PMIP Documentation for GISS-IIP
Goddard Institute for Space Studies: Model GISS Model IIp (4x5 L9) 1997
Mr. Richard J. Healy, Marine Chemistry & Geochemistry, MS 25, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA; Phone: +1-508 289 3514; Fax: +1-508 457 2193; e-mail:. firstname.lastname@example.org
World Wide Web URL: http://www.giss.nasa.gov
GISS Model IIp MY02 - (4x5L9) 1997 for 6fix
GISS Model IIp MY06M9 - (4x5L9) 1998 for 21fix
Number of days in each month: 31 28 31 30 31 30 31 31 30 31 30 31
Differences relatives to AMIP and Hansen et al 1997: corrected calculation
of insolation as a function of orbital forcing.
|infinite depth modified-Ekman layer, crude bulk formula for surface fluxes & exchange coeffients||finite modified-Ekman layer drag & mixing coefficents based on similarity theory||alters fluxes under unstable conditions; affects latitudinal precipitation distribution|
|Convection||50 percent of grid box mass elevated when moist static energy profile unstable||variable mass flux proportional to instability, entraining & non-entraining plume, down draft mass 1/3 updrafts||limits vertical mass exchange|
|Large Scale Clouds||diagnostic cloud parameterization as the saturated fraction of the grid box, optical thickness based on type, altitude, thickness||liquid water and ice prognostic variables, microphysical sources/sinks of cloud water, interactive optical thickness function of cloud particle size||atmospheric dynamics associated with cloud radiative forcing, positive feedback on low latitude warming|
|Land Surface||variable size two-layer bucket model specificed as a function of vegetation cover||physically-based calculation of evapotranspiration, evaporation of intercepted precipitation, dew, and over bare soil, infiltration, soil water movement, & runoff||more realistic partitioning of latent & sensible heating, stomatal control of transpiration|
|Heat and Moisture Advection||second-order-centered differencing scheme||quadratic upstream scheme||preserve strong gradients, alter eddy available potential energy & moisture/ radiative forcing|
|Momentum Advection||second-order B grid scheme||fourth-order B grid scheme||propagate waves faster affecting Intrahemispheric transport & high-frequency synoptic variations|
|Surface Radiative fluxes||radiative flux calculation for an average albedo of surface types in grid box||individual radiative flux calculations for albedo of each surface type in grid box|
Dim_longitude*dim_latitude : 72*46
For a surface pressure of 1000 hPa, 2 levels are below 800 hPa and 3 levels are above 200 hPa.
The 21fix simulation was run on a Silicon Graphics multiprocessor R10000 cpu / R10010 fpu.
CH4 and N2O concentrations are of 1643 ppb and 306 ppb for control, 650 ppb and 280 ppb for 6k, 350 ppb and 190 ppb for 21k.
The shortwave radiative fluxes are calculated after Lacis and Hansen (1974) , but with modifications to obtain accurate results at all solar zenith angles and optical thicknesses. Gaseous absorbers include water vapor, carbon dioxide, ozone, oxygen, nitrous oxide, and nitric oxide. Multiple scattering computations are made of 12 k-profiles, with strong line (exponential) absorption of the direct solar beam computed separately for water vapor, carbon dioxide, and oxygen. Absorption and scattering by aerosols are also included using radiative properties obtained from Mie calculations for the global aerosol climatology of Toon and Pollack (1976) . The spectral dependence of Mie parameters for clouds, aerosols, and Rayleigh scattering is specified in 6 intervals; these are superimposed on the 12 k-profiles to account for overlapping absorption.
In the longwave, the k-distribution method is used to model absorption by water vapor, carbon dioxide, ozone, nitrous oxide, nitric oxide, and methane (but with scattering effects neglected). A single k-distribution is specified for each gas, with 11 k-intervals for water vapor, 10 for carbon dioxide, and 4 for ozone.
Shortwave optical thickness of large-scale cloud is based on the prognostic cloud water path (see Cloud Formation). The required droplet effective radius is diagnosed from the cloud water content by assuming constant number concentration with different values for land/ocean cloud and liquid/ice cloud (cf. Del Genio et al. 1993) . The shortwave optical thickness of a convective cloud is proportional to its pressure depth. Cloud particle phase function and single-scattering albedo are functions of spectral interval, based on Mie computations for cloud droplet data of Squires (1958) and Hansen and Pollack (1970) . For purposes of the radiation calculations, partial cloud cover of a grid box is represented as full cloud cover that occurs for a percentage of the time (implemented via a random number generator--cf. Hansen et al. 1983) . Longwave effects of clouds are treated by an emissivity formulation, where the longwave cloud properties are self-consistent with the shortwave properties as a result of the application of Mie theory.
The amount of convective mass flux is obtained from a closure assumption
that the cloud base is restored to neutral buoyancy relative to the next
higher layer. The mass of the rising convective plume is changed by the
entrainment of drier environmental air, with associated decreases in buoyancy.
The entrainment rate is prescribed for an ensemble of two convective cloud
types (entraining and non-entraining). Heating/cooling of the environment
occurs through compensating environmental subsidence, detrainment of cloud
air at cloud top, a convective-scale downdraft whose mass flux detrains
into the cloud base layer, and evaporation of falling condensate (see Precipitation).
(Latent heat release serves only to maintain cloud buoyancy.) The convective
plume and subsiding environmental air transport gridscale horizontal momentum,
under the assumption that exchanged air parcels carry with them the momentum
of the layer of origin see Del Genio and Yao (1993).
The local convective cloud fraction is given by the ratio of convective mass flux to the total atmospheric mass of the grid box (see Convection and Atmospheric Dynamics). Large-scale clouds are predicted from a prognostic cloud water budget equation, where fractional cloudiness is an increasing function of relative humidity above a 60-percent threshold. (Upper-tropospheric convective condensate is also detrained into large-scale anvil cloud.) Cloud top entrainment instability is accounted for using a restrictive instability criterion. see Del Genio et al. (1996) for further details.
Evaporation of both the large-scale and convective condensate is modeled, with the residual falling to the surface as rain or snow (see Snow Cover). Evaporation of large-scale precipitation occurs in all unsaturated layers below its origin. The fraction of convective condensate that evaporates below cloud base is equated to the ratio of the convective mass flux to the total air mass of the grid box, while the fraction that evaporates above the cloud base is taken as half this value.
The topography used was based on Peltier?s modern and 21k topographic data. The procedure used to generate it took the difference between 21k and 0k and applied that to giss?s 0k which is based on the ETOPO5 data (5 minutes resolution). 106 meters are subtracted for the change in sea level between modern and 21k.
For 6 fix: SSTs and sea-ice prescribed at their present day value from the AMIP 10 year, as in the control run.
For 21 fix : The change in SST (LGM minus present-day) given by CLIMAP (1981) available at NGCD rather than the LGM absolute values in order to avoid differences due to differences in present day climatologies, was used. To obtain seasonally varying SSTs and sea ice edge from data for February and August, a simple sinusoidal variations, with extrems in February and August, is used.
The original 4x5 GISS CLIMAP LGM SSTs for Model II were generated from
the CLIMAP LGM SSTs at 1x1 degree resolution from a tape provided by LDGO.
In September 1995 we regridded the Model II data for the Model II? grid.
The difference between the Model II and the Model II? 4x5 grid is a 2.5
shift degrees in longitude.
For 21fix, the sea-ice edge of CLIMAP data is used for February and August.
Over land, surface roughness is a fit to the data of Fiedler and Panofsky (1972) as a function of the standard deviation of the orography (see Orography). The maximum of this roughness and that of the local vegetation (including a "zero plane displacement" value for tall vegetation types--cf. Monteith 1973 determines the roughness over land. Over sea ice, the roughness length is a uniform 4.3 x 10-4 m, after Doronin (1969) . Over ocean, the surface roughness is a function of the momentum flux and is used to compute the neutral drag coefficient as well as the Stanton and Dalton numbers (see Surface Fluxes).
The surface albedo is prescribed for ocean, land ice, sea ice, and for eight different land surface types after the data of Matthews (1983 , 1984 ). The visible and near infrared albedos are distinguished, and seasonal variations in the albedo for vegetated surfaces are included. The albedo of snow-covered ground is the snow-free value modified by factors depending on snow depth, snow age, and vegetation masking depth, which varies with vegetation type. Ocean albedo is a function of surface wind speed and solar zenith angle after Cox and Munk (1956) .
The spectral dependence of longwave emissivity for deserts is included from data of Hovis and Callahan (1966) , and for snow and ice from data of Wiscombe and Warren (1980) . The emissivity of the ocean is a function of the surface wind speed and of the albedo.
Soil properties are specified from Webb et al. (1993)
The surface wind stress is expressed as a product of air density, a drag coefficient, and the surface wind speed and velocity (see Planetary Boundary Layer). The surface drag coefficient is a function of both roughness length (see Surface Characteristics) and vertical stability.
Over land, the latent heat flux is computed separately for bare and vegetated surfaces (see Land Surface Processes), following Penman (1948) and Monteith (1981) . Over ocean and ice surfaces, the latent heat flux is expressed by a bulk formula that includes the product of the surface air density, a transfer coefficient, the surface wind speed, and the difference between the saturation value of mixing ratio at the ground temperature and the surface atmospheric value (see Planetary Boundary Layer).
Over land, the sensible heat flux is determined as the residual of the total net heat flux computed by the surface model (see Land Surface Processes) minus the latent heat flux. (The partitioning of the latent and sensible heat fluxes, or Bowen ratio, implies the surface temperature--see Land Surface Processes.) Over ocean and ice surfaces, the sensible heat flux is calculated from a bulk aerodynamic formula as a product of the surface air density and heat capacity, a transfer coefficient, a surface wind speed, and the difference between the skin temperature and the surface air temperature (see Planetary Boundary Layer). (The transfer coefficient is a stability-dependent function of the drag coefficient and is different from that used for the latent heat flux over ocean and ice.).
Land-surface hydrology is treated after the physically based model of Abramopoulos et al. (1988) . The scheme includes a vegetation canopy, a composite over each grid box from the vegetation types of Matthews (1983 , 1984 ), that intercepts precipitation and dew. Evaporation from the wet canopy and from bare soil is treated, as well as soil-moisture loss from transpiration according to moisture availability and variable vegetation resistance and root density. Diffusion of moisture is predicted in the six soil layers, accounting for spatially variable composite conductivities and matric potentials that depend on soil type and moisture content. Infiltration of precipitation and snowmelt is explicitly calculated, with surface runoff occurring when the uppermost soil layer is saturated; underground runoff that depends on topographic slope is also included. See Rosenzweig and Abramopoulos et al. (1997) for additional details.