Model CSIRO: Elaborations
Model CSIRO is an entry in both the CMIP1 and CMIP2 intercomparisons.
The procedure for spinup/initialization to the simulation starting
of the CSIRO coupled model is as follows (reference: Gordon
and O'Farrell 1997):
- The atmosphere model was coupled to the dynamic sea ice model and
for many decades (without restoring forces toward climatology or
stress adjustments) to ensure equilibrium of the ice. The surface
fluxes from the final 10 years of the integration were averaged to
a monthly climatology for use in calculating the flux adjustments for
- The ocean model started from rest and with temperature and
each level specified from global average values after Levitus
(1982). It was forced at the surface with the climatological
wind stress data of Hellerman
and Rosenstein (1983). In the first phase, the ocean model
integrated for 1000 years with annual-mean forcing using asynchronous
stepping (2 days for T,S and 20 minutes for U,V); then it was
for an additional 100 years with annual mean forcing a reduced
(1 day for T,S and 20 minutes for U,V); then it was integrated for an
1070 years with monthly mean forcing and further-reduced asynchronicity
(1 day for T,S and 40 minutes for U,V); finally it was integrated for
years with monthly mean forcing using synchronous time stepping (1 hour
- Flux adjustments for heat, salinity, and surface stresses were
following the method of Sausen et al.
and the ocean model's SST and SSS were further corrected at each time
before being passed to the atmospheric model in order to reduce the
in coupled mode (cf. Gordon and
1997 for details).
- The atmosphere/sea ice and ocean models then were coupled and run
years with these flux adjustments and corrections. The resulting
climate drift was modest (~ -0.4 deg K in 2m surface air
and ~ -0.3 deg K in SST).
Land Surface Processes
The land surface scheme is after Kowalczyk
et al. (1991,1994).
- The scheme includes a single-layer vegetation canopy that acts as
or sink of water vapor and sensible heat. Monthly-varying values of
fractional vegetation type, unconstrained stomatal resistance,
length, and albedo are specified from data of Dorman
and Sellers (1989). The canopy shields the ground from solar
and intercepts precipitation and dew up to a maximum depth that depends
on the leaf area index (LAI); the intercepted moisture subsequently
at the potential rate. Following the
of Noilhan and Planton (1989),
evapotranspiration from the dry portion of the canopy is regulated by a
variable stomatal resistance that depends on the unconstrained value
by factors associated with incident radiation, air temperature,
of water in the root zone, and atmospheric vapor pressure deficit. Cf. Kowalczyk
et al. (1991,1994) for further details.
- Precipitation falls as rain and, if the surface air temperature
0 deg C, freezes at the surface with release of its latent heat. The
of snow modifies the roughness length and the radiative/thermal
of the surface. The latter depend on the fractional snow depth: the
albedo is assigned a maximum value of 0.8 when a snow scaling depth of
0.14 m is exceeded and the surface temperature is calculated using a
snow/soil layer. Changes in the total snow mass are calculated from a
- Soil temperature is calculated following McGregor
et al. (1993). Three layers are used, their thicknesses (0.03,
and 2.5 m) chosen to enable a reasonable representation of both diurnal
and seasonal temperature waves in the soil. The temperature of the i-th
layer is incremented by the diffusive heat flux into and out of the
Soil moisture is calculated using two layers, as described by Kowalczyk
et al. (1994). The chief source of soil moisture is surface
which depends on precipitation, snowmelt, surface runoff, and soil
properties; moisture sinks include evaporation from bare soil, root
associated with canopy transpiration, and
gravitational drainage when the local field capacity is exceeded.
runoff and gravitational drainage do not contribute to the flux of
to the ocean model, however.
- In a departure from common practice, the representation of sea
1998) is associated with the atmospheric component of the coupled
physics (radiation, clouds, moisture and heat budget, etc.) are
separately for ice and open-water fractions of each grid box, and then
are combined as an area-weighted average before computing the dynamics
(advection, mixing, etc.). The ice concentration and the percentage of
open water in each grid box are determined from the energy budget of
ocean mixed layer adjacent to the ice: once this water reaches the
freezing point (-1.85 deg C), new ice forms at increments of 4 percent
of the grid box. If ice advecting outside the existing area encounters
a mixed layer temperature close to the freezing point, the advected ice
is allowed to remain.
- Both vertical and lateral accretion of ice are
the existing ice floe is less than 0.25 m thick, new ice is added to
base; otherwise, ice is added to its side. Additional ice also is added
to the base of the floe when the ice concentration is greater than
in the Antarctic or 99.5% in the Arctic. In addition,
ice " (cf. Ledley 1985) forms from
overlying snow pack when it is massive enough to depress the floe
the ocean surface. Maximum ice thickness is constrained to 5 m,
only when ice is trapped in embayments where grid resolution
the calculation of ice velocities.
- Vertical and lateral ice ablation also are represented.
is calculated after the thermodynamic model of Semtner
(1976) which divides the ice system into three layers--two
ice and one for overlying snow, with constant albedos specified for the
two media; simulated brine pockets also provide an internal heat
No lateral melting occurs for a system temperature between -1.85 and
deg C. For a temperature between -1.0 and 0.5 deg C, any
heat is partitioned between the ocean mixed layer and lateral ice
All additional heat goes to lateral ice melt when the temperature is
than 0.5 deg C. Basal melt also results from ocean-to-ice heat
which is proportional to the difference between the surface
of the ocean and the ice bottom at the saltwater freezing
(The diffusion coefficient K = 1.5*10-6 m2 s-1
permits the ocean temperature to rise 1 to 2 degrees above the
freezing point without excessive melting of ice.) In the Antarctic an
fixed ocean-to-ice flux of 10 Wm-2 is applied to represent
higher transfer rates that occur during convection and the passage of
- The model also incorporates the cavitating fluid rheology of Flato
and Hibler (1990), wherein the sea ice has a finite
to compression, but can diverge freely. Convergence of ice,
with thickening by the formation of ridges, occurs when the internal
pressure is exceeded. Divergence increases the area of open water
in leads (i.e. reduces the ice compactness) rather than affecting the
thickness. The ice also is assumed to offer no resistance to
stress; its motion is determined by a dynamical balance between the
ice presssure, the atmospheric and oceanic surface stresses, and the
force. Ice advection is treated numerically with an upstream
- The ocean component of the coupled model is based on the
(cf. Cox 1984), and is described in Hirst
et al. (2000). Temperature/salinity grid points are coincident with
the physics grid of the atmospheric model, hence the model has a
resolution of 3.2° latitude by 5.2°
longitude. The model features three "islands" (Antarctica,
Guinea, New Zealand), as well as the major Eurasian-African-American
mass. There is no Bering Strait throughflow. Nonzero net circulation is
permitted around each island. The Mediterranean Sea and Hudson Bay are
included, and the effect of water exchange between these features and
North Atlantic is parameterized via a mixing of tracers across the
Strait of Gibraltar and Hudson Strait.
- The ocean component has 21 levels in the vertical, with level
from 25 m near the surface to 450 m in the deep ocean.
- Tracers (temperature, salinity) are diffused strongly along the
density surfaces, and weakly in the vertical. The isoneutral diffusion
is performed by the Cox (1987) implementation
the Redi (1982) scheme using a uniform
tracer diffusivity of 1* 103 m2 s-1.
vertical diffusivity is specified in terms of local static stability
to the formula of Gargett (1984), however
a minimum allowed default value of 3 * 10-5 m2 s-1.
Larger default minima of 2 * 10-3 m2 s-1
and 1 * 10-4 m2 s-1 for vertical
are set between the first and second, and second and third model
respectively, to roughly simulate some wind forced mixing effect.
is simulated by applying an enhanced vertical diffusivity, of 103
m2 s-1, in regions of static instability.
- The model includes the advective form of the Gent
and McWilliams (1990) scheme for eddy-induced transport of tracers
(cf. Gent et al. 1995). The isoneutral
diffusivity is set at 1 * 103 m2 s-1
270 m depth, above which this diffusivity is tapered linearly to zero
the surface. Horizontal tracer diffusivity is set to zero, as is common
in models employing the Gent and
- The horizontal viscosity is set at 9 * 105 m2
and the vertical viscosity is set at 2 * 10-3 m2
s-1, which are standard choices for the present resolution
cf. Bryan et al. 1975).
Details of the CMIP Integrations
- The CMIP 2 integration is performed using the same model version
CMIP 1, and is initialized from the model fields at the end of year 130
of the CMIP 1 simulation. Thus the 80 years of the transient CMIP 2
corresponds in time to the final 80 years of the period included in
1 from the control integration. The response of the model to the 1% per
annum compounding CO2 concentration in the CMIP 2 integration is
to that shown in Hirst et al. (1996)
(1999) for the same model version forced according to the IPCC
scenario. See also Gordon and
(1997) for a detailed discussion of an earlier 1%-compounding
experiment performed with the CSIRO model (prior to inclusion of the Gent
and McWilliams 1990 scheme).
Chief Differences from the Closest AMIP Model
In contrast to AMIP model CSIRO
CSIRO9 Mark 1 (R21 L9), the representation of atmospheric moisture
transport is semi-Lagrangian.
A dynamic/thermodynamic sea ice model after O'Farrell
(1998) replaces the formulation of the AMIP
Land Surface Processes
Land surface processes related
to hydrology and vegetation differ from those in the AMIP
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Princeton , N.J., 141 pp. [Available from Geophysics Fluid Dynamics
Princeton University, Princeton, NJ 08542.]
Cox, M.D., 1987: Isopycnal diffusion in a z-coordinate
ocean model. Ocean Modelling, 74, 1-15.
Dorman, J.L., and P.J.
1989: A global climatology of albedo, roughness length and stomatal
for atmospheric general circulation models as represented by the Simple
Biosphere Model (SiB). J. Appl. Meteor., 28, 833-855.
Flato, G.M., and W.D. Hibler
On a simple sea-ice dynamics model for climate studies. Ann.
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in the ocean interior. J. Mar. Res., 42, 359-393.
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Gent, P.R., J. Willebrand, T.J.
and J.C. McWilliams, 1995: Parameterising eddy-induced tracer
in ocean circulation models. J. Phys. Oceanogr., 25,
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Hellerman, S., and M.
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Hirst A.C., 1999: The Southern Ocean
to global warming in the CSIRO coupled ocean-atmosphere model. Environ.
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Hirst, A.C., S.P. O'Farrell and H.B.
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oceanic eddy-induced advection. Part I: Ocean Spinup and control
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Kowalczyk, E.A., J.R. Garratt,
P.B. Krummel, 1991: A soil-canopy scheme for use in a numerical model
the atmosphere--1 D stand-alone model. CSIRO Division of Atmospheric
Technical Paper No. 23, 56 pp.
Kowalczyk, E.A., J.R. Garratt,
P.B. Krummel, 1994: Implementation of a soil-canopy scheme into the
GCM-Regional aspects of the model response. CSIRO Division of
Research Technical Paper No. 32, 59 pp.
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McGregor, J.L., H.B. Gordon, I.G.
Watterson, M.R. Dix, and L.D. Rotstayn, 1993: The CSIRO 9-level
general circulation model. CSIRO Division of Atmosheric Research
Paper No. 26, 89 pp.
Noilhan, J., and S. Planton,
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model. J. Geophys. Res. (Oceans), 103, 15751-15782.
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Dyn., 2, 145-163.
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growth of sea ice in numerical investigations of climate. J.
Oceanogr., 6, 379-389.
CMIP Documentation Directory
Last update 15 May, 2002. This page is maintained by Tom