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Effect of Mesoscale Eddies on the Taiwan Strait Current
YU-LIN CHANG
National Taiwan Normal University, Taipei, Taiwan

YASUMASA MIYAZAWA
Japan Agency for Marine-Earth Science and Technology, Yokohama, Kanagawa, JapanXINYU GUO
Ehime University, Matsuyama, Ehime, Japan(Manuscript received 10 December 2014, in final form 22 March 2015)
ABSTRACTThis study shows that mesoscale eddies can alter the Taiwan Strait Current. The 20-yr data-assimilated
Japan Coastal Ocean Predictability Experiment 2 (JCOPE2) reanalysis data are analyzed, and the results are
confirmed with idealized experiments. The leading wind-forced seasonal cycle is excluded to focus on theeffect of the eddy. The warm eddy southwest of Taiwan is shown to generate a northward flow, whereas thecold eddy produces a southward current. The effect of the eddy penetrates onto the shelf through the joint
effect of baroclinicity and relief (JEBAR). The cross-isobath fluxes lead to shelfward convergence and divergence, setting up the modulation of the sea level slope. The resulting along-strait current anomaly
eventually affects a wide area of the Taiwan Strait. The stronger eddy leads to larger modification of the crossshelf flows and sea level slope, producing a greater transport anomaly. The composite Sea-Viewing Wide
Field-of-View Sensor chlorophyll-a (Chl-a) serves as an indicator to show the change in Chl-a concentrationin the strait in response to the eddy-induced current. During the warm eddy period, the current carries thesouthern water of lower concentration northward, reducing Chl-a concentration in the strait. In contrast,
Chl-a is enhanced because the cold eddy–induced southward current carries the northern water of higher
concentration southward into the strait.

1. Introduction
The Taiwan Strait lies between southeastern China
and Taiwan (Fig. 1) and is oriented in southwest andnortheast directions. The water depths in the Taiwan
Strait are generally shallower than 50 m, except for the
deeper Penghu Channel to the southwest of Taiwan.
With a length of ;350 km, the Taiwan Strait is a passage
connecting two large marginal seas: the South China Sea
to the south and the East China Sea to the north. Although the strait is shallow, it plays an essential role in
the water exchange between the two marginal seas andthe Pacific Ocean (Wyrtki 1961; Liu et al. 2000).
The volume transport of the Taiwan Strait had been
investigated since the early 1960s. Based on sea level

Corresponding author address: Yu-Lin Chang, S609, Lecture andResearch Building 88, Sec. 4, Tingzhou Rd., Taipei 10677, Taiwan.
E-mail: ychang@ntnu.edu.tw
DOI: 10.1175/JPO-D-14-0248.1
Ó 2015 American Meteorological Society

station data, Wyrtki (1961) estimated the Taiwan Straittransport to be less than 1 Sverdrup (Sv; 1 Sv [ 106 m3 s21)
in July and December. Fang et al. (1991) combined several
in situ current meter datasets to obtain transport values of
3.1 Sv in summer and 1.0 Sv in winter and an annual-meantransport of 2 Sv. Using the observed shipboard acoustic
Doppler current profiler (sb-ADCP), Wang et al. (2003)
calculated a mean transport of 1.8 Sv, which was stronger
in July (2.7 Sv) and weaker in January (0.9 Sv). Jan et al.
(2006) used cluster analysis for the ADCP data from 1999
to 2001 to derive a transport of 1.16–2.34 Sv from March toAugust that changes into a weak transport of 20.1 Sv
southward in winter. Negative transport was also observedby Lin et al. (2005) using a bottom-mounted ADCP during
October–December 1999 under a strong northeasterly
monsoon. The modeling study by Wu and Hsin (2005)
proposed a transport of 1.09 6 1.67 Sv based on a 5-yr run.
Their simulation also shows stronger transport in summer
than that in winter.

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FIG. 1. (a) Bathymetry (m) of the study region and (b) the time series of mean along-strait velocity (m s21) takenfrom the area between the two black dotted lines in (a). (c) Power spectra (x axis represents days, and y axis is thepower in Watts) for the along-strait velocity in (b).

The Taiwan Strait transport reveals strong seasonal
signals that are stronger in summer than those in winter.
The annual-mean transport in the Taiwan Strait isnorthward, flowing along the strait. The net northwardtransport under the absence of wind is induced by thepersistence of sea level slope (Yang 2007). The East
Asia monsoon is considered to be the main forcing responsible for the seasonal variation (Wang et al. 2003;
Wu and Hsin 2005; Jan et al. 2006; Wu et al. 2007). The
summer southwesterly wind together with the persistent
sea level slope causes a negative pressure gradient thatdrives the stronger along-strait flow. In winter, thenortheasterly wind competes with the backgroundpressure gradient, resulting in weaker northward oreven southward flow (Oey et al. 2014). On the basis of
empirical orthogonal function (EOF) results, a recent
study by Oey et al. (2014) offered a different mechanismof the wind-driven, fluctuating pressure gradient on atime scale of days. They determined that the northeasterly wind sets up the high sea level to the southwestern strait such that the flow moves northward while
the wind relaxes. Moreover, W.-Z. Zhang et al. (2014)
used a 4-yr numerical simulation and suggested that theeffect of typhoons can reduce the Taiwan Strait transport by up to 0.45 Sv and the annual mean by 0.09 Sv.

The wind is considered to be the force affecting the
Taiwan Strait Current variation for time scales of days toseasons.
In the present study, a new candidate other than windis proposed. Oceanic eddies originate almost everywhere in the world’s oceans. Mesoscale eddies play important roles in transporting heat and in biochemical
processes (Volkov et al. 2008; Chelton et al. 2011;
Mahadevan et al. 2012). The mass transport by global
mesoscale eddies of 30–40 Sv has a magnitude comparable to that of large-scale, wind-driven, and thermohaline circulation (Z. Zhang et al. 2014). Eddies arecharacterized by their distinct sea surface height fromthe ambient water. The effect of warm-core eddies on
the along-shelf pressure gradient in the Middle Atlantic
Bight was demonstrated by Xu and Oey (2011), whoshowed that the impingement of warm-core eddies upon
the shelf break can produce sea level changes on the
shelf. Mesoscale eddy activities are often observed insouthwestern Taiwan (Nan et al. 2011). The mesoscalefeatures to the southwest of Taiwan are attributed to
locally formed eddies, westward penetration subtropicalcountercurrent eddies, Kuroshio-shedding eddies, or the
Kuroshio intrusion through the Luzon Strait (Li et al.
1998; Nan et al. 2011; Sheu et al. 2010). The present

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FIG. 2. (a) 20-yr mean EKE (m2 s22) with 180-day high pass and(b) time series of the EKE of the black box in (a).

study aims to demonstrate that eddies can create external
forcing to affect the variation of the Taiwan Strait Current, even though they cannot intrude into the shallow
coastal water owing to their deeper vertical structures.

2. Data and methods
The analysis in this study is based on the 20-yr(1993–2012) model reanalysis results from Japan
Coastal Ocean Predictability Experiment 2 (JCOPE2;
Miyazawa et al. 2009). JCOPE2 is a data-assimilating
ocean model based on the Princeton Ocean Model(POM) and includes a generalized coordinate system(Mellor et al. 2002). The domain encompasses thewestern North Pacific Ocean (10.58–628N, 1088–1808E)
with a horizontal resolution of 1/ 128 and 46 vertical layers.
The lateral boundary conditions are determinedfrom a basinwide model using a one-way nesting1653

method (Guo et al. 2003). The model is forced bywind stresses, heat, and salt fluxes from the 6-hourly
National Centers for Environmental Prediction–
National Center for Atmospheric Research (NCEP–
NCAR) reanalysis data. In situ temperature and salinity datafrom the Advanced Very High Resolution Radiometer/
multichannel sea surface temperature (AVHRR/MCSST), U.S. Global Ocean Data Assimilation Experiment (USGODAE) project (http://www.usgodae.org/),
Colorado Center for Astrodynamics Research (CCAR)
near-real-time altimetry data (http://eddy.colorado.edu/
ccar/data_viewer/index), and Global Temperature andSalinity Profile Programme (GTSPP) are assimilated into
the model. Details of the model setup and descriptions
have been reported by Miyazawa et al. (2009).
Idealized experiments were conducted to isolate theeffect of eddies. The model is based on the message
passing interface (MPI) version of POM (Blumberg
and Mellor 1987; Jordi and Wang 2012; Oey et al.
2013). The domain covers the North Pacific Ocean
from 168S to 708N and from 988E to 738W at 0.18 3 0.18
horizontal resolution with 41 sigma levels. All surface
fluxes are zero. Climatology World Ocean Atlas(WOA) data from the National Oceanographic Data
Center (NODC; http://www.nodc.noaa.gov/OC5/WOA05/
pr_woa05.html) for temperature T and salinity S werefirst used to diagnostically integrate the model tosteady state, which yielded current fields in balancewith the corresponding densities. Starting from thebalanced state, the NCEP–NCAR wind climatology isconstantly applied to maintain the Kuroshio. The windis turned off west of 1308E to eliminate the local windeffect. The basic run is continued from the balanced
state. For the eddy cases, we inject a warm/cold eddy
with a radius of 150 km, centered at 218N, 1198E. The
eddy injection method follows that reported by Shaw
(1994) and Xu and Oey (2011). The warm/cold eddy is
gradually ramped up over a period of 10 days. To
conserve heat, the same heat is removed by specifying
a uniform upward/downward surface heat flux over themodel domain.

3. Resultsa. Signals other than seasonal
In this study, the x and y directions are rotated 308
clockwise from true north; positive y points along-strait
northeastward. Figure 1b plots the time series of the
Taiwan Strait mean along-strait current. The samplingarea is enclosed by the two dotted lines in Taiwan Strait(Fig. 1a). The along-strait current reveals strong seasonal variation as well as higher-frequency signals

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FIG. 3. EOF of sea surface elevation (m) for (a) mode 1 and (b) mode 2. (top) Eigenvectors and (bottom) PCs. Gray
contours in the top panels indicate 200- and 2000-m isobaths.

(Fig. 1b). The seasonal cycle is shown as the largest
peak in the power spectra analysis. The secondary
peaks capture the intraseasonal signals with periods
shorter than 180 days (Fig. 1c). W.-Z. Zhang et al.
(2014) suggested that the short-time fluctuation is induced by intense weather events such as storms andwinter wind bursts. In addition to wind, the effects ofeddies are considered in the present study. The 20-yr
mean eddy kinetic energy [EKE; (u0 2 1 y 0 2)/2] is plotted in Fig. 2. Slower varying signals are excluded byapplying a 180-day high-pass filter. The large EKE
appears to the southwest of Taiwan (Fig. 2a). This active eddy activity southwest of Taiwan was also reported by Nan et al. (2011), who used 17-yr satellite
altimetry data. The high EKE region, from 208 to 228N
and from 1188 to 120.58E, is chosen to represent theeddy activity following Nan et al. (2011). The time series of domain-averaged EKE is plotted in Fig. 2b. The
decadal signal shift at approximately 2002 may be related to Pacific decadal oscillation (PDO) (Soeyanto
et al. 2014). Time-scale variation longer than seasonal
is not considered in this study. The link between eddy

activity and the Taiwan Strait flow is next demonstrated through EOF and singular value decomposition(SVD) analyses.

b. EOF and SVD analyses
The 3.5-month (1 February–15 May 2012) EOF results reported by Oey et al. (2014) obtained the firstmode as the monsoon wind–driven, geostrophic mode.
The second mode was the shorter time-scale, windforced, fluctuating pressure gradient mode. We extended the concepts of Oey et al. (2014) by using EOF
analysis (Kutzbach 1967) with a larger spatial domain
and longer time scale. The EOF domain is set from 198
to 278N and from 1148 to 1228E to include the area ofhigh EKE southwest of Taiwan. The time series isevery 5 days for 20 yr, covering 1993–2012. The stericheight computed from temperature and salinity(Marshall and Plumb 2007) is removed. The high energy Kuroshio is masked out for focusing on the Taiwan Strait Current and eddies. Note that both EOF
and SVD analyses are conducted on the basis of theanomaly with the 20-yr mean removed. The leadingJUNE 2015

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FIG. 4. EOF of surface currents (m s21) for (a) mode 1 and (b) mode 2. (top) Eigenvectors and (bottom) PCs. Twoscales are used for shelf (red) and off-shelf (blue) regions.

modes of EOF on the sea surface height anomaly(SSHA) and surface current are shown in Figs. 3 and 4,
respectively. The first mode of SSHA and surface
current are both dominated by the seasonal signal(Figs. 3a and 4a, respectively). The first mode of SSHA
accounts for 62.7% of the total variance. The surface
current has a lower percentage, explaining 29.8% of
the total variance. The surface current is nearly geostrophic (Figs. 3a, 4a). This geostrophic relation explains the along-strait current moving southward in
winter and northward in summer (Figs. 3a, 4a). The
leading mode is consistent with that reported in earlier
studies such that the along-strait current flows southward under the northeasterly winter monsoon; during
the southwesterly summer monsoon, the current
moves northward. The following leading mode of
SSHA and the surface current reveals faster variationsin principal components (PCs) than those in seasonal
components (Figs. 3b, 4b). The mode-2 SSHA showsthat the warm eddy southwest of Taiwan corresponds
to the positive cross-strait SSHA gradient (Fig. 3b),which is balanced by the northward flow based on thegeostrophic relation. Conversely, the negative crossstrait SSHA gradient corresponding to the southwardcurrent is related to the cold eddy during the negative
PC. The mode-2 surface current also captures an eddy
southwest of Taiwan together with the along-strait
current in the southern strait (Fig. 4b). The currents tothe north are weaker than those to the south and are
not geostrophic. The geostrophic and seasonal variations, shown in the first mode of SSHA and the surface
current, agree with the first EOF mode reported by
Oey et al. (2014). The wind-forced fluctuated pressure
gradient mode reported by Oey et al. (2014) does notappear in the leading modes in our analysis, which may
be attributed to the different time scales between the
two analyses. The EOF analysis suggests a link between mesoscale eddies and the along-strait current;
SVD analysis will next show the coupling between
the two.
The analysis shown above indicates that the eddy
signal has faster fluctuation than seasonal. To focus on
the effect of eddies, the seasonal cycle is removed in thefollowing analysis. The SVD analysis (Bretherton et al.

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FIG. 5. (top) Mode-1 SVD for the (left) SSHA (m) and the (right) surface current (m s21). (middle) Expansion
coefficients for the SSHA (red) and surface current (blue). (bottom) The lag correlations between the two
expansion coefficients. The value exceeding the 95% significant level is plotted in solid black.

1992) applied on two variables seeks to identify the pairof coupled spatial patterns, with each pair explaining afraction of the covariance between the two fields. One of
the advantages in SVD analysis is that the two variables
do not require the same spatial grid, although they musthave the same time length. The currents inside the eddy
are stronger than the shelf currents induced by the eddy.
If the same area is used for both SSHA and shelf current,
the coupling is captured within the eddy itself (i.e., eddy’s elevation and its current field) rather than being a

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FIG. 6. (a) SSHA (m) taken from the black box in Fig. 2a. Composite (b) warm and (c) cold eddy events on (left)
day zero, (middle) day 10, and (right) day 25 are shown. Colors represent the composite SSHA (m), and vectors
represent the composite surface current anomaly (m s21). Shelf currents follow the black unit scale, and off-shelf
flows are represented by the green unit scale.

target of the shelf current. To address the coupling between the eddy and the shelf current, different domains
for the two variables are required. The domain of SSHA
is chosen off the shelf, south of 238N, where eddies areobserved. The area for the current is set on the shelf,
north of 218N, covering the Taiwan Strait. The firstmode of SVD accounts for 29.1% of the total variance,
as shown in Fig. 5. The SVD spatial patterns are similar
to mode 2 of the EOF (Fig. 4), showing a strong, warm
eddy southwest of Taiwan coupled with the northwardcurrent south of 238N. The current inside the Taiwan
Strait is weak and southward. The maximum coupling
coefficient is 0.48, with SSHA leading the shelf current

by 10 days. The EOF and SVD analyses reveal that themonsoon serves as a primary forcing to drive the seasonal shelf flow. The secondary effect by the mesoscaleeddies to the southwest of Taiwan is responsible for the
shorter time-scale fluctuation. SSHA EOF results showthat while a warm eddy appears to the southwest of
Taiwan, a positive cross-strait SSHA gradient is observed; the negative cross-strait SSHA gradient appears
during the existence of the cold eddy to the southwest of
Taiwan. Similar results are detected in the surface current in EOF and SVD analysis. However, the warm eddy
coupling with the northward current exists only in thesouthern strait; the current inside the strait is nearly

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FIG. 7. Lag correlation coefficient between eddy index, which is the domain-averaged SSHA of that shown in theblack box in Fig. 2a and SSHA. Values lower than 95% significant levels are in white. Minus n day indicates eddy
index lags of SSHA by n days.

quiet. The response in the strait with a 10-day time lag is
illustrated using composite analysis.

c. Composite events
The composite method is used to show the manner inwhich the along-strait current varies with the eddy insequence. The eddy index is first defined by taking thedomain-averaged SSHA (black box in Fig. 2, high EKE
region) with the seasonal cycle and interannual signal
removed. The eddy index (Fig. 6a) shows highfrequency fluctuation. The period of 90–100 days is determined through power spectra analysis. The maximum
eddy strength exceeding the criterion of 60.05 (m) is
chosen for the composite. We have detected 47 warm
eddy events and 44 cold eddy events exceeding thethreshold. Warm and cold eddy events are sorted into
two groups for the composite. The composite day zero isdefined as the time at which the eddy index reaches

maximum. Figures 6b and 6c show the composite SSHA
and the surface current anomaly, respectively. A warm
eddy is shown to the southwest of Taiwan on day zero(Fig. 6b, left panel). The along-strait current is southward in the northern strait, and the current west of theeddy moves northward. Then, 10 days later, the alongstrait current turns northward in the entire strait(Fig. 6b, middle panel). The warm eddy becomes much
weaker and dissipates on day 25, and a cold eddy appears to the south of the warm eddy. The along-strait
current remains northward near the shelf break, and thecurrent inside the strait becomes weaker (Fig. 6b, rightpanel). The case of the cold eddy is similar to the warm
eddy event, although the sign is reversed. The alongstrait current is northward on day zero, whereas the cold
eddy appears southwest of Taiwan (Fig. 6c, left panel).
On day 10, the eddy moves slightly northward, and thealong-strait current flows southward (Fig. 6c, middle

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FIG. 8. (a) Diagnostic SSHA (m, color) and surface current (m s21, white vector) from the idealized experiments.
The pink circle indicates the location of the injected eddy. The green line marks 1308E for the border of wind switch.
(b) Flowchart and list of experiments.

panel). The cold eddy weakens on day 25, and a warm
eddy is observed to the south. The current weakens andeven changes direction in the southern strait (Fig. 6c, rightpanel). The criteria of 0 (m) and 0.1 (m) are repeated forthe composite, and the conclusion remains the same.
To determine the source of the eddies, spatial lag correlation is conducted between the eddy index (Fig. 6a) andthe SSHA (Fig. 7). In this study, the seasonal cycle is removed to focus on the intraseasonal time scale. The lagcorrelation shows that the correlation increases locally.
The eddies in the intraseasonal time scale are locallyformed rather than propagating from the Pacific Ocean.

d. Mechanisms
To examine the mechanism of the eddies’ effect on the
Taiwan Strait Current, idealized experiments are conducted. The model details are given in section 2. The experiment list and flowchart are summarized in Fig. 8.
Figure 8a plots the diagnostic sea surface height and surface
current obtained from climatology. A preexisting warm
eddy appears southwest of Taiwan. To distinguish the effect of the injected warm/cold eddy to the Taiwan Strait,
the basic experiment (Exp_0; Fig. 8b) is subtracted fromthe eddy experiments (i.e., SSHAExp_w1 5 SSHExp_w1 2
SSHExp_0).
Figure 9a shows the Taiwan Strait transport anomaly(with Exp_0 subtracted) for the eddy experiments. The
transport is integrated from the sea surface to the bottom across 24.28N near the center of the strait. The

transport begins to change around days 15–20. Thechange in transport lasts until the injected eddy dissipates(approximately 40 days). The transport strengthens in thecase of the warm eddy. Conversely, the transport
weakens under the cold eddy conditions. A stronger eddy
is related to greater changes in transport.
We then examine the cause of the transport changewith the existence of eddies. As reported by Oey et al.
(2014, their appendix) for the case of the Taiwan Strait
Current, the linearized, depth-integrated [from bottomz 5 2h(x) to surface z 5 0] dynamical equations for ahomogeneous fluid are as follows:›h
1 t ox 2 t bx ,›x›hyfua h 5 2gh 1 t yo 2 t b , and›y

2f y a h 5 2gh›(ua h) ›(ya h)
1
5 0,
›x
›y

(1)
(2)
(3)

where x and y are in the cross-strait and along-strait
directions, respectively; h is the water depth; h is thesurface elevation; ua 5 (ua, ya) is the depth-averaged
velocity; f is the Coriolis parameter; to 5 (t ox , tyo ) is theykinematic wind stress vector; and tb 5 (t bx , tb ) is the
kinematic bottom stress vector, which is parameterized ast b 5 rua ,

(4)

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FIG. 9. Time series of (a) transport anomaly (Sv) and (b) sea level slope anomaly (m) from the idealized experiments. The x axis represents days. The inset shows the locations from which the transports were taken (24.28N; blackline) and sea level slope using north (dots north of 258N) minus south (dots south of 238N).

where r is a constant friction coefficient. In the Taiwan
Strait, the cross-strait bottom friction and wind stress are
negligible in comparison to along-strait components.
Equations (1) and (2) then respectively become
›h
, and
›x
›h
fu 1 ry a 5 2gh 1 t yo .
›y
2f y a h 5 2gh

(5)
(6)

The cross-strait momentum is the geostrophic balance. In
the idealized experiment, the wind stress is dropped
according to the model setting, and the cross-strait velocity
is weak. The along-strait momentum is approximately
ry a ’ 2gh

›h
.
›y

(7)

Next, the contributions from eddies and the large-scalepressure gradient are separated, and Eq. (7) is rewritten as!
›h›hr(y a_pg 1 y a_eddy ) ’ 2gh.
1›y pg ›y eddy(8)

By subtracting Exp_0, the anomaly in the eddy case isry a_eddy ’ 2gh›h.
›y eddy(9)

In Eq. (9), the along-strait current is proportional to theminus sign of the along-strait sea level slope. The
northward flow corresponds to a negative sea level
slope, and the southward current is associated with apositive sea level slope. The along-strait sea level slopes
from the eddy experiments are computed. Following
Wu and Hsin (2005), the sea level slope takes the meansea level at the northern strait (258–268N; inset in Fig. 9)
minus the mean sea level at the southern strait (228–
238N). The results show that the warm eddy induces anegative sea level slope, leading to the northward flow
[red curves in Fig. 9; Eq. (9)]. On the contrary, thepositive sea level slope is balanced by the southwardcurrent in the case of the cold eddy [blue curves in Fig. 9;
Eq. (9)]. A stronger (positive/negative) eddy corresponds to a larger sea level tilt, resulting in a stronger
along-strait flow [Fig. 9; Eq. (9)].

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FIG. 10. As in Fig. 9, but for JCOPE2.

The transport anomaly and sea level slope for
JCOPE2 are repeated following the same composite asthat shown in Fig. 6; results similar to those of theidealized experiments are shown in Fig. 10. When thewarm eddy approaches the strait, the sea level slope
becomes negative (higher sea level to the southernstrait) and the transport shifts northward (red curves inFig. 10). This process can last for 40–50 days until theeddy weakens. The case of the cold eddy results in the
opposite sign, which makes the sea level slope positive
and the along-strait current flow southward (bluecurves in Fig. 10). Note that with the existence of windstress in JCOPE2, the balance follows Eq. (6); thetransport and sea level slope are nearly out of phase.
Because of the modification from the wind stress, thevariation in transport does not closely follow thechange in sea level slope shown in the idealized
experiment.
The vorticity equation is next analyzed to determine
the mechanism wherein the eddy effects are transferred
onto the shelf to produce a sea level slope. Following

Xu and Oey (2011) and Oey et al. (2014), thedepth-averaged vorticity equation derived from thedepth-averaged momentum equations is





 
›z
f 1z
›h f 1 z
1
1U  $
2
5 k  $(x) 3 $
›t
D
›t
D
D
t 2 t 
o
b
1$3
,
D
(10)
where z is the curl of the depth-averaged velocity ua;
D 5 h 1 h is the total water depth; U 5 Dua is thetransport
vector per unit length; k is the z-unit vector;
Ð0x 5 2h zb dz; and b 5 g(r/ro ) is the buoyancy. By approximating D ’ h, the time-dependent terms for thecomposite are negligible, and Eq. (10) becomes
 

 
 
f
z
1
’ 2U  $1 k  $(x) 3 $
U$
h
h
h
t 
t 
1$3 o 2$3 b .
h
h

(11)

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FIG. 11. Composite (a),(c) PV flux and (b),(d) JEBAR for cases of (top) warm eddy and (bottom) cold eddy.
Units are 10211 s22. The superimposed black and white contours represent composite SSHAs (m). The contour
interval is 0.05 m. The gray curve represents the 200-m isobath.

The left-hand side of Eq. (11) is the potential vorticity(PV) flux, which represents the horizontal advection of
the geostrophic PV. The first term on the right-hand sideis the advection of relative vorticity. The second term isthe joint effect of baroclinicity and relief (JEBAR) term
(Mertz and Wright 1992). The third term is the curl ofdepth-distributed wind stress, and the last term is thenegative curl of depth-distributed bottom stress. All theterms are computed in the idealized experiments todetermine which ageostrophic term on the right-handside of the equation dominates the cross-isobath PV
flux. The 30-day periods with large sea level slopechanges are selected for the composite (Fig. 9). Days 27–
57 are chosen for the warm eddy case, and days 20–50
are selected for the cold eddy experiment. The standard

cases (Exp_w1 and Exp_c1) are demonstrated because
the eddy SSHA in these cases is similar to thatin JCOPE2.
Figure 11 plots the composite dominate terms, PV
flux, and the JEBAR term together with the composite
SSHA. Since all other terms are approximately one order of magnitude smaller, the figures are not shown. In
the case of the warm eddy, the water is more buoyant
toward its center. In the southern half of the eddy, $(x)points northeastward, and $(1/h) points onshore toward
the coast. The resulting JEBAR is positive (Fig. 11b),
which is balanced by the positive PV flux (Fig. 11a),corresponding to the onshore flow. Similarly, in thenorthern half of the eddy, $(x) changes direction topoint southwestward, and the resulting JEBARJUNE 2015

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FIG. 12. Composite SSHA (m) for (a),(b) days 11–15 and (c),(d) 30-day composite for cases of (left) warm eddyand (right) cold eddy. The contour interval for the black contours on the shelf is 0.005 m. The gray curve represents
the 200-m isobath.

becomes negative. The balanced negative PV flux,
therefore, leads to the offshore flow. Thereafter, theshelfward convergence and divergence produce anorthward sea level tilt (higher to the south). This phenomenon is shown in Fig. 12. On days 11–15, after the
injection of the eddy and before the transport begins tochange, a sea level anomaly dipole is detected on the
shelf, west of the eddy (Fig. 12a). The onshore and offshore movements continue to increase and decrease sea
level over time. The water then spreads out, and thesignal ultimately expands into the strait. With the effectof Coriolis force, the sea level contours tilt in southwest
and northeast directions (Fig. 12c). The sea level becomes higher in the south and east, leading to thenorthward transport (Fig. 9a).

We note that the sea level contour is concentrated to
the right of the strait, which is similar to the concept of
the arrested topography wave proposed by Csanady(1978). Combining Eqs. (3), (5), and (6) by eliminating u
and y, the equation for surface elevation is›2 h f ›h ›h1
5 0.
›x2 r ›x ›y(12)Equation (12) is identical to Eq. (3) in Csanady (1978).
The eddy imposes the sea level gradient at the shelf break,serving as the boundary condition. The solution therefore
has exponential decay away from the eddy, which resultsin the condensed sea level to the right side of the strait.
The same analogy applies in the case of the cold eddy. The

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FIG. 13. Composite Sea-Viewing Wide Field-of-View Sensor Chl-a (mg m23) for (a),(b) day zero and (c),(d) day 16
for cases of (left) warm eddy and (right) cold eddy. The gray curve represents the 200-m isobath.

center of the cold eddy is less buoyant. With the same
topography, the JEBAR sign is opposite that in the warm
eddy case (Figs. 11c,d). The resulting shelfward divergence and convergence produce a southward sea leveltilt that is higher to the north of the eddy (Fig. 12b). Days
later, the lower sea level south and east of the strait isbalanced by the southward current (Figs. 12d, 9a). The
right-side concentration signal is also apparent.
The stronger eddy cases (Exp_w2 and Exp_c2) show
similar results with more significant responses (Fig. 9).
The larger buoyancy gradient in the stronger eddy

results in greater JEBAR. The corresponding stronger
cross-shelf flow leads to a higher sea level tilt, which
later results in a more enhanced transport anomaly.
Sensitivity experiments are conducted to test theresults. The run with a perturbed initial condition leads
to the same conclusion. The other experiment applies
wind throughout the domain. Since the basic experiment (Exp_0_wind; with wind in this case) is subtracted to show the effect of the eddy, the cases withwind show eddy effects that are highly similar to thosein the original sets.

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4. Summary and discussion
The present study, based on the 20-yr JCOPE2 reanalysis data, explores the roles of eddies in modifying the
Taiwan Strait Current. Apart from the well-documented
monsoon-forced seasonal variation, eddy signals withhigher-frequency fluctuation are also observed. The effects of the eddy penetrate onto the shelf through the
JEBAR effect and eventually spread into a wide area ofcoastal water. The opposite signs of buoyancy gradientssouth and north of the eddy result in opposite directionsof cross-isobath fluxes, which lead to the shelfward
convergence and divergence and set up the modulationof the sea level slope. The warm eddy induces a northward sea level tilt, resulting in a northward, along-strait
flow tendency. Conversely, the cold eddy tends to produce a southward sea level slope, leading to a southwardtransport tendency. The stronger eddy results in a larger
modulation of cross-shelf flow and sea level slope,
producing a greater transport anomaly.
Eddies contribute to the changes in Taiwan Straittransport. Changes in ocean current can also influence
the biochemical environment. Sea-Viewing Wide Fieldof-View Sensor chlorophyll-a (Chl-a) data (O’Reilly
et al. 1998) are used to show the changes due to eddies.
These data were documented as overestimation in the
coastal region (Hyde et al. 2007). In the present study,
Chl-a is used as an indicator; the absolute value is notaddressed. The same composite method (Fig. 6) is applied. The dataset used for the composite has 9-km
resolution, every 8 days for 1998–2010; the seasonal
cycle is removed. On day zero (Figs. 13a,b), the composite shows negative (positive) anomalies southwest of
Taiwan, where the warm (cold) eddy is located. On day
16, Chl-a becomes less than that on day zero in the caseof the warm eddy (Fig. 13c). For the cold eddy composite,
Chl-a increases in the southern and central regions of thestrait (Fig. 13d). The chlorophyll concentration is generally higher in the northern China Seas (i.e., Yellow Sea
and East China Sea) than that in the South China Sea(Tang et al. 1998). The chlorophyll content in the Yangtze
River estuary is particularly significant. The warm eddy
leads to a northward along-strait current, carrying thewater of lower concentration from the south to reduce
Chl-a concentration in the strait. During the effective
period of the cold eddy, the resulting southward current
carries the water of higher concentration southward, enhancing Chl-a in the strait.
This research explains the mechanism by which
mesoscale eddies affect Taiwan Strait transport in an
intraseasonal time scale; the same concept can be used forseasonal or longer time scales. Considering the leading
seasonal variation in the Taiwan Strait, additional factors

are involved. The seasonal Kuroshio intrusion, winterChinese coastal water, and monsoon wind may all playroles in modifying the pressure field in the Taiwan Strait.
Detailed dynamics for seasonal variation are a topic
worthy of future study.
Acknowledgments. We thank the reviewers, whose
comments improved the manuscript. Useful discussion
with Dr. Tsubasa Kodaira is appreciated. AVHRR/MCSST, USGODAE project, CCAR near-real-time
altimetry data, and GTSPP are acknowledged for providing data for model assimilations. YLC is supportedby the Ministry of Science and Technology of Taiwan
under Grant 102-2611-M-003-003-MY3. XG and YM
are supported by JSPS KAKENHI (26287116).
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