## Abstract

A simplified climate model is presented which includes a fully 3-D, frictional geostrophic (FG) ocean component but retains an integration efficiency considerably greater than extant climate models with 3-D, primitive-equation ocean representations (20 kyears of integration can be completed in about a day on a PC). The model also includes an Energy and Moisture Balance atmosphere and a dynamic and thermodynamic sea-ice model. Using a semi-random ensemble of 1,000 simulations, we address both the inverse problem of parameter estimation, and the direct problem of quantifying the uncertainty due to mixing and transport parameters. Our results represent a first attempt at tuning a 3-D climate model by a strictly defined procedure, which nevertheless considers the whole of the appropriate parameter space. Model estimates of meridional overturning and Atlantic heat transport are well reproduced, while errors are reduced only moderately by a doubling of resolution. Model parameters are only weakly constrained by data, while strong correlations between mean error and parameter values are mostly found to be an artefact of single-parameter studies, not indicative of global model behaviour. Single-parameter sensitivity studies can therefore be misleading. Given a single, illustrative scenario of CO_{2} increase and fixing the polynomial coefficients governing the extremely simple radiation parameterisation, the spread of model predictions for global mean warming due solely to the transport parameters is around one degree after 100 years forcing, although in a typical 4,000-year ensemble-member simulation, the peak rate of warming in the deep Pacific occurs 400 years after the onset of the forcing. The corresponding uncertainty in Atlantic overturning after 100 years is around 5 Sv, with a small, but non-negligible, probability of a collapse in the long term.

This is a preview of subscription content, access via your institution.

## References

Annan JD, Hargreaves JC, Edwards NR, Marsh R (2005) Parameter estimation in an intermediate complexity earth system model using an ensemble Kalman filter. Ocean Model 8:135–154

Boville BA, Gent PR (1998) The NCAR climate system model, version one. J Clim 11:1115–1130

Edwards NR, Shepherd JG (2001) Multiple thermohaline states due to variable diffusivity in a hierarchy of simple models. Ocean Model 3:67–94

Edwards NR, Shepherd JG (2002) Bifurcations of the thermohaline circulation in a simplified three-dimensional model of the world ocean and the effects of interbasin connectivity. Clim Dyn 19:31–42

Edwards NR, Willmott AJ, Killworth PD (1998) On the role of topography and wind stress on the stability of the thermohaline circulation. J Phys Oceanogr 28:756–778

Goosse H, Selten FM, Haarsma RJ, Opsteegh JD (2001) Decadal variability in high northern latitudes as simulated by an intermediate-complexity climate model. Ann Glaciol 33:525–532

Gordon C, Cooper C, Senior CA, Banks H, Gregory JM, Johns TC, Mitchell JFB, Wood RA (2000) The simulation of SST, sea-ice extents and ocean heat transports in a version of the Hadley Centre coupled model without flux adjustments. Clim Dyn 16:147–168

Griffies SM (1998) The Gent-McWilliams skew flux. J Phys Oceanogr 28:831–841

Hall MM, Bryden HL (1982) Direct estimates and mechanisms of ocean heat transport. Deep-Sea Res 29:339–359

Hargreaves JC, Annan JD, Edwards NR, Marsh R (2005) Climate forecasting using an intermediate complexity Earth System Model and the Ensemble Kalman Filter. Clim Dyn (in press).

Hibler WD (1979) Dynamic thermodynamic sea ice model. J Phys Oceanogr 9:815–846

Hogg AMcC, Dewar WK, Killworth PD, Blundell JR (2003) A quasi-geostrophic coupled model: Q-GCM. Mon Weather Rev 131:2261–2278

Holland DA, Mysak LA, Manak DK (1993) Sensitivity study of a dynamic thermodynamic sea ice model. J Geophys Res 97:5365–2586

Jia Y (2003) Ocean heat transport and its relationship to ocean circulation in the CMIP coupled models. Clim Dyn 20:153–174

Josey SA, Kent EC, Taylor PK (1998) The Southampton Oceanography Centre (SOC) Ocean-Atmosphere Heat, Momentum and Freshwater Flux Atlas. Southampton Oceanography Centre Rep. 6, Southampton, United Kingdom, 30 pp + figures

Killworth PD (2003) Some physical and numerical details of frictional geostrophic models. Southampton Oceanography Centre internal report 90

Knutti R, Stocker TF, Joos F, Plattner G-K (2002) Constraints on radiative forcing and future climate change from observations and climate model ensembles. Nature 416:719–723

Levitus S, Boyer TP, Conkright ME, O’Brien T, Antonov J, Stephens C, Stathoplos L, Johnson D, Gelfeld R (1998) Noaa Atlas Nesdis 18, World ocean database 1998, vol. 1, Introduction, US Government Printing Washington DC, 346pp

Marsh R, Yool A, Lenton TM, Gulamali MY, Edwards NR, Shepherd JG, Krznaric M, Newhouse S, Cox SJ (2005) Bistability of the thermohaline circulation identified through comprehensive 2-parameter sweeps of an efficient climate model. Clim Dyn (in press)

McPhee MG (1992) Turbulent heat flux in the upper ocean under sea ice. J Geophys Res 97:5365–5379

Millero FJ (1978) Annex 6, freezing point of seawater. Unesco technical papers in the marine sciences 28:29–35

Oort AH (1983) Global atmospheric circulation statistics, 1958–1973:NOAA Prof Pap 14

Petoukhov V, Ganopolski A, Brovkin V, Claussen M, Eliseev A, Kubatzki C, Rahmstorf S (2000) CLIMBER-2: a climate system model of intermediate complexity. Part I: model description and performance for present climate. Clim Dyn 16:1–17

Roemmich D, Wunsch C (1985) Two transatlantic sections: meridional circulation and and heat flux in the subtropical North Atlantic Ocean. Deep-Sea Res 32:619–664

Semtner AJ (1976) Model for thermodynamic growth of sea ice in numerical investigations of climate. J Phys Oceanogr 6:379–389

Sinha B, Smith RS (2002) Development of a fast Coupled General Circulation Model (FORTE) for climate studies, implemented using the OASIS coupler. Southampton Oceanography Centre Internal Document, No 81:67

Thompson SL, Warren SE (1982) Parameterization of outgoing infrared radiation derived from detailed radiative calculations. J Atoms Sci 39:2667:2680

Trenberth KE, JM Caron (2001) Estimates of meridional atmosphere and ocean heat transports. J Climate 14:3433–3443

Weaver AJ, Eby M, Wiebe EC, Bitz CM, Duffy PB, Ewen TL, Fanning AF, Holland MM, MacFadyen A, Matthews HD, Meissner KJ, Saenko O, Schmittner A, Wang H, Yoshimori M (2001) The UVic Earth System Climate Model: model description, climatology, and applications to past, present and future climates. Atmos-Ocean 39:361–428

Wright DG, Stocker TF (1991) A zonally averaged ocean model for the thermohaline circulation. Part I: model development and flow dynamics. J Phys Oceanogr 21:1713–1724

Zaucker F, Broecker WS (1992) The influence of atmospheric moisture transport on the fresh water balance of the Atlantic drainage basin: General Circulation Model simulations and observations. J Geophys Res 97:2765–2773

## Acknowledgements

We thank J.D. Annan for helpful comments on statistical analysis and Jeff Blundell for help in processing ETOPO5 data. The modification to Hibler’s sea ice-area equation was suggested by Masakazu Yoshimori. NRE is supported by the Swiss NCCR-Climate programme. RM acknowledges the support of the UK NERC Earth System Modelling Initiative.

## Author information

### Authors and Affiliations

### Corresponding author

## Appendix, the sea-ice model

### Appendix, the sea-ice model

Ocean surface fluxes are everywhere partitioned between ocean- and ice-covered fractions, the total heat flux *Q*_{t} into the ocean or ice surface being

where *E* is the rate of evaporation or sublimation, calculated as in Weaver et al. (2001), *L* is the latent heat of evaporation *L*_{v}, or sublimation *L*_{s}, over ocean or ice respectively and the planetary albedo over sea ice is assumed to decrease linearly with air temperature within a given range following Holland et al. (1993):

The sea ice has no heat capacity, thus the heat flux *Q*_{t} from the atmosphere is assumed to be equal to the vertical heat flux through the ice given by

where ν_{
i
} is the vertical conductivity of heat within the sea ice and *T*_{f} is the local salinity-dependent freezing temperature of sea water, assumed to be the temperature at the base of the sea ice, given by

where the values for the constants γ_{
i
}, *i*=1,2,3, are as given by Millero (1978). From Eqs. 6 and 8 we obtain a single equation, Φ(*T*_{
i
})=0, say, for the ice-surface temperature *T*_{
i
}, which is solved by the Newton–Raphson iteration.

Having obtained the ice-surface temperature, and hence the heat flux from atmosphere to sea ice, we calculate the heat flux from the sea ice into the ocean (normally negative) as

where *T*_{o} is the temperature at the ocean surface; in the model *T*_{o} is the temperature of the uppermost (mixed) layer, Δ*z* is the thickness of this layer, *C*_{pi} is the specific heat of sea ice at constant pressure, ρ_{
i
} its density and τ_{
i
} is a timescale for the relaxation of the ocean surface temperature to the freezing temperature. McPhee (1992) suggests a physical value for the ratio *Q*_{b}/(*T*_{f}− *T*_{o}) although the most appropriate value is liable to depend on model temporal and spatial resolution. However, we retain the value of 17.5 days for τ_{
i
} implied by McPhee’s parameterisation with the present value of Δ*z*.

We are now in a position to calculate the growth rate *G*_{
i
} of sea-ice height in the ice-covered ocean fraction, which is given by the deficit of heat fluxes into and out of the sea ice, minus the latent heat loss of sublimation. Snow is not considered in the model, and all the precipitation over the ocean or sea ice is added directly to the ocean surface layer. Thus

where *L*_{
f
} is the latent heat of fusion of ice. In the open-ocean fraction we take −*Q*_{b}, from Eq. 7, to be the largest possible heat flux out of the ocean. Thus if the ocean-to-atmosphere heat flux is greater than this, the deficit leads to ice growth in the open water fraction. In general the growth rate of sea ice in the open-ocean fraction is

We can thus calculate the net growth rate *G* and the rate of change of the average sea-ice height *H*, which is also subject to advection by the surface ocean velocity and a diffusive term, which takes the place of a detailed representation of unresolved sea-ice advection and rheological processes. Note that *H*
represents the height of sea ice averaged over both open ocean and ice-covered fractions.

where κ_{hi} is a horizontal diffusivity.

The rate of change of sea-ice area *A* is given by

The first term on the right-hand side parameterizes the possible growth of ice over open water. The effect of this term is that, if *G*_{o} is positive, the open water fraction decays exponentially at the rate *G*_{o}/*H*_{0}, where *H*_{0} is a minimum resolved sea-ice height. The second term parameterizes the possible melting of sea ice and corresponds (by simple geometry) to the rate at which *A* would decrease if all the sea ice were uniformly distributed in height between 0 and 2*H*/*A* over the sea-ice fraction *A*. Note that this represents a small modification to the original sea ice-area equation proposed by Hibler (1979) in that the decay term is proportional to the (negative) growth rate *AG*_{
i
} over the sea-ice fraction as opposed to the total (negative) growth rate *G*. The two formulations differ only if sea ice is forming in the open water fraction and simultaneously melting in the ice-covered fraction, in which case using the total growth rate means that melting affects both terms; a form of “double counting”. This is not as unlikely as it sounds since the ice fraction is normally much colder than the water fraction, thus the heat flux deficit causing sea-ice growth in the model is normally smaller over sea ice and can easily be of opposite sign to the deficit over open water.

We can now define the flux of heat into the ocean as

The flux of fresh water into the ocean is given by

The first two terms represent precipitation and runoff, the third term evaporation, *E*_{e}, over open water and the last two ice melting and freezing. The net fresh water source due to melting being proportional to the sink of average sea-ice height minus the net rate of sublimation, denoted *E*_{s} here for clarity. The flux of salinity into the ocean is simply taken to be *F*_{
S
}=-*S*_{0}*F*_{
W
}, where *S*_{0} is a constant reference salinity (using the local salinity would complicate the calculation of the global freshwater budget).

Finally, we have to ensure, at each timestep, that the calculated sea-ice height *H* is positive. In practice we ignore the presence of thin ice and set *H*=0 whenever the calculated value of *H* is less than *H*_{0}. This means that the heat and freshwater fluxes must be modified accordingly for consistency. For positive or negative *H*, if *H* is set to zero, the corresponding amount of heat to be added to the ocean is −*H*ρ_{
i
}*L*_{f}, while the corresponding amount of fresh water to be added is *H*ρ_{
i
}/ρ_{o}. In addition, we have to ensure that numerical truncation error does not result in unphysical values of *A*, *A*<0 or *A*>1, but these latter adjustments carry no further implications for the conservation of heat and fresh water.

## Rights and permissions

## About this article

### Cite this article

Edwards, N.R., Marsh, R. Uncertainties due to transport-parameter sensitivity in an efficient 3-D ocean-climate model.
*Clim Dyn* **24**, 415–433 (2005). https://doi.org/10.1007/s00382-004-0508-8

Received:

Accepted:

Published:

Issue Date:

DOI: https://doi.org/10.1007/s00382-004-0508-8

### Keywords

- Antarctic Circumpolar Current
- Couple Model Intercomparison Project
- Freshwater Flux
- Initial Ensemble
- Deep Temperature