Mindanao Current and Undercurrent: Thermohaline Structure and Transportfrom Repeat Glider Observations

MARTHA C. SCHÖNAU AND DANIEL L. RUDNICK

Scripps Institution of Oceanography, University of California, San Diego, La Jolla, California

(Manuscript received 15 December 2016, in final form 3 May 2017)

ABSTRACT

Autonomous underwater Spray gliders made repeat transects of the Mindanao Current (MC), a low-

latitude western boundary current in the western tropical North Pacific Ocean, from September 2009 to

October 2013. In the thermocline (,26 kgm23), the MC has a maximum velocity core of 20.95m s21,

weakeningwith distance offshore until it intersects with the intermittentMindanaoEddy (ME) at 129.258E. Inthe subthermocline (.26 kgm23), a persistent Mindanao Undercurrent (MUC), with a velocity core of

0.2m s21 andmean net transport, flows poleward.Mean transport and standard deviation integrated from the

coast to 1308E is 219 6 3.1 Sv (1 Sv[ 106m3 s21) in the thermocline and 236 12 Sv in the subthermocline.

Subthermocline transport has an inverse linear relationship with the Niño-3.4 index and is the primary in-

fluence of total transport variability. Interannual anomalies during El Niño are greater than the annual cycle

for sea surface salinity and thermocline depth. Water masses transported by the MC/MUC are identified

by subsurface salinity extrema and are on isopycnals that have increased finescale salinity variance (spice

variance) from eddy stirring. The MC/MUC spice variance is smaller in the thermocline and greater in the

subthermocline when compared to the North Equatorial Current and its undercurrents.

1. Background and introduction

TheMindanaoCurrent (MC) is a low-latitudewestern

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).

TheMC flows equatorward along the coast ofMindanao

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 im-

pacts 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 In-

termediate Waters equatorward while a Mindanao

Undercurrent (MUC; Hu et al. 1991) flows poleward.

The confluence of subtropical, tropical, and inter-

mediate water masses of North and South Pacific origin

and their role in salt and heat exchange of the tropical

Pacific and IndianOceans has led to several initiatives in

the last two decades to observe andmodel theMC/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/MUCwatermass

transport, coastal structure, and variability are not well

understood owing to strong currents, eddy activity, and

large annual and interannual changes in surface forcing.

From hydrographic surveys, the MC current is narrow

at 108N, then widens, shoals, and increases in speed as it

flows equatorward until it separates into the ITF and

NECC between 78N and 68N (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 78N, 1308E, causes the MC

transport to increase as it flows equatorward (Lukas

et al. 1991). At 88N, the mean MC is in geostrophic

balance, extending approximately 200km offshore,

with a subsurface velocity core of 0.8 to 1ms21 (Wijffels

et al. 1995; Qu et al. 1998). Although the ME has beenCorresponding author: Martha Schönau, mcschonau@ucsd.edu

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DOI: 10.1175/JPO-D-16-0274.1

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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 polewardMUC is roughly 50

to 80km offshore of the coast of Mindanao near 500-m

depth and potential density of 27.2kgm23 (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 con-

nectivity of the MUC, and its relationship to the thermo-

cline 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 transportwhen theNECbifurcates 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 IntermediateWater (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 In-

termediate 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).

FIG. 1. (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 amean line from 8.158N, 126.618E to 8.778N, 1308E (thick black line). Themean line is perpendicular to coastal

bathymetry and MDT near the coast.

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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 parameteri-

zations 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 toOctober 2013 under theOrigins

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

2000km during all seasons and one ENSO cycle. Ob-

jective maps of geostrophic velocity, salinity, and po-

tential 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 ob-

servations 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 ther-

mohaline structure.

2. Data collection and analysis

To observe theMC/MUC, the autonomous underwater

glider Spray (Sherman et al. 2001; Rudnick et al. 2004)

made repeat transects from Palau to the coast of Mind-

anao. 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 1000m and back over a horizontal distance of

6km in a cycle time of 6h. Data are collected during as-

cent at a vertical resolution of about 1m 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 (7.58N, 134.58E) until reaching the 1000-m isobath

along the coast ofMindanao at a latitude between 78 and108N 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 suf-

ficient to observe theMC 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.

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 9.38N, 126.318E to

roughly 1308E. 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 objec-

tivelymapped on depth surfaces (Fig. 3b). Alternatively,

to maintain the horizontal resolution of 6 km, in-

terpolating data to density surfaces also filters out in-

ternal wave motion (Fig. 3c).

To make a mean section, glider observations are lin-

early projected and objectively mapped onto a mean

line extending from the coast of Mindanao, from 8.158N,

126.618E to 8.778N, 1308E (Figs. 1b, 2a), hereinafter re-

ferred 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 calcu-

lated 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 com-

pares favorably to ADCP measurements from a moor-

ing (Lien et al. 2014), satellite sea surface height

(Rudnick et al. 2015a), and to Argo climatology refer-

enced to a level of no motion at 2000m in the tropical

Pacific (Schönau and Rudnick 2015).

The mean line extends from the 1000-m isobath ap-

proximately 380 km offshore and was chosen at an angle

808E of north to be perpendicular to isobaths and con-

tours 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 (1/68; ScrippsInstitution of Oceanography and International Argo

Program 2009; Roemmich and Gilson 2009) indicates

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that there is alongshore shoaling of the thermocline of

20m 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. Re-

moving 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 to-

gether averages out the linear alongshore gradient, as

gliders tended to traverse north then south of the mean

line.On four of themissions, the glider trajectory exceeded

220km 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

10m. 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 re-

flects spatial and temporal variability. However, compari-

sons 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 ob-

jectively 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

FIG. 2. (a) Trajectories of the 10 gliders deployed fromPalau to the 1000-m isobath along the coast ofMindanao. Arrows show themean

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 obser-

vations 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.

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poleward (red) cores offshore that may be expressions

of theMUC. The range is extended to 1348E 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 (18)and annual (1/28) 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 avail-

able from CPC Merged Analysis of Precipitation

(CMAP; CPC 1997; Xie and Arkin 1997), each gridded

at 2.58 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 c(x)5 ei2pkxe2x2/2, with k 5 1. For

each deployment, the transform is performed on the

along-track isopycnal salinity that is detrended and

interpolated to a 1-km grid. Performing the wavelet

transform on the entire mission avoids cutting off low-

wavenumber variance near the coast of Mindanao where

the glider turns around. The resulting salinity variance is

projected for each of the 16 sections and averaged in 10-km

bins along the mean line. The mean salinity variance is

an average over all transects. To separate length scales,

different ranges of wavelengths are included in the

wavelet transform for finescales (10 to 80km) and large

scales (120 to 200 km). The finescale scale ranges from

the nearest spacing of glider observations to the decor-

relation length scale, and the large scale begins at the e-

folding envelope of the chosen Morlet wavelet.

3. Mean structure

a. Water masses

Water masses are identified by their salinity extrema.

Themean salinity has two haloclines (Fig. 4a): first, from

the surface fresh layer that overlies themean section to a

subducted subtropical salinity maximum and second, in

FIG. 3. 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 9.38N, 126.318E to 8.228N, 1308E: (a) vertically binned on depth surfaces,

(b) objectively mapped onto depth surfaces, and (c) interpolated to potential density surfaces. Tickmarks 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 kgm23, respectively. The range is extended to 9.158N, 1348E to

show the relatively weak velocities east of 1308E, denoted by the dashed line. Negative (positive) velocity is equatorward (poleward).

(e) Objectively mapped depth-averaged velocity for mission 11A03601.

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FIG. 4. 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.

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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 .288C; Wyrtki and Kilonsky 1984),

formed by local precipitation and advection (Delcroix

and Hénin 1991). NPTSW is influenced by annual vari-

ability of the East Asian monsoon and interannual

variability associated with ENSO (Delcroix and Hénin1991; Li et al. 2013). NPTSW creates a large vertical

salinity gradient within the upper 50m 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 po-

tential density 23.5 kgm23. 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 kgm23 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 kgm23 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 path-

way 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 1308E(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).

b. Geostrophic velocity

The mean absolute geostrophic velocity of the MC,

referenced to the depth-averaged velocity, is equator-

ward, with a subsurface velocity maximum of21.0m s21

at 100-m depth (22 kgm23) and 50km offshore (1278E;Figs. 4b,c). Near the coast, where the MC is the stron-

gest, 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 250km offshore (128.758E).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 evi-

dence of a mean ME. The width of the possible ME is

FIG. 5. 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 0.28C 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.

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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 1000m near the coast, decreasing in velocity with

increasing depth and distance offshore (Fig. 4b). Off-

shore of the equatorward subthermocline MC is a

poleward MUC. The MUC has a maximum velocity of

0.17m s21 located 75km offshore (127.758E) and 650-m

depth (27.2 kgm23).

East of theMUC there is weak equatorward flow until

128.58E; then mean poleward flow exists through the

offshore edge of the mean line. The coastal structure of

the equatorwardMCandmean location andwidth of the

MUC referenced to an absolute velocity are a funda-

mental improvement to previous observations, which

have either been time series at single locations or ref-

erenced 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 kgm23 is 222.4 Sv

(1Sv 5 106m3 s21), with a thermocline transport (surface

to 26kgm23) of 219Sv and subthermocline transport

(26 to 27.3kgm23) of23Sv. Potential density 27.3kgm23

is the densest isopycnal that is present in all sections. Po-

tential density 26kgm23, 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 228.3Sv during this time period, and hydrographic

estimates of themean transport from the surface to 1000m

and coast to 1308E are around227Sv (Wijffels et al. 1995;

Qu et al. 1998). These estimates exceed the glider trans-

port as they do not consider or resolve the poleward un-

dercurrent. To compare, glider transport summed from

the coast approaches 230Sv before decreasing offshore

(Fig. 6a). The sumof equatorward transport from the coast

to 128.58E and surface to 27.3kgm23 is 230.4Sv. In the

subthermocline, the poleward transport of the inshore

MUC (127.258E) is 3.6Sv, and the poleward sub-

thermocline transport across the mean line is 8.9Sv.

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 128.58E (Fig. 6b). The possible

FIG. 6. (a) Total transport (black) summed along the mean line and separated into thermocline (blue; surface to

26 kgm23) and subthermocline (red; 26 to 27.3 kgm23) 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 func-

tion 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.

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ME, centered at 129.258E, 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 equatorwardMC, and then becomes

poleward between 1278 and 127.758E at the mean loca-

tion of the MUC. As a function of density, mean trans-

port is greatest at the surface between potential density

21 and 22kgm23 and uniform through the thermocline.

Deeper than 27kgm23, the large standard error in-

dicates that the direction of mean net transport is in-

conclusive (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 in-

tegrates to the average transport over the 10 missions

from the surface to 1000-m depth. Both NPTW and

NPIW have equatorward transport. In the sub-

thermocline, there is a net poleward transport of water

with potential density 27 to 27.5 kgm23 and salinity

greater than 34.5 psu, typical of AAIW. The poleward

subthermocline transport shallower than 27.3 kgm23 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 vari-

ability and the responsible forcing.

a. Thermocline depth

The standard deviation in potential temperature is

greatest in the thermocline, between depths of 100 to

200m, 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 80m, isotherms tend to follow iso-

pycnals, causing isopycnal displacement to appear as

temperature variability. The thermocline depth (mea-

sured as the depth of the layer 23–24 kgm23) is at a

minimum in March 2010, coinciding with El Niño, andat 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.

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 1508Ebetween the MC, NEC, and NECC (Masumoto and

Yamagata 1991). TheMDdecays inMay as theEastAsian

monsoon shifts to southwesterlies and the annual down-

welling Rossby wave arrives from the central Pacific

(Tozuka et al. 2002).Annual thermocline depth variability,

taken as the depth of the 218C isotherm from Argo cli-

matology at 8.58N is consistent with this forcing (Fig. 8b).

At 128.58E, 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.

To assess annual and interannual forcing, glider ob-

servations can be compared to Argo climatology and a

linear, 1.5-layer linear model that allows the propaga-

tion of Rossby waves (Kessler 2006; Qiu and Chen

2010). Following Kessler (2006), the depth anomaly of

the thermocline, h (positive upward), is governed by

›h

›t1 c

r

›h

›x1Rh5=3

�t

fr

�.

Here, R is a damping coefficient time scale of roughly

9 months, f is the Coriolis parameter, t the wind stress,

and r the density of seawater. The long Rossby wave

speed is cr 5 2bc2/f, where b is the meridional de-

rivative of f, and c is the internal gravity wave speed,

which is taken to be;2.6–3ms21 across the Pacific basin

(Chelton et al. 1998). The solution is given by

h(x, t)51

cr

ðxxE

e2(R/cr)(x2x0)=3tfx0, [t2 (x2 x0)/c

r]g

frdx0 1 e2(R/cr)(x2x0)h

E[t2 (x2 x0)/c

r] ,

where hE is the depth anomaly on the eastern boundary,

and h is the depth anomaly on the western boundary.

The thermocline depth thus depends on local wind, the

propagation of Rossby waves forced by wind stress curl

in the central and eastern Pacific, and Rossby wave ra-

diation from the eastern boundary from coastally trap-

ped Kelvin waves; however, the dominant term is the

propagation of Rossby waves generated in the central

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Pacific (Qiu and Chen 2010). At 128.58E, the model shows

depth variance consistent with annual forcing: deep in

summer and fall and shallow during winter (Fig. 8c).

The model, Argo climatology, and glider results are

compared at 8.58N and 128.58E, 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 thermo-

cline than the climatology. The model and Argo clima-

tology have roughly the same magnitude but differ in

FIG. 7. 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 kgm23 at 1 kgm23 increments and at

(c) 26 kgm23. The standard deviation includes both temporal and spatial variability over the

range of the glider missions in Fig. 2b.

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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 cli-

matology and gliders observed shoaling of the thermo-

cline 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 100m (.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 ac-

count 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 LaNiña, withlow variability in intermittent years (Fig. 9a).

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, consis-

tent 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 sur-

face salinity are each large following El Niño and La

Niña. The precipitation anomaly lags the Niño-3.4

FIG. 8. 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 kgm23 for each glider

mission. Shading on date axis indicates observations duringElNiño (red) and LaNiña (blue) according to theNiño-3.4 index. (b) Annual cycle of 218C isotherm depth from Argo climatology (Roemmich and Gilson 2009) along

8.58N. (c) Depth anomaly of the 218C isotherm for each a linear, 1.5-layer, wind-forced model that allows Rossby

wave propagation, monthly Argo climatology, and glider observations, all at 8.58N, 128.58E. The mean from June

2009 to October 2013 has been removed from each but not an annual cycle. The linear model andArgo 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.

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index: negative following the 2009/10 El Niño and pos-

itive following the 2010/11 La Niña (Fig. 9c). The in-

terannual anomaly from the glider (where the annual

cycle from Argo climatology has been removed) is cor-

respondingly positive during lack of rainfall and nega-

tive 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 pre-

cipitation 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 ad-

vected away or from upwelling and mixing with un-

derlying 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

FIG. 9. Sea surface salinity and precipitation variability. (a) Salinity anomaly (0–50m) from each glider mission.

(b) Annual anomaly of precipitation and surface salinity (0–50m) 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 8.58N, 127.58E,and precipitation is averaged over the region 3.758–13.758N and 126.258–136.258E, 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 (0–50m, 127.58E), 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 20.75, with the Niño-3.4 index leading the salinity

anomaly and precipitation by 1 month. The precipitation anomaly and salinity anomaly have a maximum corre-

lation of 20.65 with a zero time lag. Results depend on the area over which precipitation is averaged.

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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.

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

theMC 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 theME (129.258E), thevariability is roughly equal to the mean, suggesting the

ME was intermittently observed. There is a vertical

minimum in standard deviation near potential density

26 kgm23 that separates the thermocline and sub-

thermocline 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 ex-

ceeds 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 127.58and 129.58E, and becomes poleward in roughly half the

sections (Fig. 10a). The poleward transport may be an

expression of theME 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 dif-

ficult to definemultiple cores of theMUC, as it is unclear

if each has net water mass transport or recirculates. For

example, mission 11A03601 has symmetric isohalines in

FIG. 10. Transport per distance for each glider mission summed from (a) the surface to 26 kgm23 and (b) 26 to 27.3 kgm23, where

26 kgm23 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 1288E to encompass the

equatorwardMC and first polewardMUC core; (d) as in (c), but from the coast to 1298E to encompass the equatorwardMC, first poleward

MUC core, and offshore equatorward flow; and (e) as in (c), but from the coast to 1308E. 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).

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the subthermocline between poleward and equatorward

velocities, centered at 1298E, suggesting water mass re-

circulation (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

1288E has equatorward transport of NPIW and a pole-

ward transport of water typical of AAIW (.34.5psu,

27.2kgm23; Fig. 10c). Summing transport from the coast

to 1298E (Fig. 10d) now includes saltier equatorward

subthermocline transport (;27 kgm23, .34.5 psu).

However, thisT–S range has net poleward transport when

integrating to 1308E (Fig. 10e). Thus, the subthermocline

equatorward flow between 1288 and 1298E (Fig. 10b) is

likely a partial cyclonic circulation from farther offshore,

as confirmed by cyclonic depth-averaged velocity cen-

tered at 1298E (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). TheMUC

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 sub-

thermocline layers (Fig. 11) provides a transport time

series. The relative stability of the thermocline transport

and variability of the subthermocline transport is

FIG. 11. Transport through the mean line for each glider mission from the surface to

27.3 kgm23 (black) and divided into thermocline (blue) and subthermocline (red) transport by

potential density 26 kgm23. 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.

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apparent in their ranges: about 10Sv (224.0 to213.6Sv)

for the thermocline and 40Sv (223.3 to 19.6Sv) for the

subthermocline (4 times as large). The total transport

ranges from 2 to 237Sv, fluctuating with the sub-

thermocline (Fig. 11a). This is in accordance with pre-

vious ranges of transport from hydrographic sections of 8

to240Sv (Toole et al. 1990; Lukas et al. 1991). Unlike in

the NEC, where thermocline and subthermocline trans-

port were highly correlated (Schönau andRudnick 2015),

there does not appear to be a coherent relationship be-

tween these transports in the MC/MUC.

The MC/MUC transport may have annual and in-

terannual 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 re-

lationship exists between the Niño-3.4 index and sub-

thermocline transport (Fig. 11c). Transport in the

subthermocline is strongly poleward during La Niña andequatorward during El Niño. During ENSO neutral

conditions the direction of subthermocline transport

tends to be equatorward and less than 8Sv. The corre-

lation coefficient between the Niño-3.4 index and the

subthermocline transport is 20.92 and that between

the Niño-3.4 index and the total transport is 20.87.

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 ob-

servations only describe circulation near the coast.

However, during this time period subthermocline fluc-

tuations were correlated with the Niño-3.4 index and

were the leading cause of total transport variability.

The depth penetration of the equatorward MC ap-

pears to be the source of the subthermocline transport

difference between El Niño and La Niña. Both di-

rections 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 trans-

port. 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 thermo-

cline transports and notably different surface and sub-

thermocline 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 sub-

thermocline transport was thus negative during El Niñobecause of net equatorward transport of NPIW and pos-

itive during La Niña because of net poleward transport of

water typical of AAIW, a significant insight into the in-

terannual transport variability of the MC/MUC system.

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 , 80km; Fig. 13a) salinity variance per kilometer is

large from the surface to 25 kgm23, encompassing the

FIG. 12. 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.

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NPTW, a minimum at the base of the thermocline from

25.2 to 25.5 kgm23, and large in intermediate water

deeper than 26kgm23. 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.

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 con-

stant 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 dia-

pycnal diffusion. Salinity decreases and density in-

creases 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.

Changes in finescale salinity variance can be used to

track diffusion of water masses. Finescale salinity vari-

ance 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 varianceminimumnear 25.5kgm23 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 25.4kgm23,

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

FIG. 13. (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.

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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 ob-

served in this isopycnal range in 2012 but was likely ob-

servable 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 ther-

mohaline 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 1000m 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 be-

tween the thermocline and subthermocline transport and

nodiscernable 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 in-

termediate 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 de-

viation compare favorably to observations by a single

mooring with an ADCP, located at 88N, 127.38E from

2010 to 2012 (Zhang et al. 2014). The mooring was off-

shore of the MC core and had a mean velocity and

standard deviation of 0.7 6 0.2m s21 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

FIG. 14. (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, Rr , 10, where Rr is defined as auz/bSz, and a and b are the thermohaline expansion

and saline contraction coefficients, respectively. (b) Finescale (10, 80 km) isopycnal salinity variance for the MC

andNEC.Variance is smaller (larger) for the thermocline (subthermocline) transport in theMC than theNEC. For

both (a) and (b), the average for the MC is across the mean line and that for the NEC is from 98 to 158N along

134.38E.

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velocity variance. The gliders passed the mooring only

a few times, making a direct velocity comparison un-

reasonable; however, velocity comparisons between

gliders and a mooring array near the Luzon Strait show

good agreement between glider-derived absolute geo-

strophic 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 gra-

dients 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 intra-

seasonal velocity variability and transport variability

cannot be determined from the temporal resolution of

glider observations.

The relationship between the interannual sub-

thermocline transport variability and ENSO is difficult

to assess as only one ENSO cycle was observed. The

increase in equatorward subthermocline transport dur-

ing the El Niño may be related to the vertical structure

of NEC bifurcation, the strength of the NEC, re-

circulation 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 bi-

furcation occurs at the highest latitude, the sub-

thermocline 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 al-

though 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 theMC/MUC transport and its relationship to ENSO has

implications for the ITF transport, which is at a mini-

mum 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 investi-

gated 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 pow-

erful 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 re-

quest (drudnick@ucsd.edu). We thank the Instrument

Development Group at Scripps Institution of Ocean-

ography for all facets of the operations of Spray un-

derwater gliders. We thank Pat and Lori Colin at the

Coral Reef Research Foundation for their support

in Palau.

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