1. Background and introduction
The Mindanao Current (MC) is a low-latitude western boundary current (LLWBC) in the western tropical North Pacific, formed by the bifurcation of the North Equatorial Current (NEC) in the Philippine Sea (Fig. 1). The MC flows equatorward along the coast of Mindanao until splitting to contribute to the global overturning circulation through the Indonesian Throughflow (ITF; Gordon 1986) and to close the interior North Pacific Sverdrup transport via the North Equatorial Counter Current (NECC; Wyrtki 1961). The MC is the northern source of subtropical water in the western Pacific warm pool (Fine et al. 1994; Lukas et al. 1996) and thus impacts the tropical heat and freshwater budgets and the El Niño–Southern Oscillation (ENSO) phenomena (Gu and Philander 1997; Zhang et al. 1998). In the subthermocline, the MC transports North Pacific Intermediate Waters equatorward while a Mindanao Undercurrent (MUC; Hu et al. 1991) flows poleward. The confluence of subtropical, tropical, and intermediate water masses of North and South Pacific origin and their role in salt and heat exchange of the tropical Pacific and Indian Oceans has led to several initiatives in the last two decades to observe and model the MC/MUC system. Observations of the MC/MUC include sea level gauges (Lukas 1988), hydrographic surveys (Hacker et al. 1989; Toole et al. 1990; Hu et al. 1991; Lukas et al. 1991; Wijffels et al. 1995; Kashino et al. 1996; Qu et al. 1998; Firing et al. 2005; Kashino et al. 2009, 2013), moorings (Kashino 2005; Zhang et al. 2014; Kashino et al. 2015; Hu et al. 2016), and Argo floats (Qiu et al. 2015; Wang et al. 2015). However, MC/MUC water mass transport, coastal structure, and variability are not well understood owing to strong currents, eddy activity, and large annual and interannual changes in surface forcing.
(a) Schematic of circulation in the tropical northwestern Pacific. Red is circulation from the surface to the bottom of the thermocline, typically 300-m depth. Main features are the NEC, MC, the Kuroshio, ITF, NECC, and ME. Subthermocline currents (blue) that oppose the direction of thermocline flow are the LUC, MUC, and NEUCs. (b) MDT from AVISO with contours at 0.015-m intervals. Glider observations are objectively mapped onto a mean line from 8.15°N, 126.61°E to 8.77°N, 130°E (thick black line). The mean line is perpendicular to coastal bathymetry and MDT near the coast.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
From hydrographic surveys, the MC current is narrow at 10°N, then widens, shoals, and increases in speed as it flows equatorward until it separates into the ITF and NECC between 7°N and 6°N (Lukas et al. 1991). Upper-ocean cyclonic recirculation in the Philippine Sea, such as that associated with the quasi-permanent Mindanao Eddy (ME), centered near 7°N, 130°E, causes the MC transport to increase as it flows equatorward (Lukas et al. 1991). At 8°N, the mean MC is in geostrophic balance, extending approximately 200 km offshore, with a subsurface velocity core of 0.8 to 1 m s−1 (Wijffels et al. 1995; Qu et al. 1998). Although the ME has been observed by surface drifters (Lukas et al. 1991), its sporadic presence is difficult to observe in hydrographic transects (Wijffels et al. 1995; Qu et al. 1998; Firing et al. 2005; Kashino et al. 2013). The velocity variability of the MC, its structure near the coast, and the variability of the ME remain inadequately observed.
In the subthermocline, the poleward MUC is roughly 50 to 80 km offshore of the coast of Mindanao near 500-m depth and potential density of 27.2 kg m−3 (Hu et al. 1991; Lukas et al. 1991; Qu et al. 1998), with a second velocity core occasionally reported farther offshore (Hu et al. 1991; Qu et al. 1998). The MUC could be a quasi-permanent recirculation or a poleward current transporting South Pacific water masses; intermittent subthermocline eddies that populate the region on intraseasonal time scales make previous observations inconclusive (Hu et al. 1991; Qu et al. 1998, 1999; Firing et al. 2005; Dutrieux 2009; Chiang and Qu 2013; Wang et al. 2014; Zhang et al. 2014; Chiang et al. 2015; Kashino et al. 2015; Schönau et al. 2015; Hu et al. 2016). The forcing, persistence, and connectivity of the MUC, and its relationship to the thermocline circulation are areas of ongoing research.
The transport of the MC/MUC forms a closed mass system with the NEC and Kuroshio Current (KC), and their respective undercurrents, the North Equatorial Undercurrents (NEUCs) and the Luzon Undercurrent (LUC) (Toole et al. 1990; Qiu and Lukas 1996; Qu et al. 1998). The relative transports of the MC and KC depend on the bifurcation latitude of the wind-driven NEC. The bifurcation latitude, dictated by Sverdrup theory as the zero line of zonally integrated wind stress curl, is affected by local Ekman pumping and suction caused by shifts of the East Asian monsoon and the arrival of annual Rossby waves from the central Pacific (Qiu and Lukas 1996; Kim et al. 2004; Qiu and Chen 2010). Interannual variability of the MC is forced by basinwide changes during ENSO. Numerical models suggest an increase in MC transport when the NEC bifurcates at higher latitude: annually in the fall and interannually during El Niño (Qiu and Lukas 1996; Kim et al. 2004; Qiu and Chen 2010). Observations have both verified (Lukas 1988; Kashino et al. 2009) and conflicted (Toole et al. 1990) as to the relationship of transport with ENSO. Less is known about the transport variability of the MUC.
Transport variability of the MC/MUC impacts the heat and salt balance in the tropical Pacific. Water masses transported equatorward by the MC are the North Pacific Tropical Water (NPTW), a subsurface salinity maximum in the thermocline, and the North Pacific Intermediate Water (NPIW), a salinity minimum in the subthermocline (Gordon 1986; Bingham and Lukas 1994; Fine et al. 1994; Kashino et al. 1996; Schönau et al. 2015; Wang et al. 2015). Antarctic Intermediate Water (AAIW), saltier than NPIW and of South Pacific origin, is also present in the Philippine Sea and may be carried northward by the MUC (Fine et al. 1994; Wijffels et al. 1995; Qu et al. 1999; Qu and Lindstrom 2004; Schönau et al. 2015; Wang et al. 2015). The volume transport and modification of these water masses as they transit the Philippine Sea are less well understood. Isopycnal stirring of large-scale salinity gradients introduced by subsurface water masses creates finescale salinity variance that leads to isopycnal and diapycnal diffusion (Ferrari and Polzin 2005). Quantifying the relationship between stirring, which increases isopycnal salinity variance, and diffusion, which decreases it, will give insight into heat and salt transfer in the ocean interior and improve parameterizations of mixing in models.
Repeat glider observations of the MC/MUC provide mean water mass transport and variability during a 4-yr period. Spray gliders made 16 sections of the MC/MUC from September 2009 to October 2013 under the Origins of Kuroshio and Mindanao Current (OKMC) initiative (Rudnick et al. 2015b). Gliders observed temperature, salinity, and depth-averaged velocity from the surface to 1000-m depth at horizontal length scales from 6 to 2000 km during all seasons and one ENSO cycle. Objective maps of geostrophic velocity, salinity, and potential temperature resolve the mean structure and variability of the MC, the MUC, and the ME. Argo climatology is compared to glider observations to assess annual and interannual variability. The thermohaline structure is compared between the MC/MUC and the NEC (Schönau and Rudnick 2015) to infer water mass connectivity and modification. Gliders provide new observations on the stability of the MC thermocline transport and the relationship between the variable MC/MUC subthermocline transport and ENSO. The paper proceeds with a description of data and analysis methods, the mean and variability of water masses, geostrophic velocity and transport, and finescale thermohaline structure.
2. Data collection and analysis
To observe the MC/MUC, the autonomous underwater glider Spray (Sherman et al. 2001; Rudnick et al. 2004) made repeat transects from Palau to the coast of Mindanao. The Spray glider is equipped with a Sea-Bird CTD, Seapoint fluorometer, and is remotely controlled with Iridium satellite. The glider completes a dive from the surface to 1000 m and back over a horizontal distance of 6 km in a cycle time of 6 h. Data are collected during ascent at a vertical resolution of about 1 m and binned into 10-m vertical bins. The depth-averaged velocity of the glider is obtained from dead reckoning (Davis et al. 2008; Todd et al. 2011; Davis et al. 2012).
For each mission the glider profiled westward from Palau (7.5°N, 134.5°E) until reaching the 1000-m isobath along the coast of Mindanao at a latitude between 7° and 10°N and then profiled on the return to Palau (Fig. 2a). There were 10 missions, with 16 sections of the MC/MUC completed within the given latitude range (Fig. 2b). The temporal coverage of the glider was sufficient to observe the MC at all times of the year, with 10 sections from September to February and 6 sections from March to August (Fig. 2b). The ENSO phase, as measured by the Niño-3.4 index, was positive (El Niño) from September 2009 to April 2010 and negative (La Niña) from June 2010 to April 2011 and August 2011 to March 2012.
(a) Trajectories of the 10 gliders deployed from Palau to the 1000-m isobath along the coast of Mindanao. Arrows show the mean depth-averaged velocity from the objective map of glider profiles within 80 km of the mean line (dashed box). There are 968 profiles included. (b) Glider missions are coded by date and glider serial number. The first two digits are the abbreviated year, the third digit is the hexadecimal month, and the fourth and fifth digits are the glider serial number. The dashed box denotes a 220-km range where observations from each mission are linearly projected and objectively mapped onto the mean line (solid black) to assess variability between missions. Most missions included observations north and south of the mean line. Mission 11A03601 (thick magenta) is featured as an example in Figs. 3 and 10. (c) Annual coverage of glider missions colored by the Niño-3.4 index. Observations are at all times of the year and ENSO phases.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
An example section from the return of mission 11A03601 is shown in Fig. 3. Salinity with potential temperature contours are plotted on depth and density surfaces as a function of along-track distance from the nearest position to the coast at 9.3°N, 126.31°E to roughly 130°E. On depth surfaces, internal wave motion causes waviness of isotherms (Fig. 3a; Rudnick and Cole 2011). To filter out this high-frequency noise, salinity, potential density, and potential temperature are objectively mapped on depth surfaces (Fig. 3b). Alternatively, to maintain the horizontal resolution of 6 km, interpolating data to density surfaces also filters out internal wave motion (Fig. 3c).
Example section from glider mission 11A03601 during its return from near the coast of Mindanao to Palau (see Fig. 2a). Along-track salinity (color) and temperature (contours) from 9.3°N, 126.31°E to 8.22°N, 130°E: (a) vertically binned on depth surfaces, (b) objectively mapped onto depth surfaces, and (c) interpolated to potential density surfaces. Tick marks in (a) are at the location of each profile. (d) Geostrophic velocity referenced to the depth-averaged velocity (color) and salinity (contours) from an objective map of mission 11A03601 onto the mean line. Dark gray lines denote 26 and 27 kg m−3, respectively. The range is extended to 9.15°N, 134°E to show the relatively weak velocities east of 130°E, denoted by the dashed line. Negative (positive) velocity is equatorward (poleward). (e) Objectively mapped depth-averaged velocity for mission 11A03601.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
To make a mean section, glider observations are linearly projected and objectively mapped onto a mean line extending from the coast of Mindanao, from 8.15°N, 126.61°E to 8.77°N, 130°E (Figs. 1b, 2a), hereinafter referred to as the mean line. Only observations within 80 km of the mean line are included in the map. The objective map is made using a Gaussian autocovariance (Bretherton et al. 1976) with a length scale of 80 km. This length scale is based on the autocovariance calculated from observations of temperature and salinity and is identical to that used in Schönau and Rudnick (2015). Geostrophic velocity, calculated from thermal wind equations, is referenced to the depth-averaged velocity of the glider (Rudnick et al. 2004; Todd et al. 2011; Pelland et al. 2013; Lien et al. 2014; Pietri et al. 2014). The absolute geostrophic velocity from gliders compares favorably to ADCP measurements from a mooring (Lien et al. 2014), satellite sea surface height (Rudnick et al. 2015a), and to Argo climatology referenced to a level of no motion at 2000 m in the tropical Pacific (Schönau and Rudnick 2015).
The mean line extends from the 1000-m isobath approximately 380 km offshore and was chosen at an angle 80°E of north to be perpendicular to isobaths and contours of mean dynamic topography (MDT; Fig. 1b; AVISO 2014). The mean depth-averaged velocity is perpendicular to the mean line near the coast where the current is strongest (Fig. 2a), supporting the choice of the line’s direction. Because of the curvature of the coast, 80 km is the cutoff for inclusion of profiles in the objective map (Fig. 2a). The inshore coordinates of the mean line were chosen to maximize the number profiles (968 profiles). Argo climatology (⅙°; Scripps Institution of Oceanography and International Argo Program 2009; Roemmich and Gilson 2009) indicates that there is alongshore shoaling of the thermocline of 20 m within this 80-km range but that it is less than the annual thermocline displacement (Scripps Institution of Oceanography and International Argo Program 2009; Roemmich and Gilson 2009). There were roughly an equal number of profiles north and south of the mean line to accommodate for the spatial variability. Removing a linear alongshore gradient to account for shoaling does not affect the results of the objective map. The objectively mapped mean is thus a spatial and temporal average.
The variability of the MC/MUC is assessed from the standard deviation between the 10 glider missions (Fig. 2b), each objectively mapped onto the mean line. Linearly projecting and mapping each glider mission together averages out the linear alongshore gradient, as gliders tended to traverse north then south of the mean line. On four of the missions, the glider trajectory exceeded 220 km from the mean line, and only the nearer transect was included in the map. For these single transects, alongshore thermocline depth variability was less than 10 m. Since strong currents and meanders influence the glider path, excluding observations because of distance from the mean line would exclude temporal variability. The standard deviation between glider missions thus reflects spatial and temporal variability. However, comparisons to Argo climatology in section 4 and to mooring results in the conclusion indicate that temporal variance exceeds spatial variability within the region of observations.
An example of geostrophic velocity and salinity objectively mapped on the mean line is shown for mission 11A03601 in Fig. 3d. The MC flows equatorward near the coast (blue), and there are two subthermocline poleward (red) cores offshore that may be expressions of the MUC. The range is extended to 134°E to show the small velocities east of the mean line, although these may vary by section. The objectively mapped glider missions are used to make composite potential temperature–salinity (T–S) diagrams to assess water mass transport, and they are compared to monthly (1°) and annual (½°) Argo climatology (Scripps Institution of Oceanography and International Argo Program 2009; Roemmich and Gilson 2009) and forcing from wind and precipitation to assess annual and interannual variability. Monthly 10-m wind products are available from NCEP–NCAR reanalysis (Kalnay et al. 1996; NCEP 1996) and monthly precipitation data are available from CPC Merged Analysis of Precipitation (CMAP; CPC 1997; Xie and Arkin 1997), each gridded at 2.5° and available from the Earth System Research Laboratory (ESRL).
Isopycnal salinity variance, obtained using a wavelet transform, is used to examine thermohaline structure of the MC/MUC at different length scales and can be used to trace water masses (Todd et al. 2012; Schönau and Rudnick 2015). Following Todd et al. (2012), we use a Morlet wavelet 
3. Mean structure
a. Water masses
Water masses are identified by their salinity extrema. The mean salinity has two haloclines (Fig. 4a): first, from the surface fresh layer that overlies the mean section to a subducted subtropical salinity maximum and second, in the transition to fresh intermediate water below the thermocline. The subsurface salinity extrema are greatest near the coast as displayed in a T–S diagram (Fig. 5a). The surface fresh layer that overlies the section is North Pacific Tropical Surface Water (NPTSW; <34.1 psu and >28°C; Wyrtki and Kilonsky 1984), formed by local precipitation and advection (Delcroix and Hénin 1991). NPTSW is influenced by annual variability of the East Asian monsoon and interannual variability associated with ENSO (Delcroix and Hénin 1991; Li et al. 2013). NPTSW creates a large vertical salinity gradient within the upper 50 m that strongly impacts the western Pacific warm/fresh pool (Cronin and McPhaden 1998; Delcroix and McPhaden 2002). Beneath NPTSW is the subsurface salinity maximum of NPTW (>34.95 psu) near the coast and centered at potential density 23.5 kg m−3. NPTW is formed in the subtropics due to excess evaporation (Tsuchiya 1968; Katsura et al. 2013) and advected westward in the NEC and into the MC, freshening as it mixes with surface and intermediate waters along the path of flow (Lukas et al. 1991; Fine et al. 1994; Qu et al. 1999; Li and Wang 2012; Schönau and Rudnick 2015). In the subthermocline, the subsurface salinity minimum of NPIW (<34.4 psu), formed in the Okhotsk Sea (Talley 1993), is at potential density of 26.55 kg m−3 in the MC (Bingham and Lukas 1994). Like NPTW, the extrema of NPIW is near the coast; however, its source may be either from the NEC or more directly from the equatorward LUC beneath the KC (Bingham and Lukas 1994; Fine et al. 1994; Qu et al. 1997; Schönau et al. 2015). AAIW, with potential density of 27.2 kg m−3 and salinity greater than 34.5 psu, is denser and saltier than NPIW (Fig. 5a) and is commonly identified by its higher oxygen (Qu and Lindstrom 2004). The presence of AAIW and its pathway into the Philippine Sea has been recorded by several observational studies (Fine et al. 1994; Wijffels et al. 1995; Qu et al. 1999; Qu and Lindstrom 2004; Zenk et al. 2005), and a similar T–S curve for AAIW is observed here. Lower oxygen intermediate waters that have mixed in the tropics, such as the North Pacific Tropical Intermediate Water (NPTIW), are found east of 130°E (Bingham and Lukas 1995). The poleward transport of AAIW has been used to verify some connectivity between the MUC and southern North Equatorial Undercurrent (Schönau et al. 2015; Wang et al. 2015).
Mean structure of the MC/MUC from an objective map of glider profiles within 80 km of the mean line: (a) salinity with potential temperature contours, (b) geostrophic velocity referenced to the depth-averaged velocity on depth surfaces with potential density contours (gray), and (c) geostrophic velocity referenced to the depth-averaged velocity interpolated to potential density surfaces with salinity contours. In (b) and (c), equatorward velocity is negative (blue), and poleward velocity is positive (red). Subsurface salinity extrema are nearest to the coast and advected equatorward in the MC.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
Potential T–S diagrams colored by (a) mean longitude and (b) net transport over the 10 glider missions, objectively mapped onto the mean line and binned by 0.2°C and 0.01 psu. Water masses are the NPTSW, NPTW, NPIW, and AAIW. In (b) transport is summed in each bin then normalized by the bin size and number of sections. Equatorward transport is negative (blue) and poleward transport is positive (red). Water masses that had small net transport from recirculation or were measured infrequently appear as white. The MC has a net equatorward transport of NPTW and NPIW, and the MUC has a net poleward transport of water typical of AAIW.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
b. Geostrophic velocity
The mean absolute geostrophic velocity of the MC, referenced to the depth-averaged velocity, is equatorward, with a subsurface velocity maximum of −1.0 m s−1 at 100-m depth (22 kg m−3) and 50 km offshore (127°E; Figs. 4b,c). Near the coast, where the MC is the strongest, there is advection of NPTW and NPIW (Fig. 4c). The MC decreases in strength with increasing distance offshore and intersects at the surface with what could be a cyclonic ME roughly 250 km offshore (128.75°E). The core of the possible ME extends through the thermocline to 250-m depth, with similar poleward and equatorward velocities and symmetric isohalines (Fig. 4c). The mean depth-averaged velocity is cyclonic over this longitude range (Fig. 2a). Mean hydrographic sections from Wijffels et al. (1995) and Qu et al. (1998) have similar velocity structures as the glider-observed mean MC, although neither of these studies found evidence of a mean ME. The width of the possible ME is roughly 250 km, consistent with that found by surface drifters at this location (Lukas et al. 1991).
In the subthermocline, the equatorward MC extends to 1000 m near the coast, decreasing in velocity with increasing depth and distance offshore (Fig. 4b). Offshore of the equatorward subthermocline MC is a poleward MUC. The MUC has a maximum velocity of 0.17 m s−1 located 75 km offshore (127.75°E) and 650-m depth (27.2 kg m−3).
East of the MUC there is weak equatorward flow until 128.5°E; then mean poleward flow exists through the offshore edge of the mean line. The coastal structure of the equatorward MC and mean location and width of the MUC referenced to an absolute velocity are a fundamental improvement to previous observations, which have either been time series at single locations or referenced to arbitrary levels of no motion. The mean MC/MUC compares favorably with numerical results from Estimating the Circulation and Climate of the Ocean (ECCO) and the Parallel Ocean Program (POP) simulation (Schönau et al. 2015).
c. Transport
The transport of the MC/MUC across the mean line and from the surface to 27.3 kg m−3 is −22.4 Sv (1 Sv = 106 m3 s−1), with a thermocline transport (surface to 26 kg m−3) of −19 Sv and subthermocline transport (26 to 27.3 kg m−3) of −3 Sv. Potential density 27.3 kg m−3 is the densest isopycnal that is present in all sections. Potential density 26 kg m−3, at the base of the thermocline, divides the thermocline and subthermocline transport, which encompasses the intermediate water and the MUC (Figs. 4b,c). This isopycnal has been used as a reasonable separation of thermocline and subthermocline transports for the NEC (Schönau and Rudnick 2015) and NECC (Johnson et al. 2002; Hsin and Qiu 2012). The Sverdrup transport from the zonally averaged wind stress curl was −28.3 Sv during this time period, and hydrographic estimates of the mean transport from the surface to 1000 m and coast to 130°E are around −27 Sv (Wijffels et al. 1995; Qu et al. 1998). These estimates exceed the glider transport as they do not consider or resolve the poleward undercurrent. To compare, glider transport summed from the coast approaches −30 Sv before decreasing offshore (Fig. 6a). The sum of equatorward transport from the coast to 128.5°E and surface to 27.3 kg m−3 is −30.4 Sv. In the subthermocline, the poleward transport of the inshore MUC (127.25°E) is 3.6 Sv, and the poleward subthermocline transport across the mean line is 8.9 Sv.
(a) Total transport (black) summed along the mean line and separated into thermocline (blue; surface to 26 kg m−3) and subthermocline (red; 26 to 27.3 kg m−3) transports. Shading is the standard deviation between the 10 missions. (b) Transport per distance for the thermocline (blue), subthermocline (red), and total (black) as a function of longitude. (c) Transport per density along the mean line. For (b) and (c), shading is the standard error, and the legend gives the mean and standard deviation of each layer.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
As a function of longitude, mean transport is greatest in the thermocline at the location of the MC core and then decreases to zero at 128.5°E (Fig. 6b). The possible ME, centered at 129.25°E, has a cyclonic circulation of roughly 2 Sv, with 0.5 Sv greater transport poleward than equatorward. The subthermocline transport has an equatorward maximum near the coast, consistent with the subthermocline equatorward MC, and then becomes poleward between 127° and 127.75°E at the mean location of the MUC. As a function of density, mean transport is greatest at the surface between potential density 21 and 22 kg m−3 and uniform through the thermocline. Deeper than 27 kg m−3, the large standard error indicates that the direction of mean net transport is inconclusive (Fig. 6c).
Binning the volume transport of each objectively mapped mission by potential temperature and salinity, summing over all sections, and normalizing by bin size and the number of sections creates a T–S diagram of water mass transport (Fig. 5b). The T–S transport integrates to the average transport over the 10 missions from the surface to 1000-m depth. Both NPTW and NPIW have equatorward transport. In the subthermocline, there is a net poleward transport of water with potential density 27 to 27.5 kg m−3 and salinity greater than 34.5 psu, typical of AAIW. The poleward subthermocline transport shallower than 27.3 kg m−3 is 5 Sv, earning the MUC the designation as a current, as there is a mean net transport of a distinct water mass.
4. Variability
Variability in the MC/MUC is forced by basinwide wind stress, transient eddies, Rossby waves, annual changes in the East Asian monsoon, and interannual changes associated with ENSO. The forcing creates water mass and geostrophic velocity variability between glider missions. Variability is greatest in thermocline depth, surface salinity, and subthermocline transport. The following sections assess the magnitude of variability and the responsible forcing.
a. Thermocline depth
The standard deviation in potential temperature is greatest in the thermocline, between depths of 100 to 200 m, where the vertical temperature gradient is also a maximum (Fig. 7a), making heaving of isopycnals the likely cause. Below the temperature mixed layer, which extends to 80 m, isotherms tend to follow isopycnals, causing isopycnal displacement to appear as temperature variability. The thermocline depth (measured as the depth of the layer 23–24 kg m−3) is at a minimum in March 2010, coinciding with El Niño, and at a maximum in 2013 (Fig. 8a). Here, a positive anomaly is the shoaling of the thermocline, whereas a negative anomaly is the deepening of the thermocline.
Standard deviation of the objectively mapped glider missions (color) with mean contours (black) for each (a) potential temperature, (b) salinity, and (c) geostrophic velocity. Potential density contours (white) are from (a) 22–27 kg m−3 at 1 kg m−3 increments and at (c) 26 kg m−3. The standard deviation includes both temporal and spatial variability over the range of the glider missions in Fig. 2b.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
Observations and model of thermocline depth variability. For each, the depth anomaly is positive for shoaling and negative for deepening. (a) Depth anomaly of potential density layer 23–24 kg m−3 for each glider mission. Shading on date axis indicates observations during El Niño (red) and La Niña (blue) according to the Niño-3.4 index. (b) Annual cycle of 21°C isotherm depth from Argo climatology (Roemmich and Gilson 2009) along 8.5°N. (c) Depth anomaly of the 21°C isotherm for each a linear, 1.5-layer, wind-forced model that allows Rossby wave propagation, monthly Argo climatology, and glider observations, all at 8.5°N, 128.5°E. The mean from June 2009 to October 2013 has been removed from each but not an annual cycle. The linear model and Argo climatology were smoothed with a 3-month running mean. The correlation coefficient is 0.6 between the glider and Argo climatology, 0.59 between the glider and model, and 0.52 between the model and Argo climatology.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
Annual changes in thermocline depth are forced by the local wind and by Rossby waves generated by wind over the central Pacific (Kim et al. 2004). During the winter, strong northeasterlies over the region create positive wind stress curl and Ekman suction that shoals the thermocline, leading to a cold layer at 100-m depth known as the Mindanao Dome (MD) that extends east to around 150°E between the MC, NEC, and NECC (Masumoto and Yamagata 1991). The MD decays in May as the East Asian monsoon shifts to southwesterlies and the annual downwelling Rossby wave arrives from the central Pacific (Tozuka et al. 2002). Annual thermocline depth variability, taken as the depth of the 21°C isotherm from Argo climatology at 8.5°N is consistent with this forcing (Fig. 8b). At 128.5°E, the minimum depth is from January to April and the maximum is from May to November. The annual Rossby wave arrives at this longitude in late summer.
The model, Argo climatology, and glider results are compared at 8.5°N and 128.5°E, where the glider and Argo climatology have the same mean thermocline depth (Fig. 8c). Nearer to the coast, the glider better resolves the coastal structure and has a deeper thermocline than the climatology. The model and Argo climatology have roughly the same magnitude but differ in phase. The model has a stronger annual cycle than Argo climatology (Fig. 8c) but weaker interannual variability. The glider has the greatest extremes in shoaling and deepening as is expected of synoptic measurements when compared to gridded climatology. Both Argo climatology and gliders observed shoaling of the thermocline during the El Niño in 2009/10 and deepening in 2013. The shoaling in 2009/10 is consistent with forcing during El Niño, but the deepening of the thermocline in 2013 is of unknown origin.
b. Salinity
The standard deviation in salinity is greatest in the upper 100 m (>0.2 psu), with a secondary variance maximum in the subsurface halocline between the NPTW and NPIW (Fig. 7b). Where the vertical salinity gradient is large, isopycnal heaving is sufficient to account for the observed salinity variability, as calculated from the depth variance of isopycnals and their mean salinities. The vertical minimum in salinity variance coincides with the low vertical salinity gradient between the two haloclines.
The surface salinity is saltiest during the 2010/11 El Niño and freshest at the end of the 2010/11 La Niña, with low variability in intermittent years (Fig. 9a).
Sea surface salinity and precipitation variability. (a) Salinity anomaly (0–50 m) from each glider mission. (b) Annual anomaly of precipitation and surface salinity (0–50 m) from Argo climatology (Roemmich and Gilson 2009), each with a 3-month running mean and averaged from 2004 to 2014. The Argo anomaly is at 8.5°N, 127.5°E, and precipitation is averaged over the region 3.75°–13.75°N and 126.25°–136.25°E, encompassing rainfall over the NEC, MC, and local recirculation from the NECC. (c) Niño-3.4 index and interannual anomalies of precipitation and glider surface salinity (0–50 m, 127.5°E), where the glider mean and annual Argo anomaly are removed from glider observations. Precipitation is averaged over the same region as in (b) and filtered with a 5-month running mean. The maximum, lagged correlation coefficient between the Niño-3.4 index and the salinity anomaly is 0.8 and that between the Niño-3.4 index and precipitation anomaly is −0.75, with the Niño-3.4 index leading the salinity anomaly and precipitation by 1 month. The precipitation anomaly and salinity anomaly have a maximum correlation of −0.65 with a zero time lag. Results depend on the area over which precipitation is averaged.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
The annual cycle in surface salinity is less than 0.1 psu (Argo climatology), fresh from August to February and salty from July to August (Fig. 9b). As evaporation tends to be weaker than precipitation in this region (Cronin and McPhaden 1998), the surface salinity is compared to the regional precipitation (Fig. 9b). A positive (negative) salinity anomaly lags anomalously low (high) regional precipitation, consistent with the annual migration of the ITCZ.
The relatively small magnitude of the annual cycle suggests that glider-observed extremes are related to interannual precipitation forcing. Removing the annual cycle, interannual anomalies of precipitation and surface salinity are each large following El Niño and La Niña. The precipitation anomaly lags the Niño-3.4 index: negative following the 2009/10 El Niño and positive following the 2010/11 La Niña (Fig. 9c). The interannual anomaly from the glider (where the annual cycle from Argo climatology has been removed) is correspondingly positive during lack of rainfall and negative with excessive rainfall. The magnitude of the interannual surface salinity anomaly is roughly 5 times greater than that of the annual anomaly. During ENSO neutral conditions, when the Niño-3.4 index indicates neither an El Niño nor La Niña state, interannual anomalies of precipitation and surface salinity are small by comparison.
Although these results assess the phase of surface salinity anomalies to precipitation, they are not entirely independent of mixing and advection caused by wind and wind stress curl. Annually, precipitation is in phase with wind forcing (Masumoto and Yamagata 1991). In winter and early spring the ITCZ is south of the region, precipitation is small, and positive wind stress curl forces upwelling. In the summer, the ITCZ creates large precipitation over the region, wind stress curl changes, and the annual Rossby wave arrives to force downwelling (Fig. 8b). Interannually, low precipitation and upwelling each have extremes during the El Niño in 2009/10 (Figs. 8c, 9c). Thus, salty anomalies from “outcropping” of denser, saltier water occur when a fresh layer is advected away or from upwelling and mixing with underlying isopycnals. The relationship between surface salinity, upwelling, and advection thus requires a more complete salinity budget than presented here.
c. Geostrophic velocity and transport
The MC/MUC geostrophic velocity and transport variability are assessed by the standard deviation of geostrophic velocity (Fig. 7c) and transport as a function of longitude (Fig. 10) and time (Fig. 11). The MC core velocity and transport are stable in the thermocline, whereas the subthermocline has a large transport range. Interannual transport variability is more pronounced than annual variability, and a relationship between the subthermocline transport and ENSO emerges.
Transport per distance for each glider mission summed from (a) the surface to 26 kg m−3 and (b) 26 to 27.3 kg m−3, where 26 kg m−3 separates the thermocline and subthermocline transports. Note the magnitude of the color bars. Black bars in (b) denote glider mission 11A03601, used in (c)–(e). (c) T–S transport diagram summed along the mean line from the coast to 128°E to encompass the equatorward MC and first poleward MUC core; (d) as in (c), but from the coast to 129°E to encompass the equatorward MC, first poleward MUC core, and offshore equatorward flow; and (e) as in (c), but from the coast to 130°E. See Fig. 3 for a geostrophic cross section and depth-averaged velocity of mission 11A03601. See Fig. 5b for bin size and normalization of (c)–(e).
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
Transport through the mean line for each glider mission from the surface to 27.3 kg m−3 (black) and divided into thermocline (blue) and subthermocline (red) transport by potential density 26 kg m−3. Transports for each are shown as a function of (a) time, (b) day of year, and (c) Niño-3.4 index. The thermocline transport is relatively constant compared to the subthermocline transport, which has an inverse linear relationship with the Niño-3.4 index.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
The standard deviation of geostrophic velocity is large where the mean current velocity is strongest, approaching 50% of the mean velocity near the core of the MC and also large in the subthermocline by the coast (Fig. 7c). However, the standard deviation is less than the mean equatorward MC in both the thermocline and subthermocline. Near the core of the ME (129.25°E), the variability is roughly equal to the mean, suggesting the ME was intermittently observed. There is a vertical minimum in standard deviation near potential density 26 kg m−3 that separates the thermocline and subthermocline transport, occurring even near the coast where the current is strong. In the subthermocline, the standard deviation is large on either side of the MUC, typical of changes in location and breadth. The velocity variability in the subthermocline east of the MUC exceeds the mean.
Velocity is integrated over density surfaces to yield thermocline and subthermocline transport for each glider mission (Fig. 10). In the thermocline, transport is greatest near the coast, tapering to zero between 127.5° and 129.5°E, and becomes poleward in roughly half the sections (Fig. 10a). The poleward transport may be an expression of the ME as it is at the same location of large variability near the ME core (Fig. 7c). The transport in the subthermocline is variable: persistently equatorward near the coast with poleward transport in all sections (Fig. 10b). Roughly half the sections have multiple poleward cores that previously have been reported as double MUC cores (Hu et al. 1991). However, it is difficult to define multiple cores of the MUC, as it is unclear if each has net water mass transport or recirculates. For example, mission 11A03601 has symmetric isohalines in the subthermocline between poleward and equatorward velocities, centered at 129°E, suggesting water mass recirculation (Fig. 3d). T–S transport diagrams are used to assess the net transport of water masses (Figs. 10c,d,e). Subthermocline transport summed from the coast to 128°E has equatorward transport of NPIW and a poleward transport of water typical of AAIW (>34.5 psu, 27.2 kg m−3; Fig. 10c). Summing transport from the coast to 129°E (Fig. 10d) now includes saltier equatorward subthermocline transport (~27 kg m−3, >34.5 psu). However, this T–S range has net poleward transport when integrating to 130°E (Fig. 10e). Thus, the subthermocline equatorward flow between 128° and 129°E (Fig. 10b) is likely a partial cyclonic circulation from farther offshore, as confirmed by cyclonic depth-averaged velocity centered at 129°E (Fig. 3e). Such recirculation, typical of eddies, occurs during other glider missions. However, all glider missions had a net poleward subthermocline transport of intermediate water that is saltier than NPIW and typical of that found in the mean (Fig. 5b). The MUC is thus a persistent current by its net poleward transport of a distinct water mass even if at times it meanders, partially recirculates, or interacts with eddies.
Integrating the velocity over the mean line for each glider mission and for each of the thermocline and subthermocline layers (Fig. 11) provides a transport time series. The relative stability of the thermocline transport and variability of the subthermocline transport is apparent in their ranges: about 10 Sv (−24.0 to −13.6 Sv) for the thermocline and 40 Sv (−23.3 to 19.6 Sv) for the subthermocline (4 times as large). The total transport ranges from 2 to −37 Sv, fluctuating with the subthermocline (Fig. 11a). This is in accordance with previous ranges of transport from hydrographic sections of 8 to −40 Sv (Toole et al. 1990; Lukas et al. 1991). Unlike in the NEC, where thermocline and subthermocline transport were highly correlated (Schönau and Rudnick 2015), there does not appear to be a coherent relationship between these transports in the MC/MUC.
The MC/MUC transport may have annual and interannual variability (Lukas 1988; Toole et al. 1990; Qiu and Lukas 1996; Qu et al. 2008; Kashino et al. 2009). Plotting transport by day of year (Fig. 11b), the annual cycle is not discernable. The observed range of transport is larger than a model-estimated annual transport range of 10 Sv (Qiu and Lukas 1996). Thus, the annual cycle may be masked by variability at other time scales. Plotting by the Niño-3.4 index, an inverse linear relationship exists between the Niño-3.4 index and subthermocline transport (Fig. 11c). Transport in the subthermocline is strongly poleward during La Niña and equatorward during El Niño. During ENSO neutral conditions the direction of subthermocline transport tends to be equatorward and less than 8 Sv. The correlation coefficient between the Niño-3.4 index and the subthermocline transport is −0.92 and that between the Niño-3.4 index and the total transport is −0.87. Assuming each glider mission is an independent degree of freedom, the correlation coefficient is significant within a 1% confidence interval. It should be noted that only one ENSO cycle was observed, and these observations only describe circulation near the coast. However, during this time period subthermocline fluctuations were correlated with the Niño-3.4 index and were the leading cause of total transport variability.
The depth penetration of the equatorward MC appears to be the source of the subthermocline transport difference between El Niño and La Niña. Both directions of subthermocline flow were observed during each event (Fig. 10b), but net equatorward transport was 24 Sv greater on average during El Niño than during La Niña, compared to a 9-Sv difference in poleward transport. Composite T–S transport diagrams during each ENSO phase show the distribution of water mass transport (Fig. 12). During neutral conditions (Fig. 12a) transport is similar to the mean (Fig. 5b). El Niño (Fig. 12b) and La Niña (Fig. 12c) have similar thermocline transports and notably different surface and subthermocline transports. During El Niño there is a lack of fresh NPTSW and greater equatorward than poleward transport in the subthermocline. Although there was poleward subthermocline velocity during El Niño (June and December 2009; Fig. 10b), the lack of net poleward transport (Fig. 12b) suggests that there is recirculation or that poleward transport moved offshore of the mean line. During La Niña there was weak equatorward flow of NPIW and large poleward transport of saltier intermediate water across the mean line. The subthermocline transport was thus negative during El Niño because of net equatorward transport of NPIW and positive during La Niña because of net poleward transport of water typical of AAIW, a significant insight into the interannual transport variability of the MC/MUC system.
Composite T–S transport diagrams for glider observations during ENSO (a) neutral, (b) positive, and (c) negative phases. The greatest transport difference between El Niño and La Niña states was in the subthermocline, where equatorward (poleward) transport dominated during El Niño (La Niña). See Fig. 5b for bin size and normalization.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
5. Isopycnal salinity variance
The thermohaline structure of the MC/MUC at finescales can be related to large-scale gradients by comparing salinity variance on isopycnals, separated by length scale with a wavelet transform. At finescales (10 < 80 km; Fig. 13a) salinity variance per kilometer is large from the surface to 25 kg m−3, encompassing the NPTW, a minimum at the base of the thermocline from 25.2 to 25.5 kg m−3, and large in intermediate water deeper than 26 kg m−3. The large-scale salinity variance (120 < 200 km) has a similar vertical structure (Fig. 13b). At all scales, salinity variance is greatest near the coast, where the subsurface salinity extremes of the NPTW and the NPIW create large-scale horizontal salinity gradients.
(a) Finescale (10 < 80 km) and (b) large-scale (120 < 200 km) isopycnal salinity variance from the wavelet transform projected, binned, and averaged over glider missions. Black contours are mean salinity. Salinity variance is large in the subtropical and intermediate water masses, separated by a salinity variance minimum. (c) Average over the mean line for each the finescale and large-scale salinity variance and their ratio.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
The ratio of finescale to large-scale salinity variance, a type of horizontal Cox number (Osborn and Cox 1972), is relatively constant through the water column, with the absolute value dependent on the range of wavelengths included in the transform (Fig. 13c). A vertically constant horizontal Cox number was also observed across the NEC (Schönau and Rudnick 2015). Both results confirm a direct relationship between the finescale and large-scale thermohaline gradients, where the stirring of large-scale gradients causes finescale salinity variance.
6. Water mass modification
Water masses are modified as the NEC advects them into the MC and the MUC advects them into the southern NEUC (Schönau et al. 2015; Wang et al. 2015). Changes in salinity are caused by horizontal and diapycnal diffusion. Salinity decreases and density increases for the subsurface salinity maximum of NPTW (Fig. 14a). Similarly, the salinity minimum of the NPIW becomes saltier and shifts to a denser isopycnal. The downward flux of salt and density between the warm, salty NPTW and cool, fresh NPIW is typical of double diffusion (Schmitt 1981). However, the well-mixed layer between the two (shaded) has a density ratio such that turbulence caused by internal wave shear is likely a dominant mechanism of diapycnal mixing.
(a) Mean salinity averaged on isopycnal surfaces for the MC (Fig. 4) and the mean NEC (Schönau and Rudnick 2015). Salinity extremes of NPTW and NPIW are smaller in the MC than in the NEC. Gray shaded area indicates a density ratio 5 < Rρ < 10, where Rρ is defined as αθz/βSz, and α and β are the thermohaline expansion and saline contraction coefficients, respectively. (b) Finescale (10 < 80 km) isopycnal salinity variance for the MC and NEC. Variance is smaller (larger) for the thermocline (subthermocline) transport in the MC than the NEC. For both (a) and (b), the average for the MC is across the mean line and that for the NEC is from 9° to 15°N along 134.3°E.
Citation: Journal of Physical Oceanography 47, 8; 10.1175/JPO-D-16-0274.1
Changes in finescale salinity variance can be used to track diffusion of water masses. Finescale salinity variance decreases along the isopycnals of NPTW between the NEC and MC (Fig. 14b). In the subthermocline, salinity variance is greater in the MC/MUC than in the NEC/NEUCs. This corresponds to finescale salinity gradients that are introduced by the stirring of the fresh NPIW with saltier intermediate water.
A salinity variance minimum near 25.5 kg m−3 between subtropical and intermediate water masses is a notable feature in both the NEC and MC (Fig. 14b). The salinity variance minimum decreases along the path of flow, similar to NPTW, but shifts from 25.6 to 25.4 kg m−3, opposite the direction of flux by the salinity extrema, which shifts to denser isopycnals (Fig. 14a). It is unclear if this is a remnant of sampling or would apply in other regions. The cause of the variance minimum is unknown. It is within an isopycnal range that has a low horizontal salinity gradient across much of the North Pacific, as observed by Argo climatology, and where interannual climatic variability at the surface has been observed to cause temporal spice variance (Yeager and Large 2007; Sasaki et al. 2010). A step in salinity of 0.08 psu was observed in this isopycnal range in 2012 but was likely observable only because of the small spatial variance along this isopycnal. The persistence of the variance minimum, and its shift to a less dense isopycnal, would make it an interesting layer to study mixing.
7. Conclusions and discussion
Repeat glider observations resolve the mean thermohaline structure and transport of the MC/MUC and their variability during a 4-yr period. The results provide new insight into the structure of the currents near the coast, water mass transport, variability, and finescale structure. The MC is a strong current with a persistent transport of subtropical water masses, extending to a depth of 1000 m near the coast and decreasing in velocity with distance offshore. In the subthermocline, the MUC is a persistent poleward flow that meanders offshore of the MC, with mean transport of a distinct water mass.
The range of subthermocline transport is 4 times greater than that of the thermocline and strongly influences the total transport variability. During the ENSO cycle observed, subthermocline transport along the mean line had an inverse linear relationship with the Niño-3.4 index. There was no evident relationship between the thermocline and subthermocline transport and no discernable annual cycle for total transport, although it may have been dwarfed by strong interannual variability. The subsurface salinity extrema of NPTW and NPIW decrease in magnitude between the NEC and MC along the path of flow, as does the salinity variance in the thermocline. The salinity variance of subthermocline water in the MC/MUC is greater than that in the NEC/NEUCs, as fresh NPIW is stirred with saltier intermediate water in the tropical Pacific. The relationship between isopycnal salinity at large and small scales is consistent with stirring and diffusion of salinity extrema.
Glider-observed mean velocity and standard deviation compare favorably to observations by a single mooring with an ADCP, located at 8°N, 127.3°E from 2010 to 2012 (Zhang et al. 2014). The mooring was offshore of the MC core and had a mean velocity and standard deviation of 0.7 ± 0.2 m s−1 at 100-m depth, nearly identical to that in the glider (Figs. 4b, 7c). The inclusion of spatial variability between glider missions thus had minimal impact on the mean velocity and velocity variance. The gliders passed the mooring only a few times, making a direct velocity comparison unreasonable; however, velocity comparisons between gliders and a mooring array near the Luzon Strait show good agreement between glider-derived absolute geostrophic velocity and mooring ADCP (Lien et al. 2014).
Single moorings in the MC/MUC have observed intraseasonal velocity variability on time scales of 50–80 days (Kashino 2005; Zhang et al. 2014; Kashino et al. 2015). In the subthermocline, mooring observations and model results suggest the MUC velocity varies on even shorter time scales and is also affected by pressure gradients from Rossby waves and ENSO (Hu et al. 2016). Subthermocline eddies that have been observed and modeled in the region at these time scales (Firing et al. 2005; Dutrieux 2009; Chiang and Qu 2013; Wang et al. 2014; Chiang et al. 2015) were observed by the glider (Fig. 10). However, the relationship between intraseasonal velocity variability and transport variability cannot be determined from the temporal resolution of glider observations.
The relationship between the interannual subthermocline transport variability and ENSO is difficult to assess as only one ENSO cycle was observed. The increase in equatorward subthermocline transport during the El Niño may be related to the vertical structure of NEC bifurcation, the strength of the NEC, recirculation of the NECC around the MD, or forcing from the equator. The bifurcation latitude of the NEC increases with increasing depth, leading to a more northerly origin of the MC subthermocline water (Qu and Lukas 2003). As the bifurcation latitude shifts, the vertical structure of the bifurcation also changes, with greater changes in the subthermocline. When the bifurcation occurs at the highest latitude, the subthermocline bifurcates at higher latitude than at the surface (Qu and Lukas 2003; Kim et al. 2004), possibly causing poleward transport to increase more in the subthermocline than in the thermocline. Preliminary investigations with Argo climatology verify that although this occurs during strong ENSO events, there was only a small northward shift of bifurcation latitude in the subthermocline and no change in the thermocline during the 2009/10 El Niño. The vertical structure of the MC/MUC transport and its relationship to ENSO has implications for the ITF transport, which is at a minimum during El Niño and maximum during La Niña (Gordon et al. 1999; Hu et al. 2015). Regional transport changes over a longer temporal range could be investigated with Argo but would be limited near the coasts.
The high horizontal resolution and absolute reference velocity from repeat gliders’ observations have provided accurate water mass transport and exchange in the tropics. These observations are useful in a western boundary current where deep, strong currents near the coast pose a challenge to moorings, floats, and satellite altimetry. Combining such glider observations with Argo and satellite observations, and incorporating these observations into numerical models, would be a powerful tool to examine large-scale water mass transport and heat budgets in the tropical Pacific.
Acknowledgments
Glider observations and analysis were funded by the Office of Naval Research as part of the Origins of Kuroshio and Mindanao Current (OKMC) project through Grants N00014-10-1-0273 and N00014-11-1-0429, and through Flow Encountering Abrupt Topography (FLEAT) project Grant N0014-15-1-2488. Glider observations are available upon request (drudnick@ucsd.edu). We thank the Instrument Development Group at Scripps Institution of Oceanography for all facets of the operations of Spray underwater gliders. We thank Pat and Lori Colin at the Coral Reef Research Foundation for their support in Palau.
REFERENCES
AVISO, 2014: Combined mean dynamic topography. CLS Space Oceanography Division/CNES AVISO, accessed 14 May 2014. [Available online at https://www.aviso.altimetry.fr/en/data/products/auxiliary-products/mdt.html.]
Bingham, F. M., and R. Lukas, 1994: The southward intrusion of North Pacific Intermediate Water along the Mindanao Coast. J. Phys. Oceanogr., 24, 141–154, doi:10.1175/1520-0485(1994)024<0141:TSIONP>2.0.CO;2.
Bingham, F. M., and R. Lukas, 1995: The distribution of intermediate water in the western equatorial Pacific during January–February 1986. Deep-Sea Res. I, 42, 1545–1573, doi:10.1016/0967-0637(95)00064-D.
Bretherton, F. P., R. E. Davis, and C. B. Fandry, 1976: A technique for objective analysis and design of oceanographic experiments applied to MODE-73. Deep-Sea Res. Oceanogr. Abstr., 23, 559–582, doi:10.1016/0011-7471(76)90001-2.
Chelton, D. B., R. A. Deszoeke, and M. G. Schlax, 1998: Geographical variability of the first baroclinic Rossby radius of deformation. J. Phys. Oceanogr., 28, 433–460, doi:10.1175/1520-0485(1998)028<0433:GVOTFB>2.0.CO;2.
Chiang, T.-L., and T. Qu, 2013: Subthermocline eddies in the western equatorial Pacific as shown by an eddy-resolving OGCM. J. Phys. Oceanogr., 43, 1241–1253, doi:10.1175/JPO-D-12-0187.1.
Chiang, T.-L., C.-R. Wu, T. Qu, and Y.-C. Hsin, 2015: Activities of 50-80 day subthermocline eddies near the Philippine coast. J. Geophys. Res. Oceans, 120, 3606–3623, doi:10.1002/2013JC009626.
CPC, 1997: Merged analysis of precipitation (CMAP). NOAA/OAR/ESRL PSD, accessed 5 August 2016. [Available online at http://www.cpc.ncep.noaa.gov/products/global_precip/html/wpage.cmap.html.]
Cronin, M. F., and M. J. McPhaden, 1998: Upper ocean salinity balance in the western equatorial Pacific. J. Geophys. Res., 103, 27 567–27 587, doi:10.1029/98JC02605.
Davis, R. E., M. D. Ohman, D. L. Rudnick, and J. T. Sherman, 2008: Glider surveillance of physics and biology in the southern California Current System. Limnol. Oceanogr., 53, 2151–2168, doi:10.4319/lo.2008.53.5_part_2.2151.
Davis, R. E., W. S. Kessler, and J. T. Sherman, 2012: Gliders measure western boundary current transport from the South Pacific to the equator. J. Phys. Oceanogr., 42, 2001–2013, doi:10.1175/JPO-D-12-022.1.
Delcroix, T., and C. Hénin, 1991: Seasonal and interannual variations of sea surface salinity in the tropical Pacific Ocean. J. Geophys. Res., 96, 22 135–22 150, doi:10.1029/91JC02124.
Delcroix, T., and M. McPhaden, 2002: Interannual sea surface salinity and temperature changes in the western Pacific warm pool during 1992–2000. J. Geophys. Res., 107, 8002, doi:10.1029/2001JC000862.
Dutrieux, P., 2009: Tropical western Pacific currents and the origin of intraseasonal variability below the thermocline. Ph.D. thesis, University of Hawai‘i at Mānoa, 278 pp.
Ferrari, R., and K. L. Polzin, 2005: Finescale structure of the T–S relation in the eastern North Atlantic. J. Phys. Oceanogr., 35, 1437–1454, doi:10.1175/JPO2763.1.
Fine, R. A., R. Lukas, F. M. Bingham, M. J. Warner, and R. H. Gammon, 1994: The western equatorial Pacific: A water mass crossroads. J. Geophys. Res., 99, 25 063–25 080, doi:10.1029/94JC02277.
Firing, E., Y. Kashino, and P. Hacker, 2005: Energetic subthermocline currents observed east of Mindanao. Deep-Sea Res. II, 52, 605–613, doi:10.1016/j.dsr2.2004.12.007.
Gordon, A. L., 1986: Inter-ocean exchange of thermocline water. J. Geophys. Res., 91, 5037–5046, doi:10.1029/JC091iC04p05037.
Gordon, A. L., R. D. Susanto, and A. Ffield, 1999: Throughflow within Makassar Strait. Geophys. Res. Lett., 26, 3325–3328, doi:10.1029/1999GL002340.
Gu, D., and S. G. Philander, 1997: Interdecadal climate fluctuations that depend on exchanges between the tropics and extratropics. Science, 275, 805–807, doi:10.1126/science.275.5301.805.
Hacker, P., E. Firing, R. Lukas, P. L. Richardson, and C. A. Collins, 1989: Observations of the low-latitude western boundary circulation in the Pacific during WEPOCS III. Proc. Western Pacific Int. Meeting and Workshop on TOGA COARE, Nouméa, New Caledonia, Institute Français de Recherche Scientifique pour le Développement en Coopération/Centre ORSTOM, 135–143.
Hsin, Y.-C., and B. Qiu, 2012: The impact of eastern-Pacific versus central-Pacific El Niños on the North Equatorial Countercurrent in the Pacific Ocean. J. Geophys. Res., 117, C11017, doi:10.1029/2012JC008362.
Hu, D. X., M. C. Cui, T. D. Qu, and Y. X. Li, 1991: A subsurface northward current off Mindanao identified by dynamic calculation. Oceanography of Asian Marginal Seas, K. Takano, Ed., Elsevier Oceanography Series, Vol. 54, Elsevier, 359–365.
Hu, D. X., and Coauthors, 2015: Pacific western boundary currents and their roles in climate. Nature, 522, 299–308, doi:10.1038/nature14504.
Hu, S., D. Hu, C. Guan, F. Wang, L. Zhang, F. Wang, and Q. Wang, 2016: Interannual variability of the Mindanao Current/Undercurrent in direct observations and numerical simulations. J. Phys. Oceanogr., 46, 483–499, doi:10.1175/JPO-D-15-0092.1.
Johnson, G. C., B. M. Sloyan, W. S. Kessler, and K. E. McTaggart, 2002: Direct measurements of upper ocean currents and water properties across the tropical Pacific during the 1990s. Prog. Oceanogr., 52, 31–61, doi:10.1016/S0079-6611(02)00021-6.
Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-Year Reanalysis Project. Bull. Amer. Meteor. Soc., 77, 437–471, doi:10.1175/1520-0477(1996)077<0437:TNYRP>2.0.CO;2.
Kashino, Y., 2005: Variability of the Mindanao Current: Mooring observation results. Geophys. Res. Lett., 32, L18611, doi:10.1029/2005GL023880.
Kashino, Y., M. Aoyama, T. Kawano, N. Hendiarti, Y. Anantasena, K. Muneyama, and H. Watanabe, 1996: The water masses between Mindanao and New Guinea. J. Geophys. Res., 101, 12 391–12 400, doi:10.1029/95JC03797.
Kashino, Y., N. España, F. Syamsudin, K. J. Richards, T. Jensen, P. Dutrieux, and A. Ishida, 2009: Observations of the North Equatorial Current, Mindanao Current, and Kuroshio Current System during the 2006/07 El Niño and 2007/08 La Niña. J. Oceanogr., 65, 325–333, doi:10.1007/s10872-009-0030-z.
Kashino, Y., A. Atmadipoera, Y. Kuroda, and Lukijanto, 2013: Observed features of the Halmahera and Mindanao eddies. J. Geophys. Res. Oceans, 118, 6543–6560, doi:10.1002/2013JC009207.
Kashino, Y., I. Ueki, and H. Sasaki, 2015: Ocean variability east of Mindanao: Mooring observations at 7°N, revisited. J. Geophys. Res. Oceans, 120, 2540–2554, doi:10.1002/2015JC010703.
Katsura, S., E. Oka, B. Qiu, and N. Schneider, 2013: Formation and subduction of North Pacific tropical water and their interannual variability. J. Phys. Oceanogr., 43, 2400–2415, doi:10.1175/JPO-D-13-031.1.
Kessler, W. S., 2006: The circulation of the eastern tropical Pacific: A review. Prog. Oceanogr., 69, 181–217, doi:10.1016/j.pocean.2006.03.009.
Kim, Y. Y., T. Qu, T. Jensen, and T. Miyama, 2004: Seasonal and interannual variations of the North Equatorial Current bifurcation in a high-resolution OGCM. J. Geophys. Res., 109, C03040, doi:10.1029/2003JC002013.
Li, Y., and F. Wang, 2012: Spreading and salinity change of North Pacific tropical water in the Philippine Sea. J. Oceanogr., 68, 439–452, doi:10.1007/s10872-012-0110-3.
Li, Y., F. Wang, and W. Han, 2013: Interannual sea surface salinity variations observed in the tropical North Pacific Ocean. Geophys. Res. Lett., 40, 2194–2199, doi:10.1002/grl.50429.
Lien, R.-C., B. Ma, Y.-H. Cheng, C.-R. Ho, B. Qiu, C. M. Lee, and M.-H. Chang, 2014: Modulation of Kuroshio transport by mesoscale eddies at the Luzon Strait entrance. J. Geophys. Res. Oceans, 119, 2129–2142, doi:10.1002/2013JC009548.
Lukas, R., 1988: Interannual fluctuations of the Mindanao Current inferred from sea level. J. Geophys. Res., 93, 6744–6748, doi:10.1029/JC093iC06p06744.
Lukas, R., E. Firing, P. Hacker, P. L. Richardson, C. A. Collins, R. Fine, and R. Gammon, 1991: Observations of the Mindanao Current during the western equatorial Pacific Ocean circulation study. J. Geophys. Res., 96, 7089–7104, doi:10.1029/91JC00062.
Lukas, R., T. Yamagata, and J. P. McCreary, 1996: Pacific low-latitude western boundary currents and the Indonesian Throughflow. J. Geophys. Res., 101, 12 209–12 216, doi:10.1029/96JC01204.
Masumoto, Y., and T. Yamagata, 1991: Response of the western tropical Pacific to the Asian winter monsoon: The generation of the Mindanao Dome. J. Phys. Oceanogr., 21, 1386–1398, doi:10.1175/1520-0485(1991)021<1386:ROTWTP>2.0.CO;2.
NCEP, 1996: NCEP/NCAR Reanalysis monthly means and other derived variables, accessed 19 March 2014. [Available online at https://www.esrl.noaa.gov/psd/data/gridded/data.ncep.reanalysis.derived.surfaceflux.html.]
Osborn, T. R., and C. S. Cox, 1972: Oceanic fine structure. Geophys. Fluid Dyn., 3, 321–345, doi:10.1080/03091927208236085.
Pelland, N. A., C. C. Eriksen, and C. M. Lee, 2013: Subthermocline eddies over the Washington continental slope as observed by Seagliders, 2003–09. J. Phys. Oceanogr., 43, 2025–2053, doi:10.1175/JPO-D-12-086.1.
Pietri, A., V. Echevin, P. Testor, A. Chaigneau, L. Mortier, C. Grados, and A. Albert, 2014: Impact of a coastal-trapped wave on the near-coastal circulation of the Peru upwelling system from glider data. J. Geophys. Res. Oceans, 119, 2109–2120, doi:10.1002/2013JC009270.
Qiu, B., and R. Lukas, 1996: Seasonal and interannual variability of the North Equatorial Current, the Mindanao Current, and the Kuroshio along the Pacific western boundary. J. Geophys. Res., 101, 12 315–12 330, doi:10.1029/95JC03204.
Qiu, B., and S. Chen, 2010: Interannual-to-decadal variability in the bifurcation of the North Equatorial Current off the Philippines. J. Phys. Oceanogr., 40, 2525–2538, doi:10.1175/2010JPO4462.1.
Qiu, B., S. Chen, D. L. Rudnick, and Y. Kashino, 2015: A new paradigm for the North Pacific subthermocline low-latitude western boundary current system. J. Phys. Oceanogr., 45, 2407–2423, doi:10.1175/JPO-D-15-0035.1.
Qu, T., and R. Lukas, 2003: The bifurcation of the North Equatorial Current in the Pacific. J. Phys. Oceanogr., 33, 5–18, doi:10.1175/1520-0485(2003)033<0005:TBOTNE>2.0.CO;2.
Qu, T., and E. J. Lindstrom, 2004: Northward intrusion of Antarctic Intermediate Water in the western Pacific. J. Phys. Oceanogr., 34, 2104–2118, doi:10.1175/1520-0485(2004)034<2104:NIOAIW>2.0.CO;2.
Qu, T., T. Kagimoto, and T. Yamagata, 1997: A subsurface countercurrent along the east coast of Luzon. Deep-Sea Res. I, 44, 413–423, doi:10.1016/S0967-0637(96)00121-5.
Qu, T., H. Mitsudera, and T. Yamagata, 1998: On the western boundary currents in the Philippine Sea. J. Geophys. Res., 103, 7537–7548, doi:10.1029/98JC00263.
Qu, T., H. Mitsudera, and T. Yamagata, 1999: A climatology of the circulation and water mass distribution near the Philippine coast. J. Phys. Oceanogr., 29, 1488–1505, doi:10.1175/1520-0485(1999)029<1488:ACOTCA>2.0.CO;2.
Qu, T., J. Gan, A. Ishida, Y. Kashino, and T. Tozuka, 2008: Semiannual variation in the western tropical Pacific Ocean. Geophys. Res. Lett., 35, L16602, doi:10.1029/2008GL035058.
Roemmich, D., and J. Gilson, 2009: The 2004–2008 mean and annual cycle of temperature, salinity, and steric height in the global ocean from the Argo program. Prog. Oceanogr., 82, 81–100, doi:10.1016/j.pocean.2009.03.004.
Rudnick, D. L., and S. T. Cole, 2011: On sampling the ocean using underwater gliders. J. Geophys. Res., 116, C08010, doi:10.1029/2010JC006849.
Rudnick, D. L., R. E. Davis, C. C. Eriksen, D. M. Fratantoni, and M. J. Perry, 2004: Underwater gliders for ocean research. Mar. Technol. Soc. J., 38, 73–84, doi:10.4031/002533204787522703.
Rudnick, D. L., G. Gopalakrishnan, and B. D. Cornuelle, 2015a: Cyclonic eddies in the Gulf of Mexico: Observations by underwater gliders and simulations by numerical model. J. Phys. Oceanogr., 45, 313–326, doi:10.1175/JPO-D-14-0138.1.
Rudnick, D. L., S. Jan, and C. Lee, 2015b: A new look at circulation in the western North Pacific: Introduction to the special issue. Oceanography, 28, 16–23, doi:10.5670/oceanog.2015.77.
Sasaki, Y. N., N. Schneider, N. Maximenko, and K. Lebedev, 2010: Observational evidence for propagation of decadal spiciness anomalies in the North Pacific. Geophys. Res. Lett., 37, L07708, doi:10.1029/2010GL042716.
Schmitt, R. W., 1981: Form of the temperature-salinity relationship in the central water: Evidence for double-diffusive mixing. J. Phys. Oceanogr., 11, 1015–1026, doi:10.1175/1520-0485(1981)011<1015:FOTTSR>2.0.CO;2.
Schönau, M. C., and D. L. Rudnick, 2015: Glider observations of the North Equatorial Current in the western tropical Pacific. J. Geophys. Res. Oceans, 120, 3586–3605, doi:10.1002/2014JC010595.
Schönau, M. C., D. L. Rudnick, I. Cerovecki, G. Gopalakrishnan, B. Cornuelle, J. McClean, and B. Qiu, 2015: The Mindanao Current: Mean structure and connectivity. Oceanography, 28, 34–45, doi:10.5670/oceanog.2015.79.
Scripps Institution of Oceanography and International Argo Program, 2009: Roemmich-Gilson Argo Climatology, accessed 20 October 2016. [Available online at http://sio-argo.ucsd.edu/RG_Climatology.html.]
Sherman, J., R. E. Davis, W. B. Owens, and J. Valdes, 2001: The autonomous underwater glider “Spray.” IEEE J. Oceanic Eng., 26, 437–446, doi:10.1109/48.972076.
Talley, L. D., 1993: Distribution and formation of North Pacific Intermediate Water. J. Phys. Oceanogr., 23, 517–537, doi:10.1175/1520-0485(1993)023<0517:DAFONP>2.0.CO;2.
Todd, R. E., D. L. Rudnick, M. R. Mazloff, R. E. Davis, and B. D. Cornuelle, 2011: Poleward flows in the southern California Current System: Glider observations and numerical simulation. J. Geophys. Res., 116, C02026, doi:10.1029/2010JC006536.
Todd, R. E., D. L. Rudnick, M. R. Mazloff, B. D. Cornuelle, and R. E. Davis, 2012: Thermohaline structure in the California Current System: Observations and modeling of spice variance. J. Geophys. Res., 117, C02008, doi:10.1029/2011JC007589.
Toole, J. M., R. C. Millard, Z. Wang, and S. Pu, 1990: Observations of the Pacific North Equatorial Current bifurcation at the Philippine coast. J. Phys. Oceanogr., 20, 307–318, doi:10.1175/1520-0485(1990)020<0307:OOTPNE>2.0.CO;2.
Tozuka, T., T. Kagimoto, and Y. Masumoto, 2002: Simulated multiscale variations in the western tropical Pacific: The Mindanao Dome revisited. J. Phys. Oceanogr., 32, 1338–1359, doi:10.1175/1520-0485(2002)032<1338:SMVITW>2.0.CO;2.
Tsuchiya, M., 1968: Upper Waters of the Intertropical Pacific Ocean. Johns Hopkins Press, 50 pp.
Wang, F., N. Zang, Y. Li, and D. Hu, 2015: On the subsurface countercurrents in the Philippine Sea. J. Geophys. Res. Oceans, 120, 131–144, doi:10.1002/2013JC009690.
Wang, Q., F. Zhai, F. Wang, and D. Hu, 2014: Intraseasonal variability of the subthermocline current east of Mindanao. J. Geophys. Res. Oceans, 119, 8552–8566, doi:10.1002/2014JC010343.
Wijffels, S., E. Firing, and J. Toole, 1995: The mean structure and variability of the Mindanao Current at 8°N. J. Geophys. Res., 100, 18 421–18 435, doi:10.1029/95JC01347.
Wyrtki, K., 1961: Physical Oceanography of the Southeast Asian Waters. NAGA Rep., Vol. 2, Scripps Institute of Oceanography, 195 pp.
Wyrtki, K., and B. Kilonsky, 1984: Mean water and current structure during the Hawaii-to-Tahiti Shuttle Experiment. J. Phys. Oceanogr., 14, 242–254, doi:10.1175/1520-0485(1984)014<0242:MWACSD>2.0.CO;2.
Xie, P., and P. A. Arkin, 1997: Global precipitation: A 17-year monthly analysis based on gauge observations, satellite estimates, and numerical model outputs. Bull. Amer. Meteor. Soc., 78, 2539–2558, doi:10.1175/1520-0477(1997)078<2539:GPAYMA>2.0.CO;2.
Yeager, S. G., and W. G. Large, 2007: Observational evidence of winter spice injection. J. Phys. Oceanogr., 37, 2895–2919, doi:10.1175/2007JPO3629.1.
Zenk, W., G. Siedler, A. Ishida, J. Holfort, Y. Kashino, Y. Kuroda, T. Miyama, and T. J. Müller, 2005: Pathways and variability of the Antarctic Intermediate Water in the western equatorial Pacific Ocean. Prog. Oceanogr., 67, 245–281, doi:10.1016/j.pocean.2005.05.003.
Zhang, L., D. Hu, S. Hu, F. Wang, F. Wang, and D. Yuan, 2014: Mindanao Current/Undercurrent measured by a subsurface mooring. J. Geophys. Res. Oceans, 119, 3617–3628, doi:10.1002/2013JC009693.
Zhang, R.-H., L. M. Rothstein, and A. J. Busalacchi, 1998: Origin of upper-ocean warming and El Niño change on decadal scales in the tropical Pacific Ocean. Nature, 391, 879–883, doi:10.1038/36081.