Ocean Modelling 8 (2005) 31–54 www.elsevier.com/locate/ocemod
An eddy resolving global 1/10 ocean simulation
Mathew E. Maltrud a,*, Julie L. McClean b a Fluid Dynamics Group, Los Alamos National Laboratory, Los Alamos, NM 87545, USA b Department of Oceanography, Naval Postgraduate School, Monterey, CA 93943, USA Received 9 October 2003; received in revised form 2 December 2003; accepted 2 December 2003 Available online 23 December 2003
Abstract Initial results are presented from a global eddy resolving simulation using the Parallel Ocean Program (POP) general circulation model. The model has 1/10 horizontal spacing at the equator, employs a displaced pole grid in the Northern Hemisphere to allow the inclusion of the Arctic Ocean, and uses 40 vertical levels. The simulation was spun up from rest using a combination of daily NCEP/NCARreanalysis and monthly observational data products for the period 1979–1993 as surface forcing. As expected, the simulation exhibits extremely turbulent behavior, with eddy energy agreeing well with satellite altimetry in both location and magnitude in most high-activity areas. However, significant problems exist in the Gulf Stream/North Atlantic Current region, and relatively minor discrepancies occur elsewhere. Overall, the general circulation is well represented, with acceptable values for overturning mass and heat transports, and good agreement with transport estimates of the major current systems. This simulation represents a major step forward in high resolution ocean modeling, with applications to prediction, climate, and general ocean science. 2003 Elsevier Ltd. All rights reserved.
Keywords: Modelling; Ocean circulation
1. Introduction
The realism of ocean general circulation models (OGCMs) has increased significantly in the past decade. Improved numerics, sub-grid scale parameterizations, surface forcing, and bathymetry products have all contributed to better representations of the observed ocean
* Corresponding author. E-mail address: [email protected] (M.E. Maltrud).
1463-5003/$ - see front matter 2003 Elsevier Ltd. All rights reserved. doi:10.1016/j.ocemod.2003.12.001 32 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 circulation. A significant step forward has also been achieved with advances in supercomputing technology that allow these models to be run at eddy-resolving rather than eddy-permitting resolutions. The increased availability of both in situ and satellite observations from the World Ocean Circulation Experiment (WOCE) and other field programs has provided a means of evaluating the simulated space–time representation of important ocean processes. The success of these eddy-resolving models in a variety of applications such as state estimation, climate studies, and ocean process studies is dependent upon the statistical consistency of the model and obser- vations at the scales of interest. A number of studies, using different ocean models, have been conducted at eddy-resolving resolutions in the North Atlantic. Smith et al. (2000) performed a 1/10 , 40-level simulation using the Parallel Ocean Program (POP). They found significant improvements in the repre- sentation of the mesoscale variability as well as in the mean flow (such as the Gulf Stream separation and path of the North Atlantic Current) relative to coarser simulations such those performed by the Community Modeling Effort (CME, Bryan and Holland (1989)). McClean et al. (2002), using a separate North Atlantic 1/10 POP simulation, compared pseudo-Eulerian and Lagrangian statistics from it and surface drifting buoys; the 1/10 model produced realistic current strengths, eddy energy levels, and intrinsic scales. Hurlburt and Hogan (2000) simulated a realistic Gulf Stream separation and circulation when they increased the horizontal resolution of their Naval Research Laboratory (NRL) layer ocean model (NLOM) from 1/8 to 1/16 . Chassignet and Garraffo (2001), using a 1/12 configuration of the Miami Isopycnic Coordinate Ocean Model (MICOM), also obtained a realistic Gulf Stream separation. It was concluded from these results that the horizontal grid spacing of such models should be at least 1/10 to adequately simulate both mesoscale variability and narrow mean flows such as western boundary currents. From these basin scale studies, it was expected that the integrity of global eddy-resolving ocean solutions would likely be significantly improved relative to that of lower resolution models such as the 0.28 near-global simulations of Maltrud et al. (1998). However, running these models stretches current high performance computing resources to their limits, so few efforts have been made to do so. Smedstad et al. (2003) have run a 1/16 , near-global (up to 65 N) configuration of NLOM with only 6 layers (plus a mixed layer) as the underlying ocean component of an eddy-resolving prediction system. Sakuma et al. (2003) have run a 1/10 near-global (up to about 75 N) version of the Modular Ocean Model (MOM) that demon- strates the enormous potential of the Earth Simulator. Coward et al. (2002) have begun a 1/12 , 66 level fully global simulation using the Southampton Oceanography CentreÕs Ocean Circulation and Advanced Modelling Project (OCCAM) model that shows impressive early results. In this manuscript, we describe the 15-year spinup of a global POP simulation with 1/10 equatorial resolution and 40 vertical levels. The model has been configured on a displaced North Pole grid that includes the Arctic, so can, in fact, be considered truly ‘‘global’’ instead of having a boundary at high northern latitudes where water is conditioned to be consistent with known characteristics. We have also chosen to force the model with synoptic fluxes whenever possible, as opposed to repeating monthly (or daily) climatologies. The primary purpose of this paper is to describe, in a very descriptive sense, the general characteristics of the simulation which may subsequently be built upon for more detailed analyses. Comparisons M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 33 with observations will be used in this preliminary assessment of the integrity of the model solution. We will also be comparing with other model results, though it is important to note that it will not be a true convergence study due to the very different configurations of the models. The paper is organized in the following manner. The model setup is described in Section 2, while model analyses and comparisons with data and published findings are found in Section 3. The discussion and conclusions are contained in Section 4.
2. Model setup
2.1. Grid and bathymetry
The high-resolution global simulation described here was performed using the Parallel Ocean Program (POP, Dukowicz and Smith (1994)), a z-level primitive equation model descended from the original Bryan–Cox code (Bryan, 1969). Choices regarding the model setup were influenced by the success of an earlier 1/10 , 40-level POP simulation of the North Atlantic basin (Smith et al., 2000) which showed that this resolution appears to be the minimum necessary for achieving significant improvements in many aspects of the circulation that had been clearly deficient at lower resolution. For example, the Gulf Stream separated cleanly at Cape Hatteras, the North Atlantic Current followed a much more realistic path, and the Azores Current was well simulated. As a result, we chose the same horizontal grid spacing of 1/10 at the equator, as well as the same vertical discretization of 40 levels, varying from 10 m at the surface to 250 m at depth, with a maximum depth of 5500 m. From 79 S to the equator, the grid is mercator (d/ ¼ 1=10 Ã cosð/Þ), where / is the latitude. North of the equator, the grid is smoothly distorted so that the north pole of the grid is located in North America (Smith et al., 1995), thus allowing the Arctic region to be completely included without having to resort to high latitude filtering. The resulting grid has a logical dimension of 3600 · 2400 · 40 (over 345 million points); of these, 177 million are active ocean points. Based on an average area of the surface ocean grid cells, the average resolution is 7.8 km. While a mercator grid has the desirable property of having grid cells with a horizontal aspect ratio of unity (the ratio of the longer to the smaller cell length), this displaced pole grid has some cells with an aspect ratio as high as 4.3 near the logical poles, though 91% have an aspect ratio under 2. Bathymetry for the grid was generated using three separate topographic data sets: Smith and Sandwell (1997), the International Bathymetric Chart of the Arctic Ocean (IBCAO) (Jakobsson et al., 2000), and BEDMAP (Lythe and Vaughan, 2001) for the non-polar regions, the Arctic, and the Southern Ocean, respectively. Linear interpolation was used for the polar datasets when calculating depths at the model grid points; for the Smith–Sandwell region the model depth was defined as the average of all data points within a grid cell. In regions where the products overlap (66 Nto67 N for the Arctic, 71 Sto66 S for the Antarctic), the depths were defined using a linear weighted average. Any resulting depths shallower than 3 m were considered to be land, and the minimum ocean depth was set to be 20 m (two levels). Several person-weeks were then devoted to hand-modifying the model depth field in regions of important sills and channels that may have 34 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 been smoothed out by the interpolation process, such as the Canadian and Indonesian Archi- pelagos, mid-ocean ridge fracture zones, and the Denmark and Faroe-Shetland overflows.
2.2. Parameterizations
Biharmonic momentum and tracer diffusivity were used to parameterize subgrid scale hori- zontal mixing, as well as topffiffiffi damp out small scale noise. The values varied with the cube of the average grid length (dl ¼ A, where A is the area of the grid cell) for a given cell (see Maltrud et al. (1998)),
3 3 m ¼ m0 Ãðdl=dl0Þ ; j ¼ j0 Ãðdl=dl0Þ
17 4 where the equatorial values are denoted by the subscript 0, with m0 ¼ 27 10 cm /s for 17 4 momentum and j0 ¼ 9 10 cm /s for tracers exactly as in Smith et al. (2000). Due to the distortion of the grid, some cells became small enough to severely limit the time step due to the diffusive CFL criterion. To alleviate this problem, the values of the diffusivities were modified in these few regions (Gulf of Maine and the Canadian Archipelago) to allow for a reasonable time step, while still effectively damping noise. During the first 2 years of the simulation, vertical mixing coefficients were calculated at every timestep using the Richardson number based formulation of Pacanowski and Philander (1981) and convective instabilities were resolved using two passes through a convective adjustment scheme. For the remainder of the run, the K-profile parameterization (KPP) (Large et al., 1994) was used. Background values for tracer diffusion range from 0.1 cm2/s near the surface to 1.0 cm2/s at depth, with viscosity values an order of magnitude higher. Large values for the diffusivity and viscosity (1000 cm2/s) were used to simulate convection. It is also worth noting which parame- terizations this simulation does not use. In particular, there is no explicit sea ice model and no bottom boundary layer parameterization.
2.3. Forcing and initial conditions
In order to more accurately simulate energy input into the mixed layer, we chose to spin up the model from rest using synoptic forcing (whenever possible), as opposed to applying a repeating monthly (or even daily) climatology. Surface fluxes were calculated in a manner very similar to Large et al. (1997) using a combination of the NCEP/NCARreanalysis products (Kalnay et al., 1996) (6-h fields averaged to 1-day), as well as monthly data from various sources noted below. Wind stress was calculated offline using the formulation of Large and Pond (1982) with clima- tological sea surface temperature (SST, Shea et al. (1990)). No modifications to the strength of the wind stress was applied in ice-covered areas. Surface heat flux was calculated every timestep using the model potential temperature in the upper-most level, the daily reanalyses, and monthly International Satellite Cloud Climatology Project (ISCCP) downward shortwave flux and cloud fraction (Rossow and Schiffer, 1991), with the latter being used to calculate the longwave flux using the formulation of Fung et al. (1984). For calendar years 1984–1990, monthly varying ISCCP data was used; otherwise a climatology calculated from this seven year period was utilized. In addition, the model potential temperature is M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 35 restored to climatology over the depth of the uppermost level (10 m) with a timescale of about 5 days in regions that are specified as ice-covered as defined by the )1.8 C isotherm in the SST climatology. For the fresh water fluxes, evaporation was derived from the latent heat flux, and precipitation was specified with a blended monthly microwave sounding unit (MSU) (Spencer, 1993) and Xie– Arkin (Xie and Arkin, 1997) dataset (see Doney et al., 2003). As with the heat flux, there is a 5 day restoring of salinity over the depth of the uppermost level to climatology (Polar Hydrography Climatology (PHC), version 1, (Steele et al., 2001)) in ice-covered regions. In addition, salinity in the top level is restored to climatology in the open ocean with an 18 day timescale in order to keep it from drifting too much. To ensure a global freshwater balance, the precipitation was multiplied by a constant factor that was adjusted every year based on the change in globally integrated salinity (Large et al., 1997). This ‘‘precipitation adjustment’’ factor (typically around 1.15 in value) was increased/decreased if the previous year showed an increase/decrease in salinity. In essence, the precipitation factor makes up for the lack of explicit river runoff in the freshwater balance (but is distributed in proportion to the precipitation throughout the ocean surface), while the relative freshness of the water in outflow regions is approximated by the salinity restoring terms. The initial temperature and salinity distributions were generated using a combination of the NavyÕs Modular Ocean Data Assimilation (MODAS) 1/8 climatological product (Fox et al., 2002) and the PHC (Steele et al., 2001) dataset for the Arctic. These initial fields were also used to restore temperature and salinity in three marginal sea outflow regions that were found to be problematic. The outflow from the Mediterranean Sea, Red Sea and Persian Gulf were found to be too shallow (about 500 m in the Mediterranean case, for example, instead of 1000 m). Full depth restoring within a radius of about 350 km with a time scale of 15 days was used to try and maintain more realistic overflow water masses in these regions; otherwise no restoring was used below the surface. As noted above, daily forcing was applied from the very beginning of the run starting at January 1, 1979 (1979 being the first year with monthly precipitation data) with a timestep of 6.3 min. Using 500 processors of an IBM Power3 computer at the Navy Oceanographic Office, the model was able to complete one simulated year in about 8.5 days, generating an average of about 700 GB of output per model year. The model was run for 15 years (through 1993) which appears to be a sufficient amount of time for mesoscale processes in the upper ocean to have reached a quasi-equilibrium. Results described here will typically be based on averages over the three year period 1991–1993 (model years 13–15) or snapshots from the final year.
3. Results
The main objective of this paper will be to present an overview of the characteristics of this simulation, with more detailed analyses to follow in later manuscripts. We will rely on compar- isons with other global and regional POP runs, as well as with data, as a means of assessing the quality of this simulation. In particular, we will compare this run with the 0.28 global (except for the Arctic) simulation of Maltrud et al. (1998) (actually, an extension of this run using daily 36 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 forcing for years 1993–1997; for details see Garfield et al. (2001)), and the 1/10 North Atlantic simulation of Smith et al. (2000).
3.1. Quasi-steady state diagnostics
Time series of volume averaged kinetic energy, potential temperature and salinity are shown in Fig. 1 to demonstrate the extent to which a quasi-steady state has been reached at the end of the 15-year run. The kinetic energy is seen to increase rapidly from 10 to 30 (cm/s)2 by 1981 and then fluctuate between 26 and 30 (cm/s)2 for the remaining years of the simulation. As a result of the increased resolution, the total kinetic energy is about 20% larger than the 0.28 global run of Maltrud et al. (1998). The potential temperature is seen to decrease some 0.03 C over the first 7 years, stay roughly constant until 1989, increase slightly for the years 1989–1991, and then de- crease during 1991–1993. It is likely that these small changes in the temperature trend are due to using non-climatological forcing. The salinity trend is almost constant after 1988 as a result of using the precipitation factor to encourage global freshwater balance.
2 35 30 25 20 15 10 5 Kinetic Energy (cm/s) 1980 1982 1984 1986 1988 1990 1992 1994 year 3.62 3.61 3.60 3.59 3.58 3.57 1980 1982 1984 1986 1988 1990 1992 1994 Potential Temperature (˚C) year 1.25
4.5 Precipitation Factor 4.0 1.20 3.5 1.15 3.0 2.5 1.10 2.0 1.05 1.5 1.00 1.0 0.5 0.95
Salinity Residual (ppm) 1980 1982 1984 1986 1988 1990 1992 1994 year
Fig. 1. Time series of globally averaged kinetic energy, potential temperature, and salinity residual (1000 · [salin- ity)34.73]). Included as black squares in the salinity plot are the annual values of the precipitation factor, showing that it was accidently reset to 1 at the beginning of year 3, which had an immediate effect on the salinity field. M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 37
3.2. General circulation
To provide a broad overview of the large scale general circulation of the 1/10 model, we will present three interrelated measures of the flow for the global ocean as well as for each of the ocean basins. First, we consider the meridional overturning streamfunction (Fig. 2) for a zonally averaged view. Next, in order to point out features that can be difficult to identify in a zonal average, we will compare model mass transports with depictions of the present global ocean circulation obtained by Schmitz (1995) (hereafter referred to as S1995) from observations and by Ganachaud and Wunsch (2000) (hereafter referred to as GW2000) from an inverse model. Transports are divided into three layers: upper, deep, and bottom, with the upper layer including both surface and intermediate waters. GW2000 uses very specific definitions of these water masses using neutral surfaces, while S1995 is somewhat less rigorous but extremely detailed in a
Fig. 2. Meridional overturning streamfunction averaged over 1991–1993 for (a) Atlantic, (b) Indian + Pacific, (c) Indian (which is only defined north of the Indonesian Throughflow). The contour interval for (a) and (b) is 2 Sv (1 Sv ¼ 106 m3/s) and is 0.5 Sv for c). Negative contours are dashed and denote a counter-clockwise sense of flow. The streamfunctions were calculated with 1 latitudinal spacing, which smoothes out small scale variations. 38 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 descriptive sense. Here, we simply chose depth levels that varied a bit between ocean basins, but were in the range 0–1500 m for the upper layer, 1500–3500 m for the deep layer, and 3500–5500 m for the bottom layer. Sensitivity testing was performed by varying these ranges, and by using temperature and density classes to define the layers, but none made a major difference in the interpretations that follow. In fact, the variations were typically very similar to the quantified uncertainties resulting from the GW2000 inverse estimates. Vertical transfers are due only to advective processes in the model, while the others consider both advection and mixing. Fig. 3 shows the result of distilling a three year average of the model mass fluxes into a figure similar to Fig. 2 of GW2000 that includes both the inverse estimates and those of S1995 where appropriate. Overall, the agreement between all three estimates is quite good, while noting that both GW2000 and S1995 represent estimates of flow over a much longer time scale than the model. However, significant differences between the estimates do exist and will be addressed below for individual basins.
Fig. 3. Volume transports (in Sv) from 3 different estimates (denoted by outline around each number) divided into 3 depth ranges (denoted by color). Unenclosed numbers correspond to values from this model averaged over model years 1991–1993; numbers enclosed in boxes are from the inverse calculation of Ganachaud and Wunsch (2000); numbers enclosed in circles are from Schmitz (1995). Red numbers correspond to surface and intermediate waters (typically 0– 1250 m); blue numbers are deep waters (typically 1250–3500 m); green numbers are bottom waters (typically below 3500 m). Vertical transports are denoted by circles with either a dot (for upward) or an X (for downward) and are colored according to the source water. For example, a blue circle with an X shows deep water transforming into bottom water, while a green circle with a dot shows bottom water becoming deep water at the same depth. M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 39
The third measure of the model circulation is provided by the meridional heat transport (Fig. 4) for the global ocean and Atlantic and Indo-Pacific basins. Also noted in Fig. 4 are estimates based on atmospheric reanalyses by Trenberth and Caron (2000) and inverse modeling by Ganachaud and Wunsch (2003). Certainly the model is not in thermal equilibrium after 15 years of simulation, so the heat transports must be viewed with some caution. In fact, the area-averaged net heat flux for 1991–1993 is )1.6 W/m2, which is consistent with the downward trend in average temperature for these years seen in Fig. 1, and is likely due to interannual variability in the surface forcing. While this is an order of magnitude smaller than the amplitude of the seasonal cycle, it still amounts to 0.6 PW of imbalance. The Arctic basin will also be discussed in some detail, but since it does not lend itself well to a zonally averaged description, it will be dealt with somewhat differently than the Atlantic and Indo-Pacific basins. The Southern Ocean will not be considered in detail; we only note here that
3 2 Global 1 0 -1
heat transport (PW) -2 -80 -60 -40 -20 0 20 40 60 80 latitude 1.6 1.4 Atlantic 1.2 1 0.8 0.6 0.4 0.2
heat transport (PW) 0 -80 -60 -40 -20 0 20 40 60 80 latitude 1.5 1 0.5 Indo-Pacific 0 -0.5 -1 -1.5 -2
heat transport (PW) -2.5 -80 -60 -40 -20 0 20 40 60 80 latitude
Fig. 4. Meridional heat transport (in Petawatts) averaged over 1991–1993 for the global ocean, the Atlantic, and the Indian + Pacific. Model curves are in red. For comparison, estimates from Trenberth and Caron (2000) using NCEP and ECMWF reanalyses are shown with thick blue and green curves, respectively. The thin blue and green curves show these estimates plus and minus the uncertainty. The black circles are the inverse model estimates (with error bars) by Ganachaud and Wunsch (2003). In the Atlantic, the cyan curve is a 1989–1991 average from the 1/10 North Atlantic simulation of Smith et al. (2000). As was done for the meridional streamfunctions in Fig. 2, the heat transport from this simulation was calculated using 1 latitudinal spacing. 40 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 the transport of the Antarctic Circumpolar Current through Drake Passage (Fig. 3) shows very good agreement with GW2000 as well as the estimate of 134 ± 14 Sv by Nowlin and Klinck (1986).
3.2.1. Atlantic Fig. 2a shows the meridional overturning streamfunction for the Atlantic basin. A maximum of 23 Sv is seen at 40 N, the cross-equatorial transport is 16 Sv, and the strength of the bottom cell is 2 Sv. The strength of the model overturning in this basin agrees well with previous model studies, such as Smith et al. (2000). Fig. 3 shows that the agreement between the model and observational mass transport estimates are generally quite good in the upper and deep levels, in that the model values are within or very close to the limits of observed uncertainty. In the bottom layer however, the model transport is too weak, with only 2 Sv flowing north through the Vema channel (and a small fraction of a Sverdrup through Hunter Channel) which is partially balanced by 1 Sv flowing south through Walvis Ridge. Observational estimates put the Vema and Hunter transports at 4 and 3 Sv, respectively (Hogg et al., 1999). The model and S1995 agree that the bottom flow at 25 N is toward the north, while GW2000 show southward flow that is due to a northern source of bottom water combined with upwelling of bottom water that enters from the south. However, this is one example where the choice of depth range that defines a water mass in the model can make a difference. By moving the top boundary of the bottom layer up 500 m, the direction of the net transport changes to southward. In the Atlantic south of about 30 N, meridional heat transport values (Fig. 4b) fall in between the estimates based on ECMWF and NCEP, but are at the very low end of the inverse model estimates. To the north, the model strongly overestimates the heat transport. Almost all of the net heat loss in the model occurs in the Labrador Sea, Irminger Sea and North Atlantic Current between 50 N and 65 N, instead of between 30 N and 45 N in the Gulf Stream Extension. This is in part due to anomalous heat fluxes that result from the modelÕs poor representation of the Gulf Stream/North Atlantic Current system (which will be discussed in more detail in Section 3.4). Similar behavior can be seen in the 1/10 North Atlantic basin simulation by Smith et al. (2000) (cyan curve in Fig. 4b), though not as pronounced since it does exhibit a very realistic Gulf Stream and North Atlantic Current. It is interesting to note how close the global and North Atlantic transports agree given the very different heat flux forcing in the two models.
3.2.2. Pacific and Indian Oceans The Indo-Pacific streamfunction with its intense equatorial upwelling is shown in Fig. 2b. The bottom water flow is quite robust and due primarily to the Pacific component since the Indian also shows weak bottom flow (Fig. 2c). The Indian basin also shows a hint of a shallow (top 200 m) overturning cell (Schott et al., 2002) with northward subsurface flow compensated by southward Ekman transport. However, this circulation also appears to be very weak, with only about 1 Sv reaching 5 N, instead of an estimated 6 Sv (Schott et al., 2002). In the Pacific, the cases where the model mass transports (Fig. 3) significantly disagree with inverse results or observations are the upper layer flow compared to GW2000 and the bottom flow compared to S1995. In the model, essentially all of the bottom flow north of 25 S (10 Sv) enters along the Kermadec Ridge with very little recirculation to the south over the rest of the basin. This transport is low (but within uncertainty) compared to current meter estimates of 16 Sv M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 41
(Whitworth et al., 1999), though the model transport of 6 Sv (not explicitly noted in Fig. 3) agrees with current meter data (also 6 Sv, Rudnick (1997)) by the time the flow reaches the Samoan Passage around 10 S. For the upper layer, GW2000 shows twice as much transport at 15 S than the model or S1995. Part of this excess translates into an increased Indonesian Throughflow, but it also requires a conversion from upper to deep water in the mid-Pacific. The model predicts just the opposite, showing a conversion of deep to upper water that augments the flow into the Indian Ocean. In the northern part of the basin, the model and GW2000 agree quite well in that the upper transport of the Kuroshio Current system is balanced by the southward gyre return flow, and that the few Sverdrups of bottom water that reach this area are upwelled and returned as deep water. The modelÕs Pacific heat transport (Fig. 4) agrees fairly well with estimates in the northern hemisphere. In the southern hemisphere, itÕs southward transport is significantly lower than the estimates of Trenberth and Caron (2000), but is within the uncertainty limits of Ganachaud and Wunsch (2003). Disagreements between the model and observational estimates are more common in the Indian Ocean. The modelÕs Indonesian Throughflow is bracketed between the others and is almost certainly a reasonable value, though it is interesting to note that the transport is always westward (Fig. 9), in contrast with hydrographic estimates of Sprintall et al. (2002) who show episodes of net eastward flow. The bottom flow in the model (Fig. 3) is quite a bit too weak, though closer to the GW2000 estimate. Bottom water enters the model Indian Ocean almost entirely through the Atlantis II Fracture Zone in the Southwest Indian Ridge, with a smaller amount entering just east of Broken Ridge. Although the model topography was modified by hand, it is likely that this fracture zone is still too narrow to allow a realistic amount of transport into the Madagascar Basin. All three do agree in general that most (if not all) of the bottom water upwells to eventually cause the formation of new upper water flowing south out of the basin. The model and GW2000 agree that there is no net flow of deep water past 32 S, while the large amount of upper layer transport seen in GW2000 compared to the model is again partly due to the high Indonesian Throughflow, and partly due to twice as much bottom layer transport.
3.2.3. Arctic Earlier near-global simulations, such as the 0.28 POP (Maltrud et al. (1998)), had their northern boundaries around 75 N. Here the use of a truly global grid allows us to examine the circulation in the Arctic Ocean and the flow through the Canadian Archipelago as part of a high resolution global simulation. Fig. 5 shows time series of transports through several key straits and passages for the final three years of the spinup. The transport through Bering Strait of 1.0 Sv agrees well with the 1990–1994 estimate of 0.83 Sv by Roach et al. (1995), as does the variability with events of order 1 Sv southward and over 2 Sv northward. The net transport through Fram Strait (1.8 Sv) is only about half of the value (4.2 Sv) given by Fahrbach et al. (2001) for 1997– 1999 based on current meter moorings. It is also interesting to note that model transports of the northward West Spitzbergen Current (3.9 Sv) and the southward East Greenland Current (5.7 Sv) which make up the net transport through Fram Strait, are individually about half the observa- tional values. During the first few years of the spinup, it was noticed that the flow through the Canadian Archipelago was far too strong, with over 4 Sv entering the northern end of Baffin Bay. It is important to repeat that the model has no sea ice, and the wind stress acting on the ocean was not 42 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54
3.0 2.0 Bering 1.0 0.0
Transport (Sv) -1.0 1991 1992year 1993 1994 2.0 0.0 -2.0 -4.0 -6.0 Fram
Transport (Sv) -8.0 1991 1992year 1993 1994 -0.2 -0.4 -0.6 -0.8 -1 Lancaster
Transport (Sv) -1.2 1991 1992year 1993 1994 -0.2 -0.4 Robeson -0.6 -0.8 -1 -1.2 Transport (Sv) 1991 1992 1993 1994 year
Fig. 5. Volume transports (Sv) through several key Arctic passages. A 7 day running average filter has been applied to each time series. The long dashed line is the long term mean (1985–1993) and the short dashed lines are one standard deviation from the mean. Positive values denote flow into the Arctic. The circles are observational estimates by Roach et al. (1995) for the Bering Strait, by Fissel et al. (1988) for Lancaster Sound and Robeson Channel, and by Fahrbach et al. (2001) for Fram Strait. modified to mimic the effect of ice cover, so it is not too surprising that the transports were too large. In an attempt to reduce the transport, we made some minor topographic modifications in Lancaster Sound (between Baffin and Devon Islands) and Robeson Channel (between Greenland and Ellesmere Island) that cut this transport in half, which is closer to observational estimates (Fissel et al., 1988). Once the modifications were made, the resulting flow through the Archipelago appears to be quite realistic (Fig. 5), not just in Lancaster Sound and Robeson Channel, but throughout this region of intricate topography. There is a noticeable seasonal cycle in these transports which is out of phase with the cycle in Fram Strait, indicating a partial balance between these exits from the Arctic Ocean. Overall, the circulation in the Arctic Ocean is similar to what is believed to be the true circu- lation, though direct measurements are scarce. Fig. 6 shows average (1991–1993) velocity vectors at 268 m depth, which has many characteristics of schematics (e.g., Grotefendt et al., 1998) for flow of the subsurface Atlantic layer. There is cyclonic flow around the edge of the entire basin at this depth. The circulation is strongly coupled to the topography, with strong flow along the Lomonosov Ridge, and around the Chukchi Plateau/Northwind Ridge. Mean currents are over 10 cm/s in the more energetic regions (over 25 cm/s at the surface), with instantaneous velocities reaching values twice as large. M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 43
Fig. 6. Time mean (1991–1993) velocity vectors at about 268.5 m depth. Every 12th grid point has been plotted. The model bathymetry is shaded, with shallower depths denoted by lighter shades of grey. The region is a bit distorted (since it is plotted using the logical geometry) so latitude-longitude lines have been included.
3.3. Major current systems
One the of reasons for performing high resolution simulations is to more accurately resolve narrow mean currents, making it more likely that flows such as boundary current separations will be better represented. This was certainly the case in 1/10 North Atlantic simulation reported by Smith et al. (2000). In this section, we will look in some detail at several such current systems.
3.3.1. Gulf stream As noted above, the quite realistic simulation of the Gulf Stream and North Atlantic Current of Smith et al. (2000) was directly responsible for the choice of 1/10 resolution for this study. However, Fig. 7a shows that the global model results were not necessarily as good as expected. In particular, the Gulf Stream exhibits behavior reminiscent of lower resolution simulations by hugging the coast and separating north of Cape Hatteras. However, the circulation is somewhat different from lower resolution cases (see for example, Maltrud et al., 1998) where the separation is dominated by a quasi-permanent anticyclone which is not present here. These problems are not limited to just the flow at Hatteras and northward, as the flow through the Florida Straits is too low by at least 10 Sv (Fig. 8). Even though the flow into the Caribbean through the Lesser Antilles is within observational uncertainty, the model shows several Sverdrups of flow leaving through the Windward Passage (between the islands of Cuba and 44 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54
Fig. 7. Snapshot of model temperature ( C) at 15 m (model level 2) on November 28, 1992 for four Western Boundary Current regions.
Hispaniola) instead of 5–10 Sv of inflow. This difference makes up most of the deficit of transport through the Florida Straits. In fact, there appears to be a balance between the Antilles inflow and Windward Passage; relatively weak/strong inflow through the Antilles results in inflow/outflow through Windward Passage, while the sum of the two doesnÕt stray too far from about 20 Sv. Even though the net overturning (Fig. 2a) is similar here compared to Smith et al. (2000), the transport of the Deep Western Boundary Current (DWBC) at 26 N is quite different in this run, with 29 Sv of southward transport here compared to over 40 Sv in the North Atlantic basin simulation. While there remains some disagreement about the true magnitude of this transport (Lee et al., 1996; Chave et al., 1997), our simulated value of 29 Sv is within the range of obser- vational estimates.
3.3.2. Agulhas In contrast to the Gulf Stream, the separation point for the Agulhas Current (Fig. 7b) in both the model and observations is quite variable. Although the model current correctly leaves land at M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 45
35 30 Bahamas 25 20
Transport (Sv) 15 1986 1988 1990 1992 1994 year -10 -15 Lesser Antilles -20 -25 -30
Transport (Sv) -35 1986 1988 1990 1992 1994 year 15 10 Windward Passage 5 0 -5
Transport (Sv) -10 1986 1988 1990 1992 1994 year 4 Mona Passage 2 0 -2
Transport (Sv) -4 1986 1988 1990 1992 1994 year
Fig. 8. Volume transports (Sv) through four key Caribbean passages. A 30 day running average filter has been applied to each time series. The long dashed line is the mean over the plotted time period (1985–1993) and the short dashed lines are one standard deviation from the mean. Positive values denote flow out of the Caribbean. The square with error bars is an observational estimate based on cable data for the Bahamas (Larsen, 1992) and the circles are from Johns et al. (2002) for the others. about 27 E and follows the continental slope, it rarely extends past 22 E, where the Agulhas Bank causes it to separate. Other times, the current separates from the shelf even before it reaches 25 E. The real Agulhas current manages to typically flow around Agulhas Bank, allowing the current to extend at least 2 further west, though sometimes it, too, separates near 22 E. Strong eddies are indeed formed as the current becomes unstable, but it appears that the location of their formation and their subsequent path across the Atlantic are significantly affected. This will be discussed more in Section 3.4. Bryden and Beal (2001) estimate the Agulhas transport to be about 70 Sv, while the mean model transport is 66 Sv (Fig. 9). Also noticeable in Fig. 9 is a shift in the characteristics of the current around 1986 from relatively high mean transport and variability to lower values. This signal is not seen in the Mozambique Current so must be related to gyre scale vari- ability that affects the East Madagascar Current which ultimately joins the Agulhas. Unlike the 0.28 global model of Maltrud et al. (1998), here we find the existence of an equatorward undercurrent. However, the transport and peak speed of 1.1 Sv and 13 cm/s, respectively, are quite weak compared to the observed values of 6 Sv and 30–40 cm/s (Beal and Bryden, 1997). 46 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54
-4 -8 -12 -16 -20 Indonesian Throughflow Transport (Sv) 1980 1982 1984 1986 1988 1990 1992 1994 year -20 -40 Agulhas -60 -80 -100 -120 Transport (Sv) 1980 1982 1984 1986 1988 1990 1992 1994 year 0 Mozambique Channel -10 -20 -30 Transport (Sv) 1980 1982 1984 1986 1988 1990 1992 1994 year
Fig. 9. Volume transports (Sv) for Indonesian Throughflow, Agulhas Current, and Mozambique Channel. A 30 day running average filter has been applied to each time series. The long dashed line is the mean over the plotted time period (1979–1993) and the short dashed lines are one standard deviation from the mean. Positive values denote northward and eastward flow, negative values are southward and westward. For the Indonesian Throughflow, estimates by Ganachaud and Wunsch (2000) (circle with error bars), Schmitz (1995) (square), and Sprintall et al. (2002) (triangle) are included. An estimate by Bryden and Beal (2001) is included for the Agulhas, as are estimates for the Mozambique Channel by Ganachaud and Wunsch (2000) (circle with error bars) and DiMarco et al. (2002) (solid and unfilled diamonds).
3.3.3. East Australia current Like the Agulhas, the East Australian current (Fig. 7c) separation is variable due to shedding of eddies, occurring primarily in two locations. Sometimes the modeled current separates from the shelf in the vicinity of Cape Byron at about 28 S. At other times, the separation is about 4 further south. In both cases, the current quickly turns eastward and becomes unstable. The structure of the current at 30 S agrees very well with a combined current meter/hydrographic data analysis (Mata and Tomczak, 2000), who also note the existence of multiple separation regimes. The 1991– 1993 mean southward transport in the upper 2000 m in the model is 26.7 Sv, compared to 22.1 ± 4.6 Sv for the observational mean from November 1991 through November 1993. The magnitude of the variability is comparable to the mean in both the model and observations. The maximum mean speed in the model is 70 cm/s compared to 60 cm/s in the observations. In addition, there is a well defined northward undercurrent below 2000 m in both cases, with a transport of 3.9 Sv in the model compared to 3.2 ± 0.8 Sv in the observations.
3.3.4. Kuroshio Of all the western boundary currents, the Kuroshio appears to have the most realistic sepa- ration point, typically only about 1 north of the true separation location (Fig. 7d). However, there is a different persistent problem here. While the real Kuroshio can exhibit strong meanders as seen around 137 E in Fig. 7d, they are transient features (Kawabe, 1995) that typically last for M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 47
Fig. 10. 1991–1993 average of the zonal velocity (cm/s) constructed from model grid points near the locations of the Tokara Strait current meter moorings (Feng et al., 2000). The contour interval is 5 cm/s with negative (westward) values dashed. a few years. For the first several years of the model run, the Kuroshio showed realistic variability in this area, but the current has been locked into a state with a semi-permanent meander for the final decade of the spinup. Fig. 10 shows a cross section of the flow through Tokara Strait (south of Kyushu, Japan) which is a major pathway of the Kuroshio as it leaves the East China Sea. The average transport of 16.9 Sv is low compared to 24.3 Sv estimated by Feng et al. (2000), however, the structure of the model current is quite similar to observations. There are two distinct cores, clearly due to the influence of the Tokara Islands just upstream from the section. In the data, the southern branch extends deeper than the northern core, which is not the case in the model. In agreement with the data, the northern core is narrower, but it isnÕt clear in the data if there is such a difference in peak speed (as seen in the model) since the moorings did not extend above 200 m in the core regions. The data and model also both show the existence of a subsurface counter current on the north side below 500 m.
3.3.5. Equatorial Pacific currents Based on the Rossby radius, this 1/10 model easily resolves the important horizontal scales in the equatorial region. However, vertical resolution typically becomes more of an issue here due to the complex vertical structure of alternating jets. Fig. 11 shows cross sections of the average zonal velocity at 140 W and at the equator. Note that the average is over 1991–1993 which encom- passes the relatively mild 1992–1993 El Nino.~ 48 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54
Fig. 11. 1991–1993 average of the zonal velocity (cm/s) at (a) 140 W and (b) the equator. The contour interval is 10 cm/s with thicker lines every 50 cm/s. Negative values (westward) are dashed; regions of positive velocity have been shaded.
At 140 W, the structure looks very good, in fact, remarkably similar to that seen in Fig. 2 of Johnson et al. (2002), showing all of the currents very close to their correct location with appropriate strengths. The Equatorial Undercurrent (EUC) core is seen at about 100 m depth with an maximum of 85 cm/s (instantaneous maximum of 130 cm/s). During the height of strong El Nino~ periods, the model EUC weakens considerably (it essentially disappears during the 1982– 1983 event) in agreement with observations. Also clearly seen are the Subsurface Countercurrents (or Tsuchiya Jets) at depths between 150 m and 350 m at 7 S, 4 S, and 4 N, with peak speeds of 10–20 cm/s and the northern current being the strongest. At the surface, both branches of the South Equatorial Current can be seen, again with a stronger northern branch (59 cm/s maximum). The North Equatorial Contercurrent (NECC) maximum is found to be below the surface (around 50 m) at 6 N, which is not the case in reality. The South Equatorial Contercurrent is also characterized by an unrealistic subsurface maximum, and is at least 2 too far north. The zonal structure of the currents at the equator (Fig. 11b) is dominated by the characteristic decrease of the depth of the EUC core towards the east, though it is not as pronounced in the model as in reality, possibly due in part to El Nino.~ The influence of the Tungaru seamount chain can be seen at about 173 E, the Solomon Islands at 150 E, and the New Guinea Coastal Undercurrent at 140 E, all of which are noticeable in the data as well (Johnson et al., 2002). The relatively weak westward flowing Equatorial Intermediate Current can be seen at depths between 300 and 450 m. M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 49
3.4. Mesoscale variability
Another major reason for carrying out high resolution simulations is to investigate eddy var- iability in the model, which can then be quantitatively compared globally with satellite altimetry data. Fig. 12 shows the rms sea surface height variability from this 1/10 run compared with both data (Ducet et al., 2000) and the 0.28 global simulation. The average value of the variability for
Fig. 12. Sea surface height variability (cm) from (a) the 1/10 model for 1991–1993 (b) 0.28 model for 1995–1996 and (c) TOPEX-ERS1 for 1995–1997 (Ducet et al., 2000). 50 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 the current simulation (6.1 cm) is not much greater than that achieved with the 0.28 model (5.6 cm), both of which are much lower than the observations (8.0 cm). This is due primarily to the large areas in the central and eastern ocean basins where the models show very low variability. The increased resolution has certainly helped in regions of relatively high eddy energy. In the ACC, the agreement with data is extremely good, even when looking at individual maxima in the vicinity of New Zealand, Drake Passage and south of Australia. Energy levels associated with strong currents are also much improved with higher resolution, as are their areal extents, most notably in the Kuroshio, Gulf Stream and East Australia currents. For example, the high vari- ability region of the Kuroshio Extension in the 1/10 simulation extends out to 180 E in close agreement with data, while the 0.28 model variability only reaches to 160 E. Very good agree- ment can also be seen in the Mozambique channel, the South Madagascar current, the Agulhas Retroflection, and the Gulf of Mexico Loop current. Improvements have also been made in the eastern boundary currents (such as the Leeuwin and California currents), though the variability in these locations is typically still too low, except off the southwestern coast of Chile. The structure of the variability can be a useful indicator of both strengths and deficiencies of the mean flow. The location of the strong variability in the Kuroshio highlights that the separation is occurring at about the correct latitude. Some of problems discussed in Section 3.3 can also be seen here: model variability in the East Australia Current is several degrees too far equatorward and the maximum in the Agulhas is several degrees too far to the east. Additional deficiencies can also be seen. The narrow tongue of high variability extending to the northwest from the tip of South Africa indicates that the path of Agulhas rings is much too regular than in reality. This may be linked to the fact that the eddies are typically formed too far eastward. Also very obvious is the poor representation of the Gulf Stream/North Atlantic current system. Unlike the data (and the 1/10 North Atlantic simulation of Smith et al. (2000)), the modelÕs current does not turn northwestward around the Grand Banks, instead heading almost zonally across the Atlantic as in the 0.28 run. There is also no noticeable signature of the Azores current at about 35 N, which may be related to problems with the Mediterranean outflow discussed in Section 2.3 (Jia, 2000).
4. Discussion and conclusions
We have presented an overview from a 15 year spinup of a global 1/10 ocean circulation simulation using the Parallel Ocean Program (POP). Overall, the model results are quite satis- factory when compared to observations and previous model results. With the exception of the Gulf Stream/North Atlantic Current system, one must typically look in some detail to find problems. Despite its shortcomings, we feel that this simulation is a significant first step in exploring the behavior and fidelity of eddy resolving global models. We have presented a somewhat descriptive view of the general circulation, setting the necessary groundwork for fur- ther, more detailed studies. In fact, many of the problems described in this paper have started to be investigated and im- proved, even though exhaustive sensitivity experiments are prohibitively expensive to perform. Perhaps the most significant is the realization that even at such high resolution, care must be taken when parameterizing subgrid scale processes. In a limited series of sensitivity experiments, it has been seen that the separation of the Gulf Stream and the characteristics of the North Atlantic M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54 51
Current (including turning at the Northwest Corner) are noticeably influenced by the values of the tracer and momentum diffusivities. Another issue is the relationship of fully global to almost global and basin scale simulations that use three-dimensional restoring of properties at the domain boundaries. While investigating the problems with the Gulf Stream, we ran tests with just the North Atlantic ( 20 Sto72 N) portion of the global grid. All parameters and forcings were the same as the fully global run, except for the introduction of restoring buffer zones at the northern and southern boundaries (very similarly to Smith et al. (2000)). This North Atlantic sector run showed a much more realistic Gulf Stream/North Atlantic Current system than the fully global run, including a well defined flow around the Northwest Corner. The transport through the Florida Straits and in the DWBC are both about 10 Sv stronger than in the fully global model, again closely resembling the North Atlantic basin model. Whether the differences are due to water mass conditioning at the boundaries, the dynamics of Arctic inflow/outflow, or something else is currently under investi- gation. This result also shows that the problem does not lie with the bathymetry or the displaced pole grid. After the 15 year spinup was completed, we felt that the surface restoring of temperature and salinity was too strong. We have recently been able to continue the run several more years using time scales of about 25 and 90 days (5 times longer than the spinup) for restoring under ice and open ocean salinity restoring, respectively. It appears that relaxing the influence of restoring has allowed the permanent meander in the Kuroshio to disappear, and has resulted in more realistic shedding and propagation of Agulhas eddies. The role of interannual (and longer) variability has been ignored in many of the results pre- sented here. In particular, variability in the meridional overturning and heat transport, and in the general circulation schematic were not included because it was felt to be beyond the scope of this paper. In addition, the comparisons with current meters and ADCP could change somewhat since the data records are typically rather short. As expected, the model does a good job in simulating eddy variability when compared to surface height altimetry data, especially in regions of high eddy energy. Those areas where agreement is poor are primarily the result of problems with the mean flow (Gulf Stream) or too strong surface salinity restoring (Agulhas rings and Kuroshio meander). Energy levels are still too low in the relatively quiescent regions of the ocean which make up a large percentage of the total, resulting in global average variability value being about 2 cm lower than TOPEX/ERS-1. As a next step, there are a number of improvements in model physics that could be imple- mented that would likely improve such a simulation. Sensitivity studies on the bottom flow into the Atlantic and Indian basins through narrow channels are a high priority for future study, since this is an area where the model transport was underestimated by as much as an order of mag- nitude. At the very beginning of the run, the transport through Vema Channel is quite high (about 10 Sv), but drops within a couple of months to much lower values, suggesting that the topography of the channel should be modified. Some new model physics possibilities include partial bottom cells (Adcroft et al., 1997) for improved topographic interaction, a bottom boundary layer parameterization (Doscher€ and Beckmann, 2000) to help overflows, and realistic river runoff to remove the need for precipitation adjustment. There may be some benefits of using an adiabatic eddy mixing scheme (Gent and McWilliams, 1990) even at such high resolution. Also very important will be the addition of a fully coupled dynamic-thermodynamic sea ice model. 52 M.E. Maltrud, J.L. McClean / Ocean Modelling 8 (2005) 31–54
Acknowledgements
This work was supported by the Department of Energy Climate Change Prediction Program, the Office of Naval Research (N0001403WR20188 and N0001402IP20027), and the National Science Foundation (OCE-0221781). Computational resources were provided primarily by the Department of Defense High Performance Computing Modernization Office, particularly the Navy Oceanographic Office in Mississippi in the form of a Grand Challenge Award, and the Army Research Laboratory. Additional computational resources were provided by the DOE National Energy Research Scientific Computing Centre and the Advanced Computing Lab at Los Alamos National Laboratory. Thanks to Frank Bryan (NCAR) for many valuable interactions and comments on the manuscript. Thanks also to Steve Piacsek (NRLSSC) for providing the MODAS initial fields, to Detelina Ivanova (NPS) for preparing some of the surface forcing fields, to Kathleen Johnson for hand modifying much of the model topography, and to Rick Smith (LANL) and Yoshi Yoshida (CRIEPI) for important interactions.
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