Sea surface temperature in EBUS

To address this and other societally relevant issues, we conducted an unprecedented ensemble of high-resolution historical and future climate simulations (hereafter HR) using CESM with a nominal horizontal resolution of 0.25° for the atmosphere and land components and 0.1° for the ocean and sea-ice components. The results are compared to a parallel set of low-resolution simulations (hereafter LR) at a nominal resolution of 1° for all model components38 (also see Methods). Improvements in HR over LR are evident. HR realistically simulates the narrow structure and intensity of coastal upwelling in all EBUSs, including CUS and CCS in the Northern Hemisphere (NH), and BUS and P-CUS in the Southern Hemisphere (SH) (Fig. 1), based on a comparison between  observed and simulated annual mean sea-surface temperature (SST) in these regions during 1991-2020. Although the narrow upwelling zone along the northern Benguela coast (north of ~25°S) in the BUS is still underrepresented by HR, it is far more realistic than in LR, especially in the central Benguela, and in the southern Benguela (south of ~30°S) HR even shows cooler-than-observed SST that suggests an overestimate of the upwelling. In all other EBUSs the HR-simulated SSTs show a remarkable agreement with the observations, while the narrow intense coastal upwelling is virtually absent in LR. These results hold not only for the CESM simulations, but also for other climate model simulations within the High-Resolution Model Intercomparison Project39 (HighResMIP) (Supplementary Fig. 1). Consistent with previous studies32,34, the vastly improved representation of the coastal upwelling in HR is closely linked to the better resolved LLCJs and associated wind stress curls along the upwelling coasts, particularly in terms of the jet core structures and their proximity to the coast (Supplementary Fig. 2). Since the atmospheric model resolution in HighResMIP ranges from 0.25° to 0.5° (Supplementary Table 1), it suggests that a 0.5° or finer atmospheric grid is required to simulate low-level coastal jets, and the associated coastal wind stress and wind stress curl32,40. In addition, ocean eddies and upwelling fronts are explicitly represented in HR rather than parameterized as in LR. In fact, a comparison between satellite observed and HR-simulated eddy kinetic energy (EKE) in EBUSs shows an overall agreement, except that in the Canary and Chilean region HR tends to underestimate EKE (Supplementary Fig. 3). Together, these results suggest that HR is more credible and represents dynamical processes much more realistically for projecting future changes along EBUSs. In the following, we focus our analyses on the projected EBUS changes in HR.

Fig. 1: Eastern boundary upwelling systems (EBUSs) and sea-surface temperature (SST).
figure 1

Annual mean SST bias (°C, model minus observation) from HR and locations of five EBUSs (black boxes) (a) and difference between LR and HR SST (°C, b). Observed SST45 (°C) in CCS (c), CUS (d), P-CUS (e, f), BUS (g). HR SST (°C) in CCS (h), CUS (i), P-CUS (j, k), BUS (l). LR SST (°C) in CCS (m), CUS (n), P-CUS (o, p), BUS (q). Observed and simulated SSTs are all averaged for the period of 1991–2020. Contours in (a) and (b) are −3, −1 (dashed) and 1, 3 °C (solid).

Projections of winds, ocean temperature, and vertical velocity

To start, we examine the difference in alongshore wind stress strength between future and historical periods, since it lies at the heart of the Bakun hypothesis. From Fig. 2a–e, it is immediately clear that not all regions show strengthened winds. In the NH, the CCS shows unambiguous future weakening of the wind stress with a maximum change near 38°N just south of Cape Mendocino. In contrast, the change in the CUS is more complex; although there is an overall increase in the future alongshore wind stress, it weakens marginally near the coast. Such a change in wind stress pattern will lead to a change in near-coast wind stress curl (Fig. 2g). Consistent with previous studies8, the wind stress change in the CUS shows a poleward shift. In the SH, however, the winds off southern Chile show a substantial future strengthening, while there is also a general strengthening off the Benguela coast. The Peruvian system, however, does not exhibit significant wind stress changes (Fig. 2c). Therefore, there is clearly a regional dependence as to whether the alongshore winds follow the Bakun hypothesis. In the NH EBUSs, the coastal winds show a decrease in strength, which is at odds with the Bakun hypothesis, whereas in the SH EBUSs the winds are strengthening as predicted, except off the Peruvian coast where wind changes are mostly insignificant. Consistent with the wind stress changes, the project wind stress curl changes show a substantial future decrease in the CCS and increase off Chilean and Benguela coast, but insignificant change off Peruvian coast (Fig. 2f–j).

Fig. 2: Projected future changes of alongshore wind stress and ocean temperature in each EBUS from HR.
figure 2

Projected change (defined as the mean over the future period of 2071–2100 minus the mean over the historical period of 1991–2020) of alongshore wind stress (×10−2 N m−2) shown in longitude-latitude plane (upper) for CCS (a), CUS (b), P-CUS (c, d), BUS (e), and projected change of wind stress curl (×10−7 N m−3) (middle) for CCS (f), CUS (g), P-CUS (h, i), BUS(j). The upwelling favorable wind stress is southward (negative) in the NH and northward (positive) in the SH. Upwelling favorable wind stress curl is positive in the NH and negative in the SH. Areas without green hatching represent the regions where the mean alongshore wind stress over the historical period of 1991–2020 is upwelling favorable (see Supplementary Fig. 2). In order to avoid confusion in the sign of future change, future strengthening (weakening) in the upwelling favorable alongshore wind stress (ae) and curl (fj) is shown as positive (negative) values here, regardless of the EBUS hemisphere. Projected change of ocean temperature (°C in color) and historical annual mean temperature (°C in contours) shown in cross-shore section as a function of distance (°) from the coast and depth (m) (bottom) for CCS (k), CUS (l), P-CUS (m, n), BUS (o). The cross-shore sections of temperature in (k)–(o) are derived by averaging the temperature along each upwelling coast within the area indicated in green in (a)–(e) and (f)–(j). Areas without gray hatching represent the regions where the changes are significant at 95% level based on a Student’s t test.

Next, we examine alongshore variations of future coastal upwelling changes in EBUSs. Historically, coastal upwelling has been estimated using the so-called Bakun index2. A revision to the Bakun index has been recently proposed that includes additional upwelling dynamics35. Here, we include all the dynamic processes affecting coastal upwelling in the model by computing an upwelling index using the simulated vertical velocity at the Ekman depth (see Methods) within a narrow coastal zone (Fig. 3). Comparing HR and LR shows that the HR upwelling index for the historical period is a factor of 3-5 times stronger than that in LR and shows a much richer alongshore variability (Fig. 3 left). The CUS displays the strongest mean upwelling (2.4–7.9 × 10−5 m s−1) and the strongest alongshore variability off the coast of Morocco, which appears to be linked to the small-scale topography (Supplementary Fig. 4). However, such features are completely missing in LR that shows only a gradual northward decrease in upwelling strength from 0.6 × 10−5 m s−1 in the southern CUS to 0.4 × 10−5 m s−1 in the northern CUS. Similar differences between HR and LR are also seen in other EBUS. The projected future upwelling changes show equally striking differences between HR and LR (Fig. 3 right). While LR projects generally weak and nearly uniform future upwelling changes along the coast, which agrees with previous studies10, HR projections are much more dynamic and spatially variable. For example, while LR projects a decrease in upwelling of 10−6 m s−1 along the Peruvian coast, HR projects a local maximum upwelling decrease of one order magnitude larger near 15°−16°S in the southern Peruvian coast. Along the Chilean coast, LR again projects a nearly uniform future increase of coastal upwelling, but HR shows that the upwelling increase near 37°S is 2–3 times larger in other regions. In the northern BUS between the North and Central Namibian Upwelling Cells, LR projects a future upwelling decrease, while HR projects a substantial increase of 1–2 × 10−6 m s−1. The upwelling system showing the most consistency between HR and LR is the CCS, where HR and LR both project a nearly uniform future upwelling decrease south of 40°N, except that the magnitude of the projected decrease in HR is 2–3 times of that in LR. As in the mean upwelling, the CUS stands out as showing the strongest HR-projected future upwelling changes, as well as the strongest alongshore variations. The projected alongshore variation is so strong that the sign of the future upwelling changes can alter within a short distance. This suggests that CUS is a very dynamic system, with future changes likely involving complex multiscale interactions between coastal circulations, ocean eddies, and local fine-scale topography.

Fig. 3: Upwelling index and the associated futures change off the coast of each EBUS.
figure 3

Mean upwelling index (×10−5 m/s) averaged over the historical period of 1991-2020 (left) in HR (gray shade and purple with squares) and LR (green with triangles) and the corresponding future change (×10−5 m/s) computed as the future mean (over 2071–2100) minus the historical mean (over 1991–2020) (right) in HR (dots and purple line) and LR (green with triangles), respectively, as a function of latitude for CCS (a, b), CUS(c&d), P-CUS (e, f, g, h), and BUS (I, j). The upwelling index is computed based on simulated vertical velocity at the Ekman depth averaged in a coastal bin of 1° × 1° for LR (green with triangles) and ~0.5° (cross-shore) × 0.1° (alongshore) for HR (gray shade in left panels and color dots in right panels) to reflect model resolution difference. Red and blue dots indicate future increase and decrease in the HR upwelling index, respectively, with dark (light) colors indicating passing (failing) Student’s t test with 95% significance level. A smoothed upwelling index and future change in HR (purple) are also shown by averaging over 1° bin alongshore to compare directly with LR.

While there is a dynamical consistency between projected changes in the upwelling index and alongshore wind stress and wind stress curl in most EBUSs except off the Peruvian coast where wind changes are largely insignificant (Fig. 2c, h), future ocean temperature changes show a more perplexing picture. Here, we focus on the annual-mean vertical temperature structure averaged along each upwelling coast, as the annual-mean changes are similar to upwelling-season mean changes (Supplementary Fig. 5). In all EBUSs, the simulated vertical temperature structures during the historical period show upper ocean isotherms tilted upwards towards the coast (solid contours in Fig. 2k–o), indicative of a well simulated upwelling in these regions consistent with the simulated SST (Fig. 1). Like the winds, the projected temperature changes again show differences between the NH and SH (color in Fig. 2k–o). In the NH, the CCS shows a strong and broad upper-ocean warming, but within 100 km (~1°) of the coastal zone where the upwelling is the strongest, future warming is reduced in the upper 50 m (Fig. 2k). The CUS is also projected to have less future warming near the coast compared to the open ocean, with the projected cross-shore gradient change being more prominent than in the CCS (Fig. 2l). These projected temperature changes seemingly imply strengthened Ekman upwelling, consistent with the Bakun hypothesis, although the coastal winds in these regions are projected to weaken (Fig. 2a, b), inconsistent with strengthened Ekman upwelling. In the SH, all three EBUSs show enhanced warming within the coastal zone in response to anthropogenic forcing, which clearly contradicts the Bakun hypothesis, even though the coastal wind changes (Fig. 2d, e) and the upwelling index changes (Fig. 3h–j), especially for Chile and the Benguela, are in line with Bakun’s ideas. These results hold not only for the CESM, but also for the multi-model ensemble mean of HighResMIP models (Supplementary Fig. 6).

Heat budget for the ocean temperature increase

Clearly, there is a discrepancy between temperature in the upwelling region and coastal wind responses to future warming as portrayed by the Bakun hypothesis that is entirely based on vertical Ekman upwelling changes. This discrepancy is, to a significant extent, caused by neglecting the contribution from horizontal transport of heat from the tropics to EBUS in Bakun’s mechanism. Figure 4 shows the 2000-2100 warming trend in the upper 50 m temperature, area-averaged within the ~1° coastal zone of the EBUSs, along with contributions from each term in the heat budget (see Methods). The warming trend varies from 0.16 Wm−2 in the CUS to 0.27 Wm−2 in Peru and CCS, and the cause of the warming differs drastically between NH and SH EBUSs. For all SH EBUSs, ocean heat advection is the dominant contributor to the warming trend, while surface fluxes are the dominant terms counteracting the warming. Decomposing the heat advection into its components indicates that it is the mean horizontal advection that makes the primary contribution to warming in the SH EBUSs (Fig. 4b). The warming trend from the mean horizontal advection can arise from 1) equatorial oceans, particularly the equatorial Pacific, warming faster, and 2) stronger LLCJs leading to an increase in the upwelling-favorable wind stress curl near the coast, which in turn drives stronger poleward coastal currents through Sverdrup balance. Both processes enhance warm advection from the tropics to the upwelling regions, producing faster warming within the coastal zone than in the open ocean. Because the Peruvian system is closest to the equator and shows a weak wind stress curl change (Fig. 2h), the warm advection in this region likely comes primarily from the faster warming equatorial Pacific. Indeed, HR shows a prominent future warming of SST in the equatorial Pacific and to a lesser extent warming in the equatorial Atlantic (Supplementary Fig. 7). On the other hand, the Chilean and Benguela systems see significant strengthening in the local wind stress curl under future warming (Fig. 2i, j), suggesting that the intensified alongshore poleward flows contribute more to the warm advection. Neither of these warming mechanisms are included in the Bakun hypothesis and can only be fully captured by high-resolution climate models, because LLCJs and the associated coastal wind stress curl, both poorly represented in low-resolution models, are vital. We emphasize that the key difference between HR and LR occurs near the coast. In LR, the warm advection via Sverdrup transport is too broad and very weak near the coast. Therefore, it has little impact on the coastal warming, which is sharp contrast to the strong and narrow coastal warm advection in HR.

Fig. 4: Heat budget for the warming trend in each EBUS.
figure 4

a The warming trend (red), and contributions from surface heat flux (yellow), total ocean heat advection (cyan), and turbulent mixing (blue). b Contributions from eddy-induced advection (green) and mean-current-induced advection (olive green) to total ocean heat advection (cyan in (a)), and contributions from horizontal mean advection (seagreen) and vertical mean advection (gray) to mean-current-induced advection (olive green). c Contributions from shortwave (green), net longwave (olive green), latent (seagreen), and sensible (gray) heat flux to the total heat flux (yellow in (a)). The trend in each budget term (in W m−2) is computed based on averages within the top 50 m and over a 100 km wide strip next to the coast within each green area indicated in Fig. 2a–e, using a 3-member ensemble mean of monthly mean output from HR for the period of 2000–2100 after subtracting the preindustrial control simulation. Note that there is a large amount of cancellation between the advection and surface heating terms, such that the net trend is relatively small compared to these individual terms, but remains positive for all EBUSs.

In NH the largest contributor to the warming trend is the net surface heat flux, indicating that the coastal warming trend is driven by atmospheric heating. Ocean heat advection is primarily responsible for driving a cooling trend, in sharp contrast to the SH. The cold advection is ~3 times stronger in the CUS than in the CCS (Fig. 4). Further decomposition of the advection shows the mean vertical heat advection, rather than the horizontal advection as in the SH, is the dominant contributor to the cooling trend, and its strength in the CUS is more than twice of that in the CCS. This stronger vertical heat advection cannot be explained by increased vertical Ekman upwelling because the coastal winds are weakened under future warming (Fig. 2). As the vertical velocity in the CCS decreases (Fig. 3a, b) in the future, the enhanced cooling due to the vertical heat advection here is most likely caused by the increased vertical temperature gradient within the upwelling zone as the increased surface heat flux increases upper ocean stratification. This differs from the CUS, where despite a decrease in the alongshore wind stress, the vertical velocity shows a future increase (Fig. 3c, d). This increase can be attributed to the intensification of near-coast wind stress curl as the wind pattern changes off the CUS under future warming (Fig. 2). Indeed, there is a narrow stretch of strengthened wind stress curl along the Canary coast from 25°N to 28°N (Fig. 2g), which coincides with the large increase of the near-coast vertical velocity (Fig. 3). These results suggest that the strong increase in the vertical cold advection in the CUS depends on increases in both ocean stratification and vertical velocity driven by intensified near-coast wind stress curl through Ekman suction. Both these mechanisms are more complex than the classical Ekman upwelling mechanism invoked by Bakun.

Regarding the contrasting role of surface heat fluxes in the NH and SH, a decomposition of surface heat fluxes shows that the shortwave is always decreasing due to the aerosol scattering and absorption41 (Fig. 4c). In contrast, the net longwave radiation is seen to help warming in all EBUSs. Interestingly, turbulent heat fluxes show a different role between NH and SH, with less heat release from the ocean in NH but more heat release in SH. The changes in the turbulent heat fluxes are consistent with the wind changes, i.e., less turbulent heat loss in NH where the winds get weaker and more in SH, in agreement with the increasing winds. Thus, turbulent heat fluxes work in concert with longwave radiation to drive warming in the NH EBUSs, while in the SH EBUS they act with shortwave radiation to drive cooling .

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