1. Introduction
The meridional overturning circulation is responsible for the exchanges of heat and dissolved materials meridionally and vertically, which has a significant impact on the global climate and biogeochemical systems (e.g., Johnson et al. 2007, 2019; Chen and Tung 2014). In contrast to the Atlantic Ocean, the North Pacific Ocean lacks a local bottom water formation due to the small width and shallow depth of the Bering Strait (Talley 2013). Therefore, the Pacific deep water is transported northward from the Antarctic by the lower and upper deep branches of the Pacific meridional overturning circulation (L-PMOC and U-PMOC) through different pathways, which are also known as the deep western boundary currents (Talley 2013; Voet et al. 2015). The L-PMOC carries the Lower Circumpolar Waters (LCPW) with potential temperature (θ) less than 1.2°C, roughly corresponding to the depths greater than 3500 m in the tropical western Pacific. The U-PMOC takes the Upper Circumpolar Waters (UCPW) with θ over 1.2°–2.2°C and is located over ∼2000–3500 m in the tropical western Pacific (Johnson and Toole 1993; Kawabe and Fujio 2010). The LCPW transforms into the North Pacific Deep Water (NPDW) through upwelling in the North Pacific, which occupies the same layer of the UCPW (Siedler et al. 2004). Compared to the NPDW, the UCPW is featured with the same ranges of θ but slightly higher dissolved oxygen and lower silicate (Kawabe 1993; Reid 1997). Therefore, the boundary between the NPDW and UCPW is difficult to identify locally. We treat their mixed water with θ over 1.2°–2.2°C as the UCPW for convenience.
The general route of the U-PMOC from the Antarctic to the North Pacific has been sketched by previous studies (e.g., Kawabe et al. 2006; Komaki and Kawabe 2007; Reid 1997). The UCDW detaches from the Antarctic Circumpolar Current, enters the South Pacific subtropical anticyclonic flow system northward (Reid 1997), and then passes east of the Solomons Rise into the North Pacific (Kawabe et al. 2006; Komaki and Kawabe 2007). Based on the isobaths of 2500 m, the Yap–Mariana Junction (YMJ), Izu–Ogasawara Ridge, Mariana Ridges, and north of New Guinea (NG) contain the deep channels for the U-PMOC flowing between the Philippine Sea (PS) and the North Pacific (Fig. 1a). Wang et al. (2020, 2021b) have revealed the seasonal intrusions of the L-PMOC and U-PMOC into the PS at the YMJ based on direct mooring measurements. Through the deep channel north of NG, Kawabe et al. (2003) suggested that the U-PMOC flows westward into the PS with a volume transport of more than 3.8 Sv (1 Sv ≡ 106 m3 s−1), inferred from the distribution of dissolved oxygen and geostrophic flow data. Based on the silicate distribution and geostrophic flow along 137°E, Kaneko et al. (1998, 2001) deduce that the U-PMOC can flow into and out of the PS through different deep channels of the Izu–Ogasawara Ridge and Mariana Ridges. However, due to the lack of direct long-term measurement of deep velocity, the volume transports of U-PMOC through the above channels and their variability still have considerable uncertainty. Besides the lateral exchange of UCPW, there also exists the vertical flux of UCPW. Most of the LCPW in the PS needs to upwell into the upper deep layer due to the geographical constriction and then participates in the UCPW transport (Kawabe and Fujio 2010; Wang et al. 2021b). The net lateral transport of U-PMOC into/out of the PS will facilitate our understanding of the rate of LCPW diapycnal upwelling in the PS in combination with deep water mass observations.
(a) Schematic flow pathways of U-PMOC (red arrows, based on this study) and L-PMOC [orange arrow, based on Wang et al. (2021b)] in the western Pacific Ocean with a zoomed-in view of the deep channel north of NG in the northeastern corner. Bathymetry from ETOPO 1 is denoted by color. The purple stars indicate the mooring sites, and the cyan triangles indicate the pressure station used in section 3d. WCB: West Caroline Basin; ECB: East Caroline Basin; ER: Eauripik Rise; Sol.: Solomon; IOR: Izu–Ogasawara Ridge. (b) Latitude–depth section of the GLORYS12V1 mean zonal velocity (U) from 2014 to 2019 along the dashed line in (a). Green triangles denote the depths of moored sensors.
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
To make progress on this topic, we deployed four moorings in the deep channels north of NG from 2014 to 2019 (Fig. 1a). These moorings are a part of the Scientific Observing Network of the Chinese Academy of Sciences (CASSON) in the western Pacific Ocean. The mean status and seasonal variability of U-PMOC’s pathway and volume transport are first presented. The mechanism behind the observed U-PMOC and its linkage with upper-ocean forcing are explored with the aid of a high-resolution oceanic reanalysis product.
2. Data and methods
a. Data
1) Mooring data
The four moorings are NG1 at 0°, 142°E; NG2 at 1°S, 142°E; NG3 at 1.7°S, 141.4°E; and NG4 at 2°N, 140°E with water depths of 3378, 3032, 3583, and 4141 m, respectively. The measurement period of NG1 and NG4 is from 20 August 2014 to 29 October 2019, and that of the other two is from 1 November 2015 to 29 October 2019. Each deep mooring was equipped with several discrete JFE INFINITY, AANDERAA, and/or Nortek Aquadopp acoustic current meters, and SBE 37-SM conductivity–temperature–depth (CTDs) and RBR solo temperatures (RBRsTs) over the depth range of 1500–3500 m. The current meters returned hourly records of point velocity. The CTDs sampled temperature, salinity, and pressure every 10 min, and RBRsTs measured temperature once per second. The depths of the RBRsTs are derived using the recorded depths of CTDs and the nominal separation depths between the RBRsTs and adjacent CTDs. The value of θ corresponding to a measured in situ temperature is obtained using the salinity value from the CTD cast carried out during the mooring deployment. Due to the battery failure, the measurements in some layers experienced some data loss.
2) Reanalysis products
We use the daily and monthly velocity, temperature, and salinity data over 2014–19 from the global oceanic reanalysis product GLORYS12V1 (GL12). The GL12 from Mercator Ocean is analyzed to gain insights into the sizeable temporal–spatial variation of U-PMOC north of NG and its controlling mechanism. The GL12 is created using the Nucleus for European Modeling of the Ocean (NEMO; Madec and NEMO Team 2016) and has a horizontal resolution of 1/12° in longitude and latitude and 50 vertical layers with a resolution of 1 m at the surface and 450 m near the bottom. GL12 jointly assimilates altimeter sea surface height, satellite sea surface temperatures, and in situ temperature and salinity profiles (Lellouche et al. 2021).
We also use the monthly sea surface wind stress data over 2014–19 from the fifth-generation European Centre for Medium-Range Weather Forecasts Reanalysis (ECMWF-ERA5, Hersbach et al. 2020). The ECMWF-ERA5 wind is analyzed to investigate the annual characteristics of the Pacific equatorial zonal wind and its relationship with the Rossby waves. It has a horizontal resolution of 1/4° in longitude and latitude.
b. Methods
1) Annual harmonic analysis
To extract the seasonal and mean characteristics, we perform an annual harmonic fit on the observed deep zonal velocities (U) using the function of U = A cos(ωannualt + φ) + Um, where t is time, A, ωannual, and φ are the annual amplitude, frequency, and phase, respectively, and Um is the residual mean velocity. The phase is defined as the month of the strongest eastward current. We also calculate the seasonal explained variance (R2). Prior to calculation, the time series are smoothed by a 61-day moving average to focus on the variability longer than the intraseasonal time scale. The R2 is defined as the ratio of the variance of the fitted time series to the variance of the smoothed time series.
2) Transport calculation and error estimation
To calculate the U-PMOC transport through the deep channel north of NG, we grid the cross-channel section with a vertical resolution of 10 m and a horizontal resolution of 1/12° based on the ETOPO1 bathymetry data. The zonal velocities in the last grid nearest the bottom and the channel sidewalls are set to zeros. Then, we interpolate U in the vertical and horizontal directions using the linear interpolation method. The U-PMOC’s volume transport (VTU-PMOC) is calculated by integrating U over the depth range between the isotherms of 1.2° and 2.2°C (z axis) and over the cross-channel section (y axis) as
We further evaluate the error (ε) for the VTU-PMOC estimation due to the finite length of the time series. We first obtain the residual time series of VTU-PMOC after subtracting the seasonal variation and time-mean value. The value of ε is calculated as
3) Streamfunction calculation
3. Results
a. Observed U-PMOC
At four mooring sites, the upper deep layer with θ over 1.2°–2.2°C roughly corresponds to the depth from 1920 m to the bottom. The observed velocities in the upper deep layer exhibit a prominent seasonal variation (Fig. 2 and Table 1). We extract their seasonal and mean characteristics by annual harmonic analysis in section 2b(1). At NG1 and NG2, the values of R2 are generally larger than 50%, and the seasonal amplitude reaches 3–5 cm s−1 (Fig. 2). At NG4, the seasonal cycle of U is significant at 2078 and 2589 m with R2 ≥ 40%, but is weakened in deeper layers of 3043 and 3896 m. At NG3, the seasonal cycle of U is evident at 2023 and 3477 m with R2 > 40%, and is very weak in two layers of 2538 and 3010 m with R2 < 25%. However, R2 of meridional velocities (V) increase to 25%–50% in the layers deeper than 2000 m, possibly due to the influence of topography on the flow direction and velocity (figure not shown). The phase mostly occurs in November–February, and shows an increasing trend with decreasing depth at each mooring site (Table 1). For a similar depth, the phase month decreases from the equator to the south. The observed θ at four mooring sites also shows a seasonality, whose features are similar to the observed U and are not repeated (Fig. 3).
Time series of observed (orange) and GL12 (blue) daily deep zonal velocities (U) at various depths of (top left) NG1, (top right) NG2, (bottom left) NG3, and (bottom right) NG4. The solid green curves and black dashed lines indicate the annual harmonic fit and residual mean (Um) of observed zonal velocities, respectively. The mean depths of the observed records and seasonal explained variances (R2) are denoted in each panel.
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
Results of annual harmonic fit on observations and comparisons between GL12 and observations at four mooring sites. See text for details.
As in Fig. 2, but for the daily potential temperature (θ).
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
The U-PMOC after the seasonal cycle removed has a mean eastward velocity core at ∼2550 m of the equator (Fig. 2 and Table 1). At NG1, the mean zonal flow at the upper deep layer is strongest at 2570 m with positive Um reaching 2.7 cm s−1. This eastward flow becomes weakened with Um of 1.06 cm s−1 at a greater depth at 2947 m, and reverses to a westward flow at the upper boundary of the upper deep layer. At NG2, the observed mean U-PMOC flows eastward with a maximum speed of 1.38 m s−1 at 1987 m and a gradually weakened speed toward the bottom. At NG4, Um at 1634 and 2078 m show large values and a decreasing trend with depth. These two depths are located near the upper boundary of the upper deep layer, and Um is influenced by the eastward flowing North Intermediate Countercurrent. At greater depths from 2589 to 3896 m, mean zonal flow is very weak with a nearly zero Um, probably due to the obstruction of Eauripik Rise. At NG3, the upper deep zonal flow becomes relatively weaker compared to that at the equator, and the mean directions are westward at 2023 and 2538 m and eastward at 3010 and 3477 m.
We next estimate the integrated and layered transports in the upper deep layer and the U-PMOC volume transport (VTU-PMOC) through the deep channel north of NG along 142°E based on mooring observations at NG1, NG2, and NG3 according to the method in section 2b(2). We assume that the observed U from the three moorings represent the flow across the section of 142°E, although their longitudes are slightly different. Figures 4a and 4b show the mean integrated transport from the bottom and layered transport over different potential temperature and depth ranges. When the upper integration boundary is changed from 1.2° to 1.8°C, the integrated transport shows a continuously increasing trend, and the layered transports over this range are all positive with the maximum value at 2550 m. The integrated transport decreases when the upper integration boundary becomes shallower, and the layered transport becomes westward. The orange curve in Fig. 4c shows the time series of observed VTU-PMOC. According to the annual harmonic fit on VTU-PMOC, the time mean value is 0.76 Sv, and the seasonal amplitude and explained variance are 9.94 Sv and 63%, respectively. The standard deviations are 8.89 and 5.41 Sv when including and excluding the seasonal cycle, respectively. The maximum eastward and westward transports occur in January and July.
Volume transports in the upper deep layer through the channel north of NG obtained from mooring observations (orange) and GL12 outputs (blue), respectively. (a) Integrated transport from the seafloor to various potential temperatures (θ) or depths, and (b) layered transport over a uniform θ interval of 0.04°C and a nonuniform depth interval during the mooring measurement period. (c) Time series of daily U-PMOC volume transports with θ ≤ 2.2°C (VTU-PMOC).
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
In summary, observations reveal two notable features of U-PMOC north of NG, including that 1) the U-PMOC velocity and transport have strong seasonality, and 2) the mean U-PMOC flows eastward mainly along the equator at 142°E with a time-mean transport of 0.76 Sv.
b. Reanalyzed U-PMOC
Given that four pointwise mooring measurements are insufficient to depict the whole picture of U-PMOC through the deep channel north of NG, we resort to the GL12 results. We evaluate the performance of GL12 in depicting the mean status and seasonal variability of U-PMOC north of NG by calculating the correlation (r) and normalized root-mean-square error (NRMSE) between GL12 and the observed data. NRMSE is defined as
Figure 1b shows the latitude–depth distribution of GL12 mean zonal velocity along 142°E during the mooring measurement period. The velocity core of U-PMOC just straddles the equator, as inferred from the mooring observation. The synthesis of observed and reanalyzed data suggests that the U-PMOC along its main route north of NG flows eastward on average. In another view, the velocity core of the U-PMOC can be regarded as a portion of the equatorial deep jets. A series of latitudinally alternating zonal jets is located north and south of the U-PMOC’s core, conforming to the nature of the equatorial jets.
The GL12 volume transports in the upper deep layer through the deep channel north of NG along 142°E are also estimated (Fig. 4, blue). The vertical profiles of GL12 integrated and layered transports generally follow those from observations (Figs. 4a,b). The time series of observed and GL12 VTU-PMOC show a fairly good agreement with NRMSE less than 15% and a correlation coefficient reaching 0.81 that is significant at the 0.01 level (Fig. 4c). Besides the above similarity, there also exist some differences. Notably, the observed integrated transport over 1.7°–2.2°C is larger than GL12. This is because mooring measurements do not cover the transport over 2°–7°N, the net value of which is negative (i.e., westward transport) in GL12. Such difference also leads to residual mean VTU-PMOC being 0.76 Sv in the observation and 0.20 Sv in GL12 during the period when the observed VTU-PMOC is available. Given that the difference may be induced by deficiency in observation, we can see that the GL12 performs well in depicting VTU-PMOC in general. Considering the GL12 data continuity, we use the GL12 outputs to obtain the statistical results of VTU-PMOC during 2014–19. Based on the annual harmonic analysis, the residual mean and seasonal amplitude are 2.19 and 10.5 Sv, respectively, and the seasonality can explain 42% of the total variance. For the estimation of residual mean VTU-PMOC, the error due to finite time length is 1.62 Sv, smaller than the residual mean value of 2.19 Sv. This suggests that the estimate of GL12 mean VTU-PMOC over 2014–19 is significant at the 0.05 level. The standard deviations are 11.4 and 8.65 Sv when including and excluding the seasonal cycle, respectively.
The general routes of the U-PMOC in the upstream and downstream regions of our observation sites are depicted by the GL12 streamfunction (Figs. 1a and 5a). The U-PMOC north of NG originates from the western boundary east of Mindanao Island, and flows eastward along the NG. When approaching the Solomon Rise, it bifurcates into a northern branch, directed toward the Melanesian Basin, and a southern branch with a smaller volume transport, directed toward the Solomon Basin through the Solomon Strait. The existence of the southern branch is consistent with the result inferred from shipboard CTD and L-ADCP measurements (Germineaud et al. 2021).
The horizontal distributions of mean streamfunction (color shading) and flow direction (arrows) of GL12 velocity at 2533-m depth over (a) all months, (c) January, and (d) July. (b) The horizontal distribution of the seasonal amplitude of the GL12 velocity field at 2533-m depth. All data are from GL12 outputs over 2014–19.
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
c. Dynamics for U-PMOC’s seasonality
The seasonal variability of Pacific equatorial zonal currents is related to the westward and downward propagation of Rossby waves (Kessler and McCreary 1993; Marin et al. 2010). Given that our observed U-PMOC sits in the deep layers of the far western basin, we examine the above mechanism by conducting an annual harmonic fit on the GL12 monthly full-depth zonal velocity along the equator over 2014–19. Figures 6a and 6b show the explained variance and phase for the annual harmonic. The region with large R2 is located west of 140°W with the depth range extending from upper 1000 m to the full depth. At depths below 2000 m east of Solomon Rise, the value of R2 is intensified possibly due to the scatter of Rossby waves. The phase isolines extend from the thermocline in the central and eastern Pacific to the upper deep layer of the western Pacific, well following the ray path of the first meridional mode (l = 1) Rossby waves (red dashed curves in Figs. 6a,b; Kessler and McCreary 1993). Our observed U-PMOC is located in the region with R2 > 40%, and the phase of the strongest eastward current occurs during January–February, consistent with the observed features shown in Fig. 2. The phase increases westward and upward, also conforming to the observed phase shifting earlier with depth shown in Fig. 2. Using the linear continuously stratified model, previous studies have suggested that the vertically and westward propagating Rossby waves are forced by the westward propagation of the annual equatorial zonal wind stress (e.g., Kessler and McCreary 1993; Wang et al. 2021a). The seasonal characteristics of Rossby waves are constrained by the corresponding features of zonal wind. Figures 6c and 6d show the explained variance and phase for the annual harmonic of zonal wind stress. The source of the ray path of Rossby waves in the thermocline corresponds to the equatorial region over 135°–90°W. In this region, the annual zonal wind stress explains more than 40% of the total variance (Fig. 6c), and the seasonal phase propagates westward (Fig. 6d). The apparent zonal wavelengths inferred from the seasonal phase distributions of Rossby waves and zonal wind stress are both ∼13 000 km. The above analyses suggest that the seasonality in the observed U-PMOC is induced by the vertically propagating annual Rossby waves driven by annual equatorial zonal wind stress.
The longitude–depth sections of the seasonal (a) explained variance and (b) phase based on the GL12 zonal velocity (U) along the equator, and the red dashed lines indicate the theoretical ray path of the first meridional mode (l = 1) Rossby waves. The horizontal distributions of the seasonal (c) explained variance and (d) phase based on the ECMWF-ERA5 zonal wind stress. The time period of data is 2014–19.
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
The U-PMOC in the upstream and downstream regions of our observation site also has an evident seasonal variability. The large values of seasonal amplitude larger than 6 cm s−1 are distributed along the western boundary over 127°–152°E (Fig. 5b). In contrast, the maximum mean zonal velocity of U-PMOC is less than 4 cm s−1 (Fig. 1b). The fact that the seasonal amplitude exceeds the mean transport value would result in the seasonal reversal of U-PMOC. The horizontal patterns of the U-PMOC streamfunction during the months when the maximum eastward and westward transports north of NG occur are shown in Figs. 5c and 5d, respectively.
d. Dynamics for mean U-PMOC
Considering the weakness of the Coriolis force near the equator, the pressure gradient force would be dominant in the transport north of NG. We thus select two stations symmetrically on both sides of the channel and examine their pressure difference (ΔP = PW − PE) between the western (PW, 0°, 134°E) and eastern (PE, 0°, 150°E) sites (cyan triangles in Fig. 1a). The pressure profiles are calculated according to the hydrostatic equation
Vertical profiles of mean (a) pressure ΔP (black), density Δρ (green), (b) potential temperature Δθ (orange), and salinity ΔS (blue) difference between the western and eastern sites (0°, 134°E and 0°, 150°E, cyan triangles in Fig. 1a). (c) Time–depth variations of daily pressure difference ΔP (color shading) and mean zonal velocity over 1°S–1°N at 142°E (U, contour). (d) Vertical profile of the correlations (r) with ΔP leading U by 0–120 days. The purple curve indicates the leading days corresponding to the maximum correlation, and the gray shading indicates the correlations that are not significant at the 0.01 level. All data are from GL12 outputs over 2014–19.
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
Given that ΔP varies within a small magnitude at depths deep than 500 m, we can deduce that differences in SSH and water mass density at depths above 500 m play a decisive role in the formation of the deep positive ΔP. The value of ΔP is negative (i.e., PW < PE) in the upper 150 m, but their difference ΔP shows an increasing trend with depth up to ∼200 m. This can be attributed to two facts. First, the study region north of NG is located on the western side of the western Pacific warm pool, and the weakening of the easterly wind here leads to a decreasing trend of SSH toward the west, resulting in negative ΔP on the surface (verified by satellite altimeter data). Second, compared to the eastern site, the western site has a farther distance to the climatological core region of the western Pacific warm pool, and therefore has a lower temperature and a higher density (Figs. 7a,b). This, on the one hand, is consistent with lower SSH at the western site due to the steric effect of surface water. On the other hand, the positive Δρ (ρW − ρE) leads to gradually increasing ΔP. For the steric effect and forming Δρ, the temperature makes a dominant contribution compared to salinity. At depths of 200–500 m, ΔP gradually decreases to a smaller positive value. This is related to the transport route of South Pacific Tropical Water (SPTW), which is characterized by higher temperature and salinity. The SPTW is transported northwestward across the equator by the New Guinea Coastal Undercurrent and arrives at the western site first. Then, the SPTW turns clockwise toward eastward guided by the Mindanao Eddy or New Guinea Eddy (Kashino et al. 2007, 2013), and arrives at the eastern site. This results in higher temperature and smaller density at the western site, and therefore weakens ΔP. At depths greater than 500 m, the difference in water mass density between the two sites is minor and has a weak effect on ΔP.
Figure 7c shows the time–depth variations of GL12 mean U over 1°S–1°N at 142°E and ΔP. At the depth of U-PMOC, their lead–lag correlation reaches a peak value of 0.75 when ΔP leads U by about 2 months (Fig. 7d). This suggests that the time deviation of velocity can be roughly balanced by the pressure gradient and friction term. Given the strong seasonality in U-PMOC, the leading time is supposed to be 3 months if the friction effect is not included. Although the eastern and western sites are selected based on the U-PMOC route, the high correlation between ΔP and zonal velocity with r > 0.60 can be established over the whole depth range. The corresponding phase lags are shortened to 1 month or less in the surface and near-bottom layers due to the significant influence of the friction term there.
4. Conclusions
Based on the synthesis of results from mooring measurements and a high-resolution oceanic reanalysis product, we reveal the pathway and volume transport of U-PMOC north of NG. In the mean status, the U-PMOC flows eastward out of the western Pacific with a mean transport of 2.19 Sv and an error bound of 1.62 Sv during 2014–19. On the basis of the mean eastward flow, there also exists strong seasonality in the U-PMOC velocity with a standard deviation of 11.4 Sv. This leads to the eastward and westward transports occurring during November–April and May–October, respectively.
The observed U-PMOC is closely linked to the wind forcing and the upper-ocean processes. The underlying mechanisms are summarized in Fig. 8. For the mean eastward transport, it is driven by the local deep pressure gradient over its upstream and downstream regions, which is mainly determined by the difference in their SSH and water mass properties in the upper 500 m layer. For the seasonality, it is related to vertically and westward propagating equatorial Rossby waves forced by the remote annual zonal wind. Such connections of deep ocean variations with upper-ocean processes have been reported in previous studies (e.g., Hughes et al. 2018; Jayne and Marotzke 2001; Rao and Tandon 2021; Wang et al. 2021a). Besides seasonal variability, interannual variability is also observed in the volume transport of U-PMOC (Fig. 4c). The characteristics and underlying mechanisms of interannual variability deserve further studies. Furthermore, the linkage of U-PMOC transport with climate change and biogeochemical cycling processes deserves attention.
Schematic diagram of the dynamic mechanisms for the mean (yellow) and seasonal (green) characteristics of U-PMOC north of NG.
Citation: Journal of Physical Oceanography 53, 5; 10.1175/JPO-D-22-0180.1
Acknowledgments.
The authors thank the editor and two anonymous reviewers for the very insightful comments that greatly helped us improve the quality of this manuscript. J. Wang thanks the support from the Strategic Priority Research Program of the Chinese Academy of Sciences (Grant XDA22000000), the National Natural Science Foundation of China (NSFC, Grants 91958204 and 42222602), the Science and Technology Innovation Project of Laoshan Laboratory (LSKJ202203100 and LSKJ202202700), the Youth Innovation Promotion Association CAS, and TS Scholar Program. F. Wang thanks the support from the NSFC (Grants 42221005 and 42090040). Z. Zhang thanks the support by the NSFC (Grant 42106009). Q. Ma thanks the support by the NSFC (Grant 42006003).
Data availability statement.
The mooring data analyzed in this paper are available for download at https://doi.org/10.5281/zenodo.7024824. The GL12 product is available on the Copernicus Marine Environment Monitoring Service website (https://resources.marine.copernicus.eu/product-detail/GLOBAL_MULTIYEAR_PHY_001_030/DATA-ACCESS). The ECMWF-ERA5 wind stress data are available on the Copernicus Climate Change Service website (https://cds.climate.copernicus.eu/cdsapp#!/dataset/reanalysis-era5-single-levels-monthly-means?tab=form).
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