Equilibrium response to a change in vegetative cover

As is known, in the presence of vegetation, not all precipitated moisture reaches the underlying surface. Part of it is detained by the vegetation (the effect of interception). The evaporation of the moisture detained by vegetation increases the water vapour content in the atmosphere and (in the case of condensation of water vapour) the intensity and duration of precipitation. The moisture is extracted from the soil by the roots of plants and, by means of transpiration via leaf stomata, returns to the atmosphere. The rate of transpiration depends on density, height and the type of vegetative cover. For example, for grass cover transpiration is approximately equal to evaporation, while for a tall wood transpiration can be more than evaporation by an order of magnitude. Moreover, it can happen (as in a dense, deciduous forest) that the latent heat flux at the surface of the vegetative cover and soil is directed into the atmosphere, and the sensible heat flux if soil is in shadow is directed to the underlying surface. It is clear that such a situation occurs only in limited areas of continuous forest tracts. When increasing spatial scales, the effects of the horizontal sensible and latent heat transport and of precipitation dominate.

The mechanism of redistribution of heat and moisture between soil and vegetative cover mainly affects the formation of diurnal oscillations of meteorological parameters. But this does not exclude the fact that vegetation can affect processes with long time scales. We illustrate this by the following example. Wood usually decreases the maximum run-off during snow melting and torrential rains, and at the same time prevents rivers becoming shallow in rainless periods. According to Baumgartner (1979) the ratio between run-off and precipitation in forest tracts amounts to 0.18; in the open country it amounts to 0.42.

It is established that vegetation in droughty regions tends to minimize water loss by soil, and in regions with excessive moisture it tends to maximize biomass at the expense of increase in vegetation density. Different kinds of vegetation show this tendency in different ways: in arid regions plants turn their leaves in the direction of the sun's rays, and in regions with excessive moisture they turn their leaves perpendicular to the sun's rays, thereby decreasing transpiration and evaporation in the first case and increasing them in the second. There are cases of more complicated organization where the upper plant leaves shield the lower ones without regard to the degree of their moisture or lack of moisture in the overlying air layer. The estimations of tranpiration in arid regions and in regions with excessive moisture point to the fact that the tendency mentioned above is observed over a wide range of spatial scales from 10 to 100 km. This gives hope for a proper description of the effects of large plant associations in global climatic models.

Until recently the processes of heat and moisture exchange between the vegetative cover and the atmosphere have been described by the introduction of different kinds of dependencies of exchange intensity on the humidity and roughness of the underlying surface and/or on horizontal coordinates. The biospheric models proposed by Dickinson et al. (1986), and Sellers et al. (1986) take into account the influence of the physiological characteristics of plants on heat and moisture exchange. In these models, the temperature and mass of drop-shaped moisture at the surface of the vegetative cover, soil moisture content at certain levels and the mass of water at the soil surface are used as dependent variables. The transpiration and water transport via plant roots are parametrized in terms of the resistance law according to which the flux of the substance examined is assumed to be proportional to the difference in its values at boundaries, with the proportionality factor characterizing an exchange rate. The latter is considered to be a function of the coefficient of diffusive resistance for the underlying surface and vegetative cover and of the rate of moisture exchange inside individual plants and at their surfaces. For transpiration, the resistance coefficient is represented as the sum of the leaf diffusive resistance and stoma resistance depending on the value of solar radiation, season and amount of water extracted by the plant from the soil. The root resistance is determined by the degree of soil moistening, by the size of the root system, and by the highest possible (for a given type of plant) transpiration. Dickinson et al. (1986) singled out 18 types of underlying surface differing by the size of the root system as well as by the physiological, aerodynamic and radiative properties of vegetation. In turn soil is divided into twelve classes according to its mechanical structure, and into eight classes according to its colour. The main restrictions of the proposed models are related to determination of the average properties of plant associations and the effect of their interaction. Overcoming these restrictions is a matter for the future.

Since the time of its appearance, humanity has been continuously perfecting ways to destroy the vegetative cover. From the beginning this was by gathering roots and the fruits of wild plants, then when man had learned to make fire he deliberately provoked fires, and nowadays we see both the replacement of the natural vegetation by cultivated vegetation and the deterioration in the composition and structure of flora (the synanthropization of the vegetable kingdom). The first considerable changes in vegetative cover started as far back as the transition from hunting and fishing to cattle-breeding and agriculture. Since then these changes have continued with increasing rate. The scales of the changes can be judged from the following figures. According to Sagan et al. (1979), since the first appearance of human beings the area of the deserts has increased by 9 x 106 km2; on the other hand the area of the forests has decreased by 8 x 106 km2 at temperate latitudes and by 7 x 106 km2 at tropical latitudes; urbanization has resulted in the total destruction of the vegetation of an area of about 1 x 106 km2; the area of lands with excessive salt content has increased by 0.6 x 106 km2. Thus, anthropogenic changes have affected a total of 26 x 106 km2, or 17%, of the land area.

The most intense destruction of forest occurs in the tropics. According to data from Woodwell et al. (1983) the rate of decrease in the area of the tropical forests is from 0.16 x 106 to 0.19 x 106 km2 per year. The new estimates based on the FAO 1988 and 1990, Tropical Forest Assessments show that tropical deforestation increased from 0.132 x 106 km2 in 1980 to 0.193 x 106 km2 in 1990. If such deforestation continues after 1990, with the rate depending on growth of population in the countries of the tropical belt, then according to IPCC (1992), from 73% (14.47 x 106 km2) to 91% (16.86 x 106 km2) of tropical forests will be cleared by the end of the next century. All this points to the fact that, at present, anthropogenic changes in the vegetative cover have acquired a global character. What climatic consequences can this lead to? To answer this question we examine two extreme situations: the destruction of forest tracts and the total destruction of the vegetative cover (see Kagan et al., 1990).

Destruction of forest tracts. This is accompanied by changes in the albedo of the underlying surface, soil moisture and the concentration of C02 in the atmosphere. In their turn, these changes depend on the form of secondary vegetation and the measures undertaken for soil humus conservation. We consider the following scenarios of the economic activities of mankind: (1) wood is burned, vegetation on the cleared territories is not recovered, soil humus is not preserved; (2) wood is burned, vegetation on the cleared territories is recovered, soil humus is not preserved; (3) wood is burned, cleared lands are used for agriculture, appropriate measures are undertaken for soil humus conservation.

We define the change Scts in the albedo of the underlying surface owing to the destruction of forests as (anf — af)sf/s where anf is the albedo of cleared territory; af and s{ are the albedo and area of forests; s is the land area. According to Whittaker and Likens (1975) the total area of forests in 1950 amounted to 48.5 x 106 km2, of which 24.5 x 106 km2 formed a share of tropical forests (tropical jungles and seasonal tropical forests); 5.0 x 106 km2 of evergreen subtropical forests; 7.0 x 106 km2 of seasonal forests at temperate latitudes, and 12 x 106 km2 of boreal forests. When assuming that the albedo of tropical jungles and boreal forests is equal to 0.10, and that of seasonal tropical forests, evergreen subtropical forests and seasonal forests at temperate latitudes is equal to 0.15, then the area-weighted average value of af is equal to 0.12. Assuming now that the albedo of cleared lands, meadow vegetation and cultivated areas is equal to 0.14, 0.20 and 0.25, respectively, we obtain Sas « 0.01, 0.03 and 0.04. We note for comparison that the maximum albedo change over the last 25-30 years caused by the effects of desertification, salinization, temperate and tropical deforestation and urbanization is between 3.3 x 10~4 and 6.4 x 10~4 (Henderson-Sellers and Gornitz, 1984).

Further, according to data from Whittaker and Likens (1975), the biomass density (in units of carbon) in tropical jungles amounts to 20.25 kgC/m2, in seasonal tropical and evergreen subtropical forests it amounts to 13.5 kgC/m2, and in boreal forests it amounts to 9.0 kgC/m2. Therefore, having available information on the area of forests (see above), we can speculate that, when burning all the forests, about 0.744 TtC (1 Tt = 1012 t) is transferred into the atmosphere. Destruction of forests is also accompanied by enhancement of soil breathing. Assuming that the soil biomass density in tropical jungles is equal to 13.0 kgC/m2, that in seasonal tropical and evergreen subtropical forests is equal to 7.0 kgC/m2, that in seasonal forests of temperate latitudes is equal to 10.7 kgC/m2 and that in boreal forests is equal to 20.6 kgC/m2, then the total carbon content in the humus of forest soils amounts to 0.631 TtC. If soil breathing results in a 15% reduction in the carbon content of humus we obtain an additional gain of carbon in the atmosphere that will be equal to 0.095 TtC. By this measure the total input of carbon into the atmosphere amounts to 0.839 TtC.

An excess amount of carbon in the atmosphere has to be partly absorbed by the ocean and partly assimilated by the terrestrial biota. We will suppose that the production of primary vegetation on lands not occupied by forest is limited, not by the C02 content in the atmosphere, but by other factors (for example, by the flux of short-wave solar radiation and soil moisture content). In this case the sink of excessive amounts of atmospheric carbon has to be determined by its absorption by the ocean and secondary vegetation on cleared lands. We assume that about half of the C02 excess is absorbed by the ocean and that the sink of atmospheric carbon, owing to its assimilation by secondary vegetation, can be evaluated by the biomass density of grassland and cultivated lands which, according to Whittaker and Likens (1975), is equal to 0.72 and 0.45 kgC/m2. As a result we arrive at the conclusion that the transformation of cleared lands into grasslands and cultivated lands entails the removal from the atmosphere of 0.035 and 0.022 TtC, respectively, and that the total increase in the concentration of atmospheric C02 for the three scenarios listed above is equal to 196, 188 and 169 ppm respectively.

The results of calculations for different scenarios of human economic activity obtained within the framework of the 0.5-dimensional seasonal thermodynamic model (see Section 5.5) point to the fact that when burning forest and failing to renew vegetation (the first scenario) the increase in the concentration of atmospheric C02 dominates, and because of this the changes in the climatic characteristics coincide qualitatively with those caused by doubling of the atmospheric C02 concentration whereas they differ quantitatively by several tens of percent. The effect of an increase in albedo manifests itself only in a small decrease in the land surface temperature in the southern box. The transformation of forest tracts into grasslands and cultivated lands (the second and third scenarios) leads to opposite changes in climatic characteristics, due to the domination of the effect of the land surface albedo increase over the effect of increased concentration of atmospheric C02.

Total destruction of the vegetative cover. The series of numerical experiments described above leads to a somewhat unexpected result: destruction of forest tracts entails an increase in soil moisture content and a decrease in run-off. In reality, the destruction of vegetative cover (say, when transforming forests into savannas or savannas into deserts) is accompanied by a decrease in soil moisture content. Obviously, the result obtained is a consequence of the constancy of the hydrological cycle parameters (soil moisture capacity W0 and the fraction y of precipitation spent on the formation of run-off), which, generally speaking, depend on the type and state of the vegetative cover. Since such dependencies for the Earth as a whole are unknown there is only one thing to do: to examine an extreme situation where the usual landscape is replaced by a surface devoid of any vegetation at all and reminiscent, by its properties, of dry grey soil.

Let the albedo of this surface be equal to 0.25 and the moisture be zero. We also take into account that in the absence of terrestrial vegetation the carbon in it and in the soil humus has to be redistributed between the atmosphere and the ocean. According to Whittaker and Likens (1975), the carbon content in terrestrial vegetation is 0.828 TtC, in soil humus it is 2.9 TtC, and from the 2.9 TtC contained in soil humus only 15% of 2.0 TtC (the carbon amount in the labile part of soil humus which takes part in exchange with the atmosphere) is added to the atmosphere. Because of this, the net incoming carbon in the ocean-atmosphere system is equal to 1.128 TtC. Assuming, as before, that this is redistributed between the ocean and the atmosphere equally, we have that the concentration of atmospheric C02 has to increase by 264 ppm compared with to its present-day value. Thus, the limiting values of the albedo of the land surface, the soil moisture content and the concentration of atmospheric C02 are known.

It is clear that the climatic consequences of the total destruction of the vegetative cover can be established only in the case where all three listed factors are considered, not separately, but in combination. Their combined effect demonstrates results close to those obtained in an experiment with zero soil moisture content. This indicates that the first two factors, in the sense of their impact on the separate characteristics of the climatic system, compete between themselves and, for the majority of them, practically compensate each other. The only exception is the annual mean temperature of the land surface: the predominant effect of increased albedo leads to a decrease in the land surface temperature (compared with its value corresponding to the zero soil moisture content), followed by a decrease in the surface air temperature.

Overall, the total destruction of the Earth's vegetation causes an increase in planetary albedo, a decrease in radiative cooling of the atmosphere and in the net radiative flux at the underlying surface, attenuation of the meridional sensible and latent heat transport in the atmosphere and ocean, a fall in the mass-weighted average air temperature and a rise in the surface air temperature (as a whole for the hemisphere), a decrease in evaporation and precipitation at temperate and low latitudes and their increase at high latitudes, an increase in run-off, a displacement of the boundary between the boxes to the north, a rise in temperature in the area of cold deep water formation, the disappearance of sea ice, a rise in temperature in the deep ocean at temperate and low latitudes, weakening of upwelling, an increase in the thickness of the UML, and, finally, a decrease in temperature in the upwelling area. As this takes place, the amplitudes of the seasonal oscillations of the mass-weighted average temperature and humidity of the atmosphere, as well as precipitation and evaporation, decrease, while the amplitude of the seasonal oscillations of the surface air temperature increases at temperate and low latitudes and decreases sharply at high latitudes.

In summary, we present some estimates of the equilibrium response of the climatic system to anthropogenic changes in the vegetative cover. According to Sagan et al. (1979) the changes in the land surface albedo related to the already-mentioned increase in desert area, the destruction of forests at tropical and temperate latitudes, the salinization of soil and urbanization result in a global fall in the surface air temperature by 1 °C. According to estimates by Potter et al. (1981), an increase in the land surface albedo at the expense of continuing (during the last millennia) desertification and destruction of tropical forests leads to a global reduction in the surface air temperature and precipitation by 0.2 °C and 1.6% respectively (in the Northern Hemisphere the surface air temperature decreases by 0.6 °C). Finally, most estimates characterizing an increase in global average values of the surface air temperature and precipitation due to doubling of the concentration of atmospheric C02 are found in intervals from 1.2 to 4.5 °C, and from 0.06 to 0.20 mm/day. Thus, despite different initial prerequisites the available estimates of consequences of anthropogenic changes in vegetative cover are qualitatively similar.

0 0

Post a comment