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close this bookForests, Climate, and Hydrology: Regional Impacts (UNU, 1988, 217 pages)
close this folder8. Review of general circulation models as a basis for predicting the effects of vegetation change on climate
View the document(introductory text...)
View the documentAbstract
View the documentIntroduction
View the documentGeneral circulation models
View the documentResponse to variation in land surface properties
View the documentGCM Simulations of tropical rainfall
View the documentRecommendations for future research
View the documentSymbols and abbreviations
View the documentReferences
View the documentAssessment

Response to variation in land surface properties


Surface Solar Atbedo (a *). The surface albedo is of major importance in determining the absorption of solar energy. Large variations are possible due to vegetation. Generally albedo decreases for a given vegetation type as the height of the vegetation increases because of internal reflections, though for short, sparse vegetation this may not be true if the albedo of the soil is low relative to that of the vegetation. Albedo also decreases generally for soil and vegetation as the surface wetness increases. The range of albedo for snow-free conditions is from about 0.1 for tropical forest to about 0.4 for some dry, sandy surfaces. With incident mean daily solar fluxes typical of the tropics (300-400 Wm-2 in cloudless conditions), a variation of about 100 Wm-2 is possible. However, this would be an upper limit to spatial variations and is only conceivable at one place with an extreme climatic change or extensive human intervention such as deforestation or irrigation on highly reflective soil. Actual changes on a large scale (1,000 km and greater) seem unlikely to exceed about half this magnitude, or 50 Wm-2. These could be due to variations in surface wetness (Idso et al. 1975, Norton, Mosher, and Hinton 1979) or replacement of forests by grassland or dry soil.

Much larger variations in albedo can occur due to snow cover, which may have an albedo of over 0.9. This is of significance in the context of deforestation in middle and high latitudes where, because snow seldom covers trees for long, the effects of snow on the heat budget are less in forested regions. In lower latitudes, because of the lack of a large seasonal variation, forests are unlikely to exist above the snow-line.

Variations in surface albedo may be expected to affect local climate in two ways through their modification of the net radiation. Firstly, by equation (14), a reduction in RN(0) must decrease the energy available for upward transfers of sensible and latent heat and for downward transfer into the soil. The Penman-Monteith equation (9) may be used to estimate the variations of the components of the turbulent fluxes as (RN - G) varies. For typical tropical temperatures of 26°C,


so that


For given d q/raE, the change in LE will be a fraction x = 3/(4 + rs/raE) of the change in (RN - G). The value of x varies from about 0.75 for a moist surface to zero for a dry surface. Thus, as (RN - G) decreases, the evaporation will decrease: for incoming solar radiation of 250 W m-2, a change in albedo of 0.1 will decrease (Rx - C) by 25 W m-2 if C is unchanged, reducing evaporation by 0.65 mm d-1 if rs = 0. Whether relative humidity will be decreased or increased will depend on both temperature and rs/raE. If it is decreased, precipitation is likely to be reduced (see discussion below in Surface Moisture Availability).

The second mechanism by which surface radiative characteristics can modify tropical climate is that discussed by Charney (1975), whereby heating of a vertical column (surface and atmosphere) relative to adjacent regions can increase the ascent of air masses and vice versa. Thus the geographical distribution of heat sources and sinks leads to vertical circulations that are mainly responsible for limiting the atmospheric temperature gradients to values in line with the dynamical constraint of the low rotation rates found in the tropics. Thus an increase in surface albedo, by weakening the heat source, tends to reduce the ascent of air masses and the associated rainfall. This mechanism can be enhanced by the drying of the atmosphere (which increases upward long wave radiative flux) and drying of the land surface, which can cause surface heating and so also increase upward long wave radiation. These may be counterbalanced by a decrease in cloudiness, reducing reflected solar radiation more than increasing upward long wave radiation.

An impression of the possible impact of albedo changes is gained by noting that typically in the tropics the net radiative heating of the surface of ~ 150 Wm-2 is offset by an atmospheric radiative cooling of ~ 100 Wm-2. With clear-sky solar radiation of 400 Wm-2 a 0.1 increase in albedo will thus eliminate most of the net heating.

GCM Experiments with Surface Albedo Changes. For a more quantitative estimate of the effects of albedo changes we must consider the results of GCM experiments. However, in doing so it must be remembered that the complexity of the interactions, between radiation, cloud, and atmospheric and surface moisture content, requires a degree of realism in modelling surface and cloud processes that is probably not yet attainable. The GCM experiments made to date, therefore, cannot be expected to give more than an indication of the likely impacts.

Studies with GCMs on the effects of albedo variations, some of which include albedo changes in more than one area, are listed in table 2. All are for northern summer. Except as indicated the models included interactive surface hydrology and radiative transfer dependent on modelling water vapour and cloud.

Although the global-scale change in albedo studied by Carson and Sangster is not a realistic change, their results are instructive in showing clearly the major features of a model's response. In their first experiment (fig. 1a) albedos of all snow-free land were increased from 0.1 to 0.3; the land was kept wet to avoid confusing soil moisture-albedo interactions. The major effects were:

  1. higher surface pressure over land (by up to 12 mb in the inner northern continents);
  2. decreased average evaporation over land (from 3.6 to 2.7 mm d-1);
  3. decreased average precipitation (and ascent) over land (from 4.6 to 3.4 mm d-1) with compensating increases over oceans (from 2.9 to 3.2 mm d-1).

Note that the decrease in precipitation exceeded that in evaporation, showing it was partly due to a decrease in moisture convergence. These results are consistent with the above discussion.

Similar results were obtained in the second experiment (fig. 1b), in which soil moisture was interactive. In both experiments —but especially the second—some land areas, mostly on the western edge of continents (western North Africa, Europe), were wetter with the higher albedo, probably because the higher pressures over the land mass to the east decreased the equatorward advection of dry air.

The other experiments in table 2 all involved albedo increases over relatively small areas. The principal results are:

  1. evaporation was decreased in all cases;
  2. precipitation was decreased in all experiments except two, where changes were small: both were near mountain ranges (Rockies, Himalayas) and poleward of latitude 20°;
  3. moisture convergence was decreased in all but three cases, they being poleward of latitude 20°.

Although for Chervin's (1979) experiment there is no direct evidence concerning (i) and (iii), indirect evidence is provided by the maps, showing decreased ascent and reduced soil moisture with increased albedo, especially over the Sahara.

For the areas equatorward of 20° latitude that should be most relevant to tropical deforestation, the ratios of the fractional change in rainfall to the change in albedo (last column of table 3) average - 2.1, suggesting a 21% decrease in rainfall for a 0.1 increase in albedo. On average, for the limited area anomalies more than half the change is due to moisture convergence, though there is considerable variability in this, particularly between Charney et al.'s (1977) results and those of Sud and Fennessy (1982).

TABLE 2. Experiments on effect of surface albedo changes

Reference Area Latitude Averaging period Albedo Comments
Control Modified
Carson and Sangster 1981 Global 90 days 0.1 0.3 1, 2, 4
Global 90 days 0.2 0.3 1, 2
Charney et al. 1977 Sahel 12°-16°N 31 days 0.14 0.35 4
NW India 24°-32°N        
Great Plains 32°-48°N        
Central Africa 8°-12°N        
Bangladesh 20°-28°N        
Mississippi 32°-48°N        
Sud and Fennessy 1982 Sahel 12°-20°N 31 days 0.183 0.3  
NW India 24° 32°N 31 days 0.15 0.3  
NE Brazil 4°-24°S 31 days 0.091 0.3  
Great Plains 32°-48°N 31 days 0.129 0.3  
Chervin 1979 Sahara 7.5°- 37.5°N 60 days .08-.35 .45  
W. USA 27.5°-52.5°N 60 days .07-.17 .45  
Henderson- Sellers and Gornitz 1984 Brazil 15.6°S-7.8°N 5 years .11 .17 3, 5

1. No moisture feedback in radiation
2. No cloud feedback in radiation
3. Roughness length and ground water capacity also changed
4. Evaporation set equal to potential evaporation
5. Ocean surface temperatures and sea ice not prescribed

An important aspect of the results is the effect of albedo on cloud amount. Here also Sud and Fennessy's and Charney et al.'s results differ, the latter finding decreases in cloudiness of 7 to 24% in five of their six cases, whereas Sud and Fennessy obtained decreases of at most 4%. In consequence, radiative impacts were greater in Sud and Fennessy's experiment (30 W m-2 in net surface radiation compared with 17), yet rainfall was less affected.

FIG. 1a. Excess of precipitation with an albedo of 0.1 relative to that with an albedo of 0.3 at Days 21 to 110 (soil moisture fixed at 15 cm).

FIG. 1b. Excess of precipitation with an albedo of 0.2 relative to that with an albedo of 0.3 (interactive initially zero soil moisture).

TABLE 3. Effects of albedo

Experiment Change in albedo
(d a.)
Change in evap.
(d E) (mm/d)
Change in rainfall (d R) (mm/d) Change in moisture convergence (d R-d E) (mm/d) d R/R d R/R d a .
Charney et al. 1977
Sahel .21 - 0.9 - 3.4 - 2.5 - .46 - 2.2
NW India .21 - 0.5 - 2.6 - 2.1 - .53 - 2.5
Great Plains .21 - 1.0 - 1.5 -0.5 - .41 -2.0
Central Africa .21 - 0.7 - 3.1 - 2.4 - .62 - 3.0
Bangladesh .21 - 0.2 0 0.2 0 0
Mississippi .21 - 1.6 - 1.1 0.5 - .25 - 1.2
Sud and Fennessy 1982
Sahel .12 - 0.7 - 1.5 -0.75 - .26 - 2.1
NW India .15 -0.2 -0.5 -0.3 -.13 -0.9
NE Brazil .21 - 0.3 - 0.5 - 0.25 - .24 - 1.1
Great Plains .17 -0.2 0 0.25 .01 0.1
Chervin 1979
Sahara - .22 - ~ - 2.5 - ~ - .4 ~ - 1.8
W. USA - .33 - ~ - 1 - ~ - .15 - 0.5
H.-S. and Gornitz 1984
Brazil .06 -0.45 -0.6 - 0.15 ~ - .12 - 2.0
Carson and Sangster 1981
0.1 (r) 0.3 .2 - 0.95 - 1.22 -0.27 - .27 - 1.3
0.2 (r) 0.3 .1 - - 0.4 - - .14 - 1.4

It is to be expected that an increase in surface albedo will tend to reduce the mean global temperature. However, of the available experiments only that of HendersonSellers and Gornitz can show this since ocean surface temperatures were prescribed in the others. Although no details of the geographical distribution are available, a mean cooling of 0.10 K was obtained over the last year of their experiment (HendersonSellers, personal communication). This is a little greater than expected with the prescribed mean albedo change of 8 x 10-4 from results of experiments with one dimensional models (e.g. Sagan, Toon, and Pollack 1979), which suggest cooling of about 1 K for a surface albedo increase of 0.01. A rather larger response is to be expected from three-dimensional models (e.g. Manabe and Wetherald 1975) because of the possibility of snow/ice albedo feedback.

Carson and Sangster (as reported by Rowntree 1982) made experiments with albedo dependent on soil moisture through a simple linear relation from 0.15 for a moist to 0.30 for a dry surface. This provides a positive feedback to any tendency for change in surface moisture and may be expected to enhance contrasts between wet and dry regions. This expectation was confirmed with sharpened contrasts between southern India and north-western India/Pakistan and across the Sahel by days 21 to 50; later, however, there was a northward expansion of rainfall over western North Africa, probably associated with the higher albedo over much of Africa and Asia for the reasons suggested above in connection with Carson and Sangster's albedo experiments. Laval (1983) has also recently run experiments with soil moisture dependent albedos in the LMD model, as well as with increased prescribed Sahel albedos; with the latter she noted a tendency for weakening of the upper tropospheric easterlies in the Sahel region.

Surface Long Wave Emissivity. The effect of a change in surface long wave emissivity on the net surface radiation may be written:


Typically in the tropics

(e.g. Viswanadham 1972), so that if



A survey of data on surface emissivities was presented by Kondratyev et al. (1982), who commented on the inadequate spectral detail of the available observations. Their limited data suggest that the emissivity integrated over the long wave spectrum is about 0.92 for soils and 0.94 for vegetation, while for the 9-12µm "water vapour window," the mean for dry soils is 0.96 and for vegetation 0.98. Whilst the relative uncertainties in these figures must be substantial, the suggested value of 0.02 for d e * for both the integrated emissivity appropriate for upward radiation and the water vapour window, which is more appropriate for the net long wave radiation, is so small that one may have sufficient confidence in the consequent estimate of

to assert that the effects of deforestation on the net radiation through emissivity changes can be neglected in comparison to the changes due to surface albedo. The only relevant GCM experiment known to the writer is that reported briefly by Hansen et al. (1983), who found little effect on the general circulation or atmospheric temperatures" from the introduction of spectrally dependent emissivities for deserts, snow, and ice.