The Role of Oscillating Southern Hemisphere Westerly Winds: Global Ocean Circulation

15 MARCH 2020 C H E O N A N D K U G 2111

WOO GEUN CHEON Maritime Technology Research Institute, Agency for Defense Development, Changwon, South Korea

JONG-SEONG KUG School of Environmental Science and Engineering, Pohang University of Science and Technology (POSTECH), Pohang, South Korea

(Manuscript received 23 May 2019, in ﬁnal form 5 December 2019)

ABSTRACT

In the framework of a sea ice–ocean general circulation model coupled to an energy balance atmospheric model, an intensity oscillation of Southern Hemisphere (SH) westerly winds affects the global ocean circulation via not only the buoyancy-driven teleconnection (BDT) mode but also the Ekman-driven teleconnection (EDT) mode. The BDT mode is activated by the SH air–sea ice–ocean interactions such as polynyas and oceanic convection. The ensuing variation in the Antarctic meridional overturning circulation (MOC) that is indicative of the Antarctic Bottom Water (AABW) formation exerts a significant influence on the abyssal circulation of the globe, particularly the Pacific. This controls the bipolar seesaw balance between deep and bottom waters at the equator. The EDT mode controlled by northward Ekman transport under the oscillating SH westerly winds generates a signal that propagates northward along the upper ocean and passes through the equator. The variation in the western boundary current (WBC) is much stronger in the North Atlantic than in the North Pacific, which appears to be associated with the relatively strong and persistent Mindanao Current (i.e., the southward flowing WBC of the North Pacific tropical gyre). The North Atlantic Deep Water (NADW) formation is controlled by salt advected northward by the North Atlantic WBC.

1. Introduction Klinger and Cruz 2009). Herein, impacts of oscillating SH westerly winds on global ocean circulation are investigated This paper extends the study of Cheon et al. (2018, from two perspectives: 1) the buoyancy-driven telecon- hereinafter C18), which showed the following: 1) the nection (BDT) mode that results from SO sea ice/ocean formation rate of coastal polynyas as being in linear interaction and directly controls AABW formation and correlation with the oscillating intensity of Southern 2) the Ekman-driven teleconnection (EDT) mode that Hemisphere (SH) westerly winds, 2) the favorable con- results from northward Ekman transport and directly conditions of the Weddell Sea in relation to open-ocean trols the upwelling of Circumpolar Deep Water (CDW) polynyas, and 3) the importance of recovery time of deep and outflow of North Atlantic Deep Water (NADW), as ocean heat content, which controls the occurrence fre- well as remotely influencing the North Pacific and Atlantic quency of open-ocean polynya. Coastal polynyas and surface WBCs and the Atlantic meridional overturning open-ocean polynyas are closely linked to near-boundary circulation (MOC). convection and open-ocean deep convection, respec- The SH westerly winds are closely linked to the south- tively; therefore, they play a major role in the Antarctic ern annular mode (SAM) in association with their fluc- Bottom Water (AABW) formation. Moreover, it has tuations in the latitudinal position and strength (Gong been proposed that the SH westerly winds controlling and Wang 1999; Thompson and Wallace 2000). At the northward Ekman transport in the Southern Ocean (SO) surface, they are much stronger in the SH than in the affect the surface western boundary currents (WBCs) Northern Hemisphere westerly winds (Gille 2005)owing in the North Pacific and Atlantic (McDermott 1996; to a unique zonal band that is not interrupted by the continent in the SO. The strongest surface westerly winds Corresponding author: Woo Geun Cheon, [email protected] are located in the latitude band between 558 and 508S

DOI: 10.1175/JCLI-D-19-0364.1 Ó 2020 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses). Unauthenticated | Downloaded 10/03/21 10:44 PM UTC 2112 JOURNAL OF CLIMATE VOLUME 33

(Gille 2005; Gent 2016) and are shifted poleward (equa- Regardless of where it forms, 60%–80% of its volume torward) and intensified (weakened) in a positive (nega- eventually ends up in the Pacific Ocean, 20%–30% in tive) phase of the SAM (Gong and Wang 1999; Thompson the Indian Ocean, and 5%–10% in the Atlantic Ocean, and Wallace 2000). According to analysis of satellite ob- which is in agreement with the observation data of servation and atmospheric reanalysis data, the maximum Johnson (2008). It covers much of the ocean floor zonal wind stress over the SO increased by at least 20% globally, except for in the Arctic Ocean and the northern between 1980 and 2010 (Swart and Fyfe 2012; Bracegirdle North Atlantic Ocean, where the NADW overlies the et al. 2013; Farneti et al. 2015; Gent 2016) in association ocean bottom (Mantyla and Reid 1983; Orsi et al. 2001). with SAM variation. The SAM changed from a prolonged The NADW, whose upper water masses form in the negative phase beginning in the early 1960s to a positive Labrador Sea and lower water masses form in the phase in the 1980s and has maintained its positive phase Greenland–Iceland–Norway (GIN) Seas, flows south- until the present day (Gordon et al. 2007; Hartmann et al. ward as the deep WBC with the mean volume transport 2 2013; Cheon and Gordon 2019). As illustrated in C18,the ranging from 12 to 17 Sv (1 Sv [ 106 m3 s 1)(Schmitz SH surface westerly wind variability is highly correlated and McCartney 1993; Ganachaud and Wunsch 2000; with the SAM; the trend of this relationship is expected to Lumpkin and Speer 2003; Schott et al. 2004). Analysis be positive in the future climate (Zheng et al. 2013). of a moored array in the western basin of the equatorial Therefore, it is important to investigate the response of Atlantic (Hall et al. 1997) shows that the transition layer global ocean circulation to the SH westerly wind fluctua- between the lower NADW and the AABW rises and falls tions associated with the SAM. synchronously with the quasi-annual cycle of AABW In the study by Cheon et al. (2014), who used a sea ice– transport. This so-called ‘‘bipolar seesaw balance’’ be- ocean coupled general circulation model (GCM), SH tween two water masses was also simulated on the mul- westerly winds were proposed to trigger the formation of ticentennial time scale by the sea ice–ocean coupled open-ocean polynya in the Weddell Sea. The intensified models with salt flux directly manipulated (Brix and westerly winds induce spinup of the cyclonic Weddell Gerdes 2003) or indirectly manipulated by the dynamic Gyre, weakening and raising the pycnocline and helping air–sea ice interaction (Cheon and Stössel 2009) in the the relatively warm Weddell Deep Water (WDW) to rise Weddell Sea. The inverse correlation between the deep to the surface. The upwelled, warm WDW melts the WBCs of NADW and AABW associated with this bi- overlying sea ice, or prevents it from forming, generating polar seesaw balance was explained by Cheon and open-ocean polynya and deep convection. This hypoth- Stössel (2009) from the viewpoint of conservation of esis was clearly demonstrated by the analyses of Cheon potential vorticity. and Gordon (2019) using diverse, long-term observation The SO has a unique zonal band that is not interrupted and reanalysis data. They showed that the relatively by the continent and passes through the Drake Passage of strong Maud Rise polynya observed from September to about 2000-m depth between the Antarctic Peninsula and November of 2017 was generated by the combined effect South America. In this zonal band, zonal pressure gra- of weak stratification near the Maud Rise and the wind- dients averaged around the globe are zero above this induced spinup of the cyclonic Weddell Gyre that acti- depth; therefore, there cannot be a net geostrophically vated the WDW eddies significantly in the southwestern balanced meridional flow. The overlying SH westerly flank of the Maud Rise. As proposed in C18, in addition winds push a large amount of very cold, fresh Antarctic to contributing to the formation of coastal polynya in the Surface Water (AASW) away from the Antarctic conti- western Weddell and Ross Seas, the SH westerly winds nent via northward Ekman transport, which is ageos- play an important role in the formation of AABW by way trophically balanced and not affected by the above of open-ocean deep convection and near-boundary con- constraint. Since any subsurface return flow must be vection, which is referred to as the BDT mode in this geostrophically balanced and therefore cannot exist, the paper. Zanowski and Hallberg (2017) showed that the net poleward geostrophic flow balancing this Ekman Weddell Polynya generated by the increased diapycnal transport exists only below depths where meridional to- diffusivity in a coupled climate model enhanced the rate pographic barriers such as the East Scotia Ridge support of AABW formation, and the resulting signal propagated the zonal pressure gradients (Toggweiler and Samuels in the deep and abyssal oceans via large-scale waves and 1995). This deep southward return flow is mixed with the advection transports, affecting bottom water properties CDW, which upwells south of the Antarctic Circumpolar of the global ocean. Current (ACC). Numerous modeling studies have shown According to the modeling study of van Sebille et al. not only the impact of the SH westerly on the formation (2013), AABW that has formed in the Weddell Sea, rate and physical properties of Antarctic Intermediate Ross Sea, and Eastern Antarctica spreads northward. Water (AAIW) (Ribbe and Tomczak 1997; Sørensen

Unauthenticated | Downloaded 10/03/21 10:44 PM UTC 15 MARCH 2020 C H E O N A N D K U G 2113 et al. 2001; Oke and England 2004; Santoso and England mode, which remotely affects the WBCs of the North 2004) but also its remote influence on the Atlantic MOC Pacific and North Atlantic and the Atlantic MOC. A (Toggweiler and Samuels 1993, 1995; McDermott 1996; summary and discussion are provided in section 4. Rahmstorf and England 1997; Gnanadesikan 1999; Gnanadesikan and Hallberg 2000; Gnanadesikan et al. 2007). However, the SO eddy compensation simulated in 2. Model eddy-resolving models or in non-eddy-resolving models a. Model and experimental design with the Gent–McWilliams (GM) scheme (Gent and McWilliams 1990), whose coefficient is variable in space This study employs the Modular Ocean Model version 4 and time, is known to weaken the influence of the SH (MOM4) of the Geophysical Fluid Dynamics Laboratory westerly winds on the SO mean MOC (Farneti and Gent (GFDL), in which the primitive equation ocean model 2011; Hofmann and Morales-Maqueda 2011; Gent and (Griffies et al. 2004) is coupled with the dynamic and Danabasoglu 2011; Gent 2016) and the ACC (Hallberg thermodynamic sea ice model (Winton 2000), the two- and Gnanadesikan 2006). dimensional global atmosphere energy balance model Regardless of the SO eddy compensation effect, the (Gerdes et al. 2006), and the land model (Anderson et al. northward Ekman transport in the SO is linearly pro- 2004). The ocean model covers the global ocean from 808S portional to the strength of the overlying westerly winds to 908N with horizontal resolution of 2831/38–18 in lon- (Gent 2016), which is a key factor controlling the EDT gitude and latitude; it has 50 semivariable layers to 5500-m mode. According to previous studies using idealized depth with the bottom layer following the actual topog- ocean models with single- and two-basin geometries raphydata(Smith and Sandwell 1997). The thickness of (McDermott 1996) and a coarse-resolution ocean GCM the upper 22 layers is uniform at 10 m, and that of the 28 with realistic bottom topography (Klinger and Cruz lower layers gradually increases to about 400 m at 5500-m 2009), the signal generated by the intensified SH westerly depth. The mesoscale eddy transport is parameterized on winds can propagate in the upper ocean to the North isopycnal surfaces following the GM scheme (Gent and Atlantic and Pacific. This remotely affects each region’s McWilliams 1990). The GM and Redi diffusivity coeffi- 2 WBCs, which are mainly controlled by local wind stress cients are capped at 600 m2 s 1, which is generally used in curl. In both modeling results, the increase rate of the the standard version of GFDL climate model version 2.1 North Atlantic WBC was much larger than that of the (CM2.1), in order to simulate the realistic volume trans- North Pacific when the SH westerly winds were uniformly port of the ACC under the observed wind stress strength. intensified in the zonal direction. However, the cause of When the SH westerly winds are intensified, the GM co- this different response has not been investigated, and efficient is expected to increase, which causes the SO eddy neither has the detailed process by which the Atlantic MOC to turn in the direction opposite to the SO mean MOC remotely responds to varying SH westerly winds. flow MOC (Gent 2016). Therefore, under intensifying SH These aspects are investigated in the present paper from westerly winds, capping the GM and Redi diffusivity co- 2 the perspective of the EDT mode. Whether or not the efficients at 600 m2 s 1 can act to amplify the volume Antarctic MOC that is activated by the BDT mode can transport of the ACC owing to the underestimated eddy affect the Atlantic MOC is a further issue addressed in activity. This is assumed to decrease the meridional dis- this study. This is considered in association with the bi- sipation of potential energy reserved in the Antarctic polar seesaw balance between the AABW and NADW Polar Front and thus enhance the meridional density (Hall et al. 1997; Broecker 1998; Brix and Gerdes 2003; gradient from the surface downward across the ACC. Cheon and Stössel 2009). However, as previously stated, the northward Ekman This paper aims to investigate the responses of the transport, an important factor associated with the EDT global ocean circulation to the oscillating SH westerly mode, is in a linear correlation with the strength of SH winds by use of the BDT and EDT modes. The model surface westerly winds regardless of the effect of SO eddy configuration and experimental setup are described in compensation. Moreover, in this study, the Antarctic the following section, and the basic states of the global MOC that is indicative of the formation rate of the ocean’s properties and simulation of circulation in the AABW is regulated by the SO air–sea ice–ocean inter- absence of oscillating wind forcing in the control simu- actions associated with the BDT mode. Therefore, the lation (CTRL) are then described. In section 3 we in- uniform GM and Redi diffusivity coefficients are judged vestigate variations of the Antarctic MOC indicative of to be suitable for this type of sensitivity experiment. The the formation rate of AABW, the abyssal flows, and the sea ice model consists of a dynamic component including transition layer between the AABW and NADW in as- the viscous-plastic constitutive law (Hibler 1979)anda sociation with the BDT mode, and segues into the EDT thermodynamic component that follows the formulation

Unauthenticated | Downloaded 10/03/21 10:44 PM UTC 2114 JOURNAL OF CLIMATE VOLUME 33 of Winton (2000), that is, with one snow layer and two sea et al. 2013) and the reanalysis data derived from the Simple ice layers and run on the same grid as the ocean model. Ocean Data Assimilation version 3.3.1 (SODA3.3.1, Whereas the atmosphere energy balance model provides downloaded from http://www.atmos.umd.edu/;ocean/ thermodynamic forcing for the sea ice–ocean model, the index_files/soda3.3.1_mn_download.htm). Thus, we as- wind field and precipitation are derived directly from the sess whether or not the simulated global ocean is suffi- ERA-15 reanalysis data of the European Centre for ciently close to reality and is appropriate for this study. The Medium-Range Weather Forecasts (ECMWF). potential temperature and salinity at the surface and 4500- As stated in C18, one standard deviation of the anom- m depth are averaged over 200 years in CTRL; these are alous zonal wind stress derived from the monthly mean of compared with those derived from WOA13 (Fig. 1). The the ECMWF ERA-40 and ERA-Interim data corresponds warm pools and cold tongues in the equatorial Pacific and to about 49% of the climatological range. The current SO Atlantic are well simulated, as are the high-salinity regions has experienced a prolonged positive SAM with its pre- in the subtropics. In the midlatitudes the simulated surface ceding negative mode beginning around 1965 (Gordon temperature and salinity are lower overall than the ob- et al. 2007). This indicates that the most recent SAM pe- servations, except for in the northwestern Pacific and riod is longer than 50 years, whereas in the early twentieth Atlantic north of 408N, which show positive surface tem- century the periods were less than 30 years. In this study, perature and salinity biases (Figs. 1c,f). As shown in Fig. 2, we employ two experimental designs: SW050P02, in which this results from WBCs such as the Kuroshio, Gulf Stream, the zonal wind stress oscillating within the range of 650% and Brazil Current being simulated as slightly stronger over 50 years (two cycles per century) is added to its than the reanalysis data. Since the surface WBCs play a original value; and SW050P04, in which the zonal wind key role in transporting the warm and saline water masses stress oscillating with the same amplitude over 25 years from subtropics to the high latitudes, the relatively strong (four cycles per century) is added to its original value. simulation of WBCs in CTRL causes the surface water These oscillations are applied only to positive values of the to be colder and fresher in most midlatitudes but to be SH zonal wind stress that is indicative of the SH westerly warmer and saltier in the northwestern Pacific and Atlantic winds (refer to Fig. 2 of C18). After 600 years of spinup, the northof408N, as compared with the observations. Owing model was run with each wind forcing for another 200 to the limitations of the coarse-resolution model, the years. In C18, six experimental designs were employed to simulated ACC in CTRL does not show the meandering investigate the detailed responses of coastal and open- feature that is captured in SODA 3.3.1 of the eddy- ocean polynyas in the Weddell and Ross Seas to diversely permitting model. oscillating wind forcing consisting of two amplitudes and Compared with the observations, the simulated abyssal three periods. However, the present study employs only water is clearly warmer and saltier in the Atlantic and is the two experimental designs (among the six designs used colder and fresher in the Weddell Sea (Figs. 1i,l), which in C18) that are closest to the main characteristics of the results from the enhanced formation of the NADW SAM. Here we focus on the diverse responses of the global (characterized by the relatively warm and salty water ocean circulation from the perspective of BDT and EDT. mass) and of the AABW (characterized by the relatively As presented in SW050P02 and SW050P04 in C18,thein- cold and freshwater mass). As shown in Fig. 3,thisisdueto tensifying (weakening) SH westerly winds increase (de- the Atlantic and Antarctic MOCs being simulated as crease) the formation rate of coastal polynya and the stronger in CTRL than they are in the reanalysis data. ensuing ice-to-ocean salt flux in the western Weddell and Although persistent, larger-scale open-ocean polynya and Ross coasts, thereby enhancing (weakening) near-boundary deep convection (as confirmed in C18) never occurred in convection. Open-ocean polynyas occur with the wind CTRL, some small episodes of convection associated with forcing being in a positive phase. This leads to open-ocean the convection scheme of Rahmstorf (1993), together with 2 deep convection, which has much larger impact on the the GM and Redi coefficients being capped at 600 m2 s 1, formation rate of AABW than near-boundary convection are the causes of the relatively strong Antarctic MOC. The associated with coastal polynyas. The SO convection trig- simulated Atlantic MOC of about 25 Sv is 2 Sv stronger gered by the oscillating SH westerly winds is a major forcing than the reanalysis product but is intermediate within the of the BDT mode. range of modeled historical MOC in phase 5 of the Coupled Model Intercomparison Project (CMIP5) (Cheng b. Model evaluation et al. 2013). Its maximum core is located farther north, In this section the basic state of the global ocean’s which is more reasonable than the reanalysis data because properties and circulation simulated in CTRL are com- the NADW predominantly forms in the Labrador Sea and pared with the observation data derived from the World Greenland Sea (Killworth 1983; Marshall and Schott Ocean Atlas 2013 (WOA13; Locarnini et al. 2013; Zweng 1999). The NADW outflow passing through 308S reaches

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FIG. 1. Horizontal distributions of annul mean (left) potential temperature and (right) salinity at (a)–(f) the surface and (g)–(l) 4500-m depth, derived from (a),(d),(g),(j) the World Ocean Atlas 2013 (WOA13) and (b),(e),(h),(k) CTRL. The variables derived from CTRL are averaged over the whole period (200 yr), and (c),(f),(i),(l) the differences between CTRL and WOA13 are presented. Note that even for the same variable the color bars for the surface and 4500-m depth are different.

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FIG. 2. Surface current ﬁelds derived from (a) the reanalysis data (SODA3.3.1) averaged over 36 years (1980–2015) and (b) CTRL averaged over 200 years, and (c) their difference.

18.7 Sv, which is in good agreement with the observed es- used in this study is appropriate for this type of sensitivity timate [18 Sv according to Dong et al. (2009)]. Although not experiment. shown here, the simulated ACC of about 192.8 Sv is larger than the observed estimates, which span from 110 to 150 Sv (Whitworth et al. 1982; Orsi et al. 1995; Cunningham et al. 3. Results 2003), but it is in line with the simulated estimates (Hallberg a. BDT mode and Gnanadesikan 2006; Kuhlbrodt et al. 2012). According to the criteria of England (1993), namely a variation in The intensiﬁcation of SH westerly winds plays an global mean bottom water properties of less than 0.018C important role in generating open-ocean polynyas and 2 2 (100 yr) 1 for temperature and 0.001 psu (100 yr) 1 for deep convection in the SO (Cheon et al. 2014, 2015, salinity, the bottom water mass has not completely reached 2018; Cheon and Gordon 2019). Figure 4 illustrates the an equilibrium state during the 600-yr spinup period; seesaw balance between the Antarctic MOC and up- however, most of the main features of the global thermo- welling of the CDW, as well as the relationship between haline circulation are within reasonable ranges. The SO sea the occurrence of open-ocean polynyas in the Weddell ice also reached an equilibrium state and was within a and Ross Seas and the strength of the Antarctic MOC. reasonable range as described in C18. Therefore, the model The clockwise MOC of the SO is often described as a

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FIG. 3. The (left) Antarctic and (right) Atlantic MOCs derived from (a),(b) the reanalysis data (SODA3.3.1) averaged over 36 years (1980–2015) and (c),(d) CTRL averaged over 200 years.

Deacon cell (hereafter denoted as ‘‘Deacon MOC’’ in this the immediate area where open-ocean polynyas occur paper) and consists of upwelling of the CDW (Figs. 3a,c) decreases below 40%. In SW050P02, the Weddell Sea controlled by the SH westerly winds and subduction of the open-ocean polynya (WSOP) occurs regularly around the Antarctic Intermediate Water driven by the buoyancy loss peak of each cycle, whereas the Ross Sea open-ocean (Speer et al. 2000). As the SH westerly winds intensify polynya (RSOP) occurs only once around the first peak. (weaken), the underlying northward Ekman transport In SW050P04, the period is half that of SW050P02, and increases (decreases), which intensifies (weakens) up- the WSOP occurs in every cycle except the second cycle; welling of the CDW in order to balance the Ekman whereas the RSOP, as it does in SW050P02, occurs just transport. Since the northern limb of the Antarctic MOC once, in the first cycle. As discussed in C18,theunfa- and the southern limb of the Deacon MOC are in contact vorable conditions for the RSOP are due to the relatively and rotate in opposite directions, there may exist a seesaw strong stratification of the Ross Sea and the Ross gyre balance between these MOCs. Figure 4a shows the in- being compressed by the southward shift of ACC fronts. ternal variations of 1) the Antarctic MOC calculated using Moreover, owing to intensified negative wind stress curl, the minimum value of the counterclockwise SO over- the spinup of the Weddell Gyre activates the WDW turning cell and 2) the Deacon MOC calculated using the eddies in the southwestern flank of the Maud Rise of the maximum value of clockwise SO overturning in CTRL. seamount rising up to 2000-m depth (Cheon and Gordon The Deacon MOC precedes the Antarctic MOC by about 2019), which subsequently weakens the pycnocline and 2–3 years, and the two time series are positively correlated. provides the favorable conditions for the WSOP. However, since strengthening of the Deacon (Antarctic) The response of the Antarctic MOC to the oscillating MOC corresponds to a positive (negative) anomaly, such wind forcing consists of two impacts: 1) oceanic con- positive correlation supports the seesaw balance (i.e., an- vection associated with coastal and open-ocean po- ticorrelation) between the Antarctic MOC and upwelling lynyas and 2) the seesaw balance driven by upwelling of of the CDW. The occurrence of open-ocean polynya is the CDW. The coastal polynyas in the Weddell and Ross estimated by the criterion that the sea ice concentration in Seas are, as discussed in C18, in a linear correlation with

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FIG. 4. (a) Variation in the detrended Deacon (green solid) and Antarctic (blue solid) MOCs of CTRL. The Deacon MOC is estimated by the maximum values over 608–308S, and the Antarctic MOC is estimated by the minimum values over 808–608S. The positive (negative) anomalies of the Deacon (Antarctic) MOC indicate an intensifying MOC, whereas its negative (positive) anomalies indicate a weakening MOC. Variation in the winter- mean sea ice concentration of the Weddell Sea (red solid) and Ross Sea (red dashed) and the Antarctic MOC (blue solid) for (b) SW050P02 and (c) SW050P04. The sea ice concentration is averaged over 808–608S, 508–208W in the Weddell Sea and 808–688S, 1808–1608W in the Ross Sea. Time series of the maximum zonal-mean wind stress between 808 and 308S (thin dashed black lines) are presented for reference. the oscillating wind forcing. Since their impact on the intensifying upwelling; however, it is significantly inten- Antarctic MOC is smaller than the impact of open-ocean sified not only by the occurrence of open-ocean polynya polynyas, an assessment of the coastal polynyas is not but also by the increased formation of coastal polynyas. presented here. In the first cycle, as the SH westerly winds When the SH westerly winds weaken, the Antarctic MOC intensify, the Antarctic MOC weakens in both SW050P02 is intensified by the combined effect of the seesaw balance and SW050P04 in association with the seesaw balance driven by weakened upwelling with the open-ocean po- driven by the intensifying upwelling. However, its weak- lynyas and the deep convection remaining. This over- ening is insignificant because the oscillating wind forcing powers the near-boundary convection, which is weakened being in a linear correlation with the formation of coastal by the decreased formation of coastal polynyas and acts in polynyas simultaneously intensifies near-boundary con- the opposite direction. This process is repeated through- vection and thus the Antarctic MOC. The Antarctic out the entire period and provides the signal to be prop- MOC suddenly begins intensifying drastically in year 9 of agated northward through the abyssal ocean, which SW050P02 and in year 7 of SW050P04 as a result of the initiates the BDT mode. occurrence of open-ocean polynyas and deep convection. Figure 5a shows the current and age of water (AOW) It then again weakens with the intensifying winds until in CTRL at depths between 4400 and 5500 m where the the WSOP reoccurs in the second cycle. In summary, when AABW masses are present. The AOW indicates how the SH westerly winds intensify, the Antarctic MOC is old the water mass is after sinking from the ocean sur- generally weakened by the seesaw balance driven by the face. Since the Pacific and Indian Oceans do not have

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FIG. 5. (a) Horizontal distributions of the annual-mean abyssal current and AOW of CTRL averaged over 4400–5500-m depths over 200 years, and the differences between CTRL and SW050P02 at years (b) 12, (c) 35, (d) 50, (e) 62, (f) 85, and (g) 100.

Unauthenticated | Downloaded 10/03/21 10:44 PM UTC 2120 JOURNAL OF CLIMATE VOLUME 33 any predominant locations for deep water formation, (van Sebille et al. 2013; Zanowski and Hallberg 2017). In very old bottom water masses occupy the North Pacific year 50 (Fig. 5d), when the first cycle finishes, the en- and the eastern Indian Oceans. Relatively young bottom hanced bottom circulation substantially diminishes from water masses indicative of the AABW are observed in the south, whereas the signal reaches farther north to the southwestern Indian and Pacific Oceans as well as 308N in the Pacific. In the Atlantic and Indian Oceans in the South Atlantic Ocean. The Mid-Atlantic Ridge the signals are also propagated farther north. In year 62 plays an important role in determining the pathways of the oscillating wind forcing reaches the peak of its sec- the AABW in the Atlantic (i.e., one in the western re- ond cycle, and the Antarctic MOC begins intensifying gion of the ocean divided by the ridge and the other in with the reoccurrence of WSOP (Fig. 4b). Since the the eastern region); this is in broad agreement with signal is just recently activated and thus remains south of previous studies (Mantyla and Reid 1983; Stephens and the ACC fronts, the enhanced bottom circulation keeps Marshall 2000; Morozov et al. 2010). Orographic fea- diminishing (Fig. 5e). As shown in Figs. 5f and 5g, the tures also have a determining impact on the propagation bottom circulation is again intensified as the newly and of abyssal waters of the Pacific. The AABW flowing substantially forming AABW around Antarctica sinks to northward along the Kermadec-Tonga and Louisville the bottom and pushes the preexisting AABW north- Ridges systems bifurcates at the southern tip of the ward. The change in bottom circulation is not as large as Manihiki Plateau, with one portion passing through the that in the first cycle because only the WSOP occurs and Samoan Passage and the other passing anticlockwise its scale is relatively small (Fig. 4b). around the eastern flank of the Manihiki Plateau. They The abyssal WBCs passing through the equator show join at the southern tip of the Gilbert Ridge and again the BDT mode-induced oscillation with lags of about 30 bifurcate. The Pacific bottom current fields simulated years in the Pacific and of about 40 years in the Atlantic in CTRL are in broad agreement with observations (Figs. 6a,c). Owing to the NADW occupying the deep (Mantyla and Reid 1983; Roemmich et al. 1996; Kawabe Atlantic, the response of the AABW passing through the et al. 2003). equator to the BDT mode is slower and smaller in the The variations in AOW and current in SW050P02 Atlantic than it is in the Pacific. The variations in MOC (Figs. 5b–g) illustrate representative responses of the estimated at the equatorial Atlantic clearly illustrate the abyssal ocean to the BDT mode during the first 100 BDT mode-induced bipolar seesaw balance (Figs. 6b,d), years. In year 12 (Fig. 5b), when the wind forcing reaches which is in phase with the variation of the abyssal WBC. its first positive peak, although the Antarctic MOC is Therefore, under the oscillating SH westerly winds, the intensifying with the occurrence of WSOP and RSOP, bipolar seesaw balance between deep and bottom waters the signal activated by the BDT mode does not reach the at the equatorial Atlantic is controlled by the BDT mode- abyssal ocean north of 408S. In year 35 (Fig. 5c), when induced AABW formation. the wind forcing approaches its first negative peak, most Figure 7a shows the marked southward deep WBCs at of the AABW masses that begin newly forming from 2000-m depth in the equatorial Atlantic. To show their year 10 enter the Pacific and flow along the bottom representative responses to the oscillating SH westerly western boundary, affecting the abyssal water south of winds, the 16th and the 40th years in SW050P02 were 208N via advection. Although not shown here, the newly selected for the positive phase and negative phase, re- formed AABW with intensifying SH westerly winds in spectively. As shown in Figs. 7b and 7c, the intensifying the Weddell and Ross Seas occupies most of the SO (weakening) SH westerly winds modulate the Ekman bottom south of the ACC fronts. Moreover, the rela- divergence-driven upwelling that increases (decreases) tively young (about 80–250 years old) water masses oc- the NADW outflow. The temporal variations in the cupying the SO bottom before the application of the Atlantic deep WBC and the NADW outflow are pre- oscillating wind forcing come out through the ACC sented in Figs. 7d and 7e, with the meridional velocity fronts, reaching the South Equatorial Pacific at about estimated at 2000-m depth near the northeastern coast of 108S 25 years after activation of the BDT mode. The South America (58N, 458W) and the maximum Atlantic abyssal circulation in the Pacific is drastically intensi- MOC estimated at 308S. The positive (negative) anomaly fied, e.g., the northward velocity of the abyssal currents of the meridional velocity indicates a weakening (inten- averaged over the eastern flank of Manihiki Plateau sifying) southward deep WBCs in the equatorial Atlantic, (108–58S, 1608–1558W) quadruples from about 0.006 to whereas the positive (negative) anomaly of the NADW 2 0.024 m s 1. In contrast to the Pacific, the signals in the outflow indicates an increase (decrease) in the NADW Atlantic and Indian Oceans are propagated only to 308S. passing through 308S. When the experimental setup em- The northward AABW flows prevailing over the Pacific ployed in this study was designed, the BDT mode- are consistent with those of previous modeling studies induced change in the AABW formation was expected

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FIG. 6. Variation in the meridional velocities averaged over 4400–5500-m depths in the Atlantic (58S–58N, 508– 308W; red solid) and in the Pacific (58S–58N, 1708–1808E; blue solid), and time–depth cross sections of meridional overturning streamfunctions estimated at 308S for (a),(b) SW050P02 and (c),(d) SW050P04. Time series of the maximum zonal-mean wind stress between 808 and 308S (thin dashed black lines) are presented for reference. to affect not only the NADW outflow but also the the majority of its volume transport. It then changes Atlantic MOC. However, the NADW outflow is modu- direction, moving northward as the Guiana Current, lated by the oscillating wind forcing with the lag of about most of which moves to the Gulf of Mexico and a part of 2–3 years, rather than by the seesaw balance driven by the which bifurcates to become the Gulf Stream together northward AABW flow varying with the BDT mode (cf. with the North Equatorial Current (NEC) (Fig. 8a). In Figs. 7d and 7e with Figs. 6a and 6c). As will be discussed the North Pacific, most of the SEC does not reach the in the following subsection, the Atlantic MOC varies in- eastern coast of Indonesia, instead turning eastward and dependently of not only the AABW-driven bipolar see- joining the North Equatorial Countercurrent (NECC). saw balance but also the NADW outflow. The Pacific NEC, which is much stronger than the Atlantic NEC, reaches the eastern coast of Philippines b. EDT mode and bifurcates, with one portion turning counterclock- In the North Atlantic, the South Equatorial Current wise and joining the NECC, and the other portion (SEC) reaches the eastern coast of South America with moving northward as the Kuroshio (Fig. 8d). In year 16

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FIG. 7. (a) Deep current field of CTRL in the subtropical Atlantic, the differences between CTRL and SW050P02 at years (b) 16 and (c) 40, and variation in the deep meridional velocity (red solid) and NADW outflow (blue solid) anomalies for (d) SW050P02 and (e) SW050P04. The deep meridional velocity is estimated at 58N, 458W, 2000-m depth; the NADW outflow is calculated using the maximum values between 1000 and 3000-m depths at 308S. The negative (positive) anomaly indicates the intensifying (weakening) current, whereas the positive (negative) NADW outflow anomaly indicates an increase (decrease). The 10-yr running mean is applied to these variables. Time series of the maximum zonal-mean wind stress between 808 and 308S (thin dashed black lines) are presented for reference. when the oscillating wind forcing in SW050P02 passes South America) because the Gulf Stream is significantly through its first positive peak, the northward flowing affected by the gyres that vary with oceanic deep con- surface WBCs such as the Guiana Current of the Atlantic vection and are associated with NADW formation. and the Kuroshio of the Pacific intensify (Figs. 8b,e). Figures 8g and 8h illustrate the variations in surface However, in year 40, when the wind forcing is around its WBCs: one, indicative of the Guiana Current in the first negative peak, these WBCs weaken (Figs. 8c,f). As Atlantic, is estimated by the meridional velocity near the noted in Figs. 8a–f, the scale of variation is much smaller northeastern coast of South America (88–108N, 608– than that of the basic current field in CTRL. Moreover, 558W); the other, indicative of the Kuroshio in the the variation in WBC is smaller in the North Pacific than North Pacific, is estimated by the meridional velocity near it is in the North Atlantic. A comparison of the deep the east coast of Taiwan (238–258N, 1228–1308E). Since (Figs. 7a–c) and bottom (Fig. 5) current variations indi- both currents flow northward, the positive (negative) cates that the EDT mode-induced variation in the surface anomalies are indicative of strengthening (weakening). WBCs is much smaller than the BDT mode-induced The North Atlantic surface WBC anomaly is highly cor- variation in the deep and bottom WBCs. The Gulf related with the oscillating SH westerly winds, and their Stream does not respond linearly to the oscillating SH time lag is dependent on the oscillation period of each westerly winds (in contrast to the Guiana Current, which wind forcing. In SW050P02, the two factors are in good shows a clear correlation near the northeastern coast of correlation without any time lag, whereas in SW050P04

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FIG. 8. The surface current fields of CTRL in the western (a) Atlantic and (d) Pacific, the differences between CTRL and SW050P02 at years (b),(e) 16 and (c),(f) 40, and variation in the North Atlantic (red solid) and Pacific (blue solid) WBCs for (g) SW050P02 and (h) SW050P04. The North Atlantic (North Pacific) WBC is averaged over 88–108N, 608– 558W(238–258N, 1228–1308E). Since the surface WBC is basically northward, the positive (negative) anomaly indicates the intensifying (weakening) surface WBC. The 10-yr running mean is applied to these variables. Time series of the maximum zonal-mean wind stress between 808 and 308S (thin dashed black lines) are presented for reference. the response of the Guiana Current to the oscillating wind GCM, Klinger and Cruz (2009) also showed that the signal forcing is delayed by about 5–7 years. The Kuroshio generated by the 50%-intensified SH westerly winds was anomaly is also highly correlated with the SH westerly quickly propagated to the Northern Hemisphere in 3–10 winds, with a time lag of about 6–8 years in SW050P02 years, which is in good agreement with our results. As and of about 10–12 years in SW050P04. Without any at- such, it is important to elucidate how the change in the SH mospheric teleconnection, this EDT mode-generated westerly winds remotely affects the WBCs in the Northern signal is very quickly propagated to the North Atlantic Hemisphere. According to the study of McDermott and North Pacific. Using a coarse-resolution ocean-only (1996), who used the single- and two-basin simple ocean

Unauthenticated | Downloaded 10/03/21 10:44 PM UTC 2124 JOURNAL OF CLIMATE VOLUME 33 models with a re-entrant channel indicative of the Drake limb and the eastward NECC in their southern limb. In Passage, the 50%-intensified SH westerly winds enhance the Pacific, the NEC reaching the western boundary the anticyclonic subtropical gyre, which can play a role along the Philippine coast near 1258E bifurcates with one in exciting baroclinic Kelvin-like signals that propagate portion turning to the north to become the Kuroshio and northward to the equator along the western boundary. At the other veering south to become the Mindanao Current the equator, these signals turn east and propagate to the (Qiu et al. 2015). The southward flowing Mindanao eastern boundary like equatorial Kelvin modes. While Current constitutes the western limb of the Pacific trop- some of the energy splits into two hemispheres at the ical gyre and separates into the Indonesian Throughflow eastern boundary, some is reflected to the west and and the eastward flowing NECC between 68 and 78N propagates as Rossby modes. At the western boundary, (Lukas et al. 1991). On average, the NECC transports the the reflected Rossby modes act to enhance the north- warm pool water of more than 30 Sv to the eastern Pacific ward WBC. This process, which is termed the ‘‘equa- basin (Zhao et al. 2013), and the mean transport of the torial buffer’’ by Johnson and Marshall (2002), plays a Mindanao Current measured over the thermocline is role in restricting abrupt changes in the thermohaline about 19 Sv. In the Atlantic, although there are negative overturning in the North Atlantic. barotropic streamfunctions (implying cyclonic circulation It should be noted that the WBC variations are gen- between 08 and 208N; Fig. 9c), the Guiana Current actu- erally larger in the Atlantic than in the Pacific, which is ally splits from the North Brazil Current and continues to an important issue requiring our attention. The North flow northward along the western boundary between the Atlantic has major locations where deep water masses NEC and NECC; this contrasts with the southward form and thus plays a crucial role as a ‘‘main engine’’ in flowing Mindanao Current in the tropical Pacific (Condie driving the so-called Atlantic Ocean conveyor belt; this 1991). Moreover, the NECC is robust in the boreal contrasts with the North Pacific, which exhibits no deep summer and fall but is very weak or even absent in the water formation. Since the Atlantic Ocean conveyor winter and spring (Yang and Joyce 2006). This difference belt consists of 1) the northward Ekman transport as- between tropical gyres located in the North Atlantic and sociate with the EDT mode, 2) the NADW formation in Pacific may be a crucial cause of the EDT-generated the Labrador and GIN Seas, 3) its outflow toward the signal having a larger remote impact on the North SO, and 4) upwelling of the CDW, the EDT mode- Atlantic than it does on the North Pacific. As illustrated generated signal may be propagated on this conveyor by the minimum barotropic streamfunction anomalies belt to the Northern Hemisphere faster in the Atlantic calculated in the north equatorial Atlantic and Pacific than it is in the Pacific. However, this study considers (Fig. 9d), the tropical gyre in the Pacific responds in- another cause for the signal moving more rapidly in the versely to the oscillating wind forcing, whereas the re- Atlantic. The SH westerly winds generally give rise to sponse in the tropical Atlantic is negligible. That is, the northward Ekman transport uniformly over the under- distinct, cyclonic gyre in the equatorial Pacific weakens lying zonal band; therefore, the signal is propagated (intensifies) with the increasing (decreasing) northward northward uniformly in all longitudes. Figure 9a illus- Ekman transport under the intensifying (weakening) SH trates the variations of the equatorial sea surface height westerly winds and thus plays a role as a barrier to the (SSH) near the respective western and eastern bound- EDT mode-generated signal. Therefore, in the Atlantic, aries of the North Atlantic and Pacific. The equatorial the Ekman transport can keep moving northward without SSHs oscillate in the four different longitudes with being blocked by the tropical gyre, which is a crucial nearly the same amplitude and phase, which implies that cause of the response of the surface WBC to oscillating the EDT mode-generated signal passes through the SH westerly winds to be larger in the Atlantic than it is in equator uniformly in all longitudes. the Pacific. On the other hand, the lack of topographic As such, it is important to consider why the surface waveguide in the upper ocean of the Indo-Pacific is also a WBCs in the North Atlantic and Pacific respond differ- potential cause of the North Pacific WBC responding ently to the EDT mode even under this zonally uniform belatedly to the SH wind forcing. propagation of the signal. Figure 9b shows the horizon- As such, we should consider how the transported tal barotropic streamfunctions calculated in the North mode affect the Atlantic MOC associated with NADW Atlantic and Pacific. Anticyclonic subtropical gyres con- formation. Figure 10 shows the annual mean, surface trolled by the overlying wind stress curl dominate both horizontal circulation, and mixed layer depth in the oceans north of about 208N, whereas the cyclonic gyre northern North Atlantic. The mixed layer depth is in- south of about 208N is much stronger in the Pacific than it dicative of the strength of oceanic deep convection. The is in the Atlantic. The tropical cyclonic gyres in both relatively very deep mixed layer of the Labrador Sea oceans consist of the westward NEC in their northern and GIN Sea is in good agreement with previous studies

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FIG. 9. (a) Variation in the SSH in the western (solid) and eastern (dashed) boundaries of the North equatorial Atlantic (red) and Pacific (blue), horizontal barotropic streamfunctions of the (b) North Pacific and (c) North Atlantic, and (d) variations in the tropical Atlantic (dashed green) and Pacific (solid green) gyre strength. The Atlantic SSH in the western (eastern) boundary is estimated at 58–108N, 608W(58–108N, 208W), and the Pacific SSH in the western (eastern) boundary is estimated at 58–108N, 1308E(58–108N, 1008W). The Atlantic (Pacific) tropical gyre anomaly is estimated by the minimum values of barotropic streamfunction over 08–208N, 608–208W (1208– 1808E). The positive (negative) gyre anomaly indicates the weakening (intensifying) gyre. The 10-yr running mean is applied to all variables, which are detrended. Time series of the maximum zonal-mean wind stress between 808 and 308S (thin dashed black lines) are presented for reference.

(Killworth 1983; Marshall and Schott 1999). The Gulf near the Labrador Sea. These are also generally in phase Stream, the Norwegian Current, the East/West Greenland with the oscillating wind forcing, except for during the first Currents, and the Labrador Current are well reproduced in cycles of both experiments. Although the upper ocean CTRL, as is the southward eastern boundary current con- temperature anomaly near the Labrador Sea is also in nected to the Canary Current. To identify the factors con- positive correlation with the oscillating wind forcing (not trolling oceanic stratification in the vicinity of the Labrador shown here), the potential density anomaly is determined Sea, a major locations for NADW formation with the GIN by the salt anomaly. These anomalies show overall de- Sea, variations in the upper ocean (0–100 m depths) salt creases for the first 80 years, before increasing. It seems to levels and potential density at the surface (sz50)werees- be associated with the inherent dynamic features of the timated over 608–508W/558–608N(Figs. 11a,c). The Guiana Labrador and GIN Sea cyclonic gyre circulations, which Current (the WBC in the North Atlantic) transports heat require an adjustment period for the EDT mode. Although and salt from subtropics northward and fluctuates in phase not as extensive as these anomalies, the NADW outflow with the oscillating SH westerly winds (Figs. 8g,h), which (Figs. 7d,e) and the Guiana Current (Figs. 8g,h)alsoshow causes upper ocean salt and potential density anomalies some decreases in the early period. The Atlantic MOC

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via the EDT mode-generated signal. Although the signal propagated northward passes through the equator uniformly at all longitudes, perturbation on the WBC is much stronger in the Atlantic than it is in the Pacific, which is due to the blocking effect of the relatively strong and persistent Mindanao Current, the southward flowing WBC of the North Pacific tropical gyre. Northward salt advection by the North Atlantic WBC appears to control the Atlantic MOC associated with the formation rate of NADW, which is independent of the BDT mode-generated signal. Using the 640-day observation data estimated by a FIG. 10. Horizontal distributions of surface current and mixed moored array at the equator in the western Atlantic, Hall layer depth of CTRL averaged over 200 years in the northern North Atlantic. The green dashed-line box indicates the area of et al. (1997) first showed that the rise and fall of the tran- NADW formation in the vicinity of the Labrador Sea (558–608N, sition layer between the AABW and the lower NADW is 608–508W). associated with the bipolar seesaw balance between op- posing changes in the AABW and NADW. The mecha- anomaly estimated by the maximum values of meridional nism governing this seesaw balance was explained in streamfunction over 458–608N/500–1500 m shows a clear relation to the conservation of potential vorticity by Cheon correlation with the upper ocean salt and potential density and Stössel (2009). For the climate community monitoring anomalies near the Labrador Sea except for during the first the Atlantic MOC, the variation in Arctic sea ice is an cycles of both experiments and the third cycle of SW050P04 important factor because Arctic freshwater export directly (Figs. 11b,d). This indicates that the oscillating SH westerly determines the surface buoyancy and thus ocean stratifi- winds remotely affect the Atlantic MOC associated with cationinthemainNADWformationregioninthe the NADW formation via the EDT mode-generated signal, Labrador and GIN Seas (Jungclaus et al. 2005; Jahn and thereby modulating the North Atlantic WBC, northward Holland 2013; Sévellec et al. 2017). In this paper, we con- salt transport from the tropics and thus stratification of the sider the effects of SAM on the Atlantic MOC. Since the Labrador and GIN Seas. Note that the Atlantic MOC SAM index has become more positive since the 1950s varies independently of the abyssal WBCs passing through (Hartmann et al. 2013; Cheon and Gordon 2019)and the equatorial Atlantic (Figs. 6a,c), which implies that the showed a significant positive trend in the representative BDT mode-generated signal does not affect the NADW concentration pathway scenario 8.5 applied to fully cou- formation. pled climate models by Zheng et al. (2013), the negative wind stress curl over the Weddell and Ross Seas may intensify in future climates, which implies an increase in the 4. Summary and discussion potential for activating the BDT mode. Therefore, the Using a sea ice–ocean GCM coupled to an energy BDT mode-generated bipolar seesaw balance was ex- balance atmospheric model, the diverse responses of pected to affect not only the NADW outflow but also the global ocean circulation to sinusoidally oscillating SH Atlantic MOC when the experimental setup was first de- westerly winds are investigated in the context of the signed for this study. However, the NADW outflow ap- BDT mode (starting from the SO polynya events) and peared to be directly controlled by the oscillating SH the EDT mode (starting from the northward Ekman westerly winds and the ensuing Ekman divergence-driven transport). The Antarctic MOC that is indicative of the upwelling of the CDW. Moreover, the Atlantic MOC ap- formation rate of AABW varies not only linearly with peared to be affected by the remote effect of the EDT oceanic convection associated with coastal and open- mode-generated signal. That is, a positive trend in the ocean polynyas but also in an inverse correlation with SAM, which is expected in a warming climate, is not able upwelling of the CDW. Owing to the NADW occupying to weaken the Atlantic MOC by the BDT mode. the deep North Atlantic, the AABW passing through The formation of the NADW is hypothesized to play a the equator and flowing northward is highly predomi- key role in driving the Atlantic Ocean conveyor belt. In nant in the Pacific. In the Atlantic, the seesaw balance this study, the North Atlantic surface WBC and the between deep and bottom waters at the equator appears NADW outflow corresponding to the upper and lower to be controlled by BDT mode-induced AABW for- parts of the Atlantic Ocean conveyor belt appear to be mation. The SH westerly winds play a further role in highly correlated with the oscillating SH westerly winds, remotely affecting the North Atlantic and Pacific WBCs which directly control the Ekman divergence-driven

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FIG. 11. Variation (a),(c) in the upper ocean (0–100 m depths) salt (red) and potential density at the surface (sz50, blue) near the Labrador Sea (558–608N, 608–508W), and (b),(d) in the Atlantic MOC (green), for (a),(b) SW050P02 and (c),(d) SW050P04. The Atlantic MOC is estimated by the maximum values of meridional overturning streamfunction over 458–608N, 500–1500-m depth. The 10-yr running mean is applied to all variables, which are detrended. Time series of the maximum zonal-mean wind stress between 808 and 308S (thin dashed black lines) are presented for reference. upwelling of the CDW (i.e., the southern portion of the AABW variation). Since the north–south migration of SH conveyor belt). With lags of about a decade the Atlantic westerly winds is another important characteristic of the MOC variation is correlated with the oscillating SH SAM that is not considered in this study, the oscillating westerly winds, which modulate the transport of salinity position and its combination with oscillating strength are from the subtropics to the northern North Atlantic. In worthy of investigation in terms of the responses of sea this regard, the SH westerly wind anomalies associated ice/ocean interactions such as polynyas, oceanic deep with the SAM can drive the Atlantic Ocean conveyor convection, and global ocean circulation. In comparison belt variations. with a coarse-resolution ocean GCM, a high-resolution Using an ocean-only GCM, Oke and England (2004) ocean GCM that can explicitly resolve mesoscale eddies showed the impact of poleward shifting SH westerly winds does not show large ACC responses to the intensifying on the ocean circulation in association with the positive SH westerly winds because of the role that eddies play trend in SAM (i.e., without sea ice–ocean interaction and in diffusing heat and salt meridionally (Hallberg and

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Gnanadesikan 2006; Gent 2016); this is the reason that the Cunningham, S. A., S. G. Alderson, B. A. King, and M. A. ACC response was not analyzed in this study. However, Brandon, 2003: Transport and variability of the Antarctic the sea ice–ocean coupled GCM can reproduce AABW Circumpolar Current in Drake Passage. J. Geophys. Res., 108, 8084, https://doi.org/10.1029/2001JC001147. variations due to sea ice–ocean interactions and thus af- Dong, S., S. L. Garzoli, M. O. Baringer, C. S. Meinen, and G. J. fect the ACC via the thermal wind balance. Therefore, Goni, 2009: Interannual variations in the Atlantic meridional developing an eddy-resolving high-resolution sea ice– overturning circulation and its relationship with the net ocean coupled GCM forced by oscillating SH westerly northward heat transport in the South Atlantic. Geophys. Res. winds is a challenge in terms of investigating the rela- Lett., 36, L20606, https://doi.org/10.1029/2009GL039356. England, M. H., 1993: Representing the global-scale water masses tionship between the ACC and SAM. in ocean general circulation models. J. Phys. Oceanogr., 23, 1523–1552, https://doi.org/10.1175/1520-0485(1993)023,1523: Acknowledgments. We thank the editor and four RTGSWM.2.0.CO;2. anonymous reviewers for their critical constructive com- Farneti, R., and P. R. Gent, 2011: The effects of the eddy-induced ments and suggestions that helped improve this study advection coefficient in a coarse-resolution coupled climate model. Ocean Modell., 39, 135–145, https://doi.org/10.1016/ greatly. The contribution of Jong-Seong Kug was spon- j.ocemod.2011.02.005. sored by a research grant from the Korean Ministry of ——, and Coauthors, 2015: An assessment of Antarctic Circumpolar Oceans and Fisheries (KIMST20190361; PM19020). Current and Southern Ocean meridional overturning circulation during 1958–2007 in a suite of interannual CORE-II simulations. Ocean Modell., 93,84–120,https://doi.org/10.1016/ REFERENCES j.ocemod.2015.07.009. Anderson, J. L., and Coauthors, 2004: The new GFDL global at- Ganachaud, A., and C. Wunsch, 2000: Improved estimates of the mosphere and land model AM2-LM2: Evaluation with pre- global ocean circulation, heat transport and mixing from hy- scribed SST simulations. J. Climate, 17, 4641–4673, https:// drographic data. Nature, 408, 453–457, https://doi.org/10.1038/ doi.org/10.1175/JCLI-3223.1. 35044048. Bracegirdle, T. J., and Coauthors, 2013: Assessment of surface Gent, P. R., 2016: Effects of Southern Hemisphere wind changes on winds over the Atlantic, Indian, and Pacific Ocean sectors of the meridional overturning circulation in the ocean models. the Southern Ocean in CMIP5 models: Historical bias, forcing Annu. Rev. Mar. Sci., 8, 79–94, https://doi.org/10.1146/annurev- response, and state dependence. J. Geophys. Res. Atmos., 118, marine-122414-033929. 547–562, https://doi.org/10.1002/JGRD.50153. ——, and J. C. McWilliams, 1990: Isopycnal mixing in ocean circu- Brix, H., and R. Gerdes, 2003: North Atlantic Deep Water and lation models. J. Phys. Oceanogr., 20, 150–155, https://doi.org/ , . Antarctic Bottom Water: Their interaction and influence on 10.1175/1520-0485(1990)020 0150:IMIOCM 2.0.CO;2. the variability of the global ocean circulation. J. Geophys. ——, and G. Danabasoglu, 2011: Response to increasing Southern Res., 108, 3022, https://doi.org/10.1029/2002JC001335. Hemisphere winds in CCSM4. J. Climate, 24, 4992–4998, Broecker, W. S., 1998: Paleocean circulation during the last de- https://doi.org/10.1175/JCLI-D-10-05011.1. glaciation: A bipolar seesaw? Paleoceanography, 13, 119–121, Gerdes, R., W. Hurlin, and S. M. Griffies, 2006: Sensitivity of a global https://doi.org/10.1029/97PA03707. ocean model to increased run-off from Greenland. Ocean Modell., Cheng, W., J. C. H. Chiang, and D. Zhang, 2013: Atlantic meridi- 12, 416–435, https://doi.org/10.1016/j.ocemod.2005.08.003. onal overturning circulation (AMOC) in CMIP5 models: RCP Gille, S. T., 2005: Statistical characterization of zonal and meridi- and historical simulations. J. Climate, 26, 7187–7197, https:// onal ocean wind stress. J. Atmos. Oceanic Technol., 22, 1353– doi.org/10.1175/JCLI-D-12-00496.1. 1372, https://doi.org/10.1175/JTECH1789.1. Cheon, W. G., and A. Stössel, 2009: Oceanic response to interactive Gnanadesikan, A., 1999: A simple predictive model for the struc- momentum flux over Southern Hemisphere sea ice. J. Geophys. ture of the oceanic pycnocline. Science, 283, 2077–2079, Res., 114,C12002,https://doi.org/10.1029/2008JC005068. https://doi.org/10.1126/science.283.5410.2077. ——, and A. L. Gordon, 2019: Open-ocean polynyas and deep ——, and R. W. Hallberg, 2000: On the relationship of the Circumpolar convection in the Southern Ocean. Sci. Rep., 9, 6935, https:// Current to Southern Hemisphere winds in coarse-resolution doi.org/10.1038/s41598-019-43466-2. ocean models. J. Phys. Oceanogr., 30, 2013–2034, https://doi.org/ ——, Y.-G. Park, J. R. Toggweiler, and S.-K. Lee, 2014: The rela- 10.1175/1520-0485(2000)030,2013:OTROTC.2.0.CO;2. tionship of Weddell Polynya and open-ocean deep convection ——, A. M. de Boer, and B. K. Migone, 2007: A simple theory of to the Southern Hemisphere westerly winds. J. Phys. Oceanogr., the pycnocline and overturning revisited. Ocean Circulation: 44, 694–713, https://doi.org/10.1175/JPO-D-13-0112.1. Mechanisms and Impacts—Past and Future Changes of ——, S.-K. Lee, A. L. Gordon, Y. Liu, C.-B. Cho, and J. J. Park, 2015: Meridional Overturning, Geophys. Monogr., Vol. 173, Amer. Replicating the 1970s’ Weddell Polynya using a coupled ocean– Geophys. Union, 19–32. sea ice model with reanalysis surface flux fields. Geophys. Res. Gong, D., and S. Wang, 1999: Definition of Antarctic Oscillation Lett., 42, 5411–5418, https://doi.org/10.1002/2015GL064364. index. Geophys. Res. Lett., 26, 459–462, https://doi.org/10.1029/ ——, C.-B. Cho, A. L. Gordon, Y. H. Kim, and Y.-G. Park, 2018: 1999GL900003. The role of oscillating Southern Hemisphere westerly winds: Gordon, A. L., M. Visbeck, and J. C. Comiso, 2007: A possible Southern Ocean coastal and open-ocean polynyas. J. Climate, link between the Weddell polynya and the southern annu- 31, 1053–1073, https://doi.org/10.1175/JCLI-D-17-0237.1. lar mode. J. Climate, 20, 2558–2571, https://doi.org/10.1175/ Condie, S., 1991: Separation and recirculation of the North Brazil JCLI4046.1. Current. J. Mar. Res., 49, 1–19, https://doi.org/10.1357/ Griffies, S. M., M. J. Harrison, R. C. Pacanowski, and A. Rosati, 002224091784968620. 2004: A technical guide to MOM4. GFDL Ocean Group Tech.

Unauthenticated | Downloaded 10/03/21 10:44 PM UTC 15 MARCH 2020 C H E O N A N D K U G 2129

Rep. 5, NOAA/GFDL, 342 pp., https://www.gfdl.noaa.gov/ Mantyla, A. W., and J. L. Reid, 1983: Abyssal characteristics of the bibliography/related_files/smg0301.pdf. World Ocean waters. Deep-Sea Res., 30A, 805–833, https:// Hall, M., M. McCartney, and J. Whitehead, 1997: Antarctic bot- doi.org/10.1016/0198-0149(83)90002-X. tom water flux in the equatorial western Atlantic. J. Phys. Marshall, J., and F. Schott, 1999: Open-ocean convection: Oceanogr., 27, 1903–1926, https://doi.org/10.1175/1520- Observation, theory, and models. Rev. Geophys., 37, 1–64, 0485(1997)027,1903:ABWFIT.2.0.CO;2. https://doi.org/10.1029/98RG02739. Hallberg, R., and A. Gnanadesikan, 2006: The role of eddies in McDermott, D. A., 1996: The regulation of northern overturning determining the structure and response of the wind-driven by Southern Hemisphere winds. J. Phys. Oceanogr., 26, Southern Hemisphere overturning: Results from the Modeling 1234–1255, https://doi.org/10.1175/1520-0485(1996)026,1234: Eddies in the Southern Ocean (MESO) Project. J. Phys. TRONOB.2.0.CO;2. Oceanogr., 36, 2232–2252, https://doi.org/10.1175/JPO2980.1. Morozov, E. G., A. N. Demidov, R. Y. Tarakanov, and W. Zenk, Hartmann, D. L., and Coauthors, 2013: Observations: Atmosphere 2010: Deep water masses of the South and North Atlantic. and surface. Climate Change 2013: The Physical Science Basis, Abyssal Channels in the Atlantic Ocean: Water Structure and T. F. Stocker et al., Eds., Cambridge University Press, 159– Flows, Springer, 25–50. 254. Oke, P., and M. England, 2004: Oceanic response to changes in the Hibler, W. D., III, 1979: A dynamic thermodynamic sea ice model. latitude of the Southern Hemisphere subpolar westerly winds. J. Phys. Oceanogr., 9, 815–846, https://doi.org/10.1175/1520- J. Climate, 17, 1040–1054, https://doi.org/10.1175/1520-0442(2004) 0485(1979)009,0815:ADTSIM.2.0.CO;2. 017,1040:ORTCIT.2.0.CO;2. Hofmann, M., and M. A. Morales-Maqueda, 2011: The response of Orsi, A. H., T. Whitworth III, and W. D. Nowlin, 1995: On the Southern Ocean eddies to increased midlatitude westerlies: A meridional extent and fronts of the Antarctic Circumpolar non-eddy resolving model study. Geophys. Res. Lett., 38, Current. Deep-Sea Res. I, 42, 641–673, https://doi.org/10.1016/ L03605, https://doi.org/10.1029/2010GL045972. 0967-0637(95)00021-W. Jahn, A., and M. M. Holland, 2013: Implication of Arctic sea ice ——, S. S. Jacobs, A. L. Gordon, and M. Visbeck, 2001: Cooling changes for North Atlantic Deep convection and the meridional and ventilating the abyssal ocean. Geophys. Res. Lett., 28, overturning circulation in CCSM4-CMIP5 simulations. Geophys. 2923–2926, https://doi.org/10.1029/2001GL012830. Res. Lett., 40, 1206–1211, https://doi.org/10.1002/grl.50183. Qiu, B., D. L. Rudnick, I. Cerovecki, B. D. Cornuelle, S. Chen, M. C. Johnson, G. C., 2008: Quantifying Antarctic Bottom Water and Schönau, J. L. McClean, and G. Gopalakrishnan, 2015: The Pacific North Atlantic Deep Water volumes. J. Geophys. Res., 113, North Equatorial current: New insights from the origins of C05027, https://doi.org/10.1029/2007JC004477. the Kuroshio and Mindanao Currents (OKMC) Project. Prog. Johnson, H. L., and D. P. Marshall, 2002: A theory for the surface Oceanogr., 28, 24–33, https://doi.org/10.5670/oceanog.2015.78. Atlantic response to thermohaline variability. J. Phys. Oceanogr., Rahmstorf, S., 1993: A fast and complete convection scheme for 32, 1121–1132, https://doi.org/10.1175/1520-0485(2002)032,1121: ocean models. Ocean Modell., 101, 9–11. ATFTSA.2.0.CO;2. ——, and M. H. England, 1997: Influence of Southern Hemisphere Jungclaus, J. H., H. Haak, M. Latif, and U. Mikolajewicz, 2005: winds on North Atlantic deep water flow. J. Phys. Oceanogr., 27, Arctic–North Atlantic interactions and multidecadal vari- 2040–2054, https://doi.org/10.1175/1520-0485(1997)027,2040: ability of the meridional overturning circulation. J. Climate, IOSHWO.2.0.CO;2. 18, 4013–4031, https://doi.org/10.1175/JCLI3462.1. Ribbe, J., and M. Tomczak, 1997: On convection and the formation Kawabe, M., S. Fujio, and D. Yanagimoto, 2003: Deep-water cir- of Subantarctic Mode Water in the fine resolution Antarctic culation at low latitudes in the western North Pacific. Deep- model. J. Mar. Syst., 13, 137–154, https://doi.org/10.1016/ Sea Res. I, 50, 631–656, https://doi.org/10.1016/S0967-0637(03) S0924-7963(96)00119-4. 00040-2. Roemmich, D., S. Hautala, and D. Rudnick, 1996: Northward Killworth, P. D., 1983: Deep convection in the World Ocean. Rev. abyssal transport through the Samoan Passage and adjacent Geophys., 21, 1–26, https://doi.org/10.1029/RG021i001p00001. regions. J. Geophys. Res., 101, 14 039–14 055, https://doi.org/ Klinger, B. A., and C. Cruz, 2009: Decadal response of global cir- 10.1029/96JC00797. culation to Southern Ocean zonal wind stress perturbation. Santoso, A., and M. H. England, 2004: Antarctic Intermediate J. Phys. Oceanogr., 39, 1888–1904, https://doi.org/10.1175/ Water circulation and variability in a coupled climate model. 2009JPO4070.1. J. Phys. Oceanogr., 34, 2160–2179, https://doi.org/10.1175/ Kuhlbrodt, T., R. S. Smith, Z. Wang, and J. M. Gregory, 2012: The 1520-0485(2004)034,2160:AIWCAV.2.0.CO;2. influence of eddy parameterizations on the transport of the Schmitz, W. J., Jr., and M. S. McCartney, 1993: On the North Antarctic Circumpolar Current in coupled climate models. Ocean Atlantic circulation. Rev. Geophys., 31, 29–49, https://doi.org/ Modell., 52–53,1–8,https://doi.org/10.1016/j.ocemod.2012.04.006. 10.1029/92RG02583. Locarnini, R. A., and Coauthors, 2013: Temperature. Vol. 1, World Schott, F. A., R. Zantopp, L. Stramma, M. Dengler, J. Fischer, and Ocean Atlas 2013, NOAA Atlas NESDIS 73, 40 pp., http:// M. Wibaux, 2004: Circulation and deep-water export at the data.nodc.noaa.gov/woa/WOA13/DOC/woa13_vol1.pdf. western exit of the subpolar North Atlantic. J. Phys. Oceanogr., Lukas, R., E. Firing, P. Hacker, P. L. Richardson, C. A. Collins, 34, 817–843, https://doi.org/10.1175/1520-0485(2004)034,0817: R. Fine, and R. Gammon, 1991: Observation of the Mindanao CADEAT.2.0.CO;2. Current during the Western Equatorial Pacific Ocean Circulation Sévellec, F., A. V. Fedorov, and W. Liu, 2017: Arctic sea-ice Study. J. Geophys. Res., 96, 7089–7104, https://doi.org/10.1029/ decline weakens the Atlantic meridional overturning circu- 91JC00062. lation. Nat. Climate Change, 7, 604–610, https://doi.org/ Lumpkin, R., and K. Speer, 2003: Large-scale vertical and hori- 10.1038/nclimate3353. zontal circulation in the North Atlantic Ocean. J. Phys. Smith, W. H. F., and D. T. Sandwell, 1997: Global seafloor topog- Oceanogr., 33, 1902–1920, https://doi.org/10.1175/1520-0485(2003) raphy from satellite altimetry and ship depth soundings. Science, 033,1902:LVAHCI.2.0.CO;2. 277, 1957–1962, https://doi.org/10.1126/science.277.5334.1956.

Unauthenticated | Downloaded 10/03/21 10:44 PM UTC 2130 JOURNAL OF CLIMATE VOLUME 33

Sørensen, J. V. T., J. Ribbe, and G. Shaffer, 2001: Antarctic water in a Southern Ocean state estimate. Geophys. Res. Lett., 40, Intermediate Water mass formation in ocean general circulation 2177–2182, https://doi.org/10.1002/grl.50483. models. J. Phys. Oceanogr., 31, 3295–3311, https://doi.org/10.1175/ Whitworth,T.,III,W.D.Nowlin,andS.J.Worley,1982:Thenet 1520-0485(2001)031,3295:AIWMFI.2.0.CO;2. transport of the Antarctic Circumpolar Current through Drake Speer, K., S. R. Rintoul, and B. Sloyan, 2000: The diabatic Deacon Passage. J. Phys. Oceanogr., 12, 960–971, https://doi.org/10.1175/ cell. J. Phys. Oceanogr., 30, 3212–3222, https://doi.org/10.1175/ 1520-0485(1982)012,0960:TNTOTA.2.0.CO;2. 1520-0485(2000)030,3212:TDDC.2.0.CO;2. Winton, M., 2000: A reformulated three-layer sea ice model. Stephens, J. C., and D. P. Marshall, 2000: Dynamical pathways of J. Atmos. Oceanic Technol., 17, 525–531, https://doi.org/10.1175/ Antarctic Bottom Water in the Atlantic. J. Phys. Oceanogr., 30, 1520-0426(2000)017,0525:ARTLSI.2.0.CO;2. 622–640, https://doi.org/10.1175/1520-0485(2000)030,0622: Yang, J., and T. M. Joyce, 2006: Local and equatorial forcing of DPOABW.2.0.CO;2. seasonal variations of the North Equatorial Countercurrent in Swart, N. C., and J. C. Fyfe, 2012: Observed and simulated changes in the the Atlantic Ocean. J. Phys. Oceanogr., 36, 238–254, https:// Southern Hemisphere surface westerly wind-stress. Geophys. Res. doi.org/10.1175/JPO2848.1. Lett., 39, L16711, https://doi.org/10.1029/2012GL052810. Zanowski, H., and R. Hallberg, 2017: Weddell Polynya transport Thompson, D. W. J., and J. M. Wallace, 2000: Annular modes in mechanisms in the abyssal ocean. J. Phys. Oceanogr., 47, 2907– the extratropical circulation. Part I: Month-to-month vari- 2925, https://doi.org/10.1175/JPO-D-17-0091.1. ability. J. Climate, 13, 1000–1016, https://doi.org/10.1175/ Zhao, J., Y. Li, and F. Wang, 2013: Dynamical responses of the west 1520-0442(2000)013,1000:AMITEC.2.0.CO;2. Paciﬁc North Equatorial Countercurrent (NECC) system to Toggweiler, J. R., and B. Samuels, 1993: Is the magnitude of the El Niño events. J. Geophys. Res. Oceans, 118, 2828–2844, deep outﬂow from the Atlantic Ocean actually governed https://doi.org/10.1002/jgrc.20196. by Southern Hemisphere winds? The Global Carbon Cycle, Zheng, F., J. Li, R. T. Clark, and H. C. Nnamchi, 2013: Simulation M. Heimann, Ed., Springer, 303–331. and projection of the Southern Hemisphere annular mode in ——, and ——, 1995: Effects of Drake Passage on the global CMIP5 models. J. Climate, 26, 9860–9879, https://doi.org/ thermohaline circulation. Deep-Sea Res. I, 42, 477–500, https:// 10.1175/JCLI-D-13-00204.1. doi.org/10.1016/0967-0637(95)00012-U. Zweng, M., and Coauthors, 2013: Salinity. Vol. 2, World Ocean van Sebille, E., P. Spence, M. R. Mazloff, M. H. England, S. R. Rintoul, Atlas 2013, NOAA Atlas NESDIS 74, 39 pp., http://data. and O. A. Saenko, 2013: Abyssal connections of Antarctic bottom nodc.noaa.gov/woa/WOA13/DOC/woa13_vol2.pdf.

Unauthenticated | Downloaded 10/03/21 10:44 PM UTC