). The

particular experiments we ran included a 10% increase inwater vapor everywhere in the column with no cloud layer,doubling the CO2 concentration with no cloud layer, a 50%reduction in atmospheric O3 concentration with no cloudlayer, a liquid water cloud layer with base of 1.0 km andoptical thickness of 10.0, and a spherical ice cloud layerwith base of 10.0 km and optical thickness of 10.0.

3. Results

a. Standard Cases

One of the most notable differences between the tropi-cal case and the polar case is that the amount of incoming

solar radiation is about 300 W/m2

greater at the top of the atmosphere for the tropical case than it is for the polarcase. As seen in Table 1, the net shortwave flux at both thetop of the atmosphere and at the surface is much greaterfor the tropical case than for the polar case which agreeswith the fact that the tropical atmosphere is warmer. Theupward longwave flux is much greater for the tropical casethan for the polar case, which makes sense because of thelarger amounts of water vapor in the tropical atmospherecompared to the polar atmosphere. As expected, there ismuch more upward and downward longwave radiation atthe surface due the warmer surface temperatures and the“warmer“ atmosphere above due to absorption of solar ra-

diation by ozone, water vapor, carbon dioxide, and othergreenhouse gases. Overall, the incoming shortwave andlongwave combined is greater than the outgoing combinedfor the tropical case at both the surface and the top of theatmosphere with 28.74 W/m2 at the top of the atmosphereand 165.87 W/m2 at the surface. The opposite happensfor the polar case, where the net combined shortwave andlongwave fluxes are negative at the top of the atmosphereand at the surface. This means that overall the outgoingshortwave and longwave fluxes are greater than the incom-ing shortwave and longwave fluxes. The net longwave fluxat the surface is -.17 W/m2 while the net long flux at thetop of the atmosphere is -86.10 W/m2.

What I see in Figure 1 is that in both tropical and po-lar base cases have a shortwave and a longwave peak of spectral fluxes at the surface and at the top of the atmo-sphere. Spectral flux includes energy contributions fromone wavelength in W/m2logµm (Petty (2006)). For thetropical case at the surface, I see a positive net spectral fluxof 250 W/m2logµm in the shortwave spectrum with abouta peak of about 100 W/m2logµm of upward spectral flux

Table 1. Calculated shortwave, longwave, and combinedbroadband fluxes for top of atmosphere (TOA) and sur-face. Broadband fluxes are in units of W/m2 and top of atmosphere is 100 km. Net fluxes are calculated using F ↓

- F ↑, or incoming flux outgoing flux. Therefore, a pos-itive net flux suggests more incoming flux and thereforeexpected heating, while a negative net flux suggests moreoutgoing flux and therefore expected cooling.

Tropical (Top) and Polar (Bottom)

SW ↑ SW ↓ SW netTOA 122.30 441.10 +318.80SFC 69.37 297.48 +228.12

LW ↑ LW ↓ LW netTOA 290.06 0 -290.06SFC 458.22 395.98 -62.24

SW + LW ↑ SW + LW ↓ SW + LW netTOA 412.36 441.10 +28.74SFC 527.59 693.46 +165.87

SW ↑ SW ↓ SW netTOA 56.73 141.62 +84.89SFC 20.69 80.22 +59.53

LW ↑ LW ↓ LW netTOA 170.99 0 -170.99SFC 194.25 134.55 -59.70

SW + LW ↑ SW + LW ↓ SW + LW netTOA 227.72 141.62 -86.10SFC 214.94 214.77 -.17

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and about 300 W/m2logµm of downward spectral flux. Inthe longwave spectrum, I see a negative net spectral flux of about -100 W/m2logµm in the longwave spectrum with apeak of about 350 W/m2logµm of upward spectral flux anda peak of about 310 W/m2logµm of the downward spectralflux. In the top of the atmosphere plot for the tropical at-mosphere, I see a positive net spectral flux of 270 W/m2µm

in the shortwave spectrum with about 100 W/m

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µm up-ward spectral flux and about 300 W/m2logµm downwardspectral flux. In the longwave spectrum, I see a negativenet spectral flux of -300 W/m2logµm in the longwave spec-trum with about 200 W/m2logµm upward spectral flux and0 W/m2logµm downward spectral flux. This makes senseto have no downward longwave emission from outer space,with a temperature of only about 2.7 Kelvin!

For the polar case at the surface, there is a positivenet spectral flux of about 50 W/m2logµm in the short-wave spectrum while there is a negative net spectral fluxof about -100 W/m2logµm in the longwave spectrum. Inthe shortwave spectrum, incoming spectral flux is about

50 W/m2

logµm and outgoing spectral flux is about 25W/m2logµm. In the longwave spectrum, incoming spec-tral flux averages about 50 W/m2logµm and the outgoingspectral flux averages at about 100 W/m2logµm. For thetop of the atmosphere in a polar environment, there is apositive net spectral flux of about 75 W/m2logµm in theshortwave spectrum while there is a negative net spectralflux of about -125 W/m2logµm in the longwave spectrum.In the shortwave spectrum, the incoming spectral flux isabout 100 W/m2logµm and outgoing spectral flux is about50 W/m2logµm. In the longwave spectrum, incoming spec-tral flux is 0 W/m2logµm and outgoing spectral flux isabout 100 W/m2logµm.

The heating profiles from Figure 2 of both tropical andpolar environments show that the majority of the heat-ing and cooling in the atmosphere occurs at the top of the atmosphere. The polar atmospheric heating profileslook to be muc h smoother with height than the tropicalatmospheric heating profiles. I speculate that this maybe caused by the abundant moisture in the tropical atmo-sphere and the fact that water vapor is not as well-mixedthroughout the atmosphere as other gases such as oxygenand nitrogen, for example.

b. Experiments

10% increase in water vapor amount everywhere in the column

An increase in water vapor, a very important green-house gas that absorbs longwave radiation and thus helpsto trap heat, has an effect on the total broadband fluxes inshortwave and longwave wavelengths at the surface and atthe top of the atmosphere. As can be seen from Table 2, anincrease is water vapor leads to less outgoing shortwave ra-

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Fig. 1. Upward, downward, and net spectral fluxes inunits of W/m2logµm at the surface and the top of the at-mosphere for standard tropical (top) and polar (bottom )cases. The x-axis is logarithmic so as to make the areasunder the curves comparable between shortwave and long-wave bands.

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Fig. 2. Heating profiles for shortwave, longwave, and com-bined for standard tropical (top) and polar (bottom ) atmo-spheres.

diation and also less incoming shortwave radiation for boththe tropical and polar cases. So with less outgoing short-wave flux than originally, the net shortwave broadband fluxat the top of the atmosphere is slightly more with the 10%increase in water vapor. As for the effect of water vapor onthe amount of incoming and outgoing longwave fluxes, inboth tropical and polar cases the outgoing longwave flux

at the top of the atmosphere was less when water vaporwas added to the atmosphere. There was no change outgo-ing longwave flux at the surface. The downward longwaveflux from the atmosphere to the surface was greater for theincrease in water vapor. Overall, the net longwave flux atthe top of the atmosphere and at the surface was greaterwith the increased water vapor.

Table 2. Calculated shortwave, longwave, and combinedbroadband flux differences for top of atmosphere (TOA)and surface for a 10% increase in water vapor amount ev-erywhere in the column. Broadband flux differences arein units of W/m2 and top of atmosphere is 100 km. Dif-ferences are calculated by subtracting baseline fluxes fromthe experimental fluxes. Therefore, a positive flux differ-ence suggests that the baseline flux was less than the ex-perimental flux. A negative flux difference suggests thatbaseline flux was greater than the experimental flux.


SW ↑ SW ↓ SW netTOA -.41 0 +.41SFC -.43 -1.18 -.76

LW ↑ LW ↓ LW net

TOA -1.32 0 +1.32SFC 0 .71 +.71

SW + LW ↑ SW + LW ↓ SW + LW netTOA -1.73 0 +1.73SFC -.43 -.47 -.04

SW ↑ SW ↓ SW netTOA -.12 0 +.12SFC -.09 -.27 -.18

LW ↑ LW ↓ LW netTOA -.34 0 +.34

SFC 0 .97 +.97SW + LW ↑ SW + LW ↓ SW + LW net

TOA -12.87 0 +.46SFC -.09 .7 +.78

Based on Figure 3, an increase in water vapor evidentlycauses a slight decrease in all shortwave fluxes and increases

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incoming and net longwave fluxes at the surface. At thetop of the atmosphere, upward flux from below is decreasedand so longwave net spectral flux is increased because lesslongwave radiation leaves the atmosphere. The polar casealso shows a decrease in shortwave net spectral flux andan increase in longwave net spectral flux at the surface aswell as less upward solar and infrared at the top of the

atmosphere, all suggesting a heating of the atmosphere.The magnitudes of the differences in the polar case aregenerally smaller than the magnitudes of the differences inthe tropical case. The differences in spectral fluxes seemto occur around 1 µm, around 6.7 µm, and around 20-70µm. Based on Figure 6, these are the wavelengths (withexception to 1 µm) where water vapor absorption is thestrongest. Therefore it would make sense to see the largestchanges in spectral flux at these wavelengths with a 10%increase in water vapor amount.

It is evident from Figure 4 that an increase in heating,albeit a small increase, occurs with an increase in watervapor in the lower atmosphere below about 10 kilometers

for both tropical and polar atmospheres. There is also avery small decrease in heating above 10 kilometers. Thisis perhaps because most of the water vapor is in the loweratmosphere and so that is where the majority of the long-wave heating occurs, so a 10% increase in water vapor inthe upper atmosphere is virtually negligible because thereis so little water vapor to begin with.

Doubling CO 2 concentration

By doubling CO2, or increasing the carbon dioxide by100%, we see a similar effect to what increasing water va-

por did to the shortwave and longwave fluxes. The changesin shortwave spectral fluxes at the top of the atmosphereand at the surface are generally less than the changes thatoccurred in the first experiment. In both tropical and po-lar atmosphere, though, the changes in longwave spectralfluxes are greater for the doubling of CO2 than they are forthe 10% increase in water vapor. For the combined short-wave and longwave spectral fluxes at the top of atmosphereand at the surface, there is an increased positive spectralflux for both model atmospheres.

Based on Figure 5, it appears that there is virtually nochange in the shortwave spectral fluxes for both tropicaland polar atmospheres and for both the top of the atmo-

sphere and the surface. In the tropical plot with the spec-tral flux differences at the surface, it looks as though thereis a slight decrease of upward spectral fluxes in the short-wave and a decrease in downward spectral fluxes in near-IRwavelengths around 2.5 µm. Near-IR radiation (between1 µm and 10 µm) is a continuation of the visible band inthat the primary source of this radiation in the atmosphereis the Sun (Petty (2006)). Based on Figure 6 taken from

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Fig. 3. Changes in upward, downward, and net spectralflux in units of W/m2logµm at the surface and at the topof the atmosphere relative to the baseline tropical (top)and polar (bottom ) cases. The x-axis is logarithmic so asto make the areas under the curves comparable betweenshortwave and longwave bands.

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Fig. 4. Associated changes in heating profiles for short-wave, longwave, and combined for standard tropical (top)and polar (bottom ) atmospheres with a 10% increase inwater vapor amount everywhere in the column.

Table 3. Calculated shortwave, longwave, and combinedbroadband flux differences for top of atmosphere (TOA)and surface for a doubling of carbon dioxide. Broadbandflux differences are in units of W/m2 and top of atmosphereis 100 km. Differences are calculated by subtracting base-line fluxes from the experimental fluxes. Therefore, a pos-itive flux difference suggests that the baseline flux was lessthan the experimental flux. A negative flux difference sug-gests that baseline flux was greater than the experimentalflux.



LW ↑ LW ↓ LW netTOA -2.36 0 +2.36SFC 0 .46 +.46

SW + LW ↑ SW + LW ↓ SW + LW netTOA -2.49 0 +2.49SFC -.12 -.05 +.06


LW ↑ LW ↓ LW netTOA -.87 0 +.87SFC 0 1.77 +1.77

SW + LW ↑ SW + LW ↓ SW + LW netTOA -.94 0 +.94SFC -.05 1.51 +1.56

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A First Course in Atmospheric Radiation by Grant W.Petty, right around 2.5 µm there is virtually no transmis-sion through the atmosphere due to CO2. In other words,CO2 strongly absorbs wavelengths around 2 µm. This islikely why we see a decrease of about 2 W/m2logµm of in-coming near-IR radiation at the surface with a doubling of CO2. We also see a slight signature of decreased incoming

near-IR radiation in the surface plot for the polar atmo-sphere, but not nearly as significant as a change that wesee in the tropical case.

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Fig. 5. Changes in upward, downward, and net spectralflux difference in units of W/m2logµm at the surface andthe top of the atmosphere relative to the baseline tropical(top) and polar (bottom ) cases. The x-axis is logarithmic so

as to make the areas under the curves comparable betweenshortwave and longwave bands.

In the longwave spectrum, we see most changes of spec-tral flux occur right around 15 µm. This makes sense be-cause again according to Figure 6, CO2 strongly absorbsradiation at 15 µm. This increased absorption due to dou-bling of CO2 causes an increase in downward infrared ra-diation at the surface and thus an increase in net longwave

spectral flux. At the top of the atmosphere around 15 µm,there is a decrease in upward spectral flux from the atmo-sphere below and thus less cooling of the atmosphere whichhelps increase the net spectral flux at the top of the atmo-sphere. The differences in spectral flux from the baselinecases are quite significant. For the tropical atmosphere, wesee an increase of about 5 W/m2logµm to net longwave

spectral flux at the surface and an increase of about 20W/m2logµm at the top of the atmosphere. For the polaratmosphere, we see an increase of about 15 W/m2logµm tonet longwave spectral flux at the surface and an increaseof about 6 W/m2logµm at the top of the atmosphere.

Fig. 6. Zenith transmittance through a cloud-free andaerosol-free atmosphere for conditions typical of mid-latitude summer. Absorption contribution due to individ-ual gases is shown, with the bottom chart showing totalcontribution by all gases. (Figure obtained from A First Course in Atmospheric Radiation by Grant W. Petty)

Based on Figure 7, the changes in heating profiles inthe tropical and polar atmospheres after a doubling of CO2

look very similar. There is virtually no change from the sur-face to about 20 kilometers, if anything a slight increasein longwave heating that is more evident in the tropicalcase. A decrease in longwave heating (and thus increase inlongwave cooling) occurs above 20 kilometers for both thetropical and polar cases.

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Fig. 7. Associated changes in heating profiles for short-wave, longwave, and combined for standard tropical (top)and polar (bottom ) atmospheres with a 50% increase inCO2 concentration.

50% reduction in ozone concentration

In this experiment, a greenhouse gas is now being re-duced to see what kind of changes occur in broadbandfluxes, spectral fluxes, and heating profiles. As seen in Ta-ble 4, we get the opposite effect where upward shortwavebroadband fluxes have slightly increased at the top of the

atmosphere, and at the surface, implying less absorptionand thus more reflection. We also see that downward short-wave broadband fluxes have increased at the surface, againimplying less absorption and more transmission. For long-wave broadband fluxes, we see an increase in upward long-wave emission, which agrees with less absorption by theatmosphere. In other words, more emission of thermal in-frared radiation from the surface and atmosphere is makingit through to the top of the atmosphere instead of getting

absorbed by greenhouse gases. Therefore, we see a decreasein net longwave at the top of atmosphere because now moreradiation is leaving than is coming in from outer space. Atthe surface, we see slightly less incoming longwave radia-tion from the atmosphere above because the atmosphere isemitting less because it is absorbing less. We see similarchanges occurring for the polar model, with comparable

magnitudes in shortwave broadband flux changes but withmuch smaller magnitudes than the changes we see in thetropical longwave broadband fluxes.

Table 4. Calculated shortwave, longwave, and combinedbroadband flux differences for top of atmosphere (TOA)and surface for a 50% in ozone concentration. Broadbandflux differences are in units of W/m2 and top of atmosphereis 100 km. Differences are calculated by subtracting base-line fluxes from the experimental fluxes. Therefore, a pos-itive flux difference suggests that the baseline flux was lessthan the experimental flux. A negative flux difference sug-gests that baseline flux was greater than the experimentalflux.


SW ↑ SW ↓ SW netTOA .91 0 -.91SFC .06 1.19 +1.12

LW ↑ LW ↓ LW netTOA 2.47 0 -2.47SFC 0 -.78 -.78

SW + LW ↑ SW + LW ↓ SW + LW netTOA 3.38 0 -3.38

SFC .06 .41 +.35

SW ↑ SW ↓ SW netTOA .89 0 -.89SFC .07 .99 .92

LW ↑ LW ↓ LW netTOA .71 0 -.71SFC 0 -.86 -.86

SW + LW ↑ SW + LW ↓ SW + LW netTOA 1.6 0 -1.6SFC .07 .13 +.06

Based on Figure 8, in the tropical atmosphere we see anincrease in incoming shortwave spectral fluxes at the sur-face with a reduction in ozone concentration. The amountof shortwave reflecting off of the surface is about the same.We see similar changes in shortwave spectral fluxes at thesurface in the polar atmosphere. For the top of the atmo-

8



sphere, there are very small changes in shortwave spectralfluxes. At the top of the polar atmosphere, we see changesof the same magnitude in shortwave spectral fluxes. We seean increase in upward shortwave at the top of the atmo-sphere and no change in downward shortwave from the Sun,thus leading to an overall decrease in net shortwave. Wefind much larger changes in the longwave spectral fluxes,

with a decrease of about 10 W/m

2

logµm in downwardfluxes at the surface for about 10 µm wavelength. Forthe polar atmosphere, we see about the same amount of decrease in downward longwave radiation. For the top of the atmosphere, we see a general increase in upward long-wave radiation of about 30 W/m2logµm in the tropicalatmosphere and only an increase of about 8 W/m2logµmin the polar atmosphere. In both cases, however, we seethe amount of longwave leaving the atmosphere increaseswith a reduction in ozone.

Based on Figure 9, the changes in heating profiles inthe tropical and polar atmospheres after a 50% reductionin ozone concentration look very similar. There is virtually

no change in heating profiles from the surface to about 10kilometers. Above 10 km, it appears shortwave heating isdecreased and that longwave heating is slightly decreasedas well before it starts to increase above 32 kilometers inthe tropical atmosphere and above 20 kilometers in the po-lar atmosphere. Overall, the differences in heating profilesare greater in the tropical case than the polar case. Thecombined shortwave and longwave heating profiles throughthe atmosphere are less after the decrease in ozone. Thismakes sense because a decrease in a greenhouse gas shouldin fact help to decrease temperatures since there should beless absorption in the atmosphere.

Add liquid water cloud layer with optical thickness 10.0 and base at 1.0 km

Now we have added a low, warm cloud into the atmo-spheres. The associated changes in broadband fluxes in Ta-ble 5 are much more significant than in the previous caseswith changes in atmospheric constituents in cloud-free at-mospheres. In the tropical case, we see a large increasein upward shortwave broadband fluxes of 125.95 W/m2

reaching the top of the atmosphere. This increase in outgo-ing shortwave fluxes and the fact that incoming shortwavefluxes have not changed leads to a decrease in net short-

wave fluxes at the top of the atmosphere. At the surfacein the tropical case, we see an incredible decrease in short-wave fluxes of about 165.66 W/m2 and thus a decreasein upward shortwave at the surface leading to an overallshortage of net shortwave broadband fluxes. The low cloudessentially blocks a large portion of sunlight from gettingto the surface, and at the same time reflects a significantamount of solar radiation back out to space. These pro-

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Fig. 8. Changes in upward, downward, and net spectralflux difference in units of W/m2logµm at the surface andthe top of the atmosphere relative to the baseline tropical(top) and polar (bottom ) cases. The x-axis is logarithmic soas to make the areas under the curves comparable betweenshortwave and longwave bands.

9



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Fig. 9. Associated changes in heating profiles for short-wave, longwave, and combined for standard tropical (top)and polar (bottom ) atmospheres with a 50% reduction inO3.

cesses caused by adding a low cloud layer all contribute tocooling of the Earths climate. In the polar atmosphere, wesee similar increases and decreases in shortwave radiation,but of much smaller magnitudes than we see in the trop-ical atmosphere. In the tropical atmosphere, the upwardlongwave broadband fluxes at the top of the atmosphereare less than what was seen without the cloud layer, likely

because the cloud is colder than the underlying atmosphereand surface and so is emitting less thermal infrared thanthe surface would if it were not blocked by a cloud. Atthe surface, we see an increase in the downward longwavebroadband fluxes of 49.2 W/m2 due to the low, warm cloudemitting radiation back down to the surface. We see some-thing different going on in the polar case, where we seeupward longwave slightly increased at the top of the atmo-sphere. We can attribute this to the fact that over polarregions, the atmosphere is often warmer than the underly-ing surface! The downwelling longwave radiation seen atthe surface is quite higher than we see in the tropical sit-uation with an increase of about 77.48 W/m2. This again

may have to do with the fact that the surface is colder thanthe atmosphere in polar regions and so the cloud may bequite a bit warmer than the surface instead of very closeto surface temperature in the tropical case.

We can easily see from Figure 11 that the most signif-icant changes in heating profiles occur at 1 km where wehave a considerable decrease in longwave heating in boththe tropical and polar atmospheres. The decrease in heat-ing is by about 5 K/day for the tropical atmosphere andabout 7 K/day for the polar atmosphere. So this exper-iment helps to show how low clouds help to cool surfacetemperatures.

From Figure 10 we can see that for the tropical at-mosphere with a cloud layer at 1 km, incoming shortwavespectral fluxes at the surface are significantly decreased byabout 150 W/m2logµm and outgoing shortwave spectralfluxes are decreased as well. At the top of the atmosphere,upward shortwave spectral fluxes increased by about 150W/m2logµm due to the increased reflection off of the newcloud layer. The incoming longwave spectral fluxes at thesurface have increased due to emission from the ”warm”cloud. At the top of the atmosphere, upward longwavespectral fluxes have decreased. For the polar atmospherewith a cloud layer at 1 km, incoming shortwave spectral

fluxes at the surface also decreased but only by about 50W/m2logµm. We see the opposite situation happening forupward longwave spectral fluxes at the top of the polaratmosphere. The upward longwave spectral fluxes have in-creased from the case without any cloud layer.

We can easily see from Figure 11 that the most signif-icant changes in heating profiles occur at 1 km where wehave a considerable decrease in longwave heating in both

10



Table 5. Calculated shortwave, longwave, and combinedbroadband flux differences for top of atmosphere (TOA)and surface for a liquid cloud layer at 1km and opticalthickness of 10.0. Broadband flux differences are in unitsof W/m2 and top of atmosphere is 100 km. Differences arecalculated by subtracting baseline fluxes from the experi-mental fluxes. Therefore, a positive flux difference suggeststhat the baseline flux was less than the experimental flux.A negative flux difference suggests that baseline flux wasgreater than the experimental flux.


SW ↑ SW ↓ SW netTOA 125.95 0 -125.95SFC -37.51 -165.66 -128.16

LW ↑ LW ↓ LW netTOA -14.57 0 +14.57SFC 0 49.2 +50.33

SW + LW ↑ SW + LW ↓ SW + LW netTOA 111.38 0 -111.38SFC -37.51 -116.46 -78.95


LW ↑ LW ↓ LW netTOA 8.31 0 -8.31SFC 0 77.48 +77.48

SW + LW ↑ SW + LW ↓ SW + LW netTOA 42.81 0 -42.82SFC -12.13 30.52 +42.64

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the tropical and polar atmospheres. The decrease in heat-ing is by about 5 K/day for the tropical atmosphere andabout 7 K/day for the polar atmosphere. So this exper-iment helps to show how low clouds help to cool surfacetemperatures.

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Fig. 11. Associated changes in heating profiles for short-wave, longwave, and combined for standard tropical (top)and polar (bottom ) atmospheres with a liquid water cloudlayer at 1 km and with an optical thickness of 10.0.

Add ice cloud layer with optical thickness 10.0 and base at 10.0 km

We see quite a different situation when an ice cloudlayer at 10 km is added to the tropical and polar atmo-spheres. Looking at Table 6, for the tropical atmospherethe upward shortwave broadband fluxes have increased atthe top of the atmosphere, although not as much of anincrease as what we saw for the 1 km cloud layer. Thedownward shortwave broadband fluxes have decreased atthe surface by about the same amount as in the 1 km cloud

layer case. Upward longwave broadband fluxes at the topof the atmosphere have decreased a lot more in this highcloud case since high clouds tend to better trap longwaveradiation emitting from the surface. The increased incom-ing longwave fluxes at the surface are much less than the1 km cloud layer case, since these higher, colder cloudswill emit less radiation back towards the surface than the

1 km cloud layer will. For the polar atmosphere, we havequite a different situation with the longwave fluxes. Theupward longwave fluxes at the top of the atmosphere are51.8 W/m2 less than what they were in the baseline case.This is probably because of the high clouds being so coldthat they dont emit too much radiation up towards thetop of the atmosphere. At the surface in the polar region,we have an increase in the downward longwave fluxes byabout 36.11 W/m2, not as much as the 77.48 W/m2 fluxincrease that occurred with the inclusion of the 1 km cloudlayer.

In Figure 12 we see that for the tropical atmosphere atthe surface, the shortwave spectral flux changes look very

similar to the changes we saw with the 1 km cloud layer.For the longwave spectral fluxes, however, we see less of anincrease in downward fluxes at the surface for the tropicalatmosphere than we saw in the 1 km cloud layer experi-ment. At the top of the atmosphere for the tropical case,we see that upward longwave spectral flux decreases insteadof increasing like in the 1 km cloud layer case. This is be-cause the emission will be less from the top of the 10 kmcloud layer than it would be for a cloud-free atmospherethat includes warmer emission closer to the surface. Wesee similar changes for the polar atmosphere, only smallermagnitudes of the fluxes at the top of the atmosphere. Wegenerally see larger positive differences in the downward

longwave spectral fluxes at the surface for the polar casethan we see for the tropical case. For the polar atmosphere,the low clouds seem to cause a larger increase in downwardlongwave radiation at the surface than the high clouds do.Also, the upward longwave radiation at the top of the at-mosphere increased in the 1 km cloud layer case but thendecreased in the 10 km cloud layer case (likely due to thediffering cloud-layer temperatures).

For the 10 km cloud layer experiment, we see from Fig-ure 13 that heating profiles have generally increased belowthe 10 km level in both tropical and polar atmospheres.In the tropical case at 10 km, we have an increase in bothshortwave and longwave heating. In the polar case at 10

km, we have increased heating in the shortwave and signif-icantly decreased heating in the longwave. Again I thinkthe main reason why there is an increase in longwave heat-ing in the tropical case and a decrease in longwave heatingfor the polar case may have to do with the polar regionswarmer atmosphere and colder surfaces. Both cases show,however, that high clouds tend to increase the overall com-bined heating below the cloud layer.

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Table 6. Calculated shortwave, longwave, and combinedbroadband flux differences for top of atmosphere (TOA)and surface for an ice cloud layer at 10 km and opticalthickness of 10.0. Broadband flux differences are in unitsof W/m2 and top of atmosphere is 100 km. Differences arecalculated by subtracting baseline fluxes from the experi-mental fluxes. Therefore, a positive flux difference suggeststhat the baseline flux was less than the experimental flux.A negative flux difference suggests that baseline flux wasgreater than the experimental flux.




SW + LW ↑ SW + LW ↓ SW + LW netTOA 15.7 0 -15.7SFC -36.78 -148.36 -111.58



SW + LW ↑ SW + LW ↓ SW + LW netTOA -9.72 0 +9.72SFC -11.71 -9.5 +2.2

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Fig. 13. Associated changes in heating profiles for short-wave, longwave, and combined for standard tropical (top)and polar (bottom ) atmospheres with an ice cloud layer at10 km and with an optical thickness of 10.0.

4. Conclusions

In conclusion, it seems quite clear to me that a criticalfactor in our climate is the amount of cloud cover. Petty(2006) Not only does the extent of cloud cover matter, butthe height of the clouds seems to have considerably differ-ent impacts on the climate. Low, warm clouds seem to

help cool the surface temperatures, while high, cold cloudstend to warm the surface temperatures. Low clouds tendto reflect more incoming solar radiation and so cause anincrease in upward shortwave radiation. Low clouds alsoprevent solar radiation from reaching the surface, while atthe same time low clouds can help to emit more longwaveradiation to the surface due to warmer temperatures emit-ted from the cloud. The longwave radiation seen at the top

of the atmosphere will be from the cloud layer below, and sothe level of these cloud layers are of the utmost importancein determining the amount of longwave radiation emittedfrom that layer. It was found that low clouds cooled sur-face temperatures and lower atmospheric temperatures byapproximately 5 K/day. High clouds appeared to causeless combined heating at 10 kilometers in polar regions but

caused more combined heating at 10 kilometers in tropicalregions. For both the tropical and polar cases in the highcloud experiment, we see less shortwave reflection off of thecloud layer (thus implying more sunlight gets through theclouds) and thus we see heating below the cloud layer. Wealso see much less longwave radiation at the top of the at-mosphere with the high clouds, likely because high cloudsare cold and thus emit less longwave and also because highclouds help to trap longwave emission from the atmospherebelow.

A change in the amount of an atmospheric gases af-fects longwave fluxes more so than shortwave fluxes. Thisis because certain gases such as water vapor, CO2, and O3

absorb radiation of certain long wavelengths. Water vaporabsorbs at wavelengths around 2.5 µm, 6.7 µm, and at 20µm. CO2 absorbs at wavelengths around 4 µm and 15 µmand O3 absorbs at 9.6 µm and a little bit at 5 µm. There-fore, an increase in greenhouse gases helps to decrease out-going longwave fluxes out of the atmosphere, thus increas-ing longwave radiation that is trapped in the atmosphere.A decrease in greenhouse gases does the opposite in thatit leads to less absorption by gases and thus less trappedheat. Less absorption leads to cooler surface temperaturesthan what would be if there were more greenhouse gases.

Of course in the real atmosphere we cant simply turn onone constituent and turn off another. We have all the con-

stituents filling the atmosphere and we have variable watervapor amounts and variable cloud cover every day. Theever-changing atmosphere and even the greenhouse gasesthat humans add to the atmosphere make it difficult topredict how our climate will respond. These experimentsprovide some insight into how the radiative balance changesdue to a particular change in individual atmospheric gasesand due to both water and ice clouds (i.e. low and highclouds, or warm and cold clouds). Of course these are alsoresults that came from a numerical model which always hasits setbacks due to approximations and the issue of missingsmall-scale atmospheric processes due to the resolution of the model. Nevertheless, the more experiments that are

performed to see how the atmosphere responds to changesin certain parameters can only give atmospheric scientistsmore information to try to interpret and understand. Notonly do we want to see what radiative transfers changeswould take place with certain atmospheric changes, butwe want to try to predict what those atmospheric changeswill be. In other words, do we expect to have more highclouds or low clouds in our future? Do we expect signifi-

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cant increases in greenhouse gases in our future? As modelscontinue to improve and the amount of observations con-tinues to increase into the future, we can know with morecertainty what Mother Nature has in store for us.

REFERENCES

Petty, G., 2006: A first course in atmospheric radiation .Sundog Pub., URL http://books.google.com/books?

id=Q5sRAQAAIAAJ.

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