Upwelling in the Ocean Basins North of the ACC: 2. How Cool

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RESEARCH ARTICLE Upwelling in the Ocean Basins North of the ACC: 2. How 10.1029/2018JC014795 Cool Subantarctic Water Reaches This article is a companion to Toggweiler et al. (2019), https://doi.org/ the Surface in the Tropics 10.1029/2018JC014794. J. R. Toggweiler1 , Ellen R. M. Druffel2 , Robert M. Key3 , and Eric D. Galbraith4,5

1 2 Key Points: Geophysical Fluid Dynamics Laboratory, NOAA, Princeton, NJ, USA, Department of Earth System Science, University 3 • Cool subantarctic water is drawn up of California, Irvine, CA, USA, Atmospheric and Oceanic Sciences Program, Princeton University, Princeton, NJ, USA, to the surface in the tropical oceans 4ICREA, Barcelona, Spain, 5ICTA, Universitat Autonoma de Barcelona, Barcelona, Spain • The upwelling tends to be poorly expressed in ocean general circulation models Abstract Large volumes of cool water are drawn up to the surface in the tropical oceans. A companion • A new mechanism is proposed to explain how the upwelling is tied paper shows that the cool water reaches the surface in or near the upwelling zones off northern and southern into the ocean's large‐scale Africa and Peru. The cool water has a subantarctic origin and spreads extensively across the Atlantic and overturning Pacific basins after it reaches the surface. Here, we look at the spreading in two low‐resolution ocean general circulation models and find that the spreading in the models is much less extensive than observed. The problem seems to be the way the upwelling and the spreading are connected (or not connected) to the Correspondence to: ocean's large‐scale overturning. As proposed here, the cool upwelling develops when warm buoyant water in J. R. Toggweiler, the western tropics is drawn away to become deep water in the North Atlantic. The drawing away [email protected] “ ” shoals the tropical thermocline in a way that allows cool subantarctic water to be drawn up to the surface along the eastern margins. The amounts of upwelling produced this way exceed the amounts generated by Citation: the winds in the upwelling zones by as much as 4 times. Flow restrictions make it difficult for the warm Toggweiler, J. R., Druffel, E. R. M., Key, R. M., & Galbraith, E. D. (2019). buoyant water in our models to be drawn away. Upwelling in the ocean basins north of the ACC: 2. How cool subantarctic Plain Language Summary A companion paper uses the radioactive isotope carbon‐14 to map water reaches the surface in the tropics. the upwelling in the ocean basins north of the Antarctic Circumpolar Current (ACC). It shows that deep Journal of Geophysical Research: water drawn up to the surface to the south of the ACC is drawn up to the surface a second time in a Oceans, 124. https://doi.org/10.1029/ 2018JC014795 number of upwelling areas in the tropics. In this paper, we attempt to simulate the tropical upwelling in an ocean circulation model and find that it is largely missing. Other circulation features are missing as Received 21 NOV 2018 well. The other features operate on the opposite sides of the ocean basins from the upwelling areas and Accepted 28 FEB 2019 help carry warm water out of the tropics ( outflows ). The combination of missing elements leads to a Accepted article online 4 MAR 2019 “ ” hypothesis about the upwelling mechanism. Our hypothesis is that the tropical upwelling is driven by the outflows and the fact that the warm water being carried away ultimately becomes deep water again in the North Atlantic.

1. Introduction The overturning of North Atlantic Deep Water (NADW) has been associated in recent years with the wind‐driven upwelling in the Southern Ocean (Marshall & Speer, 2012). The key to this association is the open channel through Drake Passage and the westerly winds directly above and north of the channel. In this setting, the westerlies induce upwelling within the channel and a northward transport out of the channel in the surface Ekman layer. The flow into the channel, meanwhile, is constrained to be a deep geostrophic flow. The net result is a massive conversion of deep water into upper‐ocean water in the south that is balanced by the formation and sinking of NADW in the North Atlantic (Gnanadesikan, 1999; Toggweiler & Samuels, 1995). While this description accounts for the overall mass balance, it does not account for the temperature and buoyancy changes that are known to take place. The deep water that is drawn up to the surface in the south has a temperature of about 2 °C. It is warmed and freshened as it is pushed to the north and is subsequently converted into 7–10° subantarctic water in the area north of the Antarctic Circumpolar Current (ACC). The 7–10° water then flows back to the North Atlantic (Gordon et al., 1992; Rintoul, 1991; Talley, 2008). If the ©2019. American Geophysical Union. 7–10° water were to return to the North Atlantic directly, without alteration, the temperature changes All Rights Reserved. associated with the overturning would be limited to 8°.

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Schmitz and Richardson (1991) showed years ago that the return flow of NADW in the Atlantic takes place via the coolest and the warmest parts of the Florida Current. The cool part has temperatures of 7–10 °C, like the inflow from the ACC. The warm part, on the other hand, has temperatures in excess of 24°. This shows (1) that the return flow is strongly differentiated in the vertical and (2) that part of the cool inflow in the south undergoes a massive transformation as it passes through the tropics. Wunsch (1984) examined the return flow through the tropical Atlantic with an inverse model and determined that 7–10 Sv of the inflow from the south is upwelled to the surface in the equatorial zone. He did not say how the upwelling occurs but suggested that it occurs off the coast of southern Africa. More recently, Lumpkin and Speer (2007) carried out a global inversion of the meridional flow and determined that 60% of the flow through the Atlantic, about 10 out of 16 Sv overall, is transformed into a near‐surface flow, in accord with the results above. Sloyan et al. (2003), in a Pacific inversion study, determined that ~9 Sv of subantarctic water reaches the surface in the eastern equatorial Pacific, water that is presumed to flow back to the Atlantic through the Indonesian Seas and the Indian Ocean. So, while the upwelling in the Southern Ocean is important, much of the deep water upwelled in the Southern Ocean is drawn up to the surface again in the tropics. The additional upwelling has major implications for the heat uptake and heat transport associated with the overturning of NADW. Unfortunately, the inversion studies do not shed much light on the location or the mechanism behind the upwelling. 14 A companion paper (Toggweiler et al., 2019) uses the surface distribution of Δ C to show where the upwel- 14 ling north of the ACC occurs. The present paper compares the Δ C distributions in the companion paper with the simulated distributions in two low‐resolution ocean general circulation models. The model simulations are wide of the mark, but they lead in a roundabout way to a possible explanation of how the upwelling in the real ocean occurs. The present paper begins with a review of the findings and protocols of the companion paper (section 2). The models are described in section 3, and their circulation issues are previewed in section 4. Our data‐model comparison is presented in sections 5–7. Section 8.1 outlines our proposed mechanism. The main flaw in our models and our full hypothesis are described next (sections 8.2–8.4). The paper ends with a discussion about the tendencies in more highly resolved ocean models (section 8.6).

2. Brief Synopsis of the Findings and Protocols of the Companion Paper The deep waters of the ocean are deficient in 14C because they are isolated from the atmosphere. When deep water is drawn back up to the surface, the deficits in the deep water can persist at the surface for 10 years or more due to the slow gas exchange rate for 14C (Broecker & Peng, 1974). Thus, low levels of 14C at the surface 14 are a fairly indelible sign of upwelling. Carbon‐14 values are reported as Δ C, the per mil departure of the measured 14C/12C ratio from a standard reference ratio after a correction for isotopic fractionation (Stuiver & Polach, 1977). Subantarctic Mode Water (SAMW) is a water mass that is found just below the thermocline in all three ocean basins (McCartney, 1977). It is a mixture of upwelled deep water from the area south of the ACC and subtropical water from the subtropical gyres (Iudicone et al., 2008). Toggweiler et al. (1991), hereon TDB91, 14 determined that the Δ C of the SAMW in the South Pacific was about −76‰ to −81‰ during prebomb time. The companion paper shows that water with this composition was reaching the surface off Peru, Northwest Africa, and in the northern Arabian Sea before 1955. 14 Figure 1, from the companion paper, is a map of the surface Δ C deficits in the Atlantic Ocean during 1990– 14 14 1994. A “deficit” is defined as the measured local Δ C minus the contemporary Δ C at Okinawa in the western North Pacific. The map shows that the largest deficits south of 40°N are located near the upwelling 14 zones off Northwest Africa and Namibia along the African margin. Low‐Δ C water seems to spread across 14 the entire tropical Atlantic from these areas. Low‐Δ C water seems to spread across the equatorial Pacific from Peru in much the same way. The extent of the spreading is surprising because the volumes of upwelling in these areas are thought to be fairly modest. Messié et al. (2009), for example, determined that the along- shore winds in the upwelling zones draw 2 Sv or less up to the surface.

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14 The surface Δ C has been altered by the burning of fossil fuels and the nuclear weapons tests of the 1950s and early 1960s. The burning of fossil 14 fuels brought about a slow decline in the atmospheric Δ C that began around 1900 and continues to the present day. The bomb tests, mean- 14 14 while, added a large spike of new C atoms that elevated the Δ C. The 14 companion paper uses the Δ C at Okinawa to help neutralize the impact of the bomb and fossil fuel transients. 14 Because the measurement errors for Δ C are not small, the companion 14 paper also identiﬁes 143 “regions of interest” where multiple Δ C measurements are available during one of three time windows. All the measurements from a given region and time window are averaged. The 14 average Δ C from Okinawa is then subtracted from the regional averages to produce the deﬁcits in Figure 1 and those in the companion paper.

14 3. Surface Δ C in CM2Mc CM2Mc is a low‐resolution version of GFDL's coupled climate model CM2.1 (Delworth et al., 2006). It was set up by Galbraith et al. (2011) to study the ocean's biogeochemical cycling. The ocean component is based 14 on a 2.5° grid with 28 levels in the vertical. The overturning of NADW in Figure 1. Surface Δ C deficits in the Atlantic during 1990–1994 relative to Okinawa. The original figure appears in the companion paper (Toggweiler CM2Mc is nominally quite strong, 22 Sv. CM2Mc has separate tracers for 14 12 13 14 14 et al., 2019). Data sources are given in the supporting information of the CO2 and CO2. C is not simulated, so Δ C is simply ( CO2/ companion paper. The deficits are color coded according to their magnitude, 12 CO2 − 1) · 1,000. small (red), medium (orange), large (green), and extra large (blue). The 14 Okinawa reference value for 1990–1994 is 116‰. WOCE = World Ocean Galbraith et al. (2011) described the Δ C distribution in CM2Mc at the Circulation Experiment. end of a long pre‐industrial run. The surface distribution is shown in the 14 top panel of Figure 2. As will be shown below, the Δ C deficits along the eastern margins of the ocean basins are much smaller than observed—a clear indication that the upwelling in the real ocean is not well represented. The weak upwelling could be due to winds in the coupled model that are too weak. So, a second run was carried out in an uncoupled model, OM1p7, with wind stresses derived from observations. The surface 14 Δ C from this run is shown in the bottom panel of Figure 2. OM1p7 was developed at the same time as CM2Mc and is based on the same grid (Duteil et al., 2013; Galbraith et al., 2010). The forcing for OM1p7 is from the normal year Coordinated Ocean‐ice Reference Experiment (CORE)‐1 forcing described in Large and Yeager (2004) and Griffies et al. (2009). 14 Both preindustrial runs were extended to introduce the fossil fuel and bomb perturbations. The Δ C values from the area near Okinawa were then subtracted to create deficit maps like those in the companion paper. Detailed data‐model comparisons are given below for three time periods, the prebomb period (1940–1954), the World Ocean Circulation Experiment (WOCE) era (1990–1994), and the Climate and Ocean: Variability, Predictability and Change (CLIVAR) era (2003–2006). The measured deficits are overlaid on the model output below as colored dots. When the model and the observations are in agreement the colors in the dots are the same as the colors on the map. The same color scale is used for all three time periods. The data files used to put the dots on the maps are included in the supporting information for the companion paper.

4. Overview of the Model Simulations Okinawa is located at 26.5°N on the offshore side of the Kuroshio in the western North Paciﬁc. One 14 would expect the surface Δ C in this area to be well simulated in a model. Indeed, the simulated prebomb surface values for the Okinawa area are −43‰ in the coupled model and −40‰ in the CORE‐ forced model. These values agree very well with the measured value of −41‰ (Table S4 in the companion

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paper). The problems with our models show up mainly along the eastern margins and are associated with a spurious shallow overturning circulation. Figure 3 shows the meridional overturning circulations from the two models in the IndoPacific sector. The overturning in the top panel is from the coupled model. It features a robust shallow overturning in the North Pacific (centered at 500 m and 35°N) in which ~5 Sv of water from the equatorial zone reaches the subpolar zone north of 50°N. The water appears to sink to about 1000 m and is drawn up to the surface again near the equator. The same shallow overturning is present in the CORE‐forced model in the bottom panel but is not as strong. 14 The shallow cells in Figure 3 are readily apparent in the surface Δ C distributions simulated by our models. Their impact is most easily seen off North America, where the orange shades in Figure 2 track the flow 14 of salty high‐Δ C subtropical water toward California and then northward into the Gulf of Alaska. As can be seen in the Gulf of Alaska, the northward flow is stronger in the coupled model with the stronger shallow overturning. Both models also have a complimentary flow in the South Pacific, as seen in the darker orange shades in Figure 2 that extend across 40°S along the coast of Chile. The shallow overturning is thereby present south of the equator as well but is not as easily differentiated from the overturning below. As will be shown in more detail below, the poleward flows into the Gulf of Alaska and off the coast of Chile simply do not exist in the real ocean. 14 Figure 2. Surface Δ C from the fully coupled version (top) and the The shallow cells in Figure 3 are related to the shallow overturning Coordinated Ocean‐ice Reference Experiment (CORE)‐forced version (bot- circulations (“subtropical cells”) described by McCreary and Lu (1994) tom) of CM2Mc at the end of preindustrial time. and Liu et al. (1994). The tropical easterlies, in these descriptions, induce

Figure 3. Meridional overturning stream functions for the IndoPaciﬁc region from the coupled version of CM2Mc (top) and the Coordinated Ocean‐ice Reference Experiment‐forced version (bottom).

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Figure 4. Deﬁcit patterns from the coupled version of CM2Mc (top) and Coordinated Ocean‐ice Reference Experiment‐ forced version of CM2Mc (bottom) during prebomb time (1940–1954). Observations from the companion paper are overlaid as colored dots.

upwelling along the equator, which is balanced by downwelling and subduction in the areas to the north and south. The downwelling and subduction in the real ocean take place in the subtropical gyres. The downwelling/subduction/sinking in our models, on the other hand, takes place in the subpolar gyres. The shallow cells in our models are therefore expanded versions of the shallow cells described by McCreary and Lu (1994) and Liu et al. (1994). The expanded cells are perhaps best described as mini‐ thermohaline circulations that deliver warm equatorial water into the subpolar areas north and south of the equator and bring cool subpolar water back into the equatorial zone below.

5. Data‐Model Comparison for the Prebomb Period (1940–1954) 14 The prebomb period provides the most straightforward comparison between our models and the Δ C deficits in the companion paper. Figure 4 shows the simulated deﬁcits for the years 1940–1954 for the coupled model (top) and in CORE‐forced model (bottom). The colored dots show the measured deﬁcits from the observations.

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14 The quantity being mapped is the Δ C difference from Okinawa. Areas with small differences between 0 14 and −5 are given a reddish‐orange shade. Areas with positive differences, that is, areas with higher Δ C values than Okinawa, have a deeper red shade. Deficits greater than −30 are shaded blue. The largest data‐model mismatches in Figure 4 are associated with three blue dots off the west coast of North America. The simulated deficits in these areas are small while the measured deficits in these areas are quite large. The water with the small deficits in the models is derived from the subtropical zone along ~25°N. Salty, 14 high‐Δ C subtropical water from this area flows toward the east and then northward into the Gulf of Alaska. As mentioned above, the northward flow into the Gulf of Alaska is mirrored by a southward flow off the coast of Chile. The latter produces an additional mismatch (a pale blue dot in an area with orange and red shades in the models). The poleward flows in both areas are completely spurious. A map of the surface salinity (Figure S2 in the Supporting Information) shows that fresh subpolar waters tend to flow equator- ward off North America and Chile instead. The maps in Figure 4 also have a blue dot off Peru (−38) and a white dot (−27) in the Galapagos Islands. As emphasized in the companion paper, the measured deficit off Peru, −38, is identical (as far as we can tell) to the prebomb deficit in SAMW, −35 to −40. The simulated deficits in the eastern Pacific, in contrast, are about −20. The measured deficits in the western equatorial Pacific, meanwhile, fall into the −14 to −20 range (pale orange shades) while the simulated deficits are close to 0. The observed deficits reflect the spreading of high‐deficit water from east to west. The near‐zero deficits in the models reflect a lack of spreading. These 14 differences are also reflected in the Δ C gradients across the Pacific. The E‐W gradient near the equator in the observations is small, only −12‰, due to more advection. The simulated E‐W gradient is larger, −20‰. The same distinction is seen below during WOCE in Figure 5, when the observed E‐W gradient is again half as large, −32‰, as the simulated gradient, −65‰. Something very similar is seen in the Atlantic, where the measured deficits off Florida (−17) and Puerto Rico (−19) are similar to those in the western Pacific. The simulated deficits in these areas are again near 0, reflecting a general lack of spreading from the east. The weak spreading is also apparent off northern Brazil where the measured deficit of −23 (pale yellow dot) stands out against simulated deficits of −5 and +5. 14 The weak spreading of low‐Δ C water from the upwelling zones cannot be explained by inadequate winds. The integrated divergence of the surface flow off Peru in our CORE‐forced model is about 2 Sv, the same volume of upwelling determined by Messié et al. (2009) for the Peru upwelling zone in the real ocean. There is instead a more clear‐cut explanation. The water drawn up to the surface in the eastern Pacific in our models is part of the shallow overturning described above in section 4. After reaching the surface, this water tends to flow away from the equatorial zone to the north and south. There is therefore less incentive for the water upwelling in the east to spread across the basin to the west. The subantarctic water drawn up to the surface in the real ocean, on the other hand, spreads naturally to the west because it leaves the Pacific en masse via the Indonesian Seas on its way back to the North Atlantic (Kessler, 2006). The same distinction would appear to hold in the Atlantic, where the subantarctic water drawn up to the surface off Africa flows into the Caribbean Sea on its way to the North Atlantic. In the Indian Ocean, four pale blue dots in the Arabian Sea stand out against the pale orange shades generated by the models. As discussed in the companion paper, the four observations are consistent with a subantarctic source for the water upwelled in this area. The same cannot be said for the water drawn up to the surface in our models. The simulated deficits in the Bay of Bengal, meanwhile, are notably larger than those in the Arabian Sea. This suggests that the upwelling in the Indian Ocean in our models has been shifted from the Arabian Sea to the Bay of Bengal.

6. Data‐Model Comparison for the WOCE Era (1990–1994) The results from the WOCE era in Figure 5 provide more observations to ﬁll in key details. Like the maps in Figure 4, the maps in Figure 5 have a number of blue dots in the Gulf of Alaska and off the west coast of

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Figure 5. Deﬁcit patterns from the coupled version of CM2Mc (top) and Coordinated Ocean‐ice Reference Experiment‐ forced version of CM2Mc (bottom) during the World Ocean Circulation Experiment (1990–1994). Observations from the companion paper are overlaid as colored dots.

North America that stand out against the color patterns in the models. The largest data/model mismatch is off Southern California where the measured deficit of −89 stands out against a deficit in our coupled model of +5! A similar model tendency seems to exist near 40°S off southern Chile as well. Perhaps the most important new observations in Figure 5 are the red and orange dots in the subtropics that 14 show where the surface waters with the highest Δ C levels were located. The red and orange dots are found 14 near 30°N and 30°S in all three oceans. The highest Δ C levels in our models, meanwhile, are found near 25°N and 25°S. Thus, there is a clear tendency for the subtropical surface waters in our models to be displaced toward the equator. 14 The displacement is due to the fact that the low‐Δ C subpolar waters in our models tend to impinge on the subtropical zones from both the north and south as part of the shallow overturning. In the Pacific, this

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tendency is especially evident east of Japan and north of New Zealand where the blue shades in the models extend nearly to 30°N and 30°S. Similar displacements are seen in the Atlantic and Indian Oceans. The overall color pattern during WOCE is also quite different from the pattern during prebomb time. Deep blue shades dominate in the North Atlantic and in the eastern Pacific. The Indian Ocean also switches from orange and yellow shades to mostly blue shades. This pattern holds in both the real ocean and in the models. The blue shades in the North Atlantic are found in areas with deep winter mixing. During prebomb time, the 14 surface waters and deep waters of the North Atlantic had Δ C values that were not very different. Under these circumstances, winter mixing produces modest surface deficits (Figure 4). But the large reservoir of deep water is not easily filled in with bomb 14C. The same winter mixing therefore leads to larger surface deficits and the deep blue shades during WOCE. In the eastern Pacific, the simulated deficits in our models flip from too small during prebomb time to too large during WOCE. More specifically the deficits in our models increase from −20 during prebomb time to −75 or so during WOCE. The measured deficits off Peru, in contrast, increase more modestly from −38 to −60. These patterns seem to reflect a difference in transit times to the upwelling areas. The SAMW upwelling off Peru in the real ocean seems to have a transit time of about 30 years (Rodgers et al., 2003). As such, the water upwelling off Peru during WOCE would have picked up some bomb 14C when it was last in contact with the atmosphere in the 1960s. This would account for the relatively small measured deficit off Peru (−60). The water upwelling in the eastern Pacific in our models, on the other hand, seems to have little or no bomb 14C. This is presumably because the upwelled water was last in contact with the atmosphere before the bomb era began. The lack of any bomb 14C in the water would account for the larger simulated deficits and the deep blue shades in the eastern Pacific in Figure 5. The simulated deficits off Namibia in the South Atlantic are similar to those in the eastern Pacific. Just offshore, the simulated deficits are actually in good agreement with the measured deficits. But again, the weak spreading in our models leads to large data‐model differences downstream in the western Atlantic, as seen in the yellow dots off northeast Brazil, in the equatorial zone, and off Puerto Rico and Florida.

7. Data‐Model Comparison for the CLIVAR Era (2003–2006) The simulated and measured deficits for the CLIVAR era (2003–2006) are shown in Figure 6. Where observations are available, the deficits during CLIVAR reveal some rather striking post‐WOCE changes. The simulated deficits, meanwhile, tend to be similar to those during WOCE, with blue shades in the eastern Pacific and North Atlantic. 14 14 Sometime around the year 2000 the Δ C of the atmosphere fell below the Δ C at Okinawa due to the ongoing burning of fossil fuels (Druffel et al., 2010). The period between WOCE (1990–94) and CLIVAR (2003–2006) therefore marks a changeover from the “bomb era” to the “fossil fuel” era for 14C. So, if bomb 14 14 C continues to reach a given upwelling area while the Δ C at Okinawa is declining (due to the fossil fuel effect), one would expect a smaller deficit in the upwelling area during CLIVAR than during prebomb time. The most telling data‐model comparison in this regard is in the South Atlantic where the measured deficit off southern Africa was indeed smaller during CLIVAR, −15 (pale orange dot), than during prebomb time, −27 (yellow dot). The deficits through the center of the tropical Atlantic were also smaller during CLIVAR. The deficits in our models, meanwhile, are larger during CLIVAR, −30, than during prebomb time, −15. The same tendency is also seen in the eastern Pacific, where the simulated deficits are larger during CLIVAR, −50 to −55, than during prebomb time, −20. The changeover in our models is therefore quite muted, as the water reaching the surface seems to have remained largely out of contact with the atmosphere through the entire bomb era. Another telling comparison is in the eastern North Atlantic where the measured deficits south of Iceland are relatively small during CLIVAR, −14 (orange dot) and −24 (pale yellow dot) compared with the deficits in the models (blue shades). The small deficits in the real ocean are consistent with the fact that the return flow of NADW in the real ocean is drawn up to the surface in the tropical Atlantic (Schmitz & Richardson, 1991; Wunsch, 1984). The return flow is thereby able to pick up bomb 14C and transport it into the North Atlantic, where it gradually fills‐in the deficits seen during WOCE.

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Figure 6. Deﬁcit patterns from the coupled version of CM2Mc (top) and Coordinated Ocean‐ice Reference Experiment‐ forced version of CM2Mc (bottom) during Climate and Ocean: Variability, Predictability and Change (2003–2006). Observations from the companion paper are overlaid as colored dots.

The simulated deficits in the eastern North Atlantic, in contrast, remain quite large. This is because the return flows of NADW in our models tend to pass through the tropical Atlantic without reaching the surface. The return flows therefore pick up and transport relatively little bomb 14C into the North Atlantic.

8. Discussion The tropical overturning circulations in our models and in the real ocean have orthogonal tendencies. The overturning in our models has a north‐south orientation that links the equatorial, subtropical, and subpolar 14 regions. As seen in the surface Δ C, the overturning in the real ocean has more of an east‐west orientation, in which water upwelling near the eastern margins ﬂows out of the equatorial zones in the west. An under- standing of how the model tendency develops can perhaps shed some light on the overturning in the real ocean. The discussion to follow will be focused on the Paciﬁc, where the divergent tendencies are most easily seen.

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8.1. Upwelling in the Eastern Pacific 14 As shown by TDB91, low‐Δ C subantarctic water from the area north of the ACC is drawn up to the surface again in the eastern Pacific. So, how does one understand this process? The usual approach is to view the upwelling as a response to the southerly winds off Peru and/or the easterly winds along the equator, but the upwelling produced this way does not necessarily draw up subantarctic water from the ACC. Our alternative is summarized in Figure 7. Each of the panels depicts the positions of four isotherms in a zonal section near the equator. The top panel shows the positions of the isotherms in a closed basin. The easterlies, in this case, push water across the basin and, in so doing, take water from the east and pile it up in the west. The easterlies thereby depress the isotherms in the west and elevate the isotherms in the east. The easterlies also produce a divergence along the equator that draws up water from the Equatorial Undercurrent (EUC). Thermocline water flows into the equatorial zone from the north and south to feed the EUC (Liu et al., 1994; McCreary & Lu, 1994). The lowermost isotherm in each panel depicts the coolest water that flows into the equatorial zone in this way. It would have a temperature of about 19 °C. It would also have a relatively high 14 Δ C, owing to its subtropical origin. The Pacific basin is not closed, however. Some of the warm water piled up in the west is drawn away (Lu et al., 1998). Accordingly, the middle panel of Figure 7 includes a stippled area that encompasses the warm water in the west. Warm water from the stippled area is shown being drawn away to the North Atlantic. With the removal of the warm water, the isotherms in the bottom panel of Figure 7 are shifted up toward the surface. The shift displaces the 19° isotherm away from the eastern boundary and exposes the cooler water below. With the shift, SAMW reaches the surface in the upwelling zones that are just off the equator along the eastern margin. The withdrawal in the west puts a different conceptual frame around the upwelling in the eastern Pacific. Without the withdrawal in the west, all the water pushed across the Pacific by the easterlies returns to the east 14 via the EUC. None of this water is very cool or low in Δ C. With the with- Figure 7. Development of the cool upwelling in the eastern Pacific. The drawal, on the other hand, the amount of water that returns via the EUC three panels depict the response of four thermocline isotherms to the is reduced—because the withdrawal in the west lowers the pressure head easterly winds near the equator. The top panel depicts the response in a that drives the EUC. More water is therefore pushed across the Pacific by closed basin (after Philander et al., 1987). The middle panel depicts an open basin where the warmest water is drawn away in the west. With the the easterlies than returns via the EUC. The difference, we claim, is the withdrawal (bottom panel) the four isotherms are shifted up toward the SAMW upwelled in the east. surface and the coolest isotherm is displaced from the eastern margin. From this perspective, the total amount of upwelling in the east is set by Subantarctic Mode Water (SAMW) from the Antarctic Circumpolar Current the withdrawal in the west. This amount exceeds (by perhaps 4–5 times) then resides near the surface in its place. The easterly wind stress, τx, is shown at the top of the figure. EUC = Equatorial Undercurrent. what one would expect from the winds off Peru. In this context, the local winds are a secondary influence. They may dictate where the upwelling of subantarctic water is focused or concentrated (as shown by McCreary et al., 2002), but they do not determine the amount. The upwelling in its entirety is presumed to be more widespread, that is, offshore from the upwelling zones and further up and down the South American coast.

8.2. Vertical Distribution of the Indonesian Throughflow The transport through the Indonesian Seas in CM2Mc is 13.5 Sv (Galbraith et al., 2011). This is a very reason- able transport in regard to the expected values (Godfrey, 1996). The problem is that the Throughflow in CM2Mc is weakly differentiated in the vertical. Half of the Throughflow in the model takes place between 300 and 600 m with an average temperature of 6.5 °C. The flow near the surface is quite weak. As a result, very little of the warm water piled up by the easterlies is drawn away.

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Figure 8. Comparison of the dynamic height at the surface relative to 1,500 m from CM2Mc (top), the Coordinated Ocean‐ice Reference Experiment (CORE)‐forced model (middle), and the observations (bottom) in the vicinity of the Indonesian Seas. The observations in the bottom panel feature a prominent trough along 8°N to the east of Mindanao that is missing in CM2Mc. The models, in turn, have a 0.3‐m step in the surface height off the northwest tip of Australia that is not present in the observations. The topography in all three maps is the bottom topography in the models at 1,500 m. NECC = North Equatorial Countercurrent.

The transport in the real ocean is skewed toward the surface. The peak flow through Makassar Strait, for example, is between 110 and 140 m (Gordon et al., 2008) where the water has an average temperature of about 23 °C (Ffield et al., 2000). The flow speed at the surface is also fairly strong, about 75% of the peak flow at 110–140 m. The water at the surface has an average temperature of 28 °C (Boyer et al., 2013). During La Niña summers, the peak flow shifts from 110–140 m into the upper 100 m and is about 50% stronger (Figure 4 in Sprintall et al., 2014). So, why does CM2Mc fail to produce a strong near‐surface outflow? Figure 8 shows the dynamic height at the surface (relative to 1,500 m) from CM2Mc (top), the CORE‐forced model (middle), and the real ocean (bottom), in the vicinity of the Indonesian Seas. The observed dynamic height in the bottom panel features a prominent trough along ~8°N to the east of Mindanao in the Philippines. The near‐surface outflow in the real ocean is known to take place along the northern edge of the trough (Gordon et al., 2008; Lukas et al., 1991; Qiu et al., 2015). The North Equatorial Countercurrent (NECC) flows away to the east along the southern edge of the trough.

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Figure 9. (top) Dynamic height at the surface relative to 2,000 m (copied from Figure 2.8 in Tomczak and Godfrey (1994). Our hypothesis is that low pressure from the North Atlantic reaches into the tropics in all three ocean basins, where it teams up with the local winds to maintain outflows of warm near‐surface water. Key features are the troughs of low pressure east of Mindanao in the Pacific and along 6°S off Africa in the Indian Ocean. (bottom) The return flow of North Atlantic Deep Water (red arrows) is overlaid on the contours from the map at the top.

The trough in the bottom panel of Figure 8 is completely missing in CM2Mc. Our CORE‐forced model has only the slightest hint of a trough. What our models have instead is a pair of ~0.3 m steps in the surface height between the Pacific and Indian Oceans. The steps can be seen off the northwest tip of Australia in the top and middle panels of Figure 8 where the Throughflow emerges from the Indonesian Seas. The steps in the surface height are an indication that the Throughflows in our models are maintained by an artificial pressure difference between the Pacific and Indian Oceans, which is presumed to be due to an inadequate resolution of the area where the Throughflow emerges from the Indonesian Seas. (The poor resolution leads to flow restrictions that back up the flow.) In the real ocean, cross‐stream pressure differences maintain a transport that is more geostrophiccally balanced. A strong near‐surface outflow is achieved when the cross‐stream pressure differences diminish with depth. The weak near‐surface outflows in our models provide a simple explanation for the north‐south orientation of the shallow overturning circulations in Figure 3. The subantarctic water entering the Pacific in the real ocean is cool; the water leaving the Pacific through the Indonesian Seas is warm. The net result is a buoyancy loss that offsets the buoyancy added by the Sun and atmosphere in the tropics. In contrast, the

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water leaving the Pacific in our models tends to have the same temperature as the inflow in the south. As a result, the buoyancy added from the Sun and atmosphere accumulates in relation to the real ocean. As this pool of buoyant water expands, some of it reaches the subpolar zones. Cooling of the buoyant water leads to sinking and the spurious shallow overturning. 8.3. Full Hypothesis The top panel of Figure 9 is a map of the dynamic height for the whole ocean copied from Tomczak and Godfrey (1994). (The dynamic height, in this case, is referenced to 2,000 m, and the height contours have units of meters squared per second squared, the dynamic height in meters mul- tiplied by the gravitational acceleration, 9.8 m/s2). The ACC stands out as the area of tightly packed contours along the bottom edge of the map. North of the ACC is a belt of relatively high pressure (high surface height) as seen in the area between the 17 and 21 contours in the South Pacific. This is the convergence zone where subantarctic water is piled up by the southern westerlies. The formation of NADW, meanwhile, produces an area of low pressure (low surface height) in the North Atlantic (Reid, 1961). Subantarctic water flows from the area of high pressure north of Figure 10. Schematics showing how areas of low pressure in the western the ACC to the area of low pressure in the North Atlantic. Pacific (top) and Indian Ocean (bottom) guide the currents that carry warm near‐surface water out of the Pacific and Indian basins. SECC = South Our hypothesis starts from the fact that the subantarctic water piled up in Equatorial Countercurrent; SEC = South Equatorial Current; the South Pacific is too warm and too light to flow directly into the NECC = North Equatorial Countercurrent; NEC = North Equatorial Atlantic via Drake Passage. So, it loops back around through the Current. Indonesian Seas and joins the ACC in the Indian Ocean instead (Schmitz, 1995; Talley, 2008). The looping flow is sketched out in the bottom panel of Figure 9 over the dynamic height map in the top panel. It begins as a subthermocline flow in the South Pacific, which turns to the east in the equatorial zone and shoals to the surface. Once at the surface, it turns back to the west and leaves the Pacific along the northern edge of the trough. It then enters the Indian Ocean through a series of narrow passages among the Lesser Sunda Islands to the east of the island of Java. For a geostrophic near‐surface outflow to exist there must be some continuity in the surface height field. In particular, there must be an area of low surface height to the east and south of Java that connects with the trough off Mindanao. The sort of height field that we have in mind is sketched out in the top panel of Figure 10. The hatched areas with red Ls denote areas of low pressure (low sea surface height) east of Mindanao and south of Java. The stippled area with a blue H is an area of high pressure. The water that is piled up in this area is pushed across the basin from the eastern Pacific. The gist of our hypothesis is that the areas of low pressure in the top panel of Figure 10 are connected to the area of low pressure in the North Atlantic. With the pressure field set up this way, the warm water piled up by the easterlies in the western Pacific can be drawn away to the North Atlantic. The warm water flows initially back to the east as part of the NECC and then circles back to the west along the northern flank of the trough. It then flows out of the Pacific and down the coast of Africa, as shown in the bottom panel of Figure 10. This process resembles the global seiching of upper‐ocean water between the Atlantic and IndoPacific basins described in Cessi et al. (2004) and Goodman (2001). The seiche, in these examples, is a wave that develops in response to an abrupt change in the surface height in the North Atlantic due to a shutdown or a switch‐on of NADW formation. The wave travels down the east coast of North America, across the equator, and around the tip of Africa. It then travels up the east coast of Africa, across the Indian Ocean, and through the Indonesian Seas into the equatorial Pacific. The seiche acts to restore the surface height difference between the Atlantic and Pacific basins following a disturbance. In the present context, the formation of NADW creates a negative disturbance in the surface height, while the buoyancy added by the Sun and the easterlies creates a positive disturbance in the western Pacific. The two disturbances come into balance when information about the negative disturbance is able to

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pass from the Atlantic to the Paciﬁc along the path followed by the seiche wave. This pathway appears to be blocked in our models by the pressure buildup across the Indonesian Seas.

8.4. Upwelling in the Eastern Atlantic The upwelling in the eastern Atlantic can be explained in a similar way. The Atlantic has easterlies and an EUC, and the flow that leaves the tropics off Florida is strongly differentiated in the vertical (Schmitz & Richardson, 1991). Low pressure from the formation of NADW, in this case, extends down the east coast of North America, around the tip of Florida and into the northern Gulf of Mexico (Figure 9). It then extends down the coasts of Central and South America into the area off French Guiana near 6°N. The Atlantic version of our hypothesis is that the low pressure from the North Atlantic encounters the warm water piled up by the easterlies off French Guiana near 6°N. The North Brazil Current carries warm equatorial water along the coast into this area. Some of the equatorial water continues up the coast into the Caribbean Sea (Johns et al., 1998) and becomes the near‐surface outflow in the Florida Current (Schmitz & Richardson, 1991). The rest retroflects back to the east into the NECC. The partitioning of the North Brazil Current between the Florida Current and the NECC should be directly related to the low pressure along the coast: a deeper low would be an indication that more warm equatorial water is flowing northward into the Caribbean Sea and then on to Florida. Data‐model differences in the surface height can be seen in the three panels in Figure 8 (along the right‐hand sides of the plots). The areas of low surface height in the western Gulf of Mexico and western Caribbean are poorly developed in our models in relation to the observations. We would argue that the surface height in these areas is elevated due to flow restrictions in the narrow passages between Florida, Cuba, and the Bahamas (the Florida Straits). With the flow restrictions, an artificial pressure difference develops from one side of Florida to the other that prevents the low pressure from the North Atlantic from reaching the tropics.

8.5. Upwelling in the Arabian Sea The Indian Ocean does not have easterlies like the Atlantic and Pacific and its main upwelling zone does not lie along an eastern margin. The northern Arabian Sea is, nevertheless, a major upwelling zone. So, how does one understand the upwelling in this area? As seen in the two panels of Figure 9, there is a trough in the Indian Ocean along ~6°S that is much like the trough along ~8°N in the western North Pacific. The South Equatorial Countercurrent (SECC) flows to the east along the northern edge of the trough while the South Equatorial Current flows to the west along the southern edge of the trough, as shown in the bottom panel of Figure 10. The outflow from the Pacific crosses the Indian Ocean along the southern edge of the trough as well. Like the trough in the western Pacific, the trough along 6°S may be instrumental in maintaining a flow of warm equatorial water back to the North Atlantic. The winds of the summer monsoon are known to push a large volume of warm water away from the Arabian Sea (across 9°N), which is balanced by a northward flow of cooler water in the Somali Current (Shi et al., 2002). The winds of the winter monsoon then drive a southward flow away from this area that becomes part of the SECC along the northern edge of the trough (Schott & McCreary, 2001). Without the trough, these flows would presumably reverse during the opposite seasons and not have much impact. But there clearly seems to be a convergence into the SECC from the north that persists over the annual cycle (Figure 5 in Peter & Mizuno, 2000). The inflow from the north then joins with the outflow from the Pacific and becomes part of the Agulhas Current along the coast of Africa, as illustrated in the bottom panel of Figure 10. As in the discussion above, the impetus for the circulation is the accumulation of buoyancy in the Pacific and Indian Oceans. With the two troughs, some of the warm buoyant water can be drawn away. The accumulation is thereby reduced with respect to the accumulation seen in our models.

8.6. More Highly Resolved Models The near‐surface outflows called for in our hypothesis take place through narrow passages like those in the Indonesian Seas and the Florida Straits. It should come as no surprise that the flows through these passages might be restricted in our models in a way that leads to ageostrophic down‐gradient flows.

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More highly resolved models should be less subject to these restrictions. To test this idea we examined the surface height fields in the three coupled models in Griffies et al. (2015). The three models are nearly identical except for their ocean resolutions, nominally 1°, 0.25°, and 0.10°. We compared the surface height during each model's spin up with the surface height fields from AVISO (https://www.aviso.altimetry.fr/). Our expectation was that the surface height fields in the models would become reoriented to support more geos- trophically balanced flows through the Indonesian Seas. The 1° model in Griffies et al. (2015) is able to resolve islands that CM2Mc cannot. Its flow paths are therefore more realistic. But more islands give rise to narrower passages. As it turns out, the surface height fields in all three of the Griffies et al. models depart from the observed surface height in much the same way. In particular, the trough to the east of the Mindanao weakens as the models are run out and height differences build in the outflow region in the south. As a result, the increased resolution in the 0.25° and 0.10° models does not have the expected impact. We also examined the models ESM2M and ESM2G in Dunne et al. (2012). The ocean components of both models, in this case, are built on 1° grids. The ocean component in ESM2M is based on the Modular Ocean Model, like the three models in Griffies et al. (2015). The ocean component in ESM2G is based on the Generalized Ocean Layer Dynamics model developed by Hallberg (1995). Among other differences, Generalized Ocean Layer Dynamics uses a horizontal grid discretization (the Arakawa C grid) that is better able to describe the flows through narrow channels. The C grid in ESM2G appears to make a difference in the Pacific. The trough to the east of the Philippines is deeper and the height fields east of the Philippines and south of Java are connected in a more gradual way without an obvious step. As a result, ESM2G has a stronger near‐surface outflow than ESM2M. It also has cooler sea surface temperatures (SSTs) in the equatorial Pacific and slightly cooler SSTs offshore from Peru (see Figure 5 in Dunne et al., 2012). ESM2G still has a warm bias, however, right along the coast of Peru. This may be due to model winds that do not concentrate the upwelling sufficiently right along the coast. On the other hand, ESM2G and ESM2M have similar SST biases in the tropical Atlantic. The SSTs in both models are too warm by the same amounts off southern Africa. Thus, the upwelling in the tropical Atlantic is not notably improved by the C grid in ESM2G. At the end of the companion paper, we show empirically that there are two kinds of upwelling at work in the upwelling areas north of the ACC. One is a response to the winds in the upwelling zones; the other is associated with the overturning of NADW. The second type is missing in CM2Mc and is largely missing, it would seem, in the models considered above. The second type is arguably more important volumetrically than the first and should therefore be very important for the ocean and climate system. The second kind of upwelling alters how we think about the ocean in three ways. First, the extra volumes of upwelling could significantly enhance the uptake of heat and the export of heat away from the tropics in both hemispheres. Ocean heat uptake, though fairly well known, is not well represented in ocean models, and the extra upwelling could easily explain why the models fall short, for example, Wang et al. (2014). Second, the extra upwelling should enhance the biological productivity of the oceans by increasing the supply of nutrients from below. Modern estimates of the biological productivity tend to be based on nutrient inputs derived from models, for example, Dunne et al. (2013). These may be artificially low, especially in the upwelling areas. Third, the extra volumes of upwelling imply that more of the ocean's overturning is routed through the surface layers of the ocean in low latitudes. As with the heat uptake above, this is known in principle, for example, Schmitz and Richardson (1991), but is poorly represented in ocean models.

9. Conclusions We propose that cool subantarctic water is drawn up to the surface in the eastern tropics in response to out- ﬂows of warm water from the western tropics. The driving forces are the buoyancy added by the Sun and tropical atmosphere, the easterly winds near the equator, and the formation of NADW in the North Atlantic. As argued here, the low pressure (low surface height) from the formation of NADW is able to reach into the tropics in a way that allows the warm water in the western tropics to be drawn away. As the warm water is drawn away in the west, cool subantarctic water is drawn up to the surface in the east. The

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overturning of NADW is thereby brought directly up to the surface in a way that greatly enhances the heat uptake by the ocean and the transport of heat out of the tropics. The outﬂows in the real ocean occur via geostrophic currents that pass through narrow channels in the Indonesian Seas and Florida Straits. These currents are poorly developed in the models examined in the paper. The culprit may be ﬂow restrictions that lead to along‐stream steps in the surface height that prevent the low pressure from the North Atlantic from reaching into the tropics. As a result, the overturning of NADW in our models tends to pass through the tropics without reaching the surface. This shortcoming has implications for how the ocean's heat transport and the nutrient supply are represented in ocean models and for the way that models are used to help us understand the ocean's overturning.

Acknowledgments References This study is based on hundreds of Δ14C measurements made by other Boyer, T. P., Antonov, J. I., Baranova, O. K., Coleman, C., Garcia, H. E., Grodsky, A., et al. (2013). World Ocean Database 2013. In S. Levitus investigators. The authors would first & A. Mishonov (Eds.), NOAA Atlas NESDIS 72 (209 pp.). Silver Spring: MD. like to acknowledge the important Broecker, W. S., & Peng, T.‐H. (1974). Gas exchange rates between air and sea. Tellus, 26, 21–35. contributions made by the people who Cessi, P., Bryan, K., & Zhang, R. (2004). Global seiching of thermocline waters between the Atlantic and Indian‐Pacific Ocean basins. produced these measurements. The Geophysical Research Letters, 31, L04302. https://doi.org/10.1029/2003GL019091 documentation for these contributions Delworth, T. L., Broccoli, A. J., Rosati, A., Stouffer, R. J., Balaji, V., Beesley, J. A., et al. (2006). 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