2.5. Momentum, Sensible Heat, and Latent Heat Fluxes¶
The zonal and meridional momentum fluxes (kg m^{1} s^{2}), sensible heat flux (W m^{2}), and water vapor flux (kg m^{2} s^{1}) between the atmosphere at reference height (m) [where is height for wind (momentum) (), temperature (sensible heat) (), and humidity (water vapor) (); with zonal and meridional winds and (m s^{1}), potential temperature (K), and specific humidity (kg kg^{1})] and the surface [with , , , and ] are
(2.5.1)¶
(2.5.2)¶
(2.5.3)¶
(2.5.4)¶
These fluxes are derived in the next section from MoninObukhov similarity theory developed for the surface layer (i.e., the nearly constant flux layer above the surface sublayer). In this derivation, and are defined to equal zero at height (the apparent sink for momentum) so that is the aerodynamic resistance (s m^{1}) for momentum between the atmosphere at height and the surface at height . Thus, the momentum fluxes become
(2.5.5)¶
(2.5.6)¶
Likewise, and are defined at heights and (the apparent sinks for heat and water vapor, respectively
are the aerodynamic resistances (s m^{1}) to sensible heat and water vapor transfer between the atmosphere at heights and and the surface at heights and , respectively. The specific heat capacity of air (J kg^{1} K^{1}) is a constant (Table 2.2.7). The atmospheric potential temperature used here is
(2.5.7)¶
where is the air temperature (K) at height and K m^{1} is the negative of the dry adiabatic lapse rate [this expression is firstorder equivalent to (Stull 1988), where is the surface pressure (Pa), is the atmospheric pressure (Pa), and is the gas constant for dry air (J kg^{1} K^{1}) (Table 2.2.7)]. By definition, . The density of moist air (kg m^{3}) is
(2.5.8)¶
where the atmospheric vapor pressure (Pa) is derived from the atmospheric specific humidity
(2.5.9)¶
2.5.1. MoninObukhov Similarity Theory¶
The surface vertical kinematic fluxes of momentum and (m^{2} s_{2}), sensible heat (K m s ^{1}), and latent heat (kg kg^{1} m s^{1}), where , , , , and are zonal horizontal wind, meridional horizontal wind, vertical velocity, potential temperature, and specific humidity turbulent fluctuations about the mean, are defined from MoninObukhov similarity applied to the surface layer. This theory states that when scaled appropriately, the dimensionless mean horizontal wind speed, mean potential temperature, and mean specific humidity profile gradients depend on unique functions of (Zeng et al. 1998) as
(2.5.10)¶
(2.5.11)¶
(2.5.12)¶
where is height in the surface layer (m), is the displacement height (m), is the MoninObukhov length scale (m) that accounts for buoyancy effects resulting from vertical density gradients (i.e., the atmospheric stability), k is the von Karman constant (Table 2.2.7), and is the atmospheric wind speed (m s^{1}). , , and are universal (over any surface) similarity functions of that relate the constant fluxes of momentum, sensible heat, and latent heat to the mean profile gradients of , , and in the surface layer. In neutral conditions, . The velocity (i.e., friction velocity) (m s^{1}), temperature (K), and moisture (kg kg^{1}) scales are
(2.5.13)¶
(2.5.14)¶
(2.5.15)¶
where is the shearing stress (kg m^{1} s^{2}), with zonal and meridional components and , respectively, is the sensible heat flux (W m^{2}) and is the water vapor flux (kg m^{2} s^{1}).
The length scale is the MoninObukhov length defined as
(2.5.16)¶
where is the acceleration of gravity (m s^{2}) (Table 2.2.7), and is the reference virtual potential temperature. indicates stable conditions. indicates unstable conditions. for neutral conditions. The temperature scale is defined as
(2.5.17)¶
where is the atmospheric potential temperature.
Following Panofsky and Dutton (1984), the differential equations for , , and can be integrated formally without commitment to their exact forms. Integration between two arbitrary heights in the surface layer and ( ) with horizontal winds and , potential temperatures and , and specific humidities and results in
(2.5.18)¶
(2.5.19)¶
(2.5.20)¶
The functions , , and are defined as
(2.5.21)¶
(2.5.22)¶
(2.5.23)¶
where , , and are the roughness lengths (m) for momentum, sensible heat, and water vapor, respectively.
Defining the surface values
and the atmospheric values at
(2.5.24)¶
the integral forms of the fluxgradient relations are
(2.5.25)¶
(2.5.26)¶
(2.5.27)¶
The constraint is required simply for numerical reasons to prevent and from becoming small with small wind speeds. The convective velocity accounts for the contribution of large eddies in the convective boundary layer to surface fluxes as follows
(2.5.28)¶
where is the convective velocity scale
(2.5.29)¶
is the convective boundary layer height (m), and .
The momentum flux gradient relations are (Zeng et al. 1998)
(2.5.30)¶
The sensible and latent heat flux gradient relations are (Zeng et al. 1998)
(2.5.31)¶
To ensure continuous functions of , , and , the simplest approach (i.e., without considering any transition regimes) is to match the relations for very unstable and unstable conditions at for and for (Zeng et al. 1998). The flux gradient relations can be integrated to yield wind profiles for the following conditions:
Very unstable
(2.5.32)¶
Unstable
(2.5.33)¶
Stable
(2.5.34)¶
Very stable
(2.5.35)¶
where
(2.5.36)¶
and
.
The potential temperature profiles are:
Very unstable
(2.5.37)¶
Unstable
(2.5.38)¶
Stable
(2.5.39)¶
Very stable
(2.5.40)¶
The specific humidity profiles are:
Very unstable
(2.5.41)¶
Unstable
(2.5.42)¶
Stable
(2.5.43)¶
Very stable
(2.5.44)¶
where
(2.5.45)¶
Using the definitions of , , and , an iterative solution of these equations can be used to calculate the surface momentum, sensible heat, and water vapor flux using atmospheric and surface values for , , and except that depends on , , and . However, the bulk Richardson number
(2.5.46)¶
is related to (Arya 2001) as
(2.5.47)¶
Using for unstable conditions and for stable conditions to determine and , the inverse relationship can be solved to obtain a first guess for and thus from
(2.5.48)¶
Upon iteration (section 2.5.3.2), the following is used to determine and thus
(2.5.49)¶
where
The difference in virtual potential air temperature between the reference height and the surface is
(2.5.50)¶
The momentum, sensible heat, and water vapor fluxes between the surface and the atmosphere can also be written in the form
(2.5.51)¶
(2.5.52)¶
(2.5.53)¶
(2.5.54)¶
where the aerodynamic resistances (s m^{1}) are
(2.5.55)¶
(2.5.56)¶
(2.5.57)¶
A 2m height “screen” temperature is useful for comparison with observations
(2.5.58)¶
where for convenience, “2m” is defined as 2 m above the apparent sink for sensible heat (). Similarly, a 2m height specific humidity is defined as
(2.5.59)¶
Relative humidity is
(2.5.60)¶
where is the saturated specific humidity at the 2m temperature (section 2.5.5).
A 10m wind speed is calculated as (note that this is not consistent with the 10m wind speed calculated for the dust model as described in Chapter 2.31)
(2.5.61)¶
2.5.2. Sensible and Latent Heat Fluxes for NonVegetated Surfaces¶
Surfaces are considered nonvegetated for the surface flux calculations if leaf plus stem area index (section 2.2.1.4). By definition, this includes bare soil and glaciers. The solution for lakes is described in Chapter 2.12. For these surfaces, the surface may be exposed to the atmosphere, snow covered, and/or surface water covered, so that the sensible heat flux (W m^{2}) is, with reference to Figure 2.5.1,
(2.5.62)¶
where , , and are the exposed, snow covered, and surface water covered fractions of the grid cell. The individual fluxes based on the temperatures of the soil , snow , and surface water are
(2.5.63)¶
(2.5.64)¶
(2.5.65)¶
where is the density of atmospheric air (kg m^{3}), is the specific heat capacity of air (J kg^{1} K^{1}) (Table 2.2.7), is the atmospheric potential temperature (K), and is the aerodynamic resistance to sensible heat transfer (s m^{1}).
The water vapor flux (kg m^{2} s^{1}) is, with reference to Figure 2.5.2,
(2.5.66)¶
(2.5.67)¶
(2.5.68)¶
(2.5.69)¶
where is the atmospheric specific humidity (kg kg^{1}), , , and are the specific humidities (kg kg^{1}) of the soil, snow, and surface water, respectively, is the aerodynamic resistance to water vapor transfer (s m^{1}), and is the soil resistance to water vapor transfer (s m^{1}). The specific humidities of the snow and surface water are assumed to be at the saturation specific humidity of their respective temperatures
(2.5.70)¶
(2.5.71)¶
The specific humidity of the soil surface is assumed to be proportional to the saturation specific humidity
(2.5.72)¶
where is the saturated specific humidity at the soil surface temperature (section 2.5.5). The factor is a function of the surface soil water matric potential as in Philip (1957)
(2.5.73)¶
where is the gas constant for water vapor (J kg^{1} K^{1}) (Table 2.2.7), is the gravitational acceleration (m s^{2}) (Table 2.2.7), and is the soil water matric potential of the top soil layer (mm). The soil water matric potential is
(2.5.74)¶
where is the saturated matric potential (mm) (section 2.7.3.1), is the Clapp and Hornberger (1978) parameter (section 2.7.3.1), and is the wetness of the top soil layer with respect to saturation. The surface wetness is a function of the liquid water and ice content
(2.5.75)¶
where is the thickness of the top soil layer (m), and are the density of liquid water and ice (kg m^{3}) (Table 2.2.7), and are the mass of liquid water and ice of the top soil layer (kg m^{2}) (Chapter 2.7), and is the saturated volumetric water content (i.e., porosity) of the top soil layer (mm^{3} mm^{3}) (section 2.7.3.1). If and , then and . This prevents large increases (decreases) in for small increases (decreases) in soil moisture in very dry soils.
The resistance to water vapor transfer occurring within the soil matrix (s m^{1}) is
(2.5.76)¶
where is the thickness of the dry surface layer (m), is the molecular diffusivity of water vapor in air (m^{2} s^{2}) and (unitless) describes the tortuosity of the vapor flow paths through the soil matrix (Swenson and Lawrence 2014).
The thickness of the dry surface layer is given by
(2.5.77)¶
where is a parameter specifying the length scale of the maximum DSL thickness (default value = 15 mm), (mm^{3} mm^{3}) is the moisture value at which the DSL initiates, (mm^{3} mm^{3}) is the moisture value of the top model soil layer, and (mm^{3} mm^{3}) is the ‘air dry’ soil moisture value (Dingman 2002):
(2.5.78)¶
where is the porosity (mm^{3} mm^{3}), is the saturated soil matric potential (mm), mm is the air dry matric potential, and is a function of soil texture (section 2.7.3.1).
The soil tortuosity is
(2.5.79)¶
where (mm^{3} mm^{3}) is the air filled pore space
(2.5.80)¶
depends on temperature
(2.5.81)¶
where (K) is the temperature of the top soil layer and (K) is the freezing temperature of water (Table 2.2.7).
The roughness lengths used to calculate , , and are , , and . The displacement height . The momentum roughness length is for soil, glaciers, and for snowcovered surfaces (). In general, is different from because the transfer of momentum is affected by pressure fluctuations in the turbulent waves behind the roughness elements, while for heat and water vapor transfer no such dynamical mechanism exists. Rather, heat and water vapor must be transferred by molecular diffusion across the interfacial sublayer. The following relation from Zilitinkevich (1970) is adopted by Zeng and Dickinson 1998
(2.5.82)¶
where the quantity is the roughness Reynolds number (and may be interpreted as the Reynolds number of the smallest turbulent eddy in the flow) with the kinematic viscosity of air m^{2} s^{1} and .
The numerical solution for the fluxes of momentum, sensible heat, and water vapor flux from nonvegetated surfaces proceeds as follows:
An initial guess for the wind speed is obtained from (2.5.24) assuming an initial convective velocity m s^{1} for stable conditions ( as evaluated from (2.5.50) ) and for unstable conditions ().
An initial guess for the MoninObukhov length is obtained from the bulk Richardson number using (2.5.46) and (2.5.48).
The following system of equations is iterated three times:
Potential temperature scale ((2.5.37) , (2.5.38), (2.5.39), (2.5.40))
Roughness lengths for sensible and latent heat ((2.5.82) )
Virtual potential temperature scale ( (2.5.17))
Wind speed including the convective velocity, ( (2.5.24))
MoninObukhov length ((2.5.49))
Aerodynamic resistances , , and ((2.5.55), (2.5.56), (2.5.57))
Sensible heat flux ((2.5.62))
Water vapor flux ((2.5.66))
2m height air temperature and specific humidity ((2.5.58) , (2.5.59))
The partial derivatives of the soil surface fluxes with respect to ground temperature, which are needed for the soil temperature calculations (section 2.6.1) and to update the soil surface fluxes (section 2.5.4), are
(2.5.83)¶
(2.5.84)¶
where
(2.5.85)¶
The partial derivatives and , which cannot be determined analytically, are ignored for and .
2.5.3. Sensible and Latent Heat Fluxes and Temperature for Vegetated Surfaces¶
In the case of a vegetated surface, the sensible heat and water vapor flux are partitioned into vegetation and ground fluxes that depend on vegetation and ground temperatures in addition to surface temperature and specific humidity . Because of the coupling between vegetation temperature and fluxes, NewtonRaphson iteration is used to solve for the vegetation temperature and the sensible heat and water vapor fluxes from vegetation simultaneously using the ground temperature from the previous time step. In section 2.5.3.1, the equations used in the iteration scheme are derived. Details on the numerical scheme are provided in section 2.5.3.2.
2.5.3.1. Theory¶
The air within the canopy is assumed to have negligible capacity to store heat so that the sensible heat flux between the surface at height and the atmosphere at height must be balanced by the sum of the sensible heat from the vegetation and the ground
(2.5.86)¶
where, with reference to Figure 2.5.1,
(2.5.87)¶
(2.5.88)¶
(2.5.89)¶
where
(2.5.90)¶
(2.5.91)¶
(2.5.92)¶
where is the density of atmospheric air (kg m^{3}), is the specific heat capacity of air (J kg^{1} K^{1}) (Table 2.2.7), is the atmospheric potential temperature (K), and is the aerodynamic resistance to sensible heat transfer (s m^{1}).
Here, is the surface temperature at height , also referred to as the canopy air temperature. and are the exposed leaf and stem area indices (section 2.2.1.4), is the leaf boundary layer resistance (s m^{1}), and is the aerodynamic resistance (s m^{1}) to heat transfer between the ground at height and the canopy air at height .
Equations (2.5.86)  (2.5.89) can be solved for the canopy air temperature
(2.5.93)¶
where
(2.5.94)¶
(2.5.95)¶
(2.5.96)¶
are the sensible heat conductances from the canopy air to the atmosphere, the ground to canopy air, and leaf surface to canopy air, respectively (m s^{1}).
When the expression for is substituted into equation (2.5.88), the sensible heat flux from vegetation is a function of , , and
(2.5.97)¶
Similarly, the expression for can be substituted into equation to obtain the sensible heat flux from ground
(2.5.98)¶
The air within the canopy is assumed to have negligible capacity to store water vapor so that the water vapor flux between the surface at height and the atmosphere at height must be balanced by the sum of the water vapor flux from the vegetation and the ground
(2.5.99)¶
where, with reference to Figure 2.5.2,
(2.5.100)¶
(2.5.101)¶
(2.5.102)¶
where
(2.5.103)¶
(2.5.104)¶
(2.5.105)¶
where is the atmospheric specific humidity (kg kg^{1}), is the aerodynamic resistance to water vapor transfer (s m^{1}), (kg kg^{1}) is the saturation water vapor specific humidity at the vegetation temperature (section 2.5.5), , , and are the specific humidities of the soil, snow, and surface water (section 2.5.2), is the aerodynamic resistance (s m^{1}) to water vapor transfer between the ground at height and the canopy air at height , and ((2.5.76)) is a resistance to diffusion through the soil (s m^{1}). is the total resistance to water vapor transfer from the canopy to the canopy air and includes contributions from leaf boundary layer and sunlit and shaded stomatal resistances , , and (Figure 2.5.2). The water vapor flux from vegetation is the sum of water vapor flux from wetted leaf and stem area (evaporation of water intercepted by the canopy) and transpiration from dry leaf surfaces
(2.5.106)¶
Equations (2.5.99)  (2.5.102) can be solved for the canopy specific humidity
(2.5.107)¶
where
(2.5.108)¶
(2.5.109)¶
(2.5.110)¶
are the water vapor conductances from the canopy air to the atmosphere, the leaf to canopy air, and ground to canopy air, respectively. The term is determined from contributions by wet leaves and transpiration and limited by available water and potential evaporation as
(2.5.111)¶
where is the fraction of leaves and stems that are wet (section 2.7.1), is canopy water (kg m^{2}) (section 2.7.1), is the time step (s), and is a soil moisture function limiting transpiration (Chapter 2.9). The potential evaporation from wet foliage per unit wetted area is
(2.5.112)¶
The term is
(2.5.113)¶
where is the fraction of leaves that are dry (section 2.7.1), and are the sunlit and shaded leaf area indices (section 2.4.1), and and are the sunlit and shaded stomatal resistances (s m^{1}) (Chapter 2.9).
When the expression for is substituted into equation (2.5.101), the water vapor flux from vegetation is a function of , , and
(2.5.114)¶
Similarly, the expression for can be substituted into (2.5.84) to obtain the water vapor flux from the ground beneath the canopy
(2.5.115)¶
The aerodynamic resistances to heat (moisture) transfer between the ground at height ( ) and the canopy air at height () are
(2.5.116)¶
where
(2.5.117)¶
is the magnitude of the wind velocity incident on the leaves (equivalent here to friction velocity) (m s^{1}) and is the turbulent transfer coefficient between the underlying soil and the canopy air. is obtained by interpolation between values for dense canopy and bare soil (Zeng et al. 2005)
(2.5.118)¶
where the weight is
(2.5.119)¶
The dense canopy turbulent transfer coefficient (Dickinson et al. 1993) is
(2.5.120)¶
The bare soil turbulent transfer coefficient is
(2.5.121)¶
where the kinematic viscosity of air m^{2} s^{1} and .
The leaf boundary layer resistance is
(2.5.122)¶
where ms^{1/2} is the turbulent transfer coefficient between the canopy surface and canopy air, and is the characteristic dimension of the leaves in the direction of wind flow (Table 2.5.1).
The partial derivatives of the fluxes from the soil beneath the canopy with respect to ground temperature, which are needed for the soil temperature calculations (section 2.6.1) and to update the soil surface fluxes (section 2.5.4), are
(2.5.123)¶
(2.5.124)¶
The partial derivatives and , which cannot be determined analytically, are ignored for and .
The roughness lengths used to calculate , , and from (2.5.55), (2.5.56), and (2.5.57) are , , and . The vegetation displacement height and the roughness lengths are a function of plant height and adjusted for canopy density following Zeng and Wang (2007)
(2.5.125)¶
(2.5.126)¶
where is canopy top height (m) (Table 2.2.2), and are the ratio of momentum roughness length and displacement height to canopy top height, respectively (Table 2.5.1), and is the ground momentum roughness length (m) (section 2.5.2). The fractional weight is determined from
(2.5.127)¶
where and (m^{2} m^{2}) is a critical value of exposed leaf plus stem area for which reaches its maximum.
Plant functional type 
(m) 


NET Temperate 
0.055 
0.67 
0.04 
NET Boreal 
0.055 
0.67 
0.04 
NDT Boreal 
0.055 
0.67 
0.04 
BET Tropical 
0.075 
0.67 
0.04 
BET temperate 
0.075 
0.67 
0.04 
BDT tropical 
0.055 
0.67 
0.04 
BDT temperate 
0.055 
0.67 
0.04 
BDT boreal 
0.055 
0.67 
0.04 
BES temperate 
0.120 
0.68 
0.04 
BDS temperate 
0.120 
0.68 
0.04 
BDS boreal 
0.120 
0.68 
0.04 
C_{3} arctic grass 
0.120 
0.68 
0.04 
C_{3} grass 
0.120 
0.68 
0.04 
C_{4} grass 
0.120 
0.68 
0.04 
Crop R 
0.120 
0.68 
0.04 
Crop I 
0.120 
0.68 
0.04 
Corn R 
0.120 
0.68 
0.04 
Corn I 
0.120 
0.68 
0.04 
Temp Cereal R 
0.120 
0.68 
0.04 
Temp Cereal I 
0.120 
0.68 
0.04 
Winter Cereal R 
0.120 
0.68 
0.04 
Winter Cereal I 
0.120 
0.68 
0.04 
Soybean R 
0.120 
0.68 
0.04 
Soybean I 
0.120 
0.68 
0.04 
2.5.3.2. Numerical Implementation¶
Canopy energy conservation gives
(2.5.128)¶
where is the solar radiation absorbed by the vegetation (section 2.4.1), is the net longwave radiation absorbed by vegetation (section 2.4.2), and and are the sensible and latent heat fluxes from vegetation, respectively. The term is taken to be the latent heat of vaporization (Table 2.2.7).
, , and depend on the vegetation temperature . The NewtonRaphson method for finding roots of nonlinear systems of equations can be applied to iteratively solve for as
(2.5.129)¶
where and the subscript “n” indicates the iteration.
The partial derivatives are
(2.5.130)¶
(2.5.131)¶
(2.5.132)¶
The partial derivatives and , which cannot be determined analytically, are ignored for and . However, if changes sign more than four times during the temperature iteration, . This helps prevent “flipflopping” between stable and unstable conditions. The total water vapor flux , transpiration flux , and sensible heat flux are updated for changes in leaf temperature as
(2.5.133)¶
(2.5.134)¶
(2.5.135)¶
The numerical solution for vegetation temperature and the fluxes of momentum, sensible heat, and water vapor flux from vegetated surfaces proceeds as follows:
Initial values for canopy air temperature and specific humidity are obtained from
(2.5.136)¶
(2.5.137)¶
An initial guess for the wind speed is obtained from (2.5.24) assuming an initial convective velocity m s^{1} for stable conditions ( as evaluated from (2.5.50) ) and for unstable conditions ().
An initial guess for the MoninObukhov length is obtained from the bulk Richardson number using equation and (2.5.46) and (2.5.48).
Iteration proceeds on the following system of equations:
Aerodynamic resistances , , and ((2.5.55), (2.5.56), (2.5.57))
Magnitude of the wind velocity incident on the leaves ((2.5.117) )
Leaf boundary layer resistance ((2.5.136) )
Aerodynamic resistances and ((2.5.116) )
Sunlit and shaded stomatal resistances and (Chapter 2.9)
Sensible heat conductances , , and ((2.5.94), (2.5.95), (2.5.96))
Latent heat conductances , , and ((2.5.108), (2.5.109), (2.5.110))
Sensible heat flux from vegetation ((2.5.97) )
Latent heat flux from vegetation ((2.5.101) )
If the latent heat flux has changed sign from the latent heat flux computed at the previous iteration (), the latent heat flux is constrained to be 10% of the computed value. The difference between the constrained and computed value ( ) is added to the sensible heat flux later.
Change in vegetation temperature ((2.5.129) ) and update the vegetation temperature as . is constrained to change by no more than 1ºK in one iteration. If this limit is exceeded, the energy error is
(2.5.138)¶
where . The error is added to the sensible heat flux later.
Water vapor flux ((2.5.133) )
Transpiration ((2.5.134) if , otherwise )
The water vapor flux is constrained to be less than or equal to the sum of transpiration and the water available from wetted leaves and stems . The energy error due to this constraint is
(2.5.139)¶
The error is added to the sensible heat flux later.
Sensible heat flux ((2.5.135) ). The three energy error terms, , , and are also added to the sensible heat flux.
The saturated vapor pressure (Chapter 2.9), saturated specific humidity and its derivative at the leaf surface (section 2.5.5), are reevaluated based on the new .
Canopy air temperature ((2.5.93) )
Canopy air specific humidity ((2.5.107) )
Temperature difference
Specific humidity difference
Potential temperature scale where was calculated earlier in the iteration
Humidity scale where was calculated earlier in the iteration
Virtual potential temperature scale ((2.5.17) )
Wind speed including the convective velocity, ((2.5.24) )
MoninObukhov length ((2.5.49) )
The iteration is stopped after two or more steps if and where , or after forty iterations have been carried out.
Sensible heat flux from ground ((2.5.89) )
Water vapor flux from ground ((2.5.102) )
2m height air temperature , specific humidity , relative humidity ((2.5.58) , (2.5.59), (2.5.60))
2.5.4. Update of Ground Sensible and Latent Heat Fluxes¶
The sensible and water vapor heat fluxes derived above for bare soil and soil beneath canopy are based on the ground surface temperature from the previous time step and are used as the surface forcing for the solution of the soil temperature equations (section 2.6.1). This solution yields a new ground surface temperature . The ground sensible and water vapor fluxes are then updated for as
(2.5.140)¶
(2.5.141)¶
where and are the sensible heat and water vapor fluxes derived from equations and for nonvegetated surfaces and equations and for vegetated surfaces using . One further adjustment is made to and . If the soil moisture in the top snow/soil layer is not sufficient to support the updated ground evaporation, i.e., if and where
(2.5.142)¶
an adjustment is made to reduce the ground evaporation accordingly as
(2.5.143)¶
The term is the sum of over all evaporating PFTs where is the ground evaporation from the PFT on the column, is the relative area of the PFT with respect to the column, and is the number of PFTs on the column. and are the ice and liquid water contents (kg m^{2}) of the top snow/soil layer (Chapter 2.7). Any resulting energy deficit is assigned to sensible heat as
(2.5.144)¶
The ground water vapor flux is partitioned into evaporation of liquid water from snow/soil (kgm^{2} s^{1}), sublimation from snow/soil ice (kg m^{2} s^{1}), liquid dew on snow/soil (kg m^{2} s^{1}), or frost on snow/soil (kg m^{2} s^{1}) as
(2.5.145)¶
(2.5.146)¶
(2.5.147)¶
(2.5.148)¶
The loss or gain in snow mass due to , , , and on a snow surface are accounted for during the snow hydrology calculations (Chapter 2.8). The loss of soil and surface water due to is accounted for in the calculation of infiltration (section 2.7.2.3), while losses or gains due to , , and on a soil surface are accounted for following the subsurface drainage calculations (section 2.7.5).
The ground heat flux is calculated as
(2.5.149)¶
where is the solar radiation absorbed by the ground (section 2.4.1), is the net longwave radiation absorbed by the ground (section 2.4.2)
(2.5.150)¶
where
(2.5.151)¶
and and are the sensible and latent heat fluxes after the adjustments described above.
When converting ground water vapor flux to an energy flux, the term is arbitrarily assumed to be
(2.5.152)¶
where and are the latent heat of sublimation and vaporization, respectively (J (kg^{1}) (Table 2.2.7). When converting vegetation water vapor flux to an energy flux, is used.
The system balances energy as
(2.5.153)¶
2.5.5. Saturation Vapor Pressure¶
Saturation vapor pressure (Pa) and its derivative , as a function of temperature (ºC), are calculated from the eighthorder polynomial fits of Flatau et al. (1992)
(2.5.154)¶
(2.5.155)¶
where the coefficients for ice are valid for and the coefficients for water are valid for (Table 2.5.2 and Table 2.5.3). The saturated water vapor specific humidity and its derivative are
(2.5.156)¶
(2.5.157)¶
water 
ice 


6.11213476 
6.11123516 

4.44007856  5.03109514 

1.43064234  1.88369801 

2.64461437  4.20547422 

3.05903558  6.14396778 

1.96237241  6.02780717 

8.92344772  3.87940929 

3.73208410  1.49436277 

2.09339997  2.62655803 
water 
ice 


4.44017302 
5.03277922 

2.86064092 
3.77289173 

7.94683137 
1.26801703 

1.21211669 
2.49468427 

1.03354611 
3.13703411 

4.04125005 
2.57180651 

7.88037859 
1.33268878 

1.14596802 
3.94116744 

3.81294516 
4.98070196 