SURFACE PROCESSES OVER FORESTED AND CULTIVATED AREA IN WEST AFRICA THEIR IMPORTANCE AS COMPONENT OF

CLIMATIC MODELS . :/

B.k MONTEM Laboratoire de Bioclimatologie, INRA 78850 Thivemal-Grignon, France

ABSTRACT

The tropical rain forest contributes largely to a climatic equilibrium which is maintained by the general atmospheric circulation. However, due' to the population pressure, the current evolu- tion of the forest area in West Africa requires that an agricultural development should be stimulated in accordance with the characteristics of this zone, which urged us to contribute to this comprehensive programme.

Micrometeorological data from a rubber tree plantation in the Ivory Coast were used to define a simple, empirical model of its evapotranspiration, expressed in terms of equilibrium evapora- tion, available soil water, precipitation and incident solar energy. This model is assumed to be representative of the regional evaporation. When incorporated into a hydrological model, it is found to give a satisfactory description of the watet balance of a catchment covered with undis- turbed natual forest, some distance away. On this basis, the average energy balance given by the model is assumed to be broadly representative and used to show that the water vapour content of the southwesterly trade wind does not decrease when passing over the large-scale forest, in spite of the heavy rainfall.

The large-scale introduction of annual crops would alter the regional energy equilibrium. In West Africa, the forest transfers to the atmosphere as water vapour the equivalent of 75 to 90% of available energy. Annual crops such as rice, corn and cassava transfer only 45 to 55% of the net radiation. Modifications of the land surface characteristics in the forest area affect the hydrologi- cal balance, particularly the infiltration (thus the soil water storage) and overland flow com- ponents. The impact will be observed on areas situated north of th'e forest zone in relation to the Atmospheric Circulation particularly the Intertropical Convergence Zone motion (ICZ).

!INTRODUTION

L Macroscale hydrological modelling of land surface processes provides climatic models with inputs. Recent improvements in the modelling of large scale climatic dynamics (Walker et al. 1977; Shukla et al. 1982; Eagleson, 1982, 1986; Manabe, 1982; Rind, 1984; Dickinson, 1983, 1986, 1988; Sellers et al. 1986; Wilson et aL 1987; Druyan, 1987) show the importance of the earth sur- face characteristics on atmospheric conditions.

Regional anomalies in the surface characteristics such as albedo (reflection), temperature and wetness have local and sometimes far-reaching effects upon the atmospheric temperature, water vapour concentration (humidity) and therefore on rainfall.

171 - -

prec i p i t a t i on

I a i r temperature I

solar energy E 6T: /ii\ sensible heat ' I I I I

reflection I I I .. .... :.:..::':..: :.:i::.: I surface temDerature:::'

.. . .

l a ten t .heat d .. . .. . ,

Main land surface processes

In order to couple the effect of physical atmospheric parameters on the earth surface with the feed-back effect of the forcing function of the surface state, the large grid area of a General Atmospheric Circulation model (GCM) must be subdivided into square subgrid units. It would be easier to take account of the area variability of the vegetation, of hydrological and climatic conditions (Abbot et al. 1986; Becker et al. 1987; Girard et al. 1981, 1984; Klemes, 1983). Each subunit'is modelled separately but interacts with one another inside the global grid network

Hydrological studies in the tropical forest ecosystems have been conducted by ORSTOM for more than three decades (Girard, 1976; Roche, 1982; Casenave et al. 1982; Dubreuil 1985, 1986). Instrumented catchments gave water input-output budgets, to determine the average evaporation loss as opposed to short term evapotranspiration estimates. However, until recently, few quantita- tive data about tropical forest evapotranspiration and interception losses were available (Hutte1 1975; Pinker et al. 1980; Calder et al. 1986; Lloyd et al. 1988a, b, Shuttleworth 1988).

rWater budget methods associated with the energy budget, allow to better understand the importance of the quantitative water vapour exchange processes in different sites such as forests surface : Pinker et al. 1980; Shuttleworth et al. 1984a; Calder et al. 1986; Monteny et al. 1987; Shuttleworth 1988 or crops : Lhomme et al. 1982; Perrier 1982; Monteny et al. 1976, 1981, 1984, 1985.

This study aims first at modelling the evapotranspiration rate based on plant physiological activities, climatic and soil properties. This information has alFeady been inferred from previous studies on Hevea plantations and on rich and cassava field crops using the energy budget-Bowen ratio method (Lhomme et al. 1982; Monteny et al. 1985,1986b; Monteny, 1987).

Secondly, it aims at modelling the hydrological budget of a subgrid square, corresponding to a catchment area completely covered by tropical forest or by annual covers. The structure of the model is based on physical equations of land surface processes and on empirical equations derived from experimental research concerning some of the hydrological balance components.

Simulation of the catchment hydrological budget with natural forest or with annual crops after deforestation will give a general view of the importance of the modifications in surface characteristics on water flows and therefore on the atmospheric parameters.

172

It

METHODS

Synoptic situation of the forest zone in West Africa ' In the Ivory Coast, the tropical rain forest, situated along the gulf of Guinea, is influenced by

oceanic effects. Seasonal changes in climate can be considerable due to the movement of the 'Meteorological Equator (M.E.). The motion of the ME over the forest zone determines dry and rainy seasons (Wauthy, 1983; Leroux, 1983; Collinet et al. 1984). The tropical forest is subjected to' "a 1-3 month dry season with a few weeks when the northern wind is blowing intermittently (the Harmattan). This continental air reduces drastically the humidity, temperature and visibility. The wet season is associated with the tropical rain belt The rain characteristics are associated with the structure of the ME (Leroux, 1988) : O in front of the ME, the Inter Tropical Front (ITF) separates two air masses differing in tem-

perature and humidity. This structure, with high precipitated water potential, gives frequent , rainstorms from squall lines (March-May and November-December);

O behind the ITF, the Inter-Tropical Convergence Zone (ITC) corresponds to the tropical rain belt with moisture laden southwesterly winds and is responsible for continuous and abun- dant rainfall (i.e. rainy season).

L

Energy balance method Micrometeorological measurements allow to determine the mean water vapour transfer from

the vegetated surface to the atmosphere. The evapotranspiration rate has been measured by the energy budget-Bowen ratio method expressed as :

Rn= y E + H + G + P f & (W.m-20rKIm-2d-') ' (1)

where Rn : net radiation; yE : latent heat flux density, H : s5hsible heat flux density, G : soil heat flux density; IJ : i energy absorbed for photosynthesis; AS : change in energy sto+ in the air and the biomass between the levels of

measurement of G and Rn. Generally, most of the radiant energy absorbed by the canopy is converted into sensible and

latent heat flux densities on a daily basis. Rn is positive when there is an energy gain by the canopy; E and H are positive when there is an enegg loss by the surface. Assuming that the fluxes are constant with height above the canopy, the introduction of the Bowen ratio (ratio of the sensible heat flux to the latent heat flux : B) in a simplified form of equation (1) gives the instantaneous evapotranspiration rate of the surface :

U

E - (Rn - G) /y * (1 i- B) (2) Direct determination of evapotranspiration (micrometeorological techniques) is based on

point measurements taken in the atmosphere and soil at selected sites. Data analysis shows the control exerts by plants on water loss (leaf area and stomatal-root

resistances) which is also incorporated into the evapotranspiration rate models. Micrometeorological measurements were undertaken in a Heveu plantation (area 70 km2) at

Dabou, and on rice and cassava crops at Adiopodoum;, near Abidjan (5"19"4"13W). Watershed (area 38 km2) measurements were conducted on an tindisturbed forest at Tai, in the south western region (5"45'N-7"23W).

-

Hydrological measurements

Rainfall into a landscape gives rises to changes in the soil-vegetation-water continuum. Little work has been undertaken so far on the quantification of the hydrological processes in the humid forest regions of the tropical zone (Bernard, 1953; Pereira et al. 1960; Dubreuil, 1985,

1 73

.7

1986). Water budget studies give baseline data on which analyses of climatic characteristics depend.

The hydrological budg2t of a catchment area is given by : P + B = Q + O + Pi 4- E * ASW (mm per unit time) (3) .

where P : precipitation (mean for the catchment); B : head waters flowing into the subgrid area; in this case, B is not - considered; Q : drainage to water table which runs over as base flow; O : overland flow; Q+O = R : runoff; Pi : pre- cipitation intercepted by the canopy; E : actual evapotranspiration rate; ASW : variation of the volumetric soil moisture in the root zone.

Generally, for a subgrid area, watershed data give water input-output budgets, and the net watershed loss is due to evapotranspiration and interception rates.

tion in solar energy Rg and rain P inputs induces changes in soil moisture storage which affect evapotranspiration. Whereas the monthly water budget can be informative, shorter budget evaluations (5- 10 days) give a more detailed understanding of the hydrological processes.

Precipitation is the only water input in the studied catchment area A network of fifteen recording rain gauges gives sufficient accuracy for the estimated precipitation volume over the experimental catchment area (Casenave et al. 1980, 1984). The individual’ gauges give the data from which the isohyetal map can be drawn. They describe the spatial rainfall variability and give the average amount of rainfall over the catchment area. Because of the weakness”of the rain gauge network in the African forest zone and in order to minimize any error by extrapolating measured precipitation to a larger scale, these mean values are compared with those measured at a central meteorological station considered as representative of the subgrid square.

The stream discharge is measured at the outlet catchment stream equipped with a calibrated V-notch weir and a recording stage gauge monitors continuously the water height of the river. Rainfall of sufficient intensity and duration exceeding the soil surface infiltration capacity induces overland flow after the soil wa-tep holding capacity has been recharged. Streamflow Q and intermittent overland flow (after rain), R are deduced from the graph. (Casenave et a l 1984; Collinet et al. 1979). Streamflow and overland flow are converted to millimetre depth, giving unit area runoff values.

The changes in soil water storage (ASw) in the root zone of the forest vegetation is pre- dominantly seasonal. In southern Ivory Coast, the maximum storage occurs during May-June until late July and from mid September to November, depending mainly upon the ME motion which induces rainfall. The minimum soil water storage is measure$ generally during February- March, at the end of the dry season. The-difference between these two values of % volumetric soil moisture represents the maximum available moisture Swc stored in the root zone, rainfall equivalent of 220” for forest (3.5m deep root zone) and 125 mm for rice and grasses (1.0m root zone deep) (Talineau et al. 1974). The ratio of actual available soil moisture (Swd) to its maximum value represents the soil moisture wetness.

I Most of the hydrological balance components are measured on a daily time scale. The varia-

RESULTS AND DISCUSSION

Primary components of the land phase of the hydrological cycle are modelled: radiation budget, canopy rain interception, evapotranspiration, overland flow and soil moisture storage. Secondary details such as annual vegetation growth (roots and aerial leaf developments for annual crops) or water flow in unsaturated soil zone are not explicitly modelled at the present time.

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1

hdirtioa budget

In the tropics, net radiation measurements at the canopy interface are not usually available even though they are of primary interest, Rubber, oil palm and tropical rain forests have higher regression factors (a) than all other canopies in this area, mainly due to the b w reflectivity and weak thermal fluxes (Table 1). The long wave radiation balance has little influence on the net radiation flux due to the high atmospheric water vapour concentration (21g.m-3) in the coastal region. Fugure 1 presents the evolution of the net radiation values obtained over different canopies.

r

Fig. 1 :

lOOO-, ’ Radiation

_ _

global

Rn forest 9 8 0 0 E

6 0 0

4 0 0 Rn rice C O - c a 2 0 0 Rn bare soil

K O 5

9 11 13 15 l7 nov-dec 1979 local hours

Daily evolution of the available radiation amount over different land surfaces : rubber forest, rice crop and dry bare SoiL

The following relations were obtained between netiradiation Rn and incident solar radiation , - . - .Rg expressed in W.m-2 (Table 1).

TABLE-I Forest albedo and rehtionsbip between net radiation Ra m d global d a r radiation Rg for different tropical forests

albedo: R n = a R g + b (5) (W.m-2)

a b r

Forest rubber forest (humid air) 14 0.72 - 0.8 0.98

oil palm forest (humid air) 13 0.7 1 - 12 0.98

r

rubber forest (dry air) 14 0.78 - 50 0.98

Fldd crop rice 19 ! 0.68 - 5 0.99 CaSSaVa 18 0.64 - 9 0.99

fallow field 21 0.58 - 36 O.% !So$ surfmce

bumtout aera 9 0.54 4 2 0.97 barè soil : dry 23 0.64 - 42 0.98

wet 12 0.73 - 5 0.97

The slope u of the relationship reflects the short wave radiation budget of tall vegetation as forests (oil palm and rubber) than for the other vegetated area mainly due to the albedo values. b represents the long wave radiation balance which is high in the case of short and low vegetated

1 75

surfaces due to the surface temperature. This is no more the case when the M.E. is south Ofthe equator when dry air is blowing over the forest area : long wave radiation budget (b) increases modifying the slope of the relation.

Following some previous measurements (Monteny et al. 1979, 1981; Lhomme et al. 1982) net radiation can however, be derived from the total solar radiation Rg or sunshine duration which are more readily available.

The yearly evolution of the net radiation must be known more precisely in relation to land surface characteristics which affect the available energy used in 'different processes.

Figure 2 summarizes the evolution of the ava'ilable radiation amount over two different vegetation surfaces :

0 . 8 0 1 f

m,wL::V:wdm*3?% - Sx-RGssSs 0 . 5 0 -

* dry soil covered + wet soil - burn + wet dried crops

0 . 4 0 1 1 1 l l l l l l l l I 1

J F M A M J J A S O N D

Ng- 2 : Variation of the available radiation amount (Rn/Rg) during the year over forest CO) and cultivated CmPS (A: rice). I

0

0

for the forest ecosystem, slight variationsagreobserved throughout the year since it is ever green; for the cultivated or grass area, surface modifications such as the growing crop period, bare soil, dry vegetation, bumtout so il... are to be taken into account because they affect the reflection (albedo) and the emitted long wave radiation (EUT:).

. -

4

Water budget

Raiqfall distribution over the catcbment area w

The amount of rainwater within the boundaries of the catchment must l?e known with accuracy because rainfall is highly variable in time and space. The uncertainty in the precipita- tion distribution seems to be a limiting factor in the successful development of hydrological mod- els (Hromadka, 1988). Spatial resolution of rainfall has a dominant influence on the reliability of computed runoff (Troutman, 1983). Table 2 gives some results concerning the observed variance of rainfall over a catchment area of 38 km2 measured with 15 recording rain gauges.

TABLE-2 Variance ¡o &e individual event of rainfall over &e catchmeat are8

rainfall at the station mean value mm mm

variance standard deviation mm mm

16.4 37.6 58.2

18.9 42.8 63.2

18.8 21.8 18.9

f 4.4 f 4.1 f 4.5

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Spatial variability of precipitation introduces an error in one of the most important forcing function of all models. It varies from a general space sampling errors of 30% for a mean pre- cipitation input over the catchment area to 10% for more higher but exceptional rainfalls.

The general assumption stating that precipitation is equally distributed in a subgrid is not necessarily confirmed. It is one of the main parameters that reduce the accuracy of any hyd-

. rologica1 model. In another way, the representativeness of observed precipitation values at a single point such

- as the meteorological station and the averaged values from a small number of points in the catchment must be compared Figure 4 illustrates this relationship from July 1979 to June 1980.

While there is some daily variation, a fairly good correlation exists which allows to use the precipitation data of the meteorological station in order to determine rain input in the hydrologi- cal catchment model.

~

Taï 7.79-6.80 1 0 0 ,

n c c i c o m o c

v

- I Q c m K

r .-

L

8 0

6 0

4 0

2 0

O O 2 0 4 0 ; ' - 6 0 8 0 1 O0

Rainfall measured at the meteorological station (mm) i Fig. 3 : Relationship between daily precipitation values at the metemlogical station and the mean rainfall over the

catchment area (Tai, Ivory Coast). I I

a

Interception and stemflow components RainfalT intercepted by a forest canopy is considered to evaporate during or immediately after

the rain storm. The amount of intercepted water is calculated as the difference between gross pre- cipitation (measured at the top of the canopy) and the through-fall reaching the ground. Because only rainfall totals are radily available, a simple relationship is obtained.

In the forest area, through-fall was measured from 39 linear rain gauges of 0,05m2, arranged in 3 lines of 110m each (direction ; N-E, S-W and S-N) on the forest floor close to the gross rain- fall gauge site. It was possible to take readings M i e a day (at seven in the morning and one hour after rainfalls during the day). The calculated standard deviation of the through fall represents

I

I ' I

24% of the gross precipitation (Cardon, 1979). I

Intercepted rainwater by a dry forest canopy and the daily indiddual gross precipitation gives

Pi = O. 773* In (1 i- p ) + 0.138 (mdday) r2 = 0.86 (4)

I the following regression equation :

I

- 1 The maximum canopy (leaf and branches) storage capacity is evaluated at 2.8-3.0 mm (Car- don, 1979). The intercepted quantity depends on the precipitation characteristics and on the verti- cal leaf distribution, leaf area index (LAI) varying from 6 to 8 (Alexandre 1981). The variability of

177

I_-

the intercepted amount is enhanced by the canopy movement due to the wind during rainstorms. The evaporatiod resulting from this canopy water storage is assumed to &cur at the Penman potential evaporation rate which is 25-30% higher than the actual rate for wellwatered canopy in those climatic conditions. This is due to the forest.surface resistance IV which decreases from 80- 100s.m-1 for dry canopy to nearly zero with intercepted water.

Stemflow, expressed as a depth of rainfall over the canopy, is very small and neglected But if expressed as a volume of water, stemflow can represent a considerable input to the soil since it is accumulated at the foot of some tree species, depending upon their structure.

Stem water is collected by an adhesive plastic collar put on 16 adjacent trees covering an area of 300 m2-(Huttel, 1975). During two years (1970 and 1971) the water volumes in the containers were recorded and the following expression was obtained :

r2 = 0.92

~

-

stemflow: 0.00033* P'-68 (midday) The percentage of gross rainfall which reached the soil surface as stemflow represents 0,83%

of the annual precipitation (1800mm), compared with 1,8% for the Amazonian forest (Lloyd and Marques 1988a).

Rain interception by annual low crops or grass is difficult to evaluate, and is generally ignored. Sellers et al. (1981b) consider this fraction to be 30 to 65% of the forest canopy intercep- tion rates depending on their aerial structure.

*

Evapotranspiration

In order to know accurately the water vapour transfer from the forest to the atmosphere, the evapotranspiration rate needs to be evaluated using a biophysical formula which takes into account both climatic and biological reactions.

Priestley et al. (1972) proposed a formuia'to-calculate the regional evapotranspiration rate. It is based on large scale data rather than on' micrometeorological results : the amount of energy required for evaporation comes predominantly from net radiation Rn, and the effect of vapour pressure deficit is introduced by a coefficient C in the equation depending on the surface wetness :

(5a) The equilibrium evaporation rate Eo = (o/h+U* Rn as defined by Davies et al (1973) is more representative of the climatic demand in this humid atmosphere. The ratio between the measured value of the actual evapotranspiration rate of the rainforest (measured by micrometeorological techniques) and Eo gives the coefficien! C, also called crop coefficient (Katerji et al. 1983). Generally, in daytime conditions, C is taken as 1.26 for well watered crops in temperate regions (Priestley et al. 1972) corresponding to a value of B = 0.05. The soil heat flux G is generally neglected because it represents only 1 to 2% of Rn in the case of rain forests. The value of C results from different processes which control the water vapour transfer from the leaf surface to the atmosphere (Jarvis et al. 1976; Pemer 1980, 1985; McNaughton et al. 1983, 1988): aerodynamic ra and climatic rc resistances and mean canopy resistance N which depend on leaf physiological activities through different growing stages :

(5b)

E = C EO = C(UA+v* (Rn-Gj

C = / I 4- (Y/('A-kyI) ?W"]//l 4- (Y/A 4- 8) n/r a) From the micrometeorological measurements, figure 4 gives the diurnal variation of C for

04.28.82 : youg leaves (2 months old) without soil moisture limitation 04.15.81 : young leaves (2 months old) with a soil water deficit 12.22.81 : old leaves (10 months old) and well watered.

daytime hours, in relation to plant phenology and soil conditions :

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0 9 2 j .._

0 . 0 I I I I , I I

04.1 5.81

14.00 16.00 1 1 1 1 1 I I I ) , 1 1 1 1

’ 10.30 12.00 Hour

Fig 4 : Diurnal variation of the coefficient C for different plant and soil conditions

It shows that : O a well watered forest canopy with young leaves has the maximum transpiration rate (0.9 < C

< 1.1) O the soil water deficit affects the actual evapotranspiration rate of a well developed canopy

O

Figure 5 illustrates the relationship between C %nd the calculated canopy resistance values. The reduction of the evapotranspiration rate is maidy Mated to an increase of the canopy resis- tance TV (equation Sb), the climatic resistance chafiges only slightly from one day to another. C vanes from 0.3 to 1.1 depending on leaf physiological activities in relation to the soil moisture availability.

surface as noted by the reduction of the coefficient C to a mean value of 0.55; leaf ageing reduces C as shown during 12.22.81.

Higher values of C could be measured during-the rainy season due to the evaporation of- intercepted water. Without soil water limitations, C varies between 0.80 and 1.1 for dry canopy

Rg.5:

0 . 0 1 I 1

10 1 O0 1 O00 Canopy resistance (s.m-i) Relationship between the coefficient C and the forest canopy resistance (dry surface) for different soil water contents

1 79

conditions depending on leaf surface properties and atmospheric demand Forest actual e v a p transpiration rate under dry surface conditions is generally near the equilibrium evaporation rate Eo. Values of C are higher for vegetation with lower canopy resistances such as grass and agricultural crops (C - 1.26).

Because of small variations in the atmospheric water vapour deficit in this climatic zone, C is more characteristic of the stomatal regulation (plant physiological properties) in relation to water availability in the soil volume explored by the root system. Soil moisture depletion affects stoma- tal aperture (rv) and thus the canopy behaviour. It gives a water component relatively easy to evaluate.

Average weekly volumetric soil moisture content from a great deal of coresampling, pre- cipitation amounts and the fraction intercepted by the canopy were measured during the year 1983 to evaluate the actual evapotranspiration rates. The general response of the forest is given, using C, in figure 6. Soil moisture depletion affects the relative evaporation rate only when nearly 40% of the total soil moisture content Swc in the root zone has been extracted The point distribu- tion in the graph is mainly due to changes in the climatic parameters (variability of incoming solar energy and precipitation).

- 1 . o - 0 . 0 -

- -

0 . 6 - - 0 . 4 - - 0 . 2 -

0 . 0 -

1 . 2 -41 m

0.91 *(SWd/SWc) i ’’ -- C= 0.44 e

1 I l I I I 1 1 1 l I I 1 1 1 1 1 1 ~ 1

m m m

o

fig. 6 : Forest r e d v e evapotranspiration rate in relation to the fraction of depleted soil moisture by the root system. (depth : 3.5m).

‘5

The fact that the canopy behaviour depends on soil water availability contributes to an

For a well developed forest (LAI 2 4), the evapotranspiration rates were evaluated from the expression of forest evapotranspiration rate in the tropical humid zone.

relationship shown in figure 6 : w . = c = 0.44 wp (0.91 SWdSWC) (6a)

which gives : 10.44 eXp (o. glIsWd/sW)] (4/a+V

The influence of soil water depletion on canopy transpiration is incorporated in the model as an average value of the soil moisture availability over the whole catchment Because the transpiration rate in humid regions is mainly due to radiative energy, this process is based on the radiation budget as in the Priestley et aL (1972) formula. Little meteorological information is available for the forest regions and the Penman-Monteith combination method (Monteith, 1%5) must be used carefully at the catchment scale for estimating water vapour transfer to the atmosphere (McNaugthon et al. 1983; Mawdsley et al. 1985).

180

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Dunin et al. (1985% b) found the same kind of relation with an Eucalyptus forest as in figure 6. They show the effct of the leaf area index evolution and the variability of the climatic demand on coefficient C. In the case of the tropical rain forests, the leaf area index does not change con- siderably during the year, leaf fall or growth being not simulataneous for all species. C is thus not, influenced in the same way as for temperate deciduous forests (Singh et al. 1980; Sharma, 1984).

Our regional evaporation model is consistent .with the concept that soil moisture depletion has an influence on the canopy resistance which controls the water use.

Runoff

Streamflow (base flow + overland flow) is one of the most common runoff surface processes. Overland flow occurs when the hydraulic conductivity of the forest soil surface is saturated; The 'floor surface parameters depend on certain soil characteristics such as the multichannel mac- ropore system and the soil horizon moisture.

Daily precipitation versus overland flow of more than 1 mm equivalent rainfall at the weir discharge (38,000 m3.d-1) are plotted in figure 7. Data equivalent to less than lmm were discarded.

n

E E

3

Y

O

m t Q

I c

- t

401 rainy season

/ O

ö 20 4 0 =' 6 0 8 0 100

Precipitation (mm . day -1) f r

Fig. 7 :' Relationship between daily rainfall and overland flow for discharge. flow at the weir, higher &an the equivalent of lmm during dry and rainy seasons

A distinction is made between dry and humid climatic conditions, giving two curves in rela-

. Under dry conditions, the soil surface of a forested catchment plays a major role in the rapid tion to the soil water status which affects an overland flow volume.

absorption of stormflow due to the empty macropore system. The corresponding equation is :

Under wet conditions when precipitation or successive rainfalls of sullïcient intensity and duration exceed the soil surface infiltration capacity, overland flow is the main surface process This gives the following equation :

ovedandfrow : O = 0.85* (P"0nr.') r2 = 0.81 Va)

o = 1.012* (eaM3.P) r2 = 0.72 mo

181

I ,

'1 I '

When the soil surface macropores are near saturation such as is the case during the rainy season, precipitation discharge induces more overland flow. The previous soil moisture con- ditions and the rainfall intensity/duration are of critical importance on the water speed infiltra- tion (Collinet, 1983).

Modification of the landscape characteristics such as deforestation in the whole catchment area affects the hydrological cycle and particularly the overland flow component. Roche (1982) and Fritsch (1986) measure the overland flows after the removal of tropical forest over a catch- ment area and obtained the following equation :

.

.

'

Overland : O = 040* P-2.8 mm r2= 0.88 (7c) For a global evaluation, these equations have the advantage to integrate the soil permeability

differences from different sub-basins of the catchment area. On the basis that the catchment does not collect water from outside the topographic surface

boundaries, the forest water losses and the changes in the soil water storage can be estimated *

using the hydrological budget A daily water budget model is developed and simulates satisfac- tory evapotranspiration and interception losses by tall forests and low vegetations as well as runoff flows.

HYDROLOGICAL CATCHMENT MODEL

Water balance of the atmosphere over a region depends on : @ O

O

the regional transfer of water vapour from the soil-canopy to the atmosphere; the condensation-precipitation of this atmospheric water vapour to the soil-vegetation surface; 1 the budget between incoming-outgqidg :water vapour across the air volume boundaries above the region.

Only the first two points will be considered by the model to show the importance of the West African forest zone in maintaining climatic stability in this region.

The model presented here is used to simulate the forested or cultivated waterphed behaviour and to obtain estimates of the long-term evapotranspiration losses from the whole catchment sur- face. It also determines the range of soil-moisture deficits which could occur during the dry geason. But on a regional basis, difficulties arise in relation to the data needed to define the flows and storage levels for detailed modelling. Our hypothesis is that methods which are proven satis- factory on a catchment of 38 W are also gppropriate for describing processes over larger areas.

The daily transpiration rate, E, is estimated as a function of available energy. The fraction of water intercepted on the canopy evaporates at a potential rate, transpiration is assumed to be zero. The evaporation of intercepted rain is 25-309ó higher than the equilibrium evapotranspira- tion rate Eo on a daily basis. Therefore, during rainy days the total represents the actual evapo- transpiration rate. The catchment model has only a single soil storage compartment retaining a fraction of rain water. The excess water drains towards the watertable which runs over as base flow.

Some specific characteristics must be known such as the mean depth of the root zone of forest trees or of annual crops, the evolution of the leaf area index, soil hydraulic conductivity, water storage capacity, The different components of the model are as follows : O 0

'

-

'

interception is relaced to the amount of daily precipitation (equation 4); evapotranspiration is related to potential evaporation (a function of the net radiation) which is controlled by soil moisture wetness in the root zone (equation 6b);

1 82

!' ) i >

0 overland flow depends on the amount of daily precipitation (equation 7% b, or c); rainfalls Of sufficient intensity and duration when exceeding the soil surface infiltration capacity induce overland flows; soil water status of the catchment determines the volume of drainage flow generated after the rainfall.

From the data of the central meteorological station and the different equations based on ' averhge values from a representative regional catchment area, the hydrological components have

been deduced from data from August 1979 to July 1981. One of the most important temq the evapotranspiration rate, will be discussed here in more details.

Studies on the water catchment balance give an indirect measurement of the total aerial evapotranspiration rate. By subtracting catchment losses (stream and overland flows : R) from average catchment rainfall, P, the total amount of aerial evapotranspiration rate and the changes in soil water storage are estimated (E-kPi-kASw). The comparison between the estimated e v a p transpiration losses by simulation (E+Pi) and the results obtained from the hydrological balance measurements shows some discrepancy in relation to the inertia of the catchment flows (fig. 8) particularly during rainless period and after it when water is stored

0

2 5 0 , œ

C O

l8

.- c

i

v) C

- n

L c O P l8 * w

08.79 01.1980 . 07.80 01.1981 07.81

month .-year k 8 : amparison between measured (E+Pi+AS = o) and estiamted total evapotranspiration rates (E+Pi = a) of

th'e undisturbed forested catchment area during 2 years.

The water loss by the forested catchment, estimated by the hydrological equation; depends on l e frequency of moderate-heavy rainfall in this forest environment, During the rainy season, the hest evapotranspiration rate is higher than the equilibrium evaporation due not O&' to banspiration, the major souse of water loss, but also to a significant amount of intercepted rain *ch evaporates directly from the forest canopy. During the dry season, the evapotranspiration

depends on the available solar radiation. But, after a few weeks, the soil moisture wetness *Wd/SWc) affects the canopy resistance and hence the transpiration rate which decreases while *&e Same time the equilibrium evaporation increases.

Evapotranspiration depletes the soil moisture during the dry season, reaching sometimes the Point at the end ofithe season. The evapotranspiration difference rates from November to

kbruaV (double dotted line, fig. 8) corresponds to the contribution of the soil water content in 'e zone. At this time, the catchment baseflow is very low. During March-April, the higher rues of measured evapotranspiration result partially from the fill upaby infiltrated water of the

183

Storage is a combination of soil moisture changes and water table fluctuations. The vanations '

in soil moisture in the r&t zone are predominantly seasonal, depending on the rainfall distribu- - tion. The mean available soil water for the tree roots from the ground surface to 3.5m depth 220" rain equivaltnt But one must be careful : a lot of tropical forest trees do not have a deep , rooting system but a large subsurface one. Rain distribution over the year is an important factor in maintaining theecosystem equilibrium. Without soil moisture measurements, catchment water - balance is not precise enough to evaluate the forest evapotranspiration rate accurately d u h g cer- tain periods of the dry and the rainy seasons because of the variations in soil moisture storage ASw.

Generally changes in the watershed storage term are neglected on a yearly basis. In this case, the annual hydrological budget. simulation results agree with the experimental work (Morton, 1983; Holmes, 1984).

Precipitations in the tropical rain belt (ITC structure) of the ME are characterized by high intensity associated with large drops particularly during March to May *and September-October. Due to the rapid saturation of the surface soil, the infiltration rate is reduced and the rain watm flows on the forest floor. This mechanism is exacerbated by the change of the vegetated cover.

Figure 9 presents the results of the simulation of the hydrological catchment balance with two different land surface vegetation.

17

1 5

1 2

1 0

7

5

2 5 0

- Uf;,,,,,,,,,,,,,,,,l,,,,,,,,,,,,,,l,i

08.79 ' 01.1980 07.80 decade - year

Rg. 9 : - Accumulated values of the simulated total evaporation (E) and runoff rates (R) for an undistuhd forest (9 and an agricultural vegetation cover (c) for the same catchment area

- .si-

The evaporation rates of forested surfaces are generally higher than annual crops : e

O

tropical forests are never leafless and interception losses are high, depending on the dis- tribution of daily precipitation; a number of forest trees have deep rooting system which provide them with soil water du& most of the dry season; after removal of forest trees, soil is generally bare and overland flows are especially high due to the impacts of rain drops which give rise to a driving crust on the soil surface (Valentin, 1987); annual crops grow during 4-5 months with a varying leaf area index : the highest transpira- tion rates occur when soil is completely covered but unfortunately their interception iates are low.

- 1

O

184

t

On a yearly basis, the West Aftican tropical forest and crop surfaces or grassland inject res- pectively the equivalent of 60 to 75% and 35 to 50% of rainfall amounts as water vapour into the atmosphere if similar precipitation characteristics are found. It confirms the results obtained by Hutte1 (1975) at Banco, a site in the Ivory Coast forest zone. The same conclusions have been @ven for the Amazonian forest by Lettau et al. (1979); Salati et al. (1984) and Schuttleworth (1988) and by numerical simulation by Walker et al. (1977); Shukla et al. (1982); Dickinson et al. . (1988) and by Sellers et al. (1981% b) for low vegetated surfaces. This decrease in water vapour exchange is linked to the short vegetative growing period of main agricultural crops, to rapid soil water depletion in the shallow rooting depth and to heavy precipitation characteristics.

During the wet season, more than 85% of the net radiation input is transferred as latent heat to the atmosphere due to the rainfall frequency and the interception rate of canopies (fig 10). Calder et a l (1986) found 100% in the West Java forest and Shuttleworth (1988) found that it amounted to90% in the Amazonian forest The variation of the transfer amount is due mainly to the Precipitation di+ tribution at the different forested sites

- 2 5 0 - 6

E E

O F 2 0

,l 5

1 O0

5 0

O

O

O

...................... 08.79 01.1980 ,

decade - year

, o

. a

. 6

0 . 4

0 . 2

0 . 0 07.80

FTp. 10 : Evolution of the relative energy transfer as water va&ur to the atmosphere by the forest surface ( O ) and by annual crops (a) in relation to the rainfall distribution, at Tai, Ivory Coast,

5

With the reduction of soil moisture content, the availaeases : (E+Pi)/Rn = 0.50 for forest and up to 0.1 for annual crops. It increases the atmospheric temperature and therefore its water vapour pressure deficit

Under significant soil moisture stress, annual crops dry up (December-January).powerfd rain drops and to solar rays : the destruction of the cloud structure induces a watertight crust which dries out, affecting considerably the water infiltra tion. The decrease of water vapour exchange by annual vegetated covers is compensated by an increase of sensible heat transfer which affects the air temperature (fig. 1 I).

Figure 11 illustrates the daily evolution of dew point (Tr) and air temperature (Ta) over well watered and stressed rice crops. The reduction of the latent heat flux due to soil water depletion is counterbalanced by a sensible heat flux increase. The warming of leaves and soil surfaces give rise to an increase in the air temperatures disturbing the atmosphereic conditions. This process could be the beginning of the desertification of the forest zone as presented by Kandel (1984); Henderson-Sellers et al. 1984 and Gomitz 1985.

185

o Ta 28.03.83 Rg : 20 MJm-2j-1 0 Tr . ----

O O Ta 15.04.83 O Tr +sol humide O 0

o o O 3 2 O

n

o Y

3 0 W U

I I I I I I I I I 1 I 1 O 4 8 1 2 1 6 2 0 2 4

hours Fig. 11 : Evolution ofthe atmospheric parameters (dew point Tr and air temperature Ta) 2m above vegetated rice sur-

face before and after irrigation (incoming radiation = 20 MJ.m-2.d-1).

Simulation of the catchment hydrological budget shows that land use changes on a large scale in the tropical forest zones will have some drastic consequences on physical atmospheric parameters as such temperature and water vapour concentration.

The forested land acts as a water source for the air mass layer of the ITF within the Meteorological Equator. The depletion of tht atmospheric water vapour content by precipitation is reduced by the evapotranspiration rates,,\aith the mean monthly vapour pressure decreasing

The tropical rain forest accounts for a large turnover of the precipitated water back to the atmosphere during the shift of the Meteorological Equator. From April to November, the impor- tance of the latent heat transfer from forested surface to the air mass (monthly evapotranspira- tion rates between 95 and 150") depends mainly on the available solar energy. Especially in July and August, the ITC structure over the forest area consists mainly of different cloud stratifie- tiens which absorb and reflect a large amount of solar energy : only 25 to 35% of extraterrestrial radiation are transmitted to the forest canopy (Monteny et al. 1981; Lhomme et al, 1982). This climatic parameter is the main limiting factor which could restrict the atmospheric moisture s u p ply by the vegetated land Strong upwelling in June-July in the oceanic equatorial zone (Bakun, 1978; Merle, 1983; Hisard et al. 1986; Lamb et al. 1986) contributes to the condensatiori of this atmospheric water vapour forming large scale cloudiness stratocumulus, altostratus) Monteny 1986 a, b).

The large forested area induces steady &ate in the atmospheric characteristics : water vapour content and temperature which are determined by the available energy and the surface resistan- ce. These characteristics lead to an equilibrium value for the evapotranspiration rate (Peier , 1982). The forest zone tends to maintain itself in a moist state : water vapour pressure over the forested area from the coast to 4OOkm inside the continent does not decrease despite variations in monthly rainfall. Water vapour transport between the surface and 850mb (equivalent to 30" high) over West Africa shows a progressive penetration of moisture from the coast to the north (Cadet et al. 1987). The climate of the continental Amazonian forest has been demonstrated by Salati et al. (1984) to be partially self regulating : in turnover of water vapour between air mass and forest is rapid, maintaining a moist and cloudy overlying atmosphere. The residence time of

slowly.

,

186

1’

water vapour in the atmosphere may range from a few hours to a few weeks with an averige of five days (Penman, 1970).

Without soil water limitation, the forest actual evaporation is at the equilibrium rate, but dur- ing oceanic upwelling periods, the sea surface - atmosphere interaction affects the air vapour pressure and induces important modifications in the global atmospheric circulation (GAC), affecting some climatic parameters in this region such as rain distribution, incident radiation and air temperature.

CONCLUSIONS

Water cycle is one of the most basic of all the biogeochemical cycles. Regional anomalies in the surface state have local effects on the atmospheric parameters and affects the global circula- tion in the atmosphere over the continent. In tropical regions, forest zones play a considerable role of the hydrological cycle, thus on the surf-ace-energy budget

In order to know accurately the water vapour transfer from the forest or low vegetated cover to the atmosphere, the evapotranspiration rate needs to be evaluated using the biophysical formula which takes into account both climatic demand and biological reactions.

*

Estimated forest evapotranspiration rates, obtained from soil and micrometeorological relationships, give a general outline of the dynamics of water movement through catchment areas. Soil water content in the root zone is one factor which affects the actual evapotranspiration in the humid tropical zone. The evaluation of the different processes according to the soil- climatic conditions allows to be remodelled of the micrometeorological dynamics of the forest zone. The comparison of the evaporative water losses obtained using the hydrological budget or estimated with the hydrological catchment model $ows some differences due to the soil water depletion and to the inertia of the catchment flow with time.

Forest vegetation in the west African tropical rigion transfers the equivalent of 75-80% on an average of the yearly net radiation as latent heat It varies from 50% for the drier month of the year to 90% during the rainy season. This re-evaporation represents 60 to 75% of measured rain- fall. An experiment conducted on the low vegetated cover and focused on the land surface pro- cesses has shown that generally less energy is removéd by latent heat The increase in surface and . soil temperatures give rise to an increase in the air temperature and the water vapour deficit On a large-sple, this warming effect such as the reduction water vapour transfer modifies the atmospheric water cycle.

Large scale modification’s of the forest cover affect some surface characteristics disturbing the energy exchange equilibrium. The feedback effect could be a slight reduction in local precipita- tion and an important one in northern zones such as the Sahel, via the General Atmospheric Circulation.

If globally, the continental land surfaces are net sinks for atmospheric moisture picked up %over the ocean, the forest zone plays the same role as a large energy converter. It acts as an important water vapour source rder of the forest zone despite the importance of the precipitation amount. The main limiting factor is the available solar energy. Strong upwelling during June-July in the oceanic equatorial zone reduces the atmospheric instability and contributes to an impor- tant cloud cover, intercepting a large amount of solar energy.

The hydrological catchment model presented here could be easily applied to linkeGeneral Circulation Models with regional catchment behaviour. Small scale hydrological studies can be applied on large scale (Eagleson, 1986). The climatic impact assessment could be devaluated from any local modification in land surfaces due mainly to human activities.

187

t

ACKNOWLEDGEMENTS P

We thank Mrs Cavalozzi for checking the text I

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'

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, T?~?UL& JC et

, International Atomic Energy Agency.

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PROCEEDINGS OF THE INDO-FRENCH WORKSHOP

ON

Mukhteswar, March 1 4 , 1989

ORGANISED BY

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PROMOTION OF ADVANCED RESEARCH

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Jointly with

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