Model ECHAM3+LSG: Elaborations
Note: The ECHAM3 + LSG model was sponsored by the Deutsches Klimarechenzentrum
Model ECHAM3+LSG is an entry in both the CMIP1 and CMIP2 intercomparisons.
The spinup/initialization procedure for the CMIP I intercomparison experiment
was as follows (personal communication U. Cubasch):
The atmospheric model was integrated with prescribed climatological SSTs
and sea ice extents derived from the AMIP data until quasi-equilibrium
The ocean model was integrated first with prescribed Hellerman
and Rosenstein (1983) wind stresses, and was relaxed toward surface
salinity data of Levitus (1982), together with
a feedback to effective monthly mean surface air temperatures constructed
from the COADS data set (cf. Woodruff et
al. 1987). In a second stage, the diagnosed freshwater fluxes from
this run, together with the same surface wind stresses and surface air
temperatures then were used as external forcing (along with a mild relaxation
to AMIP SSTs). In a third stage, the ocean model was integrated further
with the forcings of the second stage, except for the additional application
of monthly mean values of the daily forcing anomalies. The total length
of the three stages was ~27,000 years.
The atmospheric model then was integrated again to quasi-equilibrium (a
20-year run), with SSTs and sea ice extents derived from the third stage
of the ocean-only run. Monthly mean flux corrections for heat, freshwater,
momentum (surface wind stresses) were computed following the method of
et al. (1988). The SSTs were also flux-corrected to adjust for the
mean bias between the AMIP and COADS climatologies.
The atmospheric and oceanic models then were coupled and integrated with
the application of these fixed monthly mean flux corrections. Evidence
of subsequent model drift was generally slight, except for the melting
of about 80% of the Antarctic sea ice between years 50 to 150 of the coupled
Land Surface Processes
Soil temperature is determined after Warrilow
et al. (1986) from the heat conduction in 5 layers (proceeding downward,
layer thicknesses are 0.065, 0.254, 0.913, 2.902, and 5.70 m), with net
surface heat fluxes as the upper boundary condition and zero heat flux
as the lower boundary condition at 10 m depth.
Snow pack temperature is also computed from the soil heat equation using
heat diffusivity/capacity for ice in regions of permanent continental ice,
and for bare soil where water-equivalent snow depth is <0.025 m. For
snow of greater depth, the temperature of the middle of the snow pack is
solved from an auxiliary heat conduction equation (cf. Bauer
et al. 1985). The temperature at the upper surface is determined by
extrapolation, but it is constrained not to exceed the snowmelt temperature
of 0 degrees C.
There are separate prognostic moisture budgets for snow, vegetation canopy,
and soil reservoirs. Snow cover is augmented by snowfall and is depleted
by sublimation and melting. Snow melts (augmenting soil moisture) if the
temperatures of the snow pack and of the uppermost soil layer exceed 0
degrees C. The canopy intercepts precipitation and snow (proportional to
the vegetated fraction of a grid box), which is then subject to immediate
evaporation or melting.
Soil moisture is represented as a single-layer "bucket" model (cf. Manabe
1969) with field capacity 0.20 m that is modified to account for vegetative
and orographic effects. Direct evaporation of soil moisture from bare soil
and from the wet vegetation canopy, as well as evapotranspiration via root
uptake, are modeled. Surface runoff includes effects of subgrid-scale variations
of field capacity related to the orographic variance; in addition, wherever
the soil is frozen, moisture contributes to surface runoff instead of soil
moisture. Deep runoff due to drainage processes also occurs independently
of infiltration if the soil moisture is between 5 and 9 percent of field
capacity (slow drainage), or is larger than 90 percent of field capacity
(fast drainage). When the model atmosphere is coupled to a dynamical ocean,
this source of freshwater is discharged at coastal points by means of a
river transport model that uses local runoff as input. Cf. Dümenil
and Todini (1992) and Sausen et al. (1994)
for further details.
Only the thermodyamics of sea ice are represented. Ice forms when a net
heat loss from the ocean would cause the SST to fall below the freezing
point. Existing ice thickens/melts in relation to the available heat
of fusion when the net heat flux is upward/downward. The heat flux is calculated
from a surface energy balance which includes a conductive heat flux between
the ocean and sea ice, based on the assumption of a linear temperature
profile. The surface albedo of the ice is assumed to remain constant;
solar penetration, brine pockets, or other internal heat capacities are
Snow cover is not represented and a modification of the heat flux by leads
or partial ice cover is not accounted for. Salinity changes are associated
with changes of the volume of the sea ice on the asumption of constant
Bauer, H., E. Heise, J. Pfaendtner, and
V. Renner, 1985: Development of an economical soil model for climate simulation.
In Current Issues in Climate Research (Proceedings of the EC Climatology
Programme Symposium, held 2-5 Oct. 1984, in Sophia Antipolis, France),
A. Ghazi and R. Fantechi (eds.), D. Reidel, Dordrecht, 219-226.
Cubasch, U., K. Hasselmann, H. Hock,
E. Maier-Reimer, U. Mikolajewicz, B.D. Santer, and R. Sausen, 1992:
Time-dependent greenhouse warming computations with a coupled ocean-atmosphere
model. Climate Dynamics, 10, 55-69.
Dümenil, L., and E. Todini,
1992: A rainfall-runoff scheme for use in the Hamburg climate model. In
in Theoretical Hydrology: A Tribute to James Dooge, J.P. O'Kane (ed.),
European Geophysical Society Series on Hydrological Sciences, Vol. 1, Elsevier
Press, Amsterdam, 129-157.
Hellerman, S., and M. Rosenstein,
1983: Normal monthly wind stress over the world ocean with error estimates.
Phys. Oceanogr., 13, 1093-1104.
Levitus, S., 1982: Climatological atlas of
the world's oceans. NOAA Professional Paper 13, 173 pp.
Manabe, S., 1969: Climate and ocean circulation.
1. The atmospheric circulation and the hydrology of the Earth's surface.
Wea. Rev., 97, 739-774.
Sausen, R., K. Barthels, and K. Hasselmann,
1988: Coupled ocean-atmosphere models with flux correction. Climate
Dyn., 2, 154-163
Sausen, R., S. Schubert, and L. Dümenil,
1994: A model of the river runoff for use in coupled atmosphere-ocean models.
Hydrology, 155, 337-352.
Warrilow, D.A., A.B. Sangster, and
A. Slingo, 1986: Modelling of land surface processes and their influence
on European climate. DCTN 38, Dynamical Climatology Branch, United Kingdom
Meteorological Office, Bracknell, Berkshire RG12 2SZ, UK.
Woodruff, S., R.J. Slutz, R.L. Jenne,
and P. Steurer, 1987: A comprehensive ocean-atmosphere dataset. Bull.
Amer. Meteor. Soc., 68, 1239-1250.
CMIP Documentation Directory
Last update 15 May, 2002. This page is maintained by Tom Phillips