Institute of Atmospheric Physics (of the Chinese Academy of Sciences): Model IAP IAP-2L (4x5 L2) 1993
 and Zeng et al. (1987)  combined with physics similar to that of the Oregon State University model described by Ghan et al. (1982) . .  and Zeng et al. (1987)  is applied on a staggered C-grid (cf. Arakawa and Lamb 1977) . Vertical Representation and Vertical Resolution. Vertical Representation). For a surface pressure of 1000 hPa, these are at 800 hPa and 400 hPa (with the dynamical top at 200 hPa). Ocean and Sea Ice. . The pressure gradient force terms are also smoothed (cf. Schuman 1971)  to permit use of a longer time step, which is 6 minutes for dynamics, 30 minutes for diffusion, and 1 hour for physics (including radiation). The vertical flux of atmospheric moisture is also computed hourly, and it is recomputed if conditional instability of a computational kind occurs (cf. Arakawa 1972) , as evidenced by relative humidities in excess of 100 percent.
- Orography is smoothed (see
Orography). Poleward of 70 degrees
latitude, the wave-selected damping technique of Arakawa and Lamb (1977)
 and the Fast Fourier Transform algorithm of Lu (1986)
 are applied; from 38 to 70 degrees latitude,
the recursive operator of Fjortoft (1953)
is also used. In addition, a Shapiro (1970)
 smoothing operator is applied to
perturbation values of surface pressure, temperature, and water vapor mixing
ratio, and zonally on the wind field once per hour (cf. Liang 1986)
. A 9-point horizontal areal smoothing of the
lapse rate and a three dimensional smoothing of the diabatic heating are also
performed. See also Time Integration Scheme(s).
- Filling of spurious negative values of moisture is accomplished by application of the numerical scheme of Liang (1986)  to the advection of atmospheric water vapor.
- Nonlinear horizontal diffusion of heat, momentum, and moisture following
 and Washington and
 is applied on the
modified sigma surfaces (see Vertical Representation).
- There is no vertical diffusion as such, but momentum may be redistributed between the two atmospheric layers (see Vertical Resolution) by either eddy viscosity or convective friction. When the latter dominates, the friction coefficient depends on whether midlevel or penetrative convection occurs (see Convection). Cf. Zeng et al. (1989)  for further details.
- Shortwave radiation is calculated in ultraviolet (wavelengths <0.5
micron) and visible (wavelengths 0.5-0.9 micron) spectral intervals, employing
a delta-Eddington approximation (cf. Cess 1985  and Cess et al. 1985
). The shortwave calculations include
treatment of Rayleigh scattering, absorption by water vapor using the
exponential sum-fit method of Somerville et al. (1974)
, absorption by ozone following Cess and
 and Lacis and Hansen
 and scattering/absorption by cloud
droplets. The optical properties (single-scattering albedo, asymmetry factor,
and optical depth) of these droplets depend on cloud type (see
- Longwave absorption by carbon dioxide and water vapor, with empirical transmission functions after Katayama (1972) , is calculated for five spectral intervals with wavelengths >0.9 micron. Cirrus clouds or cirrus anvils on convective clouds (see Cloud Formation) are treated as graybodies (emissivity of 0.5, expressed as a modified fractional cloudiness), and other clouds as blackbodies (emissivity of 1.0). For purposes of the radiation calculations, blackbody clouds overlap fully in the vertical, but graybody clouds only partially. Cf. Zeng et al. (1989)  for further details.
- The parameterization of convection after Arakawa et al. (1969)
 includes these elements: dry convective
adjustment; shallow, midlevel, and penetrative convection; determination of the
cumulus mass flux; and modification of the planetary boundary layer (PBL). If
either atmospheric layer (see Vertical Resolution) is dry adiabatically
unstable, the potential temperatures are adjusted to a common value, such that
dry static energy is conserved. The layers that display evidence of moist
convective instability determine whether (mutually exclusive) shallow,
midlevel, or penetrative convection occur. Precipitation is associated with
midlevel and penetrative convection (see
Precipitation), and changes in
PBL temperature and humidity are associated with shallow and penetrative
convection (see Planetary Boundary Layer). Convective clouds are assumed
to form a steady-state ensemble, with cloud air being saturated; the vertical
profile of cloud temperature is computed assuming that moist static energy is
- Cumulus mass flux is computed (differently) for midlevel and penetrative convection, assuming that the convective instability is removed with an e-folding time of 1 hour. (Cumulus mass flux is not associated with shallow convection in the model.) The mass flux from the PBL and that entrained from the free atmosphere are mixed in the lower part of the convective cloud, with detrainment of this mixture at the level of nonbuoyancy. The magnitude of the mass flux determines the amount of cumulus friction (with different coefficients for midlevel and penetrative convection), which brings about momentum transfer between the two atmospheric layers (see Diffusion). Cf. Zeng et al. (1989)  for further details.
- Clouds may result either from large-scale condensation or from convection
(see Convection). Large-scale clouds form when the relative humidity
exceeds 90 percent in the lower vertical layer or 100 percent in the upper
layer (see Vertical Resolution). Clouds are assumed to fill a grid box
(cloud fraction = 1). (For purposes of longwave radiation calculations,
however, graybody cirrus and cirrus anvil cloud fractions are 0.5.)
- Four basic cloud types are modeled. Type 1 is formed by either midlevel or
penetrative convection, type 2 when the relative humidity of the lower vertical
layer exceeds 90 percent, type 3 as a result of shallow convection, and type 4
when large-scale precipitation occurs in the upper vertical layer (see
Precipitation). (Precipitation is limited to cloud types 1, 2, and 4.)
Types 2, 3, or 4 are defined as graybody cirrus clouds if they are
nonprecipitating or if the average of their base and top temperatures is
<-40 degrees C; convective cloud type 1 is also capped by a graybody cirrus
anvil if it satisfies this temperature criterion.
- In addition, a cloud type 5 is formed by the coexistence of types 2 and 4, and a cloud type 6 by the coexistence of types 3 and 4. (Types 1 and 3 cannot coexist; type 1 also overrides the formation of types 2 and 4, while type 2 overrides type 3.) Both types 5 and 6 are treated as low-level clouds for radiation purposes. Cf. Zeng et al. (1989)  for further details. See also Radiation.
- Precipitation may result from midlevel or penetrative convection (see
Convection). The amount of precipitation is equal to the net water vapor
entrained from the environment, and falls at a rate that is a function of the
cumulus mass flux. There is no subsequent evaporation of convective
- Precipitation also results from large-scale condensation. When the upper atmospheric layer becomes supersaturated, the resulting precipitation is assumed to evaporate completely in passing through the lower layer, thereby cooling and moistening the environment in proportion to the amount of evaporation (cf. Lowe and Ficke 1974) . If the lower layer then becomes supersaturated, the resulting precipitation falls to the surface.
- Roughness lengths are not specified, since these are not required for the
calculation of turbulent fluxes (see Surface Fluxes).
- The surface albedo is prescribed for water, land ice, sea ice, and for six
land surface types. Following Manabe and Holloway (1975)
, the albedo of snow-covered surfaces varies
as the square root of snow mass up to a maximum value that is assigned for a
critical snow mass of 10 kg/(m^2) (see Snow Cover). Albedos for
the land surfaces (with and without snow cover) are assigned from data of Posey
and Clapp (1964)
. Over water, the albedo
for diffuse solar flux is taken as 0.07, and that for the direct beam is a
function of solar zenith angle (cf. Zeng et al. 1989)
- Longwave emissivity is prescribed as unity (blackbody emission) for all surface types.
- The absorbed surface solar flux is determined from the surface albedo, and
surface longwave emission from the Planck function with constant surface
emissivity of 1.0 (see Surface Characteristics).
- Turbulent surface fluxes of momentum, heat, and moisture are calculated
from bulk aerodynamic formulae. The momentum flux is proportional to the
product of a drag coefficient and the effective surface wind, which is
determined by extrapolating the wind at the two vertical levels and multiplying
by a factor of 0.7 (but constrained to be at least 2 m/s in
magnitude). The drag coefficient is not a function of vertical stability but it
does depend on surface elevation. Over the oceans, the drag coefficient is a
function of the surface wind speed, but it is constrained to be at most 2.5 x
- The surface heat and moisture fluxes also depend on a product of the same drag coefficient and effective surface wind speed, as well as on the difference between the ground and surface atmospheric temperatures (for heat fluxes) or specific humidities (for moisture fluxes). These surface values of atmospheric temperature and humidity are determined by equating, respectively, the surface sensible heat and evaporative flux to corresponding fluxes at the top of the constant-flux surface layer. The latter are parameterized following K-theory (cf. Arakawa 1972) , where the eddy diffusivity depends on vertical stability but is constrained to be < 15 m^2/s over water surfaces, and to be <100 m^2/s elsewhere. The surface moisture flux also depends on an evapotranspiration efficiency factor beta, which is taken as unity over snow, ice, and water, and in areas of dew formation. Over land, beta is a function of soil moisture. Cf. Zeng et al. (1989)  for further details. See also Diffusion, Planetary Boundary Layer, and Land Surface Processes.
- Following Priestly (1959)
 the average ground
temperature at the diurnal skin depth is computed from the net surface energy
fluxes (see Surface Fluxes), taking account of the thermal conductivity,
and the volumetric and bulk heat capacities of snow, ice, and soil.
- Soil moisture is predicted as a fraction of a uniform field capacity of 0.15 m in a single layer (i.e., a "bucket" model). Fractional soil moisture is determined from a budget that includes the rates of precipitation, snowmelt, surface evaporation, and runoff. The evapotranspiration efficiency beta over land (see Surface Fluxes) is specified as the lesser of twice the fractional soil moisture or unity. Runoff is given by the product of the fractional soil moisture and the sum of precipitation and snowmelt rates. If the predicted fractional soil moisture is in excess of unity, the excess is taken as additional runoff. Cf. Zeng et al. (1989)  for further details.
Last update April 19, 1996. For further information, contact: Tom Phillips ( email@example.com )