Model LMD/IPSL1: Elaborations

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Participation

Spinup/Initialization

The CMIP I simulation starting point was the tenth year of a previous coupled model integration which was initialized as follows:

• The atmospheric model was started from 1 January 1980 of an (uncoupled) AMIP simulation which began on 1 January 1979.

• The ocean model was initialized from the Levitus (1982) temperature and salinity fields for climatological January.

• The atmospheric and ocean models were coupled without any further spinup and without application of flux adjustment (i.e., a "cold start").

Land Surface Processes

• Soil thermodynamics is determined by a 7-layer heat transfer model. The 7 layers are of uneven depths and are spaced between 0.02 m and 3.0 m below the surface, providing for resolution of thermal forcing at periods from 0.5 hour to 2 years. A zero-flux condition is imposed at the model's lower boundary, and the thermal insulation of snow is accounted for at its upper boundary. Introduction of the 7-layer model impacts ground temperature, snow mass/melt, and the seasonal change in prescribed vegetation that is tied to the soil temperature at a depth of 0.4 m. In turn, changes in snow cover and vegetation affect the albedo and roughness length over land. Cf. Polcher (1994) for further details.

• Soil hydrology is simulated using the land surface scheme SECHIBA (Schématisation des Echanges Hydriques à l' Interface entre la Biosphère et l'Atmosphère). The total depth of the soil column (corresponding to the vegetation root zone) is a constant 1.0 m. Soil moisture is computed in two layers, the upper layer being the most reactive: when precipitation exceeds evapotranspiration, the upper layer fills first; when the reverse is true, it empties first. Runoff occurs whenever the soil column is completely saturated (water depth 0.15 m). The remaining prescribed parameters for bare soil are a constant evaporative resistance and an empirical constant used to compute surface relative humidity for calculation of evaporation.

• In SECHIBA, each of the 7 prescribed vegetation classes interact individually with the soil hydrology and contribute individually to the surface evaporative flux. All the vegetation is assumed to have a single-story canopy that transpires or intercepts precipitation, but the canopy moisture capacity varies with the leaf area index, which is prescribed differently for each vegetation class. Different architectural and canopy resistances for evaporation/transpiration also are prescribed for each vegetation class. Cf Ducoudre et al. (1993) for further details. See also Surface Fluxes.

Sea Ice

• Sea ice extents are diagnostically determined, as described by Braconnot et. al (1997).  Ice forms in an ocean grid box with surface temperature less than the freezing point for salt water. The underlying ocean is assumed to supply the ice with a a constant heat flux of 2 W m^-2 in the Arctic and 4 W m^-2 in the Antarctic.

• The upper surface temperature of the ice is determined from the net atmospheric heat flux and the heat conduction from the ocean below, where the latter is inversely proportional to the ice thickness (assumed to be a constant 3 m).

• Sea ice dynamics and rheology are neglected.

Chief Differences from Closest AMIP Model

Radiation

The interaction of radiation and cloud has been subject to some tuning. As in the AMIP model, cloud optical thickness and emissivity are determined from the cloud liquid water path W (in kg m^-2), the effective radius r of cloud droplets (in m), and the absorption coefficient k (in m² kg^-1); however, the prognostic cloud liquid water content (LWC) (and therefore W) as well as r and k are determined differently:
 
• The LWC changes because of the introduction of a different parameterization of precipitation. Also, in place of the AMIP model's specification of r as a linear function of LWC, the effective droplet radius in the CMIP I experiment is prescribed as 10 microns for warm (water) clouds (with cloud-top temperatures > 0 C) and as 40 microns for cold (ice) clouds (with cloud-top temperatures < -25 deg C). (In the CMIP II experiments, r is set to somewhat different values.)
• The absorption constant k is also different for warm and cold clouds: 130 m² kg^-1 and 50 m² kg^-1, respectively, rather than a constant 130 m²kg^-1for all clouds, as in the AMIP model.
• In each grid box, warm clouds constitute a fraction X, and cold clouds a fraction 1-X of the total cloud cover, where X = 0 for temperatures below -25 deg C and X = 1 for temperatures above 0 deg C; X is assumed to vary linearly for intermediate temperatures. (Somewhat different temperature bounds are specified in the CMIP II experiments.)


Cf. Le Treut et al. (1994) for further details. See also Precipitation.

Precipitation

In place of the AMIP model's sharp distinction between precipitation in warm vs cold clouds, a smoother transition is specified in model LMD5.3:
 
• For warm (water) clouds, the precipitation rate is parameterized after Sundqvist (1981) as the product of a characteristic precipitation time scale T (value = 5.5 x 10^-4 s^-1), the prognostic cloud liquid water mixing ratio m, and an exponential function of (m/C)², where C is a prescribed precipitation threshold value = 2 x 10^-4 kg/kg.
• For cold (ice) clouds, the precipitation rate is determined by the ratio of m to a different timescale t = z/v, where z is the depth of the vertical layer and v is the terminal velocity of the water droplets, which is determined as an empirical function of m after Heymsfield and Donner (1990).
• The proportion of liquid and ice water mass within the clouds varies linearly between 0 and -25 deg C.


As in the AMIP model, evaporation of falling convective and large-scale precipitation is not explicitly modeled, but evaporation of small stratiform cloud droplets is simulated stochastically.

Surface Characteristics

The formulation of surface characteristics in model LMD5.3 is the same as in the AMIP model, except for the following features:
 
• Fractional sea ice and free ocean conditions may be specified within a single grid box, with roughness lengths and surface fluxes computed separately over ocean and ice. Roughness lengths over ice surfaces are a uniform 1.1 x 10^-3 m. Over ocean, roughness lengths are evaluated as a function of wind speed according to Charnock (1955).
• Over ocean, the surface albedo is specified according to Larson and Barkstrom (1977).

Surface Fluxes

In its treatment of surface fluxes, model LMD5.3 differs in the following respects from the AMIP model:
 
• The surface drag coefficient for the momentum, sensible heat, and evaporative fluxes depends on surface roughness and stability, following the approach of Louis (1979).

 
• In addition, over the continents, the surface evaporative flux is determined by the SECHIBA land surface scheme. The evaporative flux is calculated separately for each of the 8 coexisting surface types (bare ground plus 7 vegetation classes with fractional areas specified according to grid box) that are also present in the AMIP model. The total evaporative flux in each grid box then is computed as an area-weighted average of the individual fluxes. The total flux includes sublimation from snow, evaporation from bare soil and from moisture intercepted by the canopy of each vegetation class, and transpiration from the dry foliage of each class. Sublimation and evaporation from intercepted canopy moisture occur at the potential rate, while canopy transpiration and evaporation from bare soil depend on the surface relative humidity, which is parameterized in terms of soil moisture. Evaporation from sub-canopy soil is neglected.

• In SECHIBA also, the surface evaporative flux is computed by a bulk method that depends on the moisture gradient between the surface and the overlying air, and on resistances of different kinds (aerodynamic, soil, architectural, and canopy) that vary according to surface type and/or the nature of the moisture flux (sublimation, evaporation, transpiration). Cf. Ducoudre et al. (1993) for further details. See also Land Surface Processes.

Land Surface Processes

Model LMD5.3 includes the SECHIBA land surface scheme, which provides a more complex representation of land surface processes than that of the AMIP model.

References

Charnock, H., 1955: Wind stress on a water surface. Quart. J. Royal Meteor. Soc., 81, 639.

Ducoudre, N., K. Laval, A. Perrier, 1993: SECHIBA, a new set of parameterizations of the hydrologic exchanges at the land/atmosphere interface within the LMD atmospheric general circulation model. J. Climate, 6, 248-273.

Harzallah, A., and R. Sadourny, 1995: Internal versus SST forced atmospheric variability as simulated by an atmospheric general circulation model. J. Climate, 8, 474-498.

Heymsfield, A.J., and L.J. Donner, 1990: A scheme for parameterizing ice-cloud water content in general circulation models. J. Atmos. Sci., 47, 1865-1877.

Larson, J.C., and B.R. Barkstrom, 1977: Effects of realistic angular reflection laws for Earth's surface upon calculations of the earth-atmosphere albedo. In Proceedings of the Symposium on Radiation in the Atmosphere, Science Press, 451-453.

Le Treut, H., Z.-X. Li, and M. Forichon, 1994: Sensitivity of the LMD general circulation model to greenhouse forcings associated with two different cloud water parameterizations. J. Climate, 7, 1827-1841.

Levitus, S., 1982: Climatological atlas of the world's oceans. NOAA Professional Paper 13, 173 pp.

Louis, J.F., 1979: A parametric model of vertical eddy fluxes in the atmosphere. Bound. Layer Meteor., 17, 187-202.

Polcher, J., 1994: Etude de la sensibilité du climat tropical a la deforestation. Ph.D. dissertation, Université Pierre et Marie Curie, Paris, France.

Sundqvist, H, 1981: Prediction of stratiform clouds: Results of a 5-day forecast with a global model. Tellus, 33, 242-253.
 

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