Japan Meteorological Agency: Model JMA GSM9603 (T63 L30) 1998a

Contact Information

Experimental Implementation

Model Output Description

Model Characteristics

Contact Information

Modeling Group

AMIP Representative(s)

• Ken-ichi Kuma (kumaken@naps.kishou.go.jp), Hiroto Kitagawa (kitagawa@naps.kishou.go.jp) and Masayuki Kyouda (kyouda@naps.kishou.go.jp)
• Mail address: Japan Meteorological Agency, 1-3-4 Ote-machi, Chiyoda-ku, Tokyo 100-8122 Japan.
• Phone number: +81-3-3212-8341 (ext. 3315)
• Fax number: +81-3-3211-8407
• Web address: http://www.jma.go.jp/JMA_HP/jma/indexe.html

Experimental Implementation

Simulation Period

Earth Orbital Parameters

Calendar

Radiative Boundary Conditions

Ocean Surface Boundary Conditions

Orography/Land-Sea Mask

• The recommended U.S. Navy 10' x 10' orography data set (cf. Joseph 1980 [2]) is used.  This orography is spectrally truncated at the T63 horizontal resolution of the model (global-average orographic height is 237.11 m).

• Subgrid-scale orographic variances required for parameterization of gravity wave drag are also computed on the T63 Gaussian grid from the U.S. Navy data set (see Gravity Wave Drag).

• The land-sea mask of the model, derived from the Global Vegetation and Land-Use Data Bases (1x1 degree resolution) of Matthews (1983) [3], indicates whether each grid box is completely land or completely ocean.

Atmospheric Mass

Spinup/Initialization

• The recommended method described in Taylor et al. (1997) [4] is used for spin-up of the model to quasi-equilibrium: previous model solutions with supplementary (climatological and artificially generated) SST and sea-ice boundary conditions are used to initialize the model's atmosphere, snow cover/depth and soil moisture/temperature at the nominal start time of 00Z 1 January 1979.

• Following the AMIP II recommendation, there is no deep infinite soil moisture reservoir and the temperature of the deepest soil layer is not prescribed.

Computer/Operating System

Computational Performance

Model Output Description

Calculation of Standard Output Variables

• The recommended procedure of Boer (1986) [5] is used to calculate the percentage of time that a pressure surface is below ground.

• In order to avoid degradation of vertical profiles by interpolation, monthly mean tendencies and cloud properties are saved on the model levels (see Vertical Resolution).

• The following monthly mean tendencies are not supplied because the model does not treat them explicitly:

• Temperature tendency due to dry convection
• Eastward/northward momentum tendencies due to convection

 
• The recommended procedure of Hess et al. (1995) [6] is used to calculate surface variables (10 m winds, 2m specific humidity and 2m temperature).

• By assuming the lapse rate of temperature to be 0.005 K/m, mean sea-level pressure is calculated from the model orography, and the surface pressure and temperature of the lowest  atmospheric level.

• The recommended (method II) procedure of Potter et al. (1992) [7]  is used to calculate clear-sky radiation.

• Potential vorticity and planetary boundary layer height are not supplied.

Sampling Procedures

Interpolation Procedures

• As recommended, all upper-air standard output variables are vertically interpolated to the 17 WMO standard pressure surfaces every sampled time step.

• The algorithm used for treatment of variables on pressure surfaces below ground is as follows:

• For winds, vertical motion, and humidity, values on pressure surfaces below ground are set equal to those at the lowest atmospheric level.
• For temperature and geopotential height, values below ground are extrapolated from their values at the lowest atmospheric level by assuming the lapse rate of temperature to be 0.005 K/m.

Output Data Structure/Format/Compression

Model Characteristics

AMIP II Model Designation

Model Lineage

Model Documentation

Numerical/Computational Properties

Horizontal Representation

Horizontal Resolution

Spectral triangular 63 (T63), roughly equivalent to 1.875 x 1.875 degrees latitude-longitude.

Vertical Domain

Vertical Representation

Vertical Resolution

Time Integration Scheme(s)

• A leapfrog semi-implicit scheme with an Asselin (1972) [12] time filter is used for the time integration (cf. Jarraud et al. 1982 [13]). The condensation processes are not included in the leapfrog scheme, but are computed implicitly as adjustment terms.

• A variable time step length (about 15.9 minutes on the average for the AMIP II simulation) is used for dynamics and physics, except for radiation/cloud calculations. It is reset every 6 hours to satisfy the Courant-Friedrichs-Lewy (CFL) condition for the advection terms.

• Shortwave radiation and cloud properties are recalculated hourly, and longwave radiation every 3 hours (but with corrections made at every time step for diurnal variations in the shortwave fluxes and in the surface upward longwave fluxes).

 

Smoothing/Filling

• Orography is truncated at the model's spectral T63 horizontal resolution. (see Orography/Land-Sea Mask)

• Spurious negative values of atmospheric specific humidity (due to truncation errors in the discretized moisture equation) are reset to zero without any other correction to conserve the local or global moisture budgets.

 

 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 

Dynamical/Physical Properties

Equations of State

Diffusion

• The representation of horizontal diffusion is as follows:

• A linear fourth-order (del**4) horizontal diffusion is applied to vorticity, divergence, temperature and specific humidity on the hybrid sigma-pressure surfaces in spectral space.

 
• In order to reduce spurious mixing along steep mountain slopes, a first-order correction to approximate diffusion on constant pressure surfaces is also applied to the temperature and moisture equations

 
• Diffusion coefficients are chosen so that the enstrophy power spectrum coincides with that expected from two-dimensional turbulence theory.

• The representation of vertical diffusion is as follows:

• Within the PBL (see Planetary Boundary Layer), a stability-dependent local formulation of the vertical diffusion of momentum, heat and moisture follows the level-2 turbulence closure scheme of Mellor and Yamada (1974) [14].

 
• The vertically variable diffusion coefficient depends on stability (bulk Richardson number) after Blackadar (1962) [15], following standard mixing-length theory.

Gravity Wave Drag

• Orographic gravity wave drag is parameterized by two schemes that differ mainly in the vertical partitioning of the momentum drag, depending on the wavelength of the gravity waves (cf. Iwasaki et al. (1989) [18] for further details).

• Longwaves (wavelengths >100 km) are assumed to exert drag mainly in the stratosphere (type A scheme), and shortwaves (wavelengths approximately 10 km) only in the troposphere (type B scheme). In both schemes,  gravity wave stress is a function of atmospheric density, wind, the Brunt-Vaisalla frequency and subgrid-scale orographic variance obtained from the 10'x10' U.S. Navy data set (see Orography/Land-Sea Mask). For the type B scheme, orographic variance is calculated as an average difference of maximum and minimum heights within each 10'x10' grid box.

• Since the momentum drag (stress) due to shortwave gravity waves decreases with height as a result of nonhydrostatic effects (cf. Wurtele et al. 1987 [17]), the type B scheme assumes the wave stress to be quadratic in pressure and to vanish near the tropopause.  In the type A scheme, the vertical structure of the momentum stress induced by gravity waves is simulated after a modified Palmer et al. (1986) [16] amplitude saturation hypothesis.

Chemistry

• The carbon dioxide concentration is the AMIP II-specified value of 348 ppmv.

• The recommended monthly climatological zonal mean ozone profiles specified by Wang et al. (1995) [1] are linearly interpolated to obtain intermediate values once each day.

• Radiative effects of prognostic water vapor and globally uniform oxygen (at 0.20949 volume mixing ratio), but not of aerosols, are also included (see Radiation).

Radiation

• In order to save computation time, radiation/cloud properties are calculated on a grid with reduced horizontal resolution.

• Shortwave radiation is simulated by separate parameterizations for spectral wavelength intervals 0.20-0.70 micron (ultraviolet and visible) versus 0.70-5.00 microns (near-infrared). In the UV/visible, absorption by ozone, Rayleigh scattering by air molecules and Mie scattering by cloud droplets are treated after the method of Lacis and Hansen (1974) [20]. In the near-infrared, scattering/absorption are modeled by a two-stream formulation using the delta-Eddington approximation (Joseph et al. 1976 [21] and Coakley et al. 1983 [22]) which is applied in the 10 spectral intervals, as described by Briegleb (1992) [23]. (There are 7 intervals between 0.70 and 5.00 microns to capture oxygen A-band and water vapor absorption, and 3 intervals between 2.63 and 4.55 microns to capture carbon dioxide absorption). See also Surface Characteristics.

 
• Longwave radiation is treated by a broad-band flux emissivity method over 4 spectral bands with boundaries at wavenumbers of 4.0 x 10³, 5.5 x 10⁴, 8.0 x 10⁴,  1.2 x 10⁵ and 2.2 x 10⁵ m^-1.  Absorption due to water vapor, ozone and carbon dioxide are calculated from the transmission functions of Rodgers and Walshaw (1966) [24], Goldman and Kyle (1968) [25], and Houghton (1977) [26], respectively. Pressure broadening corrections of the absorption lines are also included. Continuum absorption due to water vapor is treated by the method of Roberts et al. (1976) [27]. Transmission in the 4 spectral bands includes overlapping effects of different absorbers. See also Surface Characteristics.

• Shortwave scattering and absorption due to clouds are treated by the delta-Eddington approximation. Cloud optical properties (optical depth, single-scattering albedo and asymmetry factor) are linked to diagnostic cloud water and effective droplet radius (specified differently for liquid droplets (Slingo 1989 [28]) and for ice particles (Ebert and Curry 1992 [29])). Clouds are treated as graybodies in the longwave, and cloud absorption is modeled by an emissivity formulation from cloud water path. For the radiation calculations, vertical overlap of clouds is assumed to be mixed (changing from full to random overlap with increasing vertical distance). See also Cloud Formation.

• Radiative interactions with aerosols are not treated.

Convection

• The vertical profile of upward mass flux is assumed to be a linear function of height, as proposed by Moorthi and Suarez (1992) [31]. The mass flux at the cloud base (fixed near 950 hPa in the model) is determined by the prognostic equation of Randall and Pan (1993) [32].

• The major contributions of the cumulus clouds are made through compensating downward motion, detrainment of mass, and reevaporation of the convective rainfall. See also Precipitation.

• The effects of an ensemble of multiple-type cumuli are considered. Each type is defined by the level of the cloud top, where the updraft cloud mass loses its buoyancy and detrainment occurs. During the upward movement of the cloud air mass, entrainment of environmental air is considered. The lower limit of the entrainment rate is proposed by Tokioka et al. (1988) [33].

• The convective downdraft associated with the cumulus clouds affects the environment by decreasing the net upward mass flux, and by the detrainment of the downdraft in the subcloud layer. The downdraft mass flux at the cloud base is specified so that the reevaporation of rainwater moistens the downward flux below the cloud base to saturation. In the extratropics, where moist convection is not always initiated at the surface, the mass flux is determined so that the large-scale moisture increase is spent by the convection. The effect of the cumulus convection on the large-scale tendency is calculated from a budget equation.

Cloud Formation

• There is no explicit determination of sub-gridscale convective cloud fraction.  Stratiform cloud fraction is diagnosed as a quadratic function of the excess of the layer relative humidity above a empirical threshold value that varies with height (cf. Saito and Baba 1988 [34]). Low-level (below 850 hPa) clouds associated with temperature inversions are treated.

 
• Cloud water content is parameterized as a function of temperature after Heymsfield (1977) [35]. All clouds form completely as liquid clouds if the temperature is higher than 273.15 K, and completely as ice clouds if the temperature is lower than 233.15 K. In between, the liquid/ice phase ratio of cloud water is linearly interpolated by temperature. The effective radius of a cloud liquid droplet is fixed at 15 microns, while the radius of an ice particle varies in the range 20-50 microns, according to the temperature. See also Radiation for the treatment of cloud-radiative interactions.

Precipitation

• Large-scale precipitation forms as a result of condensation when the predicted local specific humidity exceeds the saturation value at ambient temperature/pressure. The precipitation amount depends on the new equilibrium specific humidity, so as to account for the associated latent heat release. Moisture and temperature are mutually adjusted in three iterations.

• For convective precipitation, all the condensed water in the updraft is assumed to be carried to the cloud top. Here part of the condensed water is converted from cloud droplets to raindrops, while the remainder is evaporated, moistening the environment through the detrainment process. See also Convection.

• Subsequent evaporation of falling large-scale and convective precipitation is simulated following the formula proposed by Ogura and Takahashi (1971) [36]. The effect of the wind shear, which is responsible for the vertical tilting of the cumulus, is taken into account for the reevaporation process using the Fritsch and Chappel (1980) [37] relationship of precipitation efficiency and shear.

• Condensation is instantaneously converted to rain or snow (see Snow Cover). Melting processes of precipitation are considered when the mixed or ice phase precipitation encounters a model layer with temperature higher than 273.15 K.

Planetary Boundary Layer

Sea Ice

• The AMIP II monthly sea ice extents and concentrations are used as the model's sea ice, with intermediate daily values determined by linear interpolation (see Ocean Surface Boundary Conditions).  Because the model does not account for fractional sea ice, a threshold concentration of 50 % is used to determine whether a grid cell is covered by sea ice.

• Sea ice is assumed to have a constant depth of 2 m, and the ocean temperature below the ice is specified to be that for sea ice formation (about -2 degrees C). The surface temperature of the ice is prognostically determined by the force-restore method of Deardorff (1978) [38]. The forcing includes the net energy balance of surface heat fluxes as  well as the conduction heating from the ocean below.  Snow does not accumulate on sea ice.

Snow Cover

• Precipitation may fall as snow if the temperature at the lowest atmospheric level is less than the freezing point temperature of water (273.16 K). Snow depth (measured in meters of equivalent liquid water) is prognostically determined from a budget equation that accounts for accumulation (allowed only on land surfaces) and melting. Snow melt (which contributes to soil moisture) may occur if either the ground temperature exceeds the freezing point or the heat quantity of rainfall on snow exceeds the latent heat of fusion.

• A snow density value of 0.2 is used for conversion of snow mass to snow depth (water equivalent). Fractional snow coverage of a grid square without vegetation is given by the ratio of snow depth to a critical water-equivalent depth (0.02 m), or is set to unity if the snow depth exceeds this critical value. In a vegetation-covered grid square, critical depth varies with vegetation type and coverage.

• Sublimation of snow does not contribute to the total surface evaporative flux. Roughness of the surface decreases with increasing snow depth, the minimum value being 5 % of the snow-free value.  Snow cover also alters the surface albedo and the heat capacity/conductivity of the soil.  See also Surface Characteristics.

Surface Characteristics

• In the model, surface types are distinguished as ocean, sea ice, continental ice, bare ground (desert, half desert, tundra) and  vegetated ground, where the 8 vegetation types of the Simple Biosphere (SiB) model of Sellers et al. (1986) [39] are specified at monthly intervals.

• Over open ocean,  roughness lengths for the surface momentum flux are determined from variable surface wind stress after the method of Charnock (1955) [40], while those for surface heat and moisture fluxes are specified as a constant 1.52 x 10^-4 m (cf. Kondo 1975 [41]). Roughness length over sea ice is prescribed as a uniform 1 x 10^-3 m.  Spatially varying roughness lengths over land vary monthly according to the seasonal changes in vegetation (cf. Dorman and Sellers 1989 [42]) and are altered by snow cover.

• Wavelength-independent surface albedos over open ocean are prescribed to be 0.12347 for the direct-beam (with sun overhead) and 0.0419 for the diffuse-beam component of radiation; the direct-beam albedo varies with the solar zenith angle. Albedo of sea ice (independent of the solar zenith angle) is a constant 0.80 for the UV/visible and 0.40 for the near-infrared spectral intervals. Over land, albedo is a function of the solar zenith angle and is independent of the soil moisture; it is specified separately for the UV/visible and near-infrared spectral intervals. Snow-free land albedos vary monthly according to the seasonal changes in vegetation (cf. Dorman and Sellers 1989 [42]). Following Sellers et al. (1986) [39], land albedo is altered by snow cover; it is an average of the background albedo and the snow albedo, weighted by the fractional snow coverage. Snow albedo (maximum 0.80 for UV/visible and 0.40 for near-infrared) varies depending on the snow mass temperature (see Snow Cover).

• Longwave emissivity is prescribed to be unity (blackbody emission) for those sufaces without vegetation. There is graybody longwave emission (emissivity less than unity) from the vegetated surfaces.

Surface Fluxes

• For the treatment of surface radiative fluxes, albedo and longwave emissivity are determined for the distinguished surface types (see Surface Characteristics).

• Treatment of turbulent eddy fluxes of surface momentum, heat and moisture follows the Monin-Obukhov similarity theory implemented as bulk formulae. The required values of wind, temperature and humidity are taken to be the values at the lowest atmospheric level. Over land, surface moisture flux includes evapotranspiration from the dry vegetation (reflecting the presence of stomatal and canopy resistances) as well as direct evaporation from the wet canopy and bare soil. See also Land Surface Processes.

• Following Louis et al. (1981) [43], turbulent drag coefficients depend on the vertical stability and surface roughness (see Surface Characteristics). Whenever surface temperature and humidity are determined over land, the surface roughness of grassland is applied to calculations for all surface types.

Land Surface Processes

• To parameterize land surface processes, the Simple Biosphere (SiB) scheme of Sellers et al. (1986) [39] is used, as implemented by Sato et al. (1989a [44], b [45]).

• Vegetation in each grid box may consist both of the ground cover and upper-story canopy, with the spatial pattern of ground cover varying monthly. Precipitation interception by the canopy (with large-scale and convective precipitation distinguished) is included, and infiltration of moisture into the ground is limited to less than the local hydraulic conductivity of the soil. Direct evaporation from the wet canopy and from the bare soil is treated, and evapotranspiration from the dry leaves is also included by the detailed modeling of stomatal and canopy resistances. Within the canopy, evaporative fluxes are computed by the Penman-Monteith method (cf. Monteith 1973 [46]).  See also Snow Cover, Surface Characteristics, and Surface Fluxes.

 
• Soil heating is computed by the Deardorff (1978) [38] method of force-restore. Soil liquid moisture is computed from budget equations in three layers. This moisture is increased by the infiltrated precipitation and snowmelt, and is depleted by the direct evaporation and evapotranspiration.  Both surface runoff and subsurface runoff associated with gravitational drainage are treated.

AMIP I/AMIP II Model Differences

References

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