2. Irrigation. Key words: right amount at right time What if it s too little too late? Too much too often?

Size: px

Start display at page:

Download "2. Irrigation. Key words: right amount at right time What if it s too little too late? Too much too often?"

Millicent Fletcher
6 years ago
Views:

1 2. Irrigation Key words: right amount at right time What if it s too little too late? 2-1 Too much too often? To determine the timing and amount of irrigation, we need to calculate soil water balance. Mass balance concept IN - OUT = Storage Change Q in - Q out = S

2 Major components in soil water balance 2-2 control volume We need to evaluate each term. Terminology for soil moisture storage Consider a block of soil. mass (M) volume (V). particle density (ρ m ) = M solid / V solid total gas water solid particles void dry bulk density (ρ b ) = M solid / V total porosity (n) = V void / V total volumetric water content (θ) = V water / V total Storage of water in the control volume is expressed by volumetric water content. A soil is saturated when water fills up the void space.

Sands have lower field capacity than clays. wilting Wilting point is usually considered 15 atm. 1 atm is equivalent to the pressure of 10 m of water.

3 2-3 Field capacity (FC) : the amount of water that remains in soil after gravitational drainage. Wilting point (WP): the amount of water that is so tightly held in the soil that plants can not utilize it. Field capacity depends on the type of soil. Sands have lower field capacity than clays. wilting Wilting point is usually considered 15 atm. 1 atm is equivalent to the pressure of 10 m of water. drainage Plant roots can pull the water against the soil s binding force (suction) as large as the force corresponding to the pressure of 150-m water column! The objective of irrigation is to keep θ reasonably well above the wilting point. FC WP optimal range

4 Porosity of most natural soils = 2-4 Field capacity depends on the soil texture. Why? Dunne and Leopold (1978, Fig. 6-9) Note that the heavy textured soil may have a fair amount of water, but this water is not accessible to plants.

5 Evapotranspiration 2-5 Our approach is to first determine potential evaporation and multiply it by a crop coefficient. Evaporation, potential (PE) and actual (AE) PE: the rate that would occur under plentiful supply of water. e.g. lake evaporation AE: the actual rate. e.g. evaporation from soil surface Potential evaporation There are several methods to estimate PE. (1) If a gross estimate over a large region (>100 km 100 km) is needed, use a published evaporation map or an empirical formula based on air temperature. PE = f (T a ) T a : annual mean temperature e.g. Thornthwaite method (D&L, p.137) (2) If site specific data is needed, use a Class-A pan (D&L, p.100). (3) If standard meteorological data is available, use Penman method (described later)

6 Evaporation map 2-6 Compilation of evaporation in map forms. General information: Hydrological Atlas of Canada Specific information: published journal articles. e.g. Lake evaporation map by Fred Morton. Annual lake evaporation (mm) Morton, F.I., Operational estimates of lake evaporation. Journal of Hydrology, 66:

7 Penman method 2-7 This method combines the energy balance and mass transfer methods and eliminates the need for highly specialized data collection. Energy balance method Consider an isolated water body Q s -Q rs -Q lw = Q h + Q e + Q θ [1] Q s Q rs Q lw Q h Q e Q s : incoming short-wave radiation Q θ Q rs : reflected short-wave radiation Q lw : net long-wave radiation Q h : sensible heat transfer Q e : latent heat transfer All units are in [J m -2 d -1 ] Q θ : heat storage change This is identical to D&L (p.103, 4-2), except that Q v and Q ev are assumed negligible. It is possible to measure radiation components using radiometers. Data are available from meteorological stations. It is also reasonably easy to estimate Q θ. It is very difficult to measure Q h and Q e separately.

8 However, it is relatively easy to estimate the ratio Q h / Q e. 2-8 Each parcel contains lots of molecules. Parcels near the water surface contain more water vapor molecules than the ones far from the surface. Random motion of the parcels lead to the net upward transfer of water vapor molecules. The same principle applies to heat transport. If the water surface is warmer than air, random motion of the parcels lead to the net upward transfer of warm air parcels. Under a given wind condition, Q h T s - T a Q e e s - e a T s : water surface temperature ( C) T a : air temperature at specified height ( C) e s : vapor pressure at the water surface (mb) e a : vapor pressure in the air at specified height (mb)

9 2-9 R = Q h / Q e (T s - T a ) / (e s - e a ) R is called Bowen ratio. More precisely (D&L, p.104); R = Q h / Q e = p (T s - T a ) / (e s - e a ) [2] where p is barometric pressure in mb. If we know R, we can write Q h = RQ e. Then substitute this into the energy balance equation [1] to get; Q s -Q rs -Q lw = Q e (1+R) + Q θ Q e = (Q s -Q rs -Q lw -Q θ ) / (1+R) [3] Now we need to convert Q e [J d -1 m -2 ] to E 0 [m d -1 ]. Suppose a water surface having an area A [m 2 ]. The volume of water that evaporates from this surface is equal to AE 0 [m 3 d -1 ]. How much energy is required to this much of water? It is well known that J is needed to vaporize 1 kg of water. This quantity (L) is called latent heat of vaporization. Therefore, we have AE 0 ρl = AQ e ρ : density of water [kg m -3 ] E 0 = Q e /(ρl) [4] In theory, we can determine E 0 from [2]-[4]. However, it is difficult to measure T s and e s.

10 Mass transfer method 2-10 As we have seen, E 0 ρl = Q e e s - e a or E 0 e s - e a Can we directly determine the proportionality coefficient? The coefficient clearly depends on the wind speed u [m s -1 ]. E 0 = Nf(u) (e s - e a ) [5] f(u) is called wind function. N is a constant. N and f(u) are experimentally determined. It is common to use u 2, measured 2 m above the ground surface. S Dunne and Leopold (1978, Fig. 4-6) Note that Eq. [5] requires e s, which is difficult to measure. Penman recognized this problem in 1940 s in an agricultural research station in England, and combined [3] and [5] to derive an equation for estimating E 0 without using T s and e s.

11 Penman equation Penman used a well-defined relationship between air temperature (T a ) and the saturation vapor pressure (e sa ). The slope of the curve is [mb ºC -1 ]. e sa (mb) Penman equation can be written as: E 0 100Qn + Eaγ ρl = + γ where Q n = Q s - Q rs - Q lw, called net radiation slope = T a ( C) Eq.[4-20] in DL γ = 0.66 [mb ºC -1 ], called psychrometric constant. E a = ( u 2 )(e sa - e a ) called aerodynamic evaporation u 2 : wind speed [km d -1 ] measured 2 m above the surface. Note that Q n is given in J m -2 d -1, and 100 converts the unit of evaporation from m d -1 to cm d -1. The coefficients in the equation for E a are only valid when the specified units are used for all parameters.

Actual ET: Effects of soil moisture 2-12 Plants can transpire at the potential rate, limited only by meteorological factors, when the soil is reasonably wet.

12 Actual ET: Effects of soil moisture 2-12 Plants can transpire at the potential rate, limited only by meteorological factors, when the soil is reasonably wet. wind transpiration radiation The ET rate is reduced when the soil water becomes a limiting factor. actual evapotranspiration (AET) Water in drier soil is more strongly bound to the soil particles, resulting in a higher resistance to root uptake. Root uptake Definition of available water Available water (AW) = (water content - wilting point) rooting depth Available water capacity (AWC) = (field capacity - wilting point) rooting depth

13 2-13 Actual evapotranspiration (AET) is commonly assumed to be a fraction of the potential evapotranspiration (PET). The relationship between AET and PET is given by: AW AET = PET f AWC where f ( ) is a function of the ratio AW/AWC. The function depends on many variables, most important of which is the soil type. AET/PET is also called crop coefficient. AET/PET AW/AWC Dunne and Leopold (1978, Fig. 5-6) After a rain or irrigation event, when soil is at near field capacity, plants transpire at a rate close to PET. The rate of ET decreases as available water decreases with time, depending on the rooting depths. Dunne and Leopold (1978, Fig. 5-9)

14 Soil water balance (P - O) + I r - E - G = S/ t P - O E I r 2-14 S = θ av z r θ av : average water content z r : depth of root zone z = 0 S = θ av z r Storage at field capacity G z = z R S FC = θ FC z r Any addition of water in excess of S FC will generate G. In irrigation, we want to keep S between S FC and wilting point (S WP ) by adjusting the timing and amount. For practical methods of calculation, see the handout for laboratory exercise.

15 Development of soil salinity 2-15 Groundwater is generally recharged at high-elevation areas and discharged into streams in low-elevation areas. discharge recharge In semi-arid regions there are many groundwater discharge areas without streams. discharge area evapotranspiration groundwater Groundwater transports dissolved species, particularly sulfate salts, up to the soil zone. Evapotranspiration removes water from the soil zone leaving salts behind. High soil salinity may have adverse effects on plants. Improper irrigation practice often leads to increased groundwater discharge and exacerbates soil salinity. High salinity areas can be identified by the growth of indicator plants such as foxtail barley or patchy occurrence of vegetation ( field trip).

GEOG415 Mid-term Exam 110 minute February 27, 2003

GEOG415 Mid-term Exam 110 minute February 27, 2003 1 Name: ID: 1. The graph shows the relationship between air temperature and saturation vapor pressure. (a) Estimate the relative humidity of an air parcel