Effect of wind forcing on the meridional heat transport in a coupled climate model: equilibrium response
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The effect of the ocean surface winds on the meridional heat transports is studied in a coupled model. Shutting down the global surface winds causes significant reductions in both wind-driven and thermohaline ocean circulations, resulting in a remarkable decrease in the poleward oceanic heat transport (OHT). The sea surface temperature responds with significant warming in the equator and cooling off the equator, causing an enhancement and equatorward shift in the Hadley cell. This increases the poleward atmospheric heat transport (AHT), which in turn compensates the decrease in the OHT. This compensation implies a fundamental constraint in changes of ocean–atmosphere energy transports. Several other compensation changes are also identified. For the OHT components, the changes in the Eulerian mean and bolus OHT are compensated with each other in the Southern Ocean, since a stronger wind driven Ekman transport is associated with a stronger meridional density gradient (stronger bolus circulation) and vice versa. For the AHT components, the changes in the dry static energy (DSE) and latent energy transports are compensated within the tropics (30°N/S), because a stronger Hadley cell causes a stronger equatorward convergence of moisture. In the extratropics, the changes in the mean and eddy DSE transports show perfect compensation, as a result of the equatorward shift of the Ferrell Cell and enhancement of atmospheric baroclinicity in mid-high latitudes, particularly over the North Atlantic. This work also shows how the Earth’s climate is trying to maintain the balance between two hemispheres: the ocean in the Northern Hemisphere is colder than that in the Southern Hemisphere due to much reduced northward heat transports cross the Equator in the Atlantic, therefore, the atmosphere responds to the ocean with temperature colder in the Southern Hemisphere than in the Northern Hemisphere by transporting more heat northward cross the equator over the Pacific, in association with a southward shift of the intertropical convergence zone.
KeywordsCoupled climate model Atmospheric heat transport Oceanic heat transport Hadley cell Atlantic meridional overturning circulation Bjerknes compensation
Assessing the meridional heat transport (MHT) and its oceanic and atmospheric partition is a classical question of the climate research (Carissimo et al. 1985; Vallis and Farneti 2009). As early as in 1970s, Vonder Haar and Oort (1973) have estimated the MHT using the satellite observation of radiative fluxes. One robust feature of the total MHT is that the Earth system is maintained by a hemispherically antisymmetric poleward heat transport with a peak value of about 5.5 PW (1 PW = 1015 W) at 35°N/S (e.g., Trenberth and Caron 2001). Another robust picture of the MHT is that the atmospheric heat transport (AHT) dominates poleward of about 30°N/S while the oceanic heat transport (OHT) dominates in the deep tropic (Held 2001; Wunsch 2005; Czaja and Marshall ). However, there is a highly debatable question on the relationship between the changes in the AHT and OHT. Bjerknes (1964) first suggested that if the top of atmosphere (TOA) fluxes and the oceanic heat storage did not vary too much, the total energy transport in the climate system would not vary too much either. This implies that any large variations of oceanic and atmospheric energy transports should be equal and opposite. This simple scenario has become known as Bjerknes compensation (BC). It would be a critical constraint that might reduce the degree of freedom or uncertainty of the climate system. Previous studies have shown the BC is highly dependent on the timescale, latitudes as well as the models (Shaffrey and Sutton 2006; Vellinga and Wu 2008; Farneti and Vallis 2013).
The relationship between the changes in the AHT and OHT has been studied broadly. It has been examined in the frame of internal climate variability (Shaffrey and Sutton 2006; Swaluw et al. 2007; Farneti and Vallis 2013), or under significant external forcings (Zhang and Delworth 2005; Cheng et al. 2007; Vellinga and Wu 2008; Laurian et al. 2009; Drijfhout 2010; Zhang et al. 2010). Both scenarios show the robust feature of the BC in the changes of MHT. However, the compensation “structure” and mechanisms in these two scenarios are radically different. The former (latter) suggests that the BC is valid in the high (low) latitudes. Even in the latter scenario, although the responses of the atmosphere and ocean circulations as well as their thermal structures are more or less the same in the water hosing experiments of different models, the compensation structures are also different among differ models. For example, in Vellinga and Wu (2008) (VW08) the compensation is very efficient at low latitudes and near complete at the equator, but is incomplete farther north across the northern midlatitude storm tracks. The maximum AHT change occurs right on the equator while the maximum OHT change occurs near the 20°N. In Zhang and Delworth (2005) (ZD05) the compensation situation is similar to that in VW08, but the maximum AHT change occurs around 15°–20°N. The overcompensation of AHT to OHT is insignificant in VW08 and ZD05, but significant in Zhang et al. (2010) due to strong cloud feedback in the tropics. In Cheng et al. (2007) (CBC07) the compensation is quite good from the low latitudes to 40°N. The maximum changes in AHT and OHT occur around 20°–30°N. Instead of overcompensation, there is a weak (strong) undercompensation in the tropics (poleward of 40°N).
These modelling studies suggest that the BC remains an open question because of the fundamental controversies on its applicability. The most serious problem is that currently the BC has not been validated from the observations. Even the data are available, the direct calculation of ocean and atmosphere heat transports from velocity and temperature fields would contain big errors, which might be even bigger than the absolute values of AHT and OHT (Wunsch 2005) and may result in practical infeasible to validate the BC in the real world. However, the relationship between changes of the AHT and OHT is deserved to explore extensively and thoroughly, because it might suggest some fundamental mechanisms in maintaining the stability of the Earth’s climate. Furthermore, it could help to understand the so-called reversibility of climate change after sudden big natural disasters or anthropogenic forcing.
A coupled climate model is used in this work to study the BC. Different from previous studies, this work investigates the wind effect on the changes of the MHT. A series of wind perturbation experiments are performed to examine how the wind change affects the wind-driven and thermohaline circulations and thus the OHT; and how the oceanic changes in turn affect the atmosphere circulation and the AHT. In the coupled model system, all heat transport components in the atmosphere and ocean are calculated directly from the velocity and temperature (Yang et al. 2014a). This work focuses on the equilibrium response. The transient changes in the MHT will be studied in our next work.
It is found that changes in AHT and OHT compensate very well in the mid-low latitudes when altering the ocean surface wind stress. This is different from previous studies, in which the BC is valid only in either the tropics (Vellinga and Wu 2008) or the high latitudes Atlantic (Shaffrey and Sutton 2006; Swaluw et al. 2007). In the wind experiments, the TOA flux remains nearly unchanged. Both the wind-driven and thermohaline circulations are weakened rapidly in response to the change in surface winds. The tropical (extratropical) oceans are warmed (cooled) due to the slowdown of poleward OHT. The atmospheric Hadley cell (HC) is thus enhanced and transports more heat poleward, compensating the decreased OHT. The good compensation in the tropic should be attributed to the cloud feedback, which affects the shortwave (SW) and longwave (LW) radiation at the TOA oppositely. Several other compensation changes are also identified in this work. For example, for the OHT components, the changes in the Eulerian mean and Bolus OHT are compensated with each other in the Southern Ocean. For the AHT components, the changes in the dry static energy (DSE) and latent energy transports are compensated within the tropics (30°N/S). In the extratropics, the changes in the mean and eddy DSE transports show a perfect compensation. An very important finding is that this work shows how the Earth’s climate is trying to maintain the balance between two hemispheres: the ocean in the Northern Hemisphere (NH) is colder than that in the Southern Hemisphere (SH) due to the much reduced northward heat transport cross the Equator in the Atlantic, therefore, the atmosphere responds to the ocean with colder temperature in the SH than in the NH by transporting more heat northward cross the equator over the Pacific, in association with southward shift of the intertropical convergence zone (ITCZ). This is also identified in a recent work by Marshall et al. (2013).
This paper is arranged as follows. The second section introduces the coupled model and experiments. The third section discusses the spin-up process of the model climate. The fourth section examines the equilibrium response in the mean meridional overturning circulations in both the atmosphere and ocean, and quantifies the different components in the MHT and their changes. The fifth section summarizes the compensation processes. The last section is the conclusion and discussion.
2 Model and experiments
The model used in this study is the Community Earth System Model (CESM, version 1.0) of the National Center for Atmospheric Research (NCAR). CESM is a fully coupled global climate model that provides state-of-the-art simulations of the Earth’s past, present, and future climate states (http://www2.cesm.ucar.edu/). CESM1.0 consists of five components and one coupler: the Community Atmosphere Model (Neale et al. 2013), the Community Land Model (CLM4, Lawrence et al. 2012), the Community Ice CodE (CICE4, Hunke and Lipscomb 2008), the Parallel Ocean Program (POP2, Smith et al. 2010), the Community Ice Sheet Model (Glimmer-CISM) and CESM Coupler (CPL7). CESM1.0 has been widely used and validated by the community aerosols(http://journals.ametsoc.org/page/CCSM4/CESM1).
The model grid employed in this study is T31_gx3v7. The atmospheric component CAM5 has 26 vertical levels, with the finite volume nominal 3.75° × 3.75° in the horizontal. The CAM5 is essentially a new atmospheric model with more realistic formulations of radiation, boundary layer, and aerosols (Neale et al. 2013; Meehl et al. 2013). The General features of the model formulation are given by Neale et al. (2010, 2013). The CLM4 has the same horizontal resolution as CAM5. The ocean POP2 uses the grid gx3v7, which has 60 vertical levels. The horizontal grid has a uniform 3.6° spacing in the zonal direction and non-uniform spacing in the meridional direction. It is 0.6° near the equator, extending to the maximum 3.4° poleward of 35°N/S and then decreasing towards higher latitudes. The model physics is described in details in Danabasoglu et al. (2012). The sea ice component CICE4 has the same horizontal grid as POP2. No flux adjustments are used in CESM1.0.
In the CESM POP2, the physical transport terms are partitioned into resolved and unresolved components. The resolved component, such like Eulerian mean mass and heat transport, can be easily obtained using model output temperature and Eulerian mean velocities. The unresolved components, which results from meso-scale and submeso-scale processes, are parameterized using well-recognized schemes. The meso-scale eddy transport is parameterized according to Gent and McWilliams (1990). The eddy-induced advection coefficient varies in space and time (Danabasoglu and Marshall 2007). A variable coefficient provides a better representation of changes in eddy activity resulting from variable surface momentum forcing than a constant value (Long et al. 2013), allowing the model to more realistically capture the circulation response to changing winds, particularly in the Southern Ocean (Farneti and Gent 2011). Submeso-scale eddies is parameterized by the FFH scheme (Fox-Kemper et al. 2008, 2011), depicting the restratifying effect on the mixed layer.
3 Transient responses in global climate
The Earth’s climate as a whole is determined by the net radiation flux at the TOA, which reaches a quasi-equilibrium in about 50 years, roughly equal to the transient time scale of the global ocean (Fig. 1c). It is seen that the change in the net downward SW radiation and the net outgoing LW radiation at the TOA tend to be in phase in most latitudes. That is, a decrease (increase) in the downward SW can be roughly compensated by a decrease (increase) in the upward LW, leaving a trivial change in the net radiation flux. Conceptually, a decrease in the outgoing LW corresponds to a lower surface temperature, which can result in a larger planetary albedo (for example, more sea ice in the mid-high latitudes, Fig. 2b), and then a decrease in the net downward SW. The compensation effect between the LW and SW at the TOA leads to a reduced change in the net radiation flux, suggesting that the Earth’s climate as a whole tends to be steady.
The MHT of the Earth system also reaches quasi-equilibrium at the same pace with the radiation flux at the TOA (Fig. 1d). The local balance of heat flux determines the MHT. Therefore, the timescale of MHT adjustment is also very short and well consistent with that of the net radiation flux in Fig. 1c. The OHT is calculated directly from the meridional velocity and potential temperature, while the AHT here is deduced from the difference of the net radiation flux at the TOA and the net surface heat flux over the ocean, which is not explicitly related to the ocean internal dynamics. It is seen that during the whole evolution period, the changes in the OHT and AHT tend to be out of phase in most regions. They have significant negative correlation and compensate with each other to a great extent. There is also non-trivial undercompensation between the AHT and the OHT in high latitudes. The compensation problem is our focus in this work.
The mechanisms for the AMOC change (Fig. 3) have been examined in details in Yang et al. (2014b). Briefly, when the wind-stress is reduced, the vertical convection and diffusion are weakened immediately, resulting in a salt deficit in the northern North Atlantic that prevents the deep water formation there. This triggers the AMOC reduction. As the AMOC weakens, the sea ice expends southward and melts (Fig. 2c), freshening the upper ocean that weakens the AMOC further. There is a positive feedback between the sea ice melting and AMOC weakening, which eventually results in the AMOC shutdown. The salinity advection from the south, however, plays a contrary role to the sea ice melting, salting the upper ocean in the North Atlantic between 40° and 60°N (Figure not shown). This is different from previous studies (e.g., Timmermann and Goosse 2004) which emphasize the important role of salinity advection from lower latitudes in the AMOC.
4 Equilibrium responses
4.1 Radiation balance at the TOA
4.2 Changes in the MHT
The compensation change between the AHT and OHT is consistent with that between the atmospheric and oceanic overturning circulations shown in Fig. 3. Here the components of OHT are examined. It is seen that the OHT changes in the Indo-Pacific and Atlantic are comparable (Fig. 5c). In 0.1 W run, the total OHT change peaks about 1.3 PW near 20°N (Fig. 5b), in which about 0.8 PW reduction occurs in the Indo-Pacific, representing a maximum 80 % reduction in the Indo-Pacific OHT (Fig. 5c). The remaining 0.5 PW reduction occurs in the Atlantic, which is almost equivalent to the shutdown of the Atlantic OHT. These reductions are consistent with the significant weakening of the STC and AMOC (Fig. 3). Actually, the Indo-Pacific (Atlantic) OHT can be roughly thought as the contribution by wind-driven (thermohaline) circulation. In 0.1 W run, the OHT reduction in the tropics (0.8 PW) is mainly attributed to the change in wind-driven circulation. While in mid-high latitudes, the thermohaline OHT change (0.5 PW) dominates (Fig. 5c).
4.3 Changes in the ocean temperature, salinity and density
The sea surface salinity (SSS) is also reduced remarkably throughout the whole ocean surface (Fig. 6c, d). Overall the ocean gains fresh water in the reduced wind experiments. The SSS is reduced by more than 6 psu in the North Atlantic in 0.1 W run (Fig. 6d). This eventually determines the density in the high latitudes. The sea ice plays a key role here in the ocean freshening. The total sea ice melting to ocean is nearly doubled in 0.1 W run, increasing from 0.43 Sv in the CTRL to 0.81 Sv in 0.1 W run (Figure not shown). The enhanced melting results from the more equatorward ice transport from the high latitudes North Atlantic (Fig. 2b). The significant weakening in the ocean circulations results in more sea ice output to the lower latitudes and thus melting. The freshening in the poleward of 40°S is due to less saline water upwelling in the Southern Ocean. In the deep tropics 10°S–10°N, the freshening is mainly due to more precipitation, associated with the stronger convection due to warmer SST.
The surface density change shows a similar pattern to that of SSS (Fig. 6f). The high latitude freshening effect overcomes the cooling effect on the density, resulting in a decrease in the surface density. In the deep tropics (10°S–10°N), the surface density reduces due to both the freshening and warming. In the subtropics (10°–30°N/S), the surface density increases slightly due to colder SST. In general, it is the surface salinity change that eventually contributes to the weakening of the AMOC and thus the reduced poleward OHT. This might imply a critical role of the hydrological cycle in the OHT, and further in the stability of the Earth energy balance. We will pursue this problem in our future work.
The density changes in the ocean interior (Fig. 7c, d) almost resemble the surface density changes (Fig. 6f), which is roughly barotropic kind. The density changes in the Atlantic and Indo-Pacific are similar. The extratropical density (30°N/S poleward) decreases significantly from the surface to around 500 m depth, which is mainly caused by the freshening in the upper oceans as discussed above. The meridional density gradient is thus weakened significantly. This is particularly clear in the Southern Ocean region (40°–80°S) (Fig. 7c, d), where the isopycnal levels become nearly flat in the upper ocean in 0.1 W run due to 2–3 kg/m3 decrease. The meridional density gradient is closely related to the bolus overturning circulation associated with meso-scale eddies (Marshall and Radko 2003), which is opposite to the Eulerian mean overturning circulation. The flattening of the isopycnal level will reduce the bolus circulation and thus the associated OHT.
4.4 Changes in the MOC and OHT
In 0.1 W run, all four components of the MOC are reduced significantly (Fig. 8e, f). The Eulerian mean circulation is reduced by the 80 %, in which the wind-driven part is decreased from 32 to 8 Sv and the AMOC is reduced from 20 to <4 Sv. The bolus circulation is also reduced by 80 % from the maximum 8 to <2 Sv, as a result of the flattening of the isopycnal level in the Southern Ocean. A similar situation occurs to the mixed layer submeso scale circulations, which is due to the weakening of wind-induced vertical mixing.
4.5 Changes in atmosphere static energy, mass transport and water vapour transport
The mass transport and water vapour transport are shown in Fig. 10b. The mean meridional mass transport, represented mostly by the HC, is opposite to the water vapour transport in direction, with the magnitude of the former about 100 times of the latter. The water vapour is transported by the equatorward convergent flow in the lower troposphere, that is, the lower branch of the HC in the tropics, enhancing the tropical convection through the latent heat release. It is worth noting that although the water vapour mass transport is two-order smaller in magnitude than total air mass transport, the water vapour is 100 times more efficient than dry air in transporting heat (Huang 2005). One kilogram of water vapour can deliver 2.5 × 106 J of heat (the latent heat content of water vapour L e ~ 2.5 × 106 J/kg), but one kilogram of air (water) can only release 104 J (4.18 × 104 J) of heat for a 10 °C cooling. This makes the atmospheric latent energy transport significant in the climate system. The water vapour transport in Fig. 10 is directly calculated based on the model output of vq. In the tropics it is accomplished by the mean circulation (i.e., the HC), while in the extratropics (poleward of 25°N/S) it is fulfilled mostly by the eddy activities, because the v in the vq represents mainly the HC in the tropics and then eddies in the extratropics. It is seen that the mean water vapour transport in the tropics is opposite to and also much weaker than the eddy water vapour transport in the extratropics (Fig. 10b).
The atmospheric becomes drier and its baroclinicity becomes stronger in 0.1 W run (Fig. 10c, e). These changes are very clear in the off-equator, where the lower atmosphere is about 20 % drier and the atmospheric baroclinicity (∂T/∂y) is enhanced by more than 30 %. This implies that the poleward latent heat transport will be decreased in the extratropics while the poleward eddy heat transport will be enhanced. Figure 10e shows the changes in the atmospheric potential temperature and the specific humidity in 0.1 W run. The stronger convection at the equator increases the humidity from the surface to the upper troposphere, but this is constrained within 5°N/S of the equator. Where else the water vapour decreases significantly, consistent with the colder SST that reduces the evaporation. The lower level atmospheric temperature is reduced by as much as 10 K in the mid-high latitudes, which at the same time enhances the atmospheric baroclinicity greatly. The HC is enhanced by 80 % and its ascending branch shifts southward to the equator (Fig. 10d, f) in 0.1 W run, showing a more symmetric structure in the atmospheric circulation. Consequently, the equatorward water vapour transport is enhanced. The eddy water vapour transport in the mid-high latitude is weakened slightly (Fig. 10f).
4.6 Changes in the AHT
The eddy AHT change dominates the poleward AHT anomaly in the mid-high latitudes (Fig. 11f), which is caused by the enhanced baroclinicity as shown in Fig. 10e. It is well-established that the mean circulation dominates in the tropics and eddies dominates in the mid-high latitudes (Fig. 11d, e). The eddy LE transport AHTELE is always poleward. It is small and opposite to the mean LE transport AHTMLE in the tropics. The AHTELE increases towards the higher latitudes, exceeds the AHTMLE poleward of 30°N/S and peaks at around 40°N/S. The eddy DSE transport AHTEDSE is nearly zero in the tropics and increases rapidly towards the higher latitudes, in accompany with the rapid diminishing of the mean DSE transport AHTMDSE. The AHTEDSE accounts to more than 50 % (80 %) of the total AHT in the NH (SH) mid-latitudes. Here the eddy transport includes those from both the stationary waves and transient eddies. The detailed calculations of these terms as well as the analyses can be found in Yang et al. (2014a). In 0.1 W run, the eddy contribution to the AHT comes mainly from the AHTEDSE while the AHTELE is hardly changed (Fig. 11f). The enhanced AHTEDSE is due to the stronger baroclinicity as shown in Fig. 10e. The negligible change in the AHTELE is due to drier lower atmosphere, despite the stronger eddy activities. It is interesting to see the changes in the AHTEDSE and AHTMDSE tend to compensate with each other in the off-equator (20°N/S poleward), which makes the total AHT changes in the extratropics negligible. And this eventually results in under-compensation to the OHT change (Fig. 5b).
4.7 Changes in the ITCZ and atmosphere circulations
The AHT change occurs mainly over the tropical Pacific and extratropical Atlantic. This is deduced from the changes in the surface winds and pressure, which provide clues on the relative roles of the different basins in the global AHT change (Fig. 12b). In response to the SST change, the change in surface winds is strongest over the tropical Pacific within 20°N/S and the change in surface pressure is strongest over the high latitudes Atlantic. This suggests that the AHT change associated with the HC shift occurs mainly over the tropical Pacific, and that associated with stronger eddy activities occurs mainly over the North Atlantic. This can be further confirmed by the net heat flux change in the Atmosphere (Figure not shown), which illustrates two important regions: one is located in the eastern tropical Pacific (cold tongue region) with a maximum change of 80 W/m2; the other is located over the North Atlantic (50°–70°N) with a −80 W/m2 change. This enhances our fundamental understating on the ocean basin in the global atmospheric heat budget. Furthermore, the atmospheric change over the tropics appears to be barotropic-like, since the wind change in the upper troposphere is similar but of opposite sign to that over the surface (Fig. 12b, d). The change in eddy activities over the North Atlantic is mainly contained in the lower troposphere, which is consistent with the strong baroclinic change in the wind and pressure fields. It is seen that the geopotential height over the 200 hPa is generally lowered in response to the overall cooling in the SST. In addition, Fig. 12b and d also illustrate a weakening in the zonal Walker Circulation, with a stronger anomalous convergence (divergence) over the eastern equatorial Pacific surface (200 hPa) in 0.1 W run, in accompany with the warming SST there and thus the enhanced HC.
5 Compensation and un-compensation
From the point of view of dynamical change in the atmospheric and oceanic circulations, the undercompensation of the AHT to the OHT in the NH is due to big changes in the atmospheric Ferrell Cell (Fig. 10b, d, f). The heat transport by the Ferrell Cell tends to be out of phase with the eddy heat transport. This is true for both the mean state and their changes in the reduced wind run (Figs. 11d–f, 13c). The increase in the atmospheric eddy activities is offset by the change due to the Ferrell cell, resulting in an undercompensation of the AHT to the OHT decrease (Fig. 13c). In fact, the enhancement of the Ferrell Cell is connected to the equatorward shift of the ITCZ, that is, the HC. Therefore, the changes in the tropical atmosphere also play roles in the undercompensation of the AHT to the OHT in the NH mid-high latitudes. In the SH mid-high latitudes, however, the AHT change over-compensates the OHT change. This attributes to the big decreases in the Eulerian mean OHT as well as the bolus OHT in the ACC region. The atmosphere still feels the strong ocean surface cooling and responds with enhanced baroclinicity and thus stronger southward eddy AHT (Fig. 13c), which is not offset by the mean AHT change since the Ferrell Cell is roughly unchanged (Fig. 10f).
This paper focuses on the compensation. Since both the mean MHT and their changes in the high latitudes are small, the compensation issue itself over there is not that critical. We have seen that the MHT changes at the 40°N/S poleward are less than one-third of those in the tropics even in our extreme no-wind experiment. We have also seen the consistent pictures between the dynamical changes in the atmosphere and ocean circulations and the changes in the Earth’s radiation budget in the un-compensation region. The less heat gain due to less poleward heat transport results in the less heat loss out of space. In the tropics, the good compensation should also be attributed to the clouds, which tends to affect the SW and LW oppositely, mitigating the poleward AHT change by enhancing the equatorward LE transport.
Finally the detailed cloud changes are examined in Fig. 14. The cloud amount increases over the equator and decreases along the 10°N (Fig. 14a), as a result of the southward shift of the ITCZ. Figure 14b further shows that it is the high cloud that decreases the most, which occurs mainly in the central-western Pacific of northern equator. The high cloud increase occurs mainly in the eastern Pacific to the south of equator (Fig. 14c). This is consistent with the weakening of the Walker circulation (Fig. 12). In the subtropics, the cloud change is trivial. In the high latitudes, the low cloud is generally reduced (Fig. 14a).
6 Conclusions and discussions
The opposite change in the atmospheric and oceanic meridional circulations is one of the key mechanisms to contain the Earth’s climate variability. In this work the ocean meridional circulation changes, including both the STC and the AMOC, are mainly determined by the dynamical changes in the Ekman pumping, the Ekman advection as well as the sea ice advection in the mid-high latitudes. We also see that the salinity change determines the density change in the extratropics and eventually plays an important role in the AMOC change. This implies a critical role of hydrological cycle in the global energy adjustment. The ocean cooling in the high latitudes in response to the weakening of surface wind forcing tends to increase the density in the high latitudes. So the salinity decrease has to play a critical role in the AMOC weakening. Otherwise, the AMOC could be enhanced, and the changes in the AHT and OHT would be in phase instead of out of phase. There would be no so-called BC, and thus the global climate may drift dramatically due to lack of this restoring mechanism in the global energy adjustment. The salinity change in the high latitudes can be caused by the changes in the meridional salinity advection associated with the AMOC, the sea ice melting/formation, river runoff and the fresh water flux due to evaporation minus precipitation. The detailed mechanisms are examined in Yang et al. (2014b).
The good compensation between the AHT and OHT changes in the tropics should be also attributed largely to the good cloud feedback there. The big temperature change in the tropics causes dramatic change in the atmospheric convection and thus the clouds, which affects the SW and LW oppositely and tends to maintain the stability of the net radiation flux at the TOA. This ensures the perfect out of phase change in the AHT and OHT (Bjerknes 1964). Koll and Abbot (2013) have shown that the cloud feedback is critical for maintaining constant equatorial SST as OHT changes. We show in this work that the cloud feedback is critical for maintaining the constant net radiation flux at the TOA in the tropics, even with big SST change. In the mid-high latitudes, the cloud feedback is very weak so that the compensation in the OHT and AHT is imperfect. Cloud feedback is a critical stabilizing factor for the Earth’s climate. We will further the study in our next work.
The atmospheric eddies play very important roles in energy balance in the mid-high latitudes. Without eddies’ contributions, the energy deficit will be much severer. This also suggests a critical role of the atmospheric eddies in maintaining the high latitude Earth’s climate stability, particularly the activities of eddies over the North Atlantic (Nakamura et al. 1994). By examining the changes in the atmospheric wind and pressure, we have found clues on the relative roles of the different ocean basins in the global atmospheric energy transport. The AHT change occurs mainly over the tropical Pacific and extratropical Atlantic, which are related to the mean circulation in the tropics and eddies’ activities in the mid-high latitudes, respectively. In this work, the atmospheric eddies’ contributions may be underestimated since the monthly data is used and the eddy AHT is obtained as the residual of the total AHT minus the mean AHT. This may also contribute to the un-compensation between the AHT and OHT changes. The role of eddies in global energy balance is worth investigating further.
This work focuses only on the equilibrium response in the global energy adjustment, providing the detailed analyses on 0.1 W run. The 0.3 and 0.5 W runs have the similar results that are not shown here for simplicity. Our next work is to explore the transient response of the AHT and OHT, in which we may disclose the interaction mechanism between the STC in the Pacific and the AMOC in the Atlantic. Thus the roles of different ocean basins in the different stages of the climate change will be illuminated. Ensemble wind perturbation experiments focusing on the transient responses have been performed and the thorough analyses are underway.
We are grateful to Profs. Z. Liu and R. Huang for invaluable suggestions and discussions. This work is jointly supported by the NSF of China (Nos. 91337106, 41376007, 41176002, 40976007), the National Basic Research Program of China (No. 2012CB955200), the Special Fund for Meteorological Scientific Research in the Public Interest of CMA (No. GYHY201006022) and the Norwegian Research Council through the East Asian DecCen project (No. 193690/S30). All the experiments are performed on the supercomputer at the LaCOAS, Peking University.
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