State University of New York at Albany: Model SUNYA CCM1-TG (R15 L12) 1990

AMIP Representative(s)

Dr. Wei-Chyung Wang and Dr. Xin-Zhong Liang, Atmospheric Sciences Research Center, State University of New York at Albany, 100 Fuller Road, Albany, NewYork 12205; Phone: +1-518-442-3816; Fax: +1-518-442-3360; e-mail: wang@climate.asrc.albany.edu (Wang) and liang@climate.asrc.albany.edu (Liang); World Wide Web URL: http://www.atmos.albany.edu/das.html

Model Designation

SUNYA CCM1-TG (R15 L12) 1990

Model Lineage

The SUNYA model is identical to the standard version 1 of the NCAR Community Climate Model (CCM1), except for the addition of radiatively active trace gases other than carbon dioxide. The resulting modified CCM1 is designated as CCM1-TG (see Model Designation).

Model Documentation

Key documents for the standard CCM1 model are Williamson et al. (1987) [1], Kiehl et al. (1987) [2], and Bath et al. (1987a [3], b [4]) Wang et al. (1991a [5], b [6]) describe the treatment of radiative effects of trace gases that are added to CCM1.

Numerical/Computational Properties

Horizontal Representation

Spectral (spherical harmonic basis functions) with transformation to a Gaussian grid for calculation of nonlinear quantities and some physics.

Horizontal Resolution

Spectral rhomboidal 15 (R15), roughly equivalent to 4.5 x 7.5 degrees latitude-longitude.

Vertical Domain

Surface to 9 hPa; for a surface pressure of 1000 hPa, the lowest atmospheric level is at 991 hPa.

Vertical Representation

Sigma coordinates, with energy-conserving vertical finite-difference approximations (cf. Williamson 1983 [8], 1988 [9]).

Vertical Resolution

There are 12 unevenly spaced sigma levels. For a surface pressure of 1000 hPa, 3 levels are below 800 hPa and 5 levels are above 200 hPa.

Computer/Operating System

The AMIP simulation was run on a Cray 2 computer using a single processor in a UNICOS environment.

Computational Performance

For the AMIP experiment, about 1.2 minutes Cray 2 computer time per simulated day.

Initialization

For the AMIP experiment, initial conditions for the atmospheric state, soil moisture, and snow cover/depth were specified from the NCAR CCM1 model's standard January initial dataset (cf. Bath et al. 1987a [3]). The model then was "spun up" for 210 days in a perpetual January mode. The resulting climate state was then taken as the 1 January 1979 starting point for the AMIP simulation.

Time Integration Scheme(s)

Time integration is by a semi-implicit Hoskins and Simmons (1975) [27] scheme with an Asselin (1972) [7] frequency filter. The time step is 30 minutes for dynamics and physics, except for full (at all Gaussian grid points and vertical levels) radiation calculations which are done once every 12 hours (see Solar Constant/Cycles.

Smoothing/Filling

Orography is smoothed (see Orography). Negative values of atmospheric specific humidity (which arise because of numerical truncation errors in the discretized moisture equation) are filled by horizontal borrowing of moisture in a globally conserving manner. See also Convection.

Sampling Frequency

For the AMIP simulation, the model history is written once every 12 hours.

Dynamical/Physical Properties

Atmospheric Dynamics

Primitive-equation dynamics are expressed in terms of vorticity, divergence, potential temperature, specific humidity, and surface pressure.

Diffusion

Fourth-order (del^4) horizontal diffusion of vorticity, divergence, temperature, and specific humidity is computed locally on (approximately) constant pressure surfaces in grid-point space, except at stratospheric levels, where second-order (del^2) horizontal diffusion is applied (cf. Boville 1984 [10]).
Stability-dependent vertical diffusion is computed locally in grid-point space at all levels. Cf. Williamson et al. (1987) [1] for further details.

Gravity-wave Drag

Gravity-wave drag is not modeled.

Solar Constant/Cycles

The solar constant is the AMIP-prescribed value of 1365 W/(m^2). A seasonal, but not a diurnal cycle, in solar forcing is simulated.

Chemistry

The carbon dioxide concentration is the AMIP-prescribed value of 345 ppm. The vertical distribution of zonal-mean mixing ratios of ozone is specified from monthly data of Dütsch (1978) [11], updated by linear interpolation every 12 hours. The radiative effects of water vapor and oxygen, as well as methane, nitrous oxide, and chlorofluorocarbon compounds CFC-11 and CFC-12 also are included, but not those associated with aerosols (see Radiation).

Radiation

Shortwave radiation is treated in two spectral intervals--ultraviolet/visible (0.0 to 0.9 micron) and near-infrared (0.9 to 4.0 microns). Shortwave absorption by ozone, water vapor, carbon dioxide, and oxygen is modeled. Direct-beam absorption by water vapor is after the method of Kratz and Cess (1985) [12]; the reflected-beam absorption (as well as Rayleigh scattering by gases) follows Lacis and Hansen (1974) [13]. Oxygen absorption is treated as in Kiehl and Yamanouchi (1985) [14], and near-infrared absorption by carbon dioxide is after Sasamori et al. (1972) [15]. Gaseous absorption within clouds is included. Cloud albedo depends on optical depth and solar zenith angle, with multiple scattering effects included.
Longwave radiation is calculated in 5 spectral intervals (with wavenumber boundaries at 0.0, 5.0 x 10^4, 8.0 x 10^4, 1.0x10^5, 1.2 x 10^5, and 2.2 x 10^5 m^-1). Absorption/emission by water vapor (cf. Ramanathan and Downey 1986 [16]), carbon dioxide (cf. Kiehl and Briegleb 1991 [17]), and ozone (cf. Ramanathan and Dickinson 1979 [18]) is treated; the standard CCM1 radiation code is modified to include absorption/emission by methane, nitrous oxide, and chlorofluorocarbon compounds CFC-11 and CFC-12 (cf. Wang et al. 1991a [5], b [6]). The emissivity of nonconvective cloud is a function of diagnostic liquid water content. For purposes of the radiation calculations, cloud is treated as randomly overlapped in the vertical. Cf. Kiehl et al. (1987) [2] and Wang et al. (1991a [5], b [6]) for further details. See also Cloud Formation.

Convection

Moist convective adjustment after the method of Manabe et al. (1965) [19] performs several functions: removal of negative atmospheric moisture values (operating with a scheme for horizontal borrowing of moisture--see Smoothing/Filling); dry convective adjustment of unsaturated, unstable layers in the model stratosphere, with vertical mixing of moisture; and moist static adjustment of saturated unstable layers and of supersaturated stable layers.

Cloud Formation

Cloud forms in layers where the relative humidity exceeds 100 percent. If the vertical lapse rate of the layer also exceeds the moist adiabatic value, convective cloud forms (see Convection); otherwise, the cloud is nonconvective, and the fractional cloud cover is set to 0.95 in the layer. Convective cloud cover depends on the depth of the vertical instability, with the cloud amount in each layer adjusted so that the total fractional area is at most 0.30. If there is no associated precipitation (see Precipitation), a minimum convective cloud fraction of 0.01 is specified in each layer. Cloud is not allowed to form in the lowest model layer or in the top 3 layers, but clouds form together in the second and third layers above the surface if either of these layers is supersaturated. Cf. Kiehl et al. 1987 [2] for further details. See also Radiation for treatment of cloud-radiative interactions.

Precipitation

Precipitation results from application of convective adjustment (see Convection), if the vertical column is supersaturated with a lapse rate exceeding moist adiabatic. Precipitation also results if the column is supersaturated but with a stable lapse rate. There is no subsequent evaporation of precipitation before it falls to the surface.

Planetary Boundary Layer

The height of the PBL top is assumed to be that of the first level above the surface (sigma = 0.991), except for the calculation of a bulk Richardson number (see Surface Fluxes). In that case, the PBL top is computed from the temperature at the first sigma level but is constrained to be at least 500 m.

Orography

After interpolation of 1 x 1-degree Scripps Institution surface height data (cf. Gates and Nelson 1975 [20]) to the model grid, the data are smoothed using a Gaussian filter with 1.5-degree radius. The resulting heights are transformed into spectral space and truncated at the R15 model resolution. Cf. Pitcher et al. (1983) [21] for further details.

Ocean

AMIP monthly sea surface temperature fields are prescribed, with daily values determined by linear interpolation.

Sea Ice

Monthly AMIP sea ice extents are prescribed. The ice thickness is assumed to be a uniform 2 m, and the sub-ice ocean temperature is specified as a fixed 271.2 K. The surface temperature of the sea ice is computed prognostically by determining heat conduction from the underlying ocean through the ice, following the method of Holloway and Manabe (1971) [22]. Sea-ice surface temperature is constrained to be always < 0 degrees C (ice melting is not treated), and snow is not allowed to accumulate on the ice (see Snow Cover).

Snow Cover

Precipitation falls as snow if the temperatures of the surface and the first two atmospheric levels above it are all < 0 degrees C. Snow cover is determined from a combination of a monthly latitude-dependent climatology (cf. Bath et al. 1987a [3]) and prognostic snow accumulation (on land only) that is determined from a budget equation. A surface temperature > 0 degrees C triggers snowmelt, which augments soil moisture (see Land Surface Processes). Snow cover is also depleted by sublimation, which is calculated as part of the surface evaporative flux (see Surface Fluxes).

Surface Characteristics

The surface roughness length is specified as a uniform 0.25 m over land, sea ice, and snow cover, and as 1.0 x 10^-3 m over ocean.
Surface albedos for land surfaces are derived from the Matthews (1983) [23] 1 x 1-degree soil/vegetation dataset, but with distinguished vegetation types reduced to 10 and aggregated to the model resolution (see Horizontal Resolution). Land albedo also depends on solar zenith angle and spectral interval (ultraviolet/visible vs near-infrared--see Radiation). Snow cover alters the land albedo; the composite value is determined from equally weighted combinations of the local background albedo and that of the snow (which depends on surface temperature for the diffuse beam and on solar zenith angle for the direct beam). Over the ocean, surface albedos are prescribed to be 0.0244 for the direct-beam (with sun overhead) and 0.06 for the diffuse-beam component of radiation; the direct-beam albedo varies with solar zenith angle. The albedo of ice is a function of surface temperature. Cf. Briegleb et al. (1986) [24] for further details.
Longwave emissivities are set to unity (blackbody emission) for all surface types.

Surface Fluxes

Surface solar absorption is determined from the albedos, and longwave emission from the Planck equation with prescribed surface emissivity of 1.0 (see Surface Characteristics).
Surface fluxes of momentum, sensible heat and moisture are determined from bulk aerodynamic formulae, following the formulation of Deardorff (1972) [25]. Surface drag/exchange coefficients are a function of roughness lengths (see Surface Characteristics) and bulk Richardson number (see Planetary Boundary Layer). For computing these fluxes, the surface wind speed is constrained to be at least 1 m/s.
The surface moisture flux also depends on the evapotranspiration efficiency beta, which is unity over ocean, snow, and sea ice, but which over land is a function of soil moisture (see Land Surface Processes).

Land Surface Processes

Land surface temperature is determined from the balance of surface energy fluxes (see Surface Fluxes) by the diagnostic method of Holloway and Manabe (1971) [22]. (That is, there is no heat diffusion/ storage within the soil.)
Soil moisture is represented by the single-layer "bucket" model of Budyko (1956) [26] and Manabe (1969) [19], with field capacity a uniform 0.15 m of water. Soil moisture is increased by both precipitation and snowmelt. It is decreased by surface evaporation, which is determined from the product of the evapotranspiration efficiency beta and the potential evaporation from a surface saturated at the local surface temperature/pressure (see Surface Fluxes). Over land, beta is given by the ratio of local soil moisture to the field capacity, with runoff occurring implicitly if this ratio exceeds unity.

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