European Centre for Medium-Range Weather Forecasts: Model ECMWF ECMWF Cy36 (T42 L19) 1990

AMIP Representative(s)

Dr. Laura Ferranti and Dr. David Burridge, European Centre for Medium-Range Weather Forecasts; Shinfield Park, Reading RG29AX, England; Phone: +44-1734-499000; Fax: +44-1734-869450; e-mail: Laura.Ferranti@ecmwf.INT; World Wide Web URL: http://www.ecmwf.int/

Model Designation

ECMWF ECMWF Cy36 (T42 L19) 1990

Model Lineage

Cycle 36, one of a historical line of ECMWF model versions, first became operational in June 1990.

Model Documentation

Key documents for the model are ECMWF Research Department (1988 [1], 1991 [2]) and a series of Research Department memoranda from 1988 to 1990 that are summarized in ECMWF Technical Attachment (1993) [3].

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 triangular 42 (T42), roughly equivalent to 2.8 x 2.8 degrees latitude-longitude.

Vertical Domain

Surface to 10 hPa. For a surface pressure of 1000 hPa, the lowest atmospheric level is at about 996 hPa.

Vertical Representation

Finite differences in hybrid sigma-pressure coordinates after Simmons and Burridge (1981) [4] and Simmons and Strüfing (1981) [5].

Vertical Resolution

There are 19 irregularly spaced hybrid levels. For a surface pressure of 1000 hPa, 5 levels are below 800 hPa and 7 levels are above 200 hPa.

Computer/Operating System

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

Computational Performance

For the AMIP experiment, about 15 minutes of Cray 2 computation time per simulated day.

Initialization

For the AMIP simulation start date of 1 January 1979, the model atmosphere, soil moisture, snow cover/depth are initialized from ECMWF operational analyses for 15 January 1979 that are interpolated from spectral T106 resolution to T42 (see Horizontal Resolution).

Time Integration Scheme(s)

A semi-implicit Hoskins and Simmons (1975) [6] scheme with Asselin (1972) [7] frequency filter is used for the time integration, with a time step of 30 minutes for dynamics and physics, except for radiation/cloud calculations, which are done once every 3 hours.

Smoothing/Filling

Orography is smoothed (see Orography). Negative values of atmospheric specific humidity (due to truncation errors in the discretized moisture equation) are filled by borrowing moisture from successive vertical levels below until all specific humidity values in the column are nonnegative. Any borrowing from the surface that may be required does not impact the moisture budget there.

Sampling Frequency

For the AMIP simulation, the model history is written every 6 hours.

Dynamical/Physical Properties

Atmospheric Dynamics

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

Diffusion

Fourth-order (del^4) horizontal diffusion is applied in spectral space on hybrid vertical surfaces to vorticity, divergence, moisture, and on pressure surfaces to temperature.
Second-order vertical diffusion (K-closure) operates above the planetary boundary layer (PBL) only in conditions of static instability. In the PBL, vertical diffusion of momentum, heat, and moisture is proportional to the vertical gradients of the wind, specific humidity, and dry static energy, respectively (see Planetary Boundary Layer). The vertically variable diffusion coefficient depends on stability (bulk Richardson number) as well as the vertical shear of the wind, following standard mixing-length theory.

Gravity-wave Drag

Drag associated with orographic gravity waves is simulated after the method of Palmer et al. (1986) [8], as modified by Miller et al. (1989) [9], using directionally dependent subgrid-scale orographic variances obtained from the U.S. Navy dataset (cf. Joseph 1980) [10]. Surface stress due to gravity waves excited by stably stratified flow over irregular terrain is calculated from linear theory and dimensional considerations. Gravity-wave stress is a function of atmospheric density, low-level wind, and the Brunt-Vaisalla frequency. The vertical structure of the momentum flux induced by gravity waves is calculated from a local wave Richardson number, which describes the onset of turbulence due to convective instability and the turbulent breakdown approaching a critical level.

Solar Constant/Cycles

The solar constant is the AMIP-prescribed value of 1365 W/(m^2). Both seasonal and diurnal cycles in solar forcing are simulated.

Chemistry

The carbon dioxide concentration is the AMIP-prescribed value of 345 ppm. The ozone profile is determined from total ozone in a column (after data by London et al. 1976[11]) and the height of maximum concentration (after data by Wilcox and Belmont 1977) [12], and depends on pressure, latitude, longitude, and season. Mie radiative parameters of five types of aerosol (concentration depending only on height) are provided from WMO-ICSU (1984) [13] data. Radiative effects of water vapor, carbon monoxide, methane, nitrous oxide, and oxygen are also included (see Radiation).

Radiation

Atmospheric radiation is simulated after the method of Morcrette (1989 [14], 1990 [15], 1991 [16]). For clear-sky conditions, shortwave radiation is modeled by a two-stream formulation in spectral wavelength intervals 0.25-0.68 micron and 0.68-4.0 microns using a photon path distribution method to separate the effects of scattering and absorption processes. Shortwave absorption by water vapor, ozone, oxygen, carbon monoxide, methane, and nitrous oxide is included using line parameters of Rothman et al. (1983) [17]. Rayleigh scattering and Mie scattering/absorption by five aerosol types are treated by a delta-Eddington approximation.
The clear-sky longwave scheme employs a broad-band flux emissivity method in six spectral intervals from wavenumbers 0 to 2.6 x 10^5 m^-1, with continuum absorption by water vapor included from wavenumbers 3.5 x 10^4 to 1.25 x 10^5 m^-1. The temperature/pressure dependence of longwave gaseous absorption follows Morcrette et al. (1986) [18]. Aerosol absorption is also modeled by an emissivity formulation.
Shortwave scattering and absorption by cloud droplets is treated by a delta-Eddington approximation; radiative parameters include optical thickness, single-scattering albedo linked to cloud liquid water path, and prescribed asymmetry factor. Cloud types are distinguished by defining shortwave optical thickness as a function of effective droplet radius. Clouds are treated as graybodies in the longwave, with emissivity depending on cloud liquid water path after Stephens (1978) [19]. Longwave scattering by cloud droplets is neglected, and droplet absorption is modeled by an emissivity formulation from the cloud liquid water path. For purposes of the radiation calculations, clouds of different types are assumed to be randomly overlapped in the vertical, while convective cloud and nonconvective cloud of the same type in adjacent layers are treated as fully overlapped. See also Cloud Formation.

Convection

The mass-flux convective scheme of Tiedtke (1989) [20] accounts for midlevel and penetrative convection, and also includes effects of cumulus-scale downdrafts. Shallow (stratocumulus) convection is parameterized by means of an extension of the model's vertical diffusion scheme (cf. Tiedtke et al. 1988) [21]. The closure assumption for midlevel/penetrative convection is that large-scale moisture convergence determines the bulk cloud mass flux; for shallow convection, the mass flux is instead maintained by surface evaporation. Entrainment and detrainment of mass in convective plumes occurs both through turbulent exchange and organized inflow and outflow. Momentum transport by convective circulations is also included, following Schneider and Lindzen (1976) [22].

Cloud Formation

Cloud formation follows the diagnostic method of Slingo (1987) [23]. Clouds are of three types: shallow, midlevel, and high convective cloud; cloud associated with fronts/tropical disturbances that forms in low, medium, or high vertical layers; and low cloud associated with temperature inversions.
The height of midlevel/high convective cloud is determined by the level of non-buoyancy for moist adiabatic ascent (see Convection), and the cloud amount (fractional area 0.2-0.8) from the scaled logarithm of the convective precipitation rate. If this convective cloud forms above 400 hPa and the fractional area is > 0.4, anvil cirrus cloud also forms. Shallow convective cloud amount is determined from the difference between the moisture flux at cloud base and cloud top.
Frontal cloud is present only when the relative humidity is > 80 percent, the amount being a quadratic function of this humidity excess. Low frontal cloud is absent in regions of grid-scale subsidence, and the amount of low and middle frontal cloud is reduced in dry downdrafts around subgrid-scale convective clouds. In a temperature inversion, low cloud forms if the relative humidity is > 60 percent, the amount depending on this humidity excess and the inversion strength. See also Radiation for treatment of cloud-radiative interactions.

Precipitation

Freezing/melting processes in convective clouds are not considered. Conversion from cloud droplets to raindrops is proportional to the cloud liquid water content. No liquid water is stored in a convective cloud, and once detrained, it evaporates instantaneously with any portion not moistening the environment falling out as subgrid-scale convective precipitation. Evaporation of convective precipitation is parameterized (following Kessler 1969) [24] as a function of convective rain intensity and saturation deficit (difference between saturated specific humidity and that of environment).
Precipitation also results from gridscale condensation when the local specific humidity exceeds the saturated humidity at ambient temperature/pressure; the amount of precipitation depends on the new equilibrium specific humidity resulting from the accompanying latent heat release. Before falling to the surface, gridscale precipitation must saturate all layers below the condensation level by evaporation. See also Convection and Cloud Formation.

Planetary Boundary Layer

The PBL is represented typically by the first 5 vertical levels above the surface (at about 996, 983, 955, 909, and 846 hPa for a surface pressure of 1000 hPa, or at approximate elevations of 30 m, 150 m, 400 m, 850 m, and 1450 m, respectively). The PBL height is diagnostically determined as the greater of the height predicted from Ekman theory versus a convective height that depends on dry static energy in the vertical.

Orography

Orography is obtained from a U.S. Navy dataset (cf. Joseph 1980) [10] with resolution of 10 minutes arc on a latitude/longitude grid. The mean terrain heights are then calculated for a T106 Gaussian grid, and the square root of the corresponding subgrid-scale orographic variance is added. The resulting "envelope orography" (cf. Wallace et al. 1983) [25] is smoothed by application of a Gaussian filter with a 50 km radius of influence (cf. Brankovic and Van Maanen 1985) [26]. This filtered orography is then spectrally fitted and truncated at the T42 resolution of the model. See also Gravity-wave Drag.

Ocean

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

Sea Ice

AMIP monthly sea ice extents are prescribed. The surface temperature of the ice is specified from monthly climatologies. Snow is not allowed to accumulate on sea ice (see Snow Cover).

Snow Cover

Grid-scale precipitation may fall as snow if the temperature of the layer of its formation is <0 degrees C. Convective precipitation changes to snow only if the surface air temperature is <-3 degrees C, and over land only if the ground temperature is <0 degrees C. Snow depth (measured in meters of equivalent liquid water) is determined prognostically from a budget equation, with accumulation allowed only on land surfaces. The fractional area of snow coverage of a grid square is given by the ratio of the snow depth to a critical water-equivalent depth (0.015 m), or is set to unity if the snow depth exceeds this critical value. Sublimation of snow is calculated as part of the surface evaporative flux (see Surface Fluxes). Snow cover also alters the surface albedo (see Surface Characteristics) and the heat conductivity of the soil (see Land Surface Processes). Melting of snow (which contributes to soil moisture) occurs whenever the ground temperature exceeds +2 degrees C.

Surface Characteristics

The fractional area of vegetation (undistinguished by type) on each grid square is determined from Matthews (1983) [27] 1 x 1-degree data, as modified by Wilson and Henderson-Sellers (1985) [28]
The roughness length is prescribed as 1 x 10^-3 m over sea ice. It is computed over open ocean from the variable surface wind stress by the method of Charnock (1955) [29], but is constrained to be at least 1.5 x 10^-5 m. Over land, the roughness length is prescribed as a blended function of local orographic variance, vegetation, and urbanization (cf. Tibaldi and Geleyn 1981 [30], Baumgartner et al. 1977 [31], and Brankovic and Van Maanen 1985 [26]) that is interpolated to the model grid. The logarithm of local roughness length then is smoothed by the same Gaussian filter used for the orography (see Orography).
Annual means of satellite-observed surface albedo (range 0.07 to 0.80) from data of Preuss and Geleyn (1980) [32] and Geleyn and Preuss (1983) [33] are interpolated to the model grid and smoothed by the same Gaussian filter as for orography (see Orography). Snow cover alters this background albedo: snow albedo (maximum 0.80) varies depending on depth, masking by vegetation, temperature, and the presence of ice dew (see Snow Cover). Sea ice albedo is prescribed as 0.55, and ocean albedo as 0.07. Albedos do not depend on solar zenith angle or spectral interval.
Longwave emissivity is prescribed as 0.996 on all surfaces. Cf. ECMWF Research Department (1991) [2] for further details.

Surface Fluxes

Surface solar absorption is determined from surface albedo, and longwave emission from the Planck equation with prescribed constant surface emissivity (see Surface Characteristics).
Surface eddy fluxes of momentum, heat, and moisture are expressed as bulk formulae, following Monin-Obukhov similarity theory. The near-surface wind, temperature, and moisture required for the bulk formulae are taken to be the values at the lowest atmospheric level (at about 996 hPa for a surface pressure of 1000 hPa). The drag and transfer coefficients are functions of stability (bulk Richardson number) and roughness length (see Surface Characteristics), following the method of Louis (1979) [34] and Louis et al. (1981) [35], but with modifications by Miller et al. (1992) [36] for calm conditions over the oceans. The transfer coefficient for moisture is the same as that for heat.
The surface specific humidity over the ocean and snow-covered areas is the saturated value for the local surface temperature and pressure; over bare soil it is the product of the local saturated value and the surface relative humidity. The moisture flux over vegetation is given by the vertical difference of the specific humidity at the lowest atmospheric level and the saturated value at the surface temperature and pressure, all multiplied by an evapotranspiration efficiency factor beta (cf. Budyko 1974) [37]. This efficiency is the inverse sum of the aerodynamic resistance (surface drag) and the stomatal resistance, which depends on radiation stress, canopy moisture, and soil moisture stress in the vegetation root zone (cf. Sellers et al. 1986 [38], Blondin 1989 [41], and Blondin and Böttger 1987[39]). See also Land Surface Processes.

Land Surface Processes

Soil temperature and moisture are predicted in two layers of thicknesses 0.07 m and 0.42 m that overlie a deep layer (of thickness 0.42 m) in which temperature and moisture are prescribed from monthly climatologies (cf. Blondin and Böttger 1987 [39], Brankovic and Van Maanen 1985 [26], and Mintz and Serafini 1981 [40]). The upper boundary condition for the soil heat diffusion is the net surface energy balance (see Surface Fluxes). Soil heat capacity and diffusivity are functions of snow cover, and the diffusivity is also a function of vegetation canopy area.
The vegetation canopy also intercepts a fraction of the total precipitation (which is subject to potential evaporation) that would otherwise infiltrate the soil. The infiltrated soil moisture obeys a simple diffusion equation modified by gravitational effects (Darcy's Law), and is also affected by evaporation from the bare soil portion of each grid box as well as evapotranspiration by vegetation (see Surface Fluxes). Runoff occurs if the maximum soil moisture capacity of the surface layer (0.02 m) or middle layer (0.12 m) is exceeded; the fraction of infiltrated moisture associated with the surface runoff due to sloping terrain is also simulated using orographic variance data (see Orography).

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