the PMIP models (Bonfils et al. 1998)
PMIP Documentation for BMRC
Bureau of Meteorology Research Centre:
Model BMRC (R21 L9) Version 3.2 1993
Dr. Bryant McAvaney, Bureau of Meteorology Research Centre, Box 1289K,
GPO Melbourne, Victoria 3001,
Australia; Phone: +61-3-9669-4000; Fax: +61-3-9669-4660; e-mail: firstname.lastname@example.org
WWW URL: http://www.bom.gov.au/bmrc/clchhp.htm
BMRC 3.2 (R21 L9) 1993
Model Identification for PMIP
Number of days in each month: 31 28 31 30 31 30 31 31 30 31 30 31
The BMRC model is a descendant of a spectral general circulation model
first developed in the 1970s (cf. Bourke et al. 1977 , and McAvaney et
al. 1978 ). This version is more recent than that used for the first BMRC
AMIP experiment (Version 2.7)
The model configuration for the PMIP experiment is described by :
McAvaney B.J., and Colman R.A., 1993: The BMRC Model : AMIP Configuration.
BMRC Research Report No 38.
Colman R.A., and McAvaney B.J., 1995: JGR 100, 3155-3172.
McAvaney B.J., and G.D. Hess, 1996: The Revised Surface Fluxes Parametrisation
in BMRC: Formulation. BMRC Research Report No 56.
Key documentation of the BMRC model is provided by Bourke (1988) , Hart
et al. (1988 , 1990 ),
Colman and McAvaney (1991) , McAvaney et al. (1991) , and Rikus (1991)
Colman R.A., and McAvaney B.J., 1995: JGR 100, 3155-3172.
Spectral (spherical harmonic basis functions) with transformation to a
Gaussian grid for calculation of nonlinear quantities and physics.
Spectral rhomboidal 21 (R21), roughly equivalent to 3.2 x 5.6 degrees latitude-longitude.
Surface to about 9 hPa. For a surface pressure of 1000 hPa, the lowest
atmospheric level is at about 991 hPa.
Conservative finite differences in sigma coordinates.
There are 9 unevenly spaced sigma levels with the following levels : 0.991
0.926 0.811 0.664 0.500 0.336 0.189 0.074 0.009
For a surface pressure of 1000 hPa, 3 levels are below 800 hPa and 3
are above 200 hPa.
The PMIP simulation was run on a Cray Y/MP 3E computer using 1 processor
in a UNICOS 6.1.6 environment.
For the PMIP, about 3 minutes Cray Y/MP computation time per simulation
The control experiment was started from an AMIP simulation. For the AMIP
simulation, the model atmosphere is initialized from the ECMWF III-B analysis
for 12 UT of 1 January 1979, with nonlinear normal mode initialization
operative. Soil moisture is initialized from the January data of Mintz
and Serafini (1989) . Snow cover is initialized from the albedo data of
Hummel and Reck (1979) : albedos greater than 40 percent define areas of
seasonal snow with initial depth of 5 m; in areas of permanent snow (i.e.,
Antarctica and Greenland) the initial depth is set to 250 m.
Time Integration Scheme(s)
A semi-implicit scheme with an Asselin (1972) frequency filter is combined
with a split implicit scheme for the vertical diffusion component of the
model physics. A time step of 15 minutes is used for both dynamics and
physics, except that full calculations of radiative fluxes and heating
rates are done once every 3 hours.
Orography is smoothed (see Orography). Filling of negative atmospheric
moisture values is performed by a combination of local horizontal and vertical
borrowing, and global borrowing following the method of Royer (1986) .
A mass adjustment scheme is also used to prevent a slow drift in surface
pressure during long integrations. Cf. McAvaney et al. (1991) for further
For the PMIP simulation, the model history is written every 6 hours. (However,
fields such as convective and total precipitation are accumulated over
a 24-hour period; caution should therefore be exercised in interpreting
such fields at subintervals of a day.)
Primitive-equation dynamics are expressed in terms of vorticity, divergence,
temperature, surface pressure, and specific humidity.
Linear second-order (del-squared) horizontal diffusion is applied for wave
numbers n > 21 in the upper part of the spectral rhomboid, with a first-order
sigma coordinate correction applied near topography.
Stability dependent vertical diffusion after Louis (1979) is only applied
for sigma levels > 0.5 in stable layers, but it operates in all unstable
layers with no separate removal of dry superadiabats, and with a minimum
wind speed difference of 1 m/s assumed between model levels.
Momentum transports associated with gravity waves are simulated by the
method of Palmer et al. (1986) , using directionally dependent subgrid-scale
orographic variances. 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
The solar constant is the AMIP-prescribed value of 1365 W/(m2).
The orbital parameters and seasonal insolation distribution are calculated
after PMIP recommendations. Both seasonal and diurnal cycles in solar forcing
Carbon dioxide is assumed to be well mixed at the PMIP-prescribed concentration
of 345 and 280 ppm for control and 6ka. Zonally averaged seasonal mean
ozone distributions are prescribed from the data of Dopplick (1974), with
linear interpolation for intermediate times. No other trace gases or aerosols
are present, but the radiative effects of water vapour are included (see
Shortwave Rayleigh scattering and absorption in ultraviolet (< 0.35
micron) and visible (0.5-0.7 micron) spectral bands by ozone, and in the
near-infrared (0.7-4.0 microns) by water vapour and carbon dioxide follow
the method of Lacis and Hansen (1974) . Pressure corrections and multiple
reflections between clouds are treated. The radiative effects of aerosols
are not included directly.
Longwave radiation follows the simplified exchange method of Fels and
Schwarzkopf (1975) and Schwarzkopf and Fels (1991) , with slight modifications.
(The parent code is compared against benchmark computations by Fels et
al. 1991 .) Longwave calculations follow the broad-band emissivity approximation
in 8 spectral intervals (with wavenumber boundaries at 0, 1.6 x 104,
5.6 x 104, 8.0 x 104, 9.0 x 104, 9.9 x
104, 1.07 x 105, 1.20 x 105, and 2.20
x 105 m-1). Another 14 bands are accounted for in
the cooling-to-space corrections. Included in the calculations are Fels
and Schwarzkopf (1981) transmission coefficients for carbon dioxide, the
water vapour continuum of Roberts et al. (1976) , and the effects of water-carbon
dioxide overlap and of a Voigt line-shape correction. , and the effects
of water-carbon dioxide overlap and of a Voigt line-shape correction.
The treatment of cloud-radiative interactions is as described by Rikus
(1991) and McAvaney et al. (1991) . Shortwave cloud reflectivity/absorptivity
is prescribed for ultraviolet-visible and near-infrared spectral bands
and depends only on the height class of the cloud (see Cloud Formation).
In the longwave, all clouds are assumed to behave as black bodies (emissivity
of 1). For purposes of the radiation calculations, all clouds are assumed
to be randomly overlapped in the vertical.
Deep convection is simulated by the mass-flux scheme of Tiedtke (1989),
but without inclusion of momentum effects. The scheme accounts for midlevel
and penetrative convection, and also includes effects of cumulus-scale
downdrafts. The closure assumption for midlevel/penetrative convection
is that large-scale moisture convergence determines the bulk cloud mass
flux. That is, a condition of moisture balance is specfified in the layer
between the surface and cloud base wherby convergence by surface evaporation,
turbuence or large scale advection is equated to cloud base mass flux.
Entrainment and detrainment of mass in convective plumes occurs both through
turbulent exchange and organized inflow and outflow. Reference should be
made to Colman and McAvaney (1995) and McAvaney et al. (1995) for further
details on the consequences of the use of the Tiedtke scheme in the BMRC
Simulation of shallow convection is parameterized in terms of the model's
vertical diffusion scheme, following the method of Tiedtke (1983 , 1988
). Shallow convection is triggered when the lower layers are conditionally
unstable and have a relative humidity greater than 75%.
Stratiform cloud formation is based on the relative humidity diagnostic
form of Slingo (1987) . Clouds are of 3 height classes: high (sigma levels
0.189-0.336), middle (sigma levels 0.500-0.664), and low (sigma levels
0.811-0.926). The fractional amount of each type of cloud is determined
from a quadratic function of the difference between the maximum relative
humidity of the sigma layer and a threshold relative humidity that varies
with sigma level; for high and low cloud the threshold is 60 percent humidity,
while for middle cloud it is 50 percent. In addition, following Rikus (1991)
, low cloud forms when the relative humidity at the lowest atmospheric
level (sigma = 0.991) exceeds 60 percent, and is capped by strong static
stability in the layer immediately above (i.e., a temperature inversion
is present). In this case, the amount of low cloud increases with the strength
of the inversion. There is no parameterization of clouds associated with
convective towers. (See Convection for the treatment of convective cloud
and Radiation for cloud-radiative interactions.)
Precipitation from large-scale condensation occurs if the relative humidity
exceeds 100 percent. The convective precipitation rate is determined from
the Tiedtke (1989) convection scheme (see Convection). No evaporation of
precipitation is simulated. See also Snow Cover.
Planetary Boundary Layer
The height of the PBL is assumed to be that of the lowest prognostic vertical
level (sigma = 0.991). Winds, temperatures, and humidities for calculation
of turbulent eddy surface fluxes from bulk formulae are taken to be the
same values as those at this lowest atmospheric level (see Surface Fluxes).
See also Diffusion and Surface Characteristics.
Orography from a 1 x 1-degree U.S. Navy dataset is grid-point smoothed
using a Cressman (1959) area-averaged weighting function with a radius
of influence of 3 degrees for the spectral R21 model resolution (cf. Bourke
For control: PMIP datasets will be used for SSTs and sea-ice have been
prepared at PCMDI and were calculated by averaging the 10-year AMIP datasets
For 6 fix: SSTs and sea-ice prescribed at their present day value, as
in the control run.
For 0fix and 6fix, monthly AMIP sea ice extents are prescribed via a Cressman
(1959) weighting function with a 3-degree radius of influence; these monthly
ice extents are updated by interpolation every 5 days. The thickness of
the sea ice is held fixed at 1 m for the Antarctic region and 2 m for the
Arctic. Snow is permitted to accumulate or to melt on the ice surface,
but there is no conversion of snow to ice. The surface temperature of the
sea ice/snow is determined from a heat balance calculation (see Surface
Fluxes) with inclusion of a conduction term from the ocean (at a fixed
temperature of 271 K) below the ice.
When the weighted average of the air temperature at the lowest two levels
(sigma = 0.991 and 0.926) falls below 273.16 K, precipitation falls to
the surface as snow. Prognostic snow mass with accumulation and melting
over both land and sea ice is modelled. Snow cover affects the surface
albedo and the surface roughness (see Surface Characteristics), but there
is no explicit allowance for the effects of fractional snow cover. Melting
of snow, which occurs when the surface temperature exceeds 0 degrees C,
contributes to soil moisture (see Land Surface Processes), but sublimation
of snow is not calculated as part of the surface evaporative flux (see
Distinguished surface types include ocean, land, land ice, and sea ice,
and the presence of snow cover is also accounted for on the latter three
surfaces. Soil or vegetation types are not distinguished.
The roughness length over oceans is determined from the surface wind
stress, following Charnock (1955) , with a coefficient of 0.0185 assigned
after Wu (1982) ; the ocean roughness is constrained to a minimum value
of 1.5 x 10-5 m. Roughness lengths are prescribed uniform values
over sea ice (0.001 m) and land surfaces (0.168 m), but the presence of
snow cover changes the roughness to a new (fixed) value (0.001m).
Over oceans, the surface albedo depends on solar zenith angle, following
Payne (1972). Seasonal climatological surface albedos of Hummel and Reck
(1979) are prescribed over land. The surface albedos of sea ice and snow-covered
land follow the temperature-ramp formulation of Petzold (1977) , with different
values of albedo limits and a lower temperature range for sea ice and snow,
as described by Colman and McAvaney (1992)
Longwave emissivity is set to unity for all surfaces (i.e., black body
emission is assumed).
Surface solar absorption is determined from surface albedos, and longwave
emission from the Planck equation with prescribed constant surface emissivity
of 1.0 (see Surface Characteristics).
The surface turbulent eddy fluxes of momentum, heat, and moisture follow
Monin-Obukhov similarity theory, and are formulated in terms of bulk formulae
with stability-dependent drag/transfer coefficients determined as in Louis
(1979) . The momentum flux is given by the product of the air density,
a neutral drag coefficient, wind speed and wind vector at the lowest prognostic
level (sigma = 0.991), and a transfer function that depends on roughness
length (see Surface Characteristics) and stability (bulk Richardson number).
Surface wind speed is constrained to a minimum of 1 m/s. The flux of sensible
heat is given by a product of a neutral exchange coefficient, the wind
speed at the lowest prognostic level, the difference in temperatures between
the ground and the first prognostic atmospheric level, and a modified form
of the transfer function for unstable conditions (cf. Louis 1979) .
The flux of surface moisture is given by a product of the same transfer
coefficient and stability function as for sensible heat, an evapotranspiration
efficiency beta, and the difference between the specific humidity at the
first prognostic level and the saturation specific humidity at the surface
temperature and pressure. For calm conditions over the oceans, evaporation
also is enhanced following the approximation of Miller et al. (1992) for
the transfer coefficient. Over oceans, sea ice, and snow, beta is prescribed
to be unity; over land, beta is a function of the ratio of soil moisture
to a constant field capacity (see Land Surface Processes).
Land Surface Processes
Soil temperature is computed from heat storage in two layers with a climatological
temperature specified in a deeper layer. The upper boundary condition is
the surface energy balance (see Surface Fluxes). The heat conductivity
of soil is fixed under all conditions.
Prognostic soil moisture is represented by a single-layer "bucket" model
with uniform field capacity of 0.15 m after Manabe and Holloway (1975).
. Both precipitation and snowmelt contribute to soil moisture. The evapotranspiration
efficiency beta (see Surface Fluxes) is a function of the ratio of soil
moisture to the field capacity. Runoff occurs implicitly if this ratio
Last update November 9, 1998. For further information, contact: Céline