The UK Universities' Global Atmospheric Modelling Programme: Model UGAMP UGCM1.3 (T42 L19) 1993

AMIP Representative(s)

Dr. Mike Blackburn and Dr. Julia Slingo, Department of Meteorology, University of Reading, 2 Earley Gate, Whiteknights, PO Box 239, Reading RG6 2AU, England; e-mail: M.Blackburn@reading.ac.uk (Blackburn) and swssling@swssner1.rdg.ac.uk (Slingo); World Wide Web URL: http://ugamp.nerc.ac.uk .

Model Designation

UGAMP UGCM1.3 (T42 L19) 1993

Model Lineage

The UGAMP model is based on the ECMWF (cycle 27) model (cf. Tiedtke et al. 1988 [1] and Simmons et al. 1989 [2]), but with modifications principally in the treatment of radiation, convection, surface fluxes, vertical advection, and lateral and vertical dissipation.

Model Documentation

Documentation for the ECMWF(cycle 27) predecessor model is provided by Tiedtke et al. (1988) [1] . Subsequent modifications are described by Slingo et al. (1994) [3] and references therein.

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. The transform grid is sufficient to prevent aliasing of quadratic quantities, with 128 equispaced longitudes and 64 Gaussian latitudes. The full radiative calculations are performed on a reduced longitudinal grid, retaining only the first 16 Fourier modes (see Radiation).

Vertical Domain

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

Vertical Representation

Hybrid sigma-pressure coordinates after Simmons and Burridge (1981) [4] and Simmons and Strüfing (1981) [5]. To avoid oscillations in the profile of an advected quantity with rapidly changing gradient, vertical advection is treated by the Total Variation Diminishing (TVD) scheme of Thuburn (1993) [6].

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 a UNICOS environment.

Computational Performance

For the AMIP experiment, about 8 minutes Cray 2 computer time per simulated day (including data-archiving and storage time).

Initialization

For the AMIP experiment, the model atmosphere, soil moisture, and snow cover/depth were initialized from the ECMWF operational analysis for 12Z on 15 January 1987. These initial conditions were then designated as for 12Z on 15 December 1978, and the model was (partially) "spun up" to the AMIP start time by integrating it to a simulated state for 00Z on 1 January 1979 with prescribed (and fixed) AMIP sea surface temperatures for January 1979. See also Ocean.

Time Integration Scheme(s)

The time integration is by a semi-implicit Hoskins and Simmons (1975) [7] scheme with an Asselin (1972) [8] time filter. Advection of vorticity and moisture by a zonally symmetric flow is also treated implicitly. The time step is 30 minutes for dynamics and physics, except for full radiation/cloud calculations once every 3 hours (on a reduced grid, at every fourth point in longitude only--see Radiation). To ensure mass conservation, the global mean value of the logarithm of surface pressure is rescaled at each time step (but with mass sources/sinks associated with evaporation/precipitation neglected).

Smoothing/Filling

Orography is smoothed (see Orography). Negative values of atmospheric specific humidity (due to numerical 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 moisture which must be borrowed from the surface does not affect the hydrological 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, the logarithm of surface pressure, and specific humidity. Variations of the gas constant and specific heat capacity with water vapor content are also included.

Diffusion

Sixth-order (del^6) hyperdiffusion is applied in spectral space to vorticity, divergence, temperature, and moisture on the hybrid coordinate surfaces (see Vertical Representation). A correction is also applied to the temperature term to approximate dissipation on constant pressure surfaces. The diffusion time scale is 4 hours at the horizontal truncation limit (see Horizontal Resolution), but this is successively halved on the top four model levels, beginning at approximately 73 hPa.
Second-order vertical diffusion is applied below a hybrid model coordinate level of 0.650 to parameterize the PBL (see Planetary Boundary Layer). In addition, the TVD vertical advection scheme (see Vertical Representation) includes some dissipation of kinetic energy where sharp changes in gradient are encountered.

Gravity-wave Drag

Momentum transports associated with gravity waves are simulated by the method of Palmer et al. (1986) [9], using directionally dependent subgrid-scale orographic variances obtained from the U.S. Navy dataset (cf. Joseph 1980 [10] and see Orography). 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. See also Orography.

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. (The correct annual calendar is used, including Leap Years 1980, 1984, and 1988.)

Chemistry

Carbon dioxide concentration is the AMIP-prescribed value of 345 ppm. The specified ozone profile depends on pressure, total ozone in a column, the height of maximum concentration, latitude, longitude, and season. Total ozone is obtained from London et al. (1976) [11] data, and the altitude of maximum concentration from Wilcox and Belmont (1977) [12]. Mie radiative parameters of five types of aerosol are provided for (concentration depending only on height) 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 Morcrette (1989 [14], 1990 [15], 1991 [16]). Absorption by water vapor, ozone, carbon monoxide, carbon dioxide, methane, nitrous oxide, and oxygen is accounted for, with shortwave/longwave absorption coefficients calculated from line parameters of Rothman et al. (1983) [17].
For clear-sky conditions, shortwave radiation is modeled by a two-stream formulation in two spectral wavelength intervals (0.25 to 0.68 micron and 0.68 to 4.0 microns), using a photon path distribution method to separate the contributions of scattering and absorption processes to radiative transfer. Rayleigh scattering and Mie scattering/absorption by five aerosol types (see Chemistry) are treated by a delta-Eddington approximation.
The clear-sky longwave scheme employs a broad-band flux emissivity method in six spectral intervals between wavenumbers 0 and 2.6 x 10^5 m^-1, with continuum absorption by water vapor included between 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 are 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 also 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 in terms of the cloud liquid water path. For purposes of the radiation calculations, clouds of different types are treated as randomly overlapped in the vertical; convective cloud and the same type of nonconvective cloud in adjacent layers are treated as fully overlapped.
The full radiation calculations are performed every 3 hours on a reduced horizontal grid (every fourth point in longitude only), but with effective transmissivities and emissivities returned on the T42 Gaussian grid (see Horizontal Resolution). For intermediate time steps, the effective transmissivities are scaled by the instantaneous incoming solar radiation to represent correctly the diurnal cycle; the effective emissivities are scaled by the instantaneous Planck function to treat temperature variations. However, the influence of clouds remains fixed between full-radiation steps. See also Cloud Formation.

Convection

Convection follows the scheme of Betts and Miller (1993) [20], and consists of a relaxed convective adjustment towards calculated temperature and humidity reference profiles based on observations. The relaxation times are 2 hours for precipitating deep convection and 4 hours for nonprecipitating shallow convection, regarded as mutually exclusive processes. The convection is treated as shallow if the cloud top, defined by the level of nonbuoyancy, is below about 725 hPa for land or 810 hPa for ocean (for a surface pressure of 1000 hPa), or if there is insufficient moisture for precipitation to form with deep convection. Otherwise, the deep-convection scheme (including the possibility of midlevel convection with a cloud base above 725 hPa) is operative.
The temperature and humidity reference profiles for deep convection are based on relevant observational data (cf. Betts 1986 [21]). The temperature reference profile is a lapse rate that is slightly unstable with respect to the wet virtual adiabat below the freezing level, and that returns at cloud top to the moist adiabat of the cloud base. For energy conservation, this reference profile is corrected (with a second iteration) in order to remove the vertically integrated difference between the total moist enthalpy of the environment and that of the reference profile. The humidity reference profile is derived from the temperature reference by linearly interpolating between the humidities for specified values of subsaturation pressure deficit at cloud base, freezing level, and cloud top. Below the cloud base, cooling/drying by convective downdrafts is parameterized by specifying reference profiles for air parcels originating near 850 hPa that descend at constant subsaturation and equivalent potential temperature.
Nonprecipitating shallow convection is parameterized as a mixing of enthalpy and moisture of air below cloud base with air at and just above the capping inversion top. The reference profile is a mixing line structure joining the conserved saturation pressure and potential temperature points of all mixtures of the two sources of air (cf. Betts 1983 [22], 1986 [21]). Reference temperature and humidity profiles are computed after specifying a partial degree of mixing within the cloud, and mixing that is a function of the inversion strength at cloud top. Cf. Betts and Miller (1993) [20] and Slingo et al. (1994) [3] for further details. See also Cloud Formation and Precipitation.

Cloud Formation

Cloud formation is simulated following the diagnostic method of Slingo (1987) [23]. Clouds are of three types: shallow and deep convective cloud (see Convection); stratiform cloud associated with fronts/tropical disturbances that forms in low, middle, or high vertical layers; and low cloud associated with temperature inversions.
The fraction of shallow convective cloud (typically about 0.30) is related to the moisture tendencies within the cloud layer (cf. Betts and Miller 1993 [20]). The fraction of deep convective cloud (ranging between 0.20 to 0.80) is determined from the scaled convective precipitation rate (see Precipitation). If deep convective cloud forms above 400 hPa and the fractional area is > 0.4, anvil cirrus and shallow convective cloud also form.
Stratiform cloud is present only when the local relative humidity is > 80 percent, the amount being a quadratic function of this humidity excess. Low stratiform cloud is absent in regions of grid-scale subsidence, and the amount of low and middle stratiform cloud is reduced in dry downdrafts around subgrid-scale convective clouds. Low cloud forms below a temperature inversion if the relative humidity is > 60 percent, the cloud amount depending on this humidity excess and the inversion strength. See also Radiation for treatment of cloud-radiative interactions.

Precipitation

Precipitation is obtained from deep convection as part of the relaxed adjustment to the reference temperature and humidity profiles (see Convection). Subsequent evaporation of this precipitation is implicitly treated through inclusion of effects of convective downdrafts in the lowest three atmospheric layers. There is additional evaporation below elevated convective cloud bases that are situated above these downdraft layers.
In the absence of convective adjustment, precipitation also results from gridscale condensation when the local specific humidity exceeds the saturated value at the ambient temperature and pressure; the amount of precipitate depends on the new equilibrium specific humidity resulting from the accompanying latent heat release. Before falling to the surface, grid-scale precipitation must saturate all layers below the condensation level by evaporation. Melting of falling snow (see Snow Cover) occurs for air temperatures > +2 degrees C.

Planetary Boundary Layer

Vertical diffusion of momentum, heat, and moisture (proportional, respectively, to the vertical gradients of the wind, the dry static energy, and the specific humidity) is operative only below a hybrid-coordinate vertical level of 0.650 (about 650 hPa for a surface pressure of 1000 hPa). The vertically varying diffusion coefficient depends on stability (bulk Richardson number) and the vertical shear of the wind, following standard mixing-length theory (cf. Louis 1979 [33] and Louis et al. 1981 [34]). See also Diffusion, Surface Characteristics, and Surface Fluxes.

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 subgridscale orographic variance is added. The resulting "envelope orography" (cf. Wallace et al. 1983 [24]) is smoothed by applying a Gaussian filter with a 50 km radius of influence (cf. Brankovic and Van Maanen 1985 [25]). 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 and interpolated linearly in time at each time step. (These temperatures are uncorrected for nonzero surface heights associated with the spectral fitting of the topography--see Orography).

Sea Ice

AMIP monthly sea ice extents are prescribed, but ice surface temperatures are specified from the Alexander and Mobley (1976) [26] dataset. (Points with surface temperatures < -2 degrees C that are not on land are identified as sea ice; the masking procedure is described by Brugge 1993 [27].) Snow does not accumulate on sea ice.

Snow Cover

Grid-scale precipitation falls as snow if the temperature of the cloud layer is below 0 degrees C and that of intervening layers is below +2 degrees C (thereby inhibiting the melting of falling snow--see Precipitation). Snow depth (in meters of equivalent liquid water) is determined prognostically from a budget equation, but with accumulation only on land. The fractional area of a grid box covered by snow is given by the ratio of the snow depth to a critical depth (0.015 m), or is set to unity if the depth exceeds the critical value. Sublimation of snow is calculated as part of the surface evaporative flux (see Surface Fluxes), and snowmelt (occurring for ground temperatures > 0 degrees C) contributes to soil moisture (see Land Surface Processes). Snow cover also alters the surface albedo (see Surface Characteristics).

Surface Characteristics

The land surface is modeled as bare or with snow cover. Vegetation is not explicitly specified, but is accounted for in the prescribed surface properties described below.
Roughness length is prescribed as 1.0 x 10^-3 m over sea ice. Over open ocean the roughness is computed from the surface wind stress following Charnock (1955) [28], but it is constrained to be at least 1.5 x 10^-5 m. The roughness length over land is prescribed as a blended function of local orographic variance (Tibaldi and Geleyn 1981 [29]), vegetation (Baumgartner et al. 1977 [30]), and urbanization (from the U.S. Navy data set described by Joseph 1980 [10]) that is interpolated to the model grid; the logarithm of local roughness length is also smoothed by the same Gaussian filter used for orography (see Orography).
Annual means of satellite-observed surface albedo (range 0.07 to 0.80) from data of Preuss and Geleyn (1980) [31] and Geleyn and Preuss (1983) [32] are interpolated to the model grid and smoothed by the same Gaussian filter as used for orography (see Orography). Snow cover alters this background albedo, with a limiting value of 0.80 for snow depths > 0.01 m equivalent water. Sea ice albedo is prescribed as 0.55, and ocean albedo as 0.07. All albedos are also functions of solar zenith angle.
Longwave emissivity is prescribed as 0.996 for all surfaces. See also Sea Ice, Snow Cover, Surface Fluxes, and Land Surface Processes.

Surface Fluxes

Surface solar absorption is determined from surface albedos, and longwave emission from the Planck equation with prescribed constant surface emissivity (see Surface Characteristics).
Surface turbulent eddy fluxes are simulated as stability-dependent diffusive processes, following Monin-Obukhov similarity theory. Fluxes of momentum/heat/moisture are calculated from bulk formulae that include the product of a drag/transfer coefficient, the low-level wind speed, and the vertical difference between winds/dry static energy/specific humidity at the surface and their values at the lowest atmospheric level (996 hPa for a surface pressure of 1000 hPa). The low-level wind speed includes an imposed minimum of 3 m/s and an additional 3 m/s (added quadratically) in the presence of convection. (The former quantity increases surface fluxes in the limit of low wind speed, while the latter accounts for subgrid-scale convective circulations--cf. Slingo et al. 1994 [3] .) The surface drag/exchange coefficients are functions of stability (bulk Richardson number) and roughness length (see Surface Characteristics) following the formulation of Louis (1979 [33]) and Louis et al. (1981) [34]. The same transfer coefficient is used for the surface heat and moisture fluxes.
The surface moisture flux is also equivalent to the potential evaporation from a saturated surface multiplied by an evapotranspiration efficiency factor beta (cf. Budyko 1974 [35]). The factor beta is specified as unity over oceans and regions of dew formation (where the lowest atmospheric level is supersaturated); otherwise, beta varies with the snow cover and soil moisture content (see Snow Cover and Land Surface Processes).

Land Surface Processes

Soil temperature and moisture are determined by a model consisting of a surface layer 0.07 m thick, and middle and deep layers each of thickness 0.42 m. Temperature and moisture are prescribed from monthly climatologies in the deep layer (cf. Brankovic and Van Maanen 1985 [25] and Mintz and Serafini 1981 [36]), but vary prognostically in the surface/middle layers in response to diurnal and longer-period forcings.
Soil temperature is determined by simulating heat diffusion with an upper boundary condition specified by the net balance of surface energy fluxes (see Surface Fluxes). Soil heat capacity and diffusivity are prescribed constants: the density weighted heat capacity is 2.4 x 10^6 J/(m^3 K) and heat diffusivity is 7.5 x 10^-7 m^2/s).
Soil moisture also obeys a diffusion equation (with diffusivity one-seventh that of the heat diffusivity). The upper boundary condition is specified from the combined rainfall and snowmelt, and from surface evaporation that is reduced by the presence of (fractional) snow cover. Runoff occurs if the soil moisture exceeds the layer capacity (scaled according to thickness: 0.02 m for the surface layer and 0.12 m for each of the other layers). The evapotranspiration efficiency factor beta (see Surface Fluxes) is a composite of values determined for the snow-covered and bare-land fractions of a grid box. For snow-covered surfaces (see Snow Cover), beta is unity. Over bare land, beta is the ratio of the surface layer moisture to a prescribed fraction (0.75) of field capacity, but is constrained to be at most unity. There is also a temperature-dependent correction to account for limitation of evaporation due to lack of shortwave radiation.

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