impact of global warming on evapotranspiration

Northcott, G. P. (1992) A re-assessment of the highest temperatures during July 1959. Meteorol. Mag., 120, pp. 230-232

Ratcliffe, R .A. S. (1976) The hot spell of late June- early July 1976. Weather, 31, pp. 355-357

Rowe, M. W. (1990) What is Bntain's highest temperature? Climatol. Obs. Link Bull., 243, p. 19

Shaw, M. S. (1977) The exceptional heatwave of 23 June to 8 July 1976. Meteorol. Mag. , 106, pp. 329- 346

Tinn, A. B. (1938) Local temperature variations in

the Nottingham district. Q. J. R. Meteorol. SOC., 64, pp. 391-405

Webb, J. D. C. (1983) Further remarks on the July 1868 maxima at Tonbridge, Kent. J. Meteorol. UK, 8, pp. 272-273 - (1984) An extraordinary heatwave in Shetland.

J. Meteorol. UK, 9 , pp. 211-213 Webb, J. D. C. and Meaden, G. T. (1991) Britain's

highest temperature for every date of the year, 1 January to 31 December; 1991 amendment. J. Meteorol. UK, 16, pp. 73-75

Impact of global warming on evapotranspiration

J. G. Lockwood University of Leeds

The United Kingdom Climate Change Impacts Review Group (1991) comment that there are some important parameters which help to give the United Kingdom its characteristic climate and which will not change with global warming. As an example, it is sometimes incorrectly inferred that since, due to global warming, average temperatures in the UK may approach those of, say, Spain or Italy, the UK can look forward to a Mediterranean climate of long, hot, dry summers with high evaporation rates, and mild winters. This may not be so because the geometrical path of the Sun across the sky is different in the UK from that in Spain, and it is unlikely that the UK will be dominated in summer by the subtropical high pressure systems located under the subtropical westerly jet stream that give the Mediterranean its dry summer. Similarly, the notion that UK evapotranspiration rates will necessarily increase with global warming needs further investigation.

Energy is required for evaporation, therefore the evaporative loss from a uniform land surface can be expressed (Oke 1978) in terms of the following energy balance equation:

Q , = A E + H

where Q, is the available energy (the difference between incoming and outgoing radiation and what is absorbed by the ground) at the effective evaporating surface which is the vegetated land surface, AE is the latent heat flux (energy released to the air by evaporation), E is the evaporation rate, A is the latent heat of vapo- risation of water, and H is the sensible heat flux (conduction to the air).

Similarly, the water balance of the land surface (down to a depth where vertical moisture movement is absent) can be expressed as:

p = E + Arm + As,

where p is the precipitation, A,, is the net runoff, and A,, is the change in soil moisture storage.

Equations describing evapotranspiration from vegetation are normally also expressed in terms of conceptual resistances (Dickinson 1992), using Ohm's law in electricity as a direct analogue. Ohm's law gives the relationship between current (amps) in a circuit to the elec- trical potential (volts) and the resistance of the wire:

Current = (Potential difference) /(Wire resistance).

29 1

~~ ~

For entities such as water vapour, heat and carbon dioxide this may be rewritten to read:

Flux rate = (Concentration difference of property) i (Resistance to flow exerted by the system, r)

where r represents the appropriate system resistance in seconds per metre (sm ’).

In the case of transpiration, one resistance applies to the passage of water through the vegetation to the atmosphere, the other resistance to the passage of water vapour from the effective canopy evaporating surface to a standard level (often at 2m) in the overlying atmosphere. The value of the first (the bulk canopy resistance, r,) depends on such factors as the vegetation type, soil moisture deficit, sunlight, atmospheric humidity, etc. The value of the second (the aerodynamic resistance, r,) depends on such factors as vegetation height, wind speed, etc. It is therefore also useful to consider the following equation (McNaughton and Jarvis 1983) which is a rearrangement of the standard Penman-Monteith equation:

where

and terms not previously defined are: 6e is the vapour pressure deficit (saturation vapour pressure minus actual vapour pressure at mean air temperature), C, is the specific heat capacity of air at constant pressure, s is the slope of curve of saturated vapour pressure versus temperature at mean air temperature, is the psychrometric constant, and p is the air density.

The first term in the last equation, QE,, is the rate of evaporation that would obtain if the heat budget of a surface were dominated by the diabatic (radiative) term. This condition tends to be satisfied when short, well watered vegetation (such as grass pasture) is exposed to bright sunshine, humid air and a light wind. The evaporation rate is then effectively independent of the saturation deficit of the atmosphere, be, but strongly dependent on the

available energy, (2,. Under these conditions the surface is often described as being ‘decoupled’ from the prevailing atmospheric conditions and the value of the weighting factor, Q, is nearly 1. The term ‘decoupled’ should not be taken too literally, because Q, depends on the interception of radiation by the vegetation, and radiative conditions depend on both astronomical factors and cloud distribution, that is the prevailing weather.

In contrast, the second complementary adia- batic term, (1 - Q)AE,, is the rate of evaporation ‘imposed’ by the environment when the surface is ‘fully coupled’ to the prevailing weather. This term dominates if the value of Q approaches zero. These conditions are satisfied when r, is large, such as over a forest or under drought conditions.

It is seen, therefore, that vegetation canopies may exhibit at least two responses to climatic change, dependent on whether they are coupled to atmospheric conditions or not. Grass pasture can be taken as an example of a decoupled canopy while coniferous forest is an example of a coupled canopy. The responses of the hydrology of these two sample canopies to small temperature changes are illustrated using a very simple model of the energy and water fluxes in the soil-vegetation-atmosphere system.

The simple soil-vegetation-atmosphere model

Changes in surface vegetation, perhaps due to climatic change, cause variations in the surface evaporation rate for a number of biophysical reasons (Dickinson 1992; Sellers 1992). The stomata1 properties, and therefore bulk canopy resistances, vary between vegetation types. Vegetation changes may alter the surface albedo and therefore the amount of energy absorbed, and also the surface temperature and therefore the amount of long-wave energy emitted. These radiation changes influence the amount of energy available for evaporation and therefore the total evaporative loss. Alternative vegetation covers also generate different amounts of turbulence and differing humidity deficits in the atmosphere. It is therefore necessary to incorporate into the model a very simple

292

description of the surface layers of the atmosphere to allow some adjustment for these factors.

The layer of the atmosphere between the Earth's surface and approximately 2km, in which atmospheric conditions are dominated by surface processes, is known as the planetary boundary layer (Kiehl 1992). The planetary boundary layer is further delineated into a number of sublayers, the lowest of which is the surface layer (typically the lowest 100 to 200m), defined by conditions of near-constant fluxes of heat, momentum and moisture with height. Complex models of the planetary boundary layer and its sublayers exist (e .g . Garratt 1992) and may be used in association with vegetation models (McNaughton and Spriggs 1986). The computer model used in this research is based on the bridging technique suggested by McNaughton and Jarvis (1983) to convert meteorological observations over one vegetation cover to those appropriate to a differerent vegetation cover (e .g . from grassland to forest). This allows a relatively simple computer model of soil-vegetation-atmosphere interactions to be developed.

In the simple model used here, adjustment was made for surface conditions by generating a dataset for the lOOm level in the lower atmosphere that was independent of the immediate vegetation cover. Standard profile techniques allowing for atmospheric stability (e.g. see Mon- teith and Unsworth 1990) were used to generate hourly values of dry- and wet-bulb temperatures and wind speeds at 100 m in the mixed layer of the atmosphere from a complete surface energy and water balance derived from hourly observations taken by an automatic weather station (AWS) (Strangeways and Smith 1985). An iterative scheme is used to estimate the surface temperature and hence the surface long- wave radiative emission. Since the incoming short-wave radiation is measured by the AWS and the surface albedo has been previously determined, it is possible to estimate the surface independent downwelling long-wave radiative flux from atmosphere and clouds. Thus it is possible to generate a set of meteorological input data which is independent of the immediate surface vegetation cover. These data then form the basis of the climatic change studies

and are applied to alternative hypothetical vegetation covers considered to be superimposed on the standard site.

Since the AWS was sited on an upland moor, two contrasting vegetation covers are considered in the form of moorland grass pasture and coniferous forest. For both vegetation types, iterative and standard profile techniques are used to obtain new hourly estimates of the surface net radiation and then a complete energy and water balance. Brief descriptions of the vegetation covers are as follows.

Grass pasture

The description used is largely based on the work of Thompson et al. (198 1) and Lockwood et al. (1989). The amount of foliage in a vegetation canopy is conventionally specified as a leaf area index, that is the area of leaves per unit area of ground taking one side of each leaf into account. For the months May to August the leaf area index was set at 5, the canopy height at 0.15m and the albedo at 25 per cent. Minimum bulk canopy resistance was set as follows: May 50sm-'; June, July 60smp1; August 70sm-'. For this moorland grass pasture, when the soil moisture deficit (the amount of rainfall needed to saturate the soil) increased above 28 mm the bulk canopy resistance increased in an exponential manner until the transpiration rate became zero with a soil moisture deficit of 70 mm. For grass, bulk canopy resistance is assumed to be independent of atmospheric humidity deficit and temperature, but transpiration is restricted to daylight.

Coniferous forest

Unlike grass, many plants show a strong stomatal response to atmospheric conditions. This bulk canopy resistance scheme is meant to be typical of such a response, as shown, for example, by a coniferous forest (Thompson et al. 1981). Leaf area index is set at 6, canopy height at 10 m and albedo at 12 per cent. Bulk canopy resistance is a function of both temperature and atmospheric vapour pressure deficit as shown in Fig. 1. When the soil moisture deficit exceeds 70 mm, bulk canopy resistance is increased exponentially until the soil

293

30 -

Atmospheric vapour pressure deficit (mbar)

Fig. 1 vapour pressure deficit. Soil moisture ts assumed to be non-limrttng.

Values of bulk canopy reststance ( sm ‘)for a conzferous forest as determined by air temperature and atmosphac

100 125 75

moisture deficit reaches a maximum value of 175mm.

Approach to estimating impact of warmer climates

An hourly dataset for 100m in the mixed boundary layer was generated as described in the previous section using hourly AWS data from a bracken-covered moorland site near Pateley Bridge, North Yorkshire, for days 140 (19 May) to 209 (27 July) 1984. The site is described in Lockwood et al. (1989). The period was a relatively rainless one when soil moisture changed from near saturation to almost complete dryness.

A number of problems arise in using the dataset for simulations of a climatic warming. As explained earlier, the astronomical factors controlling the Sun’s path across the sky are unaffected by global warming. Figure 2 shows hourly values of incoming short-wave radiation, SW, for a clear sky generated from:

SW = (0.477 + 0.363 sine) S sin0

where 8 is the solar angle (angle between hori-

zon and Sun), and S is the solar constant. This cloudless sky composite is based on short-wave radiation observations made near Harrogate (Lockwood 1979). Also shown in Fig. 2 are actual values of incoming short-wave radiation. On the first relatively cloudless day there are only small differences between the observed radiation and the clear sky composite, but on the second cloudy day the differences are substantial. While the clear sky radiation will be unaffected by global warming, associated changes in cloudiness could cause substantial changes in short-wave radiation received at the surface.

In general, more than half the incoming green- house long-wave radiation received at the ground from the atmosphere comes from gases (water vapour mainly, but also carbon dioxide, etc.) in the lowest lOOm and roughly 90 per cent from the lowest kilometre. The presence of low cloud increases the flux of atmospheric long- wave radiation at the surface because the radiation from the water vapour and carbon dioxide in the lower atmosphere is supplemented by emission from clouds in the waveband which gaseous emission lacks, i . e . 8- 1 3 ym .

294

900 .

800

700

600

Nf E

500 3, c 0

U m n

._ c 400

300

200

100

0

27 June Time (GMT) 28 June

Fig. 2 Values of incoming short- and long-wave radiation for actual and cloudless sky conditions for a site near Pateley Bridge, North Yorkshire, for two days in June, 1984. See text for explanation. During the night 27/28 June, 2mm of rain fell in two events of I mm each.

Figure 2 shows estimates, produced by the simple soil-vegetation-atmosphere model, of the downwelling atmospheric long-wave radiation from the atmosphere. Monteith and Unsworth (1990) suggest that under cloudless skies the mean climatological net long-wave radiative loss from the surface in north-west Europe is about 100Wm-2. The model radiative scheme was modified so as to set the surface net long-wave radiative loss at this value, and then used to produce estimates of the downwelling long-wave radiation under clear skies. These estimates are also shown in Fig. 2. It is seen that, while cloud makes effectively no contribution to the downwelling long-wave radiation across the midday period of 27 June (a relatively cloudless day), it contributes around 90Wm-2 throughout the rainfall of the night

of 27/28 June and only slightly less during the relatively cloudy 28 June.

Under actual global warming all meteorological input parameters may change. Since the impact of increasing temperatures is the focus of this particular study the observed incoming short-wave radiation is used, since the cloud distribution is assumed to be unchanged, as are the other meteorological inputs such as rainfall, wind speed and atmospheric humidity deficit. Downwelling atmospheric long-wave radiation is more of a problem, since it could be enhanced slightly with increasing temperature. Monteith and Unsworth (1990) suggest that downwelling long-wave radiation from a clear sky is given by:

L D = 213 + 5.5 T,

where T, is the screen temperature (“C).

Therefore estimates of evapotranspiration under global warming were made with downwelling long-wave radiation both kept at the observed values and also increased by 5.5 Wm-2 per deg C. It should be noted that the intensification in the downwelling long-wave radiative flux due to temperature increases is small compared with that due to changes in cloud.

Results

Estimates of both transpiration and interception loss (evaporation from rain-water inter- cepted by the canopy) from grass for the nearly rainless sunny summer of 1984 are shown in Figs. 3 and 4(a). Soil moisture deficit measure- ments made with a neutron probe show that the soil at the moorland site was near saturation during days 140-170 and that soil moisture deficit values reached a maximum around day 200 (Lockwood et al. 1989). Figure 3 indicates that lack of soil moisture is severely limiting the grassland transpiration losses from days 180 to 210, causing the value of the weighting factor, &, to decrease sharply. The fall in the value of Q implies that transpiration losses become relatively less dependent on radiation input and more dependent on bulk canopy resistance which is largely due to lack of soil moisture. Suppose that the estimated tem-

295

0.8 '"i h

i i o 150 160 i i o 180 190 200 2 i o zio Day of the year

Fig. 3 hstzmated values of weighting factor, (2, for a moorland grass pasture near Patelev Bndge. North Yorkshtre, 1984. Czrcled values are estimates of soil moisture deficits ( m m ) at ten-dav zntenuls Also shown are estimated grass trunspzratton values i mm) over ten- day penods.

peratures at lOOm in the lower atmosphere are increased by a small amount, say 4 deg C (about twice the projected average increase in UK summer temperatures for ZOSO), while the other meteorological input conditions remain constant. It is seen from Fig. 4(a) that during the period when soil moisture is non-limiting there is a slight reduction in transpiration loss. The long-wave emission from the grass inten- sifies as the canopy becomes warmer, and therefore the available energy, Q,, is reduced causing the transpiration to decrease even though the proportion of the available energy used for transpiration is actually greater. Similarly the interception loss from wet grass is reduced. If the atmospheric downwelling long-wave emission is also increased by 5.5 W m * per deg C, then the enhanced long-wave emission from the warmer canopy is approximately compensated by the intensified downwelling long-wave radiation from the atmosphere. Under these conditions it is seen from Fig. 4(a) that there is a slight increase in both the interception loss and also the transpiration loss when the soil moisture is non-limiting. In contrast, if the lOOm temperatures are decreased by 4 deg C, with a corresponding reduction in the down-

welling long-wave radiation, then there is a slight diminution in both the transpiration and interception losses. During the latter part of the investigation period, the differences in transpiration rates between the various climatic scenarios are dominated by soil moisture rather than temperature.

The bulk canopy resistance scheme for coniferous forest is strongly sensitive to both temperature and atmospheric humidity deficit. Since in this case the value of Q is small (that is, the evaporative loss is dominated by the canopy resistance and atmospheric conditions), increasing temperature will lower the bulk canopy resistance and increase the transpiration loss. Also, since the soil moisture available to the forest is relatively large, soil moisture never strongly limits transpiration loss. Figure 4(b) shows that increasing the lOOm temperatures by 4 deg C , with all other inputs constant, causes an enhancement in the transpiration loss because the lower bulk canopy resistance more than compensates for the decrease in available energy at the canopy surface. In contrast, the interception loss, which is not influenced by bulk canopy resistance, decreases slightly with a simple rise in temperature as is observed in the case of grass. If the lower atmosphere temperatures are increased by 4 deg C , together with a corresponding intensification in the downwelling long-wave radiation, there is a further significant enhancement in the transpiration loss from coniferous forest and a slight increase in the interception loss. Similarly, if the lower atmosphere temperatures are decreased by 4 deg C with a corresponding reduction in downwelling long-wave radiation, there is a significant diminution in transpiration loss.

The results are summarised in Table 1, which shows that over the investigation period evapotranspiration from grass pasture does not vary greatly with temperature, in contrast to that for coniferous forest which shows significant variations. The coniferous forest variations arise from changes in the transpiration loss, so that the interception loss is a larger proportion of the total evapotranspiration at the lower temperatures, but its importance decreases with increasing temperature. At the lower temperatures both the transpiration loss and the

296

Soil moisture deficit (mrn) for observed meterological conditions (a) 1; 3 7 25 44 63 68 63

140 150 160 170 180 190 zoo 2 i o

35

30

25 x

E E 20 0 m ._ +a

15 m W

10

5

0

Day of the year

Ten-day transpiration (T) and interception values

A B C D A B C D A B C D A B C D A B C D A B C D A B C D 140 - 49 150 - 59 160 - 69 170 - 79 180 - 89 190-99 200 - 09

D a y of the year

Soil moisture deficit (mrn) for observed meterological conditions 1 4 13 24 40 56 65

140 150 160 170 180 190 200 210 (b) :

35

30

25 --.. E E - 20 0 .- c

15 m W

10

5

0

140-49 150 - 59 160-69 170-79 180-89 190 - 99 200 - 09 D a y of the year

Fig. 4 Ten-day totals of transpiration and interception loss for (a ) grass pasture and (b ) coniferous forest. Meteorological data for I984 from an automatic weather station sited near Pateley Bridge, North Yorkshire. Upper parts of diagrams contain estimated soil moisture deficits at ten-day intervals. Four estimates under dgfering meteorological input conditions are given of transpiration and interception loss for each ten-day period. The first value marked T' is the transpiration and the second value is the interception loss. The meteorological input conditions are: A. Observed meteorological conditions. B. Lower atmosphere temperatures increased by 4 deg C. C. Lower atmosphere temperatures decreased by 4 deg C, together with downwelling long-wave radiation reduced by 22 Wm-2. D. Lower atmosphere temperatures increased by 4 deg C, together with downwelling long-wave radiation intensajied by 22 Wm-2.

Day of the year

Ten-day transpiration (T) and interception values

297

Table 1 Evapotranspiration (mm) from grass pasture and coniferous forest for days 140-209 ~~

Temperature decreased by 4 deg C with radiation correction Observed conditions

Transpiration Interception Total Transpiration Interception Total

Grass 114.9 18.1 133.0 117.4 18.5 135.9 Forest 76.1 28.1 104.2 100.6 28.7 129.3

Temperature increased by 4 deg C with radiation correction

Transpiration Interception Total

Grass 120.0 19.1 139.1 Forest 122.7 29.3 152.0

Total rainfall 136.5 mm.

total evapotranspiration from coniferous forest are less than that from grass, but at the higher temperatures the coniferous forest values exceed those from grass. Temperatures fall with increasing elevation, therefore the relationship between evapotranspiration from forest and grassland should also change with altitude. On mountains with high and frequent rainfall, differences in interception loss may dominate the water balances of the two vegetation types.

Concluding comments

Global warming may produce a climate in the UK for which there are no recent analogues. The path of the Sun across the sky will not change with global warming, therefore unless there are substantial changes in cloud distribution the energy available to drive evapotranspiration loss will not vary significantly under global warming. Part of the energy for evaporative loss comes from downwelling long- wave radiation from the atmosphere and clouds and this may be intensified in a warmer atmosphere. Calculations presented for a typical short vegetation cover such as a moist grass pasture suggest that transpiration and interception loss will not change greatly with global warming, and could even decrease under certain cir- cumstances. In contrast, for a tall vegetation cover such as coniferous forest, where stomatal resistance is strongly dependent on atmospheric conditions, transpiration could be signlficantly enhanced under global warming. The balance of evapotranspiration loss between grass pasture and forest will therefore change with global warming. Temperature falls with elevation in

mountainous areas, therefore a similar difference in the balance of evapotranspiration loss between grassland and forest should be observed in these regions.

It is demonstrated that vegetation stomatal responses to global warming will cause dis- similar changes in evapotranspiration rates for contrasting vegetation types. In reality the situation is much more complex than the simple example given in this paper. Global warming will probably result in changes in cloud type and altitude distribution (e .g . perhaps more convective cloud) and this could strongly influence both the long- and short-wave radiation received at the surface and therefore the evaporative loss. Also, atmospheric humidity deficits influence bulk canopy resistances and it is not clear how these deficits will change in a warmer atmosphere, or indeed how they may be further modified by transpiring vegetation. Lastly, changes may occur in both rainfall and the length of the growing season, leading to modifications in both vegetation canopy struc- ture and composition. Thus summers may appear to become drier with global warming, not because evapotranspiration rates are necessarily higher for many short vegetation covers, but because vegetation growth commences earlier in the spring with a corresponding earlier start to the evaporation season, resulting in soil moisture exhaustion earlier in the summer.

References Dickinson, R. E. (1992) Land surface. In:

Trenberth, K. E. (Ed.) Climate system modeling, Cambridge University Press, pp. 149-171

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Garratt, J. R. (1992) T h atmospheric boundary layer. Cambridge University Press, 316 pp.

Kiehl, J. T. (1992) Atmospheric general circulation modeling. In: Trenberth, K. E. (Ed.) Climate system modeling, Cambridge University Press, pp. 3 1%369

Lockwood J. G. (1979) Causes of climate. Edward Arnold, London, 260 pp.

Lockwood, J. G., Jones, C. A. and Smith, R. T. (1989) The estimation of soil moisture deficits using meteorological models at an upland moorland site in northern England. Agric. For. Meteorol., 46, pp. 41-63.

McNaughton, K. G. and Jarvis, P. G. (1983) Pre- dicting effects of vegetation changes on transpiration and evaporation. In: Kozlowski, T. T. (Ed.) Water deficits and plant growth Vol. 7, Aca- demic Press, New York, pp. 1-47

McNaughton, K. G. and Spriggs, T. W. (1986) A mixed-layer model for regional evaporation. Boundary-layer Meteorol., 34, pp. 243-262

Monteith, J. L. and Unsworth, M. H. (1990) Prin- ciples of environmental physics. Edward Arnold, London, 291 pp.

Oke, T. R. (1978) Boundary layer climates. Methuen, 272 pp.

Sellers, P. J. (1992) Biophysical models of land surface processes. In: Trenberth, K. E. (Ed.) Climate system modeling, Cambridge University Press, pp. 45 1-490

Strangeways, I. C. and Smith, S. W. (1985) Devel- opment and use of automatic weather stations. Weather, 40, pp. 277-285

Thompson, N. Barrie, I. A. and Ayles, M. (1981) The Meteorological Office Rainfall and Evap- oration Calculation System: MORECS. Met 0 8 Hydrol. Memo. No. 45, Meteorological Office, Bracknell

United Kingdom Climate Change Impacts Review Group (1991) The potential effects of climate change in the United Kingdom. HMSO, London, 124 pp.

Weather forecasting in Britain, 1939-80*

R. A. S. Ratcliffe Ascot, Berkshire

Before the Second World War there was much less information of all kinds available to the forecaster in the Meteorological Office than there is now. To begin with there was virtually no upper-air information on a regular basis. There were upper-air ascents done by aircraft based at Bircham Newton (Norfolk) and exper- imental radiosondes had been flown at Larkhill on Salisbury Plain shortly before the War but these did not become routine until 1939-40.

Upper-wind information came from two main sources. Firstly, hydrogen-filled balloons (which were supposed to rise at 300ftmin-') were followed by means of a theodolite and the winds were worked out from the azimuth, elevation and assumed height of the balloon. At some stations larger balloons, assumed to

* Based on a talk given at the Royal Meteorological Society Summer Meeting at Keele University in August 1991.

rise at lOOOftmin-', were used. A good assist- ant (or forecaster!) could follow a balloon single- handed, work out the winds between sightings and come back into the office with the message already coded for transmission to Meteoro- logical Office Headquarters on the teleprinter. The other method of upper-wind measurement was by using a nephoscope (see Fig. 1). This was a kind of inverted rake which revolved about a vertical axis. You aligned the rake so that the cloud was moving along the tips of the tines; that gave you the direction of motion of the cloud. You then timed the passage of a prominent portion of the cloud past a few tines of the rake and, knowing the height of the tines above eye-level and assuming the height of the cloud, you could deduce the speed of the cloud. This technique was mainly used to determine the approximate speed of cirrus cloud to give an idea of the high-level wind.

With such primitive data it is not surprising

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impact of global warming on evapotranspiration

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